uga-geosite-loope-yellow-knolls.indd utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors david b. loope earth & atmospheric sciences, university of nebraska lincoln, ne 68588-0340 dloope1@unl.edu cover image: southward-sloping crossbeds in the navajo sandstone at yellow knolls. hexagonal fracture patterns on navajo sandstone crossbeds at yellow knolls, washington county 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e geosites refl ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: loope, d.b., 2019, hexagonal fracture patterns on navajo sandstone crossbeds at yellow knolls, washington county, in milligan, m., biek, r.f., inkenbrandt, p., and nielsen, p., editors, utah geosites: utah geological association publication 48, 5 p., https://doi.org/10.31711/geosites.v1i1.62. presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president d.b. loope hexagonal fracture patterns at yellow knolls 3 introduction at this geosite, the main features of interest—remarkably uniform and beautiful fracture patterns dominantly composed of linked hexagons (fi gures 1 and 2)—are present on outcrops of the jurassic navajo sandstone. th e navajo was deposited by large, southward-migrating desert dunes about 200 million years ago, but the fractures that defi ne the hexagons here are just a surfi cial veneer less than 20 inches (half a meter) deep. th e fractures are a weathering phenomenon that developed under climate conditions similar to today’s. steep thermal gradients develop in the sandstone because it is exposed to solar radiation and changing air temperature. polygonal fracturing is present in other navajo exposures in southern utah, but only in non-bedded (homogeneous) rock. th e beautiful, bedding-parallel fracture pattern developed here is very rare; it developed because the bedding planes in the rock at yellow knolls are unusually wide-spaced. driving and hiking th e yellow knolls trail lies within the red cliff s natural reserve (http://www.redcliff sdesertreserve.com/yellow-knolls). th ere is a wellsigned parking area along the unpaved portion of cottonwood springs drive 5.5 miles (9 km) north of east red hills parkway (fi gures 3 and 4) on the right side of the road. a good trail leads to an excellent outcrop (fi gure 1) about 1.2 miles (2 km) from the parking area. gps locations: parking area: n37°11’24”, w113°34’42” (wgs84). destination (fi gure 1): n37°12’15”; 113°34’23.6” general geology two very diff erent rock types are exposed along the trail leading away from the yellow knolls parking lot. first, the trail winds through a cluster of large black boulders. th ey are basalt—the boulders were eroded from broad, 1to 2-million-year-old lava fl ows that spewed from a volcanic vent 2.9 miles (4.6 km) northwest and upslope from the trailhead (fi gure 4). th e hot lava followed stream courses that had been cut into a jurassic sandstone formation called the navajo (the tan rock with lots of fractures and bedding planes). th e geologic map (fi gure 4) shows that the sandstone forming yellow knolls is surrounded by basalt; the sandstone was buried by lava, and now lies in an erosional “window” that was cut by ephemeral streams in the last million years—these were likely the same streams that had cut the valley that the lava followed. because the navajo sandstone was deposited by migrating sand dunes, its sand layers were not laid down parallel to the fl oor of the jurassic desert. sand instead was deposited on the downwind slopes of the dunes, at an angle of about 30° (the angle of repose for dry sand). as the desert fl oor slowly subsided, the wind kept blowing southward, and the dunes piled up sloping layers as they climbed over one another, burying thick crossbeds. and, also importantly, rivers kept delivering more sand. although most of the grains in the navajo are abraded crystals of quartz, there are also some abraded zircon crystals. because each zircon crystal contains figure 1. domed hexagons developed parallel to bedding in the navajo sandstone at yellow knolls. field of view is about eight feet (2.5 m) wide. figure 2. southward-sloping crossbeds in the navajo sandstone at yellow knolls. th e exposed surfaces of many diff erent beds are covered with highly ordered, polygonal fracture patterns. underlying, unexposed beds remain unfractured. n37°12’11.9”, w113°34’18.4” wgs84 113°33'00" w wgs84 113°33'00" w 113°34'00" w 113°34'00" w 113°35'00" w 113°35'00" w 113°36'00" w 113°36'00" w 37 °1 1' 00 "n 37 °1 1' 00 "n 37 °1 2' 00 "n 37 °1 2' 00 "n 37 °1 3' 00 "n 37 °1 3' 00 "n 37 °1 4' 00 "n 37 °1 4' 00 "n map created withmap created withmap created withmap created withmap created with topo!®topo!®topo!®topo!®topo!® ©2009 national geographic©2009 national geographic©2009 national geographic©2009 national geographic©2009 national geographic topo! map printed on 02/27/19 from "topo 1.tpo" 02/27/19 tn mn 11° yellow knolls cinder cones pleistocene lava �ow pl ei st oc en e la va � ow yellow knolls trailhead (parking) 1 2 km 3 0.5 1.0 1.5 mi 0 red hills parkway ~ 4 miles (6 km) n figure 3. topographic map of the yellow knolls vicinity showing the location of the trailhead and parking area. red star shows location of photo in fi gure 1. 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 uranium, these crystals can be precisely dated. conclusions from this dating are: 1) as expected, the zircons are much older than the jurassic dune fields, and are about one billion years old; 2) they came from granite and gneiss in the appalachian mountain chain in what is now eastern united states; 3) they were delivered to the navajo dune field by westward-flowing rivers (dickinson and gehrels, 2003). fracture patterns on the navajo sandstone the patterns contain polygons with four, five, six, seven, eight, or more sides, but they are dominated by hexagons (figure 5). the patterns are developed on many of the southward sloping crossbeds of the formation, and the fractures that define most of the polygons are only about 6-10 inches (15-25 cm) deep. chan and others (2008) described polygonal fractures of similar length and depth that are largely confined to outcrops of massive (unbedded) navajo sandstone. when the cracks formed, the rock surface must have been flat. through time, weathering along the cracks has given each polygon a convex, dome-like upper surface (chan and others, 2008). there are virtually no extensive flat (unfractured) surfaces visible at yellow knolls—all the broad surfaces that could crack, did crack long ago, and all polygons are now domed. because new, unfractured bedding surfaces are hard to find, it is difficult to tell if (or how often) new fractures form in the present climate, but because the fracture patterns follow many of the small-scale erosional features here, some of them must be quite young (no more than a few hundred years). highly ordered patterns—like those displayed by daisies and leopards—are common in nature, and humans (even geologists) find them aesthetically pleasing. highly ordered fracture patterns develop only in homogeneous material (materials without preferred lines of weakness) like mud, lava, granite, and some sandstone. for example, as lava flows solidify and then continue to cool, they commonly break into uniform, 5-, 6-, and 7-sided columns. vast areas in the arctic, where the average air temperature is below the freezing point of water, are underlain by permafrost (frozen st. george yellow knolls trailhead cottonwood springs drive green springs dr. red hills pwy 2 km jna jna exit 10 jc ki i-15 ut-18 las vegas salt lake city neogene basalt cretaceous iron springs fm jurassic carmel fm jurassic navajo ss jna nb ki jc vent erosion figure 4. simplified geologic map of yellow knolls site showing distribution of bedrock and access routes. 5 meters 7 5 66 5 7 5 75 6 5 6 6 6 75 7 5 6 6 6 6 6 6 6 6 6 6 7 5 66 5 6 5 5 5 5 5 5 5 5 5 5 5 5 5 5 56 6 66 6 6 6 6 6 6 6 6 6 6 6 6 6 4 7 7 7 77 7 7 77 7 7 7 7 7 7 7 7 7 7 7 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 66 6 g g g g g g g g g g g g g g g g g gg g g g g g g g g g g g g g g g g g g g g g g g g g g g g 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 6 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 7 5 5 5 5 5 5 5 5 5 5 5 5 5 5 5 5 5 5 5 8 8 8 88 8 8 8 8 8 8 8 8 8 8 8 6 6 6 6 6 6 6 6 6 6 6 65 5 5 5 5 5 5 8 8 8 8 8 7 7 7 7 7 7 6 6 6 6 6 6 6 6 66 6 6 g g g g g g g g g g 4 ?8765 number of sides ! !" 4 5 7 8 number of sides n = 303 6 shapes of polygons mapped in (a) ? area covered 10 20 40% 30 0 ba n 24 n 37° 12’ 15.54” w 111° 34’ 23.7” figure 5. polygon shapes drawn from aerial photo that was oriented perpendicular to a rock surface that slopes 24° toward the south-southwest. the fractures that define the polygons shown in light gray (?) are partially erased, probably due to erosion down to a level below the base of some (but not all) of the fractures. d.b. loope hexagonal fracture patterns at yellow knolls 5 arctic-like, sub-zero temperatures are probably not responsible for utah’s cracked rocks; here, the steepest temperature gradients in the sandstone likely develop during summer nights. after a long, sunny june or july day, the surface temperature of the navajo sandstone (a brittle material) gets very hot, and much of that heat flows into the rock. before dawn the next day, much of that heat has flowed back out of the shallowest part of the rock, but the rock at depth is still hot. like most rocks, sandstone is strong under compression, but weak under tension (stretching; lachenbruch, 1962). this quick change of air temperature generates enough contractile (tensional) stress in the exposed rock to fracture it. fracture patterns caused by this process are widespread on sandstone in southern utah (chan and others, 2008), but they are especially well exposed, very uniform, and quite beautiful at this site. but is the navajo sandstone at this site really homogeneous? the fact that it is bedded means it cannot be completely homogeneous—instead, the rocks here are repeatedly homogeneous. bedding is caused by abrupt variations in the size of sediment grains. in this case, the sand grains avalanching down the dune slopes were coarser than those falling onto the dune slope. the hexagonal fracturing here is possible because the bedding at yellow knolls is thicker and not as distinct as at other navajo outcrops (loope and others, 2019). the greater thickness allows polygonal fracturing because bedding surfaces (planes of weakness) are few and far between. when beds are thin, tensile stresses due to a thermal gradient cannot built up because sliding parallel to the stress occurs along the shallowest bedding plane and prevents rupture—in this case, “nothing has to give”. at yellow knolls, stacks of hundreds of (sufficiently) thick beds (figures 2 and 6) allowed tensile stresses to build to the point of rupture, producing linked hexagons in each one. not all of the thermally induced fractures here are polygonal. the fractured surfaces that are not aligned with bedding commonly produce a checkerboard pattern (figure 6; chan and others, 2008). in this situation, the bedding planes act as slippage surfaces that prevent fracturing; only fractures oriented perpendicular to the bedding planes can open—that’s what makes the checkerboard pattern. the beautiful hexagonal patterns developed on south-facing slopes that, like solar panels, gather optimum radiant energy. those slopes might seem the best place for thermal fractures to develop. a checkerboard pattern of similar scale, however, develops even on north-facing slopes at yellow knolls (figure 6), and those at checkerboard mesa in zion national park (biek and others, 2010; chan and others, 2008). this raises the question: with all the available solar energy and the resulting steep temperature gradients, why would any broad navajo sandstone exposure in southern utah remain unfractured? it is likely that some navajo outcrops are not strong enough to sustain large-scale fracture because tensional stress is relieved by failure at a much smaller (grain-to-grain) scale. ground) up to hundreds of feet thick. the key thing to getting fractures is to develop a strong temperature gradient within solid, brittle material. on winter nights in the arctic, when the air temperature drops far below the average air temperature, the frozen land surface gets much colder than the frozen material at depth. because the surface material is solidly connected to the deeper material, when the surface contracts “something has to give”, and the land surface ruptures, forming polygons. thermal fracturing in southern utah land surfaces underlain by homogeneous, brittle rock can also break into polygons when they are subjected to cooling. the average temperature of rocks, soil, and groundwater within a few meters of land surfaces is close to the average temperature of the air above them, but the temperature of rock cannot change as rapidly as the temperature of air can change. this means the temperature of rocks at the land surface will be out of phase with the temperature of deeper rock, generating a thermal gradient. figure 6. checkerboard pattern composed of bedding planes and surface fractures. the sloping rock surface under the man’s feet is covered with hexagons (that sloping rock surface is among the ones shown in figure 2, but the camera is here facing in the near-opposite direction). the middle 60% of this photo shows a steep, north-facing slope, and the crossbeds are sloping away from the viewer. note that in this view, bedding planes are represented by nearly horizontal, widely spaced lines, and because they are unexposed, they bear no fractures. the checkerboard pattern develops where there is a big difference between the orientations of bedding and the exposed rock surface. the fractures on this north-facing slope also demonstrate that maximal insolation is not required for thermal fracturing. n37°12’15”; w113°34’24” 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 acknowledgments clint rowe and jon mason provided helpful reviews, and gerald bryant assisted with field work and provided the photograph leading to figure 5. references biek, r.f., willis, g.c., hylland, m.d., and doelling, h.h., 2010, geology of zion national park: in sprinkel, d.a., chidsey, t.c., and anderson, p.b.,(editors) geology of utah’s parks and monuments: utah geological association & bryce canyon natural history association publication 28, third edition, p. 109-143. chan, m.a., yonkee, w.a., netoff, d.i., seiler, w.m., and ford, r.l., 2008, polygonal cracks in bedrock on earth and mars: implications for weathering: icarus, v. 194, p. 65–71. dickinson, w.r., and gehrels, g.r., 2003. u-pb ages of detrital zircons from permian and jurassic aeolian sandstones of the colorado plateau, usa: paleogeographic implications: sedimentary geology, v. 163, p. 29–66. lachenbruch, a.h., 1962, mechanics of thermal contraction cracks and ice-wedge polygons in permafrost: geological society of america special paper 70, 69 p. loope, d.b., burberry, c.m., and searles, m.l., 2019, thermal contraction cracks in utah slickrock: geological society of america; rocky mountain section meeting, manhattan, ks. uga-geosite-spangler-ricks-spring.indd utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors lawrence e. spangler u.s. geological survey 2329 orton circle, west valley city, utah 84119 spangler@usgs.gov cover image: ricks spring rises along a nearly vertical normal fault in the ordovician garden city formation. ricks spring utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e geosites refl ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: spangler, l.e., 2019, ricks spring, in milligan, m., biek, r.f., inkenbrandt, p., and nielsen, p., editors, utah geosites: utah geological association publication 48, 6 p., https://doi. org/10.31711/geosites.v1i1.64. presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president l.e. spangler ricks spring 3 introduction ricks spring is one of several major karst springs that discharge along the logan river in the bear river range in cache county, utah. the spring is located along u.s. highway 89 in logan canyon about 17 miles (27.4 kilometers) northeast of (up-canyon from) the city of logan, at mile marker 477. it lies within uintawasatch-cache national forest at an elevation of 5880 feet (1792 meters). situated at the base of a hillside, the spring is one of the largest and most scenic along the logan river (figure 1). water from the spring flows out of a large alcove, under highway 89, and into the logan river, about 150 feet (45 meters) from the spring. pullouts on both sides of the highway provide parking for visitors to the spring, and a boardwalk crosses the spring run, which allows access to the rise pool in the alcove. several signs at the spring provide information about its history and hydrology. ricks spring typically flows during the spring, summer, and fall months, but can have periods of no flow during the winter months, particularly during extended periods of cold weather. during these times, the water in the rise pit recedes to a small pool of standing water that can contain fish, which presumably originate from the nearby logan river. during the last 10 years, cave divers have explored the conduit that feeds the spring for about 2300 feet (700 meters) into the mountainside. coordinates: 41° 50' 24.55" n, 111° 35' 19.15" w (nad 83) history ricks spring is named for thomas e. ricks, a mormon pioneer who settled in cache valley in 1859. a colonel in the utah militia, ricks was commissioned by brigham young to locate a better route from cache valley over the mountains to bear lake valley. consequently, ricks and others in the community began constructing a road to bear lake valley up logan canyon. the road to ricks spring was the first section to be completed. ricks’ explorations of logan canyon were some of the first documentations of the area, and it was during this time that he discovered a “natural spring flowing from the cavity of a large rock” (web page accessed december 27, 2018, at https://en.wikipedia.org/wiki/thomas_e._ricks_ (mormon)). since its discovery, people visiting the spring have often drank from it, bringing home jugs and barrels of the “fresh spring water” under the assumption it came from a deep aquifer. they often became sick with giardia, indicating the water was not from an artesian source (web page accessed december 27, 2018, at https://en.wikipedia.org/wiki/ricks_spring). early 1900s postcards showed visitors in their sunday best dipping water from the spring, unaware that part of the water from the spring came from the logan river just upstream. at one time a water fountain was present at the spring; however, this water was sourced from a different spring nearby. that, along with an outhouse, were removed in the 1980s (scott bushman, u.s. forest service, retired, verbal communication, june 7, 2019). in the 1990s, a sign was still posted at the spring advising visitors that giardia was present in the water. geology ricks spring discharges from the ordovician garden city formation, a thin-bedded, cherty, dolomitic, argillaceous (clay-bearing) limestone about 1250 feet (380 meters) thick in the vicinity of the spring (dover, 1995). the garden city crops out in the area surrounding and immediately to the north of ricks spring, but farther north, the formation is covered by (1) the tertiary (eocene) wasatch formation, a reddish conglomeratic siltstone with interbedded limestone and marl, (2) quaternary and (or) upper tertiary unconsolidated conglomerate (diamictite), and (3) pleistocene glacial moraine deposits consisting of poorly sorted boulder till (dover, 1995; figure 2). regionally, the garden city formation crops out on the eastern flank of the logan peak syncline, as indicated by mapped bedrock dips of 10 to 20 degrees to the southwest in the vicinity of the spring (dover, 1995; figures 2 and 3). overlying silurian dolomites and devonian carbonate and clastic (shales and sandstones) rocks crop out west of the garden city and form the core of the syncline (figures 2 and 3). ricks spring rises along a northeast-trending normal fault (herein termed the ricks spring fault) that is unmapped but roughly parallels the logan river. the fault appears to extend to the northeast from the alcove where it presumably intersects the logan river upstream of the spring and could act to divert water from the river to the spring. a mapped fault that also trends northeast and parallels the rick springs fault, lies just north of the spring and is shown to intersect the logan river (figure 2). dip along this fault is about 80 degrees to the southeast, with about 200 feet (60 meters) or less of displacement within the garden city formation, based on a geologic cross section that traverses the fault southwest figure 1. ricks spring rises along a nearly vertical normal fault (dashed red line) in the ordovician garden city formation. bedding dips to the southeast at about 36° near the fault. the small cave on left side of photo could have been a former spring outlet. photograph by larry spangler, august 2011. figure 3. generalized geologic cross section showing the location of ricks spring in relation to the logan peak syncline and the mapped fault near the spring (arrow shows relative direction of movement). cross section modified from dover (1995). refer to figure 2 for location of cross section b–b’ and to explanation for geologic units and symbols. figure 2. geologic map of the area surrounding ricks spring. much of the outcrop area of the ordovician garden city formation is covered by the tertiary wasatch formation and quaternary unconsolidated deposits. note the northeast-trending fault near the spring. map modified from dover (1995), with dye-tracing results (red arrows) from spangler (1991). green arrows indicate possible flow paths based on information from u.s. forest service and fault projections. blue dots represent spring discharge points. l.e. spangler ricks spring 5 discharge and temperature discharge of ricks spring is highly variable like most karst springs in snowmelt-dominated terrains, with low flow (base flow) during the winter months and peak flow during the snowmelt runoff period from may to june. discharge of the spring normally ranges from less than 1 cubic foot per second (ft3/s) or about 450 gallons per minute (0.028 cubic meter per second [m3/s]) to 75 ft3/s (2.1 m3/s) (mundorff, 1971), although estimates as high as 150 ft3/s (4.25 m3/s) have been reported (wilson, 1976). in a thesis completed at utah state university in 2018, skyler sorsby (written communication, december 2018) reported monthly measurements of springflow from the alcove ranging from 1.77 ft3/s (0.05 m3/s) in november 2016 and march 2017, to a peak flow of 109.5 ft3/s (3.1 m3/s) in june 2017, during the snowmelt runoff. during high flow, a boil as much as 1-foot (0.3 meter) high forms on the surface of the pool from the upwelling water (figure 5). clarity of the water at high flow is exceptionally good, reflecting the input of snowmelt into the aquifer that supplies the spring. during snowmelt runoff, most water discharges from the alcove (figure 6), and the estimated combined flow from the other springs appears to be no more than 2 to 3 ft3/s (0.057 to 0.085 m3/s). during winter, when little or no flow occurs from the alcove (figure 4), base flow of the spring is maintained by the lower elevation or underflow springs, which discharge less than 1 ft3/s (0.028 m3/s). of the spring (dover, 1995, section b-b'; figures 2 and 3). the ricks spring fault could be a branch off the mapped fault or a localized separate fault. displacement along the ricks spring fault is unknown but is also within the garden city formation, appearing to phase out to the southwest. bedding along the fault in the spring alcove where it is dramatically exposed (figure 4), strikes about n 28° e and dips to the southeast at about 36 degrees. hydrology ricks spring discharges primarily from a prominent alcove as well as from three additional outlets: one along the base of the hillside on the north side of the highway, about 380 feet (115 meters) southwest of the main spring, and two along the river on the south side of the highway, about 560 feet (170 meters) and 740 feet (225 meters) from, and at a slightly lower elevation than, the main spring (figure 2). flow in the alcove is from a funnel-shaped rise pit that is about 13 to 14 feet (4 meters) deep. during the winter months water may not rise high enough to flow from the rise pit. throughout the rest of the year water fills the rise pit and overflows into a channel that conveys the water from the alcove to the logan river. a small cave adjacent to the alcove, but higher than the spring (figure 1) can be explored for about 25 feet (7.5 meters). the cave is always dry but could have been a former spring discharge point or outlet at some point in the geologic past when groundwater levels and the logan river were higher. figure 4. flow from ricks spring often ceases during the winter, leaving only a standing pool of water in the rise pit (not visible). spring rises along normal fault (dashed red line) in alcove. photograph by larry spangler, march 2002. figure 5. boil on surface of rise pool during the snowmelt runoff period in may– june. water is typically very clear. photograph by larry spangler, june 2017. figure 6. ricks spring at high flow during the spring runoff of 2017. estimated discharge is about 100 ft3/s (2.8 m3/s). photograph by larry spangler, june 2017. temperature of water from ricks spring is higher during low-flow conditions and lower during high flow. on the basis of periodic measurements made over a 3-year period (may 1994 to july 1997), and likely representative of the annual cycle, the temperature of water discharging from the main alcove spring ranged from 5.5 degrees celsius (°c) (41.9 °f) during the snowmelt runoff (may– june) when discharge is highest, to about 7.0 °c (44.6 °f) during low-flow conditions (late fall to late winter) (spangler, 2001). the lower temperature during high flow results from the input of snowmelt into the aquifer from losing streams and possibly from the snowmelt-fed river water. the snowmelt moves rapidly through the aquifer along solution-enlarged joints, faults, and (or) bedding planes to the spring. during low-flow conditions, groundwater stored in the aquifer matrix (tighter fractures) that has had a longer residence time dominates spring discharge, and water temperatures are slightly warmer. sources of water in the 1950s hydrogeologists noticed a pattern between the flows of the logan river and ricks spring, and theorized a connection between them. during the winter of 1972, severe cold weather froze the river, resulting in ice jams in some locations, which caused the water to back upstream. subsequently, ricks spring began to flow and then subsided when the river level dropped again, suggesting a link between the river and the spring. later that summer, dye was added to the logan river and was subsequently observed in ricks spring, confirming the connection between the river and the spring (u.s. forest service, logan district ranger, verbal communication, 2010; web page accessed december 27, 2018, at https://en.wikipedia.org/wiki/ricks_spring). morgan (1992, p. 17) in a geological road guide of northern utah, recognized that the river may not be the only source of springflow and stated that “there is a direct relationship between the spring’s increased discharge and snowmelt in the spring; therefore, the source of the water is most likely local spring runoff which probably seeps into the ground in the tony grove basin area to the northwest. it is possible that some of the spring’s water may also come from the logan river, since a fault intersects the river upstream from the spring and trends directly toward the spring.” beginning in 1991, spangler (2001) began further investigations into the source of water for ricks spring. results of these investigations, based on dye-tracer tests, showed that in addition to the river, water also originated from bear hollow, northeast of the spring, tony grove creek near tony grove lake, and from bunchgrass creek, north of the spring (figure 2). results of the dye tracing indicated that the source area or groundwater basin for ricks spring extends more than 5 miles (8 kilometers) to the northwest in the tony grove lake area and 2600 feet (790 meters) higher than the spring. maximum groundwater travel time from the dye-injection points to the spring was about 4 weeks during low-flow conditions. because the recharge area for the spring is largely underlain by the garden city formation, groundwater movement appears to be confined within and is generally to the southeast through this formation to the spring. although these results indicate that water originates from higher on the mountain as well as from the river, the ratio or proportion of these sources to the total flow is unknown and could be highly variable depending on flow in the river. in addition, the location of the input point(s) along the logan river is unknown. both faults appear to intersect the river upstream of the spring and either or both could act to divert river water into the ricks spring flow system. exploration of the spring the submerged conduit feeding ricks spring was first accessed by a team of utah-based cave divers (wendell nope, richard lamb, and tibby petrescu) in 2007 (figure 7). since that time, dozens of dives have pushed the known length of the underwater cave to about 2300 feet (700 meters) (figure 8). along the way, several air-filled rooms were encountered, including one at the end of the explored cave that contains a 28-foot-high (8.5 meters) waterfall cascading from an upper level passage. the largest dry room encountered in the cave is about 260 feet (79 meters) long, 40 to 50 feet (15 meters) wide, and 50 to 70 feet (20 meters) high (figure 8). although it has been documented that the river provides part of the flow to the spring, the overall trend of the underwater passage is due north (away from the river) for almost 500 feet (150 meters) before turning and heading generally southwest and west. the remaining part of the cave heads northwest for approximately 400 feet (120 meters), then abruptly turns to the northeast for the last 270 feet (82 meters) of the known cave. observations made by the cave divers indicate a side passage off the main passage about 500 feet (150 meters) from the entrance rise pool that could be the infeeder from the logan river (figure 8, “river intrusion tunnel”). figure 7. richard lamb (left) and wendell nope rising from the depths of ricks spring. photograph by larry spangler. l.e. spangler ricks spring 7 warmer water temperatures in the side passage (more than 1°c; 1.8 °f) in comparison with water in the main passage along with an increase in sediment, also indicate a possible connection with the logan river. for much of the mapped portion of the cave, the passage lies below the surface of the rise pool in the spring alcove, with a maximum surveyed depth of 85 feet (26 meters) (fi gure 8). in addition, underwater surveying has shown that before rising to the surface in the alcove, the water descends to a depth of about 68 feet (21 meters) below the rise pool, and thus, below the level of the adjacent logan river. th e initial discovery and exploration of ricks spring cave are summarized in an article published in the national speleological society (nss) cave diving section’s underwater speleology newsletter in may/june 2008. th is article as well as the history of ricks spring exploration and videos of the underwater cave can be accessed at the website www.wendellnope.com/scuba3.htm. acknowledgments th e author would like to thank ryan rowland and tom marston, u.s. geological survey, and doug sprinkel, utah geological survey, for their benefi cial reviews of this paper. jen miller, utah geological survey, draft ed the explanation for the geologic fi gures, which is much appreciated. th anks are also owed to mike hess, u.s. geological survey, for his fi eld measurements of bedding attitudes at ricks spring. scott bushman, u.s. forest service, retired, off ered useful insights into the history of ricks spring. references dover, j.h., 1995, geologic map of the logan 30' x 60' quadrangle, cache and rich counties, utah, and lincoln and uinta counties, wyoming: u.s. geological survey miscellaneous investigations series map i-2210, 1 pl., scale 1:100,000. morgan, s.k., 1992, geologic tours of northern utah: utah geological survey miscellaneous publication 92-1, 98 p. mundorff , j.c., 1971, nonthermal springs of utah: utah geological and mineralogical survey water-resources bulletin 16, 70 p. spangler, l.e., 2001, delineation of recharge areas for karst springs in logan canyon, bear river range, northern utah: u.s. geological survey water-resources investigations report 01-4011, p. 186–193, available online at https://water. usgs.gov/ogw/karst/kigconference/proceedings.htm. wilson, j.r., 1976, glaciated dolomite karst in the bear river range, utah: salt lake city, university of utah, department of geology and geophysics, ph.d. dissertation, 122 p. figure 8. generalized map of the underwater cave at ricks spring. several parts of the cave are above water, while the deepest underwater section reaches a depth of 85 feet (26 meters) below the surface of the rise pool. map used with permission of wendell nope. uga-geosite-chidsey-fold-train.indd 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors thomas c. chidsey, jr.1, and paul b. anderson2 1utah geological survey, po box 146100, salt lake city, utah 84114-6100, tomchidsey@utah.gov 2geologic consultant, po box 101, emery, utah 84522 cover image: typical fold core showing contorted bedding, brecciation, and heterogeneous lithic component. spectacular crinkled crust—a detachment fold train in the carmel formation, western san rafael swell, utah 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 t.c. chidsey, jr., and p.b. anderson detachment fold train in the carmel formation 3 introduction imagine slipping on a small rug overlying a hardwood floor. in the process of sliding along the floor the rug produces a series of small folds and the rug moves forward from its original position. the same could be said for the “crinkled crust,” or folded layers of rocks in a detachment fold train. a spectacular detachment fold train, consisting of over 100 small, regularly spaced convex-upward folds called anticlines in gypsum-rich rock layers of the middle jurassic (about 168 million years ago [ma]) carmel formation, is exposed immediately north of interstate 70 (i-70) in the san rafael swell of east-central utah (figures 1 and 2). the san rafael swell, a large anticlinal uplift, is an icon for everything that makes the colorado plateau dramatically scenic and geologically classic. however, the fold train is located in drab-colored, relatively featureless rock layers of the carmel formation in an area called reed wash along the gently dipping west flank of the swell. after passing magnificent canyons, buttes, and mesas both to the east and west along i-70, the fold train typically goes unnoticed by not only the average tourist but geologists as well. once the fold train is pointed out, the geologic observer is immediately struck with awe at this large, well-exposed, complex structural feature. literally hundreds of classic geologic sites are well displayed in the san rafael swell; many are easily accessed overlooks and viewpoints. the detachment fold train, by contrast, is chosen as a geosite for its geologic uniqueness, educational instruction, and research opportunities in structural geology. how to get there the detachment fold train is about 190 miles (300 km) or a 3-hour drive from salt lake city, utah, via i-15 and u.s. highway 6 to state highway 10. stay southbound on state highway 10, past the town of ferron for 5.3 miles (8.5 km) and make a left turn towards the town of moore onto moore road. follow the road until a “t” in the road (about 3.6 miles [5.8 km]), turn left (east) and stay on the paved road (county road 803) until it intersects with i-70, about 16.5 miles (26.6 km) beyond the “t.” enter i-70 heading westbound for about 4.6 miles (7.4 km) to the best of several views (to the north) of the detachment fold train (38°50'49" n., 110°59'57" w., elevation 6325 feet [1928 m]) (figure 1). the fold train geosite is about 46 miles (74 km) west of the junction of u.s. highway 6 and i-70 near the town of green river, utah (figure 2). upon approaching the reed wash area (there is no sign so use the odometer or the gps coordinates) carefully slow down with flashers on and pull off the interstate to the right as far as possible. depending on the exact location, the view area may require climbing up a small but steep hill adjacent to the interstate. the majority of the lands in the reed wash area are public and overseen by the u.s. bureau of land management, with some owned by the state of utah school and institutional trust lands administration, and therefore access to examine the fold train up close is open to the public; however, there are no trails; the terrain is rugged and requires negotiating a fence designed to keep livestock and wildlife from the interstate. geologic summary of the san rafael swell the san rafael swell is a broad, asymmetric, north-southto southwest-northeast-trending anticlinal structure, about 75 miles (120 km) long and 35 miles (56 km) wide, that formed in response to compressional forces of the laramide orogeny (a regional mountain-building event) between latest cretaceous time (about 70 ma) and the eocene (about 40 ma) (hintze and kowallis, 2009 and references therein) (figures 3 and 4). uplift and erosion have made it a showcase of colorado plateau geology with a colorful array of sedimentary rocks over 7000 feet (2100 m) thick, ranging in age from permian to cretaceous (299–66 ma) and exposed in spectacular cliffs along cuestas, mesas, and deep canyons (figure 5 represents those on the west flank of the swell where reed wash is located). figure 1. spectacular detachment fold train in the winsor member of the middle jurassic carmel formation in the reed wash area on the west-dipping flank of the san rafael swell; view north from i-70. photograph by michael chidsey, sqwak productions inc. 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 figure 2. the san rafael swell and vicinity, east-central utah, showing the location of the detachment fold train geosite as well as major physiographic features, surrounding towns, and highways. t.c. chidsey, jr., and p.b. anderson detachment fold train in the carmel formation 5 figure 3. generalized geologic map of the san rafael swell and location of the detachment fold train geosite within the middle jurassic carmel formation. cross section a-a' shown on figure 4. after doelling and hylland (2002). figure 4. diagrammatic cross section across the middle of the san rafael swell. the cross section is not drawn to scale, but the vertical dimension is exaggerated about eight times relative to the horizontal; the horizontal length of the cross section covers about 50 miles (80 km). symbols and colors of geologic formations correspond to those shown on figure 3; location of cross section also shown on figure 3. after doelling and hylland (2002). 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 the sedimentary formations and their many members exposed in the san rafael swell were deposited in a wide range of environments including eolian (windy desert), floodplain, stream, deltaic, swamp, tidal flat, and shallow and restricted marine (figure 5). several major unconformities represent significant periods of erosion or non-deposition. pliocene-age igneous rocks are present in the form of dikes, conduits, and sills intruded into exposed triassic to cretaceous sedimentary strata (figure 3). the rocks in the san rafael swell are folded, faulted, jointed, fractured, and uplifted. the major uplift and deformation of the san rafael swell was likely controlled by a large, blind (buried), basement-involved reverse fault (up on the west side) bounding the east flank of the structure (figure 4). three sets of high-angle normal faults are mapped on the surface: (1) northwest-southeast striking, (2) east-west striking, and (3) north-south to northeast-southwest striking (figure 3). two styles of reverse faulting are identified in the san rafael swell: (1) west-directed, blind reverse faults on the east flank, and (2) east-directed, ramp-style thrusting. sandstone beds are quartz rich and brittle, and when folded or bent, produce prominent joints and fractures. uranium, oil and gas, carbon dioxide, helium, coal, and industrial minerals (gypsum, bentonite clay, and humate) are found within the san rafael swell. gypsum (caso4·2h2o) was produced from one mine in the carmel formation near the detachment fold train and was used for manufacturing wallboard (sheetrock) and plaster (gloyn and others, 2003). the carmel contains an estimated 7.3 million tons of minable gypsum (lupton, 1913). cumulative gypsum production in the san rafael swell since 1990 is about 1.8 million tons; there is one active mine on the north end of the structure (verbal communication, andrew rupke, utah geological survey, 2018). stratigraphy: middle jurassic carmel formation the detachment fold train at reed wash is within the winsor member of the middle jurassic (bajocian through callovian [170.3 through 161.2 ma]) carmel formation. the carmel creates a major but gentle dip slope along the west flank of the san rafael swell and ranges from steep to nearly flat lying on the east flank (figures 3 and 4). it is divided into four members, which in ascending order are: co-op creek limestone (or equivalent judd hollow), crystal creek, paria river, and winsor (figure 6). the carmel ranges from 280 to as much as 1100 feet (85–330 m) thick in the san rafael swell (witkind, 1988; doelling and kuehne, 2008, 2016; doelling and others, 2015). however, all four members are not always present. doelling and kuehne (2008) suggested a possible unconformity (angular) separates the crystal creek and paria river members. the members of the carmel in the san rafael swell were mapped, measured, and described by doelling and kuehne (2008) and sprinkel, doelling, and chidsey (utah geological survey, unpublished measured sections). winsor member the winsor member (bathonian and lower callovian, based on palynomorphs [anderson and lucas, 1994; sprinkel and others, 2011]) ranges in thickness from 190 to 380 feet (58–120 m). it consists of two main informal units: the lower gypsiferous and the upper banded. the gypsiferous unit consists of interbedded red, red-brown, green-gray, or light-gray sandstone, calcarenite, calcisiltite, and siltstone, and white alabaster gypsum and a few limestone beds. sandstone is friable, fine grained, well sorted, and cemented with calcite or iron oxide. calcarenite is very fine grained, well sorted, and laminated to thin bedded with well-developed ripple marks and some bioturbation. calcisiltite appears shaly and weathers into small plates. siltstone is coarse grained, figure 5. stratigraphic column of exposed rocks along the west flank of the san rafael swell, including age, thickness, lithology, weathering profile, and depositional environment. modified from hintze and kowallis (2009). t.c. chidsey, jr., and p.b. anderson detachment fold train in the carmel formation 7 gypsiferous (often with fine laminae of gypsum), and contains small lenses of calcarenite. sandstone and siltstone beds form steep, earthy slopes. the winsor contains six to nine gypsum beds ranging in thickness from 1 to 20 feet (0.3–6 m); total thickness of gypsum ranges from 50 to 90 feet (15–30 m) in outcrops along the west flank of the san rafael swell. gypsum is silty and forms ledges as much as 20 feet (6 m) thick. the banded unit consists of interbedded sandstone, calcarenite, siltstone, and mudstone that displays colored bands of red and gray in various shades, and white gypsum. these rocks have characteristics similar to those in the underlying gypsiferous unit. gypsum veins crisscross the clastic rocks. gypsum beds produce frothy “popcorn-like” or sugary weathering on sparsely vegetated surfaces and drape into drainages (rigby and others, 1974). depositional environment the carmel formation is the result of deposition during the transgression of the shallow marine sundance sea, which extended south from canada into a narrow embayment or arm (called the utah-idaho trough) through northern, central, and southwestern utah (figure 7) (blakey and ranney, 2008; hintze and kowallis, 2009). shoreline fluctuations produced variations between restrictedand more opento marginal-marine conditions, causing significant changes in lithology. this was especially the case along the eastern margin of the marine embayment, which is now exposed on the san rafael swell. the co-op creek and paria river members correspond to marine transgressions and the crystal creek and winsor members represent regressions (doelling and others, 2010). the winsor was deposited in restricted, muddy, hypersaline marine and coastal environments during a second major regression of the sundance sea (blakey and ranney, 2008). geologic description and possible origins of the detachment fold train the detachment fold train in the reed wash area was described by royse (1996), vickye (2004), and chidsey (2013) (figure 8). the fold train consists of over 100 repeated, regularly spaced anticlines (and intervening troughs called synclines [concave-downward folds]) confined to a 62-foot-thick (19 m) gypsum-bearing interval within the winsor member of the carmel formation. the anticlines are unfaulted between the floor and roof detachments surfaces, show a gentle northward plunge to no plunge, and have axes that strike south-southwest to north-northeast. some of the anticlines are moderately asymmetric with an eastward vergence. figure 6. excellent exposure of middle jurassic carmel formation, san rafael group, west flank of the san rafael swell, devils canyon south of i-70, view to the east. the co-op creek, crystal creek, paria river, and part of the winsor members are shown. the carmel is in direct contact with the underlying navajo sandstone represented by the j-1 unconformity. photograph by michael chidsey, sqwak productions inc. 8 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 in the western part of the fold train before continuing into the subsurface, they display chevron-like shapes with detached, rounded crests, whereas in the middle the anticlines have round, box, or isoclinal shapes with vertical or slightly overturned flanks (figures 8a through 8c). at the eastern end of the fold train, the anticlinal limbs have low to moderate dips (royse, 1996). the fold cores consist of contorted bedding, breccia, and heterogeneous lithic components (figure 8d). amplitudes of these folds are about 30 feet (9 m), but decrease on the eastern end of the fold train (vickye, 2004). the distance between anticlinal hinges ranges from 15 to 85 feet (5–26 m). the floor of the entire fold train is a single detachment zone in a 2-foot-thick (0.6 m) claystone, whereas the roof has multiple detachments in a 32-foot-thick (9.8 m) gypsum unit (royse, 1996). there are two possible origins of this unusual detachment fold train. one is thought to be an easternmost manifestation of the sevier thrust faulting in central utah (royse, 1996; vickye, 2004). in that case the compressive forces responsible for its formation are similar to those that produced the east-directed, ramp-style thrusting mapped on the surface near cedar mountain and at farnham anticline (morgan, 2007; chidsey, 2013). these folds may have formed as the result of two compressive structural processes: limb rotation and hinge migration (vickye, 2004). an alternative origin of the detachment fold train, which we prefer, may be that of a gravity slide block off the san rafael swell. such a figure 7. paleogeographic map of utah during deposition of the middle jurassic (170 ma) carmel formation. modified from blakey and ranney (2008). figure 8. fold styles and other characteristics of the detachment fold train: a – round-shaped anticline, b – chevron-shaped anticline, c – box-shaped anticline, and d – typical fold core showing contorted bedding, brecciation, and heterogeneous lithic components (see inset for close-up view). photographs by michael chidsey, sqwak productions inc. a b c d t.c. chidsey, jr., and p.b. anderson detachment fold train in the carmel formation 9 slide would post-date the 70 to 40 ma laramide-age structure and may have formed following the erosion of several thousand feet of sedimentary rocks from regional uplift of the colorado plateau beginning during the miocene (23 ma) (hunt, 1956; lucchitta, 1979; hintze and kowallis, 2009). perhaps after much of the overlying thick section of rocks was removed, the carmel formation was able to naturally slide off the crest of the swell. in any case, calcium sulfate-bearing water within the lower detachment zone may have helped “grease” the incompetent claystone beds, thus assisting with the movement of the fold train (royse, 1996). acknowledgments support for this paper was provided by the utah geological survey (ugs). cheryl gustin and jay hill of the ugs drafted figures. this paper was carefully reviewed by michael d. vanden berg, stephanie m. carney, michael d. hylland, and bill keach of the ugs, along with the editors of this publication. their suggestions and constructive criticism greatly improved the manuscript. references anderson, o.j., and lucas, s.g., 1994, middle jurassic stratigraphy, sedimentation and paleogeography in the southern colorado plateau and southern high plains, in caputo, m.v., peterson, j.a., and franczyk, k.j., editors, mesozoic systems of the rocky mountain region, usa: rocky mountain section sepm (society for sedimentary geology), p. 299–314. blakey, r., and ranney, w., 2008, ancient landscapes of the colorado plateau: grand canyon, arizona, grand canyon association, 156 p. chidsey, t.c., jr., 2013, geology of the san rafael swell, east central utah, in morris, t.h., and ressetar, r., editors, the san rafael swell and henry mountains basin—geologic centerpiece of utah: utah geological association publication 42, p. 1–73. doelling, h.h., blackett, r.e., hamblin, a.h., powell, j.d., and pollock, g.l., 2010, geology of grand staircase-escalante national monument, in sprinkel, d.a., chidsey, t.c., jr., and anderson, p.b., editors, geology of utah’s parks and monuments (3rd edition): utah geological association publication 28, p. 193–235. doelling, h.h., and hylland, m.d., 2002, san rafael swell proposed as site of new national monument: utah geological survey, survey notes, v. 34, no. 2, p. 9–11. doelling, h.h., and kuehne, p.a., 2008, interim geologic map of the temple mountain quadrangle, emery county, utah: utah geological survey open-file report 541, scale 1:24,000, 7 p., 1 plate. doelling, h.h., and kuehne, p.a., 2016, interim geologic map of the eastern half of the salina 30' x 60' quadrangle, emery, sevier, and wayne counties, utah: utah geological survey open-file report 642dm, scale 1:62,500, 2 plates. doelling, h.h., kuehne, p.a., willis, g.c., and ehler, b., 2015, geologic map of the san rafael desert 30' x 60' quadrangle, emery and grand counties, utah: utah geological survey m-267dm, 24 p., scale 1:100,000, 2 plates. gloyn, r.w., tabet, d.t., tripp, b.t., bishop, c.e., morgan, c.d., gwynn, j.w., and blackett, r.e., 2003, energy, mineral, and ground-water resources of carbon and emery counties, utah: utah geological survey bulletin 132, 161 p. hintze, l.f., and kowallis, b.j., 2009, geologic history of utah: provo, utah, brigham young university geology studies special publication 9, 225 p. hunt, c.b., 1956, cenozoic geology of the colorado plateau: u.s. geological survey professional paper 279, 99 p. lucchitta, i., 1979, late cenozoic uplift of the southwestern colorado plateau and adjacent colorado river region: tectonophysics, v. 61, p. 63–95. lupton, c.t., 1913, gypsum along the west flank of the san rafael swell, utah, in contributions to economic geology (short papers and preliminary reports), 1911, part 1 – metals and non-metals except fuels: u.s. geological survey bulletin 530, p. 221–231. morgan, c.d., 2007, structure, reservoir characterization, and carbon dioxide resources of farnham dome field, in willis, g.c., hylland, m.d., clark, d.l., and chidsey, t.c., jr., editors, 2007, central utah—diverse geology of a dynamic landscape: utah geological association publication 36, p. 297–310. rigby, j.k., hintze, l.f., and welsh, s.l., 1974, geologic guide to the northwest colorado plateau, studies for students no. 9: provo, utah, brigham young university geology studies, v. 21, pt. 2, p. 53–93. royse, f., jr., 1996, detachment fold train, reed wash area, west flank, san rafael swell, utah—an example of a limb-lengthening, roll-through folding process on the eastern margin of the sevier thrust belt: the mountain geologist, v. 33, no. 2, p. 45–64. sprinkel, d.a., kowallis, b.j., and jensen, p.h., 2011, correlation and age of the nugget sandstone and glen canyon group, utah, in sprinkel, d.a., yonkee, w.a., and chidsey, t.c., jr., editors, the sevier thrust belt—northern and central utah and adjacent areas: utah geological association publication 40, p. 131–149. vickye, v., 2004, geometric analysis of the reed wash detachment fold train, west flank of san rafael swell, utah: american association of petroleum geologists datapages achives, 14 p. witkind, i.j., 1988, geologic map of the huntington 30' x 60' quadrangle, carbon, emery, grand, and uintah counties, utah: u.s. geological survey miscellaneous investigations series map i-1764, scale 1:100,000, 5 plates. uga-geosite-boden-gilsonite.indd utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors taylor boden utah geological survey, 1594 west north temple, suite 3110 p.o. box 146100, salt lake city, utah 84114 taylorboden@utah.gov independent gilsonite vein, uintah county cover image: historical open-cut mining on the southeast end of the cowboy vein. view to the southeast. 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e geosites refl ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president t. boden independent gilsonite vein, uintah county 3 introduction the uinta basin of northeastern utah (figure 1) contains a wide variety of hydrocarbon resources including vast accumulations of crude oil and natural gas deposits, one of the largest oil shale resources in the world, and the largest tar sand deposit in the united states. in addition, unique solid hydrocarbons, including gilsonite, wurtzilite, tabbyite, and ozokerite, have a long and colorful history of exploration and/or production in the region. the most abundant of these, gilsonite, occurs in distinctive swarms of subparallel, northwest-trending veins. the lateral continuity of the veins is impressive, with relatively long, straight ribbons stretching across the hills of the eastern uinta basin. the veins are also vertically continuous, extending hundreds to more than 3000 feet (900 m) below the ground, commonly having only small variations in width. the uinta basin contains the world’s largest deposit of gilsonite and is the only place in the world where this unique resource is economically produced. gilsonite is remarkable for its unusual geologic origin, chemical and physical properties, and industrial uses. industry pioneers are noted for creating innovative uses for their product and for over 100 years have solved mining, processing, transportation, marketing, and other challenges to supply gilsonite to world markets. accordingly, gilsonite has been studied and described in a large body of research dating back to the 1880s. most recently, the utah geological survey (ugs) published special study 141 (boden and tripp, 2012), which presents the latest mapping of gilsonite deposits and a compilation of existing data. to date, over 70 significant veins and vein systems, having a total combined vein length of over 170 miles (270 km), have been mapped by ugs geologists. background definitions gilsonite is classified as a member of the asphaltite group of hydrocarbon bitumens. this group of naturally occurring solid hydrocarbons are somewhat similar in appearance, occurrence, and properties. gilsonite occurs in dikes (veins), sills, fractures, and disseminated blebs, commonly in association with tertiary-age green river formation oil shale. gilsonite has a dull black appearance on weathered surfaces and a shiny black appearance on fresh surfaces. fractures vary from conchoidal to columnar (pencillated) to flaky or scaly (figure 2). pencillated textures form at right angles to vein walls and penetrate about 6 inches into the ore (verbeek and grout, 1993). historically, industry defined three major subdivisions of gilsonite based on appearance and melting temperature: selects, seconds, and jet. select material is very shiny, melts from 300° to 334°f, and tends to occur in the center of the veins. seconds are somewhat duller, melt from 306° to 361°f, and tend to occur along vein margins, sometimes having pencillated texture. jet gilsonite has a brilliantly shiny surface, a bluish-black color, and melts from 390° to 446°f (abraham, 1960). to date, jet has only been found in the cowboy vein (figure 3). figure 1. the eastern uinta basin in northeast utah is host to numerous gilsonite veins (purple lines). the red box is the area of focus and is shown on figure 3. figure 1. the eastern uinta basin in northeast utah is host to numerous gilsonite veins (purple lines). the red box is the area of focus and is shown on figure 3. figure 2. gilsonite hand samples: a gilsonite with columnar “pencillated” structure, and b “select” gilsonite from underground workings on the independent vein showing conchoidal fracturing. t. boden independent gilsonite vein, uintah county 4 qaf qal tgp tua qat qal tgp qal tgp qac qal tua tua tgp tgp qaf little emma wagon hound cowboy tabor chepetta independent little bonanza caldwell little chepetta augustine »» »»»» »»» » »»»»»» » »» » »»» » »»» » » »»» »» » » » »»» tub tub tub tua tgp tuc qaf qae tub qac tua qac tub #* 0 1 2 30.5 miles explanation geologic units (modified from sprinkel 2007, 2009) figure 3. independent vein geosite area mining activity and geology. see figure 1 for location. ü _̂! qac qac t8s r25e t10s t9s r24er23e 45 45 0 1 2 3 40.5 kilometers stream alluvium (holocene)qal stream terrace deposits (holocene and pleistocene[?])qat member b of uinta formation (eocene)tub mixed alluvium and colluvium (holocene and pleistocene[?])qac alluvial-fan deposits (holocene and pleistocene[?])qaf member a of uinta formation (eocene)tua mixed alluvium and eolian deposits (holocene)qae member c of uinta formation (eocene)tuc parachute creek member of green river formation (eocene)tgp gilsonite vein mahogany zone outcrop gilsonite vein with permitted mining _̂ independent vein geosight overlook #* american gilsonite company mill » shaft ! bonanza, utah figure 3. independent vein area mining activity and geology. t. boden independent gilsonite vein, uintah county 5 fractures filled by gilsonite veins are large-scale hydraulic extension features that resulted from the over pressurization of porespace fluid in the organic-rich source beds in the green river formation (verbeek and grout, 1993). regional stress fields from early-stage, post-laramide, regional tectonic extension is also believed to be a factor in allowing fracture formation. fractures were first forcefully propagated by over pressurized formational water in pore spaces. subsequently, fractures were widened as viscous, liquid asphaltite generated from deeply buried kerogen was injected under high pressure. the gilsonite later solidified, probably through cooling and polymerization. based on laboratory and field evidence it has been proposed that the liquid asphaltite was sourced from kerogen-rich oil shale beds in the parachute creek member of the upper green river formation, possibly beds that are deeply buried on the west side of the basin. the mahogany oil shale zone, which is a sequence of organic-rich (kerogen-rich) dolomitic marlstone in the upper parachute creek member, is a probable major source for the gilsonite (pruitt, 1961; cashion, 1967). this interval reaches up to 40% total organic carbon and was deeply buried, especially to the northwest (down to 10,000+ feet), bringing these deposits into the upper ranges of the oil generation window. history and mining uintahite, the formal scientific name for gilsonite, was first discovered in the 1860s. the material was later informally named gilsonite, after samuel h. gilson, due to his enthusiastic development and promotional efforts (kretchman, 1957; covington, 1964). the name gilsonite was further solidified in common usage when an early mining company adopted and trademarked the name. gilsonite has been mined since the late 1880s in the uinta basin. the first regular shipments began in 1888 from veins in the fort duchesne area. early mining was predominantly by open-cut excavating (figure 5) with picks, shovels, and horse-powered hoists. industrial uses of gilsonite can be grouped into five major categories: 1) asphalt paving mixes and coatings, 2) oil and gas well drilling and completions, 3) inks and paints, 4) chemical products, and 5) metal foundry (boden and tripp, 2012). physical and chemical characteristics of gilsonite are important for determining possible industrial applications. gilsonite from different veins or different parts of veins can be mixed to achieve a product with a specific melting temperature range. origins of gilsonite gilsonite deposits primarily occur as long, mostly vertical dikes (veins) (figure 4) that predominantly trend northwest to southeast and can range in width from less than an inch to more than 20 feet (6 m). horizontal gilsonite sills are occasionally associated with the gilsonite dikes. the source of the gilsonite veins and the mechanism of their emplacement have long been debated and the various theories are well summarized by verbeek and grout (1993). gilsonite float is very light and may wash far down hill from an outcropping vein. weathered gilsonite vein near the surface has a typical “pencillated” fracture structure. this ore generally has a higher melting point and is called “seconds.” unweathered massive gilsonite generally has a lower melting point and is called “select.” sandstone walls are very firm and in clean contact with the gilsonite, with very little penetration. uinta formation green river formation the typical gilsonite vein reaches its maximum width in the vicinity of the uinta-green river formation contact. debris tends to collect at the bottom of a vein where it begins to split upon entering the organic-rich shale beds, making clean mining at this level difficult. most veins gradually thin upward until they thin to extinction in the upper uinta or duchesne river formations. oil shale most veins have their “roots” in the organic-rich oil shale beds of the green river formation. a few veins extend downward into the underlying wasatch formation. figure 4. cross section of a typical gilsonite vein (modified from eldridge, 1901). figure 5. historical open-cut mining on the southeast end of the cowboy vein. view to the southeast. t. boden independent gilsonite vein, uintah county 6 the gilsonite veins actively being mined are near the bonanza area and were discovered by early prospectors who located surface exposures. permitted mining currently occurs in the cowboy, independent, little bonanza, wagon hound, and little emma veins (figure 3) via underground mining methods. mining consists of two major phases: (1) shafts are sunk at regular intervals along the veins, and (2) drifts and stopes are then extended laterally from the shafts. the top 30 feet (9 m) of the gilsonite is left intact for safety and reclamation reasons. gilsonite mining is labor intensive because of its unusual occurrence in narrow (minable widths down to 18 inches), deep, vertical veins, and the explosive hazards associated with gilsonite dust. mining is still done by hand using air-powered chipping hammers to break the gilsonite while avoiding contaminating the ore with broken wall rock, since product purity is important to customers. the broken ore enters a vacuum tube at the bottom of the underground mine and is air lifted to the surface, where it is dropped into a holding container next to the shaft headframe before being trucked to the processing plant in bonanza. over the past decade, gilsonite production from the uinta basin has ranged between 20,000 and 85,000 short tons per year, depending on market conditions, mainly associated with the boom and bust cycles of the oil and gas industry, currently one of the largest markets. the american gilsonite company has been the major producer for many years. cashion (1967) estimated that the original gilsonite resource was approximately 45 million short tons (40.8 million mt); though mining has occurred for many years, millions of tons of resource are estimated to remain. current resources are becoming more difficult and more expensive to mine as they are in deeper, thinner, and more remote veins. gilsonite vein overlook – independent vein the gilsonite veins are in a remote and rugged region and access to some areas is difficult. the independent (bonanza), tabor, and little bonanza vein system is exposed near bonanza, utah, along utah state highway 45, and is one of the easiest and most dramatic places to view the remnants of a gilsonite vein and historical mine workings (figure 3). the independent and tabor veins are a single vein, but are called by different names northwest and southeast, respectively, of the junction with the little bonanza vein. the remnants of the independent (bonanza) vein at the overlook location are exposed in a wide, open-cut trench that stretches across the landscape (figure 6). the independent (bonanza) vein is the second widest gilsonite vein in the uinta basin and has supported mining operations since the late 1800s (pruitt, 1961). in the bonanza area, these veins are exposed in the eocene-age member a and b sandstone of the uinta formation (figure 7). the independent/tabor vein is reported to be more than 7.5 miles (12 km) long, generally strikes from n. 55°–62° w., has a maximum width of 14 feet (4 m), and has an estimated maximum vertical extent of 1100 feet (335 m) (cashion, 1967). this area is accessed by traveling south on utah state highway 45 out of vernal for about 42 miles (68 km). gilsonite veins and mine workings around the bonanza area can be observed along public roads but all mining takes place on private property having no public access. the gilsonite mine workings are dangerous and should not be approached; they should always be viewed from a safe distance. figure 6. overlook of the independent (bonanza), tabor, and little bonanza vein system exposed along utah state highway 45 at bonanza, utah. view to the northwest. gps coordinates of overlook are: easting 655864, northing 4431585. figure 7. stratigraphic column of the eocene section that hosts gilsonite veins in the independent vein area (modified from hintze and kowallis, 2009). t. boden independent gilsonite vein, uintah county 7 geologic setting the uinta basin is an asymmetric, intermontane basin along the northern edge of the colorado plateau. the basin is bordered by the uinta mountains to the north, douglas creek arch to the east, uncompahgre uplift to the southeast, san rafael swell to the southwest, and wasatch range to the west. structural features in the uinta basin and uinta mountains region have a development history that is long and complex; some structures like the uinta rift basin formed during proterozoic time and other structural activity possibly occurred during the pennsylvanian-permian ancestral rocky mountains uplift (stone, 1993). the uinta basin, uinta mountains, and associated folds were formed by west-southwest to east-northeast compression during the cretaceous-early tertiary laramide orogeny (erslev, 1993; stone, 1993). on the eastern side of the uinta basin in the bonanza area, gilsonite veins are hosted in the gently dipping eocene-age strata of the green river and uinta formations (figures 3 and 7). these formations were deposited in lacustrine (green river formation) and fluvial (uinta formation) environments and range in composition from carbonate rocks to clastic rocks. contacts are gradational with complex intertonguing relationships and abrupt facies changes that reflect fluctuating paleo-lake levels. the extensive vertical and horizontal continuity of the gilsonite veins is related to the stratigraphy and lithology of the host formations (pruitt, 1961; cashion, 1967). eocene green river formation lacustrine deposits host gilsonite veins in both the basal and upper members. all significant known gilsonite veins are located above the mahogany oil shale zone in the upper green river formation, except two veins that are in the basal member in the southeastern part of the basin (boden and tripp, 2012). gilsonite veins are known to achieve their greatest thickness in the lower uinta formation. the eocene uinta formation overlies the parachute creek member of the green river formation and consists of marginal lacustrine deposits in the lower members, mixed fluvial and marginal lacustrine deposits in the middle part, and entirely fluvial deposits in the upper part. marginal lacustrine deposits in the lower part consist primarily of thick, laterally continuous, medium-bedded to massive sandstone containing interbedded siltstone and thin intervals of marlstone and tuff. fluvial beds in the middle and upper members are composed of channel-form sandstone, variegated mudstone, and minor conglomerate that tend to be laterally discontinuous and thinner than marginal lacustrine deposits down section (pruitt, 1961; cashion, 1967). gilsonite veins commonly split, become discontinuous, and/ or pinch out in the mudstone-rich upper uinta formation. summary gilsonite is a unique material that has been mined for well over a hundred years and its geological occurrence and mining history have been studied by numerous researchers in a large body of work. even though significant amounts of the approximately 45-million-short-ton original gilsonite resource have been mined, millions of tons of the valuable resource remain. additional resources are likely to be found in the deeper parts of the bonanza area veins and in thinner, more remote veins that will likely be more expensive to mine. gilsonite will continue to be mined in the uinta basin for decades, ensuring a steady supply to world markets of this unique and valuable utah resource. references abraham, h., 1960, asphalts and allied substances—their occurrence, modes of production, uses in the arts and methods of testing (6th edition), volume 1, historical review and natural raw materials: princeton, new york, d. van nostrand company, inc., p. 1–302. boden, t., and tripp, b.t., 2012, gilsonite veins of the uinta basin, utah: utah geological survey special study 141, 50 p., 1 plate, cd. cashion, w.b., 1967, geology and fuel resources of the green river formation, southeastern uinta basin, utah and colorado: u.s. geological survey professional paper 548, 48 p. covington, r.e., 1964, a brief history of early mineral exploitation in the uinta basin, in sabatka, e.f., editor, geology and mineral resources of the uinta basin, utah’s hydrocarbon storehouse: intermountain association of petroleum geologists guidebook, 13th annual field conference, p. 1–16. eldridge, g.h., 1901, the asphalt and bituminous rock deposits of the united states, in walcott, c.d., director: u.s. geological survey twenty-second annual report of the united states geological survey to the secretary of the interior, pt. 1, p. 209–364. erslev, e.a., 1993, laramide basement tectonics, in schmidt, c.j., chase, r.b., and erslev, e.a., editors, laramide basement deformation in the rocky mountains foreland of the western united states: geological society of america special paper 280, p. 339–358. hintze, l.f., and kowallis, b.j., 2009, geologic history of utah: provo, brigham young university geology studies, special publication 9, 225 p. kretchman, h.f., 1957, the story of gilsonite: salt lake city, utah, american gilsonite company, 96 p. pruitt, r.g., jr., 1961, the mineral resources of uintah county: utah geological and mineralogical survey bulletin 71, 101 p. 8 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 sprinkel, d.a., 2007, interim geologic map of the vernal 30' x 60' quadrangle, uintah and duchesne counties, utah, and moffat and rio blanco counties, colorado: utah geological survey open-file report 506dm, 3 plates, scale 1:100,000, cd. sprinkel, d.a., 2009, interim geologic map of the seep ridge 30' x 60' quadrangle, uintah, duchesne, and carbon counties, utah, and garfield and rio blanco counties, colorado: utah geological survey open-file report 549dm, 3 plates, scale 1:100,000, cd. stone, d.s., 1993, tectonic evolution of the uinta mountainspalinspastic restoration of a structural cross section along longitude 109°15', utah: utah geological survey miscellaneous publication 93-8, 19 p. verbeek, e.r., and grout, m.a., 1993, geometry and structural evolution of gilsonite dikes in the eastern uinta basin, utah: u.s. geological survey bulletin 1787-hh, 42 p., 1 plate, scale 1:250,000. uga-geosite-loope-slot-canyons.indd utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors david b. loope earth & atmospheric sciences, university of nebraska lincoln, ne 68588-0340 dloope1@unl.edu cover image: exploring a southern utah slot canyon. photo by jim elder. cut, fill, repeat: slot canyons of dry fork, kane county 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e geosites refl ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: loope, d.b., 2019, cut, fi ll, repeat—slot canyons of dry fork, kane county, in milligan, m., biek, r.f., inkenbrandt, p., and nielsen, p., editors, utah geosites: utah geological association publication 48, 8 p., https://doi.org/10.31711/geosites. v1i1.61. presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president d.b. loope slot canyons of dry fork 3 introduction the slot canyons of southern utah (figure 1) have become popular destinations for hikers, climbers, and photographers. for most of these canyons, the geology is simple: sediment carried by flowing water abrades a thick, homogeneous sandstone. as time passes, the rate of downcutting is rapid compared to the rate of cliff retreat. end of story. the strange abundance and configuration of the slot canyons along dry fork coyote (a tributary of coyote gulch and the escalante river), however, have a convoluted geologic history that is climate-driven and involves canyon cutting, canyon filling, and more canyon cutting. driving and hiking directions the hole-in-the-rock road, starting 4.3 miles (7 km) east of the town of escalante, is a graded dirt road that provides good access to the slot canyons. when dry and graded, this road is passable by passenger cars. contact the escalante interagency visitor center (435-826-5499) for a road report and a weather forecast. from the junction of highway 12 and hole-in-the-rock road, it is 26 miles (42 km) to the turnoff to dry fork (figure 2; n37°27'59.7"; w111°13'13"). after turning, go 0.7 mile (1.1 km) on this rutted road (figure 2), and then take the road that branches left and go 0.8 mile (1.3 km) to the trailhead and parking area (figure 2; n37°27'59"; w111°13'26"). the trailhead provides a good view of the canyon of dry fork, but the slot canyons are nearly invisible from there. a cairned trail (take your time, it is somewhat difficult to follow) descends about 300 feet (90 meters) to reach the canyon floor. from there, you can choose which of seven slot canyons (if any) you want to explore. the first one—just to the left as you reach the floor of main canyon—provides a shady place to admire the walls of a slot without abrading any skin; the second one (peekaboo; figure 4b) is more of a challenge. it is easy to lose your way in this tangle of canyons. take a copy of figure 2 with you. the main (trunk) canyon of dry fork carries large volumes of gravel, sand, and mud because it drains from the straight cliffs. the tributaries coming in from the north carry only sand. if you get turned around, look for mud and gravel and follow it upstream; it will take you back to the place where the trail enters the canyon. geology at the canyon rim, you are standing on the basal, red siltstone of the jurassic carmel formation. to the north, the tan, large-scale crossbeds of the jurassic navajo sandstone lie just below, and the henry mountains are visible in the distance. to the south are the straight cliffs, composed of cretaceous and jurassic strata (figure 3). landslides, many of them deposited during the ice age, cover much of the bedrock (figure 2). as you descend the trail into the canyon, you will see some large, dark masses of rock. these are weathered concretions (sandstone preferentially cemented by calcite--calcium carbonate; caco3). these concretions weather dark because the crystals of calcite cementing the sand grains contain small amounts of iron. atoms of iron are the same size and have the same (+2) charge as calcium atoms, so they fit into the calcite crystal lattice. geologic history as you descend further into dry fork canyon, look for a steep wall of loose, bedded sand and gravel (figure 4a). this is the first evidence that, after canyons were cut into the navajo sandstone, some were then filled to the rim with sediment. streams (even those that are dry nearly all the time) can cut bedrock canyons; they can also fill them up with sediment. climate is a major control on these processes (figure 5). based on optically stimulated luminescence ages from buried sand grains (loope and others., 2014), the sediment that filled some of canyons here (the broad, older ones) was deposited during the late pleistocene between about 55,000 and 48,000 years ago. the slots are a younger generation of canyons—they originated during a second canyon-cutting episode (figure 4). during this second episode of canyon cutting, as the streams started to remove the fill figure 1. exploring a southern utah slot canyon. photo by jim elder. 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 pre-55,000 yrs trunk slot #1 trunk slot #2 trunk slot #3 spooky slot brimstone trunk slot #4 0 1.0 0.5 parking present peekaboo slot a b c 370 29’ 20” n 11 10 1 1’ 07 ” w n n n coyote gulch 7a 4b 4a kilometers recent landslides older landslides debris �ow 0 10 5 straight cli�s nv id wy az co ut * geosite geosite h o le-in-th e-rock ro ad big hollow wash dry fork “orphan” slot 7b 6c abandoned water course sandy wash slot canyon subject to �ooding no �ooding �ow direction of �oods �gure in text abandoned (“orphan”) slot 0 0.3 0.6 miles hole in the rock road (to highway 12) 3 trail dry fork 0 6 3 miles kilometers straight cli�s dry fork brimstone spooky peekaboo 6b note: construction of new parking lot and additional trails will start november, 2019. figure 2. three maps of the geosite and its vicinity. all water courses are dry washes; blue arrows show flow direction. a. broad setting from the surficial geologic map of williams (1985); b and c are nested within a. b. configuration of ancient drainages along a short portion of dry fork prior to filling of the canyons with sediment. c. configuration of modern drainages, showing four slot canyons along the trunk drainage (dry fork), the two slot canyons (peekaboo and spooky) that are paired with two abandoned water-courses, and one abandoned (orphan) slot canyon paired with brimstone wash. modified from loope and others (2014). d.b. loope slot canyons of dry fork 5 of sand and gravel, they became entrapped by the rising walls of sediment. when they cut all the way through the sand and gravel, they may have reached the lowest bedrock where the old canyons were. but if they cut in places slightly different from the paths of the older canyons, they cut new canyons (figure 6). the seven slot canyons along dry fork lie within a very small area (1.5 square miles; 4 square kilometers) (figure 2). four of the slots are along the main trunk of dry fork, the other three are tributaries from the north—peekaboo, spooky, and brimstone. each of these seven slots is paired with a segment of a wider, abandoned bedrock canyon (figure 2). among the three tributaries of dry fork, brimstone canyon (marked as an “orphan” on figure 2) has a unique geologic history. unlike peekaboo and spooky, modern drainage at brimstone is through a big, old, open canyon (figure 2). the mouth of the brimstone slot looks similar to the mouths of peekaboo and spooky (figure 7a), but the head of the brimstone slot is perched high on the east wall of brimstone canyon (figure 7b). ever since the initial cutting of peekaboo and spooky, flash floods have been forced to roar through them. not so with brimstone. apparently, during a major event in the distant past, flood waters were able to escape brimstone slot and flush out the sand that once filled the last half mile of brimstone canyon (figure 2). on the basis of 14 dated samples (loope and others, 2014), we concluded two things about the timing of the events that generated this landscape: 1) sediment filled the older, wider canyons ~55,000 to ~48,000 years ago; and 2) all the slot canyons are younger than 48,000 years. the nature of the climate and other environmental conditions that caused burial and eventual cutting of the slots are not well known. we do know that, during the pleistocene, a vast glacial ice sheet formed in the hudson bay area of canada and repeatedly advanced deep into central united states. mountain glaciers formed in the rockies, and, in utah, swelling of rivers and the cooler climate caused a water body to form in the great basin (lake bonneville) that reached a depth of 1000 feet (300 m) and an area of up to 20,000 mi2 (52,000 km2; https://en.wikipedia.org/wiki/lake_bonneville). studies of lake bonneville deposits indicate that, during the last 30,000 years, lake bonneville’s fluctuations record glacial-interglacial cycles and accompanying changes in northern hemisphere storm tracks (oviatt, 1997). regional climate shifts, the oldest of which pre-dated the rise of lake bonneville, led to changes in vegetation, sediment supply, and stream power (figure 5; blum and tornquist, 2000) and thereby likely caused the canyon filling and canyon cutting that took place along dry fork. figure 3. stratigraphic section and geologic map for a portion of the straight cliffs and escalante canyons. hole in the rock road is shown in red. from hintze and kowallis (2009) and interactive geologic map of utah; 1:500,000 scale (utah geological survey). kane county gar�eld county jn jn jn qe qao jk jc morrison fm jm ktjk qlsqls qls qls kt jc jc ksc ksc see fig. 2c carmel fm navajo sandstone kayenta fm tropic shale straight cli� fm landslides mixed stream and dune deposits dune sand 5 miles 8 kilometers qe qe qao jm wahweap formation straight cli�s formation tropic shale dakota fm cedar mountain fm morrison fm entrada sandstone carmel fm navajo sandstone kayenta fm wingate ss cr et ac eo us ju ra ss ic st ra ig ht c li� s es ca la nt e ca ny on s kane county gar�eld county jn jn jn qe qao jk jc morrison fm jm ktjk qlsqls qls qls kt jc jc ksc ksc see fig. 2c carmel fm navajo sandstone kayenta fm tropic shale straight cli� fm landslides mixed stream and dune deposits dune sand 5 miles 8 kilometers qe qe qao jm wahweap formation straight cli�s formation tropic shale dakota fm cedar mountain fm morrison fm entrada sandstone carmel fm navajo sandstone kayenta fm wingate ss cr et ac eo us ju ra ss ic st ra ig ht c li� s es ca la nt e ca ny on s kane county gar�eld county jn jn jn qe qao jk jc morrison fm jm ktjk qlsqls qls qls kt jc jc ksc ksc see fig. 2c carmel fm navajo sandstone kayenta fm tropic shale straight cli� fm landslides mixed stream and dune deposits dune sand 5 miles 8 kilometers qe qe qao jm wahweap formation straight cli�s formation tropic shale dakota fm cedar mountain fm morrison fm entrada sandstone carmel fm navajo sandstone kayenta fm wingate ss cr et ac eo us ju ra ss ic st ra ig ht c li� s es ca la nt e ca ny on s kane county gar�eld county jn jn jn qe qao jk jc morrison fm jm ktjk qlsqls qls qls kt jc jc ksc ksc see fig. 2c carmel fm navajo sandstone kayenta fm tropic shale straight cli� fm landslides mixed stream and dune deposits dune sand 5 miles 8 kilometers qe qe qao jm wahweap formation straight cli�s formation tropic shale dakota fm cedar mountain fm morrison fm entrada sandstone carmel fm navajo sandstone kayenta fm wingate ss cr et ac eo us ju ra ss ic st ra ig ht c li� s es ca la nt e ca ny on s kane county gar�eld county jn jn jn qe qao jk jc morrison fm jm ktjk qlsqls qls qls kt jc jc ksc ksc see fig. 2c carmel fm navajo sandstone kayenta fm tropic shale straight cli� fm landslides mixed stream and dune deposits dune sand 5 miles 8 kilometers qe qe qao jm wahweap formation straight cli�s formation tropic shale dakota fm cedar mountain fm morrison fm entrada sandstone carmel fm navajo sandstone kayenta fm wingate ss cr et ac eo us ju ra ss ic st ra ig ht c li� s es ca la nt e ca ny on s kane county gar�eld county jn jn jn qe qao jk jc morrison fm jm ktjk qlsqls qls qls kt jc jc ksc ksc see fig. 2c carmel fm navajo sandstone kayenta fm tropic shale straight cli� fm landslides mixed stream and dune deposits dune sand 5 miles 8 kilometers qe qe qao jm wahweap formation straight cli�s formation tropic shale dakota fm cedar mountain fm morrison fm entrada sandstone carmel fm navajo sandstone kayenta fm wingate ss cr et ac eo us ju ra ss ic st ra ig ht c li� s es ca la nt e ca ny on s kane county gar�eld county jn jn jn qe qao jk jc morrison fm jm ktjk qlsqls qls qls kt jc jc ksc ksc see fig. 2c carmel fm navajo sandstone kayenta fm tropic shale straight cli� fm landslides mixed stream and dune deposits dune sand 5 miles 8 kilometers qe qe qao jm wahweap formation straight cli�s formation tropic shale dakota fm cedar mountain fm morrison fm entrada sandstone carmel fm navajo sandstone kayenta fm wingate ss cr et ac eo us ju ra ss ic st ra ig ht c li� s es ca la nt e ca ny on s 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 sand and gravel sand and gravel former position of peekaboo peekaboo slot a b figure 4. alluvial fills and slot canyons. white arrows show people for scale. see figure 2 for locations. a. northward view of thick sediment fill of the trunk of dry fork #1 as seen from entry trail. people are about to enter trunk slot #1. b. northward view of the sediment-filled former course of peekaboo and of peekaboo slot at its junction with dry fork. from loope and others (2014). d.b. loope slot canyons of dry fork 7 figure 6. cutting and filling of canyons. see figure 2 for locations of photos. a. stages of evolution involving a bedrock spur (dark gray), burial, and cutting of a young, narrow canyon (from ouimet and others, 2008). b. upstream entrance to trunk slot #3; c. upstream entrance to trunk slot #4. in both b and c, the sediment-filled, broader canyon lies to the right (south) of the slot canyon. from loope and others (2014). 1 2 3 4 a b c stream powersediment supply aggradationerosion figure 5. balance model for erosion (incision) and aggradation (filling) of stream channels, emphasizing changes in the relationship between stream power and sediment supply. erosion of channels take place when the transport capacity of the water discharge (stream power) exceeds the sediment supply. with aggradation, the reverse is true. climate and vegetation are major factors controlling both sediment supply and stream power. modified from blum and tornquist (2000). 8 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 19 m a b figure 7. a. mouth of brimstone slot at its confluence with dry fork. b. head of the slot canyon is at tip of upper white arrow, 60 feet (19 m) above the floor of brimstone wash. black arrow points to person for scale. stream incised some of its fill and then cut the slot before abandoning it when it overflowed its (west) bank and reoccupied the broad canyon. it then rapidly excavated the remaining 60 feet (19 m) of fill in that canyon, leaving an “orphan slot” high, dry, and open. from loope and others (2014). acknowledgments jon mason, jodi norris, and walter loope provided helpful reviews and jim elder kindly provided the photograph for figure 1. references blum, m.d., and tornquist, t.e., 2000, fluvial responses to climate and sea-level change—a review and look forward: sedimentology, v. 47 (supplement 1), p. 2-48. hintze, l.f., and kowallis, b.j., 2009, geologic history of utah: brigham young university geology studies, special publication 9, 225 p. loope, d.b., goble, r.j., and johnson, j.p.l., 2014, prelude to seven slots—filling and subsequent modification of seven broad canyons in the navajo sandstone, south-central utah, in maclean, j.s., biek, r.f., and huntoon, j.e., editors, geology of utah’s far south: utah geological association publication 43, p. 11–24. ouimet, w.b., whipple, k.x., crosby, b.t., johnson, j.p., and schildgen, t.f., 2008, epigenetic gorges in fluvial landscapes: earth surface processes and landforms, v. 33, p. 1993-2009. oviatt, c.g., 1997, lake bonneville fluctuations and global climate change: geology, v. 25 (2), p. 155-158. williams, v.s., 1985, surficial geologic map of the kaiparowits coal-basin area, utah: united states geological survey miscellaneous investigations series, map i-1033-l, scale 1:125,000. uga-geosite-loope-moki.indd utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors david b. loope earth & atmospheric sciences, university of nebraska lincoln, ne 68588-0340 dloope1@unl.edu cover image: permian rhizoliths in the cedar mesa sandstone at moki dugway. plant root systems preserved in the permian cedar mesa sandstone at moki dugway, southeastern utah 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. � e geosites re� ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. � e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. � e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. � is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: loope, d.b., 2019, plant root systems preserved in the permian cedar mesa sandstone at moki dugway, southeastern utah, in milligan, m., biek, r.f., inkenbrandt, p., and nielsen, p., editors, utah geosites: utah geological association publication 48, 6 p., https://doi.org/10.31711/geosites.v1i1.60. presidents message i have had the pleasure of working with many di� erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several di� erent geological provinces. we have the basin and range to the west and the central utah hingeline and � rust belt down the middle. � e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava � ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. � e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. � e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. � e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. � is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, � ank you to the american association of petroleum geologists, rocky mountain section foundation for their nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president d.b. loope plant root systems at moki dugway 3 introduction rooted green plants represent the base of the food chain for most terrestrial ecosystems, but, compared to animal burrows, root systems are relatively rarely recognized in ancient sedimentary rocks. plant roots that penetrate unconsolidated sand dunes, especially those containing not only quartz grains, but also abundant grains of calcite (caco3), are commonly replaced by ne crystals of calcite (klappa, 1980). � ese structures (known by geologists as rhizoliths from the greek for “root rock”) are one form of calcite cemented soil and sediment called caliche ( gure 1). caliche crystallizes well above the water table and its calcite crystals are tiny because of rapid evaporation of soil water. one source of the calcium (ca) and carbonate (co3) ions necessary for making the calcite of caliche is falling dust, and another source is the dissolution of calcite grains already in the soil. caliche is widespread in semi-arid regions. in regions with abundant rainfall, available calcium and carbonate ions are rapidly � ushed downward, out of the soil, preventing calcite crystals from growing in the root zone. in arid regions there is too little available soil water for crystal growth. because plant roots in modern semi-arid settings are commonly preserved by caliche ( gure 1), rhizoliths in ancient rocks are good indicators of semi-arid paleoclimates. � e early permian (245-286 million year old) root systems preserved the cedar mesa sandstone at moki dugway ( gure 2) grew on low-relief land surfaces that formed when dune elds were � attened by wind erosion. a near-surface water table may have prevented further erosion of the permian dune sand and allowed the land surface to be colonized by woody plants. driving and walking directions perhaps the best place in the u.s. to see ancient rhizoliths is in a utah 261 roadcut along the upper portion of moki dugway and on the outcrop just above the highest switchback ( gure 3). if you are approaching this geosite from the north, stop rst at the viewpoint on the south edge of cedar mesa, at the top of moki dugway ( gure 3). � is viewpoint provides a dramatic view of the cedar mesa sandstone, and in the distance, the goosenecks of the san juan river and monument valley ( gure 4). � e moki dugway is a steep (11% grade), well-maintained, dirt road with figure 1. rhizolith collected from loose, bioturbated dune sand north of moab, ut. arrow points to the boundary between lightly cemented quartz sand and sand-free caliche composed tiny calcite crystals. void spaces were made by smaller roots that penetrated the so� , chalky caliche that � lls the cylindrical space produced by growth of a large root. figure 2. permian rhizoliths in the cedar mesa sandstone at moki dugway. � ese rhizoliths have a darker color than those found in recently deposited, uncemented sand (� gure 1) because all the pore space originally in them has been � lled by calcite cement. � e lighter patches of sandstone are reduction halos—iron oxide was removed from surrounding quartz grains by organic acid from the root. to utah 95 and natural bridges to utah 163 and mexican hat to m ul ey p oi nt park on turn 37°16'30.72" n; 109°56'11.61"w 1.0 km n to v al le y of th e g od s uta h 2 61 utah moki dugway geosight viewpoint figure 3. topographic map showing the roads that provide access to this geosite and the recommended stopping places. moki dugway geosite 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 a series of switchbacks. unless it is wet, neither high-clearance nor 4wd are needed. � e state of utah recommends it only for vehicles less than 28 feet long and less than 10,000 pounds. for viewing the rhizoliths up close, the best place to park is at the uppermost sharp switchback on the highly sinuous road ( gure 3). it is possible to climb up the outcrop at the outer bend of this switchback; you can also see rhizoliths exposed along both sides of the roadcut as you walk back up the road a few hundred meters. if you approach this geosite from the south, use gures 3 and 5 to help you recognize the correct switchback for parking. gps location: n37.2752; w109.9365 (wgs84) the cedar mesa sandstone � e permian cedar mesa sandstone is one of more than a dozen thick sedimentary rock formations on the colorado plateau that were deposited by large, wind-blown sand dunes ( gure 5). all of these sandstones have large-scale crossbedding that formed as the dunes migrated downwind and accumulated vertically in subsiding sedimentary basins. because mesozoic rock formations have been eroded from the broad crest of the monument upli� ( gure 6) in southeastern utah (named for monument valley) older rocks including the cedar mesa sandstone crop out widely there, forming cedar mesa (type locality of the formation), as well as � e needles and � e maze (canyonlands national park), and the natural bridges at natural bridges national monument. halgaito shale cedar mesa sandstone muley point monument valley figure 4. view southward from the parking area at the top of moki dugway. park here * siltstone paleosol roadcut climb here figure 5. cedar mesa sandstone outcrops along upper part of moki dugway where the rhizoliths in one of the permian paleosols are best exposed. navajo ss lukachukai mbr cedar mesa ss coconino ss moenkopi fm chinle fm kaibab ls organ rock fm halgaito fm wingate ss moenave fm kayenta fm wingate ss, rock point mbr white rim ss dechelly ss guadalupean leonardian wolfcampian lower mid. upper triassic jurassic hettangian sinemurian pliensbachian toarcian permian 296 my 187 my figure 6. eolian sandstones of the colorado plateau (orange). vertical scale is geologic time. � e black bar shows the strata that crop out at moki dugway. modi� ed from blakey and others (1988). most of the sand grains that accumulated to form the cedar mesa sandstone are abraded crystals of quartz and feldspar (ultimately derived from granite or gneiss). but calcite (calcium carbonate) grains are also present, and some of these can (using a microscope) be recognized as fragments of the skeletons of marine invertebrates—brachiopods, bryozoans, and crinoids. � ese animals d.b. loope plant root systems at moki dugway 5 sand to the dune eld diminished), the dunes were eroded and silt (probably delivered as dust fall and from small streams) started to accumulate on � at surfaces. abundant rhizoliths provide evidence that these � at surfaces were colonized by terrestrial plants. because vegetation slows the wind, the plants likely caused the falling silt to accumulate instead of continuing its downwind migration. � e rock layers with rhizoliths are thus paleosols (ancient soils) preserved in the stratigraphic record. � e diameters of the rhizoliths ( gure 2) indicate that many of the plants were medium to large trees. � e roots of trees and the burrows of small invertebrates (probably insects) penetrated the uncemented sand, obliterating the bedding in the upper portions of each tabular accumulation. beyond the reach of the roots and burrows, undisturbed crossbedding was preserved in the lower portions of each tabular sand accumulation. crustaceans such as shrimp and cray sh build complex, branching burrow systems, but, unlike root systems, the burrows are lled by loose sand (never by sand-free, ne-grained calcite). � e rod-like form and coloration of the cedar mesa structures are produced by: 1) elongation and radial expansion of a cylindrical plant root that pushed sand grains aside; 2) a living root and its rootlets that lived in a shallow seaway that lay to the north and west, and were incorporated into a broad coastal dune eld by onshore winds. a� er the dunes were eroded � at (see below), the presence of these calcite grains in the dune sand became important to the excellent preservation of root systems in the cedar mesa sandstone. stacking of sediment slabs and preservation of root systems although some of the eolian sandstones of the colorado plateau form massive, uninterrupted cli� s (for example, the jurassic navajo and wingate sandstones, gures 6 and 7), the cli� s exposing the cedar mesa sandstone are discontinuous, giving the cedar mesa a tabular, “layer-cake” appearance caused by alternation of thin, red siltstones with up to 40 distinct 6 to 65 feet thick (2 to 20 meters) slabs of sandstone ( gures 4, 5, 8; loope, 1985; mountney, 2006). during the permian period, massive glaciers were dynamically forming and melting on the supercontinent gondwana. clearly, global climate was � uctuating widely. when sand was abundant and winds were relatively gentle, sand built up as dunes climbed over one another. onshore winds led to the incorporation of calcite grains (fragments of marine fossils) into the desert dunes (loope, 1984). but as the wind strengthened (or the supply of # # tr1 j1 qao qa j1 j2 tr2 k1 qe j2 qe qe qe qa qe k1 k1 jg jg jg qe qe k1 k1 j1 qao blanding montezuma creek pcm pp pi i jtr qa qe cedar mesa monument valley co m b ri dg e goosenecks n 15 mi 24 km trc trc trm pcm ut-95 ut-163 ut-261 pennsylvanian hermosa fm qa quaternary alluvium jm jurassic morrison j1 jurassic carmel-entradasummerville fm trc triassic chinle gp trm triassic moenkopi fm pcm permian cedar mesa ssp p permian-pennsylvanian halgaito fm jg jurassic glen canyon gp. i pi qe quaternary eolian sand organ rock shale ut-az border moki dugway navajo sandstone kayenta fm. wingate sandstone chinle fm. moenkopi fm. dechelly ss cedar mesa sandstone halgaito formation hermosa fm paradox fm. salt and gypsum limestone siltstone and sandstone eolian sandstone eolian sandstone eolian sandstone eolian sandstone mudstone and sandstone siltstone siltstone and sandstone ju ra ss ic tr ia ss ic pe rm ia n pe nn sy lva ni an organ rock fm. sandstone pa le oz oi c m es oz oi c figure 7. geologic map of the area surrounding moki dugway (le� ). � e older rocks (blue and purple) are exposed because of erosion of younger rocks (tan and green) from the monument upli� . � e column on the right shows strata exposed in the area. � e black bar shows the strata that crop out at moki dugway. map based on interactive geologic map of utah (utah geological survey); column based on hintze and kowallis (2009). # # tr1 j1 qao qa j1 j2 tr2 k1 qe j2 qe qe qe qa qe k1 k1 jg jg jg qe qe k1 k1 j1 qao blanding montezuma creek pcm pp pi i jtr qa qe cedar mesa monument valley co m b ri dg e goosenecks n 15 mi 24 km trc trc trm pcm ut-95 ut-163 ut-261 pennsylvanian hermosa fm qa quaternary alluvium jm jurassic morrison j1 jurassic carmel-entradasummerville fm trc triassic chinle gp trm triassic moenkopi fm pcm permian cedar mesa ssp p permian-pennsylvanian halgaito fm jg jurassic glen canyon gp. i pi qe quaternary eolian sand organ rock shale ut-az border moki dugway navajo sandstone kayenta fm. wingate sandstone chinle fm. moenkopi fm. dechelly ss cedar mesa sandstone halgaito formation hermosa fm paradox fm. salt and gypsum limestone siltstone and sandstone eolian sandstone eolian sandstone eolian sandstone eolian sandstone mudstone and sandstone siltstone siltstone and sandstone ju ra ss ic tr ia ss ic pe rm ia n pe nn sy lva ni an organ rock fm. sandstone pa le oz oi c m es oz oi c 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 produced acid that dissolved iron oxide from the surrounding sediment, producing a “reduction halo” ( gures 2 and 8). much of the dead organic tissue is eventually replaced by ne-grained calcite ( gures 1 and 2; loope, 1988). a sandstone like the cedar mesa, with abundant sand-size, marine fossils could have been interpreted as forming below sea level. and the obliteration of bedding (called “bioturbation” by many geologists) could have been interpreted as evidence for burrowing by marine invertebrates such as crustaceans. but the distinctive, large-scale crossbedding shows that onshore winds blew the marine fossils (along with abundant quartz and feldsper) onto an emergent land surface covered by sand dunes, and the sandfree calcite rods representing ancient caliche show that the desert dunes were periodically � attened and colonized by land plants growing in a semi-arid climate. acknowledgments nigel mountney, jon mason, and jodi norris provided very helpful reviews of this paper. references blakey, r.c., peterson, f., and kocurek, g., 1988, synthesis of late paleozoic and mesozoic eolian deposits of the western interior of the united states: sedimentary geology, v. 56, p. 3-125. hintze, l.f., and kowallis, b.j., 2009, geologic history of utah: brigham young university geology studies, special publication 9, 225 p. klappa, c.f., 1980, rhizoliths in terrestrial carbonates: classi cation, recognition, genesis and signi cance: sedimentology, v.27, p. 613-629. loope, d.b, 1984, eolian origin of upper paleozoic sandstones, southeastern utah: journal of sedimentary research, v. 54, p. 563-580. loope, d.b., 1985, episodic deposition and preservation of eolian sands: a late paleozoic example from southeastern utah: geology, v. 13, p. 73-76. loope, d.b., 1988, rhizoliths in ancient eolianites: sedimentary geology, v. 56, p. 301-314. mountney, n.p., 2006, periodic accumulation and destruction of aeolian erg sequences in the permian cedar mesa sandstone, white canyon, southern utah: sedimentology, v. 53, p. 789-823. calcitic rhizoliths with reduction halos in permian paleosols siltstone 5 m crossbedded eolian sandstone 16.4 ft figure 8. relationships of thin siltstones and rhizolith-bearing paleosols to tabular sheets of crossbedded sandstone in the permian cedar mesa sandstone. uga-geosite-clark.indd utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors donald l. clark utah geological survey, po box 146100, salt lake city, ut 84114-6100 donclark@utah.gov cover image: view to northwest, from quincy spring area of cedar mountains, to red bed geosite. northern cedar mountains red beds, tooele county 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e geosites refl ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: clark, d.l., 2019, northern cedar mountains red beds, tooele county, in milligan, m., biek, r.f., inkenbrandt, p., and nielsen, p., editors, utah geosites: utah geological association publication 48, 5 p., https://doi.org/10.31711/ geosites.v1i1.57. presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president d.l. clark northern cedar mountains red beds 3 introduction recent mapping for the tooele 30' x 60' quadrangle geologic map revealed more information about interesting exposures of red beds cropping out in an 11-mile (18-km) swath along the northwestern flank of the cedar mountains (clark and others, 2017, 2019 in review). they are unusual because such rocks are seldom preserved in northwestern utah, an area known for thick thrust sheets of paleozoic marine carbonate and sandstone. these rocks were first mapped by robert maurer for his ph.d. dissertation on the geology of the cedar mountains (maurer, 1970). he called them north horn (?) formation and noted the presence of fresh-water snails (gyraulus sp.) that a paleontologist said were probably of late paleocene or eocene age. the north horn formation is considered late cretaceous (maastrichtian age) to paleocene or eocene (~70 to 60 ma) in age. however, recent detrital zircon analysis from sandstone in the unit suggest a much older age, with the youngest grains (maximum depositional age) of 117 ma, or early cretaceous, aptian (ugs and o’sullivan, 2017). two samples of gray mudstone taken for fossil pollen in the unit did not produce any usable material. inspection of these rocks by don deblieux (utah geological survey paleontologist) revealed no bone or other biological material at this location. n from/to slc 0 2.5 5 miles rest area 80 56 exit 84 exit park here hike red bed exposures cedar mtns. figure 1. geosite location map. slc is salt lake city. unit description these red beds are brightly colored and stand out from the grays and browns of surrounding rocks and vegetation. the rocks (map unit ks) consist of predominantly moderate-reddish-orange mudstone and siltstone with lenses of red and gray conglomerate and sandstone. the conglomerate clasts include sandstone, limestone, and chert that are up to 18 inches (46 cm) in diameter and were likely derived from permian-age formations; northern exposures appear to be somewhat more conglomeratic. southern exposures locally include an upper part of pink to pale purple mudstones with some interbedded coarser clastics with rapid lateral lithologic changes. much of the outcrop belt is poorly exposed with cover by eolian sand, colluvium, and landslide debris, but exposures are as much as 1100 feet (335 m) thick. geologic significance assuming the cretaceous age is correct, this is an interesting rock unit as no other cretaceous strata are currently known from northwestern utah. the red beds are likely similar in age to part of the gannett group rocks (northeastern utah), kelvin formation (central wasatch mountains), and the cedar mountain formation (eastern utah). 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 also of interest is the location of the red beds; these rocks are wedged between two thrust sheets of the sevier fold-thrust belt (see, for example, willis, 1999; hintze and kowallis, 2009, p. 68). th e cedar thrust fault (carrying pennsylvanian oquirrh group rocks, map unit ) lies directly on the red beds and below are permian rocks (gerster and plympton formations, map unit pgp) belonging to a diff erent unnamed thrust sheet. a tremendous thickness of conglomerate and sandstone were shed eastward of the fold-thrust belt and are preserved in northeastern utah, but this relatively small expanse of red beds was likely deposited in a piggyback (wedgetop) basin of the orogenic belt. th e existence of the red beds constrains emplacement of the cedar thrust sheet to no earlier than early cretaceous (aptian). th rust sheets of the sevier orogenic belt were emplaced from about 130 to 50 ma (early cretaceous to early eocene) across northern utah (yonkee and weil, 2011). 112º56'30" w113º00' 40º37'30" n 40º33'30" n qes • b # # # # # # # # # # # ! ! # # # ! ! ! ! # # # # # ! ! ! # # ! ! ! ! ! ! ! ! ! ! ! ! ! # ! ! ! ! #! ! ! ! ! ! # # # ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! b ! ! ! # #! ! ! ! b b # # ! ! ! ! ! # # # # # # # # # # # # # # # # # ! # # ! ! ! ! ! ! ! ! ! ! # # # # # # # # # e # # ! !!!! ! ! ! !# #! ! ! # ! ! ! ! ! ! ! ! # # # # # ! ! ! ! ! e e e e e e e e ! ! ! ! ! ! ! e ! e ! ! ! ! ! e e e e ! ee e # # # # # # # # ! ! ! e ! ! ! ! ! ! ! ! ! ! # # # # # # # ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! # # # # # # # # # # # # # # # ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! !! ! !! ! !! ! !! j j • b b b b b b gb b b b b b b b b b b b bbbb b b b 6 6 6 b b b b b b bbb b bb b b b b b b ’ a r ed la m sp r in g fa u lt ce da r qlg *o l qal *oqal ppm *o ppg *o qafy qal *olc qalks *o qal q *o pgp ppg *o psl qal qaf ppg psl q qal *o *o qa *o ks q * *o psl q qa ks psl ppm pgp *olc *osc *osc 6 16 8 10 13 15 8 4 9 3 5016 15 11 13 16 3 15 20 40 25 45 22 18 15 12 30 10 42 t46 a a' q th rust ppm # # # # # # ## # ! ! !#!#!#!#! ! # ! # ! # ! !# !# ! # ! # ! ! ! !! ! !! ! b b# b# bbb’#’#’’’#’#’’’b’b# b#’# b# b’bb’b fa u lt q *olc *osc 15 40 2525’25’t46 b geosite psl 3000 2000 1000 -1000 0 a west el ev at io n in m et er s cedar thrust fault p p p ks ks *olc *olc psl psl psl oquirrh group, undivided p*o b < < < < a'east geologic map cross section q ua t. c re t. pe rm ia n pe nn . 0 1 1km0 mile ppg $ 112º56'30" w113º00' 40º37'30" n 40º33'30" n qes • u-pb detrital zircon sample b # # # # # # # # # # # ! ! # # # ! ! ! ! # # # # # ! ! ! # # ! ! ! ! ! ! ! ! ! ! ! ! ! # ! ! ! ! #! ! ! ! ! ! # # # ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! b ! ! ! # #! ! ! ! b b # # ! ! ! ! ! # # # # # # # # # # # # # # # # # ! # # ! ! ! ! ! ! ! ! ! ! # # # # # # # # # e # # ! !!!! ! ! ! !# #! ! ! # ! ! ! ! ! ! ! ! # # # # # ! ! ! ! ! e e e e e e e e ! ! ! ! ! ! ! e ! e ! ! ! ! ! e e e e ! ee e # # # # # # # # ! ! ! e ! ! ! ! ! ! ! ! ! ! # # # # # # # ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! # # # # # # # # # # # # # # # ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! !! ! !! ! !! ! !! j j • b b b b b b gb b b b b b b b b b b b bbbb b b b 6 6 6 b b b b b b bbb b bb b b b b b b ’ a r ed la m sp r in g fa u lt ce da r qlg *o l qal *oqal ppm *o ppg *o qafy qal *olc qalks *o qal q *o pgp ppg *o psl qal qaf ppg psl q qal *o *o qa *o ks q * *o psl q qa ks psl ppm pgp *olc *osc *osc 6 16 8 10 13 15 8 4 9 3 5016 15 11 13 16 3 15 20 40 25 45 22 18 15 12 30 10 42 t46 a a' q th rust ppm geosite psl 3000 2000 1000 -1000 0 a west el ev at io n in m et er s cedar thrust fault p p p ks ks *olc *olc psl psl psl oquirrh group, undivided p*o b < < < < a'east geologic map cross section q ua t. c re t. pe rm ia n pe nn . 0 1 1km0 mile ppg $ q 112º56'30" w113º00' 40º37'30" n 40º33'30" n qes • b # # # # # # # # # # # ! ! # # # ! ! ! ! # # # # # ! ! ! # # ! ! ! ! ! ! ! ! ! ! ! ! ! # ! ! ! ! #! ! ! ! ! ! # # # ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! b ! ! ! # #! ! ! ! b b # # ! ! ! ! ! # # # # # # # # # # # # # # # # # ! # # ! ! ! ! ! ! ! ! ! ! # # # # # # # # # e # # ! !!!! ! ! ! !# #! ! ! # ! ! ! ! ! ! ! ! # # # # # ! ! ! ! ! e e e e e e e e ! ! ! ! ! ! ! e ! e ! ! ! ! ! e e e e ! ee e # # # # # # # # ! ! ! e ! ! ! ! ! ! ! ! ! ! # # # # # # # ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! # # # # # # # # # # # # # # # ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! !! ! !! ! !! ! !! j j • b b b b b b gb b b b b b b b b b b b bbbb b b b 6 6 6 b b b b b b bbb b bb b b b b b b ’ a r ed la m sp r in g fa u lt ce da r qlg *o l qal *oqal ppm *o ppg *o qafy qal *olc qalks *o qal q *o pgp ppg *o psl qal qaf ppg psl q qal *o *o qa *o ks q * *o psl q qa ks psl ppm pgp *olc *osc *osc 6 16 8 10 13 15 8 4 9 3 5016 15 11 13 16 3 15 20 40 25 45 22 18 15 12 30 10 42 t46 a a' q th rust ppm geosite psl 3000 2000 1000 -1000 0 a west el ev at io n in m et er s cedar thrust fault p p p ks ks *olc *olc psl psl psl oquirrh group, undivided p*o b < < < < a' ks pgp ppm ppg psl *osc *olc cretaceous strata gerster and plympton fms. oquirrh group, sandstone and siltstone unit oquirrh group, sandstone and siltstone unit oquirrh group, sandstone east geologic map cross section various quaternary deposits various quaternary deposits various quaternary phosphoria fm., meade peak mbr. phosphoria fm., meade peak mbr. phosphoria fm., park city fm., grandeur mbr. park city fm., grandeur mbr. park city fm., sandstone, limestone and dolomite sandstone, limestone and dolomite sandstone, limestone oquirrh group, limestone unit oquirrh group, limestone unit oquirrh group, p q ua t. c re t. pe rm ia n pe nn . 0 1 1km0 mile ppg $ 112º56'30" w113º00' 40º37'30" n 40º33'30" n qes • contact normal fault thrust fault lineament sedimentary bedding attitude spring u-pb detrital zircon sample 15 b # # # # # # # # # # # ! ! # # # ! ! ! ! # # # # # ! ! ! # # ! ! ! ! ! ! ! ! ! ! ! ! ! # ! ! ! ! #! ! ! ! ! ! # # # ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! b ! ! ! # #! ! ! ! b b # # ! ! ! ! ! # # # # # # # # # # # # # # # # # ! # # ! ! ! ! ! ! ! ! ! ! # # # # # # # # # e # # ! !!!! ! ! ! !# #! ! ! # ! ! ! ! ! ! ! ! # # # # # ! ! ! ! ! e e e e e e e e ! ! ! ! ! ! ! e ! e ! ! ! ! ! e e e e ! ee e # # # # # # # # ! ! ! e ! ! ! ! ! ! ! ! ! ! # # # # # # # ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! # # # # # # # # # # # # # # # ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! !! ! !! ! !! ! !! j j • b b b b b b gb b b b b b b b b b b b bbbb b b b 6 6 6 b b b b b b bbb b bb b b b b b b ’ a r ed la m sp r in g fa u lt ce da r qlg *o l qal *oqal ppm *o ppg *o qafy qal *olc qalks *o qal q *o pgp ppg *o psl qal qaf ppg psl q qal *o *o qa *o ks q * *o psl q qa ks psl ppm pgp *olc *osc *osc 6 16 8 10 13 15 8 4 9 3 5016 15 11 13 16 3 15 20 40 25 45 22 18 15 12 30 10 42 t46 a a' q th rust ppm geosite psl 3000 2000 1000 -1000 0 a west el ev at io n in m et er s cedar thrust fault p p p ks ks *olc *olc psl psl psl oquirrh group, undivided p*o b < < < < a'east geologic map cross section q ua t. c re t. pe rm ia n pe nn . 0 1 1km0 mile ppg $ figure 2. part of the tooele 30' x 60' geologic map and cross section, from clark and others (2017, 2019 in review). d.l. clark northern cedar mountains red beds 5 how to get there take i-80 west from salt lake city. continue to exit 56 aragonite. at top of ramp, turn left (south) cross i-80 on paved road to the clean harbors hazardous waste incinerator facility. turn left (east) before incinerator onto gravel road. continue east, to blm signboard and sign for cedar mountains wilderness. turn right (south) on gravel road (cedar mountains road). continue 12 miles (19 km) south, cross cattle guard. continue and pass by wash bottom. the road is too narrow for safe parking, so continue 0.25 mile (0.4 km) south, turn around on dirt road to right (west) leading to a water trough. come back to the north and park along the east side of the road before the wash bottom (latitude 40.57995 °n, longitude 113.00459 °w, wgs84). enter cedar mountains wilderness area and hike east along drainage about 1 mile (1.6 km) to exposure of red beds (~ 40.58470 °n, 112.99034 °w, wgs84). acknowledgments geologic mapping was supported by the statemap component of the national cooperative geologic mapping program. i thank ugs staff zach anderson and don deblieux, and uga editors mark milligan and robert biek for reviews of this paper. thanks also to lori steadman and martha jensen (ugs) for preparation of the fi gures. basia matyjasik (ugs) compiled the geographic information system data for the tooele 30' x 60' geologic map. references clark, d.l., oviatt, c.g., and dinter, d.a., 2017, interim geologic map of the tooele 30' x 60' quadrangle, tooele, salt lake, and davis counties, utah: utah geological survey open-file report 669dm, 43 p., 3 plates, contains gis data, scale 1:62,500. clark, d.l., oviatt, c.g., and dinter, d.a., 2019 in review, geologic map of the tooele 30' x 60' quadrangle, tooele, salt lake, and davis counties, utah: utah geological survey map, scale 1:62,500. hintze, l.f., and kowallis, b.j., 2009, geologic history of utah: brigham young university geology studies, special publication 9, 225 p. maurer, r.e., 1970, geology of the cedar mountains, tooele county, utah: salt lake city, university of utah, ph.d. dissertation, 184 p., 10 plates, scale 1:43,700. utah geological survey and o’sullivan, p.b., 2017, u-pb zircon geochronology results for the angle, donkey flat, farnsworth peak, fort douglas, and quincy spring quadrangles, utah: utah geological survey open-file report 660, variously paginated, online, http://ugspub.nr.utah.gov/publications/ open_fi le_reports/ofr-660/ofr-660.pdf. willis, g.c., 1999, th e utah thrust system—an overview, in spangler, l.e. and allen, c.j., editors, geology of northern utah and vicinity: utah geological association publication 27, p. 1-9. yonkee, w.a., and weil, a.b., 2011, evolution of the wyoming salient of the sevier fold-thrust belt, northern utah to western wyoming, in sprinkel, d.a., yonkee, w.a., and chidsey, t.c., jr., editors, sevier thrust belt—northern and central utah and adjacent areas: utah geological association publication 40, 56 p. *olc ks pgp cedar thrust fault dz sample landslide figure 4. view to northwest of red beds. ks = cretaceous strata, = oquirrh group, limestone unit of cedar thrust sheet. pgp=gerster and plympton formations. dz is u-pb detrital zircon sample location. figure 3.view northwest of lower cretaceous red beds from the quincy spring area of the cedar mountains. uga-geosite-atwood-ophir-anticline.indd 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors genevieve atwood, amanda r. mcgarry, and peg l. alderman earth science education 30 u street, salt lake city, utah 84103 genevieveatwood@comcast.net cover image: cross section of the ophir anticline. ophir anticline, tooele county a 30-minute drive through 300 million years of bedrock 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e geosites refl ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: atwood g., mcgarry a.r., and alderman, p.l., 2019, ophir anticline, tooele county—a 30-minute drive through 300 million years of bedrock, in milligan, m., biek, r.f., inkenbrandt, p., and nielsen, p., editors, utah geosites: utah geological association publication 48, 10 p., https://doi. org/10.31711/geosites.v1i1.46. presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president g. atwood, a.r. mcgarry, p.l. alderman ophir anticline 3 introduction anticlines are not unusual features, but the ophir anticline is exceptional. regions shortened laterally by compressional tectonics have folds and/or faults including anticlines, synclines, monoclines, and thrust faults (figure 1). however, most anticlines are buried or partially buried, visible on land only as discontinuous snippets in road cuts or visible from above as one flies across terrain too regional to be viewed in entirety from land. the ophir anticline is unusual as its limbs and crest can be viewed in their entirety, exposed in cross-section on ophir canyon’s walls. ophir canyon provides evidence that the compressional tectonics associated with the ophir anticline long-preceded active-extensional tectonics associated with today’s basin and range topography of western utah. this geosite is the 2-mile (3.2 km) stretch of ophir canyon that exposes the anticline (figure 2). gps location for the parking area at the intersection of state road 73 and ophir canyon road: nad83 40.35089° n / -112.30965° w suggested driving directions paved roads make the geosite easily accessible (figure 3). from salt lake city, travel west on i-80, then south on utah route 36 through tooele. at “pennys junction” south of stockton, bear left (to the east) and continue south on utah route 73 to a small parking area at the intersection of route 73 and ophir canyon road. from utah county, travel west on route 73 via lehi, cedar fort, fairfield, and five mile pass to the small parking area at the intersection of route 73 with ophir canyon road. the intersection is well marked. from this parking area, one can view the western range front of the oquirrh mountains and ophir canyon. the foreground terrain of alluvial fans emerging into the valley contrasts abruptly with the layered gray bedrock of the range. the parking area at 5650 ft a.s.l. (1720 m) lies a few hundred feet above the bonneville level of lake bonneville, the lake that occupied much of rush valley and most of western utah approximately 30,000 to 13,000 years ago. figure 1. five expressions of compressional tectonics. compression shortens portions of earth’s crust. the dips (tilts) of bedding planes define the structure of an anticline (concave upward), syncline (concave downward), and monocline (a single dipping limb). thrust faults move older rock over younger rock. figure 2. location of the ophir anticline geosite and vicinity. base: u.s. geological survey, (2019). 1. green arrow: parking area at the intersection of utah road 73 and ophir canyon road, gps location, nad83 40.35089° n / -112.30965° w 2. red line: route from parking area through the ophir anticline to the community park and turnaround east of ophir. 3. ophir anticline exposed on canyon wall. 4. pha, phb, and phc indicate locations of the three photographs of figure 6 that document contrasting orientations of ophir anticline’s bedrock layers. 5. the ophir mining district surrounds ophir and extends into upper drainages of ophir creek. 6. lion hill directly south of ophir is the prominent geographic feature of figure 8, the geologic cross-section of the ophir anticline. 7. yellow-highlighted line: general location of the geologic cross-section of figure 8. figure 3. location of the ophir anticline geosite from the wasatch front. base: u.s. geological survey (2019). access by interstate and state highways to the geosite indicated by purple line from utah county and by green line from i-80 west of salt lake county. g. atwood, a.r. mcgarry, p.l. alderman ophir anticline 4 physical location ophir canyon is located on the western flank of the oquirrh mountains, a range of the eastern basin and range physiographic province. basins surround the oquirrh mountains including rush valley west of ophir canyon (figure 4). ophir canyon and the oquirrh mountains lie west of utah’s hingeline, the north-south trending structural zone also called the wasatch line, and along the projection of the east-west uinta trend or uinta axis discussed in stokes (1986), hintze and kowallis (2009), and clark and others (2019). the ophir anticline and mining district lie along an east-west mineral belt that continues east to park city and west to gold hill (krahulec, 2018). for over a century, geologists have worked the area and puzzled over its geologic relationships (figure 5). ophir anticline geosite the drive from the parking area (intersection of route 73 and ophir canyon road) through the anticline is short with attendant hazards including rockfalls, wildlife, and driving distracted by geology. consider traveling the route twice: once for the lay of the land, and again to pull onto the road’s shoulder to take photos and recognize patterns of bedrock layering, specifically the limbs and crest of the ophir anticline (figure 6). the route has three sections: (1) approximately 2 miles (3.2 km) across alluvial fans to the ophir anticline, (2) approximately 1.5 miles (2.4 km) through the west limb of the anticline to its crest opposite the community of ophir, and (3) approximately 0.5 mile (0.8 km) via a well-maintained dirt road into the east limb of the anticline. consider parking at the turnaround at open fields beyond the community park to investigate the bedrock. alluvial fans and stream deposits dominate the terrain between the parking area and the geosite. active (holocene, within approximately the past 12,000 years) normal faults cut the alluvium between the parking area and the range front (figure 7). kirby (2012) summarizes several hazards-related investigations that present evidence that active and older normal faults cut and define the range front (bucknam, 1977; everett and kaliser, 1980; barnhard and dodge, 1988). driving east into the canyon, travelers cross abruptly into gray, layered, west-dipping bedrock of the west limb of the ophir anticline. following the west-dipping strata eastward results in traveling through older and older layers of bedrock until arriving at ophir. approaching ophir, the tilting of bedrock layers diminishes to horizontal at the anticline’s axis. a half-mile (0.8 km) east of ophir, at the turnaround and fields, the rock layers dip to the east defining the eastern limb of the anticline. patterns of bedrock orientation are shown in photographs (figure 6) and in geologic figure 4. physical features of the oquirrh mountains and environs. base: u.s. bureau of land management (1993). the red “gps” marks the intersection of state road 73 and ophir canyon road. olive green indicates urban development. figure 5. geologic context of the ophir anticline geosite. base: thelin and pike (1991). colors highlight two major structural features of utah geology. the green band highlights the north-south trending utah hingeline. the red band highlights the east-west uinta trend. these structural zones have accommodated and influenced significant changes in utah’s crust (hintze and kowallis, 2009) and persisted through the 500 million-year bedrock history of the oquirrh mountains (clark and others, 2019). g. atwood, a.r. mcgarry, p.l. alderman ophir anticline 5 figure 7. geologic map of features west of the ophir anticline geosite. base: kirby (2012). the parking area at the gps location provides a view of relationships of bedrock (blue) versus sediments (yellow). patterns of layered bedrock exposed on hillslopes indicate post-depositional tilting by compressional tectonics. red lines mark mapped range-front extensional faults. the green hachured line marks the general trend of a compressional fault now buried by far-younger sediment. young (red) faults cut valley-fill sediments as well as bedrock. therefore they are younger than those units. these range-front faults cut the folded bedrock of the ophir anticline at the mouth of the canyon. therefore, range-front faulting is far younger than the compressional tectonics that folded bedrock into the ophir anticline. figure 6. contrasting orientations of ophir anticline’s bedrock layers. modified from atwood (2018). three photographs show contrasting patterns of bedrock layering in ophir canyon. the patterns define the essential structure of an anticline. at pha (see figure 2 for location) the west limb of the ophir anticline, the layers dip west, meaning, trend downward toward the west. a black line highlights the trend to contrast bedrock surfaces from utility lines. at phb, along the crest of the anticline on the face of lion hill, the bedrock layers’ orientation is nearly horizontal. at phc, east of ophir, the layers dip to the east. the canyon exposes bedrock almost continuously providing multiple places to examine layers and appreciate their structure. crest of the anticline. ophir, looking south. genevieve atwood for scale and representation of the nearly horizontal layering of the bedrock. west limb of the anticline. looking north. note peg alderman (lowest center) for scale and representation of westward tilt of bedrock. east limb of the anticline. looking north from the park east of ophir. peg alderman for scale and representation of eastward tilt of the bedrock. phbpha phc g. atwood, a.r. mcgarry, p.l. alderman ophir anticline 6 cross-section (figure 8). the cross-section’s blue-highlighted layering conveys the anticline’s simple form. the numerous faults that displace generally-gray dipping bedrock layers against generally-gray dipping layers makes correlation of rock units challenging. ophir canyon provides a window into the mountain range. near ophir it exposes cambrian limestone, the oldest bedrock exposed in the oquirrh mountains (parry, 2016). the two geologic maps of figure 9 present the coarse-scale generalized message of the anticline and the finer-scale detailed geology important for exploration geologists. both maps show the anticline’s structure and raise intriguing questions, such as, why the rock record is missing between the significantly younger bedrock of the anticline (m=mississippian) and the older bedrock (c=cambrian) that cores the anticline. ophir, historic mining town and mining district the community of ophir takes pride in its mining heritage. the historic district, marked by a flagpole at the town’s center, includes original buildings and historical markers. the old schoolhouse, notable for its history, has spectacular views south to lion hill and the eroded crest of the ophir anticline (central photograph of figure 6). evidence of the district’s mining operations include adits, tunnels, waste rock, and rail grades. the numerous mining districts of the oquirrh mountains, including the bingham ore deposits and those of ophir, have produced remarkable riches. according to w.l. stokes, “the oquirrh mountains are good contenders for the title of ‘richest mountains on earth’” (1986, p. 171). according to utah mining districts, (krahulec, 2018, p. 119): figure 8. geologic cross-section of the ophir anticline. base: tooker and roberts (1998). this cross-section has been modified substantially from tooker and roberts (1998) to emphasize the general structure of the ophir anticline. the figure does not show tooker and robert’s labels for the age of stratigraphic units. the figure deemphasizes their interpretation of thrust faults (interpretations not shared by clark and others, 2019). tooker and robert’s interpretation is shown faintly in green. what is shown, highlighted in blue are bedrock layers. their general orientation defines the ophir anticline. the yellow-highlighted lines represent younger, normal faults. these range-front faults cut the anticline and form the boundary between rush valley and the oquirrh mountains. the colors of the cross-section tie to figure 10 – geologic history of ophir and vicinity. blue lines represent bedrock of chapter 3 – shallow seas and chapter 4 – broad basins. green lines represent effects of compressional tectonics during chapters 6 – scrunch and chapter 7 – seven up that folded the bedrock layers into the ophir anticline. yellow lines represent faults and bedrock displacement due to extensional tectonics of chapter 9 – now, stretching to the west. the ophir district is about the tenth most productive in utah. ophir is credited with production of about 2.8 million tons of ore averaging 6.2% pb, 1.5% zn, 0.8% cu, 237 ppm ag, and 0.21 ppm au, recovered. total district metal production at modern metal prices is estimated at $877 million. the ophir hill mine in ophir canyon is the largest producer in the district with about 1.5 million tons of production followed by hidden treasure mine at 300,000 tons (rubright, 1978). explanation of the unusual structures and ore deposits of the ophir mining district (krahulec, 2018; parry, 2004) is beyond the scope of this paper. however, the ophir anticline, its structure, and associated fracturing of the bedrock favored later mineralization associated with igneous intrusions and fluid interactions (kirby, 2012). the mining district aligns with the crest of the anticline. the layered nature of the district’s ore prompted innovative mining methods and impressive legal implications lasting to the present (parry, 2004). processes associated with igneous activity deposited ore in the bedrock of the ophir anticline. the presence of the anticline and associated faulting may have influenced mineralization, but mineralization of the ophir district did not influence the formation of the anticline. structure, stratigraphy, and geologic history the ophir anticline has intrigued geologists for over a century (butler and others, 1920; gilluly, 1932; laes and others, 1997; tooker and roberts, 1998; handwerger and others, 1999; kirby, 2012; clark and others, 2012; clark and others, 2019). novice and g. atwood, a.r. mcgarry, p.l. alderman ophir anticline 7 expert geologists alike can enjoy the puzzle of structure, bedrock units, and present landforms. specifically, the ophir anticline contradicts a misconception that the mountain building that caused the ophir anticline and similar folded bedrock of western utah is the mountain building of today’s basin and range topography. the compressional tectonics that formed the ophir anticline predates the extensional tectonics of today’s basins and ranges by over 100 million years. the geosite’s geologic history can be told as chapters using terminology of hintze (2005) and hintze and kowallis (2009). that history is summarized in figure 10. working backwards through time, the chart summarizes how extensional tectonics of chapter 9 – now! creates the feature, the oquirrh mountains. mineralization associated with tectonics of chapter 8 – impressive igneous deposited ores of the ophir district and oquirrh mountains. uplift due to tectonic forces of chapter 7 – seven up! raised the region high. compressional-tectonic forces of chapter 6 – scrunch… created the ophir anticline thousands of feet below the land surface. relatively quiescent tectonic forces of chapter 4 – broad basins and chapter 3 – shallow seas created environments where sediments accumulated and became thousands of feet of layered-gray bedrock of today’s western oquirrh mountains. age relationships superposition of layers of bedrock and cross-cutting relationships of folding and faulting are the evidence for interpreting the geologic history of the ophir anticline geosite. retelling that history from oldest to youngest may reinforce recognition of age relationships observable in ophir canyon. specifically, the compressional tectonics that created the ophir anticline preceded present-day extensional tectonics by millions of years. figure 10 summarizes the geologic history of ophir utah three ways: as millions of years, as traditional breaks of time with traditional geologic names, and as chapters of utah’s geologic history. figure 11 tells the sequence in sketches. the sketches illustrate critical stages of development of the ophir anticline geosite: (1) processes of sedimentation and lithification formed the bedrock for the anticline during chapter 3 – shallow seas and chapter 4 – broad basins. (2) compressional tectonics folded and uplifted the anticline during chapter six – scrunch and chapter seven – seven up! (3) extensional tectonics created the topographic relief for ophir canyon and exposure of the anticline during chapter nine – now! extensional tectonics westward. key mgbu,mgbs,mgbl mh,md,mg,mdp cly, cb, ch, cap parry clark, and others figure 10 from: parry, 2016 from: clark and others, 2019 southern oquirrh mountains detail for ophir and vicinity southern oquirrh mountains from: parry, 2016 from: clark and others, 2019 detail for ophir and vicinity figure 9. geologic maps of ophir canyon and vicinity. the map on the left by parry (2016), generalizes the geology and shows, simply, how the oldest exposed bedrock of the oquirrh mountains (c=cambrian) cores the anticline. the detailed map by clark and others (2019) shows the anticline as a generally curvilinear complex of bedrock units broken and displaced by faulting. together, the maps give regional structural context plus details of interest to geologists, such as, those exploring for ore, and rockhounds hunting for fossils. the key correlates the units of the two geologic maps with figure 10. figure 10 presents the same history but uses “chapters” specific to utah’s geologic past. for example, c=cambrian of the maps of figure 9 is part of chapter 3 – shallow seas of figure 10. key mgbu,mgbs,mgbl mh,md,mg,mdp cly, cb, ch, cap parry clark, and others figure 10 from: parry, 2016 from: clark and others, 2019 southern oquirrh mountains detail for ophir and vicinity g. atwood, a.r. mcgarry, p.l. alderman ophir anticline 8 the following discussion addresses the misconception that the compressional folding of the ophir anticline also created today’s oquirrh mountains, a classic range of basin and range topography. this summary draws on hintze and kowallis (2009). bedrock of the ophir anticline: the gray bedrock of cambrian, mississippian, and pennsylvanian age (shown in photographs of figure 6 and in maps of figure 9), began as muddy sediments laid down in shallow seas. shallow-marine fossils encased in what was once limy muds indicate that utah then was below sea level and near the equator (hintze, 2005; hintze and kowallis, 2009). continued deposition resulted in thousands of feet of sedimentary bedrock layers. folding the ophir anticline: over 100 million years after the bedrock was buried, compressional tectonics during cretaceous and early tertiary time shortened and uplifted utah’s crust. in response to this mountain building of the sevier and laramide orogenies, tooele county’s topography resembled mountain ranges as high as today’s andes mountains. at depth, within the mountains’ roots, folding and faulting created compressional features including the ophir anticline. did the compressional tectonic forces that made those mountains and the ophir anticline create today’s oquirrh mountains? the evidence can be unraveled by cross-cutting relationships illustrated in figure 8 – geologic cross-section of ophir anticline and in figure 9 – geologic maps of ophir canyon and vicinity. igneous intrufigure 10. geologic history of ophir, utah and vicinity. modified from atwood and mcgarry (2019). chapter terminology from hintze (2005) and hintze and kowallis (2009). g. atwood, a.r. mcgarry, p.l. alderman ophir anticline 9 sions cut the folded bedrock of the ophir anticline about 36 my ago (kirby, 2012) dating the approximate age of mineralization of the ophir anticline. ophir canyon cuts through mineralized rock layers of the ophir anticline. therefore, the canyon is younger than the intrusions, the mineralization, the compressional tectonics of the anticline, and the bedrock units of the anticline. ophir creek cuts ophir canyon, meaning the creek cuts deep into the oquirrh mountains. streams require topography. ophir creek flows from high terrain to lower terrain, from a range to a basin. basin and range topography necessarily preceded and has evolved with the canyon. the mountains that existed in response to the compressional tectonics that folded the ophir anticline are long gone. range-front faults located at the mouth of ophir canyon cut modern alluvial fans showing that active extensional processes continue to drop the valleys with respect to the mountains (figure 7). extensional forces continue to create the oquirrh mountains, a range of the basin and range physiographic province. the ophir anticline existed before basin and range topography. the compressional tectonic forces that bent buried bedrock into the ophir anticline did not create today’s landform, the oquirrh mountains. this brief history omits discussion of several intriguing questions pertaining to ophir canyon, the ophir mining district, and the oquirrh mountains. why are entire sections of geologic time unrepresented in the rock record? why does the layered structure of ore bodies of the ophir mining district differ from others of the oquirrh mountains? why are the oquirrh mountains exceptionally mineralized? what makes this geosite worth a visit? the relationships of the ophir anticline, its bedrock layers, and how ophir creek exposes the anticline in ophir canyon are all visible from ophir canyon road. when one comprehends the relative ages of ophir canyon’s bedrock, the bedrock’s mineralization, and today’s topography all with respect to the anticline, one can appreciate the generalized geologic history of many ranges of the basin and range physiographic province. acknowledgments the authors are grateful to tooele school district for sponsoring earth science education’s earth science outside field experiences in ophir canyon. john a. canner and tamara j. wambeam assisted with maps. this paper has benefited greatly from informal and formal reviews by william t. parry and donald l. clark. references atwood, g., 2018, ophir canyon tectonics past and present for tooele school district teachers: www.earthscienceeducation. org/ accessed january and february 2019. atwood, g., and mcgarry, a.r., 2019, geologic history of a utah place – generic chart: www.earthscienceeducation.org/ geologic-history-of-utah-for-teachers/ accessed january and february 2019. barnhard, t.p., and dodge, r.l., 1988, map of fault scarps formed in unconsolidated sediments, tooele 1° x 2° quadrangle, northwestern utah: u.s. geological survey miscellaneous field studies map mf-1990, scale 1:250,000. bucknam, r.c., 1977, map of suspected fault scarps in unconsolidated deposits, tooele 2° sheet, utah: u.s. geological survey open-file report 77-495, scale 1:250,000. butler, b.s., loughlin, g.f., heikes, v.c., and others, 1920, the ore deposits of utah: u.s. geological survey professional paper 111, 672 p., 7 plates. clark, d.l., kirby, s.m., oviatt, c.g., 2012, interim geologic map of the rush valley 30' x 60' quadrangle, tooele, utah, and figure 11. critical stages of development of the ophir anticline geosite. modified from atwood (2018). numbers tie to chapters of utah’s geologic past presented in figure 10. tectonics rules! geology is cumulative. our present understanding of the architecture of western utah’s crust includes how ancient zones of weakness remobilize during subsequent chapters of tectonics. the ore of the ophir mining district was affected by the anticline although the compressional forces that folded it into the anticline long-preceded the processes that resulted in mineralization. similarly, the compressional forces that caused the folding of the anticline long-preceded the extensional forces that resulted in basin and range topography including the oquirrh mountains and rush valley. g. atwood, a.r. mcgarry, p.l. alderman ophir anticline 10 salt lake counties, utah: utah geological survey open-file report 593, plate 1, scale 1:62,500. clark, d.l., kirby, s.m., oviatt, c.g., 2019, interim geologic map of the rush valley 30' x 60' quadrangle, tooele, utah, and salt lake counties, utah: national geologic map data base accessed may 10, 2019, https://ngmdb.usgs.gov/prodesc/ proddesc_97595.htm, plate 1, geologic map, scale 1:62,500. everitt, b.l., and kaliser, b.n., 1980, geology for the assessment of seismic risk in the tooele and rush valleys, tooele county, utah: utah geological and mineral survey special study 51, 33 p. gilluly, j., 1932, geology and ore deposits of the stockton and fairfield quadrangles, utah: u.s. geological survey professional paper 173, 171 p., 1 plate, scale 1:62,500. handwerger, d.a., cerling, t.e., and bruhn, r.l., 1999, cosmogenic ¹⁴c in carbonate rocks: geomorphology, v. 27, p. 13-24. hintze, l.f., 2005, utah’s spectacular geology: provo, department of geology, brigham young university, 202 p. hintze, l.f., and kowallis, b.j., 2009, geologic history of utah: provo, brigham young university geology studies, special publication 9, 225 p. kirby, s.m., 2012, geologic map of the ophir quadrangle, tooele county, utah: utah geological survey map 257dm, p. 1-13, 2 plates, scale 1:24,000. krahulec, k.a., 2018, utah mining districts, utah: utah geological survey open-file report 695, 191 p., 1 plate, scale 1:1,000,000. laes, d.y.m., krahulec, k.a., and ballantyne, g.h., compilers, 1997, geologic map of the oquirrh mountains, utah, in john, d.a., and ballantyne, g.h., editors, geology and ore deposits of the oquirrh and wasatch mountains, utah: society of economic geologists guidebook series, v. 29, 1 plate, scale 1:62,500. parry, w.t., 2004, all veins, lodes, and ledges throughout their entire depth geology and the apex law in utah mines: salt lake city, university of utah press, 139 p. parry, w.t., 2016, geology of utah’s mountains, peaks, and plateaus: victoria, british columbia, friesen press, 244 p. stokes, w.l., 1986, geology of utah: salt lake city, utah museum of natural history, utah geological and mineral survey, occasional paper utah museum of natural history no. 6, 280 p. thelin, g.p., and pike, r.j., 1991, landforms of the conterminous united states – a digital shaded-relief portrayal: u.s. geological survey map i-2206, https://pubs.usgs.gov/imap/i2206/, accessed january 2019. tooker, e.w., and roberts, r.j., 1998, geologic map of the oquirrh mountains and adjoining south and western traverse mountains, tooele, salt lake and utah counties, utah: u.s. geological survey open-file map 98-581, 2 plates, scale 1:50,000. u.s. bureau of land management, 1993, state of utah shaded relief map: library of congress www.loc.gov/item/95682023/ accessed january 2019. u.s. geological survey, 2019, national map: https://viewer. nationalmap.gov/advanced-viewer/ accessed january and february 2019. uga-geosite-loope-calf-creek.indd m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors david b. loope and richard m. kettler earth & atmospheric sciences, university of nebraska lincoln, ne 68588-0340 dloope1@unl.edu cover image: view of a concretion along a trail in upper calf creek falls. rinded, iron-oxide concretions in navajo sandstone along the trail to upper calf creek falls, garfield county 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e geosites refl ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: loope, d.b., and kettler, r.m., 2019, rinded iron-oxide concretions in navajo sandstone along the trail to upper calf creek falls, garfi eld county, in milligan, m., biek, r.f., inkenbrandt, p., and nielsen, p., editors, utah geosites: utah geological association publication 48, 7 p., https://doi. org/10.31711/geosites.v1i1.59. presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president d.b. loope and r.m. kettler giant concretions of upper calf creek falls 3 introduction concretions are hard rock masses, usually spheroidal, but commonly oblate or discoidal, that are formed by strongly localized precipitation of minerals in the pores of an otherwise weaker sedimentary rock (see bates and jackson, 1980, for a more extensive definition). the iron-oxide-rich concretions in the jurassic navajo sandstone in southern utah are unusual in two fundamental ways. first, they are cemented by iron oxide (fe2o3, or fe(oh)3); most other concretions are cemented by silica (sio2), calcite (caco3), dolomite (camg(co3)2), or siderite (feco3). second, unlike other concretions, they are not strongly cemented throughout, but instead, the iron oxide is concentrated in a very strongly cemented, sharply defined, exterior rind or shell. in the smaller concretions, the entire interior lacks iron-oxide cement, and is similar to the rock outside the concretion; in the larger concretions, there is a central zone that is strongly cemented by iron oxide, but the sandstone between the central core and the rind has no iron-oxide cement. the processes involved in the formation of the navajo concretions have been extensively studied by geologists ever since the martian rover opportunity imaged small, spheroidal, iron-rich concretions on the surface of mars (moore, 2004). the navajo concretions were quickly recognized as possible terrestrial analogs to the martian “blueberries” (chan and others, 2004). at present there are three main explanations for how these concretions formed. the first interpretation of the navajo concretions (chan and others, 2004; beitler and others, 2005) views the iron-oxide cement and the “rinded” aspect of today’s concretions as primary features—just a thin, three-dimensional band of sandstone (circular when viewed in two dimensions) that had been cemented by iron oxide (chan and others, 2004). a complication of this interpretation is that, because iron is not transported by water containing oxygen, mixing of an iron-bearing water (devoid of oxygen) with a second, oxygenated water mass is required (chan and others, 2004). mixing is typically not required for growth of silica, calcite, dolomite, or siderite concretions. the second interpretation (loope and others, 2010; weber and others, 2012) is that solid, spheroidal concretions were originally well-cemented throughout by siderite (feco3). the rinds formed secondarily, during much later alteration of the older, primary siderite. a weakness of this interpretation is that unaltered siderite has not been found in any navajo concretions. the third interpretation (yoshida and others, 2018) holds that the original spheroidal concretions were solid spheres cemented by calcite (caco3). the rinds formed later when acidic, iron-bearing waters invaded the navajo. the calcite of the concretions buffered the acidic waters, forcing iron oxide to precipitate. a weakness of this interpretation is that it does not explain the localization of iron into a distinct, strongly cemented rind nor the presence of abundant iron in a spatially isolated core zone. all the research cited above was stimulated by the discovery, on mars, of small, iron-rich, spheroidal concretions (the aforementioned blueberries), but the concretions exposed at this geosite are neither small nor spheroidal. they are, however, iron-rich, calcite-free, rinded, and very interesting. we think that the spatial arrangement of their internal features aid interpretation of the small, rinded spheroids. in this article, we briefly present our interpretation (loope and others, 2010; loope and others, 2011; i.e., the second interpretation above) of the very large, non-spheroidal, rinded concretions that are common on the beautiful trail to upper calf creek falls. driving instructions and hiking suggestions the trailhead is 7.4 miles (11.8 km) southwest of boulder, ut and 20.5 miles (32.8 km) northeast of escalante along a portion of utah highway 12 that follows the narrow, flat surface of new home bench, above the canyons of calf creek (to the west) and dry hollow (to the east; figure 1). the turnoff for the trailhead is on the west side of the highway, between mile 80 and mile 81. the parking area is ≈ 500 feet (150 m) from the highway, but getting there can be rough for a low-clearance vehicle (you can instead park along the highway). at its start, the trail to upper calf creek falls drops about 500 feet (150 m) down a steep outcrop of navajo parking to boulder to escalante figure 1. topographic map showing location of parking area and trail head for upper calf creek falls. 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 sandstone, the most difficult part of the short 0.9 mile (1500 meter) one-way hike. take care to step only on bare sandstone, avoiding the ball-bearing-like gravel. in summer, you can depend on calf creek as a great way to cool off; soak your shirt and hat (at least) before hiking out of the canyon. if you wander off the trail, avoid stepping on the dark soil crust that stabilizes the loose sand. gps location of trailhead: 37°51'34" n, 111°26'18" w (wgs84) geologic context boulder-strewn new home bench is an example of inverted topography. about one million years ago, major flood events deposited large boulders (derived from the aquarius plateau to the west) in a stream valley cut into the navajo sandstone (marchetti and others, 2012). the boulders effectively armored the streambed, preventing further erosion of the bedrock below the boulders. soon thereafter, erosion of the bedrock commenced on both the east and west sides of the boulder accumulation; one million years of erosion produced the two deep canyons of calf creek and dry hollow, and stranded the (inverted) flat surface between them (figures 1 and 2). the navajo sandstone, well-exposed in the canyons below new home bench, was deposited in the early jurassic (200 million years ago) by very large, southward migrating sand dunes (kocurek and dott, 1983). even though the navajo has not been tilted, most of the bedding in the navajo slopes southward at about 20-25°. these layers of sand were not deposited parallel to the floor of the giant dune field. instead they were deposited on the dunes’ downwind slopes at angles as steep as 32° (the angle of repose of dry sand); compaction of the sandstone formation during burial under thousands of feet of younger rock diminished the slopes we see today by about 20% (hunter, 1981). iron-oxide-cemented concretions shortly after you reach the base of the steep descent of the trail from the parking area, you will start to see black, iron-oxide-cemented slabs (some displaying bedding, and some with connected corners like boxes; figures 3 and 4). also present along the trail are to boulder to escalante parking new hom e bench 12 jn jn qa qb qb dry hollow 1 mile quaternary alluvium quaternary boulder deposits jurassic navajo sandstone unconsolidated bedrock qa qb jn calf creek upper calf creek falls n figure 2. geologic map showing distribution of unconsolidated deposits and bedrock in the vicinity of upper calf creek falls. figure 3. slabs of rinds from broken concretions. note that some slabs connect forming “boxworks” (see figure 9, part 3 for origin). figure 4. dense piece of concretion rind that shows bedding (black arrow) and a central fracture (white arrow; see figure 9, part 3). d.b. loope and r.m. kettler giant concretions of upper calf creek falls 5 gray, discoidal masses of bedded sandstone also cemented by iron oxide (figure 5). these are texturally distinct from the dark gray (also iron-rich), non-bedded volcanic boulders that have rolled down from new home bench; you should be able to see layering and feel the evenly-sorted grains of relatively fine sand in the remains of the concretions (but not in the volcanic boulders). the slabs are fragments of the dense rinds and the discoids were the central core zones of concretions. along the trail, there are several examples of concretions that are in place and sufficiently eroded to show their interior structure (figures 6 and 7). note that some iron-rich rinds follow fractures in the concretions. recall that all three interpretations of the ironrich navajo concretions (see introduction) have them forming in bedded (explaining the texture of the rinds and discoids) rock, not loose sediment (explaining why the iron accumulations followed fractures). figure 8a shows a large, nearly intact concretion, and figure 8b shows a microscopic view of a core stone with siderite-shaped iron-oxide crystals (pseudomorphs) that provide strong evidence for the former presence of siderite. the iron that remains in the core of large concretions provides a major clue to the origin and evolution of all rinded concretions (figure 9). figures 10 and 11 explain the role that iron-oxidizing microbes played in initiation and thickening of the densely cemented rinds. many of the minerals that form in the deep, oxygen-deprived subsurface are unstable near earth’s land surface and atmosphere. before uplift of the colorado plateau, the navajo sandstone was buried by thousands of feet of younger sedimentary rocks. john wesley powell’s research team first realized that many thousands of feet of sedimentary rocks had been stripped away during uplift of the colorado plateau (dutton, 1880). we now know that canyon cutting and uplift in this area accelerated about six million years ago (marchetti and others, 2012). if siderite crystals formed the concretions while the navajo sandstone was deeply figure 5. core stone representing the central portion of a broken concretion. white arrows show bedding planes. inset shows scattered pseudomorphs (see text). figure 6. horizontal view of an in-place concretion that is composed of a boxwork of dense rinds; core has been completely removed by weathering. site is along the trail, 250 feet (75 m) from upper calf creek falls. 37°51’17” n, 111°27’04 w. figure 7. downward view of weathered, in place concretion, showing boxwork of rinded sandstone slabs. same site as figure 6. 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 a b cs cs cs cs r r r figure 8. a) large, nearly intact, rinded concretion on the rim of dry gulch (n37°50’44.9”; w111°25’29.4). rinds (r) plate the concretion’s perimeter, but some have fallen away, exposing cores (cs). note that rinds are also developed along joints (arrows). hat is 14.5 inches (37 cm) wide. b) photo taken of a thin section through a microscope showing the fine-scale composition of the core of a large calf creek concretion. white is quartz, black is iron oxide, blue is open pore space. arrows point to rhombic (diamond-shaped), iron-oxide in the shape of (pseudomorphs after) siderite crystal. width of field of view is 0.015 inches (0.37 mm; see loope and kettler, 2015, figure 2) and loope et al. (2011, figure 5). figure 10. microbes in the iron-oxide-cemented rind of a concretion. a) arrows point to fuzz in pore space between sand grains at the inner edge of a rind. white is quartz, black is iron-oxide cement. field of view ~500 microns. b) scanning electron microscopy (sem) image of microbes (about 4 microns in diameter) on the surface of a sand grain; the surface of a second sand grain (in the background) is also covered by microbes. 1 2 3 up lif t iron oxide, many joints no siderite siderite, no joints no iron oxide dense rind ox yg en at io njoints moat core stones with rhombic pseudomorphs after siderite rhombic crystals of siderite form concretion joints cut bedded sandstone matrix and concretion bedding figure 9. interpretation of the evolution of a large, rinded concretion: 1) in deep, oxygen-free pore water, siderite crystals grow, forming a concretion in bedded, unfractured navajo sandstone (size of crystals greatly exaggerated). 2) with uplift and erosion of the colorado plateau, joints break the surrounding sandstone and the concretion; 3) in shallow, oxygenated water near the land surface, siderite starts to dissolve into the only remaining oxygen-free water (inside the concretion). dissolved iron diffuses outward to the perimeter and to the joints. at these locations where oxygen-free water meets oxygenated water, rinds form and thicken. moats represent space occupied by pseudomorphs that were dissolved, providing iron for rind growth. if the concretion remains water saturated, rinds continue to thicken, the moats to widen, and the cores to shrink. when the water table dropped below large concretions, iron could no longer migrate, so siderite crystals were oxidized in place, forming pseudomorphs. in smaller concretions, all the siderite dissolved before the water table dropped below the concretion, so no pseudomorphs remain in them (they are just rind and moat; see figure 11d). modified from loope et al. (2016, figure 23), see also loope and kettler (2011, figure 3). buried, those crystals would have been oxidized about one million years ago, when they encountered oxygen-bearing groundwater (beneath the stream that carried the boulders that now cap new home bench; marchetti and others, 2012). the iron oxide in the concretions along calf creek do not bear sufficient uranium for u-th/he radiometric dating, but those along the canyon of russell gulch in western zion national park do have enough uranium (p.w. reiners, personal communication, 2015). analysis of samples from both the rinds and cores of zion concretions showed that the iron oxide crystallized about half a million years ago (loope and others, 2016). that timing fits well with canyon-cutting rates that are based on one-million-year-old lavas on the rim of russell gulch, and puts siderite alteration, rind growth and, core oxidation at shallow depth. the presence of rhombic, iron-oxide pseudomorphs in concretion cores (figure 8b) is a strong indication that iron oxide was not a primary cement: rhombic crystals of siderite that formed in the relatively deep subsurface were the initial cement of these concretions. siderite crystals were unstable in shallow groundwater and many (but not all) were dissolved to form the rind; those that didn’t contribute to the rind were oxidized in place, above the water table. further, the overall distribution of iron in the large concretions is inconsistent with alteration of an iron-free (calcite) precursor (see third interpretation in the introduction). entry of iron from outside a pre-existing concretion cannot explain the rind-moat-core structure of the concretions. in many of the largest concretions most of the iron lies in their cores (figure 8a), and each iron-rich core is surrounded by an iron-free moat, which, in turn, is surrounded by very-strongly cemented rind (figure 9). admittedly, siderite crystals have not been found in the concretions. we attribute this to their (former) location within a relatively porous and permeable matrix—this guaranteed full exposure of the crystals to oxidizing water. like all science, our interpretation (and all three hypotheses for the navajo concretions) could be wrong. d.b. loope and r.m. kettler giant concretions of upper calf creek falls 7 acknowledgments we thank peter mozley and edward simpson for providing helpful reviews of this paper. references bates, r.l., and jackson, j.a., 1980, glossary of geology, second edition: american geological institute, 749 p. beitler, b., parry, w.t., and chan, m.a., 2005, fingerprints of fluid flow: chemical diagenetic history of the jurassic navajo sandstone, southern utah, u.s.a: journal of sedimentary research, v. 75, p. 547–561. chan, m.a., beitler, b., parry, w.t., ormo, j., and komatsu, g., 2004, a possible terrestrial analogue for haematite concretions on mars: nature, v. 429, p. 731-734. dutton, c.e., 1880, report on the geology of the high plateaus of utah: geographical and geological survey of the rocky mountain region (u.s.): government printing office, 307 p. hunter, r.e., 1981, stratification styles in eolian sandstones: some pennsylvanian to jurassic examples from the western interior u.s.a., in ethridge, f.g., and flores, r.m., editors., recent and ancient nonmarine depositional environments: models for exploration: society of economic paleontologists and mineralogists special publication 31, p. 315–329. kocurek, g., and dott, r.h., 1983, jurassic paleogeography and paleoclimate of the central and southern rocky mountains regions, in reynolds, m.w., and dolly, e.d., editors., mesozoic paleogeography of the west-central united states: rocky mountain section, sepm (society for sedimentary geology), p.101–116. loope, d.b., kettler, r.m., and weber, k.a., 2010, follow the water: connecting a co2 reservoir and bleached sandstone to iron-rich concretions in the navajo sandstone of south-central utah, usa: geology, v. 38, p. 999-1002 loope, d.b., kettler, r.m., and weber, k.a., 2011, morphologic clues to the origins of iron oxide-cemented spheroids, boxworks, and pipelike concretions, navajo sandstone of south-central utah, u.s.a.: journal of geology, v. 119, p. 505-520.  loope, d.b., and kettler, r.m., 2015, the footprints of ancient co2-driven flow systems: ferrous carbonate concretions below bleached sandstone: geosphere, v. 11, p. 943-957. loope, d., kettler, r., murray, k., pederson, j., and reiners, p., 2016, sandstones and utah’s canyon country: deposition, diagenesis, exhumation, and landscape evolution, in keller, s.m., and morgan, m.l., editors., unfolding the geology of the west: geological society of america field guide 44, p. 41–71. marchetti, d.w., hynek, s.a., and cerling, t.e., 2012, gravel-capped benches above northern tributaries of the escalante river, south-central utah: geosphere, v. 8, p. 835-853. moore, j.m., 2004, blueberry fields forever: nature, v. 428, p. 711-712. weber, k.a., spanbauer, t.l., wacey, d., kilburn, m.r., loope, d.b., and kettler, r.m., 2012, biosignatures link microorganisms to iron mineralization in a paleoaquifer: geology, v. 40, no. 8, p. 747-750. yoshida, h., hasegawa, h., katsuta, n., maruyama, i., sirono, s., minami, m., asahara,y., nishimoto, s., yamaguchi, y., ichinnorov, n., and metcalfe, r., 2018, fe-oxide concretions formed by interacting carbonate and acidic waters on earth and mars: science advances, vol. 4, no. 12, eaau0872. o2 o2 o2 o2 o2 feco3 fe(iii)s organic carbon time o2 h2o h+ + fe(iii)s organic carbon feco3 + h+ o2 h2o h+ + fe(iii)s organic carbon feco3 + h+ co2 + fe(ii) n c n co2 + fe(ii) a b c d figure 11. (a) spheroidal siderite (feco3) concretion that grew in oxygen-free water (and therefore contains no iron oxide) is exposed to oxygenated groundwater; b) microbial colony starts to metabolize the siderite (gaining energy and carbon); iron-oxide rind (the waste of the microbial colony) is starting to form as the siderite recedes; c) rind further thickens as siderite recedes; d) siderite entirely dissolved, rind growth is complete, and colony dies in fully oxygenated water. in this case, no core stone (or iron-oxide pseudomorphs after siderite) remains (compare to figure 9). uga-geosite-stanczyk-sentinel-landslide.indd utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors anna m. stanczyk1, jeffrey r. moore1, brendon j. quirk1, jessica j. castleton2 1university of utah, department of geology & geophysics, frederick albert sutton building, 115 s 1460 e, room 383, salt lake city, ut 84112, anna.marie.stanczyk@gmail.com 2utah geological survey, 1594 w north temple, salt lake city, ut 84116 cover image: looking up zion canyon from mt. spry. during sentinel lake’s time, water would have stretched up the canyon farther than can be seen in this photograph. photo credit sarah meiser. . paradise from cataclysm: zion canyon’s sentinel landslide 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e geosites refl ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: stanczyk, a.m., moore, j.r., quirk, b.j., and castleton, j.j., 2019, paradise from cataclysm—zion canyon’s sentinel landslide, in milligan, m., biek, r.f., inkenbrandt, p., and nielsen, p., editors, utah geosites: utah geological association publication 48, 9 p., https://doi.org/10.31711/geosites.v1i1.65. presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president a.m. stanczyk, j.r. moore, b.j. quirk, j.j. castleton zion canyon’s sentinel landslide 3 location information gps location (nad83): 12s 325840e 4123026n suggested driving directions: park at the springdale entrance and board the zion national park shuttle bus. disembark at stop 4, court of the patriarchs. physical location description: from the shuttle stop, follow the path west to the sand bench trailhead, then proceed approximately 250 feet (approximately 75 m) to the bridge crossing the north fork virgin river to the gps coordinates given above. introduction zion canyon hosts millions of visitors each year, yet few are aware of the massive prehistoric landslide that played an important role in shaping the iconic landscape. south of the sand bench trailhead and bridge, a large hill encroaches on the canyon bottom around which the north fork virgin river fl ows (fi gure 1). north of the bridge, zion canyon’s fl at bottom stretches into the distance. th e hill is part of an enormous rock avalanche deposit known as the sentinel slide that is nearly 2 miles (3.2 km) long and more than 650 feet (200 m) thick. aft er failure, the sentinel rock avalanche dammed the north fork virgin river creating a lake (known as sentinel lake) which persisted for approximately 700 years (grater, 1945; hamilton, 1976; castleton and others, 2016). over the course of the lake’s lifetime, sediment settled at the bottom of the lake to create thick deposits of mud, clay, and sand. sediment eventually fi lled in the canyon bottom behind the landslide dam, and the lake ceased to exist. th ese sediment layers are still visible today and are responsible for the remarkably fl at fl oor of upper zion canyon (grater, 1945; hamilton, 2014; castleton and others, 2016). a rock avalanche is a variety of landslide, primarily composed of rock, with a volume greater than approximately 1 million cubic yards [106 m3] and speed greater than 16 feet per second [5 m/sec] (hungr and others, 2014)—here we will use the terms rock avalanche, landslide, and slide interchangeably to describe the event. geologic setting like many national parks in southern utah and northern arizona, zion national park (zion) represents a slice of th e grand staircase—“layer cake” geology with relatively horizontal sedimentary rock formations stacked neatly atop one another (fi gure 2) ascending northward from the grand canyon. th e formations represent a variety of depositional environments as described by biek and others (2010) and references therein. in zion canyon, the uppermost visible formation capping many of the high cliff s is the middle jurassic (approximately 173-170 million years old) temple cap formation, which consists of three members in the park: reddish-brown mudstone of the lower (sinawava) and upper (esplin point) members, separated by a cliff of sandstone of the middle (white th rone) member, collectively representing shallow-marine to coastal-dune environments (doelling and others, 2013). below the temple cap we fi nd the formation responsible for zion’s famous cliff s—the navajo sandstone (the navajo). th e navajo was deposited about 185-180 million years ago during the early jurassic period in an expansive desert similar to the modern sahara. quartz sand grains were deposited in a vast sand dune fi eld known as an erg. sand deposited on the lee sides of massive dunes created steeply dipping beds that are preserved as the inclined layers (i.e., crossbeds) that zig-zag across cliff walls. th e navajo averages 2000 feet (600 m) thick within the park. th e slope-forming kayenta formation (the kayenta) is found below the navajo. it is an early jurassic formation (195-180 million years ago) comprised of alternating layers of thin sandstones and mudstones. th e kayenta was deposited in a fl uvial fl oodplain environment where sands came to rest along river beds while muds accumulated in the adjacent fl oodplains. th e early jurassic to late triassic (approximately 205–195 million years ago) moenave figure 1: location map of zion canyon in zion national park. imagery from google earth landsat/copernicus (2016). 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 formation (the moenave) sits below the kayenta. the moenave contains sandstone, siltstone, mudstone, and limestone amassed in a complex system of rivers, floodplains, and lakes. the lowermost kayenta and the moenave did not fail in the sentinel slide as evidenced by intact outcrops of the springdale sandstone (the basal subdivision of the kayenta formation) below the rock-avalanche deposit (figures 2 and 3) (doelling and others, 2002; castleton and others, 2016). zion stands on the western edge of the colorado plateau, on a relatively stable fault block bordered to the east by the sevier fault and to the west by the hurricane fault (grater, 1945; rogers and others, 2004; biek and others, 2010). continued tectonic uplift across the hurricane fault, along with its counterpart, fluvial erosion, are responsible for the deep incision of zion’s canyons (grater, 1945). miles upstream from the sand bench trailhead, the north fork virgin river begins cutting through the navajo sandstone in the section of zion known as the narrows. flowing water exploits pre-existing joints produced during uplift, gradually widening them (rogers and others, 2004). farther downstream near the sentinel, the virgin river has cut through the entirety of the navajo and into the underlying kayenta and moenave formations. as the north fork virgin river erodes the relatively weak kayenta and moenave, it undermines the overlying navajo cliffs causing rockfall and other landslides as part of the canyon widening process. the deposit the volume of the sentinel slide deposit was originally approximately 375 million cubic yards (286 million m3) (castleton and others, 2016), which is enough to cover all of new york city’s central park with 275 feet (84 m) of debris. it has a maximum thickness of 656 feet (200 m) and stretches approximately 2 miles (3.2 km) along zion canyon (castleton and others, 2016). approximately 45% of the original deposit has been removed by the virgin river creating the steep and narrow gorge adjacent to the slide. rock-avalanche deposits are often known for their hummocky (i.e., irregular surface comprised of small mounds) appearance. the sentinel slide’s surface has significant topography, which hinders views of the entire deposit from the canyon bottom. a short hike along the sand bench trail reveals the magnitude of local relief atop the slide. the uneven surface can be attributed to the irregularity of the slide sediments themselves, which range from house-sized boulders to sand and silt, as well as the slide’s failure mechanics (paguican and others, 2014). the topography has been subdued since emplacement of the deposit due to erosion of boulders and deposition of sand (both locally derived from weathering and from aeolian [i.e., wind-driven] deposition) across the deposit’s surface. figure 2: stratigraphy of zion national park. modified from biek and others (2010). a.m. stanczyk, j.r. moore, b.j. quirk, j.j. castleton zion canyon’s sentinel landslide 5 the sentinel slide’s composition reflects the geology of the surrounding cliffs. rock-avalanche deposits commonly preserve the stratigraphy (i.e., the ordered layers of rock) of the source area (hewitt, 2009). since the navajo overlies the kayenta in the cliffsides, we expect the navajo to overly the kayenta in the deposit and indeed, this relationship is observed. note that while the upper temple cap was likely also involved in the rock avalanche, its relative proportion is small and its sandstone blocks can be difficult to distinguish from the navajo. a short walk southwest along the park road offers a view into the internal structure of the landslide deposit revealed by river incision. here the upper deposit contains highly fractured, angular blocks of predominantly navajo sandstone of differing sizes – from enormous boulders to sand. it is the sand-size particles which lend their name to the trail, sand bench, that traverses the slide. below the highly fractured, densely-compact and jumbled blocks of navajo, large slabs of maroon and white beds of the kayenta formation are visible (figure 4). these blocks of kayenta underwent significant movement but appear less fractured and disrupted than the overlying navajo. this difference may be due to the displacement experienced by the respective blocks. the navajo’s vertical cliffs extend much higher above the river than the underlying kayenta, therefore the navajo blocks fell farther and likely experienced greater disruptive energy. additionally, the interbedded, mud-rich kayenta is more deformable than the massive navajo sandstone which is brittle and tends to break into angular blocks. the failure a 2016 study by castleton and her colleagues recounts a detailed description of the sentinel slide, including a numerical landslide runout model. to model the slide, they reconstructed the pre-rock avalanche topography including the source area and the original canyon bottom. they then employed a landslide runout model to simulate the failure (figure 5). their results showed that the rock avalanche traveled at a maximum velocity of approximately 200 figure 4: nearly intact but deformed and tilted blocks of the kayenta formation near the base of the sentinel slide deposit. trees are approximately 30 feet (9 m) tall for scale. figure 3: geologic map of the sentinel slide and surrounding area. modified from castleton and others (2016) after doelling and others (2002). map units are defined in figure 2. orange star is the bridge viewpoint. blue circles are sample locations from cosmogenic exposure age dating (castleton and others, 2016). the head scarp refers to the intact, upper limit of the surface which ruptured during the landslide. contour interval is 100 feet (30 m). see figure 5 for cross section. figure 5: rock-avalanche deposit cross profile; geological units as in figure 3. see figure 3 for cross section location. modified from castleton and others (2016). 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 mph (90 m/s), crossed zion canyon in less than 20 seconds, and then spread laterally up and down the canyon. the model indicated that the failure (i.e., initial detachment to final deposit emplacement) had a total duration of just over 1 minute. the age age estimates for the sentinel slide were previously derived from radiocarbon (14c) samples from sentinel lake sediments (hamilton, 1976, 2014; doelling and others, 2002). castleton and others (2016) later directly dated the landslide deposit itself using a method known as cosmogenic isotope surface exposure dating. this method can be used to determine the age of exposure for features at the earth’s surface, such as glacial moraines and landslides (ivy-ochs, 1996; gosse and phillips, 2001; ivy-ochs and kober, 2008). cosmic and solar rays constantly bombard earth’s surface with energetic particles. these particles interact with the atomic nuclei of elements within minerals, resulting in the creation of rare secondary, or cosmogenic, isotopes within the mineral and rock. the production of certain cosmogenic isotopes in specific minerals (such as 10be in quartz) occurs at a relatively well-known rate. by measuring the concentration of these cosmogenic isotopes in a sample from the landslide’s surface, we can calculate how long a rock has been exposed to cosmogenic bombardment and therefore when the landslide occurred (ivy-ochs, 1996; gosse and phillips, 2001; ivy-ochs and kober, 2008). castleton and others (2016) sampled 12 boulders atop the sentinel slide for dating (figure 3). the resulting mean age of the sentinel slide was found to be 4800 ± 400 years old. sentinel lake the sentinel slide immediately dammed the north fork virgin river, causing the canyon behind the deposit to slowly fill with water. based on modern flow rates measured at the north fork virgin river and reconstructions of the deposit’s original topography, it likely would have taken 5 to 10 years for the water to fill the canyon and begin to overtop the rock-avalanche dam (castleton and others, 2016). as the lake overtopped the dam, water rapidly eroded the loose upper portions, but the rushing cascade did not incise entirely through the deposit. the lake outlet stabilized, establishing sentinel lake, which persisted until it filled with sediment several hundred years later (figure 6) (hamilton, 2014; castleton and others, 2016). using modern east fork virgin river sediment yield measurements as a proxy for prehistoric north fork conditions (andrews, 2000; castleton and others, 2016), sentinel lake likely occupied zion canyon for approximately 700 years from 4800 to 4100 years ago. the lake was 4.4 miles (7 km) long at maximum and covered an area of 1.1 square miles (2.85 km2). the water level would have been approximately 175 feet (53 m) above the current ground surface at the sand bench trailhead viewpoint. the thick accumulation of sediment deposited in sentinel lake partly fills what might otherwise be a steep and narrow canyon, to create a gentle and inviting landscape. just as river erosion reveals the internal stratigraphy of the rock avalanche, it has also incised into the muds and clays deposited in sentinel lake. exposures of yellow, tan, and gray layers of fine sediment can commonly be found along river cut banks (figure 7). aquatic snail shells and uninterrupted clay beds indicate that sentinel lake was likely deep and long lasting, while overlying sand deposits indicate river deposition after the lake had filled with sediment (hamilton 1976, 2014). currently the north fork virgin river lies approximately 80 feet (approximately 24 m) below the elevation of the highest lake-bed deposits, revealing the extent of erosion into what was once a higher and broader valley floor several thousand years ago (castleton and others, 2016). other vantage points the enormity of the sentinel slide is difficult to appreciate from a single viewpoint. hiking the sand bench trail up and across the deposit allows better appreciation of its immense size; from a variety of viewpoints hikers can see clues to its original width across the valley and imagine the sentinel rock wall before and during failure (figure 8). walking southwest down the park road towards the canyon junction shuttle stop provides a sense of the deposit’s length and height and reveals outcrops of its internal structure including large blocks of the kayenta formation involved in the slide (figure 4). farther upcanyon, views from the kayenta trail highlight figure 6: sentinel lake high-stand (highest water level approximately 4659 feet [1420 m]) and stable low-stand (approximately 4413 feet [1345 m]). modified from castleton and others (2016). a.m. stanczyk, j.r. moore, b.j. quirk, j.j. castleton zion canyon’s sentinel landslide 7 the flatness of the canyon floor where sentinel lake once stood (figure 9). in this same area, exploration along the banks of the virgin river and its tributaries locally reveals excellent exposures of yellow and gray lake deposits. lastly, a hike up angels landing or observation point gives views of the overall valley shape and provides the optimal opportunity to imagine what zion canyon would have looked like filled with several hundred feet of water. landslide-dammed valleys throughout zion national park while the sentinel slide is the largest landslide dam in the park, evidence exists for 10 other hypothesized landslide-dammed canyons across zion (figure 11) (hamilton, 1976, 2014; biek and others, 2010). in the northern kolob canyons section of zion, hop valley shares many similarities to zion canyon: hop valley has a rock-avalanche deposit at its mouth and a flat valley floor caused by sediment accumulation behind the slide blockage. however, unlike zion canyon with its mud and clay lake deposits, hop valley is filled with vast amounts of sand. this is likely due to differences in water influx and hop valley’s much smaller drainage basin that taps mostly sandstone bedrock. whereas zion canyon hosts the large north fork virgin river, hop valley holds a relatively small stream which only reaches the larger la verkin creek during floods. the smaller stream flow has also better preserved the rock avalanche deposit. hop valley’s rock avalanche is older than the sentinel slide (stanczyk, in prep.), yet the deposit is significantly less incised, and it will take much longer for floods in the small catchment to cut through the remaining debris. in addition to these ancient slide deposits, landslides have dammed several canyons in zion in recent history. in the 1970’s a small rockfall occurred in mystery canyon (a minor tributary near the narrows; figure 8) (hamilton, 1976, 2014). the rockfall created a natural dam which holds a seasonal pond and impounds sediment. in 1990, a larger landslide blocked middle fork taylor creek in the kolob canyons section of the park, creating a small lake which rose for 3 years (lund and others, 2010). in march 1993 water overtopped the dam, resulting in an outburst flood that reached all the way to interstate 15 where it collided with several cars (lund and others, 2010). luckily, no injuries resulted, but this recent example emphasizes the continued hazard that landslides bring to this popular national park. figure 7: a large outcrop of sentinel lake deposits near zion lodge. gray and yellow lake-bottom clay layers are capped by red sandy layers emplaced near the end of the lake’s life once it filled with sediment. figure 8: view from the sand bench trail looking down the incised rock-avalanche deposit to the north fork virgin river and the park road below. figure 9: modern zion canyon’s flat valley floor. photo taken looking down canyon from the kayenta trail. 8 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 in zion’s steep, high-relief topography, landslides have played and will continue to play an important role in eroding and widening canyons. while running water drives the consistent incision of bedrock, landslides ranging from small-scale rockfalls to enormous rock avalanches like the sentinel slide account for abrupt changes in canyon topography, sometimes with long-lasting impact. geological studies highlight the variety of landslides in zion and historical records provide much needed context to evaluate the hazard they pose (lund and others, 2010). however, further work is needed to characterize and date ancient prehistoric landslides, especially those of large magnitude. once robust catalogs and timing information are available, trends may become apparent in landslide location and/or timing, indicating important factors contributing to triggering such as climate change and ancient earthquakes. for now, recognizing the dramatic impact that the sentinel slide has had on the geomorphology (and ecology) of zion canyon highlights both the hazard and possibility for constructive transformation brought on by large-magnitude rock avalanches in this desert paradise. acknowledgements the authors are grateful to wayne hamilton and dave sharrow for motivating discussions and field visits that opened the door to our studies of the sentinel and hop valley rock avalanches. we thank tyler knudsen and greg mcdonald for careful reviews and insightful comments that improved the text. figure 11: other landslide-dammed valleys in zion national park. (1) paria pond, (2) beatty slide and pond, (3) middle fork taylor creek slide, (4) potamogeton lake, (5) hop valley slide and landslide-dammed valley, (6) currant creek (pond), (7) cane creek (pond), (8) smith creek (pond), (9) mystery canyon slide and pond, (10) sentinel slide and sentinel lake, (11) pine creek valley (pond). modified from hamilton (2014). figure 10: looking up zion canyon from mt. spry. during sentinel lake’s time, water would have stretched up the canyon farther than can be seen in this photograph. photo credit sarah meiser. a.m. stanczyk, j.r. moore, b.j. quirk, j.j. castleton zion canyon’s sentinel landslide 9 references andrews, e.d., 2000, bed material transport in the virgin river, utah: water resources research, v. 36, p. 585–596. biek, r.f., willis, g.c., hylland, m.d., and doelling, h.h., 2010, geology of zion national park, utah, in sprinkel, d.a., chidsey, jr., t.c., and anderson, p.b. editors, geology of utah’s parks and monuments: utah geological association and bryce canyon natural history association, p. 108–143. castleton, j.j., moore, j.r., aaron, j., christl, m., and ivy-ochs, s., 2016, dynamics and legacy of 4.8 ka rock avalanche that dammed zion canyon, utah, usa: geological society of america, gsa today, v. 26, p. 4–9. doelling, h.h., willis, g.c., solomon, b.j., sable, e.g., hamilton, w.l., and naylor ii, l.p., 2002, interim geologic map of the springdale west quadrangle, washington county, utah: utah geological survey open file report, p. 11. doelling, h.h., sprinkel, d.a., and kuehne, p.a., 2013, temple cap and carmel formations in the henry mountains basin, wayne and garfield counties, utah, in morris, t.h., and ressetar, r., editors, the san rafael swell and henry mountains basin—geologic centerpiece of utah: utah geological association publication 42, p. 279-318 with appendices. gosse, j.c., and phillips, f.m., 2001, terrestrial in situ cosmogenic nuclides: theory and application: quaternary science reviews, v. 20, p. 1475–1560. grater, r., 1945, landslides in zion canyon, zion national park, utah: the journal of geology, v. 53, p. 116–124. hamilton, w.l., 2014, ancient lakes of zion national park: uga publication, v. 43, p. 1–23. hamilton, w.l., 1976, holocene and pleistocene lakes in zion national park, utah, in linn, r.m. editor, first conference on scientific research in the national parks, new orleans, la: national park service, p. 835–843. hewitt, k., 2009, catastrophic rock slope failures and late quaternary developments in the nanga parbat-haramosh massif, upper indus basin, northern pakistan: quaternary science reviews, v. 28, p. 1055–1069, doi: 10.1016/j.quascirev.2008.12.019. hungr, o., leroueil, s., and picarelli, l., 2014, the varnes classification of landslide types, an update: landslides, v. 11, p. 167–194, doi: 10.1007/s10346-013-0436-y. ivy-ochs, s., 1996, the dating of rock surfaces using in situ produced 10be, 26al, and 36cl, with examples from antarctica and the swiss alps: swiss federal institute of technology, zurich, 196 p. ivy-ochs, s., and kober, f., 2008, surface exposure dating with cosmogenic nuclides: quaternary science journal (eiszeitalter und gegenwart), v. 57, p. 179–209. lund, w.r., knudsen, t.r., and sharrow, d.l., 2010, geologic hazards of the zion national park geologic-hazard study area, washington and kane counties, utah: utah geological survey special study, v. 133, 105 p. paguican, e.m.r., van wyk de vries, b., and lagmay, a.m.f., 2014, hummocks: how they form and how they evolve in rockslide-debris avalanches: landslides, v.11, p. 67-80, doi: 10.1007/s10346-012-0368-y. rogers, c.m., myers, d.a., and engelder, t., 2004, kinematic implications of joint zones and isolated joints in the navajo sandstone at zion national park, utah: evidence for cordilleran relaxation: tectonics, v. 23, p. 1–16, doi: 10.1029/2001tc001329. uga-geosite-potter-mcintyre.indd utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors sally l. potter-mcintyre southern illinois university geology department, parkinson lab mailcode 4324 carbondale, il 62901 pottermcintyre@siu.edu crystal geyser: an unusual cold spring system, grand county cover image: crystal geyser terraces of calcium carbonate and iron oxide and iron oxyhydroxide minerals. 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e geosites refl ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: potter-mcintyre, s.l., 2019, crystal geyser—an unusual cold spring system, grand county, in milligan, m., biek, r.f., inkenbrandt, p., and nielsen, p., editors, utah geosites: utah geological association publication 48, 6 p., https://doi. org/10.31711/geosites.v1i1.63. presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president s.l. potter-mcintyre crystal geyser 3 introduction crystal geyser is a cold carbon dioxide (co2) geyser, part of a natural spring system along the little grand wash fault south of green river, utah (figure 1). the spring system hosts a series of co2-driven geysers and springs with active and fossil microbial mats and tufa deposits composed of carbonate and iron oxide and iron oxyhydroxide minerals (potter-mcintyre and others, 2017; knuth and potter-mcintyre, 2018) (figure 2). additionally, progressively older carbonate spring deposits crop out on some of the topographic highs in the area because these relatively erosion-resistant deposits armor the paleo-land surface and slow down erosion (shipton and others, 2004; burnside, 2010) (figures 1 and 2). recent radiometric u-th dating of carbonate terraces and embedded veins reveal that co2-charged fluid has constantly leaked to the surface for over 400 thousand years during the pleistocene (burnside, 2010). crystal geyser is a popular place for tourists, and it is not uncommon to see children playing in the spring. the crystal geyser conduit is actually an abandoned petroleum exploration well through which water emanates. surrounding the pipe are terraces of primarily carbonate mineral deposits (shipton and others, 2004; potter-mcintyre and others, 2017; figure 2), dull to brilliant orange in color owing to minor iron precipitated from the spring water (figure 2). these terraces cascade down to the river, and include orange and green pools depending on the microbes within them—the green color indicates photosynthesizing microbes. larger terraces are composed of multiple small terracettes that are thought to be microbially-induced structures (fouke and others, 2000). also present around the drill pipe are collections of spheroidal mineral masses called pisoids. these are formed from agitation of minerals when the geyser erupts, causing spheres of precipitate to roll around and accrete new layers of carbonate minerals. directions from i-70 head south off the east green river exit 164 and then turn east. take a right at the sign for crystal geyser and follow the road. the road is a graded dirt road that is generally in good condition. if it has been raining a lot, the road may be more difficult to navigate. about halfway between the hairpin turn and crystal geyser, an oil seep is just off the north side of the road. gps location: n38.94 w110.14 where does the water come from? the artesian spring water emanates from deep subsurface reservoirs along geologic faults that bound salt wash and ten mile graben (jung and others, 2014; figure 3). the source reservoirs are jurassic and permian units that recharge at the san rafael swell to the west (baer and rigby, 1978; mcpherson and heath, 2009; dubacq and others, 2011; kampman and others, 2014). the spring water is co2and methane-charged, saline, and of neutral ph (6.2-7; shipton and others, 2004; potter-mcintyre and others, 2017). the source of co2 is likely from decarbonation of paleozoic carbonate rocks (leadville limestone) deep below the reservoir i70green river, ut crystal geyser e. green river exit n figure 1. location map for crystal geyser. the yellow dotted line is the road. the dotted red line is the little grand wash fault. note that the rocks are tan and purple on the north side of the fault (the jurassic rocks) and dark grey on the south side (the cretaceous mancos shale). see figure 3 for stratigraphy. a b c d e f g ~2cm ~10cm figure 2. crystal geyser features. a. terraces of calcium carbonate and iron oxide and iron oxyhydroxide minerals. b. close up of pools of water at top of tufa terrace. green pools contain photosynthetic organisms. c. close up of terrace showing that the terraces are composed of small terracettes thought to be formed via the interaction of microbes during mineral precipitation. d. spheroidal calcium carbonate mineral masses called pisoids. photo is about 1 inch (2 cm) across. e. drill pipe at crystal geyser is about 2 feet (1.5 m) high. f. 100,000-year-old tufa terrace atop paleo-land surface (yellow dotted line) just northeast of crystal geyser. deposit is about 10 feet (3 m) thick. g. two photos of crystal geyser erupting; left photo reproduced from https://fotospot.com/attractions/utah/ crystal-geyser; photo on the right from https://commons.wikimedia.org/w/index. php?curid=4624320. 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 (shipton and others, 2004; heath and others, 2009; kampman and others, 2009; probst and others, 2017). th e tufa deposits are spatially dispersed and of variable volumes, suggesting that the location of co2 leakage has varied over geologic time depending on the ability of the faults to transmit fl uid (permeabilty). th e subsurface strata and faults exhibit strongly heterogeneous permeability owing to seismic activity, regional erosion, both mineral dissolution and precipitation, and changing fl uid fl ow volumes owing to variability in climate over time (burnside, 2010; burnside and others, 2013; kampman and others, 2014). how does crystal geyser work? as mentioned in the introduction, crystal geyser formed via a human-drilled wellbore. an oil seep along the little grand wash fault motivated drilling of the oil exploration well in 1935 (baer and rigby, 1978). th e well never produced oil; however, co2 dissolved and pressurized in the artesian aquifer at depth now discharges through the wellbore. th is open conduit allows rapid depressurization and discharge as episodic geyser eruptions. following each eruption, the wellbore again fi lls with water from the bottom up, with pressure building up during fi lling. th e artesian pressure ultimately exceeds the rate of refi ll and causes another eruption (watson and others, 2014). th e cycle repeats, sometimes aft er a few hours and sometimes as long as a day or more between eruptions. crystal geyser’s eruption intervals, durations, and intensities were at one time regular and consistent, but timing now is quite erratic, possibly related to vandalism. tourists have dropped rocks and even reportedly dynamite into the geyser (shipton and others, 2004). other possible factors for variable eruption rates include seismic activity (han and others, 2013) and/or interactions between recharge rates and co2 migration rates within the artesian aquifers (kampman and others, 2014). site of active scientific research crystal geyser and its related spring system are subjects of active scientifi c research on topics ranging from global warming to the search for extraterrestrial life. th is section discusses topics of recent research, including analysis of the geyser’s source aquifer as an analog to engineered carbon capture and sequestration, followed by analysis of microbial life, interactions between microbes and mineral precipitation and how these processes off er insight regarding the search for life on mars and beyond. co2 sequestration global warming of our planet is attributable to the greenhouse eff ect, specifi cally increasing concentration of anthropogenic co2 in the atmosphere that traps heat from solar radiation aft er it is refl ected from the earth’s surface (e.g., scheff er and others, 2006; eby and others, 2009; notz and stroeve, 2016; specht and others, 2016). many ideas have been proposed to reduce co2 emissions and the greenhouse eff ect, one of which is carbon capture and sequestration (e.g., yang and others, 2008; dai and others, 2013; rahman and others, 2017; rackley, 2017). carbon capture and sequestration (ccs) includes capture of co2 at point sources such as cement plants and power plants, pressurizing and condensing it to a fl uid and then injecting that fl uid into subsurface reservoirs. th e fl uid that emanates at crystal geyser comes from a natural subsurface co2 reservoir, but it leaks via migration upward along faults. th e spring water degasses its co2 at the springs and geysers and eff ectively emits co2 into the atmosphere, similar to industrial sources (but much smaller in volume). understanding how this gas moves upward, how the emissions vary from site to site along faults, and what impedes or promotes fl ow are all very important parameters to know before ccs becomes a viable mitigation strategy for anthropogenic co2 emissions (e.g., shipton and others, 2005; gouveia and friedmann, 2006; burnside and others, 2013; watson and others, 2014). astrobiology biosignatures are preserved fi ngerprints of past microbial life, which is the type of life scientists are searching for on mars and icy moons within our solar system. th ree types of biosignatures summerville fm entrada ss navajo ss kayenta fm wingate ss curtis fm carmel fm mancos sh dakota ss cedar mtn fm morrison fm fa ul t c on du it fo r � ui ds figure 3. stratigraphy at crystal geyser. th e photo is looking north-northwest (upriver). to the right, the exposed rocks are the middle jurassic section. as one drives along the little grand wash fault (see fi gure 1), the grey rocks to the south of the road are the cretaceous mancos shale. th ese rocks are younger than the exposed rocks to the north of the fault and were downthrown relative to the jurassic rocks. th e well is in green and it extends 2627 feet below the surface. it is not cased, so the co2-charged water fl ows into the pipe in both the entrada sandstone and the navajo sandstone reservoirs (watson and others, 2014). however, this fault serves as a conduit for fl uid to fl ow upward to the surface and come out at crystal geyser, and for the oil seep you passed on the way in. jurassic rocks are in yellow and the cretaceous rocks in green. s.l. potter-mcintyre crystal geyser 5 icy moons enceladus and europa are high priority targets for future exploration because of their subsurface oceans, which make them potentially habitable environments (hendrix and others, 2019). th ese moons exhibit plumes (geysers) of subsurface water that erupts to the surface. th ese plumes would make excellent targets for understanding the habitability of enceladus and europa because of their relative ease of accessibility. studies of the microbial life deep within the crystal geyser waters have found a diverse population with adaptations to reside in co2-rich, saline environments (santillan and others, 2015; emerson and others, 2016; probst and others, 2017; knuth and potter-mcintyre, 2019; fi gure 4). ongoing studies seek to fi nd ways to determine habitability from the geyser plumes to help design future missions. summary crystal geyser is a fascinating example of a rare cold spring and geyser system. it is a treasure trove of scientifi c information, as well as just a fun and scenic place to visit. spend some time hiking around and looking at the fault and the older tufa deposits, and think about how these formed throughout the millennia—and think about similar features on mars and other celestial bodies in our solar system! acknowledgements i would like to thank collaborators on the biosignature research (ms. jordan knuth) and on the icy moons research (drs. morgan cable, kate craft , michael malaska, amanda stockton, and alex patthoff ). i gratefully acknowledge dr. brian mcpherson for his review of this manuscript. th is research has been funded in part by american chemical society petroleum research fund (to potter-mcintyre). references baer, j.l., and rigby, j.k., 1978, geology of the crystal geyser and environmental implications of its effl uent, grand county, utah: utah geology, v. 5, no. 2, p. 125-130. burnside, n.m., 2010, u-th dating of travertines on the colorado plateau: implications for the leakage of geologically stored co2: scotland, university of glasgow, ph.d. dissertation, 290 p. burnside, n.m., shipton, z.k., dockrill, b., and ellam, r.m., 2013, man-made versus natural co2 leakage: a 400 ky history of an analogue for engineered geological storage of co2: geology, v. 41, no. 4, p. 471-474. cady, s.l., farmer, j.d., grotzinger, j.p., schopf, j.w., and steele, a., 2003, morphological biosignatures and the search for life on mars: astrobiology, v. 3, no. 2, p. 351-368. corkeron, m., webb, g.e., moulds, j., and grey, k., 2012, discriminating stromatolite formation modes using rare earth are described as follows: (1) carbonaceous body fossils of microbes, (2) microbially infl uenced sedimentary structures such as microbialites (laminated mineral deposits formed via microbial mats), and chemical fossils (such as organic molecules or minerals directly precipitated via the metabolisms of organisms like shells), and (3) isotopic signatures or concentrations of trace elements specifi c to sequestration by microbes (cady and others, 2003; westall, 2008; potter-mcintyre and others, 2014). crystal geyser and nearby springs all host microbial life; the water is too salty for anything else to grow in it (fi gure 4). th is section examines some ongoing research using crystal geyser as an analog to mars and then to icy moons, such as enceladus and europa. mars on earth, microorganisms commonly enhance mineral precipitation and mediate mineralogical and chemical compositions of resulting deposits (e.g., reid and others, 2000; dupraz and others, 2009; petryshyn and others, 2012; corkeron and others, 2012). many of the features at crystal geyser are thought to be created by microbes, such as the terracettes and the green color of some of the pools and even the orange color (emerson and others, 2016; potter-mcintyre and others, 2017; fi gure 2). even though some research seems to suggest abiotic precipitation plays a large part in carbonate formation at springs due to degassing of co2 (e.g., fouke and others, 2000; takashima and others, 2011; knuth and potter-mcintyre, 2019), those studies acknowledge that microbial metabolisms do aff ect precipitation, particularly in minerals forming away from the vents (fouke and others, 2000; takashima, 2011). a host of micro-organisms and mineral habits that are likely microbially induced are present in the tufa deposits (knuth and potter-mcintyre, 2019; fi gure 4). 30μm 40μm 30μm a ~5cm b c d figure 4. life in crystal geyser. a. a snake did not fare well wandering into the water that is too salty for most organisms. b. some organisms, called halophiles, love salty water. yellow arrows point to diatoms from a microbial mat at one of the springs near crystal geyser. red arrow points to exopolymeric substance that is produced by microbial life and provides a substrate onto which calcium carbonate minerals can easily precipitate. c and d. yellow arrows point to unusual mineral forms likely produced by the interaction with microbes. images b, c, and d scales are in microns (µm). a human hair is approximately 75 µm. 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from crystal geyser, ut: livermore, california, lawrence livermore national lab, no. ucrltr-221731. han, w.s., lu, m., mcpherson, b.j., keating, e., moore, j., park, e., watson, z., and jung, n.h., 2013, characteristics of co2-driven cold-water geyser, crystal geyser in utah: experimental observation and mechanism analyses: geofluids, v. 13, no. 3, p. 283-297. heath, j.e., lachmar, t.e., evans, j.p., kolesar, p.t., and williams, a.p., 2009, hydrogeochemical characterization of leaking, carbon dioxide-charged fault zones in east-central utah, with implications for geologic carbon storage, in mcpherson, b.j. and sundquist, e.t., editors, carbon sequestration and its role in global carbon cycle: washington, d.c., american geophysical union, geolphysical monograph 183, p. 147-58. hendrix, a.r., hurford, t.a., barge, l.m., bland, m.t., bowman, j.s., brinckerhoff, w., buratti, b.j., cable, m.l., castillo-rogez, j., and collins, g.c., 2019, the nasa roadmap to ocean worlds: astrobiology, v. 19, no. 1, p. 1-27. jung, n.-h., han, w.s., watson, z., graham, j.p., and kim, k.-y., 2014, fault-controlled co2 leakage from natural reservoirs in the colorado plateau, east-central utah: earth and planetary science letters, v. 403, p. 358-367. kampman, n., bickle, m., becker, j., assayag, n., and chapman, h., 2009, feldspar dissolution kinetics and gibbs free energy dependence in a co2-enriched groundwater system, green river, utah: earth and planetary science letters, v. 284, no. 3-4, p. 473-488. kampman, n., bickle, m., maskell, a., chapman, h., evans, j., purser, g., zhou, z., schaller, m., gattacceca, j. c., and bertier, p., 2014, drilling and sampling a natural co2 reservoir: implications for fluid flow and co2-fluid–rock reactions during co2 migration through the overburden: chemical geology, v. 369, p. 51-82. knuth, j.m. and potter-mcintyre, s.l., 2019, chemical, morphological, and mineralogical biosignature preservation in a cold spring system, ut, usa: insights for sample selection on mars: astrobiology, in review. mcpherson, b.j., and heath, j., 2009, self-sealing of faults by co2rich fluids: geochimica et cosmochimica acta supplement, v. 73, p. a861. notz, d., and stroeve, j., 2016, observed arctic sea-ice loss directly follows anthropogenic co2 emission: science, v. 354, no. 6313, p. 747-750. petryshyn, v.a., corsetti, f.a., berelson, w.m., beaumont, w., and lund, s.p., 2012, stromatolite lamination frequency, walker lake, nevada: implications for stromatolites as biosignatures: geology, v. 40, no. 6, p. 499-502. potter-mcintyre, s.l., chan, m.a., and mcpherson, b.j., 2014, textural and mineralogical characteristics of microbial fossils associated with modern and ancient iron (oxyhydr)oxides: terrestrial analogue for sediments in gale crater: astrobiology, v. 14, no. 1, p. 1-14. potter-mcintyre, s.l., williams, j., phillips-lander, c., and o’connell, l., 2017, taphonomy of microbial biosignatures in spring deposits: a comparison of modern, quaternary, and jurassic examples: astrobiology, v. 17, no. 3, p. 216-230. probst, a.j., castelle, c.j., singh, a., brown, c.t., anantharaman, k., sharon, i., hug, l.a., burstein, d., emerson, j. b., and thomas, b.c., 2017, genomic resolution of a cold subsurface aquifer community provides metabolic insights for novel microbes adapted to high co2 concentrations: environmental microbiology, v. 19, no. 2, p. 459-474. s.l. potter-mcintyre crystal geyser 7 rackley, s.a., 2017, carbon capture and storage (second edition): oxford, uk, butterworth-heinemann, 698 p. rahman, f.a., aziz, m.m.a., saidur, r., bakar, w.a.w.a., hainin, m., putrajaya, r., and hassan, n.a., 2017, pollution to solution: capture and sequestration of carbon dioxide (co2) and its utilization as a renewable energy source for a sustainable future: renewable and sustainable energy reviews, v. 71, p. 112-126. reid, r.p., visscher, p.t., decho, a.w., stolz, j.f., bebout, b., dupraz, c., macintyre, i., paerl, h., pinckney, j., and prufert-bebout, l., 2000, the role of microbes in accretion, lamination and early lithification of modern marine stromatolites: nature, v. 406, no. 6799, 989 p. santillan, e.f.u., shanahan, t.m., omelon, c.r., major, j.r., and bennett, p.c., 2015, isolation and characterization of a co2-tolerant lactobacillus strain from crystal geyser, utah, usa: frontiers in earth science, v. 3, p. 41. scheffer, m., brovkin, v., and cox, p.m., 2006, positive feedback between global warming and atmospheric co2 concentration inferred from past climate change: geophysical research letters, v. 33, no. 10, l10702 p. shipton, z.k., evans, j.p., kirschner, d., kolesar, p.t., williams, a.p., and heath, j., 2004, analysis of co2 leakage through ‘low-permeability’ faults from natural reservoirs in the colorado plateau, east-central utah: geological society, london, special publications, v. 233, no. 1, p. 43-58. shipton, z.k., evans, j.p., dockrill, b., heath, j., williams, a., kirchner, d., kolesar, p.t., thomas, d., and benson, s., 2005, natural leaking co2-charged systems as analogs for failed geologic storage reservoirs, in thomas, d.c. and benson, s.m., editors, carbon dioxide capture for storage in deep geologic formations—results from the co2 capture project 2: new york, elsevier, p. 699-712. specht, e., redemann, t., and lorenz, n., 2016, simplified mathematical model for calculating global warming through anthropogenic co2: international journal of thermal sciences, v. 102, p. 1-8. takashima, c., okumura, t., nishida, s., shimamoto, t., koike, h., and kano, a., 2011, microbial control on lamina formation in a travertine of crystal geyser, utah, in reitner, j., quéric, n.-v., and arp, g., editors, advances in stromatolite geobiology: berlin, heidelberg, springer, p. 123-133. watson, z., han, w.s., keating, e.h., jung, n.-h., and lu, m., 2014, eruption dynamics of co2-driven cold-water geysers: crystal, tenmile geysers in utah and chimayó geyser in new mexico: earth and planetary science letters, v. 408, p. 272-284. westall, f., 2008, morphological biosignatures in early terrestrial and extraterrestrial materials: space science reviews, v. 135, no. 1-4, p. 95-114. yang, h., xu, z., fan, m., gupta, r., slimane, r.b., bland, a.e., and wright, i., 2008, progress in carbon dioxide separation and capture: a review: journal of environmental sciences, v. 20, no. 1, p. 14-27. uga-geosite-anderson.indd utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors paul b. anderson consulting geologist, p.o. box 101, emery, utah 84522 paul@pbageo.com t o p b o tto m figure 9. unidentified plant fossil found in a boulder of the kf-6 beach sand. top to the right and bottom on the left. trekking pole for scale (3.9 feet (1.2 m )). dashed line approximates its outline. the overall length is about 12 feet (3.7 m) with the width at the upper crown at about 3.5 feet (1.1 m). location shown in figure 1. cover image: unidentifi ed plant fossil found in the kf-6 beach sand. top to the right and bottom on the left . trekking pole for scale. dashed line approximates its outline. what you can discover on an ancient cretaceous beach, a geosite in emery county, utah 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 p.b. anderson ancient cretaceous beach 3 introduction utah was prime beach country in its central to eastern portion during the middle to late cretaceous (70 to 92 million years ago). at this time, a long shallow sea extended from the artic to the gulf of mexico and from central utah east to beyond kansas. scores of ancient beach deposits that represent the shoreline along this seaway are exposed in the eastern half of utah. this paper will guide you to one of these white sandy beaches that dominated central utah during this time period. the site is located near a developed archeological site (rochester panel, smithsonian site number 42em392). the rochester panel is dissimilar to other fremont culture images and likely dates between a.d. 500 to a.d. 700 (loendorf, 1985). if you visit the panel please do not touch (oil from you skin damages the surface), deface, or climb on this precious resource. driving directions this geosite is conveniently located near an established and often visited archeological site named the rochester rock art panel. the panel has a developed parking area which will be your driving destination for this geosite. to arrive at the parking area from the wasatch front, travel south on interstate 15 to the junction of us highway 6, travel southeast to the town of price, then exit highway 6 onto state road 10 southbound toward the town of ferron. continue about 9 miles (14.5 km) past the town of ferron to the second marked left turn-off for the town of moore (county road 801). after immediately crossing a cattle guard, proceed east 0.5 miles (0.8 km) to a right turn onto a graveled road marked with a sign, “rochester panel” (county road 805). proceed south on this improved gravel road, past gravel quarries on the left (east) and at 1.8 miles (2.9 km) from the pavement, proceed past a radio tower on the left. continue south another 2.1 miles (3.4 km), following the graveled road to a dead end in a small loop with a parking area and blm sign (figure 1). the sign provides information about the archeological site. park your vehicle here. however, from this point (with the gps coordinates below) a well-marked footpath (about 1/2 miles, (0.8 km) one way) will take you to one of the state’s best fremont indian petroglyph panels. the hike is an easy walk with about 100 vertical feet (30 m) decent from the parking area to the panel. let me set the geologic picture of the surroundings before guiding you to several interesting geologic viewpoints. gps coordinates for the parking area are 38°54'25"n, 111°11'43"w; (utm 483083me, 4306444mn). stratigraphy using the parking area as our point of reference, you are parked at the approximate top of the ferron sandstone member of the mancos shale (figure 2). stratigraphically above this point is the blue gate shale member of the mancos shale, a gray-blue mudstone/shale that is visible to the northwest and continues up to the base of the tan cliff-forming unit on the wasatch plateau, in the distance. this first cliff-forming sandstone in the cliff face is the star point sandstone (hintze and kowalis, 2009, section 82). back to the parking and below the ferron sandstone is the tununk shale member of the mancos shale. look to the southeast at the tree-covered rise which follows the top of the ferron sandstone to a spectacular vertical cliff called the coal cliffs. below the cliffs are excellent exposures of the tununk shale member of the mancos shale (not visible from here). parking area m u d d y c re e k rocheste r w ash roschester panel on kf-5 shoreface kf-6 beach pinchout kf-6 beach pinchout kf-6 beach pinchout trail head to archeological site trail price castle dale green riverferron geositeemery e m er y c ou nt yhuntington a b c d 0 250 500 ft. a a’ s e c ti o n l in e s e e fi g u re 3 figure 1. index map and study area image. yellow solid line is the top of the beach sand. yellow dashed line is the approximate landward edge of the beach. a= access point to cliff base; b=viewing point for fossil plant; c=dinosaur tracks; d=point of departure from main trail. parking area m u d d y c re e k rocheste r w ash roschester panel on kf-5 shoreface kf-6 beach pinchout kf-6 beach pinchout kf-6 beach pinchout trail head to archeological site trail price castle dale green riverferron geositeemery e m er y c ou nt yhuntington a b c d 0 250 500 ft. a a’ s e c ti o n l in e s e e fi g u re 3 p.b. anderson ancient cretaceous beach 4 the ferron sandstone, below your feet at the parking area, is about 500 feet (152 m) thick (thompson, 1986) and consists of sandstone, siltstone to mudstone, and coal and almost coal (dark to black carbonaceous shale). figure 2 is a stratigraphic column of the ferron sandstone found in this locale based on nomenclature from anderson and ryer (2004). only the top portion of the ferron sandstone stratigraphy shown in figure 2 is exposed at this geosite. the sand we will focus on is labeled on figure 2 as kf-6 and outlined in red. this further subdivision of the ferron shown in figure 2 enables clear communication when discussing the details of the stratigraphy. since a large portion of the world’s population is located along the seashore, rising or falling sea level is closely monitored today. the ferron sandstone provides us a partial record of the many changes in sea level during the late cretaceous. “the sea went in and the sea went out,” has been used to describe the geologic history of the cretaceous in utah. the stratigraphic column at this site (figure 2) records five of these major shoreline fluctuations, representing about 2 million years of earth’s history. minor shoreline shifts are represented by the various sub-units (parasequences) of the kf-2 parasequence set. the stratal package names are shown to the right of the rock column in figure 2. the naming scheme is kf for cretaceous (k is commonly used to denote cretaceous because another period already claims the letter “c”), then a dash, and then a number or name. the numbered units begin at the bottom of the ferron and are sequentially numbered in ascending order. ryer (1981) originally described and numbered these units as “delta-front units” which formed as sediment deposited at the shoreline by rivers flowing generally from southwest to northeast. the unit we will focus on is kf-6, or one of the shorelines that lies near the top of the strata stack. structure this geosite lies on the west flank of a large elongated dome called the san rafael swell. looking to the southeast from the parking lot, you will notice the tree-covered slope rising at about five degrees. the ferron sandstone has been tilted in this direction by the rise of this large geologic structure which began to move slightly, soon after deposition of the ferron. the more rapid rise of the structure began about 65 million years ago (hintze, 2005), so the rocks on the san rafael swell have been eroding to look like our present landscape for millions of years. this present tilt or dip of the beds of the ferron sandstone plays an important part our story of the present landward disappearance of the kf-6 beach sand. you will note at the parking lot that the landscape in the opposite direction of the southeast ridge is very different. not only is there lack of trees (pinyon and juniper trees need sandy soils), but this sharp vegetative boundary marks the top of the ferron sandstone member and the base of the blue gate shale member. the softer shales of the blue gate erode into a wide low land or “valley,” hence the name castle valley for this portion of utah. geologic history setting the scene an important part of the story of the ferron sandstone “wedge” is the more regional history of the large continental plates of crustal material that “float” on the top of the earth and have dramatic movements when viewed on the scale of millions of years. along the western edge of the north american plate, the adjacent pacific crustal plate was growing from a spreading center in the mid-pacific but because of its growth along the spreading ridge the expanding plate was forced to dive or subduct along the bordering north american plate. we call this collision zone a subduction zone. the force of the collision caused large sheets of the north american continental crust to be pushed or thrusted over one another, like a pile of playing cards being pushed together. the result was a “pile” of over-thrusted crustal pieces in western utah and eastern nevada. the great weight of this pile of crustal sheets miller creek-b kf-7 kf-6 j-coal i-coal g-coal a-c-coal kf-5 kf-4 miller creek-a muddy creek-b muddy creek-a blue gate shale member k f-2 tununk shale member dakota formation m a n c o s s h a le f e rro n s a n d s to n e m e m b e r stratigraphy figure 2. stratigraphy of the rochester geosite. yellow hues are shoreline or beach-related sand bodies. orange are river-deposited, and brown and black are deposited in swamps. gray is offshore marine shale. red bracket – the focus of this geosite (see figure 6 for a depositional history of this interval). p.b. anderson ancient cretaceous beach 5 (called the sevier orogenic belt), loaded the north american plate and caused it to sink, sag, or subside. since the crust is semi-rigid, this sinking deformed the crust out in front of all this loading, causing a low or downwarp to form. geologists call this downwarp a foreland basin because it lies in front of or before the area of crustal loading. the subsiding foreland basin eventually dropped below sea level through a good portion of the late cretaceous age, forming the western interior seaway (figure 3) where sediments from the eroding upland area to the west began to accumulate. now we have a mechanism to explain why the ferron sandstone consists of a vertical stack of sandstone and other rocks. as each shoreline or delta-front fills in the shallow sea now occupying the foreland basin, the combination of loading from the deposited sediments and more thrusting and loading to west, causes the foreland basin to continue to subside, setting the stage for a shoreline to move back over the prior deposited sediments and add additional sandy shorelines and lower coastal-plain sediments into this continuously subsiding basin. these crustal fluctuations are considered “local tectonics” or plate motion, but world-wide sea level changes were also taking place. these sea level changes must be added to the possible causes of shoreline fluctuations (rise and fall). as a means of tracking these shoreline movements we often refer to movements of the shoreline towards the seaway or the basin as “seaward” and those in the opposite direction as “landward” since the shoreline is moving towards the land. the kf-6 beach (shoreline) edge or pinchout proceed west to the edge of muddy canyon to a point labeled “a” in figure 1 (utm nad83 482823me, 4306322mn). from here you will have a commanding view of the west side of the canyon where the kf-6 beach pinchout is superbly exposed. figure 4 is a diagrammatic cross section drawn on a line just west of the exposed outcrop and shown on figure 1. the view of the west side of muddy canyon from your vantage point should be similar to the cross section, figure 4. note the vertical exaggeration in the cross section distorts the true scale relationship, which you see in the canyon wall. the base of the cross section (figure 4) and the drainage bottom in the view to the west before you are approximately equal. the prominent gully cutting the cross section lies almost directly across the canyon. the cliff-forming sandstone, lying on the north side of the gully depicted in the cross section is the kf-6 beach or shoreline sandstone. please refer to figure 5a and b, which is a series of photographs taken from the east side of muddy canyon, and relate well to the cross section in figure 4. in figure 5a, the prominence on the north side of the small gully across muddy canyon, is located at the end of the label “i coal zone.” this is our “subject sandstone.” follow this sandstone to the left in the photo and on the outcrop, across (south) the small gully, to another prominence where it is now noticeably thinner than at our beginning point. continue to trace the cliff-forming sand to the south north a’ a kf-5 stream deposits (mostly sandstone) finer grained kf-5 stream deposits kf-5 shoreline deposits (mostly sandstone) kf-7 marine coal prospect j coal ? i coal zone kf-7 marine kf-6 beach blue gate shale/slope wash figure 4. diagrammatic cross section of the east side of muddy canyon, showing the pinchout of the beach sandstone of kf-6 subunit (darker yellow) of the ferron sandstone. line of cross section shown on figure 1. w e s te rn in te rio r s e a w a y s e v ie r o ro g e n ic b e lt figure 3. paleogeographic map of north america (modified from blakey, 2006). red dot is our geosite. p.b. anderson ancient cretaceous beach 6 landward pinchout kf-6 beach kf-6 beach kf-6 beach kf-5 stream deposits i coal zone i coal zone kf-5 shoreline kf-4 stream deposits kf-7 marine deposits kf-5 stream deposits j-coal blue gate shale figure 5. three photographs of the edge or landward pinchout of the beach sand of the kf-6 beach. a) view west across muddy canyon. partial view of the cross section in figure 3. b) close up view of part of the field of view in (a). note non-parallel bedding of the stream-laid upper portion of unit kf-5 (orange dots) and the i coal zone (blue dots). c) view southeast from (a) showing the southwest (landward) thinning of the kf-6 beach in multiple exposures along the dipping top of the ferron sandstone. the uppermost thinning does culminate at the pinchout. a b c figure 5. three photographs of the edge or landward pinchout of the beach sand of the kf-6 beach. a) view west across muddy canyon. partial view of the cross section in figure 3. b) close up view of part of the field of view in (a). note non-parallel bedding of the stream-laid upper portion of unit kf-5 (orange dots) and the i coal zone (blue dots). c) view southeast from (a) showing the southwest (landward) thinning of the kf-6 beach in multiple exposures along the dipping top of the ferron sandstone. the uppermost thinning does culminate at the pinchout. p.b. anderson ancient cretaceous beach 7 south (left) to a point where the sandstone thins and becomes obscured by overlying kf-7 sandstone boulders, which have separated from the outcrop. figure 5b is a close up of the final wedge-shaped landward edge or landward pinchout of this beach (sandstone). figure 5c is taken from across the canyon from your present position with the camera pointed to the southeast. you can turn from your present position on the outcrop and see a similar view. the photograph (figure 5c) is taken parallel to the shore of this approximately 89 million year old beach. several distinctly thinning exposures of the kf-6 sandstone cliff are visible, with the most distant (the top wedge in the photograph) also recording the un-eroded most landward position of the beach. at first glance this may seem spectacularly unimpressive, but let’s connect this to a more complete picture of what this landward pinchout and several others of the same sandstone body tell us about the paleogeography and shorelines of the ferron. paleogeography paleogeographic maps are means to express an opinion of what the landscape looked like during a specific time in the past. figure 6 is a series of paleogeographic maps that represent the time just before, during, and after the deposition of the kf-6 beach sand. figure 6a depicts the time just before the kf-6 time. the maps are drawn based on the type of rock found below the kf-6 beach sandstone, which consists of both rocks deposited by a river system, which flowed from southwest to northeast, and a partly contemporaneous/later “i” swamp. figure 5a shows these rocks and their relationship to those above and below. the swamp that formed the “i” coal zone was relatively close to sea. this swamp was the site of new sea level rise, likely due to new downwarping in this part of the foreland basin, which caused the shoreline of the western cretaceous seaway (figure 3) to move towards the land as shown in figure 6b. at the point in time when the sea level rise stopped and the crashing waves of the sea began to establish a new shoreline, a new beach was established. with a relatively stable sea level and the continuous addition of new sand at the shoreline, the new space created by a rising sea began the slow process of filling, leaving behind the sand of the kf-6 beach at this location (figure 6c). with each small (hundreds to thousands of years) increment of time (shown on figure 6c as time 1-3) more sediment was added, and, with no rise in sea level, the shoreline was forced to move seaward or toward the northeast. figure 6. maps depicting the geologic history or paleogeography (old geography) of the geosite. the stratigraphic interval depicted is shown by a red bracket in figure 2. the parking area and rock art site are labeled on (a) and shown as points of reference on all maps. a) paleogeography of the upper portion of unit kf-5. b) relative sea level rises and floods over kf-5. the first sand grains of the beach of kf-6 are deposited at the “new shoreline beach” (see figures 1 and 4). c) as sand is deposited along the shoreline through time (1-3) the shoreline slowly moves seaward (progrades). landward of the shoreline are time-equivalent deposits of near-shore bays, lagoons, and low-lands. d) as the shoreline moves seaward through time, a swamp (future coal) develops at the shoreline and follows the retreating shoreline seaward. evidence for this history is the j coal lying atop kf-6. rock art parking near-shore coastal zone (bay, lagoon,low-lands) sw am p/ flo od -p la in sw am p/ flo od -p la inseaward flowing rivers seaward flowing river (abandoned?) ? ? ?? ? ? relative rise of the sea flooding the prior land and their deposits new shoreline beach tim e 1 tim e 2 swamp (of the j coal) later tim e kf-6 shallow m arine a b c d scale 0 1 mile underlying kf-6 shoreline deposit figure 6. maps depicting the geologic history or paleogeography (old geography) of the geosite. the stratigraphic interval depicted is shown by a red bracket in figure 2. the parking area and rock art site are labeled on (a) and shown as points of reference on all maps. a) paleogeography of the upper portion of unit kf-5. b) relative sea level rises and floods over kf-5. the first sand grains of the beach of kf-6 are deposited at the “new shoreline beach” (see figures 1 and 4). c) as sand is deposited along the shoreline through time (1-3) the shoreline slowly moves seaward (progrades). landward of the shoreline are timeequivalent deposits of near-shore bays, lagoons, and low-lands. d) as the shoreline moves seaward through time, a swamp (future coal) develops at the shoreline and follows the retreating shoreline seaward. evidence for this history is the j coal lying atop kf-6. shallow sea of kf-6 tim e tim e 3 seaward-m oving shoreline/beach rock art parking near-shore coastal zone (bay, lagoon,low-lands) sw am p/ flo od -p la in sw am p/ flo od -p la inseaward flowing rivers seaward flowing river (abandoned?) ? ? ?? ? ? relative rise of the sea flooding the prior land and their deposits new shoreline beach tim e 1 tim e 2 swamp (of the j coal) later tim e kf-6 shallow m arine a b c d scale 0 1 mile underlying kf-6 shoreline deposit figure 6. maps depicting the geologic history or paleogeography (old geography) of the geosite. the stratigraphic interval depicted is shown by a red bracket in figure 2. the parking area and rock art site are labeled on (a) and shown as points of reference on all maps. a) paleogeography of the upper portion of unit kf-5. b) relative sea level rises and floods over kf-5. the first sand grains of the beach of kf-6 are deposited at the “new shoreline beach” (see figures 1 and 4). c) as sand is deposited along the shoreline through time (1-3) the shoreline slowly moves seaward (progrades). landward of the shoreline are timeequivalent deposits of near-shore bays, lagoons, and low-lands. d) as the shoreline moves seaward through time, a swamp (future coal) develops at the shoreline and follows the retreating shoreline seaward. evidence for this history is the j coal lying atop kf-6. shallow sea of kf-6 tim e tim e 3 seaward-m oving shoreline/beach rock art parking near-shore coastal zone (bay, lagoon,low-lands) sw am p/ flo od -p la in sw am p/ flo od -p la inseaward flowing rivers seaward flowing river (abandoned?) ? ? ?? ? ? relative rise of the sea flooding the prior land and their deposits new shoreline beach tim e 1 tim e 2 swamp (of the j coal) later tim e kf-6 shallow m arine a b c d scale 0 1 mile underlying kf-6 shoreline deposit figure 6. maps depicting the geologic history or paleogeography (old geography) of the geosite. the stratigraphic interval depicted is shown by a red bracket in figure 2. the parking area and rock art site are labeled on (a) and shown as points of reference on all maps. a) paleogeography of the upper portion of unit kf-5. b) relative sea level rises and floods over kf-5. the first sand grains of the beach of kf-6 are deposited at the “new shoreline beach” (see figures 1 and 4). c) as sand is deposited along the shoreline through time (1-3) the shoreline slowly moves seaward (progrades). landward of the shoreline are timeequivalent deposits of near-shore bays, lagoons, and low-lands. d) as the shoreline moves seaward through time, a swamp (future coal) develops at the shoreline and follows the retreating shoreline seaward. evidence for this history is the j coal lying atop kf-6. shallow sea of kf-6 tim e tim e 3 seaward-m oving shoreline/beach rock art parking near-shore coastal zone (bay, lagoon,low-lands) sw am p/ flo od -p la in sw am p/ flo od -p la inseaward flowing rivers seaward flowing river (abandoned?) ? ? ?? ? ? relative rise of the sea flooding the prior land and their deposits new shoreline beach tim e 1 tim e 2 swamp (of the j coal) later tim e kf-6 shallow m arine a b c d scale 0 1 mile underlying kf-6 shoreline deposit figure 6. maps depicting the geologic history or paleogeography (old geography) of the geosite. the stratigraphic interval depicted is shown by a red bracket in figure 2. the parking area and rock art site are labeled on (a) and shown as points of reference on all maps. a) paleogeography of the upper portion of unit kf-5. b) relative sea level rises and floods over kf-5. the first sand grains of the beach of kf-6 are deposited at the “new shoreline beach” (see figures 1 and 4). c) as sand is deposited along the shoreline through time (1-3) the shoreline slowly moves seaward (progrades). landward of the shoreline are timeequivalent deposits of near-shore bays, lagoons, and low-lands. d) as the shoreline moves seaward through time, a swamp (future coal) develops at the shoreline and follows the retreating shoreline seaward. evidence for this history is the j coal lying atop kf-6. shallow sea of kf-6 tim e tim e 3 seaward-m oving shoreline/beach p.b. anderson ancient cretaceous beach 8 the climate was subtropical during ferron-time and the carbon dioxide levels were estimated to be about 2000 ppm (2000 mg/kg), compared to our present 400 ppm (400 mg/kg), a good time to be a plant. not surprising, low coastal areas were the site of marshy peat swamps and that is precisely what grew on the newly created land at the seashore as the shoreline or beach retreated to the northeast. figure 6d depicts this paleogeography with the black diagonal line representing the edge or first deposits of the kf-6 shoreline and its movement seaward to the bold yellow line. in figure 5a a few black spots on the outcrop represent small exposures of the j coal, deposited in this swamp. figure 7 provides additional evidence of the swamp showing rooting of the j peat into the top of the kf-6 beach sandstone. if the top of the kf-6 sandstone is followed on outcrop to the north about one-half mile (0.8 km), the abandoned williams mine will be encountered. in the late 19th century a small amount of coal was mined here from the j coal; suppling a limited local market. dinosaur tracks and plant fossils the i coal swamp, which underlies the kf-6 beach sand, was home to dinosaurs. we know this because we find their footprints in the bottom of the kf-6 beach sand in this location. figure 8 depicts a couple of these better-preserved tracks. see figure 1 for the location of the photograph in figure 8. lumpy bottom beds of the kf-6 sandstone are common on fallen rocks in the area. many of these “lumps” are likely created by roaming dinosaurs at the end of the i peat swamp time. the three toed prints are likely produced by an ornithopod dinosaur (michael r. king, personal communication 2018). ornithopods are plant-eaters. these tracks can be accessed by departing the trail to the rock art at point “d” in figure 1, by making a scramble up a steep slope and following the base of the cliff of the kf-6 beach sandstone south to point “c” between muddy and rochester drainages. t o p b o tto m figure 9. unidentified plant fossil found in a boulder of the kf-6 beach sand. top to the right and bottom on the left. trekking pole for scale (3.9 feet (1.2 m )). dashed line approximates its outline. the overall length is about 12 feet (3.7 m) with the width at the upper crown at about 3.5 feet (1.1 m). location shown in figure 1. figure 9. unidentified plant fossil found in a boulder of the kf-6 beach sand. top to the right and bottom on the left. trekking pole for scale (3.9 feet (1.2 m )). dashed line approximates its outline. the overall length is about 12 feet (3.7 m) with the width at the upper crown at about 3.5 feet (1.1 m). location shown in figure 1. figure 7. rooting (black carbonaceous vertical “stringers”) in the top of the kf-6 beach sandstone supporting the development of vegetation of the j peat swamp which was established on the top of the sand. figure 8. dinosaur foot prints on the bottom of a fallen block of the kf-6 sandstone just west of the prominence separating the muddy and rochester drainages. trekking pole for scale. top of sandstone rootlet 26 cm figure 7. rooting (black carbonaceous vertical “stringers”) in the top of the kf-6 beach sandstone supporting the development of vegetation of the j peat swamp which was established on the top of the sand. figure 7. rooting (black carbonaceous vertical “stringers”) in the top of the kf-6 beach sandstone supporting the development of vegetation of the j peat swamp which was established on the top of the sand. figure 8. dinosaur foot prints on the bottom of a fallen block of the kf-6 sandstone just west of the prominence separating the muddy and rochester drainages. trekking pole for scale. top of sandstone rootlet 26 cm figure 8. dinosaur foot prints on the bottom of a fallen block of the kf-6 sandstone just west of the prominence separating the muddy and rochester drainages. trekking pole for scale. p.b. anderson ancient cretaceous beach 9 plant fossils, in growth position, are common in the vertical cliffs and in fallen blocks of the kf-6 beach sand in the area. figure 9 is an excellent example of one of these fossils. it lies on the face of a fallen block and can be observed in detail if one is willing to clamber along the outcrop west from the dinosaur tracks. the location of the block in figure 9 is shown on figure 1, location “b.” the fossil is viewable from the top of the cliff edge for those unwilling to scramble. many other specimens of the same plant can be found in the vertical cliff face, but seem to be limited to within about 1/3 mile (0.5 km) of the landward pinchout of kf-6. since discovering this community of fossils, this same fossil has been identified in a similar environment of deposition in the younger kf-4 at a different locality. attempts to identify the plant fossil are ongoing. some believe it is a tree with a tap root, but good modern examples that come close but are not a perfect match. the flaring “head” of the fossil complicates the match with a tree. the fossil occurs in the sandstone at various stratigraphic positions, but rarely in the bottom 1/3 of the sandstone. archeological note figure 1 shows the location of the rochester rock art panel. following the well maintained trial from the parking will lead you to the rock art. this spectacular panel is chipped or carved on the flat face of marine shoreface rocks from the kf-5 stratigraphic unit. if one looks around the canyons from the panel, this same stratigraphic unit can be followed on both sides of muddy canyon. it is distinguished by flat vertical faces and typically a dark “desert varnish” veneer in the lower portion of a vertical cliff-face. acknowledgments thomas a. ryer first introduced me and many other geologists to this locality. tom passed away in 2014 and this paper is a dedication to his vision and story-telling talents. thanks to michael (ryan) king for his interest and help with the trace and plant fossils, and julie howard for her help with archeology. the reviews of mary ann wright, kathy anderson, ryan king, and mark milligan were greatly appreciated. thanks to the editors of this volume for their dedication to education of the public about the marvels of utah’s geology. references anderson, p.b., and t.a. ryer, 2004, regional stratigraphy of the ferron sandstone, in chidsey, t.c., jr., r.d. adams, and t. h. morris, editors, regional to wellbore analog for fluvial-deltaic reservoir modeling—the ferron sandstone of utah: american association of petroleum geologists special studies in geology 50, p. 211-226. blakey, r. 2006, paleogeographic maps, website accessed in 2006. hintze, l.f., 2005, utah’s spectacular geology: provo, utah, sbrigham young university, 203 p. hintze, l.f., and kowallis, b. j., 2009, geologic history of utah: provo, utah, brigham young university geology studies, special publication 9, p. 225. loendorf, l. l., 1985, a radiocarbon date at the rochester creek site, utah. university of north dakota, grand forks, north dakota: on file at the utah state historic preservation office, salt lake city. filed under site number 42em392. ryer, t.a., 1981, deltaic coals of ferron sandstone member of mancos shale—predictive model for cretaceous coal-bearing strata of western interior: american association of petroleum geologist bulletin, v. 65, no. 11, p. 2323-2340. thompson, s.l., 1985, ferron sandstone member of the mancos shale—a turonian mixed-energy deltaic system: m.s. thesis, university of texas, austin, 165 p. uga 50:01 using pore system characterization to subdivide the burgeoning uteland butte play, green river formation, uinta basin, utah off-canvas toggle member login donate about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle thisistheplace2.jpeg subpage1.jpeg subpage2.jpeg subpage3.jpeg uga 50:01 using pore system characterization to subdivide the burgeoning uteland butte play, green river formation, uinta basin, utah riley brinkerhoff wasatch energy management, llc michael d. vanden berg utah geological survey mark millard novelstone geologic consulting doi: https://doi.org/10.31711/ugap.v50i.105 abstract over 240 horizontal wells with highly variable production results have been drilled in the uteland butte member (ubm) of the lower green river formation in the uinta basin, utah. the best wells have each produced more than 300,000 barrels of oil in the first 12 months of production, with conservative estimated ultimate recoveries (eurs) for several wells above a million barrels. conversely, the poorest ubm wells have initial production rates of less than 10,000 barrels of oil in the first year and will never recoup their drilling costs. pore pressure, oil viscosity, well length, and well completion are recognized as important controls on well productivity. less understood, but of equal importance, is the variability in reservoir types across the ubm play. the ubm can be divided into sub-plays by district using the dominant pore systems in each area. we defined four distinct sub-plays within the uinta basin: 1) intergranular-dominated porosity, 2) intercrystallinepore-dominated dolomite, 3) mixed intercrystalline-organic porosity, and 4) organic porosity. reservoirs in the intergranular-dominated porosity sub-play are mostly present in the form of nearshore siliciclastic and carbonate bars, such as ooid and ostracod shoals, fluvial mouth bars, and nearshore siliciclastic bars. the normal pressure and charge in these reservoirs are due to hydrocarbon migration out of the deeper basin. source rocks in this sub-play have an average maturity of 0.5 to 0.7 vitrinite reflectance (vro), too low for mainstage oil generation for these lacustrine shales (ruble and others, 2001), and produce highly viscous black wax with very low gas-to-oil ratios (gors). to date, horizontal wells drilled in this sub-play have not been economically successful. the intercrystalline-pore-dominated sub-play consists of thin (<2 to 8+ feet), laterally continuous, high-porosity dolomites that act as the best reservoirs. the reservoir is normally to slightly overpressured and is predominately charged with hydrocarbons that migrated out of the deeper basin. this sub-play was the first to be drilled horizontally, with the dolomite beds being identified as the highest quality targets. the ubm is thick in the intercrystalline-pore-dominated fairway, consisting of over 130 feet of carbonates and black shales, with only a fraction of the interval being high-porosity dolomites. this sub-play has an average vro of 0.6 to 0.8, still too low for mainstage oil generation in these rocks and produces a migrated black wax with low gors. the mixed intercrystalline-organic porosity sub-play is largely self-sourced and significantly overpressured. this sub-play has the thickest gross section and is dominated by thin dolomite beds, thicker argillaceous limestones, and thicker shaley beds, with the latter units contributing significant production. maturities in this sub-play range from 0.8 to 1.0 vro. this sub-play produces a yellow to gray wax with moderate gors. finally, the organic-porosity-dominated sub-play is highly overpressured and completely self-sourced. there is relatively little reservoir-quality dolomite, the limestones are more argillaceous, and the organic-rich carbonate shales are thicker. the productive reservoir in this sub-play consists of organic porosity largely contained in bitumen that has been expelled at lower maturity, then continued to thermally degrade with higher maturity, converting to zones of interconnected organic porosity. maturities range from 1.0 to 1.2 vro and the hydrocarbons produced are a bright yellow wax with relatively high gors. to be economically developed, each of these sub-plays requires individually tailored drilling, completion, and production strategies. by recognizing the important differences these pore systems exert on best development practices and then accurately mapping them across the basin, operators, interest owners, and regulatory agencies can more efficiently plan operations. click here for more information. order it ($5.00) cart your cart is empty login form login for discounted rates on document downloads and to access members only content username password show password remember me log in forgot your password? forgot your username? create an account utah geological association p.o. box 526356 salt lake city, ut 84152-0100  facebook  youtube © 2025 utah geological society | site by third sun member login off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf uga-geosite-wheatley.indd utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors david f. wheatley1,3, winston m. seiler2,3, and marjorie a. chan3 1chevron corporation, 1500 louisiana st., houston, tx 77002 ²kcsi aerial patrol, 1105 douglas st., bakersfi eld, ca 93308 ³dept. geology & geophysics, university of utah, salt lake city, ut 84112 davidfwheatley@gmail.com cover image: view into the nautilus from its upper entrance. the wind-swept nautilus, enigmatic clastic pipes, and toadstool landforms: geologic features of the paria plateau 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e geosites refl ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: wheatley, d.f., seiler, w.m., and chan, m.a., 2019, th e windswept nautilus, enigmatic clastic pipes, and toadstool landforms—geologic features of the paria plateau, in milligan, m., biek, r.f., inkenbrandt, p., and nielsen, p., editors, utah geosites: utah geological association publication 48, 11 p., https://doi.org/10.31711/geosites.v1i1.67. presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president d.f. wheatley, w.m. seiler, and m.a. chan geologic features of the paria plateau 3 introduction th e colorado plateau occupies much of the southwestern united states including portions of arizona, colorado, utah, and new mexico. th is region presents unobstructed views from mesa tops, beautifully colored soils, lone standing buttes, and canyons cut thousands of feet deep. th e colorado plateau represents a well-preserved window into the earth’s history. today, the rocks of the colorado plateau lie roughly horizontally, as they were deposited hundreds of millions of years ago. th e plateau’s rise has motivated rivers, in their downhill progress, to carve innumerable canyons. th ese river canyons allow any nature-lover the opportunity to gaze at 100s of millions of years of geologic history. within the larger colorado plateau, the paria plateau straddles the utah and arizona borders, and includes the vermilion cliff s national monument, the paria canyon-vermilion cliff s wilderness area, and the southern extent of the grand staircase escalante national monument (gsenm; pre-2018 boundaries). th e paria plateau is best known for spectacularly colored, wind-sculpted features such as coyote buttes and “th e wave,” where vivid colors accent cross-strata resembling a cresting ocean wave. th e plateau is also recognized for the geologically notable vermilion cliff s, buckskin gulch slot canyon, white pocket area, and the paria river canyon. although only two, dual-lane highways circumvent the plateau, several wash-boarded gravel and deeply mud-rutted roads allow access to its interior. from these dirt roads, a few sandy, four-wheel drive paths diminish as they extend and branch into the plateau’s interior. overall, the paria plateau is a relatively quiet and little-visited wilderness. on the northern portion of the paria plateau, a small but unique drainage into the paria river highlights some of the unique geologic character of the region (fi gure 1a). near the white house campground trailhead, a starting point for backpacking and hiking the paria river canyon, one fi nds two exceptional features: (1) the nautilus, which resembles the twisting contours of a nautilus seashell, and (2) world-class examples of peculiar sandstone columns, termed clastic pipes. further up this drainage, on the northern side of u.s. highway 89, there is a spectacular collection of hoodoos and toadstools — balanced rocks atop pedestals of sedimentary rock (fi gure 1a). th ese geosites provide both natural beauty and geologic insight into the history of this breathtaking landscape. study area and locality information th e specifi c features discussed in this paper are located on the northern end of the paria plateau. th e nautilus and several of the clastic pipe localities (here termed pipe sites 1 and 2) are south of u.s. highway 89 and east of the paria river near the white house campground (fi gure 1). th ese sites can be accessed via a short hike up the dry riverbed from the graded dirt road that runs between the paria contact station and white house campground. pipe sites 3 and 4 also occur among the washes and cliff s to the southeast along the river and can be accessed via dirt roads from big water, utah, and cottonwood wash road, about 3.1 miles (5 km) east of the paria contact station. to explore these worldclass examples head south on cottonwood wash road from u.s. highway 89 for about 5.5 miles (8.8 km) and turn right, proceed 0.4 figure 1. th e nautilus, clastic pipes, and toadstools can all be viewed in a relatively small region of the northern paria plateau near the paria contact station. th e nautilus (blue circle) lies up a dry riverbed and can be accessed from a graded dirt road that runs between the paria contact station and the white house campground. clastic pipes (orange circles) occur at numerous locations across the northern paria plateau. several accessible locations are near the nautilus in the same small canyon and the next small canyon directly to the south of the nautilus (described as clastic pipe sites 1 and 2 in the text). note that there are no marked walking paths to these locations. additional world-class clastic pipe sites can be accessed via cottonwood wash road (also a graded dirt road and described in the text as clastic pipe sites 3 and 4, see text for directions). qa = quaternary alluvium, qae = quaternary mixed eolian and alluvial sand deposits, qag = quaternary alluvial gravels. a. base image © google earth. image date: 2018. b. geologic map modifi ed from doelling and willis, 2006. 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 miles (0.7 km) until you reach a set of power lines, then turn right. pipe site 3 is along this road 1.3 miles (2.1 km) from the turn. pipe site 3 is within a dry riverbed to the left of the road at the junction with another dirt road. to access pipe site 4 continue along the road 2.1 miles (3.4 km) taking a left, then right, then left as you encounter forks in the road. the site is to the right of the road with pipes located along the cliffs and benches of the middle jurassic carmel formation. in addition to the examples on the paria plateau, kodachrome basin state park (about 31 miles [50 km] north) hosts several excellent examples of large clastic pipes (baer and steed, 2010). to access the toadstools travel about 1.5 miles (2.4 km) east along u.s. highway 89 from the paria contact station to a dirt parking lot on the north side of the highway. from the parking lot there is a short, marked trail that leads to the toadstools. gps localities the nautilus: 37.089615° -111.885812° clastic pipes: 37.088191°, -111.884580° (site 1); 37.084979° -111.885801° (site 2); 37.041435° -111.824171° (site 3); 37.055681° -111.853802° (site 4) toadstools: 37.108227° -111.870886° regional stratigraphy and geologic history the major rock units of the paria plateau include the sedimentary lower jurassic navajo sandstone, middle jurassic carmel formation, and middle jurassic entrada sandstone. these three jurassic units span roughly 30 million years of geologic history and were deposited between about 190 and 160 million years ago (figures 1b and 2) (doelling and others, 1989; peterson, 1994; blakey and others, 1996; caputo, 2010). the lower jurassic navajo sandstone the navajo sandstone is one of the most easily distinguished formations of the colo rado plateau. the navajo sandstone can form impressive cliffs over 2000 feet (600 m) tall and the formation covers an extensive region from zion national park to the vermilion cliffs through glen canyon to arches and canyonlands national parks (kocurek and dott, 1983, kocurek, 2003). the navajo sandstone was deposited about 190 million years ago (peterson and pipiringos, 1979; rahl and others, 2003) as a giant, arid, eolian (wind-blown) sand sea, or erg, much like the modern saharan desert erg. at its original extent, the navajo likely covered an area two and a half times its current reach—about 255,000 miles2 (660,000 km2) (kocurek, 2003; blakey, 2013). though reduced over the millennia by erosion, the navajo is still the world’s largest eolian deposit (kocurek, 2003). to put its size into perspective, it has been estimated that the volume of sand deposited by the navajo erg was between 14,400 to 33,600 miles3 (60,000– figure 2. regionally, the paria plateau consists of a set of sedimentary rock units dominantly deposited in dryland continental environments (e.g., eolian, sabkha, and fluvial deposits). the nautilus and clastic pipes occur within the middle jurassic carmel formation in the thousand pockets and paria river members, respectively. the toadstools occur within the middle jurassic entrada sandstone with cretaceous naturita formation (previously termed the dakota formation) caps. bolded formation names indicate formations in and around the geosites. the romana sandstone and morrison formation are absent with the naturita formation sitting directly on the entrada sandstone (see figure 1b; doelling and willis, 2006). the majority of the paria plateau is the navajo sandstone and younger units; however, older formations are exposed along the plateau’s flanks. ss = sandstone; sh = shale; ls = limestone. stratigraphic data modified from ron blakey’s colorado plateau geosystems website. d.f. wheatley, w.m. seiler, and m.a. chan geologic features of the paria plateau 5 140,000 km3), the equivalent of 6 to 14 million years of continuous deposition at the mouth of the mississippi river (kocurek, 2003). in recreating the paleogeography of the navajo erg, rahl and others (2003) envision that most of the navajo sands were derived from the ancestral appalachian mountains. they propose rivers carrying sediment from their headwaters in the ancient appalachians to the western shores of the supercontinent pangea, flowing north of the ancestral rockies. from here, the river sand was blown to the south to form the navajo erg. on the paria plateau, the navajo sandstone is a fineto medium-grained sandstone with large beds of cross-stratified laminae (internally inclined, thin sandstone layers) that represent the sloping face of the ancient dune deposits. the formation can be colored in a combination of whites, oranges, yellows, tans, and reds depending on the types and amount of iron-rich minerals present in the sandstone. the navajo sandstone is the sculpted rock of the coyote buttes and forms the prominent cliffs of the dive, steamboat rock, and white pocket. locally, the navajo sandstone has been studied extensively as a record of diagenetic (i.e., post-depositional) fluid flow as preserved in its coloration (eichhubl and others, 2004; beitler and others, 2005; nielsen and others, 2009; nielsen and chan, 2010; potter and chan, 2011; seiler and chan, 2014), as an analog to surficial weathering processes on mars (chan and others, 2008; chan and others, 2011), and for its preservation of dinosaur footprints (irmis, 2005; loope, 2006). the middle jurassic carmel formation unconformably deposited atop the navajo sandstone is the carmel formation. this contact (i.e., surface between two formations), termed the j1 unconformity, represents a significant period of erosion that formed a paleotopography. regionally, the carmel formation is composed of inter-fingered layers of limestone, sandstone, shale, and gypsum and is found throughout southern utah (caputo, 2010). at these geosites, the carmel formation crops out as small (about 100 feet [30 m]) cliffs of variegated red, pink, and white, horizontally bedded sandstone layers. the carmel formation can be subdivided into five members (i.e., groupings of similar rock types). from oldest to youngest they are: (1) the judd hollow member, (2) the thousand pockets member, (3) the crystal creek member, (4) the paria river member, and (5) the winsor member (sprinkel and others, 2011; doelling and others, 2013). for the purposes of this paper, we use the stratigraphy put forth by doelling and others (2013), which includes the thousand pockets member in the carmel formation; however, the thousand pockets member was previously included as a tongue of the navajo sandstone or the page sandstone (phoenix, 1963; pipiringos and o’sullivan, 1978) and some geologic maps and workers still use this designation. these members represent the cyclic rise and fall of an inland sea and the corresponding advance and retreat of an arid coastal plain and nearby dune field (jones and blakey, 1997). except for the judd hollow member, all carmel formation members are exposed at the main study sites. the middle jurassic entrada formation and cretaceous naturita formation the entrada formation conformably lies above the jurassic carmel formation and is another eolian unit, similar to the navajo sandstone, but with much more variability. its sandy beds represent dunes and marginal beach deposits. it is generally softer than the navajo sandstone and is not such a prominent cliff former. regionally, there are intervening units between the entrada formation and the younger cretaceous naturita formation (previously termed the dakota formation). however, locally, the intervening units are absent because of erosion resulting in isolated naturita blocks of ancient river-deposited coarse-grained to pebbly sandstone uncomfortably resting on top of spires of exposed entrada sandstone. geologic features of the paria plateau the nautilus observations the nautilus is a short, twisting slot canyon formed in the approximate shape of a logarithmic spiral, reminiscent of the cutaway of a chambered nautilus—the marine cephalopod (figures 3 and 4). it extends approximately 100 feet (30 m) in length, is 16 feet (5 m) wide at is upper opening, and 3 feet (1 m) in width in its narrowest central portion. the slot is cut to an approximate depth of 33 feet (10 m) and steepens in slope as it twists from top to bottom (note: visitors should respect the delicate nature of the nautilus and take care to preserve the feature when exploring its natural beauty). the nautilus spirals through the bleached white, eolian cross-bedded thousand pockets member of the carmel formation. several orangeto yellow-colored inchto sub-inch(mm to cm-scale) thick liesegang bands—repeating stripes of minerals (iron oxides) that streak across and overprint the original strata—and isolated iron cemented concretionary masses decorate the feature. sub-quarter inch (mm-scale) concretions are evenly dispersed in the lower portion, near the outflow of the nautilus into the main wash. though the immediate area displays fracturing in the sandstone, the nautilus is not correlated with any obvious local or regionally oriented joint set or fracture pattern. little capacity for water collection exists upstream of the nautilus feature; the drainage area above the nautilus occupies less than 2420 yards2 (2023 m2). prominent throughout the nautilus, on horizontal, vertical, and sloped faces, are inch to sub-inch (cm-scale) erosional, stepped tread and riser features. the riser portions of these features are consistently oriented towards the upper entrance of the nautilus. 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 tread and riser features visually resemble sastrugi—irregular ridges carved into hard snow by wind (loope and others, 2008). parallel-oriented fl uting in the sandstone is noted downstream of erosionally resistant concretionary features and other protuberances within the core of the nautilus. interpretations narrow and twisting slot canyons are common in sandstone on the colorado plateau. although most slot canyons are the product of repeated, powerful fl oods exploiting joints in the sandstone, wind is a less recognized, but nevertheless signifi cant erosive mechanism in a smaller subset of slot canyons. wind carrying abrasive sand grains is the primary erosional force that has shaped the nautilus. while the spiral feature may have been initiated by water exploiting a weakness in the sandstone, there is little evidence to suggest that water remains the predominant erosional force—the drainage above the entrance is extremely limited, there is no evidence of runnels or depositional outfl ow indicative of fl owing water at the exit of the nautilus, and horizontal surfaces do not display channeling by water at low points. instead, the presence of tread and riser features and fl uting downwind of concretionary protuberances are prominent ventifacts documenting the erosive force of wind in shaping the nautilus (fi gures 3 and 4). th e orientation of these features is consistent with wind channeled up the larger drainage, locally from the southwest, and being directed into the nautilus to enter from the south at its upper opening. loope and others (2008) describe sand-carrying winds as the primary active abrasive mechanism for the shaping of the wave, another photogenic wind shaped feature approximately 9 miles (14 km) southwest on the paria plateau. th ey document and defi ne erosional treads and risers cut by the impact of wind-blown sand undercutting fi ne microbial crusts lying just under the sandstone surface. th e subtle liesegang coloration and presence of iron-oxide concretions are illustrative of past changes in groundwater chemistry (e.g., a reducing/oxidizing fl uid boundary) while the rock was still buried (wang and others, 2015). th ese chemical reaction fronts and associated coloration have been described extensively by beitler and others (2005), potter and chan (2011), and seiler and chan (2014), although the precise timing of these fl uid movements remain a mystery. clastic pipes observations clastic pipes form vertical sandstone columns that typically contrast in color and texture from the host sandstone (fi gures 5a-d). th e white sandstone pipes stand out against a background of red, layered, sandstone host rocks that contain mud cracks, root casts, and cross-beds. interspersed between the layered sandstone strata, the carmel formation has several beds of volcanic clasts and fragments, and multiple thin volcanic ash layers. th e pipes are feet (10s of centimeters to meters) in diameter and are composed of quartz-rich sand and volcanic fragments similar to the surrounding host rock. however, pipes have internal, radial grading with outward coarsening (i.e., they have a consistent outward increase in grain size). many pipes have an outer rind with a 0.5 to 2 inches (1–5 cm) rind of coarser grained sediment. pipes also have significant diagenetic, fl uid-related characteristics including bleached white rings or “halos” surrounding many pipes. pipes also have numerous calcite concretions or nodules that form preferentially on the outer portion of the pipes. th e clastic pipes occur across an approximately 100 foot (30 m) thick stratigraphic interval in the paria river member of the carmel formation. two dated volcanic ash layers approximately 3 feet (1 m) below and 16 feet (5 m) above the interval provide lower and upper bounds on the sediments with ages of about 166 million years and about 164 million years, respectively (kowallis and others, 2020). pipes emanate from a single “source” bed, figure 3.view into the nautilus from its upper entrance. arrows indicate the direction of wind, funneled into the nautilus from the south. tread and riser features, indicative of erosion by wind, are found at this upper entrance. th e nautilus occurs in the th ousand pockets member of the carmel formation. d.f. wheatley, w.m. seiler, and m.a. chan geologic features of the paria plateau 7 move upward through the 100 feet (30 m) stratigraphic interval, and fl are outwards as they terminate. syn-sedimentary faulting and folding further deformed the pipe-bearing interval during to shortly aft er deposition. th e faulting is particularly concentrated around in the pipes and the folding occurs most commonly just above the source layer. interpretations th e presence of root casts, mud cracks, and cross-bedded strata along with the regional geology indicate that the pipes formed on an arid coastal plain, also called a sabkha. th e sabkha was located between an evaporative, shallow inland sea and a dune fi eld and was dissected by ephemeral streams (i.e., water only fl owed for a portion of the year). th ese marginal, coastal environments are characterized by low biological productivity, wetting and drying (e.g., mud cracks and gypsum formation), and a high water table. clastic pipes form through a process called fl uidization (fi gure 5e) (owen, 1987; wheatley and others, 2016; wheatley and chan, 2018). fluidization occurs through the rapid, violent release of pressurized, water-saturated sediments in the subsurface by a trigger such as an earthquake, sediment loading, or a sudden fl ood (obermeier, 1996; owen and moretti, 2011). th e water-saturated sabkha sediments with interbedded sands (i.e., the sand source for the pipes) and muds (i.e., the seal to allow pressurization of the sand layers) provide an ideal pipe-forming environment. pipe formation requires a specifi c set of conditions. (1) th e sandy source layer is deposited below a fi ne-grained mud layer (fi gure 5e, panel 1). th e mud layer acts as a seal above the source layer. (2) a trigger (i.e., a rapid energy release, examples include an earthquake, landslide, fl ood, etc.) then initiates liquefaction and pressurizes the source layer. (3) once the pressure in the source layer rises above the sealing capacity of the mud layer, the muddy seal fails resulting in the catastrophic release of pressure (fi gure 5e, panel 2). (4) th e fl uid moves rapidly to the surface causing fl uidization and entraining sediment. at the surfaces the pipes erupt as sand volcanoes and form the fl ared horizons observed in outcrop. th e process of fl uids moving upward through the host sediment sorts the sediment giving pipes their characteristic grading. th is process is incredibly violent and results in the faulting and folding of the surrounding host sediment. (5) aft er pipe formation the pipes de-water causing a reduction in volume of the pipe sediment and the formation of “relaxation” structures such as inward (i.e., pipe-ward) dipping faults (fi gure 5e, panel 3). (6) aft er continued deposition and erosion/reworking of the above-ground portions of the sand volcanoes, diagenetic fl uids moved through the pipes bleaching them white and leaving behind carbonate concretions/ nodules (fi gure 5e, panel 4). (7) finally, aft er the sediment turned to rock and was uplift ed, the better cemented pipes eroded out in positive relief as sandstone spires (fi gure 5e, panel 5). toadstools and hoodoos observations colorful, tall, towering toadstools and hoodoos are visible on an easy short hike farther upstream from the nautilus and clastic pipes (but most easily accessed directly from u.s. highway 89) (fi gure 6). delicately balanced hard, resistant rock caps of cretaceous naturita formation rest on weaker more easily eroded narrow stalks of entrada sandstone (fi gure 6). th e mushroom and toadstool landforms have the erosional-resistant cap, while the hoodoos are erosional towers without the caps. together, these comprise an enchanting fairyland scene. interpretations after the sedimentary deposits were buried and lithifi ed, later uplift subjected the formations to forces that created joint and fracture patterns (i.e., preferential zones of weakness), somewhat like checkerboard lines and squares. through time, wind and water worked away at the sides of the chunky cretaceous naturita formation caps resulting in isolated blocks resting on the soft er, older, white and red jurassic entrada sandstone (doelling and others, 1989; doelling and willis, 2006). th e entrada sandstone is slightly protected by the naturita caps, but over a long period of time, the soft er entrada sandstone jurassic eroded away and is undercut beneath the resistant cap. what is left are the dramatic spires that resemble a mushroom shape of perched and balanced boulders in a landscape of pedestals and towers (fi gure 6). figure 4. th e central portion of the nautilus displays fl uted ventifacts downwind of erosionally resistant concretions, while sheltered, leeward faces of the nautilus lack features indicative of erosion by wind. liesegang bands and micro concretions cross cut stratigraphy, and occur in the lowermost portions of the nautilus. th e nautilus occurs in the th ousand pockets member of the carmel formation. 8 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 figure 5. clastic pipes weather out among the washes and cliffs of the carmel formation (a, b, c, and d). the white pipes stand out against a red host rock (paria river member of the carmel formation). these images show two additional sites (sites 3 and 4) referenced in the caption of figure 1. panel e details the formation process of clastic pipes (see section on interpretation). pipes form through liquefaction and fluidization (i.e., the violent release of pressurized subsurface fluids) (modified from wheatley and others, 2016). d.f. wheatley, w.m. seiler, and m.a. chan geologic features of the paria plateau 9 conclusions the paria plateau provides an excellent outdoor classroom to explore over 200 million years of earth’s history and exposes ancient wonders through intricate erosion. the arid landscape of the jurassic period resulted in the deposition of broad eolian sand seas and coastal plains (i.e., sabkhas). these strata were subsequently deformed by fluidization to form sand volcanoes, and were later colored and cemented by iron rich fluids to form a vertical tapestry of reds, oranges, and yellows. then, as the colorado plateau was uplifted, water and wind formed the beautiful canyons and features like the nautilus and revealed the clastic pipes as sandstone spires. modern erosion acting on sediments with differing resistance to weathering has formed the towering and precarious toadstools and hoodoos. these geologic features need to be protected for generations to come so that others can come learn and marvel. acknowledgements we thank doug sprinkel, mark hansford, mark milligan, and bob biek for their helpful reviews of this manuscript. this paper is the result of geologic field studies completed at the university of utah, with support from the national science foundation, national aeronautics and space administration, the gdl foundation, the rocky mountain association of geologists, the american association of petroleum geologists, the geologic society of america, and the wheaton college scholastic honors society. references baer, j.l., and steed, r.h., 2010, geology of kodachrome basin state park, utah, in sprinkel, d.a., chidsey, t.c., jr., and anderson, p.a., geology of utah’s parks and monuments: utah geological association publication 28 (3rd edition), p. 467–482. beitler, b., parry, w.t., and chan, m.a., 2005, fingerprints of fluid flow—chemical diagenetic history of the jurassic navajo sandstone, southern utah, usa: journal of sedimentary research, v. 75, no. 4, p. 547–561. blakey, r.c., havholm, k.g., and jones, l.s., 1996, stratigraphic analysis of eolian interactions with marine and fluvial deposits, middle jurassic page sandstone and carmel formation, colorado plateau, usa: journal of sedimentary research, section b: stratigraphy and global studies, v. 66, no. 2. p. 324–342. blakey, r.c., 2013, early jurassic map, paleogeography of southwestern north america series: online, www.deeptimemaps. com, accessed feb. 6, 2019. caputo, m.v., 2010, geology and sedimentary highlights of the paria canyon–vermilion cliffs wilderness, utah and arizona, in sprinkel, d.a., chidsey, t.c., jr., and anderson, p.a., geology of utah’s parks and monuments: utah geological association publication 28 (3rd edition), p. 559–587. chan, m.a., nicoll, k., ormö, j., okubo, c., and komatsu, g., 2011, utah’s geologic and geomorphic analogs to mars—an overview for planetary exploration: geological society of america special papers, v. 483, p. 349–375. chan, m.a., yonkee, w.a., netoff, d.i., seiler, w.m., and ford, r.l., 2008, polygonal cracks in bedrock on earth and mars— implications for weathering: icarus, v. 194, no. 1, p. 65–71. doelling, h.h., davis, f.d., and brandt, c.j., 1989, the geology of kane county, utah—geology, mineral resources, geologic hazards: utah geological and mineral survey bulletin 124, p. 192. doelling, h.h., and willis, g.c., 2006, geologic map of the smokey mountain 30’x60’ quadrangle, kane and san juan counties, utah, and coconino county, arizona: utah geological survey map 213, 2 plates, scale 1:100,000. doelling, h.h., sprinkel, d.a., kowallis, b.j., and kuehne, p.a., figure 6. these toadstools are slender columns of the middle jurassic entrada sandstone with an erosionally resistant cap of cretaceous naturita formation. these whimsical features resemble toadstools or mushrooms. in some cases, the erosionally resistant cap is no longer present leaving a slender spire of rock termed a hoodoo. 10 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 2013, temple cap and carmel formations in the henry mountains basin, wayne and garfield counties, utah, in morris, t.h., and ressetar, r., editors, the san rafael swell and henry mountains basin-geologic centerpiece of utah: utah 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paleogeography symposium 2: society of economic paleontologists and mineralogists rocky mountain section, denver, co, p. 101–116. kocurek, g., 2003, limits on extreme eolian systems: sahara of mauritania and jurassic navajo sandstone examples, in chan, m.a. and archer, a., editors, extreme depositional environments: mega end members in geologic time: geological society of america special paper 370, p. 143–156. kowallis, b.j., sprinkel, d.a., christiansen, e.h., steed, s., and wheatley, d.f., 2020, rhyolite ignimbrite boulders and cobbles in the middle jurassic carmel formation of utah and arizona: age, composition, transport, and stratigraphic setting: geology of the intermountain west, v. 7, no. 3, in press. loope, d.b., 2006, dry-season tracks in dinosaur-triggered grainflows: palaios, v. 21, no. 2, p. 132–142. loope, d.b., seiler, w.m., mason, j.a., and chan, m.a., 2008, wind scour of navajo sandstone at the wave (central colorado plateau, usa): the journal of geology, v. 116, no. 2, p. 173–183. nielsen, g.b., chan, m.a., and petersen, e.u., 2009, diagenetic coloration facies and alteration history of the jurassic navajo sandstone, zion national park and vicinity, southwestern utah, in tripp, b.t., krahulec, k., and jordon, j.j., editors, geology and geologic resources and issues of western utah: utah geological association publication 38, p. 67–96. nielsen, g.b., and chan, m.a., 2010, geologic map and coloration facies of the jurassic navajo sandstone, snow canyon state park and areas of red cliffs desert reserve, washington county, utah: utah geological survey open-file report 561, scale 1:10,000. obermeier, s.f., 1996, using liquefaction-induced features for paleoseismic analysis, in mccalpin, j., editor, paleoseismology: san diego, california, academic press, p. 331–396. owen, g., 1987. deformation processes in unconsolidated sands: geological society of london special publication 29, p. 11–24, doi:10.1144/gsl.sp.1987.029.01.02. owen, g., and moretti, m., 2011, identifying triggers for liquefaction-induced soft-sediment deformation in sands: sedimentary geology, v. 235, p. 141–147, doi:10.1016/j.sedgeo.2010.10.003. peterson, f., and pipiringos, g.n., 1979, stratigraphic relations of the navajo sandstone to middle jurassic formations, southern utah and northern arizona: u.s. geological survey professional paper 1035-b, p. b1–b43. peterson, f. 1994. sand dunes, sabkhas, streams, and shallow seas—jurassic paleogeography in the southern part of the western interior basin, in caputo, m.v., peterson, j.a., and franczyk, k.j., editors, mesozoic systems of the rocky mountain region, usa: rocky mountain section, society of economic paleontologists and mineralogists, p. 233–272. pipiringos, g. n., and o’sullivan, r. b., 1978, principal unconformities in triassic and jurassic rocks, western interior united states—a preliminary survey: u.s. geological survey professional paper 1035-a, p. 29. phoenix, d.a., 1963, geology of the lees ferry area, coconino county, arizona: u.s. geological survey bulletin 1137, 86 p., 3 plates, scale 1:24,000. potter, s.l., and chan, m.a., 2011, joint controlled fluid flow patterns and iron mass transfer in jurassic navajo sandstone, southern utah, usa: geofluids, v. 11, p. 184–198. rahl, j.m., reiners, p.w., campbell, i.h, nicolescu, s., and allen, c.m., 2003, combined single-grain (u-th)/he and u/pb dating of detrital zircons from the navajo sandstone, utah: geology, v. 31, no. 9, p. 761–764; seiler, w.m., and chan, m.a., 2014, coloration and diagenetic history of jurassic navajo sandstone at coyote buttes, paria canyon-vermillion cliffs wilderness, utah and arizona, in maclean, j.s., biek, r.f., and huntoon, j.e., editors, geology of utah’s far south: utah geological association publication 43, p. 237–258. d.f. wheatley, w.m. seiler, and m.a. chan geologic features of the paria plateau 11 sprinkel, d.a., doelling, h.h., kowallis, b.j., waanders, g., and kuehne, p.a., 2011, early results of a study of middle jurassic strata in the sevier fold and thrust belt, utah, in sprinkel, d.a., yonkee, w.a., and chidsey, t.c., jr., editors, sevier thrust belt—northern and central utah and adjacent areas: utah geological association publication 40, p. 151–172. wang, y., chan, m.a., and merino, e., 2015, self-organized iron-oxide cementation geometry as an indicator of paleo-flows: scientific reports, v. 5, p. 10792. wheatley, d.f., and chan, m.a., 2018, clastic pipes and soft-sediment deformation of the jurassic carmel formation, southern utah, u.s.a.—implications for pipe deformation mechanisms and host-rock controls: journal of sedimentary research, v. 88, p. 1076–1095, doi:10.2110/jsr.2018.45 wheatley, d.f., chan, m.a., and sprinkel, d.a., 2016, clastic pipe characteristics and distributions throughout the colorado plateau—implications for paleoenvironment and paleoseismic controls: sedimentary geology, v. 344, p. 20–33, doi:10.1016/j. sedgeo.2016.03.027. uga-geosite-chidsey-arches.indd 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors thomas c. chidsey, jr., and grant c. willis utah geological survey, po box 146100, salt lake city, utah 84114-6100 tomchidsey@utah.gov cover image: double arch, consisting of two giant pothole arch spans joined at one end, is the third largest in the park. photo courtesy of ben erickson. landscape arch, delicate arch, and double arch in arches national park, southeastern utah 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e geosites refl ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: chidsey, t.j., and willis, g.c., 2019, landscape arch, delicate arch, and double arch in arches national park, southeastern utah, in milligan, m., biek, r.f., inkenbrandt, p., and nielsen, p., editors, utah geosites: utah geological association publication 48, 12 p., https://doi.org/10.31711/ geosites.v1i1.54. presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president t.c. chidsey, jr., and g.c. willis landscape arch, delicate arch, and double arch 3 introduction arches national park in southeastern utah (figure 1) has the greatest concentration of natural rock arches in the world. the park is located in a geologic region called the paradox fold and fault belt in the northern paradox basin and showcases spectacular and classic colorado plateau geology with its colorful sedimentary rocks, ancient sand dunes, cliffs, domes, fins, and pinnacles, as well as the arches. the arches in the park and the surrounding region were formed by a unique set of circumstances involving middle pennsylvanian (about 308 million years ago [ma]) to late triassic (200 ma) movement of subsurface salt layers, middle pennsylvanian to late cretaceous (about 70 ma) deposition, and tertiary and quaternary (23 ma to the present) folding, faulting, erosion, and salt dissolution. massive, hard, brittle sandstones jointed by folding, resting on or containing soft layers or partings, and located near fold structures such as salt-cored anticlines undergoing dissolution, and a dry climate, all favor the formation of arches. rarely do all these phenomena occur in one place, but they do in arches national park. the natural arch and bridge society (nabs) stated, “a natural arch is a rock exposure that has a hole completely through it formed by the natural, selective removal of rock, leaving a relatively intact frame.” they also make it clear that a natural bridge (which is at least partially formed by flowing water) is one type of natural arch (nabs website) (see a bit of perspective, below, for more explanation). using their own criteria, stevens and mccarrick (1988) catalogued over 2000 natural arches in arches national park; most have unique characteristics that could qualify them as geosites. however, the three most famous arches in the park, and perhaps the world, are landscape arch, delicate arch, and double arch, and thus these were selected as the geosites for this paper. how to get there arches national park is about 230 miles (370 km) or a 3-hour drive from salt lake city, utah, via interstate 15 (i-15), u.s. highway 6, i-70, and u.s. highway 191. enter the park about 3 miles (3.8 km) north of moab on u.s. highway 191 (figures 1 and 2). proceed through the entrance station (fee required) and follow arches scenic drive (paved) 16.7 miles (26.9 km) northeast through the park to devils garden trailhead that leads to landscape arch (figure 2). on popular summer, holiday, and weekend days, the devils garden trailhead parking lot fills early, as do the parking lots for the delicate arch and double arch trailheads. follow the relatively easy, flat, well-signed sand and gravel trail to landscape arch, 38°47'27" n., 109°36'27" w., elevation 5290 feet (1612 m). there are “rustic” restrooms at the parking lot/trailhead. the round-trip walking distance to landscape arch is 1.6 miles (2.6 km). the delicate arch trailhead and parking lot are reached by first proceeding 11.7 miles (18.8 km) from the park entrance on arches scenic drive to the intersection with delicate arch road. follow delicate arch road another 1.2 miles (1.9 km) to the trailhead and parking lot (figure 2). the 3-mile (4.8 km) round trip trail to see utah’s iconic free-standing delicate arch up close and personal is one of the more difficult hikes in the park, but well worth it. this steadily uphill trail, climbing 480 feet (146 m), also passes the historical wolfe ranch cabin and a wall of ute indian petroglyphs. the trail is well marked and heavily used as it is the most popular trail in the park. the first half mile is well-defined, but then it climbs steadily over slickrock and is marked with cairns. it has no shade—hikers should wear hats and carry at least 1 quart of water, more if it’s hot! just before getting to delicate arch, the trail traverses a narrow rock ledge for about 200 yards (180 m) and is about 5 feet (1.5 m) wide; however, it is solid rock and dips into a rock wall that can provide a secure pathway for those who may be nervous about heights. once past this short stretch is the ultimate view of the most famous and spectacular site in utah—delicate arch, 38°44'38" n., 109°29'57" w., elevation 4811 feet (1466 m). for most, the round trip should take about 2.5 to 3 hours depending how much lingering occurs at the arch. there are rustic restrooms at the parking lot/trailhead. figure 1. location of arches national park, southeastern utah, and surrounding parks, towns, and highways. 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 figure 2. geologic map of the arches national park area, grand county, utah, showing the locations of landscape, delicate, and double arches (red dots). modified from doelling (2010). the double arch trailhead and parking lot in the garden of eden area are reached by first proceeding 9.1 miles (14.7 km) from the park entrance on arches scenic drive to the intersection with the windows road. follow the windows road 2.5 miles (4.0 km) to the trailhead (figure 2) and relatively small parking lot. a relatively flat, gravel-surfaced trail leads to the base of double arch, 38°41'29" n., 109°32"25" w., elevation 5172 feet (1576 m). the round-trip distance to the arch is 0.5 miles (0.8 km). t.c. chidsey, jr., and g.c. willis landscape arch, delicate arch, and double arch 5 stratigraphy most arches in arches national park are formed in two geologic units: the slick rock member of the entrada sandstone and the moab member of the curtis formation (figure 3). landscape and double arches are in the slick rock member whereas delicate arch includes both slick rock and moab members. entrada sandstone the slick rock member of the late middle jurassic (160 ma) entrada sandstone overlies the red-brown dewey bridge member of the carmel formation (figure 3). the slick rock is a massive, well-indurated, red-orange or brown, very fine to fine-grained sandstone containing sparse and scattered medium to coarse grains (doelling, 2010). the sand grains are held together with calcite and iron-oxide cement. it commonly weathers to form smooth cliffs and bare-rock slopes. parts of the member are distinctly cross-bedded (eolian deposition), whereas other parts figure 3. lithologic column showing formations and members exposed in arches national park, including ages, map symbols, thicknesses, weathering habits, and lithologies. after doelling (2010). are planar bedded. in most places color variations form stripes or bands. in some cases, the banding is hardly noticed, in others it is pronounced. locally, indentations parallel the banding (doelling, 2010). small holes called tafoni are commonly aligned along crossbed laminae (the boundaries of ancient dune sets). tafoni are created by differential weathering and case hardening in areas where groundwater has weakened the cement in the rock, commonly near small variations or imperfections. once started, the voids provide space where water can accumulate and be protected from evaporation, further weakening cement in the rock and promoting growth. in case hardening, the mineral-laden water wicks toward the outer exposed face of the rock where the water quickly evaporates, depositing the cement, which strengthens the rock “case.” the slick rock member of the entrada sandstone is perhaps the most important geologic unit in arches national park. most of the arches in the park are positioned along its lower and upper contacts and along the easily weathered beds in the middle of the unit (doelling, 2010). where the slick rock is not exposed as a cliff, fin, or arch, the outcrop band is covered or partly covered by large irregular fields of self-derived wind-blown sand (figure 2). it was deposited in a coastal dune field and interfingers with sabkha deposits (figure 4a). the slick rock member is normally 200 to 350 (60–110 m) feet thick within arches national park (doelling, 2010). the contact between the entrada sandstone and the underlying carmel formation is generally sharp, but locally extraordinarily irregular. in some cases, the slick rock member of the entrada intertongues with the dewey bridge member of the carmel (doelling, 2010). near the arches national park visitor center, the dewey bridge member appears to be angularly planed off by the slick rock member, but this is not an unconformity; it is instead attributed to soft-sediment deformation in the dewey bridge. curtis formation the moab member of the early late jurassic (155 ma) curtis formation overlies the slick rock member of the entrada sandstone and is about 60 to 120 feet (18–37 m) thick in the park (figure 3). most of the member is pale-gray to nearly white (pale-grayorange, pale-yellow-brown, or pale-gray on seldom-seen fresh surfaces), fineto medium-grained, calcareous, thick to massive, cliff-forming sandstone that forms a prominent capping bench over entrada cliffs. the sandstone is typically well indurated, has low-angle cross-bedding, and is generally strongly jointed (doelling, 2010). the sandstone resembles the navajo in color, cementation, and differential etching of cross-bed laminae. outcrops form bare-rock, sloping, jointed benches on each side of salt valley. these strong northwest-trending joints with minor cross-joints form a “biscuit” pattern and are spectacular when viewed from an airplane (figure 5). the curtis forms the upper 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 part of delicate arch and outcrops can be viewed along the delicate arch hike. the massive sandstone is a unique eolian facies of the curtis only present in the arches area—the curtis of central and northeastern utah is characterized by shallow marine, ledgy, slope-forming, pale-green-gray, silty sandstone (figure 4b). the contact between the entrada sandstone and the curtis formation is the j-3 unconformity of pipiringos and o’sullivan (1978) and is sharp in most areas. in the north part of the park beds just above the unconformity consist of about 25 feet (7.6 m) of slope or recess-forming, red-brown, thin-bedded, silty, fine-grained sandstone, similar, except for the color, to the curtis in most parts of central utah. in the central part of the park, these beds thin to a reddish siltstone indenture a few feet thick. in the south part of the park, the contact is just a subtle line between banded slick rock member with truncated cross-beds below and monotone pale-gray moab member with no prominent bedding above. structure and geologic history during the pennsylvanian, southeastern utah and southwestern colorado subsided to form the paradox basin, a restricted shallow embayment of an open marine sea that lay to the southeast, and the uncompahgre highlands (ancestral rockies) were located to the northeast. conditions were arid and high evaporation rates led to a net inflow of seawater and elevated salinity. the structural geology in arches national park, and of the fold and fault belt of the northern paradox basin, is primarily the result of the evaporite minerals deposited in the pennsylvanian paradox formation, and their subsequent movement and dissolution (doelling, 2010). the fold and fault belt of the paradox basin extends southeast from arches national park into colorado. this structural belt is characterized by a series of northwest-trending anticlines and faults that developed in response to movement caused by crustal forces or by shallow deposits of pennsylvanian salt. salt, which was deposited as part of the middle pennsylvanian paradox formation, has a low specific gravity, is ductile and less dense than the surrounding rocks, and behaves plastically. the salt moves due to high confining pressure from the weight of the overlying column of rocks or sediments. salt movement can push up and raise overlying strata in some places, while the salt is squeezed and thinned in others. this movement (often piercing the strata above [diapiric movement]) progresses along zones of weakness or areas of low confining pressure, forming large folds (figures 2 and 6). later, these folds collapsed when groundwater dissolved the salt. the most important structural feature in the park is the salt valley–cache valley salt wall or diapir (figures 2 and 6). the diapiric salt wall formed mostly from middle pennsylvanian to late triassic time (doelling, 2010). it was covered with sediment laid down figure 4. paleogeographic maps of utah during the late middle to early late jurassic: (a) entrada sandstone – 160 ma, and (b) curtis formation – 155 ma. modified from blakey and ranney (2008). a b t.c. chidsey, jr., and g.c. willis landscape arch, delicate arch, and double arch 7 in late triassic to late cretaceous time, but slow salt movement affected formation thicknesses. thereafter, the area was folded and faulted, probably during the late cretaceous-early tertiary (about 70 to 30 ma) laramide orogeny (a regional mountain-building event), that created many large folds throughout the rocky mountain region. in the arches area, the axes of the anticlines were superimposed over the northwest-trending salt walls of the paradox fold and fault belt. many of the brittle sandstone formations were fractured (jointed) during the folding (figure 5). tertiary faults paralleled and displaced the salt walls (figures 2 and 6). with the uplift of the colorado plateau throughout the tertiary (66 to 1.8 ma), but accelerating in late tertiary time (about 15 ma) and continuing to the present, the colorado river and its tributaries eroded the sedimentary formations and formed deep canyons, slopes, and cliffs, depending on erosional resistance. after erosion cut down deep enough, groundwater via faults and joints reached the upper parts of the region’s salt walls, including those of the salt valley–cache valley feature, and dissolved the salt. the ensuing collapse created graben-valleys that overlie the salt walls today, as well displayed in arches national park (figures 2 and 6) and much of the paradox fold and fault belt. a bit of perspective a quick search of literature ranging from tourist brochures to textbooks reveals myriad, often contradictory and exaggerated claims about natural stone arches (many communities have claimed theirs as the “biggest”). in an attempt to settle many arguments, and bring some science into the equation, a group of dedicated geologists, mathematicians, physicists, geographers, and enthusiasts formed the natural arch and bridge society (nabs) and set about establishing carefully crafted rules, definitions, and criteria. as a result, much better information is available, including careful measurements following well-defined criteria, of all the largest known arches in the world. while a surprise could be lurking somewhere, we can be fairly confident of how utah arches fit in the world hierarchy. below are a few of the more important facts. the interested reader is referred to nabs’s excellent website at naturalarches.org for more detailed information. 1. nabs has chosen a simple definition for a natural arch: “a rock exposure that has a hole completely through it formed by the natural, selective removal of rock, leaving a relatively intact frame.” they provide a detailed discussion, simplified here as: (1) it must be made of rock; (2) the rock must be substantially surfigure 5. rock arch development stages in arches national park (the order generally matches the outcrops from right to left in the photo below). 1 – regional uplift by salt diapirism, then collapse due to dissolution of salt in the pennsylvanian paradox formation, created northwest-trending joint systems in brittle jurassic sandstones. 2 – following the erosion of the overlying rocks, weathering and erosion processes by wind, water, ice, and temperature variations widened and deepened joints creating rock walls or fins. 3 – continued erosion, including mass wasting (rock falls), from opposing sides of fins eventually led to breakthrough and the creation of wall arches. 4 – free-standing arches, like delicate arch, are all that remain after other arches collapse and surrounding fins erode away. the enduring arches enlarge by continued weathering and erosion processes until they too collapse. importantly, being in the open air actually slows the weathering process because rain and snow that fall on the arch evaporate quickly, leaving behind cements that running water would wash away. also, pressures that develop in the arch actually strengthen the bounds between sand grains. the bottom photo is an aerial panoramic view northwest towards the devils garden area along the northeast flank of the salt valley anticline showing extensive jointing in the moab member of the curtis formation (right area of the photo) and fins in various stages of erosion; several arches are present in the area of the fins. photo from hamblin (2004). 8 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 rounded by air; (3) the hole must conform to the mathematical, or topological, definition of a hole (they provide a long discussion of “a hole”!); (4) the hole must have formed from natural, selective removal of rock; and (5) the frame of rock that surrounds the hole must still be relatively intact. see the nabs website for much additional discussion on the nuances of arch definition. 2. nabs has chosen to not make size (width, volume, height, etc.) of the opening part of the definition. they also leave it to the viewer to decide if a feature is an arch, a tunnel, or some other feature, and qualify arches with terms such as miniature, significant, major, and giant. thus, any count of the world’s arches is highly subjective. for example, steven and mccarrick’s (1988) “count” of arches in arches national park is based on their own criteria; other researchers would come up with a different number. for their own database, nabs defines a “significant” natural arch as having a product of two orthogonal opening dimensions of about 110 square feet (10 m2) or more. but smaller openings (of which there are millions) are still arches—they just do not make it into the database. 3. nabs emphasizes that bridges are a type of arch. that they are not arches is one of the most common misconceptions that persists today. nabs actually defines at least 17 types of arches based on their shape and forming process; bridges fit into at least two of these types. furthermore, many bridge-type arches are hybrids—only partially formed by or spanning flowing water or dry channels. for comparisons, all types of arches, including bridge-types, are considered together. thus, there is no fundamental difference between a natural arch and a natural bridge. 4. everyone wants to know how “big” an arch is. nabs established the “span” as the common criteria for comparing size. their full definition is quite complex—for simplicity, here is an analogy. the span is the widest dimension through which you could slide one or more giant horizontal playing cards without overlapping or deforming the cards (the cards can step up or down, but not overlap past a vertical line) (see figure 7). the irregular third dimension (depth) of some arches can cause problems, but for most arches these simple criteria work fairly well. in the past, many arches were measured on diagonal or even twisted lines to come up with unrealistic, and non-comparable, dimensions. figure 6. cross-sectional block diagram of the arches national park area drawn approximately to scale. note the location of landscape and double arches. unit (formational) symbols correspond to those shown on figures 2 and 3. after doelling (2010). figure 7. a – delicate arch is the best-known rock arch in the world and is the symbol for arches national park. this free-standing arch is located on the north rim of cache valley. b – delicate arch has a height of 52 feet (16 m). the base and pedestals are slick rock member of the entrada sandstone. the upper part or cap is moab member of the curtis formation. photo courtesy of michael d. vanden berg. t.c. chidsey, jr., and g.c. willis landscape arch, delicate arch, and double arch 9 a few arch statistics (from nabs newsletter “arch wednesday,” february 20, 2019): 1. number of known arches in the world – 9285 (per nabs “significant” criteria above). 2. arches having a greater arch dimension (simplistically, the spanlike line tilted at any angle, even vertical) of 75 feet (23 m) or more – 305 (3.3% of worldwide count). 3. country that has the most documented arches – united states with 8181 (88.1% of worldwide count); france is second. 4. u.s. state with the most arches – utah, with 4589 (56.1% of u.s. count; 49.4% of world count); new mexico is second with 529; nevada is third with 514. 5. using the nabs definition of span, china has the largest known arches in the world (four with spans of 300 to 400 feet [90–120 m])—all massive bridge-type arches formed in limestone in karst topography. landscape arch in arches national park is the fifth largest with a span of 290 feet (88 m); kolob arch in zion national park is sixth at 287 feet (87 m). 6. however, landscape arch is the largest known sandstone arch in the world. utah has the top six largest sandstone arches in the world, and six of the eight have spans over 200 feet (60 m). after landscape and kolob arches, they are: ninth – morning glory natural bridge near moab at 243 feet (74 m); eleventh – rainbow bridge in rainbow bridge national monument at 234 feet (71 m); thirteenth – sipapu natural bridge in natural bridges national monument at 225 feet (69 m); and fourteenth – stevens arch along the escalante river at 220 feet (67 m). 7. and (this is where we get to show our bias), we consider landscape arch to be the most mind-boggling and gravity-defying arch in the world! development of natural rock arches natural rock arches are quite common. what makes utah unique is the gigantic size and number of very large arches. utah has over 60% of the world’s arches; arches national park has over half of utah’s arches. the exceptionally large arches in arches national park, natural bridges national monument, and elsewhere in utah require several criteria that must all come together in just the right way—a very rare occurrence indeed. these include: • very thick, isotropic rock—all of utah’s largest arches are in massive quartz sandstone; the homogenous nature of the sandstone allows for large conchoidal (curved) fractures that are required to produce the inherently stable arch shape. • weaknesses in or below the otherwise isotropic rock, such as bedding planes or joints, that act as “seed” sites for arch formation. • rock that is moderately, but not exceptionally, strong or brittle. • high erosion and incision rates that create many deep canyons and exposed rock faces—big arches need big cliffs, fins, or canyons. • near-vertical, subparallel, properly spaced joints (fractures)—a variety of geologic events have produced abundant joints in arches national park. • an arid climate—whereas small and medium sandstone arches are common in many environments, all of the largest sandstone arches are in areas having dry climates (giant limestone bridges are common in wetter karst environments). • in some settings, loose sand that holds moisture against the base of sandstone fins, allowing the moist basal rock to weather and erode at higher rates than exposed rock, is needed to eventually form an arch (in contrast, rare rain and snow on exposed rock in the arid climate commonly evaporates in minutes to hours). closely spaced joints, subparalleling salt-cored anticlines, in the hard, brittle rocks with associated soft bedding-plane partings, indentations, and thin, soft rock layers favor the formation of arches, fins, and alcoves in arches national park (figure 5). sandstones, typically of the entrada and curtis formations, folded over a broad hinge zone called a “roll over,” collapsed into salt-dissolution valleys or grabens along anticlinal rims (e.g., salt valley anticline). at the “roll over,” the previously formed joints opened and allowed weathering and erosion to occur at a more rapid rate, creating thin fins of sandstone where most wall arches develop. hence, most arches are located along the rims of the salt valleys where the favorable formations begin to “roll over” and collapse into the grabens (figures 5 and 6). weathering within the joints attacks weaknesses such as formational boundaries, bedding planes, and soft partings, forming alcoves and arches in and through the sandstone fins (figure 5). loose sand accumulates between the fins; the sand holds slightly acidic rainwater against the fin walls by surface tension and capillary action. the natural cement holding the sand grains together is largely calcareous and dissolves in the slightly acidic water. the action is favored along more weakly cemented rock found along bedding-plane partings and indentations. eventually, a horizontal crevice is opened through the fin along one of these weak planes. stresses develop in the overlying rock because of gravity since the weight of the rock over the crevice is now supported only by the limbs of the newly formed arch (doelling, 2010). generally, fractures develop and blocks of rock drop from the arch (upward stoping) until the classical arch form is developed. this action is responsible for free-standing and wall arches (as defined by stevens and mccarrick, 1988). free-standing arches form in isolated fins or in thin walls projecting from cliff faces. cliff-wall arches develop short 10 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 distances behind the faces of cliffs. alcoves are large arch-shaped recesses that have not yet eroded through a fin. upward stoping probably continues and eventually the arch collapses (doelling, 2010). larger, more massive arches develop in thicker sandstone fins or walls. large arches also commonly develop where a relatively thin rock wall or fin forms a prominence along a cliff face. the famous three landscape arch (devils garden area) landscape arch in the devils garden area has an incredible span of 290 feet (88 m) (77 feet [23 m] high), ranking fifth in the world (according to the nabs; wilbur, 2004; willis, 2009a, 2012) and is mind-boggling due to its gravity-defying ribbon of rock that in places narrows to only 7 feet (2 m) in thickness (figure 8)! in terms of span, this arch is the largest in arches national park and is the largest known arch formed in sandstone in the world (nabs). landscape arch was named by frank beckwith, leader of the arches national monument scientific expedition, who explored the area in the winter of 1933–1934. the arch is entirely within a fin of the slick rock member of the entrada sandstone (figure 3). it is classified as a free-standing arch, but originally began as a cliff-wall arch. on september 1, 1991, june 5, 1995, and june 21, 1995, blocks of rock 73 feet (22 m), 47 feet (14 m), and 30 feet (9 m) long, respectively, broke and fell from the underside of the span of landscape arch (figure 9) (natural arch and bridge society, 2009). out of safety concerns, the park service closed the trail beneath the arch. this concern is legitimate—nearby wall arch collapsed during the night of august 4, 2009 (willis, 2009b). landscape arch is considered an “arc type” natural arch—an arch that is old and near the end of its life cycle (natural arch and bridge society, 2009). these arches exist, however, because they are also strong, especially when the slender arc of rock is slightly arched as in the case of landscape arch (wilbur, 2007). in addition, each grain or slab that falls actually increases the compression between the remaining grains (natural arch and bridge society, 2007b). the rock falls that occurred in the 1990s indicate that it could soon (years?/hundreds of years?) collapse by natural processes. its demise could be precipitated by an earthquake (doelling, 2010). delicate arch the iconic delicate arch is located on the north side of the cache valley anticline (figures 2 and 7). the trail to delicate arch starts in the salt wash member of the upper jurassic (155 ma) morrison formation and crosses several faults (figures 2 and 3). geologic formations and members exposed along the trail include the reddish to lavender tidwell member of the morrison formation with its large hackly white chert nodules; the red-brown summerville formation; the massive, light-colored, well-jointed moab member of the curtis formation; and the orange-brown, smooth-weathering slick rock member of the entrada sandstone. along the trail, there are excellent examples of deformation bands, jointing, alcoves with springs seeping at their bases denoted by vegetation, well-defined cross-bedding, and tafoni. delicate arch is freestanding and has a horizontal span of about 32 feet (9.7 m) and a vertical span of 46 feet (14 m) as defined by the natural arch and bridge society (2007c) (figure 7). the top of the arch is about 52 feet (16 m) over the base. the moab member of the curtis formation forms the upper part of delicate arch; the slick rock member of the entrada sandstone forms the pedestals. both units are cross-bedded. the plane of weakness in this arch is the contact between the slick rock and the moab members about a third of the way below the top. this plane represents a sharp regional unconformity named the j-3 unconformity (pipiringos and o’sullivan, 1978). cross-beds in the slick rock member are truncated by the unconformity. figure 8. landscape arch, originally a wall arch and now a free-standing arch, is the largest in the park, and the largest sandstone arch in the world. figure 9. dramatic rock fall from landscape arch on september 1, 1991. from interpretive sign, arches national park; photograph taken by royce morrison. t.c. chidsey, jr., and g.c. willis landscape arch, delicate arch, and double arch 11 double arch (garden of eden area) double arch is the third largest in the park, consisting of two giant pothole arch spans that are joined at one end (figure 10). the largest opening of the two arches has a span of 148 feet (45 m) and a height of 104 feet (32 m) (wilbur, 2010). it was the site for the opening scene in the 1989 film indiana jones and the last crusade. whereas most arches in the park are wall arches, double arch is a pothole arch and formed differently from wall arches (figure 11). in this region of the colorado plateau, potholes (also called water pockets) often develop on flat-lying sandstone by a long, slow process—it starts with some small weakness in the rock, maybe only microscopic in size: (a) water is held slightly longer in the weakness, (b) it gradually dissolves the cement around sand grains, (c) the wind and rain blow and wash the loose grains away, (d) over time the tiny depression helps hold the water a little longer allowing a little extra dissolution, (e) several small depressions gradually merge, and (f) organic agents – lichen, algae, and diatoms – aid in the process, loosening more sand grains that are removed by wind and rain. if allowed to continue long enough, potholes can become amazingly large. alcoves also played a role in the formation of these two arches. alcoves form near cliff faces where groundwater percolating down through pores in the sandstone encounters a low-permeability bed that impedes continued downward water flow, in this case the contact between the porous overlying slick rock member of the entrada sandstone and the less-permeable, siltand clay-rich underlying dewey bridge member of the carmel formation (figure 3). because the infiltrating groundwater’s downward flow is restricted, it moves horizontally along a less resistant bedding plane and exits the rock near these low-permeability bedding planes, forming seeps and springs at cliff faces. the water dissolves and removes cementing minerals in the sandstone, causing the rocks there to weaken relative to the rocks elsewhere on the cliff faces. the processes of weathering and erosion proceed faster on the weakened rocks, eventually forming recesses or alcoves. continued weathering and erosion by chemical and mechanical weathering, exfoliation, and rockfalls, caused potholes nearest the cliff faces and alcoves to enlarge vertically and horizontally, respectively, until they merged creating double arch (figure 11). acknowledgments support for this paper was provided by the utah geological survey (ugs). we thank the staff at arches national park for providing access to park lands. cheryl gustin, jen miller, jay hill, and lori steadman of the ugs drafted figures. this paper was carefully reviewed by michael d. vanden berg, stephanie m. carney, michael d. hylland, and bill keach of the ugs, along with the editors of this publication. their suggestions and constructive criticism greatly improved the manuscript. references blakey, r., and ranney, w., 2008, ancient landscapes of the colorado plateau: grand canyon, arizona, grand canyon association, 156 p. doelling, h.h., 2010, geology of arches national park in sprinkel, d.a., chidsey, t.c., jr., and anderson p.b., editors, geology of utah’s parks and monuments (third edition): utah geological association publication 28, p. 11–36. figure 10. double arch, consisting of two giant pothole arch spans joined at one end, is the third largest in the park. photo courtesy of ben erickson. figure 11. pothole arch development stages. 1 – potholes developed on flat-lying sandstone at the same time an alcove is forming along the nearby cliff face. an alcove can be created where groundwater percolating down through pores in sandstone encounters a low-permeability bed that impedes continued downward flow. the groundwater then moves horizontally along the bedding plane to the cliff face where it exits the rock as seeps and springs. along the flow path the groundwater dissolve cementing minerals weakening the sandstone relative to the rock elsewhere on the cliff face. the processes of weathering and erosion on the weakened rock eventually form a recess or alcove. 2 – continued weathering and erosion cause both the pothole nearest the cliff face and the alcove to deepen vertically and horizontally, respectively, until they merge creating an arch. 12 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 hamblin, w.k., 2004, beyond the visible landscape—aerial panoramas of utah’s geology: provo, utah, brigham young university geology, p. 46–47. natural arch and bridge society, 2007a, frequently asked questions: online, https://www.naturalarches.org/archinfo/faq. htm, accessed april 23, 2019. natural arch and bridge society, 2007b, natural arch formation: online, http://www.naturalarches.org/archinfo/formation. htm#compression, accessed march 27, 2019. natural arch and bridge society, 2007c, natural arch dimensions: online, https://www.naturalarches.org/archinfo/dimensions. htm, accessed february 21, 2019. natural arch and bridge society, 2009, landscape arch—big tour 19: online, http://www.naturalarches.org/big9-1.htm, accessed february 25, 2019. pipiringos, g.n., and o’sullivan, r.b., 1978, principal unconformities in triassic and jurassic rocks, western interior united states—a preliminary survey: u.s. geological survey professional paper 1035-a, 29 p. stevens, d.j., and mccarrick, j.e., 1988, the arches of arches national park—a comprehensive study: moab and orem, utah, mainstay publishing, 169 p. wilbur, j.h., 2004, the dimensions of landscape arch—removing the uncertainty: natural arch and bridge society, online, http://www.naturalarches.org/archinfo/landscape.htm, accessed, february 25, 2019. wilbur, j.h., 2007, arch taxonomy—arc natural arch: natural arch and bridge society, online, http://www.naturalarches. org/archinfo/taxonomy-arc.htm, accessed, february 25, 2019. wilbur, j.h., 2010, selected arches by state or country—double arch, nabsqno 12s-626928-4283540: natural arch and bridge society, online, http://www.naturalarches.org/db/ arches/ut115.htm, accessed february 21, 2019. willis, g.c., 2009a, what is the biggest natural arch in the world?: utah geological survey, survey notes, v. 41, no. 2, p. 1–3. willis, g.c., 2009b, geosights, wall arch—a fallen giant: utah geological survey, survey notes, v. 41, no. 2, p. 10. willis, g.c., 2012, every record must fall—an update on the largest arches in the world: utah geological survey, survey notes, v. 44, no. 1, p. 4–5. uga 50:03 lacustrine source rocks and unconventional oil resource plays; a case study from the eocene green river formation of the uinta basin, utah off-canvas toggle member login donate about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle thisistheplace2.jpeg subpage1.jpeg subpage2.jpeg subpage3.jpeg uga 50:03 lacustrine source rocks and unconventional oil resource plays; a case study from the eocene green river formation of the uinta basin, utah david e. schmude axia energy ii llc brian r. berwick axia energy ii llc doi: https://doi.org/10.31711/ugap.v50i.107 abstract much information exists on unconventional oil resource plays, their drivers to success and sweet-spot fairways. to date, most of the successful oil resource plays are marine type ii kerogen driven systems. little to  no information is available on successful oil resource plays with lacustrine type i kerogen source rocks and in rock as young as tertiary age. this paper sets out to define the drivers and provide a case study for such plays  with the green river oil resource play of utah’s uinta basin as an example. the green river petroleum system of the uinta basin has generated vast amounts of hydrocarbons, in the form of gilsonite (~10 billion barrels), tar sands (~12 billion barrels), and oil (>700 million barrels) (johnson, 2016; schamel, 2015). in addition, the usgs estimates an additional 1.32 trillion barrels of in-place resource in the oil shales of the green  river formation. lacustrine source rocks and their type i kerogen follow a different generative path to immiscible oil generation than the more common marine derived type ii kerogens. the oil generative capacity of this  lacustrine source is much greater than that of type ii kerogen and is characterized by a distinctive highly paraffinic crude oil. these unique aspects provide exceptionally elevated formation pressures within the source rock  in the oil generative window, setting it apart from other unconventional oil resource plays. as in all unconventional oil resource plays, identification and mapping of the best source rock is key to defining sweet-spots  within the play. with an understanding of source rock distribution and thermal maturity, the extents of the play  and resource in-place can be defined aerially for the numerous productive intervals of the green river formation. finally, we will address the key attributes of this petroleum system such that it may be used as a model  for other possible lacustrine sourced unconventional oil resource plays. order it ($5.00) cart your cart is empty login form login for discounted rates on document downloads and to access members only content username password show password remember me log in forgot your password? forgot your username? create an account utah geological association p.o. box 526356 salt lake city, ut 84152-0100  facebook  youtube © 2025 utah geological society | site by third sun member login off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf uga-geosite-biek-quail-creek.indd utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors robert f. biek utah geological survey, p.o. box 146100, salt lake city, ut 84114-6100 bobbiek@utah.gov cover image: view north at a brightly colored hillside of shnabkaib (trms), upper red (trmu), and shinarump (trcs) strata just north of utah highway 9, immediately south of quail creek state park. virgin anticline and quail creek reservoir 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e geosites refl ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: biek, r.f., 2019, virgin anticline and quail creek reservoir, in milligan, m., biek, r.f., inkenbrandt, p., and nielsen, p., editors, utah geosites: utah geological association publication 48, 8 p., https://doi.org/10.31711/geosites. v1i1.52. presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president r.f. biek virgin anticline and quail creek reservoir 3 introduction the first thing most visitors to quail creek state park notice, apart from the improbably blue and refreshing waters of the reservoir itself, are the brightly colored, layered rocks of the surrounding cliffs. in fact, quail creek state park lies astride one of the most remarkable geologic features in southwestern utah. the park lies cradled in the eroded core of the virgin anticline, a long upwarp of folded rock that trends northeast through south-central washington county. the fold is breached by erosion along its crest, creating a window into the geologic past (figure 1). famous for its geology, the park is also infamous for the 1989 catastrophic collapse of the quail creek south dike, which unleashed a torrent of water and caused millions of dollars of damage. location one of the best places to see this geologic story is from the parking area at the west abutment of the quail creek south dam (37° 10.77' 113° 23.693'). the park is accessible from the south via n. 5300 w. off state route 9, or from the north via old highway 91, a frontage road parallel to i-15 south of leeds (figure 2). figure 3 shows a simplified geologic map of the park and surrounding area, modified from biek (2003a, 2003b). a kiosk with information about local geology and the problematic south dam is at the north end of the parking lot. a booklet, the geology of quail creek state park, is also available from the utah geological survey (biek, 1999). structure the greater st. george area, including quail creek state park, lies in what geologists call the transition zone between the basin and range and colorado plateau physiographic provinces. the transition zone is characterized by strata and structures common to both the basin and range province and colorado plateau. the brightly colored sedimentary rocks in the transition zone are characteristic of the generally flat-lying rocks of the colorado plateau, but they are locally deformed by both compressional forces associated with the sevier orogeny (mountain-building episode) and more recent extensional forces associated with the formation of the basin and range province. in southwestern utah, the transition zone includes two major down-to-the-west fault zones that step down from the colorado plateau to the basin and range province. the greater st. george area lies on the intermediate structural block thus created, bounded on the east by the hurricane fault zone and on the west by the gunlock-grand wash fault (figure 4). virgin anticline the virgin anticline is a 30-mile-long (50 km), northeast-trending, symmetrical fold that marks the eastern limit of significant sevier-age compressional deformation in southwestern utah. the flanks of the fold provide spectacular exposures of parts of the lower triassic moenkopi formation and overlying upper triassic chinle formation, including hogbacks of the shinarump conglomerate member that dramatically outline the fold’s shape figure 1. view northeast towards quail creek reservoir nestled in the eroded core of the virgin anticline. the flanks of the anticline are neatly outlined by the resistant shinarump conglomerate member of the chinle formation (trcs), below which are ledgy slopes of the upper red member of the moenkopi formation (trmu). the “bacon-striped” shnabkaib member of the moenkopi formation (trms) forms the eroded floor of the anticline. the snow-covered pine valley mountains are on the skyline at left and the kolob canyons part of zion national park is on the skyline at right; state route 9 cuts across the virgin anticline at the bottom of the photograph. photo courtesy of janice hayden. figure 2. location of quail creek state park, washington county, utah. the parking area at the dam’s west abutment offers a good place from which to view the geologic story of the park. from biek (1999). 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 figure 3. geologic map and cross section of the quail creek state park area. the photograph shows the view looking northeast towards quail creek reservoir. from biek (1999). figure 4. schematic block diagram showing the relationship between the hurricane and gunlock-grand wash faults. both faults are “normal” faults that formed during regional extension, allowing rocks on the west side of the faults to slip down relative to rocks on the east side. sidebar: the sevier orogeny, or mountain-building episode, began in utah by middle jurassic time, about 180 million years ago, as deformation associated with subduction of the farallon oceanic plate along the western edge of north america spread eastward into western utah (decelles, 2004). these eastward-directed compressional forces created the sevier orogenic belt and associated broad deformation zone, which consists of, from west to east, a thrust belt with wedge-top basins, foredeep basin, forebulge, and back-bulge basin (decelles, 2004; willis, 1999, 2000) (figure 5). each of these four parts of the thrust system migrated eastward over time, and each created unique environments of deposition or erosion until the end of the orogeny about 40 million years ago. r.f. biek virgin anticline and quail creek reservoir 5 (figure 6). along the northern reaches of the anticline, the carapace of shinarump stands in dramatic relief above the virgin river lowlands (figure 7). the anticline has three structural domes along its length. from south to north these are bloomington dome, washington dome, and harrisburg dome, each of which exposes gypsum-bearing lower permian harrisburg member of the kaibab formation. several west-dipping thrust faults repeat triassic and lower jurassic strata on the northwest flank of the anticline near leeds (see the silver reef mining district geosite for more detail). the fold formed during the height of the sevier orogeny, about 85 to 72 million years ago, above an east-directed blind thrust fault that soles into cambrian strata (davis, 1999; hurlow and biek, 2003). stratigraphy moenkopi formation the moenkopi formation of southwestern utah, with its alternating reddish-brown, white, and gray layers, documents shallow-marine sedimentation along the western margin of the pangean supercontinent, when what is now utah lay close to the equator. the moenkopi consists of three transgressive members (the timpoweap, virgin limestone, and shnabkaib members, which thicken westward and record an interval of sea-level rise), each of which is overlain by an informally named regressive red-bed member (the lower, middle, and upper red members, respectively, which thicken eastward and record sea-level fall and deposition in tidal-flat environments); the rock canyon conglomerate member locally forms the base of the moenkopi formation (reeside and bassler, 1921; stewart and others, 1972a; dubiel, 1994). these members thus record a series of incursions and retreats of a shallow ocean across a gently sloping continental shelf from about 245 to 240 million years ago, where sea-level changes of several feet translated into shoreline changes of many miles (blakey and others, 1993; dubiel, 1994). the moenkopi formation is collectively about 1700 feet thick in southwestern utah, but only the upper two members—the shnabkaib and upper red members—crop out in the park (biek and others, 2009). figure 5. typical parts of a thrust system. the thickened, eastward-moving, leading-edge thrust wedge on the left overloads the earth’s crust, which flexes in response, similar to loading rock on a wooden raft floating on water. in utah, the entire thrust system migrated eastward over time during the middle mesozoic to early tertiary, but this simple pattern is commonly complicated due to variations in crustal strength and pre-existing faults. from willis (1999). figure 6. view north at a brightly colored hillside of shnabkaib (trms), upper red (trmu), and shinarump (trcs) strata just north of utah highway 9, immediately south of quail creek state park. these are the bedrock units seen surrounding quail creek reservoir. note the “bacon-striped” appearance of the shnabkaib member; the yellowish-brown sandstone, locally known as the “purgatory sandstone,” near the base of the upper red member; and the cliff-forming shinarump conglomerate. a veneer of talus locally conceals bedrock. from biek (1999). figure 7. view north to the resistant shinarump conglomerate, which forms a whale-like carapace along the central part of the anticline northeast of quail creek state park. the hurricane cliffs and kolob canyons part of zion national park are in the distance at right. 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 shnabkaib member the shnabkaib (pronounced “shnab-kibe”) member forms striking redand white-banded slopes and ledges around quail creek reservoir. these rock layers are best seen above the west abutment of the quail creek south dam, and along the north shore of the reservoir. the banding is due to alternating layers of reddish-brown mudstone that weather to slopes, and slightly more resistant white gypsum and dolomite that weather to ridges and low ledges. the member is named for exposures at nearby shinob kibe butte, just southeast of washington, which derives its name from the piute words shinob (great spirit) and kaib (mountain). the shnabkaib member contains about 30 percent gypsum as white nodules and layers up to several feet thick. the gypsum was deposited in a coastal-sabkha environment, a broad, nearly level area subject to high evaporation rates and periodic inundation by warm, shallow marine water. gypsum is also present as coarser grained varieties known as selenite and satin spar that occur in cross-cutting veins. the shnabkaib member is about 600 feet (180 m) thick at quail creek state park, but thickens westward across the utah hingeline, the former continental shelf of southwestern utah. its abundant gypsum, mudstone, and microbiolites (algal mounds) indicate deposition in a restricted-marine and coastal-sabkha environment; the member contains limestone in the beaver dam mountains, indicating more open-marine conditions existed to the west (stewart and others, 1972a). upper red member the upper red member forms steep, ledgy slopes above the shnabkaib member and below the cliff-forming shinarump conglomerate. the upper red member consists of reddish-brown mudstone, siltstone, and sandstone. the base of the upper red member is marked by a prominent, cliff-forming, 100-foot-thick (30 m) yellowish-brown sandstone known locally as the “purgatory sandstone” (biek and others, 2009). the upper red member is about 400 feet (120 m) thick, but thins to the west, away from the former continental margin. common ripple marks and other features indicate that it was deposited in tidal-flat and coastal-plain environments (stewart and others, 1972a). chinle formation in southwest utah, the tr-3 regional unconformity (a pronounced gap in geologic time of about 10 million years; pipiringos and o’sullivan, 1978) separates lower triassic (moenkopi formation) and upper triassic (chinle formation) rocks and marks a change from mostly shallow-marine to continental sedimentation. this change occurred as oceanic island arcs continued to collide and accrete to the western pangean margin in what is now california and southwestern arizona during the sonoman orogeny. these collisions deformed the western continental margin, forming a large chain of mountains and volcanoes. in utah, by the late triassic, the environment changed from a passive continental margin to a large interior basin drained by northand northwest-flowing rivers. volcanic mud in rivers and airborne ash from the western volcanoes flowed and drifted east to the interior basin where it was deposited as the colorful layers of the chinle formation. the chinle formation of southwestern utah has long been divided into the lower shinarump conglomerate member and the upper petrified forest member (stewart and others, 1972b), but petrified forest strata may in fact contain beds better assigned to both younger and older chinle formation (kirkland and others, 2014). shinarump conglomerate member because of its resistance to erosion, the shinarump conglomerate forms cliffs along the central portion of the virgin anticline, including the cliffs around the west, north, and east sides of quail creek state park. the shinarump conglomerate consists of yellowish-brown sandstone and pebbly sandstone deposited by generally north-flowing braided streams (figure 8) (stewart and others, 1972b). the streams were probably similar to the modern platte river and other rivers that drain eastward from the rocky mountains and that consist of shallow, interconnected or braided channels and intervening gravel bars. when shinarump sediments were deposited in the late triassic, about 220 million years ago, this area was a gently sloping basin with mountains to the east and west. the small pebbles found in the shinarump conglomerate are mostly chert, quartzite, and quartz. petrified wood is locally common in coarser, pebbly beds; the logs and limbs are remnants of trees deposited by floods. much of the shinarump, however, lacks pebbles and instead consists of sandstone. these sandstones are figure 8. view west at a small channel (in utah highway 9 road cut just above the car) at the base of the shinarump conglomerate. channels such as these are evidence of an unconformity—a surface denoting a gap in geologic time— between the upper red member of the moenkopi formation and overlying shinarump conglomerate member of the chinle formation. the channel shows that the area was exposed and erosion stripped off layers of sediment. this unconformity marks a pronounced change in depositional environments of these two rock units. upper red strata were deposited principally in tidal flats whereas shinarump strata were deposited by braided streams. r.f. biek virgin anticline and quail creek reservoir 7 commonly stained dark brown or black by iron-manganese oxides and locally form “picture stone” or “landscape stone.” shinarump strata are about 100 feet (30 m) thick at quail creek state park. petrified forest member petrified forest strata are widely exposed along the flanks of the virgin anticline just outside the park boundaries. as might be surmised from its name, the member contains petrified wood; it also contains brightly colored clays—locally known as “blue clay”—that swell when wet and shrink when dry. these swelling clays are responsible for numerous building and road foundation problems in the area. the petrified forest member was deposited in floodplains, lakes, and meandering stream channels in a low, forested basin (stewart and others, 1972b; dubiel, 1994). amphibians, reptiles (including the crocodile-like phytosaur), freshwater clams, snails, ostracods, and fish made their home on this once vast, coastal lowland, and petrified conifer trees are common in chinle strata; fossil cycads, ferns, and horsetails are also known, as are dinosaur tracks (stewart and others, 1972b; blakey and others, 1993; dubiel, 1994; decourten, 1998). catastrophic failure of quail creek south dam although quail creek state park is located in the middle of a truly exceptional geologic setting, the park is best known for its fishing and boating. trout, bass, crappie, bullhead catfish, and bluegill attract anglers from throughout the west, and its warm waters make for a swimmer’s paradise. the reservoir itself is operated by the washington county water conservancy district for storage of virgin river water. at its maximum pool elevation of 2985 feet (911 m), the lake covers 640 acres (one square mile or 2.6 km2) and stores 40,325 acre-feet (nearly 50,000,000 m3) of water. most of the water is diverted from the virgin river east of hurricane and is piped about 8 miles (13 km) to the reservoir; this is done to avoid the salty water discharge at pah tempe hot springs between hurricane and la verkin. the annual yield of the reservoir is about 20,000 acre-feet (20,000,000 m3), an important part of the st. george basin water supply. at 12:30 a.m. on january 1, 1989, the quail creek south dike collapsed catastrophically and unleashed a torrent of water, causing millions of dollars of damage. fortunately, the downstream area was evacuated in time to avoid fatalities. approximately 25,000 acre-feet (31,000,000 m3) of water—more than half of the reservoir’s capacity—flowed through a breach in the dike over a 12-hour period (figure 9). the original quail creek south dike was a 78-foot-high, 2000-foot-long (24 m by 600 m) earthen dam constructed in 1984. the dam was poorly designed and seepage under the dam occurred immediately after filling the reservoir. this seepage eroded the dam and foundation materials over the years despite efforts to seal the leaks. ultimately, seepage and erosion of the dam and foundation materials accelerated until caving occurred on a developing opening in the dike. frantic, last-minute efforts to stem the seepage were unsuccessful and the dike finally breached. after the dike’s collapse, investigators found that participation of an engineering geologist was limited during the exploration, design, construction, and operation of the dike (o’neill and gourley, 1991; gourley, 1992). they further found that the dike’s failure was principally due to poor foundation design and construction. simply put, the dike was built mostly on the highly jointed, gypsum-bearing shnabkaib member of the moenkopi formation (except the southeast abutment which lies on sandstone and siltstone of the upper red member). the joints allowed water to rapidly infiltrate bedrock under the dike, bringing it in contact with gypsum, which slowly dissolves in water. the new dike, called the quail creek south dam, was completed in 1990 as a roller-compacted concrete gravity dam. the dimensions of the dam are basically the same as the old dike, except that it now includes a new impermeable cutoff trench as much as 75 feet (23 m) deep, which is designed to prevent water from seeping under the dam (figure 10). evidence of the flood is still visible downstream from the new quail creek south dam, where the shnabkaib member was scoured clean of overlying loose bedrock and sediment. figure 9. quail creek south dike, shortly after its catastrophic failure on january 1, 1989. note how floodwaters scoured the bedrock clean of loose, overlying sediments just below the dike. photo by ben everitt, utah division of water resources (retired). 8 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 acknowledgments my knowledge of southwestern utah geology comes from years of geologic mapping supported largely by the utah geological survey and u.s. geological survey. as the acknowledgments of those many published geologic maps attest, i am indebted to a great many people for their help over the years. thanks to grant willis, mike hylland, and bill keach (ugs) and kimm harty (ugs retired) for their insightful reviews. references biek, r.f., 1999, the geology of quail creek state park: utah geological survey public information series 6, 21 p. biek, r.f., 2003a, geologic map of the hurricane quadrangle, washington county, utah: utah geological survey map 187, 61 p., 2 pls., scale 1:24,000. biek, r.f., 2003b, geologic map of the harrisburg junction quadrangle, washington county, utah: utah geological survey map 191, 42 p., 2 pls., scale 1:24,000. biek, r.f., rowley, p.d., hayden, j.m., hacker, d.b., willis, g.c., hintze, l.f., anderson, r.e., and brown, k.d., 2009, geologic map of the st. george and east part of the clover mountains 30’ x 60’ quadrangles, washington and iron counties, utah: utah geological survey map 242, 101 p., 2 pls., scale 1:100,000. blakey, r.c., bashem, e.l., and cook, m.j., 1993, early and middle triassic paleogeography, colorado plateau and vicinity, in morales, m., editor, aspects of mesozoic geology and paleontology of the colorado plateau: museum of northern arizona bulletin 59, p. 13–26. decelles, p.g., 2004, late jurassic to eocene evolution of the cordilleran thrust belt and foreland basin system, western u.s.a.: american journal of science, v. 304, p. 105–168. dubiel, r.f., 1994, triassic deposystems, paleogeography, and paleoclimate of the western interior, in caputo, m.v., peterson, j.a., and franczyk, k.j., editors, mesozoic systems of the rocky mountain region, usa: rocky mountain section of society of economic paleontologists and mineralogists, p. 133–168. decourten, f., 1998, dinosaurs of utah: salt lake city, university of utah press, 300 p. gourley, c., 1992, geological aspects of the quail creek dike failure, in harty, k.m., editor, engineering and environmental geology of southwestern utah: utah geological association publication 21, p. 17–38. hurlow, h.a., and biek, r.f., 2003, geologic map of the pintura quadrangle, washington county, utah: utah geological survey map 196, 20 p., 2 plates, scale 1:24,000. kirkland, j.i., martz, j.w., deblieux, d.d., santucci, v.l., madsen, s.k., wood, j.r., and payne, n.m., 2014, paleontological resource inventory & monitoring, chinle and cedar mountain formations, capitol reef national park, utah: utah geological survey contract report to the national park service prepared under cooperative agreement #p13ac00601 task #p13ac01248, 128 p. plus appendices. o’neill, a.l., and gourley, c., 1991, geologic perspectives and cause of the quail creek dike failure: bulletin of the association of engineering geologists, v. 28, no. 2, p. 127–145. pipiringos, g.n., and o’sullivan, r.b., 1978, principal unconformities in triassic and jurassic rocks, western interior united states—a preliminary survey: u.s. geological survey professional paper 1035-a, 29 p. reeside, j.b., jr., and bassler, h., 1921, stratigraphic sections in southwestern utah and northwestern arizona: u.s. geological survey professional paper 129-d, p. 53–77. stewart, j.h., poole, f.g., and wilson, r.f., 1972a, stratigraphy and origin of the triassic moenkopi formation and related strata in the colorado plateau region, with a section on sedimentary petrology by r.a. cadigan: u.s. geological survey professional paper 691, 195 p. stewart, j.h., poole, f.g., and wilson, r.f., 1972b, stratigraphy and origin of the chinle formation and related upper triassic strata in the colorado plateau region, with a section on sedimentary petrology by r.a. cadigan and on conglomerate studies by w. thordarson, h.f. albee, and j.h. stewart: u.s. geological survey professional paper 690, 336 p. willis, g.c., 1999, the utah thrust system—an overview, in spangler, l.e., and allen, c.j., editors, geology of northern utah and vicinity: utah geological association publication 27, p. 1–9. willis, g.c., 2000, utah’s sevier thrust system: utah geological survey, survey notes, v. 32, no. 1, p. 1–4. figure 10. cutoff trench being excavated at the bottom of the new quail creek south dam in january 1990. view east. note seepage of water on left-side of trench. photo by bill lund, utah geological survey (retired). shelley judge1, meagen pollock1, michael williams1, krysden schantz2, kelli baxstrom3, kyle burden4, cam matesich5, emily randall1, samuel patzkowsky6, whitney sims1, william cary1, matt peppers7, addison thompson8, thomas wilch9 1the college of wooster, department of earth sciences, 944 college mall, wooster, ohio 44691; sjudge@wooster.edu 2cape fear public utility authority, 235 government center dr., wilmington, north carolina 28403 3usgs landslide hazards, 1711 illinois st., golden, colorado 80401 4north carolina department of transportation, 300 division dr., wilmington, north carolina 28401 5bethlehem-center high school, 179 crawford road, fredericktown, pennsylvania 15333 6franklin and marshall college, department of earth and environment, p.o. box 3003, lancaster, pennsylvania 17604 7chesapeake energy, 6100 n western ave., oklahoma city, oklahoma 73118 8pomona college, department of geology, 1050 n. mills ave., claremont, california 91711 9albion college, department of geology, 611 e. porter st., albion, michigan 49224 cover image: aerial photo of pocket, crescent, miter and terrace craters. view to the east (from google earth). young volcanism in millard county: ice springs volcanic field utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. the geosites reflect the interests of the many volunteers who wrote to share some of their favorite geologic sites. the list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. the utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. this is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and figures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. presidents message i have had the pleasure of working with many different geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several different geological provinces. we have the basin and range to the west and the central utah hingeline and thrust belt down the middle. the uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava flows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. the general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. the “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. the articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. this guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, thank you to the american association of petroleum geologists, rocky mountain section foundation for their financial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fill the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president judge, s., pollock, m., williams, m., schantz, k, baxstrom, k., burden, k., matesich, c., randall, e., patzkowsky, s., sims, w., cary, w., peppers, m., thompson, a., and wilch, t., 2019, young volcanism in millard county: ice springs volcanic field, in milligan, m., biek, r.f., inkenbrandt, p., and nielsen, p., editors, utah geosites: utah geological association publication 48, 5 p., https://doi.org/10.31711/ugap.v1i1.88. ice springs volcanic field 3 judge, pollock, williams, schantz, baxstrom, burden, matesich, randall, patzkowsky, sims, cary, peppers, thompson, wilch introduction geoscientists, naturalists, and rock-hound enthusiasts have explored ice springs volcanic field (isvf) for nearly 130 years because it is one of the youngest, extension-related volcanic centers in utah and the southwest u.s. ice springs received its name due to the presence of ice within its expansive a’a lava flows (davis, 2014), which have an interesting age discrepancy. previous work on ice springs dated the isvf as 660 years old (valastro and others, 1972), but we now introduce an age date of 9,800 to 11,100 years old. although the eruptive and effusive deposits (more explosive vs passive, respectively) capture our imagination today, it is possible that they were natural hazards for the early paleo-indians of the region. the isvf is an important economic resource. hoover (1974) reported that mining has occurred here since the mid-1950s. in the mid-1970s, red dome, inc. began its crushed and broken stone mining and quarrying operations at isvf. these operations continue today and supply decorative stone and road gravel to the region and beyond. from a geologic perspective, isvf is part of a zone of young volcanism—dated as less than 2 million years old and extending from northwest arizona to west-central utah (e.g., best and brimhall, 1974). this area in west-central utah is known as the black rock desert and is part of the “sevier thermal area” (mabey and budding, 1987, 1994) that contains power plants, hot springs, therapeutic baths, and tropical fish ponds (blackett and wakefield, 2002). previous workers have been instrumental in our current understanding of isvf. gilbert (1890) first described isvf in his united states geological survey (usgs) monograph on prehistoric lake bonneville. more recent work has concentrated on the age relationships, study of volcanic deposits, and geochemical trends (i.e., changes in rock chemistry) between the various basaltic lava fields of the black rock desert (condie and barsky, 1972; hoover, 1974; lynch and nash, 1980; johnsen and others, 2010). mapping for both the delta and richfield 30’ x 60’ quadrangles (hintze and davis, 2002; hintze and others, 2003) culminated in a utah geologic survey bulletin on the geology of millard county (hintze and davis, 2003). detailed geochemical studies of isvf by thompson (2009) and patzkowsky and others (2017) have improved lava field mapping. additional work at the college of wooster continues on the age and details of lava field features at isvf. these deposits of the isvf record a history of post-lake bonneville eruptive events (oviatt, 2015). this paper aims to (1) provide the physical location and geologic setting for isvf, (2) outline the unique geologic characteristics of the crater complex and associated lava fields, and (3) summarize recent research at isvf regarding lava field structure and form, age of the flows, and eruptive sequence of events. many visitors each year explore the crater remnants and their eruptive deposits and then wander among the a’a and pahoehoe lava, investigating one of the many lava tubes within the ice springs complex. it is our hope that visitors to the lava field will enjoy the wonders of utah’s youngest volcanic feature. location the isvf is located in millard county, west-central utah. it comprises one of the many volcanic fields in utah’s black rock desert, east of the cricket mountains and west of the pahvant range. the isvf is a 7.7 mi2 (20 km2) volcanic area consisting of a crater complex and basaltic lava flows that erupted onto ancient lake bonneville deposits after the lake receded from its provo-level shoreline. the volcanic crater complex and lava field are readily seen from a distance, standing in stark contrast to the agricultural fields of the valley floor. the entrance road to mining operations is ~8.7 mi (14 km) west of main street in fillmore, utah (figure 1a). although the lava fields extend outward from the several craters in all directions, the approximate center of the complex—informally referred to as “the cinders”—is at 38°57.740'n, 112°30.348'w. isvf is near two other volcanic centers: immediately to the south is the tabernacle hill lava fields and nearly 12 miles (20 km) to the northwest is pahvant butte. figure 1. a. location of ice springs volcanic field (black area at top left) in relation to tabernacle hill to the south and fillmore, utah, to the east (modified from google earth). inset map of utah counties shows the isvf location in millard county (from utah agrc). b. aerial photo of the southeast side of isvf cinder cone complex. arrows indicate possible driving routes to avoid mining operations. the “p” symbols mark suggested parking areas (modified from google earth). ice springs volcanic field 4 judge, pollock, williams, schantz, baxstrom, burden, matesich, randall, patzkowsky, sims, cary, peppers, thompson, wilch suggested driving directions the isvf is easily accessible from the north or south via i-15. from the north, take i-15 south to exit 167 (fillmore). proceed south on cedar mountain road, which becomes n. main street and continues into fillmore. once in fillmore, proceed west on 200 south toward isvf. when driving from the south, take i-15 north to exit 163 (fillmore). proceed north on state route 99, which becomes 500 south and curves to s. main street into fillmore. proceed west on 200 south toward isvf. once the edge of the lava flows is reached, the safest way to access the geology is from the south or west sides of the crater complex, and these directions follow. while on 200 south, there is a “y” junction (38°57.869'n, 112°29.088'w); proceed southwest (left fork) on s 6400 west for ~0.5 mi (1 km). at the 2nd “y” junction (38°57.552'n, 112°29.697'w), proceed west (right fork) on an unmarked unimproved dirt road. from here, you can either proceed west to go to the lava fields (38°57.458'n, 112°30.284'w) or north to the crater complex center (38°57.467'n, 112°30.263'w) at the second junction to parking areas (figure 1b). please remember: if visiting the crater complex and its nearest lava fields, please obtain permission from red dome, inc., at the office at the entrance to the quarry. all visitors, even those accessing remote lava fields adjacent to major roads, should please watch for traffic and stay clear of all mining operations. structural geology isvf is located in the thermally active black rock desert near the eastern border of the basin and range physiographic province. the basin and range is known for its north-trending alternating ranges (horsts) and valleys (grabens) that formed as a result of extensional tectonics. specifically, extension of earth’s crust in the region of the black rock desert since 10 ma has been orientated wnw-ese (christensen and yeats, 1992; nelson and tingey, 1997). structurally, the black rock desert represents an intra-graben area bounded to both the west and east by large normal faults of the cricket mountain and the pahvant range, respectively (hoover, 1974; hintze and davis, 2003). these north-south striking faults resulted from this basin and range extensional regime. fault movement facilitated migration of magma to the surface, creating the many volcanic fields that give the black rock desert its name. numerous other smaller faults and fractures within the black rock desert resulted from the same tensional stress (hintze and davis, 2002; hintze and others, 2003). geologic history volcanic fields the black rock desert was first described by gilbert (1890) and later by condie and barsky (1972) and hoover (1974). originally, condie and barsky (1972) identified seven distinct volcanic fields using age, geography, and composition; from north to south, these fields are deseret, pahvant butte, ice springs, tabernacle hill, kanosh, black rock, and cove fort (figure 2). the divisions of the black rock desert were redefined due to more detailed mapping (hoover, 1974) and subsequently by grouping into fewer volcanic subfields (johnsen and others, 2010). the volcanic fields of the black rock desert are generally pleistocene in age and are fault-controlled (hintze and davis, 2003), inferring that the volcanism was emplaced along or near to the north-south striking faults in the region. the composition of black rock desert deposits vary from basalt to andesite, with even some dacite to rhyolite (best and others, 1980). previous interpretations by hintz (2008) suggested that the fields are monogenetic volcanoes, implying small volume eruptions fed from one or more magma batches (németh and kereszturi, 2015). figure 2. generalized geologic map of the primary lava fields of the black rock desert. the seven major fields of condie and barskey (1972) are identified in italics, along with the five subfields of johnsen and others (2010). isvf is the bright red region in the center of the map (modified from johnsen and others, 2010). ice springs volcanic field 5 judge, pollock, williams, schantz, baxstrom, burden, matesich, randall, patzkowsky, sims, cary, peppers, thompson, wilch lake bonneville correlating the timing of eruptive episodes to the extent of lake bonneville is one of the more important geologic stories of the black rock desert. hoover (1974) classified the eruptions of the black rock desert into three episodes centered around the presence or absence of the lake. episode 1 consisted of eruptions before the lake’s rise, including those in the beaver ridge subfield (johnsen and others, 2010). the lake eroded (in some cases severely) some of these cinder cones in the region that existed prior to its inception. episode 2 included two more eruptions above lake level in the ice springs subfield (johnsen et al, 2010), specifically at pahvant butte, as well as underwater eruptions into lake bonneville shortly afterwards at pahvant butte (white, 1996; white 2000) and tabernacle hill (hintz, 2008). episode 3 includes eruptions of the tabernacle and ice springs volcanic fields after the lake receded. these youngest volcanic deposits lack any lake bonneville sediments on or terraces cut into them, and this lack of bonneville data demonstrates that its volcanism occurred after the lake receded from the area. this final episode represents the youngest volcanism in the black rock desert (hoover, 1974). recent work identifies lake bonneville’s lake level oscillations over a 17,000-year interval during the late pleistocene (oviatt, 2015; oviatt, 2016). formed ~30,000 years ago, lake levels fluctuated markedly but overall rose until ~18,000 years ago, the maximum extent of lake bonneville in the region (oviatt, 2015). known as the bonneville shoreline, this maximum areal extent produced observable shoreline features such as beach gravel barriers and wave-cut notches and platforms throughout utah (chen and maloof, 2017). at ~18,000 years ago, lake bonneville breached and catastrophically drained near red rock pass, idaho, causing a precipitous drop in lake levels to begin formation of the provo shoreline (oviatt, 2015). the provo shoreline ended ~15-16,700 years ago (miller and others, 2013; oviatt, 2015), with another drop in lake levels. oviatt (2015) infers that lake bonneville levels dropped to those matching the great salt lake by 13,000 years ago. this timeline suggests that lake bonneville had receded from the vicinity of isvf well before 9,800 to 11,100 years ago—our age date for its lava field. age of volcanism the age of isvf has long been intriguing because it is the youngest volcanic field in utah and one of the youngest in the southwest u.s. only one age has been reported for isvf (valastro and others, 1972), based on radiocarbon dating of a piece of root fragment found underneath a lava flow from the northeast portion of the lava field. valastro and others (1972) dated the isvf as 660 ± 170 years old, and this age has been repeated for decades even though hoover (1974) estimated the maximum age of isvf as 1000-4000 years old based on the fact that it has less soil cover and erosion compared to nearby tabernacle hill. we recently redated the isvf using cosmogenic 36cl-dating, which provides an estimate for the length of time a rock has been exposed at the surface (patzkowsky and others, 2017). five samples of pahoehoe (lava with a smooth ropy texture) were collected from isvf lava flows, yielding an age estimate of approximately 10,000 years old (table 1). the lava flows appear to be older than previously estimated. we noticed that the surfaces of the pahoehoe samples are no longer iridescent and glassy. when a lava flow erupts, a layer of volcanic glass forms on the surface. the iridescent glass cracks as the lava cools and is easily eroded by mechanical weathering. for comparison, the mccarty’s flow in the zuni-bandera volcanic field, which is in a similarly arid climate, has been dated at ~4000 years old and still has some of its glassy, iridescent surface (dunbar, 1999). furthermore, large areas of the isvf lava flows are covered with a thin layer of soil and meter-high sagebrush, suggesting the lava flows are older than previously suspected. overall, the black rock desert captures the fascinating interactions between volcanic eruptions and changing shorelines. the isvf 36cl ages are older than anticipated, but they still adhere to the established timeline of black rock desert volcanic events. because there are no lake terraces cut into the isvf, we know that isvf was erupted after lake bonneville’s provo-level shoreline receded ~13,000 years ago (milligan and mcdonald, 2017). our age estimates also show that the isvf is younger than nearby tabernacle hill, which has been carved by lake bonneville and has an absolute 14c age of 17,440 (±120) years (cerling and craig, 1994; oviatt, 1991). isvf is clearly younger than other black rock desert volcanic fields, but older than we thought. it was a short-lived, complex eruption that produced volcanic features worthy of exploration. stratigraphy cinder cone stratigraphy the center of isvf is a complex of originally four distinct cinder cones (figure 3a): crescent, miter, pocket, and terrace (hoover, 1974; lynch and nash, 1980). crescent, miter, and terrace were named by gilbert (1890), and pocket was named by hoover (1974). crescent and terrace are the largest and least well-preserved, each with large sections of their original cinder cones removed by eruptive activity or decades of mining. although pocket is the smallest and most completely preserved (i.e., least affected by mining), miter best reveals the geologic history of the crater complex. while crescent may have once been the largest crater (diameter estimate: 2600 feet; 800 m), less than half of the original crater remains because of mining on the eastern flank since the mid-1950s (hoover, 1974). crescent’s exposed eastern crater flank contains red and black beds of cinder up to 50 ft (15.2 m) thick, sloping ice springs volcanic field 6 judge, pollock, williams, schantz, baxstrom, burden, matesich, randall, patzkowsky, sims, cary, peppers, thompson, wilch table 1. cosmogenic 36cl data and dates for the isvf. major elements were measured at the college of wooster by xrf following the methods of pollock and others (2014) and are reported in weight percent (wt.%). trace elements were measured at actlabs by icp-ms and are reported in parts per million (ppm). 36-chlorine was measured at purdue prime lab by ams and is reported in atoms per gram of sample. one sample was analyzed in duplicate to assess analytical reproducibility. the age calculation in the cronus 36cl exposure age calculator (v2.0; marrero and others, 2016) uses the following values: an average measured bulk density of 2.24 g/cm3, 0 ppm co2, and the lifton-sato-dunai (sf) scaling framework (lifton and others, 2014). assuming a realistic erosion rate of 1 mm/kyr, our age estimates range from 9.8 (±1.6) 11.1 (±1.7) kyr. these ages were older than expected, so we reanalyzed the samples to make sure we removed all the meteoric 36cl and found the same results. even using high erosion rates (up to 5 mm/kyr) to estimate minimum ages, we do not calculate dates any younger than 9.2 (±1.7) 10.9 (±1.6) kyr. sample id keckis1701 keckis1707 ks15cd02a ks15cd02b ks15cd05 ks15cd05 (duplicate) latitude on 38.968 38.9597 38.959518 38.959518 38.957569 38.957569 longitude ow 112.504 112.512 112.515 112.515 112.51 112.51 elevation (ft) 4757.2 4757.2 4732.0 4732.0 4752.0 4752.0 sample thickness (cm) 4 4 9 9 10 10 36cl (atoms per gram sample) 386549 378680 334589 279622 287599 278045 sio2 (wt.%) 49.61 49.53 50.53 49.42 49.23 48.67 tio2 1.87 1.9 1.77 1.82 1.81 1.82 al2o3 14.41 14.58 14.45 15.06 15.51 15.11 fe2o3 11.45 11.68 11.04 11.34 11.48 11.55 mno 0.19 0.18 0.18 0.18 0.19 0.18 mgo 6.83 6.98 6.77 6.75 7.52 7.66 cao 11.08 10.64 11.03 11.21 10.12 11.06 na2o 2.71 2.72 2.39 2.47 2.53 2.45 k2o 1.20 1.13 1.22 1.12 1.00 0.89 p2o5 0.58 0.58 0.57 0.58 0.55 0.55 loi 2.29 1.29 3.68 3.24 1.47 2.22 cl (ppm) 251 203 144.9 122.6 112.4 100.4 b 1 1 0 0 0 0 sm 7.7 7.4 6.6 6.8 7 6.6 gd 7.8 7.9 6.3 6.6 7 6.5 u 1.1 0.9 1.7 1.3 1.1 1.1 th 3 2.5 4.9 3.6 3.7 2.8 cr 155 137 103 95.5 172 160 li 12.8 11.5 20.6 15.2 19.3 12.8 age (kyr at 1 mm/kyr erosion rate) 9.8 (±1.6) 11.1 (±1.7) 10.5 (±1.7) 9.6 (±1.4) 11.1 (±1.5) 11 (±1.3) age (kyr at 5 mm/kyr erosion rate) 9.2 (±1.7) 10.6 (±1.6) 10.2 (±1.6) 9.4 (±1.4) 10.9 (±1.6) 10.9 (±1.5) ice springs volcanic field 7 judge, pollock, williams, schantz, baxstrom, burden, matesich, randall, patzkowsky, sims, cary, peppers, thompson, wilch away from the rim at 25°-30°, capped by scoriaceous spatter (formed from frothy lava fragments; figure 3b). the scoriaceous spatter marks the evolution of crescent from a cinder cone into a spatter cone, commonly observed at ice springs. on the inner slope of crescent, welded spatter and cinders slope at 45°-50° (hoover, 1974). crescent was the source for the most northern flows in the isvf, although the lava flow outlet was obscured when the entire western side of crescent was destroyed by the formation of slightly younger cinder/spatter cones, miter and terrace, and their respective lava flows (gilbert, 1890; hoover, 1974). our research focused on miter’s volcanic history due to its better preservation when compared to crescent or terrace. miter (diameter: 900 feet; 275 m) is dominated by cinder deposits and capped by a resistant layer of welded spatter. due to its physical characteristics, we infer a strombolian eruption style, which is common for basaltic eruptions producing cinder cones. miter reveals a 27.3 ft (8.3 m) eruptive sequence at its crest (figure 3c). from bottom to top, three units were identified within the sequence: (1) a lower unit of poorly welded spatter and large bombs, (2) a middle unit of increased welded spatter but small bombs, and (3) a highly weathered top unit of scattered bombs and blocks, welded spatter, and ash lapilli. these bombs are poorly to highly vesicular (classification of houghton and wilson, 1989), with greater vesicularity typically in the rind than the core of the bombs. further, many bombs are fluidally shaped, suggesting that they were partially molten upon impact (figure 3d). pocket is a small (diameter: 330 feet; 100 m), symmetrical, cinder cone that formed after crescent (figure 3a). compared to the others in the complex, pocket produced only a small volume of lava (lynch and nash, 1980). terrace (diameter: 800 feet; 245 m) is relatively low-lying and built on the southern side of the crescent-miter complex. it is dominated by layered spatter and a’a flows (lava with a blocky, jagged texture) underneath a veneer of ejected rock fragments (tephra) (hoover, 1974). terrace is the source for the southeastward flows in the isvf. gilbert (1890) gave the terrace its name based on speculation that a molten lake left behind terraces, although mining activities have destroyed most evidence of this. lava field stratigraphy lava flow fields gilbert’s (1890) field observations are especially important for understanding the eruptive history of isvf because they preceded mining activity. he identified lava channels emanating from the western edge of the crater complex that connected the cones to specific lava flows. the lava channels are partially disrupted by mining, but building on gilbert’s (1890) observations, hoover (1974) provided a detailed summary of eruptive events (figure 4). the eruption began with lava flows from crescent advancing primarily to the north. the crescent lava flows were followed by eruptions from miter and pocket, producing lava flows that were initially channeled to the north then to the west and southwest. toward the end of the miter eruptions, lava flows from terrace advance to the south. ice springs lava flows are dominated by a’a lava except near the crater complex, where pahoehoe can be identified by its smooth, ropy surfaces. in hand sample, the lavas are very fine-grained basalt with few to no visible crystals, although green olivine crystals up to four millimeters in diameter can be found in the a’a flow west (38°57.905'n, 112°31.143'w) of the crater complex. granitic rock fragments up to about an inch (several centimeters) across are sparsely distributed in the lava flows; and, light-colored, glassy, highly vesiculated fragments (xenoliths) are found in the scoria at the crater complex (condie and barsky, 1972; hoover, 1974; thompson, 2009). although isvf lava flows appear similar to the eye, they can be separated into two groups based on their geochemical compositions (figure 4). the northern field, which formed during crescent figure 3. a. aerial photo of pocket, crescent, miter and terrace craters. view to the east (from google earth). b. the east flank of crescent showing volcanic stratigraphy with its red and black cinders. view to the west. c. volcanic stratigraphy at the eastern rim of miter crater. d. close-up of c, illustrating the vesicularity (small cavities) of the volcanic bombs. ice springs volcanic field 8 judge, pollock, williams, schantz, baxstrom, burden, matesich, randall, patzkowsky, sims, cary, peppers, thompson, wilch figure 5. lava flow features. a. boundary between a’a lava flow (left) and pahoehoe lava flow (right; 38°57.974'n, 112°31.119'w). b. ropy pahoehoe surface adjacent to a’a lava (38°57.815'n, 112°30.777'w). c. a pressure ridge (tumulus) in lava channel with a smooth floor; students on lava channel walls (38°57.590'n, 112°30.873'w). d. striae on tumulus. e. distal cinder mound in an area of rugged volcanic topography (38°57.968'n, 112°30.770'w). f. breached northern wall of miter cone (38°57.807'n, 112°30.408'w). and early miter eruptions, comprises lava that is relatively higher in silica (sio2) compared to the southern and western fields, which formed during late miter and terrace eruptions (condie and barsky, 1972; lynch and nash, 1980; thompson, 2009). the high silica lavas are also consistently higher in potassium (k2o) and lower in iron (feo*) and titanium (ti o2). we used the geochemical “fingerprints” of the lava flows to better understand the eruption history at isvf. our sampling focused on the boundaries between lava flow fields, the area immediately west of the cones where most of the flows originate, and the interior of the lava flows, which are less accessible and represented a gap in the existing data set (table 2). our work nearly doubles the number of geochemical data points for isvf lava flows. overall, the additional geochemical data overlaps with the high-silica and low-silica compositional groups, but some samples fall into groups that are opposite to what was expected based on previous mapping (figure 4). for example, the olivine-rich a’a lava west of the crater complex was previously mapped as a late miter lava flow, but it falls in the same high-silica compositional group as early miter and crescent flows (sims and others, 2012). additionally, samples in the northern field that were previously mapped as early miter flows fall within the same low-silica compositional group as late miter and terrace lava flows (patzkowsky and others, 2017). our results suggest that the chemistry of isvf is more complex than previously thought. we modified flow boundaries based on our recent geochemistry and mapping of lava field features (figure 4). figure 4. location of crescent (tan), miter (light gray), and terrace (dark gray) lava flows according to lynch and nash (1980). proposed boundaries for the miter lava flow (red dashed line) and terrace lava flow (black solid line) are based on our work. crosshatched area is disturbed by mining activity. inset shows sio2 vs mgo in weight percent for samples mapped as crescent and early miter lava flows (tan circles) and as late miter and terrace lava flows (black squares). some samples mapped as a late miter lavas fall into the high-silica compositional group while some samples mapped as crescent and early miter lavas fall into the low-silica compositional group. data from this study and lynch and nash (1980), nelson and tingey (1997), and thompson (2009). ice springs volcanic field 9 judge, pollock, williams, schantz, baxstrom, burden, matesich, randall, patzkowsky, sims, cary, peppers, thompson, wilch table 2. representative geochemical analyses of samples from isvf lava flows. all samples are from this study except for i-27, which is from nelson and tingey (1997). feo* is total iron calculated as feo. sample id keck17-07 m13-cm-19 m-12-ws-13 i-27 17-01 mapped lava flow terrace miter miter crescent crescent low/high silica group low low high high low sio2 50.36 48.71 52.15 52.01 50.38 mgo 7.43 7.52 7.49 7.44 7.49 feo* 11.06 10.46 8.49 9.34 11.10 tio2 2.00 1.88 1.44 1.66 2.00 k2o 1.14 0.92 1.50 1.45 1.25 na2o 2.86 2.72 3.05 2.97 2.82 al2o3 14.99 15.88 16.06 15.55 14.74 cao 9.30 11.11 9.24 8.99 9.35 p2o5 0.60 0.56 0.37 0.44 0.62 mno 0.18 0.18 0.15 0.15 0.19 bao 0.07 0.06 0.07 0.07 sum 100.00 100.00 100.00 100.00 100.00 general lava field features: visitors to isvf can view breached craters, a’a and pahoehoe, inflated lava features such as pressure ridges (tumuli), lava channels and topographically depressed areas containing a’a and mounds of cinders on the lava flows (figure 5). each of these are discussed briefly below. breached crater walls: pocket, miter, and terrace all show evidence of breached crater walls. the most complete illustration of the four cinder cones prior to mining was published by gilbert (1890; figure 6). in his illustration, gilbert (1890) interpreted breached crater walls from (1) the west side of pocket, (2) the north and west sides of miter, and (3) the west and south sides of terrace (figure 3a, 6a). mining operations have disrupted all the breaches except those on the west side of pocket and the north side of miter, which are clearly visible today. these breached crater walls represent the break-out of lava flows at the base of cinder cones, causing flank collapse (figure 5f). lava morphology and inflated lava features: at isvf, pahoehoe is commonly located near the crater complex, something also noted by condie and barsky (1972) (figure 5b). pahoehoe is widely exposed to the west and north of miter in elevated regions of the flow field associated with inflation. inflation of a sheet flow results when liquid lava is pumped underneath a chilled, solid crust, which then rises (self and others, 1997; gary and others, 2012). at isvf, these pahoehoe flows are evidence of inflation events, which occur during eruptions with steady low rates (bernardi and others, 2015). figure 6. illustrations of isvf, contained in lake bonneville, u.s. geological survey monograph 1, by g.k. gilbert (1890). a. the crater complex at isvf before mining operations. left to right: pocket, crescent (arch remnant in back), miter, and terrace. note the breached crater walls on the west side of both pocket and miter, in addition to the channels in the lava fields extending from each of these breaches. view to the east. b. view of crescent as seen from the top of miter crater. the top of pocket can be seen in the left-center of the illustration. view to the north. ice springs volcanic field 10 judge, pollock, williams, schantz, baxstrom, burden, matesich, randall, patzkowsky, sims, cary, peppers, thompson, wilch if flow rate and slope conditions are favorable, inflated flows can contain lava tubes and surface features such as vertical striae, pressure ridges, and tumuli (figures 5c, 5d; hon and others, 1994; self and others, 1998; anderson and others, 2012). the presence of numerous lava tubes within isvf suggests a complex transport network from the source. internal fluid in the tubes eventually would have drained, causing local collapse of the tube ceilings. lava tubes locally contain structural drips and striae on their internal walls, suggesting fluctuating flow volumes during emplacement. at isvf, pressure ridges and tumuli are apparent but unevenly distributed across the lava fields. pressure ridges are best viewed to the southwest of miter, where the direction of lava flow changes from northeast to due west. tumuli, however, are more widespread and occur both to the southwest and north of miter. lava channels: there are several channels or topographically depressed areas in isvf (figure 5c). two examples of note are north and southwest of miter. both areas have direct ties to breached crater walls in miter. the depression to the southwest of miter has a smooth floor, whereas the channel extending north of miter is filled with a’a, which typically results from “vigorous eruptions at high discharge” (rowland and walker 1990). a’a is consistently present in several areas of the flow field: (1) at the base of sinuous channels, (2) at the base of other depressed, ponded areas, and (3) at the extreme edges of the lava field (figure 5a). condie and barsky (1972, p. 340) suggested that the “basalts were very fluid at the time of eruption (forming pahoehoe), losing their volatiles and increasing in viscosity as they moved outward, thus developing a’a structure.” this agrees with rowland and walker (1990), who suggested that the main controls on the formation of pahoehoe and a’a are lava viscosity, shear rate, volumetric flow rate, and changes accompanying degassing. mapping a lava flow channel: a lava flow north of miter crater was surveyed using dgps mapping techniques as an example of the complexities within the flow field. the purpose of the mapping was to outline features including cinder cone craters, tumuli and pressure ridges, and flow margins, and to create topographic profiles across the flow to determine heights, lengths, slopes, areas, and volumes of features. the lava field north of miter can be divided into a laterally extensive, terraced, topographic upper unit that surrounds a narrow, sinuous, depressed lower unit (figure 7a). the terraced surfaces dip gently to the north and are dominated highly vesiculated platy rubble, interspersed with flat sheets or poorly developed columnar-jointed basalt and occasional rounded pahoehoe lobes. partially collapsed, anastomosing lava tubes are present along the margins of the terraces, subparallel to the central, sinuous depression. the central depression trends 0.62 mi (1 km) north from a breakout in miter’s crater wall, and it is bounded on the east by pocket and the remains of crescent and modern mining operations. topographic profiles were mapped across the central depression; these profiles show abrupt elevation changes at several mounds and at its edges (figures 7a, 7b, 7c). the central depression is interpreted as a transportation network—a collapsed master tube—for a single lava flow advancing northward from miter. the flow eventually ponded within the wide, semi-circular, northern basin. mounds scattered across its surface are interpreted as tumuli or rafted debris. figure 7. a. location of the channel (opaque yellow) extending from the breached crater wall of miter to the north, ending in a tumuli field (modified from google earth). b. topographic profile a-a' across the terraced upper unit and the central depressed lower unit. c. image of the topographic profile, showing the central channel and its terraced sides. view to the northwest. ice springs volcanic field 11 judge, pollock, williams, schantz, baxstrom, burden, matesich, randall, patzkowsky, sims, cary, peppers, thompson, wilch modeling the flow rate: we calculated the average lava flow velocity at isvf, and the results are similar to calculations at hawaiian lava fields (dragoni,1989; harris and others, 2009). using our topographic profile measurements, we calculated an average effusion rate necessary to emplace the entire edifice of ~58.6 yd3/s (~44.8 m3/s). we can use this effusion rate, along with a calculated total edifice volume, to calculate the total length of the eruption in hours: ~6.19 hours. with a 2789 ft (850 m) long flow, this implies that the flow advanced at an average velocity of 7.6 ft/min (2.31 m/min). this is within 8% of the average velocity of 8.2 ft/min (2.5 m/min) recorded during the 1983 pu’u o’o eruption of kilauea volcano, hawaii (wolfe and others, 1988). cinder mounds: west-northwest of miter there is an anomalous area of rugged topography that is very different from the surrounding a’a and pahoehoe. this area contains mounds of red, commonly bedded, scoriaceous material (figure 5e). the volcanic layers dip outward from the area’s center. the presence of bombs, lapilli, and glassy surfaces is similar to the volcanic material found at the main crater complex. the region is about the size of miter and is previously undocumented in the literature of isvf. previously, hoover (1974) described an eroded cinder mound on the northern lava flow, but none have been described from the flows to the west-northwest of miter. originally, we thought the mound was rafted material from crescent due to its similar geochemistry; however, we now interpret the region as an unrecorded cone that erupted about the same time as crescent. this is consistent with hoover’s (1974) interpretation that the cinder mounds on the northern lava flow were formed in place by eruptions along a north-south fracture system. geologic uniqueness isvf is worth a visit for several reasons: 1. the active mining along the east flank of crescent allows visitors to see the stratigraphy of a young cinder cone. understandably, visitors cannot and should not interfere with active mining operations, so some of this volcanic stratigraphy must be viewed from a distance. visitors should gain access at the mining office at the entrance to the quarry. however, in other places, such as miter and pocket, there is less mining and more opportunity to view cone stratigraphy up close. please understand, though, that mining spoil piles can often conceal cinder flanks from year to year. 2. the age of the eruptions associated with isvf are geologically young. originally thought to be 660 ± 170 years old (valastro and others, 1972), our new cosmogenic ages show it to be 9.8 (±1.6) to 11.1 (±1.7) thousand years old. isvf remains the youngest volcanic center in utah. 3. isvf is the most complex volcanic field in the black rock desert. it was constructed through a series of eruptions that produced overlapping cinder cones and lava flows of slightly different compositions. we consider isvf to be a compound polygenetic volcano (defined in németh and kereszturi, 2015) due to its numerous cinder cones and production of chemically distinct lava flows. 4. isvf is centrally located in the black rock desert. with pahvant butte and tabernacle hill to the north and south and north, respectively, visitors can investigate diverse types and ages of pleistocene volcanism. pahvant butte exhibits several eruptive stages, but it contains deposits that were clearly erupted into lake bonneville at its highest altitude, just prior to the initiation of catastrophic drainage near red rock pass, idaho (oviatt, 2015). the eruption is dated between 18,000 and 14,000 (oviatt and nash, 1989; oviatt, 1991; cerling and craig, 1994)—a time when lake bonneville covered the region. age dates from tabernacle hill deposits span the bonneville to provo shoreline period of lake bonneville. the initial tabernacle hill eruptive phase suggests magma interaction with groundwater, while later stages show deposits erupted onto the lake beds of bonneville. ice springs, the third volcanic field in the area, represents volcanism after lake bonneville had completely ceased to exist in milliard county. acknowledgments we thank red dome, inc., for graciously allowing us access to the crater complex and the lava fields. we appreciate the following wooster, albion, and keck students for their help in the field and lab: julia franceschi (w), ben hinks (a), tricia hall (w), dan minisay (w), pa nhia moua (k), ellen redner (a), kevin silver (w), adam silverstein (w), and chloe wallace (w). thanks to russell robertson for assistance with gis. our work was supported by the keck geology consortium and the college of wooster, department of earth sciences. we are grateful to d.h. elliot, a.r. hintz, and t. wilson for valuable discussions, to d.h. elliot, t. fleming, and c. millan for constructive review comments, and to m. milligan and b. biek for editorial comments that made this document better. their efforts on the manuscript are appreciated. references anderson, s.w., smrekar, s.e., and stofan, e.r., 2012, tumulus development on lava flows: insights from observations of active tumuli and analysis of formation models: bulletin of volcanology, v. 74, p. 931-946. bernardi, m.i., bertotto, g.w., jalowitzki, t.l., orihashi, y., and ponce, a.d., 2015, emplacement history and inflation evidence of a long basaltic lava flow located in southern payenia volcanic province, argentina: journal of volcanology and geophysical research, v. 293, p. 46-56. ice springs volcanic field 12 judge, pollock, williams, schantz, baxstrom, burden, matesich, randall, patzkowsky, sims, cary, peppers, thompson, wilch best, m.g., and brimhall, w.h., 1974, late cenozoic alkalic basaltic magmas in the western colorado plateaus and the basin and range transition zone, u.s.a., and their bearing on mantle dynamics: geological society of america bulletin, v. 85, no. 11, p. 1677-1690. best, m.g., mckee, e.h., and damon, p.e., 1980, space-time-composition patterns of late cenozoic mafic volcanism, southwestern utah and adjoining areas: american journal of science, v. 280, p. 1035-1050. cerling, t.e., and craig, h., 1994, cosmogenic 3he production rates from 39n to 46n latitude, western usa and france: geochimica et cosmochimica acta, v. 58, p. 249-255. christiansen, r.l., and yeats, r.s., 1992, post-laramide geology of the u.s. cordilleran region, in burchfiel, b.c., and others, editors, the cordilleran orogen: conterminous u.s.: geological society of america, geology of north america, v. g-3, p. 261–406. condie, k.c., and barsky, c.k., 1972, origin of quaternary basalts from the black rock desert region, utah: geological society of america bulletin, v. 83, p. 333-352. davis, j., 2014, geosights: volcanic features in the black rock desert, millard county: utah geological survey notes, v. 46, no. 2., p. 10-11. dragoni, m., 1989, a dynamical model of lava flows cooling by radiation: bulletin of volcanology, v. 51, p. 88–95. dunbar, n., 1999, cosmogenic 36cl-determined age of the carrizozo lava flows, south-central new mexico: new mexico geology, v. 21, p. 25-29. gary, w.b., robinson, m.s., zimbelman, j.r., bleacher, j.e., hawke, b.r., crumpler, l.s., braden, s.e., and sato, h., 2012, the origin of ina: evidence for inflated lava flows on the moon: journal of geophysical research, v. 117, no. e00h31 (doi 10.1029/2011je003981). gilbert, g.k., 1890, lake bonneville: u.s. geological survey monograph 1, 438 p. gosse, j.c., and phillips, f.m., 2001, terrestrial in-situ cosmogenic nuclides: theory and applications: quaternary science reviews, v. 20, p. 1475-1560. harris, a.j.l., favalli, m., mazzarini, f., and hamilton, c.w., 2009, construction dynamics of a lava channel: bulletin of volcanology, v. 71, p. 459-474. hintz, a.r., 2008, physical volcanology and hazard analysis of a young monogenetic volcanic field: black rock desert, utah: university of south florida, m.s. thesis, 142 p. hintze, l.f., and davis, f.d., 2002, geologic map of the delta 30' x 60' quadrangle and part of the lynndyl 30' x 60' quadrangle, northeast millard county and parts of juab, sanpete, and sevier counties, utah: utah geological survey map 184, scale 1:100,000. hintze, l.f. and davis, f.d., 2003, geology of millard county: utah geological survey bulletin 133, 305 p. hintze, l.f., davis, f.d., and rowley, p.d., 2003, geologic map of the richfield 30' x 60' quadrangle, southeast millard county and parts of beaver, piute, and sevier counties, utah: utah geological survey map 195, scale 1:100,000. hon, k., kauahikaua, j., denlinger, r., and mackay, k., 1994, emplacement and inflation of pahoehoe sheet flows: observations and measurements of active lava flows on kilauea volcano, hawaii: gsa bulletin, v. 106, p. 351-370. hoover, j.s., 1974, periodic quaternary volcanism in the black rock desert, utah, in rigby, j.k., editor, geology studies: brigham young university, v. 21(1), p. 3-72. houghton, b.f., and wilson, c.j.n., 1989, a vesicularity index for pyroclastic deposits: bulletin of volcanology, v. 51, p. 451-462. johnsen, r.l., smith, e.i., and biek, r.f., 2010, subalkaline volcanism in the black rock desert and markagunt plateau volcanic fields of south-central utah, in carney, s.m., tabet, d.e., and johnson, c.l., editors, geology of south-central utah: utah geological association publication 39, p. 109-150. lifton, n., caffee, m., finkel, r., marrero, s., nishiizumi, k., phillips, f.m., boehring, b., gosse, j., stone, j., schaefer, j., theriault, b., jull, a.j.t., and fifield, k., 2015, in situ cosmogenic nuclide production rate calibration for the cronus-earth project from lake bonneville, utah, shoreline features: quaternary geochronology, v. 26, p. 56-69. lifton, n., sato, t., and dunai, t.j., 2014, scaling in-situ cosmogenic nuclide production rates using analytical approximations to atmospheric cosmic-ray fluxes: earth and planetary science letters 386, p. 149-160. lynch, w.c., and nash, w.p., 1980, chemical trends in the ice springs basalt, black rock desert, utah: all u.s. government documents (utah regional depository, paper 31, https://digitalcommons.usu.edu/govdocs/31/. mabey, d.r. and budding, k.e., 1987, high-temperature geothermal resources of utah: utah geological and mineral survey bulletin 123, 64 p. mabey, d. r. and budding, k.e., 1994. geothermal resources of southwestern utah, in blackett, r. e. and moore, j.n., editors, cenozoic geology and geothermal systems of southwestern utah: utah geological association publication 23, p. 1-23. marrero, s.m., phillips, f.m., borchers, b., lifton, n., amur, r., and balco, g., 2016, cosmogenic nuclide systematics and the cronuscalc program: quaternary geochronology, v. 31, p. 160-187. https://scholarcommons.usf.edu/cgi/viewcontent.cgi?article=1294&context=etd https://scholarcommons.usf.edu/cgi/viewcontent.cgi?article=1294&context=etd https://digitalcommons.usu.edu/govdocs/31/ https://digitalcommons.usu.edu/govdocs/31/ ice springs volcanic field 13 judge, pollock, williams, schantz, baxstrom, burden, matesich, randall, patzkowsky, sims, cary, peppers, thompson, wilch miller, d.m., oviatt, c.g., and mcgeehin, j.p., 2013, stratigraphy and chronology of provo shoreline deposits and lake-level implications, late pleistocene lake bonneville, eastern great basin, usa: boreas, v. 42, p. 342-361. milligan, m., and mcdonald, h.g., 2017, shorelines and vertebrate fauna of pleistocene lake bonneville, utah, idaho, and nevada: geology of the intermountain west, v. 4, p. 181-214. morrison, r.b., 1966, predecessors of great salt lake, in stokes, w.l., editor, guidebook to the geology of utah: utah geological society, v. 20, p. 77-104. nelson, s.t., and tingey, d.g., 1997, time-transgressive and extension-related basaltic volcanism in southwest utah and vicinity: geological society of america bulletin, v. 109(10), p. 1249-1265. németh, k., and kereszturi, g., 2015, monogenetic volcanism: personal views and discussion: international journal of earth science, p. 2131-2146. oviatt, c.g., 1991, quaternary geology of the black rock desert, millard county, utah: utah geological and mineral survey special studies, v. 73, 27 p. oviatt, c.g., 2015, chronology of lake bonneville, 30,000 to 10,000 yr b.p.: quaternary science reviews, v. 110, p. 166171. oviatt, c.g. and jewell, p.w., 2016, chapter 5 – the bonneville shoreline: reconsidering gilbert’s interpretation, in oviatt, c.g., and shroder, j.f., editors, lake bonneville: a scientific update: developments in earth surface processes, v. 20, p. 88-104. oviatt, c.g., and nash, w.p., 1989, late pleistocene basaltic ash and volcanic eruptions in the bonneville basin, utah: geological society of america bulletin, v. 101, no. 2, p. 292-303. patzkowsky, s., randall, e., rosen, m.l., thompson, a., moua, p.n., schantz, k., pollock, m., judge, s., williams, m., and matesich, c., 2017, new cosmogenic and vml dates and revised emplacement history of the ice springs volcanic field in the black rock desert, utah: geological society of america abstracts with programs, v. 49(6), doi: 10.1130/ abs/2017am-306786. pollock, m., edwards, b.r., hauksdóttir, s., alcorn, r., and bowman, l., 2014, geochemical and lithostratigraphic constraints on the formation of pillow-dominated tindars from undirhlíðar quarry, reykjanes peninsula, southwest iceland: lithos, v. 200-201, p. 317-333, doi:10.1016/j.lithos.2014.04.023. rowland, s.k., and walker, g.p.l., 1990, pahoehoe and a’a in hawaii: volumetric flow rate controls the lava structure: bulletin of volcanology, v. 52, p. 615-628. self, s., keszthelyi, l., and thordarson, t., 1998, the importance of pahoehoe: annual review of earth and planetary sciences, v. 26, p. 81-110. sims, w., pollock, m., judge, s., peppers, m., hall, t., cary, w., silver, k., and hintz, a., 2012, petrological and geochemical analysis: mapping of miter flows in ice springs volcanic field, black rock desert, utah: geological society of america abstracts with programs, v. 44(7), p. 391. thompson, j., 2009, crustal assimilation mechanisms in continental basalts: the ice springs flow, utah: university of iowa, m.s. thesis, 203 p. valastro, s., jr., davis, e.m., and varela, a.g., 1972, university of texas at austin radiocarbon dates ix: radiocarbon, v. 14, p. 461-485. white, j.d.l., 1996, pre-emergent construction of a lacustrine basaltic volcano, pahvant butte, utah (usa): bulletin of volcanology, v. 58, p. 249-262. white, j.d.l., 2000, subaqueous eruption-fed density currents and their deposits: precambrian research, v. 101, p. 87-109. wolfe, e.w., neal, c.a., banks, n.g., and duggan, t.j., 1988, chapter 1: geologic observations and chronology of eruptive events, in wolfe, e.w., editor, the pu’u o’o eruption of kilauea volcano, hawaii: episodes 1 through 20, january 3, 1983, through june 8, 1984: u.s. geological survey professional paper 1463, p. 1-97. https://doi: 10.1130/abs/2017am-306786 https://doi: 10.1130/abs/2017am-306786 https://doi:10.1016/j.lithos.2014.04.023 https://doi:10.1016/j.lithos.2014.04.023 uga 50:05 strain partitioning between ductile and brittle stratigraphy— characterizing the sand wash fault system, uinta basin, utah off-canvas toggle member login donate about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle thisistheplace2.jpeg subpage1.jpeg subpage2.jpeg subpage3.jpeg uga 50:05 strain partitioning between ductile and brittle stratigraphy— characterizing the sand wash fault system, uinta basin, utah riley brinkerhoff wasatch energy management, llc john mcbride byu sam hudson byu douglas a. sprinkel aztec geosolutions ron harris byu kevin rey byu eric tingey byu doi: https://doi.org/10.31711/ugap.v50i.109 abstract the sand wash fault zone is a segmented and discontinuous fault system that strikes northwest to south east in the central part of the uinta basin. it is approximately 34 kilometers long with an uncommonly wide damage zone, typically 100 to 200 meters wide. due to recent, rapid, and large-scale incision by the green river and its tributaries, the sand-wash fault zone is well exposed in several closely spaced canyons. these canyon exposures allow mapping of the lateral relationships through panoramic photographs and surface kinematic descriptions.  most movement on the sand wash fault zone occurred in the late eocene, but minor, more recent movement likely occurred. evidence for fault timing includes strata-bound, syndepositional movement which occurred during lake uinta time (55 to 43 ma bp) resulting in debris flows, slump blocks, and small (>150 meters diameter) sag basins filled with poorly organized sediments. after lithification, elongate grabens formed with up to 33.5 meters of horizontal extension. two styles of deformation are present. brittle rocks, such as sandstone and limestone beds, are intensely fractured and faulted, whereas clay and organic-rich rocks are largely unfractured and unfaulted, with variably folded beds that have experienced some layer-parallel slip. laterally, deformation is distributed up to 100 meters from the fault core, which is uncommonly large for faults with short lengths and little displacement. vertically, displacement is concentrated in brittle sandstone and carbonate beds and rare in clay and hydrocarbon-rich units, such as the mahogany oil-shale zone of the eocene green river formation. the mahogany oil-shale zone mostly displays ductile flow (granular flow) commonly forming small décollements between overlying and underlying units. vertical displacement on separate fault segments is generally less than 5 meters and decreases down section, dying out completely around the top of the mahogany oil-shale zone. in this paper we show evidence for syndepositional deformation along the sand wash fault zone, strain partitioning along décollement surfaces, fault surfaces that experience multiple deformational phases, pop-up blocks, and graben development. we also show that deformation on the fault zone is related to extension above a neutral surface of a larger fold. this larger fold is associated with a basement-rooted fault zone that moved during laramide tectonism as the uncompahgre uplift developed. the sand wash fault zone appears to have many similarities to the larger, and more deeply buried, duchesne fault zone 25 kilometers to the north, and the more deeply eroded cedar ridge fault zone located 30 kilometers to the south. the high-resolution fault model, developed herein, is thus a good proxy for other complex fault zones in the uinta basin. our model will be useful to oil and gas operators as they develop horizontal wells across this and other complex fault zones in the basin. order it ($5.00) cart your cart is empty login form login for discounted rates on document downloads and to access members only content username password show password remember me log in forgot your password? forgot your username? create an account utah geological association p.o. box 526356 salt lake city, ut 84152-0100  facebook  youtube © 2025 utah geological society | site by third sun member login off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf uga 50:07 boron in the strata and waters of the eocene green river and uinta formations, uinta basin, utah off-canvas toggle member login donate about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle thisistheplace2.jpeg subpage1.jpeg subpage2.jpeg subpage3.jpeg uga 50:07 boron in the strata and waters of the eocene green river and uinta formations, uinta basin, utah david tabet utah geological survey, retired doi: https://doi.org/10.31711/ugap.v50i.111 abstract boron is a naturally occurring element that can be found in the strata and water of the eocene green river and uinta formations of the uinta basin of utah. whereas boron is suspected to be a necessary trace nutrient for proper plant and animal growth and development, higher concentrations of boron can be detrimental to living things, but the element is not known to be carcinogenic. utah has no limit on boron content in drinking water; however, for irrigation purposes a limit of 750 μg b/l (0.75 mg/l) has been established. the green river formation (grf), deposited in ancient lakes in wyoming, colorado, and utah, is well known for its oil shale and saline mineral deposits, particularly the thick, economic trona beds in wyoming. such evaporitic mineral deposits are characteristic of a saline lake environment that existed at the end stages of eocene lacustrine deposition. in utah they are found in the upper grf, and to a minor extent in the lower uinta formation, where elevated boron is evident. boron-bearing minerals are present in the parachute creek member of the grf in utah, and they occur as secondary silicate minerals in at least 12 wells across utah’s uinta basin. these minerals were first reported by the u.s. geological survey in the 1950s. this study determined that the boron mineral occurrences correlate stratigraphically and coincide with the areas delineated for the hypersaline events in the parachute creek member of the grf. this argues that boron was concentrated with other saline constituents in lake uinta and is an indicator of hypersaline conditions. water quality analyses reporting boron content from surface and subsurface samples were compiled from public databases from the utah department of environmental quality, the u.s. geological survey, and the utah geological survey. this new database consists of 155 samples from 38 sites from the upper part (parachute creek member) of the grf, and 58 samples from 26 sites from the lower part (douglas creek member) of the grf. the boron concentration was found to be different for the two grf parts. the average boron content of the 58 lower grf aquifer water analyses is 6338 μg/l, while the mean boron content for the 26 individual sites varies from 15 to 205,000 μg/l. the upper grf aquifer contains at least twice the boron content of the lower grf, and the average boron content of the 155 analyses of upper grf groundwater is 18,172 μg/l. the mean boron content for the 38 individual sites varies 40 to 480,000 μg/l. areas of highboron concentration in groundwater of the grf tend to coincide with the location of grf hypersaline paleodepocenters. groundwater boron content in the uinta formation comes from 40 analyses from 36 sites. the average of the 40 uinta formation analyses is 3251 μg b/l, while the mean boron content for the 36 individual sites varies from 40 to 24,000 μg/l. for each interval studied, less boron tends to be found in analyses from sites near the outcrop and boron content tends to increase in the studied formations as they are more deeply buried. additionally, boron content was compiled for surface waters and springs for 3955 analyses from 374 sites in the uinta basin. when the mean boron content of these surface water sites was examined by hydrologic drainage unit subareas, it was found that the tributaries in the northern uinta basin, north of the duchesne and white rivers, contain the lowest mean boron contents, whereas higher mean boron contents are common for the tributaries in southern part of the basin. the boron content of surface water from the southern uinta basin drainages also tends to increase northward toward the demarking water courses. the tributaries in the southern part of the uinta basin have higher boron contents due to their waters having contact with the boron-bearing, shallow-inclined strata of the upper grf or member b of the uinta formation. the tributaries in northern part of the basin have lower boron contents because they are not in contact with the grf, have less contact with the member b of the uinta formation, and are diluted by greater snow-melt runoff from the uinta mountains which bound the basin to the north. order it ($5.00) cart your cart is empty login form login for discounted rates on document downloads and to access members only content username password show password remember me log in forgot your password? forgot your username? create an account utah geological association p.o. box 526356 salt lake city, ut 84152-0100  facebook  youtube © 2025 utah geological society | site by third sun member login off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf uga-geosite-ege-devils-playground.indd utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors carl ege utah division of water resources, 1594 w north temple, suite 310 salt lake city ut, 84116 carlege@utah.gov cover image: miniature arch in granitic rock formed due to honeycomb weathering. devils playground, box elder county 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e geosites refl ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president c. ege devils playground 3 introduction why take your kids to the neighborhood playground, when you can visit a playground that inspires their sense of geologic adventure? devils playground is not your ordinary community playground, but a wonderland of granitic rock weathered into fantastic forms and weird shapes (figure 1). occupying an assortment of bureau of land management, state, and private land in the bovine mountains, devils playground is a relatively unknown geologic curiosity found in a remote corner of northwestern utah. devils playground is situated in the physiographic region known as the great basin province that extends across western utah, nevada, and to the sierra nevada mountains in eastern california. the area is composed mostly of granitic rocks of the emigrant pass intrusion. a combination of granitic rock, faulting, and weathering under a semiarid climate created favorable conditions for the creation of devils playground. desert plants such as sagebrush, utah juniper, pinyon pine, mormon tea, and cheatgrass are common throughout the area. geologic history the rocks of devils playground formed approximately 41 to 34 million years ago from an intrusion of magma into overlying paleozoic (300 to 240 million years old) sedimentary rocks. known as the emigrant pass pluton, this intrusion occurred in three stages and is one of at least three distinct intrusive centers in northwestern utah. the pluton is estimated to have been emplaced at depths less than 6 miles (10 km) into the country rock (egger and others, 2003). beginning approximately 13 million years ago, extensional forces thinned and fractured the earth’s crust within the great basin, creating a vast network of north-south trending faults. along these faults, mountains were uplifted and valleys were down-dropped producing the distinctive landscape of the great basin. these geological processes are still operating today throughout the great basin region. one of these north-south trending faults, the grouse creek mountains fault, led to the uplift of grouse creek and bovine mountains where devils playground is found. following uplift, the bedrock was subjected to weathering and erosion by water, ice, wind, and other agents. over millions of years, these erosional processes slowly peeled off the overlying sediments and sedimentary rocks. erosion removed roughly 3 to 6 miles (5-10 km) of rock and sediment before exposing the granitic rocks of the emigrant pass pluton. between 28,000 and 12,000 years ago, a lake called lake bonneville covered much of the area directly south of devils playground. lake bonneville was the largest late pleistocene lake in the great basin (see, for example, oviatt and jewell, 2016) (figure 2). around 14,500 years ago, the lake overflowed at red rock pass in idaho, initiating a massive flood that made its way to the snake river. the lake fell approximately 328 feet (100 m) and at its maximum discharge, the flood released an estimated 35 million cubic feet per second (cfs) of water (1 million cubic meters per second) (janecke and oaks, 2011). sediments and shorelines of lake bonneville can be found southeast of devils playground. figure 1. view of one of the many granitic rock exposures in devils playground. figure 2. location of lake bonneville in the great basin. lake bonneville is shown at its highest level (modified from oviatt, 1997). red star indicates the location of devils playground. digital compilation by carmen mcdonald. 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 stratigraphy the rocks of the devils playground area consist of granitic rock intruded into late paleozoic interbedded sandstones and limestones; these sedimentary rocks are locally metamorphosed. additionally, surficial deposits derived from the local bedrock are found throughout the area. geologic units are shown on the stratigraphic column in figure 3 and the geologic map on figure 4. mississippian rocks diamond peak formation and chainman shale (ipmdc) this bedrock is the oldest exposed in the area and consists of metamorphic rock such as dark-gray to black phyllite, fine-grained semischist, gray sandstone, and gray pebble conglomerate and marble (miller and others, 2019). the bedrock mostly crops out in the grouse creek mountains (northwest of devils playground), however two small outcrops are just northeast of the turnoff from state highway 30. permian rocks oquirrh group (poi) the oquirrh group consists of interbedded sandstone and limestone deposited in a marine environment (hintze and kowallis, 2009). this bedrock has been altered in areas by intrusion of the emigrant pass pluton (tepg). this alteration process is called contact metamorphism because a hot magma body comes in contact with country rock, altering its mineral composition and texture. good examples of this alteration can be found south and west of devils playground. eocene rocks emigrant pass pluton (tepg) the emigrant pass pluton crops out in three lobes and covers an area of approximately 10 square miles (25 km2) in the southern part of the grouse creek mountains (doelling, 1980; miller and others, 2012, 2019 in review). the rock is mostly granite containing equal parts of quartz, plagioclase, potassium feldspar, and biotite (miller and others, 2012). other intrusive rocks found in the area include granodiorite, quartz diorite, diorite, and monzonite. also found throughout the granitic rock are pegmatites (coarse-grained igneous rocks with figure 3. stratigraphic column showing formations exposed in the devils playground area, including thickness and lithology (modified from hintze and kowallis, 2009). digital compilation by carmen mcdonald. figure 4. geologic map of devils playground. geology modified from miller and others (2012) and miller and others (2019). digital compilation by john good. c. ege devils playground 5 interlocking quartz, feldspar, goethite after pyrite, and muscovite). these pegmatites formed during the late stages of the intrusion and are found throughout the pluton, usually as irregular dikes or veins several inches to 3 feet (<1 m) in thickness, and extend hundreds of feet (figure 5). they represent the last and most hydrous (water-rich) portion of magma to crystallize. the pegmatites are easy to locate because they are more resistant than the surrounding rock, thus they resemble ribs and bones sticking out in relief. some of the pegmatites contain pockets of loose quartz and feldspar crystals in the center portion of the pegmatite veins (figure 6). these crystals formed when the internal pressure of the pocket exceeds the confining pressure of the surrounding rock, thus causing the pockets to rupture in an explosive manner, fracturing many of the crystals and breaking them from the wall rock. after the breakage, mineral growth continues within the pocket forming new terminations on the broken ends of crystals. some of these quartz crystals are smoky or gray-black in color. this color is due to natural gamma irradiation and the presence of trace amounts of aluminum within the crystal structure. quaternary and holocene deposits older alluvial fans (qao) older, pre-bonneville alluvial fans are present mostly north and west of the devils playground area. the alluvial fans are deeply eroded and comprised of poorly sorted sand, mud, and boulders deposited by debris flows and derived from erosion of local bedrock. lacustrine gravel, lake bonneville (qlg) upper pleistocene lake bonneville gravel and sand is present along the access road east of devils playground. the high stand of lake bonneville, preserved as the lake bonneville shoreline, forms an arcuate, berm-like barrier across sand wash. lacustrine and alluvial coarse-grained deposits (qlag) holocene to upper pleistocene sediment found mostly southeast of devils playground. these sediments consist of mixed layers of coarse-grained lake bonneville sediments and younger alluvial sand and gravel. lacustrine and alluvial deposits, undivided over tertiary bedrock unit (qla/tepg) holocene to upper pleistocene lake bonneville gravel and sand that mostly cover emigrant pass intrusive rocks (tepg). this unit is quite extensive and found mainly east of devils playground. alluvial mud (qam) holocene mud and sandy mud that accumulated behind lake bonneville gravel barriers. younger alluvial fans (qafy) these younger holocene alluvial fan deposits consist of mud, sand, and gravel found in washes and valley bottoms. these sediments are often deposited in high-energy flash flood events from summer thunderstorms and are therefore poorly sorted. geologic uniqueness “devils,” alcoves, spires, arches, small caves, bowls, and honeycomb patterns can be found throughout the area and offer an explorer endless opportunities of discovery. these extraordinary forms of the granitic rock owe their structure to a variety of physical and chemical weathering processes. these processes began sculpting the rocks immediately after they were exposed by erosion. the two most conspicuous processes at devils playground are spheroidal (onion-skin) and honeycomb weathering. spheroidal weathering is commonly found in granitic rocks and is showcased at devils playground (figure 7). in spheroidal weathering, joints or fractures create initial openings allowing surface water to access the rock from all sides. water seeping along these fractures slowly decomposes or alters the mineral composition of the granitic rock, causing the rock to weather inward. as a result, rounded shells of decomposing rock are repeatedly loosened and peeled off the unweathered rock core like layers of an onion. the rate of weathering is greatest along the corners and edges where fractures and joints intersect because they have a greater surface-area-to-volume ratio than the rock faces. eventually, the rock core eventually becomes an isolated rounded boulder at the ground surface and is referred as a corestone or woolsack. figure 5. pegmatite vein in granitic rock. notice exposed pocket. keys for scale. figure 6. smoky quartz and feldspar specimen found in a pegmatite at devils playground. 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 however, some rock at devils playground display another type of weathering called honeycomb weathering, also known as alveolar weathering (figures 8 and 9). here, weathering creates pockets or cavities known as tafoni, giving the rock a honeycomb or stone lattice appearance. this weathering is found throughout the world in semi-arid regions, coastal areas, arctic deserts, and even on historical buildings (rodrigues-navarro and others, 1999). the weathering is likely produced by biological, chemical, and physical factors operating in the micro-environment. salt weathering is often attributed for the honeycomb features found along the coast or in deserts. nearby salt from the soil, playas, or the ocean is deposited on the surface of the rock by the wind and rain. the water allows the salt to bind to the rock, so as the salt solution evaporates, the salt begins to crystallize and grow within the pore spaces of the rock. these salt crystals then pry apart the rocks’s mineral grains and form pits in the weathered material resembling a honeycombed structure. other mechanisms that may contribute to the honeycomb pattern may include the wind, freeze-thaw cycles, humidity, bedding, and solar radiation (davis, 2012). in the case of devils playground, several of these mechanisms likely contributed to the formation of the honeycomb appearance of the rocks we see today. eventually, these weathering processes will destroy all of these current artistic forms of nature. however, these processes will continue to sculpt new features as long as weathering and erosion continues to uncover granitic rock at devils playground. figure 7. granitic rock that exhibits spheroidal weathering. figure 8. honeycomb weathering found at devils playground. sunglasses for scale. figure 9. miniature arch in granitic rock formed due to honeycomb weathering. c. ege devils playground 7 suggested driving instructions from the northern i-15/i-80 interchange in salt lake city, travel north on i-15 for 69 miles (111 km) to tremonton (exit 379). at exit 379, the freeway splits; i-84 continues to the northwest (left), and i-15 branches off and goes north (right). travel northwest on i-84 for 37 miles (60 km) to exit 5. turn left (west) on state highway 30 and travel 16 miles (28 km) to curlew junction (a junction with state highway 42). turn left (southwest) and proceed 49 miles (79 km) to the devils playground road. there is no sign at this turnoff, however there is a bright yellow cattle guard that marks the entrance into devils playground. turn right (north) and drive approximately 1 mile (1.6 km) to the first granitic outcrops on the right (northeast) side of the road. if you proceed on this road for several miles, the road will end up in the heart of devils playground. the utm coordinates (wgs84) of devils playground are: zone 12t e: 277958 n: 4599061. acknowledgments i thank the many people who provided assistance on this paper, in particular joel williams utah division of water resources (dwre) who authorized the time in the busy work week to complete the manuscript. jim davis and greg mcdonald utah geological survey (ugs) reviewed the manuscript and provided additional information to the text. also, special thanks to the adam clark and carmen mcdonald (dwre) who helped with the digital compilation of the location map and stratigraphic column for the text. also, thanks to don clark (ugs) for use of unpublished geologic mapping and john good (ugs) for the digital compilation of the geologic map. references davis, j., 2012, geosights: the honeycombs, juab county, utah: utah geological survey, survey notes, v. 44, no. 1, p. 10-11. doelling, h.h., 1980, geology and mineral resources of box elder county, utah: utah geological and mineral survey bulletin 115, 251 p. egger, a.e., dumitru, t.a., miller, e.l., savage, c.f.i., and wooden, j.l., 2003, timing and nature of tertiary plutonism and extension in the grouse creek mountains, utah: international geology review, v. 45, no. 6, p. 497-532. hintze, l.f., and kowallis, b.j., 2009, geologic history of utah: provo, utah, brigham young university geology studies special publication 9, 225 p. janecke, s.u., and oaks, r.q., jr., 2011, reinterpreted history of latest pleistocene lake bonneville: geologic setting of threshold failure, bonneville flood, deltas of the bear river, and outlets for two provo shorelines, southeastern idaho, usa, in lee, j., and evans, j.p., eds., geologic field trips to the basin and range, rocky mountains, snake river plain, and terranes of the u.s. cordillera: geological society of america field guide 21, p. 195–222. miller, d.m., clark, d.l., wells, m.l., oviatt, c.g., felger, t.j., and toda, v.r., 2012, progress report geologic map of the grouse creek 30' x 60' quadrangle, box elder county, utah, and cassia county, idaho (year 3 of 4): utah geological survey open file report 598, 25 p., scale 1:62,500. miller, d.m., felger, t.j., and langenheim, v.e., 2019 (in review), geologic and geophysical maps of the newfoundland mountains and part of the adjacent wells 30' x 60' quadrangles, box elder county, utah: utah geological survey map, scale 1:62,500. oviatt, c.g., 1997, lake bonneville fluctuations and global climate change: geology, v. 25, no. 2, p. 155-157. oviatt, c.g., and jewell, p.w., 2016, reconsidering the bonneville shoreline, in oviatt, c.g., editor, lake bonneville—a scientific update: elsevier b.v., developments in earth surface processes, v. 20, p. 88-104, http://dx.doi.org/10.1016/b978-0-44463590-7:00005-6. rodriguez-navarro, c., doehne, e., and sebastian, e., 1999, origins of honeycomb weathering: the role of salts and wind: geological society of america bulletin, v. 111, no. 8, p. 1250-1255. figure 10. location map of devils playground. digital compilation by adam clark. uga 50:02 geologic characterization and depositional history of the uteland butte member, green river formation, southwestern uinta basin, utah off-canvas toggle member login donate about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle thisistheplace2.jpeg subpage1.jpeg subpage2.jpeg subpage3.jpeg uga 50:02 geologic characterization and depositional history of the uteland butte member, green river formation, southwestern uinta basin, utah ryan d. gall utah geological survey justin e. birdwell u.s. geological survey riley brinkerhoff wasatch energy management, llc michael d. vanden berg utah geological survey doi: https://doi.org/10.31711/ugap.v50i.106 abstract the 15to 65-m-thick informal uteland butte member of the eocene green river formation represents the first widespread transgression of lake uinta in the uinta basin, utah. this study assesses the spatial and temporal variation of uteland butte member deposits along a 40-km transect in the southwestern margin of the uinta basin using detailed measured sections, organic and inorganic geochemical data, and outcrop gamma ray logs. fourteen lithofacies are identified, which comprise seven facies associations linked to with lacustrine, palustrine, and deltaic depositional settings. facies associations are traceable laterally across the study area, where five 4to 12-m-thick depositional cycles are identified. each shallowing upwards cycle is defined by a >1.5-m-thick basal package of organic-rich, argillaceous laminated mudstone, and is capped by thick packages of bedded carbonate. in the far western study area (kyune creek canyon), thick deposits of organic-rich mudstone are present and represent the most distal outcrop section; time-equivalent strata in the eastern study area (minnie maud creek canyon) are relatively organic lean with higher silt and clay content, interpreted to represent proximal lake margin deposits influenced by a nearby delta. the outcrop belt is correlated to more distal cores and well logs across the western uinta basin. similar lithological and petrophysical patterns across the western uinta basin are used to subdivide stratigraphy into nine laterally contiguous sub-units based on nomenclature from the oil-producing area of the central basin (from base to top: lower uteland butte, d bench, d shale, c bench, c shale, b bench, b shale, a bench, and a shale). siliciclastic clay-rich and carbonaterich intervals are correlated across the region and indicate distinct siliciclasticand carbonate-dominated lake phases during uteland butte member deposition. climate is interpreted to be the dominant driver of these claycarbonate cycles, in which relatively humid periods resulted in increased fluvially derived siliciclastic sediment into the basin (clay-rich periods), and arid periods resulted in evaporative conditions with decreased fluvial sediment input that favor carbonate accumulation. climatically driven depositional cycles within the uteland butte member reflect, to a smaller degree, the larger scale climatically driven depositional cycles observed at the memberand formation levels of paleocene and eocene uinta basin stratigraphy. importantly, the uteland butte member clay-carbonate cycles showcase how relatively small-scale climate shifts can impact basin-scale lacustrine deposition. order it ($5.00) cart your cart is empty login form login for discounted rates on document downloads and to access members only content username password show password remember me log in forgot your password? forgot your username? create an account utah geological association p.o. box 526356 salt lake city, ut 84152-0100  facebook  youtube © 2025 utah geological society | site by third sun member login off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf uga 50:04 paleoenvironmental reconstructions along the southern margin of eocene lake uinta during a period of climatic variability: the middle to upper green river formation, uinta basin, utah off-canvas toggle member login donate about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle thisistheplace2.jpeg subpage1.jpeg subpage2.jpeg subpage3.jpeg uga 50:04 paleoenvironmental reconstructions along the southern margin of eocene lake uinta during a period of climatic variability: the middle to upper green river formation, uinta basin, utah lauren p. birgenheier university of utah, department of geology and geophysics leah c. toms beach energy doi: https://doi.org/10.31711/ugap.v50i.108 abstract the stratigraphy preserved in the lacustrine eocene green river formation suggests a variety of depositional environments that were subject to pulses of sediment and lake level fluctuations in the uinta basin. this paper defines 13 facies within six facies associations observed in outcrop and core from the middle to upper green river formation in gate canyon, south-central uinta basin, and willow creek/indian canyon, western  uinta basin, along the margin of the ancient lake system. the facies associations are defined by dominant lithology, relative siliciclastic sediment supply, relative slope, and lake zonation, and encompass a suite of facies revealing distinct environments across a relatively small region (~35 km apart). in the lower interval, a shift  from a high energy, steeper ramp setting in the east, to a lower energy, shallower ramp setting to the west suggests a transition off the main depositional axis. correlatable siliciclastic and carbonate dominated packages  reflect the impact of early eocene hyperthermal events that followed the paleocene-eocene thermal maximum, and the diverse facies present at the two study areas suggest variable responses to these climatic events.  in the overlying stratigraphy, the facies are consistent in both regions and consist of organic rich and poor carbonate mudstones and siltstones. the interval thickens slightly from east to west, and the overlying horse  bench sandstone transitions from a wave-dominated shoreface sequence to a river-dominated deltaic sequence, indicating an overall shift in the lake depocenter towards the west. this study offers a unique glimpse  at the paleoenvironmental changes along the margin of a continental interior lake system during a time of climatic variability. it also highlights the need to use multiple depositional models to accurately characterize marginal deposits in different regions of large internally-drained lake systems which, unlike typical marine depositional models, are more sensitive to basin geometry, size, and sediment fairways. order it ($5.00) cart your cart is empty login form login for discounted rates on document downloads and to access members only content username password show password remember me log in forgot your password? forgot your username? create an account utah geological association p.o. box 526356 salt lake city, ut 84152-0100  facebook  youtube © 2025 utah geological society | site by third sun member login off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf uga 50:06 controls on organic matter variation during deposition of the mahogany oil shale zone of the parachute creek member, green river formation, utah off-canvas toggle member login donate about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle thisistheplace2.jpeg subpage1.jpeg subpage2.jpeg subpage3.jpeg uga 50:06 controls on organic matter variation during deposition of the mahogany oil shale zone of the parachute creek member, green river formation, utah amy l. elson ocean and earth science, university of southampton john marshall ocean and earth science, university of southampton jessica h. whiteside ocean and earth science, university of southampton doi: https://doi.org/10.31711/ugap.v50i.110 abstract the green river formation of utah, colorado, and wyoming represents a ~10 million-year early-middle eocene record of an unusually large, productive lacustrine system composed of several interconnected basins. this system gave rise to one of the largest oil shale deposits in the world, which preserves a rich trove of information about the climate and ecosystems that prevailed during the early eocene climatic optimum. this study uses multiple analytical approaches, including both traditional methods and novel proxies, to determine what brought about the accumulation of the unusually-rich oil shales (>40% toc) of the mahogany zone in the uinta basin portion of the green river formation. the organic-rich mudstones of the mahogany zone exhibit strong micro-scale heterogeneity, indicating complex controls on organic matter production and distribution. petrographic observations indicate highly variable amounts of terrestrial organic matter (spores, wood and plant debris), with the most abundant being amorphous organic material in thin organic-rich laminae. biomarker ratios indicate that the organic-rich laminae consist of different types of microbially derived organic matter, primarily bituminite and organo-minerallic aggregates of the fluorescent liptinite group. these largely lacustrine-derived laminated deposits are argued to have been produced by large intrato inter-annual algal blooms and other enhanced microbial productivity events paced by longer-term sub-orbital cycle fluctuations, such as el niño–southern oscillation or sunspot variations. the mahogany zone deposits appear to have sequestered enough carbon (~64.25 gt estimated) to suggest that, in aggregate, they were large enough to cause a draw-down of terrestrial co2 and exert a significant negative feedback effect on climatic warming. during the climatic cooling that occurred on the declining limb of the early eocene climatic optimum, the sedimentary provenance in the uinta basin shifted from southerly sources dominated by large ephemeral fluvial systems, to northerly sources rich in feldspathic and carbonate detritus derived from northern sites in wyoming. the latter were delivered to the uinta basin via a network of interconnected lacustrine basins. increasing volcaniclastic material delivered from these northerly basins infilled the uinta basin from east to west, and their arrival heralded the end of organic-rich deposition, most notably the prolific oil shales for which the green river formation is renowned. order it ($5.00) cart your cart is empty login form login for discounted rates on document downloads and to access members only content username password show password remember me log in forgot your password? forgot your username? create an account utah geological association p.o. box 526356 salt lake city, ut 84152-0100  facebook  youtube © 2025 utah geological society | site by third sun member login off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf uga 50:09 spatial and temporal cycle variations in the eocene lacustrine green river formation, piceance creek basin, colorado off-canvas toggle member login donate about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle thisistheplace2.jpeg subpage1.jpeg subpage2.jpeg subpage3.jpeg uga 50:09 spatial and temporal cycle variations in the eocene lacustrine green river formation, piceance creek basin, colorado kuwanna dyer-pietras department of geological sciences, binghamton university doi: https://doi.org/10.31711/ugap.v50i.113 abstract lithological cycles in the lacustrine green river formation of the piceance creek basin, colorado, vary spatially and temporally, and are challenging to correlate across the basin. in the center of the basin, lacustrine deposits contain bedded evaporites and kerogen-rich carbonate mudstone, whereas basin margin deposits contain alluvial sandstone, kerogen-poor mudstone, stromatolite, and mud cracks. because coarsening upward cycles identified at the basin center differ significantly from cycles at the basin margin, exactly how cycle boundaries extend from one location in the basin to another is commonly unclear. in this investigation, eleven lithofacies are identified based on rock composition, sedimentary textures, physical sedimentary structures, biogenic structures, and chemical features. lithofacies are grouped into four facies associations: (1) alluvial facies association; (2) shoreline facies association; (3) shallow lake facies association; and (4) deep lake facies association. lithological cycles in the green river formation vary temporally and spatially, with each stratigraphic section preserving a record of deposition in a particular area of the basin as the piceance lake evolved. to address the temporal and spatial variation of the cycles, multiple “type” cycles are described for each location in the study area. facies associations and temporal stacking of cycles suggest that the piceance lake was initially a deep, freshwater lake, following the long point transgression, as lake level rose above the douglas creek arch to fill both the uinta and piceance creek basins. lake levels then stabilized but later began a period of increased fluctuation. following basin closure, the piceance lake became hypersaline under evaporative concentration conditions, precipitating nahcolite and halite at the basin center, with exposure of mudflats at the basin margin. the lake would remain this way until the mahogany transgression deepened and diluted the lake. together, facies associations and “type” cycles described in the green river formation lacustrine deposits reveal a record of lateral and vertical facies changes that occurred in the piceance lake during its evolution. order it ($5.00) cart your cart is empty login form login for discounted rates on document downloads and to access members only content username password show password remember me log in forgot your password? forgot your username? create an account utah geological association p.o. box 526356 salt lake city, ut 84152-0100  facebook  youtube © 2025 utah geological society | site by third sun member login off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf uga 50:08 sunnyside bitumen-impregnated sandstone reservoirs at bruin point, southwest uinta basin, utah off-canvas toggle member login donate about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle thisistheplace2.jpeg subpage1.jpeg subpage2.jpeg subpage3.jpeg uga 50:08 sunnyside bitumen-impregnated sandstone reservoirs at bruin point, southwest uinta basin, utah steven schamel geox consulting inc. doi: https://doi.org/10.31711/ugap.v50i.112 abstract the south flank of the uinta basin in northeast utah is the high tavaputs plateau that dips gently northward towards the basin center and is deeply dissected by the green river and its tributaries. where not removed by canyon incision, bitumen-impregnated sandstones in the green river formation (lower eocene) occur along nearly the full length of the plateau. the bituminous sandstones encompass an area greater than 600 square miles and hold an estimated 11.6 to 14.0 billion barrels of bitumen, but the net thickness of bituminous sandstone and the ooip rarely exceed 70 feet and 80 thousand barrels per acre (mbo/ac), respectively. however, in an exceptional 4.6 square mile area centered on bruin point (elev. 10,184 ft) on the southwest basin rim, the net thickness of bituminous sandstone and ooip are measured in hundreds of feet and mbo/ac, respectively. the estimated bitumen in place in just this small area is 1.16 billion barrels. bruin point is an erosional remnant of a structural-stratigraphic trap formed by the superposition of a subtle flexure on a thick stack of deltaiclittoral sandstones. in the 1970s and 1980s, this unique area was extensively investigated to characterize the reservoirs and delineate the hydrocarbon resource. over 117 test wells with cores were drilled and analyzed. as many as 32 stacked bituminous sandstone bodies were encountered, 15 of which hold nearly all the bitumen. the sandstones were deposited in constantly shifting deltaic lobes and along inter-delta shorelines. they are encased in marsh and floodplain mudstone and littoral-lacustrine calcareous mudstone and bioclastics. the sandstones are poorly sorted, fine-grained feldspathic arenites up to 115 ft thick in distributary channels, but less than 10 ft thick in beach deposits. average porosity and permeability are 23% and 570 md, respectively, but values vary widely between, and even within, depositional settings. the bitumen at bruin point is heavy (8.6o api) and highly viscous (106 cp). just 25 miles to the north, these same amalgamated deltaic sandstones are the reservoirs in the greater monument butte conventional oil field. although many operators have attempted to exploit the bruin point site for liquid hydrocarbons using both in situ steam flood and mining with solvent extraction, so far only small-scale mining of the bituminous sandstones for road construction has been commercially successful. this paper presents seven cores, several exceeding 1000 ft in length, that graphically display the distribution of bitumen resource as related to stratigraphic heterogeneity and location within the bruin point sector of the sunnyside deposit. order it ($5.00) cart your cart is empty login form login for discounted rates on document downloads and to access members only content username password show password remember me log in forgot your password? forgot your username? create an account utah geological association p.o. box 526356 salt lake city, ut 84152-0100  facebook  youtube © 2025 utah geological society | site by third sun member login off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf uga-geosite-biek-brian-head-peak.indd utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors robert f. biek1 and peter d. rowley2 1utah geological survey, p.o. box 146100, salt lake city, ut 84114-6100, bobbiek@utah.gov 2geologic mapping inc., p.o. box 651, new harmony, ut 84757 cover image: th e type section of the brian head formation (tbh) is on the ridge just right of center (by the tbh label). brian head peak is capped by the leach canyon formation (tql), which overlies the isom formation (ti). a modern landslide (qms) is west and south of the peak. brian head peak, iron county 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e geosites refl ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: biek, r.f., and rowley p.d., 2019, brian head peak, iron county, in milligan, m., biek, r.f., inkenbrandt, p., and nielsen, p., editors, utah geosites: utah geological association publication 48, 11 p., https://doi.org/10.31711/geosites. v1i1.47. presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president r.f. biek and p.d. rowley brian head peak 3 introduction brian head peak, the highest point on the west edge of the markagunt plateau at 11,307 feet (3447 m), provides stunning views westward into the great basin. the plateau is part of the high plateaus, a subprovince of the colorado plateau. few views in southern utah so well demonstrate the huge difference between the badly broken great basin, where east-west crustal extension (pulling apart) produced north-trending faulted basins and intervening ranges, and the much less deformed and here higher colorado plateau. the southwestern flank of the peak is the type section of the brian head formation, an eocene to oligocene stream and lake deposit that is overlain by densely welded ash-flow tuff of the 27 to 26 ma isom formation. the peak itself is capped by the moderately welded, 23.8 ma ash-flow tuff of the leach canyon formation (figure 1). these two regionally extensive, upper oligocene ash-flow tuffs erupted suddenly and explosively from calderas near the utah-nevada border and made their way in minutes to their present position, devastating everything in between, millions of years before the episode of basin-range deformation formed the great basin and uplifted the high plateaus. both bear on the timing of basin-range deformation, and both can be visited on brian head. furthermore, the isom formation itself is a key player in understanding earth’s largest terrestrial landslide, the markagunt gravity slide (see sidney peaks geosite). the south side of the peak offers the best exposures of the leach canyon formation, including its vitrophyre and basal surge deposits, which are seldom exposed glassy and sandy parts of typical ash-flow tuffs. location brian head peak, just off utah highway 143 north of cedar breaks national monument, is accessible via brian head peak road (forest service road 047) during the summer to early fall; it is closed by snow the remainder of the year. the first 1.9 miles (3 km), to a forest service trailhead parking area with pit toilet, is an improved gravel road typically accessible by cars. the last 0.75 miles (1.2 km) is rougher but usually still passable by high-clearance, two-wheel drive vehicles. an open shelter, built by the ccc (civilian conservation corps) of local stone (the leach canyon formation), caps the summit plateau. the best place to see the entire leach canyon formation is at the south end of the peak at 37° 40' 48.2", 112° 49' 51", about 140 yards (125 m) south of the ccc shelter. figures 2, 3, and 4 provide a geologic map, stratigraphic column, and cross section for the area. stratigraphy regional ash-flow tuffs utah’s middle cenozoic landscape looked unimaginably different from that of today. geologists refer to that former landscape as the great basin altiplano or nevadaplano, a high-elevation region that stretched from the sierra nevada in eastern california eastward to what is now the colorado plateau (decelles, 2004; best and others, 2009, 2013). the altiplano was studded with volcanic mountains and intervening basins, analogous perhaps to the modern andean altiplano of south america. it was onto this landscape figure 1. the type section of the brian head formation (tbh) is on the ridge just right of center (by the tbh label). brian head peak is capped by the leach canyon formation (tql), which overlies the isom formation (ti). a modern landslide (qms) is west and south of the peak. 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 l 3000 ft brian head peakbrian head peakbrian head peak cedar breaks north view overlook cedar breaks cedar breaks cedar breaks cedar breaks tcp qc tcwt ti qms tcwu ti tcwm tbh tql ti tcwm qms qacf tbh qac tm(ti) tcwl qc tm qmt qc tcp qac qac qms tm qmt qacf qms qacf qms ti tcwl ti qmt qmt qacfo qmsh tbh qms tbh tm(ti) qes tm(ti) qes tbh qac tcwutcwm td tm(ti) qal black fault le dge bh071808-4, -5, -6 bh071808-7 bh071808-2, -3 bh071808-1 bh062310-1 2 6 42 53 37 14 71 16 contact normal fault gravity slide fault sinkhole bedding attitude15 figure 2. geologic map of the brian head peak area. map units are shown on figure 3 (surficial deposits are various shades of light yellow). note the large area of modern landslides (yellow with red triangle pattern) that resulted from failure of the brian head formation. the brian head ski area, and the southern end of the yankee meadows graben, is in the upper left part of the map. from rowley and others (2013). r.f. biek and p.d. rowley brian head peak 5 that dozens of widespread ash-flow tuffs accumulated—part of the middle cenozoic “ignimbrite flare-up” of southwestern north america, which represents one of the largest episodes of subduction-related, silicic volcanism known on earth (figure 5). ash-flow tuffs (ignimbrites in european parlance) are the deposits of pyroclastic flows, density currents of hot volcanic rock, ash, and gases derived from explosive volcanic eruptions. pyroclastic flows can travel rapidly more than 100 miles (160 km) across the landmiocene markagunt megabreccia tm(ti) tm(tnw) tm(tbh) tm 450 (140) 300 (90) 40 (12) 150 (45) mount dutton fm. td 10 (3) tm(tqcb) leach canyon formation tql 55–100 (17–30) isom formation ti 350 (107) oligocene wah wah springs fm. tnw 3–8 (1–3) brian head formation tbh 500 (150) eocene uppermost unit middle unit upper limestone unit tcw lower limestone unit tcwm tcwl 47 (14) tcwu tcwt w hi te m em be r 440 (135) 310 (94) te r ti a r y paleocene pink memberc la ro n fo rm at io n tcp 1000 (305) 109 (33) 45–60(14–18) 23.0 33.9 ?55.8 50+ (15+) about 20–22 ma 23.8 ma 26–27 ma 29.5 ma 35 ma 36 ma ma system and stage lithology age map unit symbol map thickness feet (meters)series stratigraphic column figure 3. stratigraphic column showing rock units of the brian head peak area. the markagunt megabreccia is the deposit of the markagunt gravity slide. from rowley and others (2013). figure 4. cross section through black ledge at sydney peaks. note that line of cross section lies just north of and covers a wider area than that shown in figure 2. the yankee meadows graben, down-dropped blocks between the black ledge and rattlesnake canyon faults, is covered with modern landslide deposits derived from the clay-rich brian head formation. from rowley and others (2013). black ledge thin surficial deposits not shown west a 11,000 10,000 9,000 8,000 7,000 6,000 navajo ridge parowan canyon utah state route 143 brian head resort well #2 (projected to section) bend in section dark hollow sidney peaks east a' 11,000 10,000 9,000 8,000 7,000 6,000 tm ksj kw tcp kgc & kwcs km tcw tbh ti ksj kw tcp kgc & kwcs km tcw tbh ti ksj kw tcp kgc & kwcs km qms qms tm(tbh) ksd navajo ridge fault ksd faulty a n k e e m e a d o w s g r a b e n brian head resort well #1 (projected to section) rattlesnake canyon fault tm cretaceous, undivided cretaceous, undivided e le va tio n in fe et m a r k a g u n t p l a t e a u scape, filling valleys that radiate away from volcanic highlands. one interesting and very useful characteristic of ash-flow tuffs, noted by mackin (1960), is that ash-flow tuffs are emplaced in a geological instant over broad areas, and thus serve as important time horizons for correlating rock formations and understanding structural development of the region. best and others (2013) summarized how our understanding of these ash-flow tuffs has evolved, beginning in the 1950s with j. hoover mackin who first realized that they were indeed the products of enormous catastrophic eruptions of volcanic ash, not simply lava flows. today, the calderas themselves are recognizable only through mapping of stratigraphic and structural relations between caldera in-fill and outflow deposits, inasmuch as 20 million years of subsequent basin-range extension, erosion, and burial under intervening basins makes the calderas all but invisible in the modern landscape. the eruption of oligocene to miocene ash-flow tuffs in nevada and utah is part of a broad pattern of volcanism that migrated southward through time across northwestern north america from about 55 to 20 million years ago (mackin, 1960; cook, 1965; armstrong and others, 1969; stewart and carlson, 1976; stewart and others, 1977; rowley, 1998; rowley and dixon, 2001). the southward migration resulted from complex plate tectonic interactions along the western margin of north america as outlined by dickinson (2006) and humphreys (2009). ultimately, these and other researchers hypothesize a tear in the relatively cold and dense oceanic crust of a subduction zone. the tear allowed the subducting slab to peel away from the less dense continental crust above, opening a window through which relatively hot upper mantle rock ascended, feeding the magmatic flare-up. best and others (2013) provided a comprehensive summary of the great basin ash-flow tuff province of nevada and western utah, where, from about 36 to 18 million years ago, more than 200 large eruptions from 42 calderas resulted in more than 16,500 cubic miles (70,000 km3) of tuff deposited over the landscape (figure 5). in the indian peak and caliente caldera complexes in the eastern part of the tuff province, more than 50 large eruptions produced 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 an estimated 7600 cubic miles (32,000 km3) of ash-fl ow tuff s now spread over an area of 15,000 square miles (63,000 km2) in east-central nevada and southwestern utah (this volume is roughly enough to fi ll the grand canyon nearly eight times over). nine of those eruptions are popularly known as “super eruptions,” each having ejected more than 240 cubic miles (1000 km3) of rock, including from older to younger the wah wah springs, lund, isom, and leach canyon formations, the latter two of which are found at brian head peak. in addition to ash-fl ow tuff s, these eruptions produced voluminous ash falls—fi ne-grained ash fall from the wah wah springs eruption, for example, is recognized in western nebraska (best and others, 2013). over time, repeated eruptions of the indian peak and caliente caldera complexes produced ash-fl ow tuff s that fi lled topographic low areas in the great basin altiplano. because ash-fl ow tuff s tend to be confi ned to former river valleys far from their source area, the distal ends of many ash-fl ow tuff s behave geomorphically like basaltic lava fl ows (see the st. george inverted valleys geosite). because it is relatively fl uid when it erupts, basaltic lava fl ows downhill until it becomes entrained in old stream valleys. th is buries and displaces the stream, and because the basalt is typically harder than enclosing rocks and sediments of the valley walls, the stream shift s to the side of the fl ow and preferentially erodes adjacent, less-resistant rock. ultimately, through a process called topographic inversion, this leaves the lava fl ow as a sinuous high ridge above the surrounding landscape—the ridge marks the location of the former stream channel. topographic inversion of the distal ends of many ash-fl ow tuff s naturally ensues as it does for basaltic lava fl ows. th e distribution of regional ash-fl ow tuff s on the markagunt plateau and sevier plateau (the next range to the east) suggests that they may have been controlled by broad, east-fl owing stream valleys (biek and others, 2015). when an ash-fl ow tuff blocked a stream valley, displaced streams stepped southward to create another valley later to be fi lled by a subsequent, younger ash-fl ow tuff . in this way, older ash-fl ow tuff s are typically found farther north in this part of southwestern utah. leach canyon formation th e leach canyon formation is a pinkish-brown, moderately welded (welding refers to the compaction of ash shards at the time of emplacement, resulting in a resistant rock even when welding is poor), moderately crystal-rich rhyolite ash-fl ow tuff that erupted from a caldera on the utah-nevada border nearly 24 million years ago. th e formation has a volume of at least 430 cubic miles (1800 km3) and a distribution of at least 6000 square miles (15,000 km2) (best and others, 1989b). for comparison, the volume of erupted products from the 1980 mount st. helens eruptions was about 1 cubic mile (4 km3). th e south side of brian head peak reveals the most complete section of the leach canyon formation on the markagunt plateau (biek and others, 2015). th ere, the classic three-part section of an ash-fl ow tuff is exposed, including an unwelded basal surge deposit, a thick vitrophyre (black glassy layer), and the moderately welded ash-fl ow tuff that caps brian head peak (fi gure 6). th e seldom exposed basal surge deposit is crystal-rich, tuff aceous sandstone with wavy and low-angle cross-bedding, resembling sand dunes. it records deposition by relatively low-concentration, turbulent density currents as the ash-fl ow tuff moved rapidly across the landscape. th e leach canyon formation contains abundant white or light-pink, collapsed pumice fragments and several percent rock fragments, many of which are reddish brown; phenocrysts of plagioclase, slightly less but subequal amounts of quartz and sanidine, and minor biotite, hornblende, fe-ti oxides, and a trace of pyroxene make up 25% to 35% of the rock. th e formation is about figure 5. southern great basin ash-fl ow tuff province that resulted from the middle cenozoic (36 to 18 ma) “ignimbrite fl are-up.” several ash-fl ow tuff s (ignimbrites) from the indian peak and caliente caldera complexes (blue lines) on the utah-nevada border spread eastward into southwest utah and as far east as the marysvale volcanic fi eld. th e coeval marysvale fi eld (purple, in southwest utah) and calderas (black) are shown to emphasize the contrasting dominance of andesitic lavas and mudfl ows (most of the marysvale fi eld) over more silica-rich ash-fl ow tuff s (most of the great basin). th is fi gure does not show pre-36 ma ash-fl ow tuff s and calderas from more northern igneous belts, nor does it show post-18 ma tuff s and calderas from more southern igneous belts. th e thick yellow line marks the western edge of the great basin altiplano; tuff s erupted from western nevada calderas fl owed mostly west down the west fl ank of the altiplano and are now preserved in exhumed paleovalleys across today’s sierra nevada mountains. th e western edge of precambrian continental basement is defi ned geochemically by the dashed 87sr/86sr = 0.706 line. from best and others (2013). 112°114°116°118°120° 42° 40° 38° 36° arizona st. george reno elko topographic barrier salt lake city austin tonopah ely 0 10050 kilometers 0 10050 miles �eld marysvale g r e a t n e va d a s ie r r a b a s i n p l at e a u c o l o r a d o las vegas fig_8.02_nevutah nevada utah cal i fornia 87 sr/ 86 sri = 0.706 andesitic lavasilicic ignimbrite western nv caldera central nv caldera indian peakcaliente caldera r.f. biek and p.d. rowley brian head peak 7 100 feet (30 m) thick where it forms the resistant caprock of brian head peak. interestingly, the leach canyon formation is petrographically and chemically similar to the haycock mountain tuff, a small, locally derived ash-flow tuff exposed to the east. this similarity led to confusion about the age of landslide deposits eventually called the markagunt gravity slide (biek and others, 2015). the 22.8 ma haycock mountain tuff unconformably overlies the markagunt gravity slide a dozen miles (20 km) east of brian head. the older, 23.8 ma leach canyon formation underlies and is locally involved in the slide. isom formation the 26 to 27 ma, crystal-poor, densely welded, trachydacitic ashflow tuff of the isom formation is present as far south as brian head peak and is spectacularly exposed for many miles along black ledge where it is about 350 feet (110 m) thick (figure 7). the isom formation is unusual in that it was so hot when erupted that it flowed like lava during its final stages (few tens of feet) of emplacement. for that reason it is commonly referred to as a tufflava or a rheomorphic ash-flow tuff—see, for example, anderson and rowley (1975, 2002), andrews and branney (2005), and geissman and others (2010). many isom outcrops reveal secondary flow characteristics, including flow breccias, contorted flow layering, and linear vesicles, such that the unit was considered a lava flow until mackin (1960) mapped its widespread distribution (300 cubic miles [1300 km3] today spread over an area of 9500 square miles [25,000 km2] [best and others, 1989a]) and found evidence of glass shards, thus showing its true ash-flow tuff nature. the isom is exposed at brian head peak where the lower part of the formation is classic tufflava about 80 feet (24 m) thick and the upper part is a flow breccia 60 to 90 feet (18-27 m) thick (biek and others, 2015). brian head formation the brian head formation is characterized by white volcaniclastic mudstone, siltstone, sandstone, volcanic ash, muddy limestone, and minor conglomerate and multi-hued chalcedony, and it is known for its propensity for landsliding. these strata, rich in volcanic ash, were deposited in low-relief fluvial, floodplain, and lacustrine environments; they record the inception of volcanism in southwest utah beginning about 37 million years ago (sable and maldonado, 1997) (figure 8). the base of the section is well exposed near the north rim overlook in cedar breaks national monument; there, a thin rhyolitic ash bed overlies a thin pebbly conglomerate likely equivalent to the conglomerate at boat mesa in bryce canyon national park. this ash bed yielded a u-pb age on zircon of 35.77 ± 0.28 ma, and several additional radiometric ages from brian head strata throughout the region show that it was deposited from about 37 to 33 million years ago (biek and others, 2015). thus, it is mostly late eocene in age, barely reaching into the oligocene. the brian head formation contains abundant trace fossils, including possible crayfish burrows and root traces (golder and wizevich, 2009; golder and others, 2009), but aside from its basal variegated interval it is surprisingly unfossiliferous (eaton and others, 1999). it also has colorful beds of chalcedony in various shades of white, gray, yellow, red, black, and brown, all typically with a white weathering rind. the chalcedony forms resistant beds as much as 10 feet (3 m) thick and is thought to have resulted from silicification of limestone beds (maldonado, 1995; sable and maldonado, figure 6. leach canyon formation on the south side of brian head peak. here, the classic three-part section of an ash-flow tuff is exposed, including an unwelded basal surge deposit, a thick vitrophyre, and moderately welded ash-flow tuff that caps brian head peak. a flow breccia of the isom formation (ti) is present in the lower left corner of the photograph. figure 7. a. the resistant isom formation at black ledge, looking south towards brian head peak. b. contorted flow layering of this densely welded ash-flow tuff, an example of its rheomorphic nature. a b 8 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 1997; schinkel, 2012). being resistant, the chalcedony commonly litters slopes developed on brian head strata; it was commonly used for tools and arrowheads by native americans. above all, however, the brian head formation is known for its swelling soils and for its susceptibility to landsliding. nearly all exposures on steep hillsides form large landslide complexes, including one that fi lls the 15-mile (24 km) length of yankee meadows graben, home to the resort town of brian head. th e formation is susceptible to landslides because of its abundant clay derived from weathered volcanic ash. a more complete section of the brian head formation is well exposed on the southwest fl ank of the sevier plateau (fi gure 9). th ere, a transitional interval as much as 160 feet (50 m) thick, not present elsewhere, of fi ne-grained, slope-forming sandstone, siltstone, and mudstone of red, pink, yellowish-brown, and purplish-gray hues forms the base of the formation. th is variegated interval yielded fossil turtles, charophytes (freshwater green algae), and fi sh suggestive of lacustrine environments of the duchesnean north american land mammal age (end of middle eocene) (feist and others, 1997; eaton and others, 1999; korth and eaton, 2004). structure th e markagunt plateau is uplift ed with respect to the great basin by high-angle normal faults, some of which create a series of horsts and grabens west and north of brian head peak. horsts are blocks that are upthrown by faults on each side, whereas grabens are blocks that are downthrown by faults on each side. th ese horsts and grabens step down from the plateau to the adjacent great basin (fi gure 10) (maldonado and others, 1997; biek and others, 2015). th e most prominent of these faults is the active hurricane fault at the base of the plateau, which has a downto-the-west vertical displacement of at least 6000 feet (1800 m) 113°00´ 112°30´ 38°00´ 37 °3 7´ 30 ˝ 0 0 5 miles 5 kilometers 143 143 148 14 r e d en o c h g r ab en h i l l s h o r s t northern cedar valley graben hur ric an e fa u l t zo ne iro n sp rin gs th ru st f a ul t p a r o w a n v a l l e y g r a b e n r ed h ills fault zo ne fi d d le r s c an yo n g r ab en m id dl e r id ge fa u l t m ap le r id g e h o r st jo ne s h ill fa ul t su m m it m o u n ta in g r ab en su m m it m ou nt ai n fa ul t parowan graben su ga rlo af m ou nt ai n fa ul t su ga rl oa f g ra be n na va jo r id ge fa ul t na va jo ri dg e ho rs t ratt les na ke can yo n fa ult yankee mead ow s g ra be n blac k le dg e fau lt m a r k a g u n t p l a t e a uparagonah parowan summit pa ra go na h fa u l t pa ro w an can yon h o r st iro n pe ak we st fa u lt ir on peak gra be n iro n peak eas t faul t w es t be ar va lle y fa ult ea st be ar va lle y fa ult be ar va ll ey graben location map normal fault, bar and ball on down-dropped side thrust fault, teeth on upper plate parowan fault 15 salt lake city brian head peak cbnm north view overlook ¯ figure 10. major faults of the western markagunt plateau and red hills, and named grabens (shaded) and horsts. cbnm = cedar breaks national monument. from biek and others (2015). figure 8. air-fall ash bed (white unit) in the brian head formation just south of haycock mountain, which yielded a u-pb age on zircon of 34.95 ± 0.83 ma. figure 9. exceptional exposures of the brian head formation on the southwest fl ank of the sevier plateau. here, brian head strata are divisible into four parts: (1) a basal variegated unit (below the hikers and so not visible), (2) a lower, light-gray, fi ne-grained volcaniclastic unit, which includes a thick, bluish-gray bentonitic mudstone at its base (tbh1), (3) a distinctive red-green-gray banded, fi ne-grained volcaniclastic unit (tbh2), and (4) a thick, upper volcaniclastic unit (tbh3). brian head strata are capped by mount dutton formation volcanic mudfl ow deposits (td), which here mark the base of the sevier gravity slide.. r.f. biek and p.d. rowley brian head peak 9 (hurlow, 2002; lund and others, 2007). overall, the collection of horsts and grabens forms a highly faulted relay ramp between the paragonah fault and the hurricane fault (fi gure 11). th e question of when this faulting began can be answered in part by the distribution of ash-fl ow tuff s. whereas modern basin-range extension that led to the topography we see today likely began about 10 to 12 million years ago in the cedar city area (rowley and others, 1981; hurlow, 2002), the initiation of such extension is much older. because ash-fl ow tuff s like the isom are widespread and were emplaced “in the blink of an eye,” they now serve as important timelines that help constrain structural interpretations of southwestern utah. th e eruption of the isom formation was followed by the eruption of three regionally distinctive ash-fl ow tuff s: the 23.8 ma leach canyon formation, the 22.8 ma bauers tuff member of the condor canyon formation, and the 22.0 ma harmony hills tuff , all part of the quichapa group. only the leach canyon is preserved as far east as panguitch lake; the bauers and harmony hills tuff s are restricted to the western part of the markagunt plateau. th e distribution of these ash-fl ow tuff s suggests that a west-facing topographic escarpment associated with early basin-range extension may have impeded eastward distribution of the latter two tuff s, as fi rst noted by rowley and barker (1978). furthermore, a graben north of nearby parowan juxtaposes bear valley and brian head strata that were later intruded by the 20 ma iron peak laccolith, showing that early extension was active prior to about 20 million years ago (biek and others, 2015). acknowledgments our knowledge of southwestern utah geology comes from combined decades of geologic mapping supported largely by the utah geological survey and u.s. geological survey. as the acknowledgments of our many published geologic maps attest, we are indebted to a great many people for their help over the years. th anks to grant willis, stephanie carney, and mike hylland (all with the utah geological survey) for their insightful reviews, and to jenny erickson (ugs) for draft ing the fi gures. references anderson, j.j., and rowley, p.d., 1975, cenozoic stratigraphy of southwestern high plateaus of utah, in anderson, j.j., rowley, p.d., fleck, r.j., and nairn, a.e.m., editors, cenozoic geology of southwestern high plateaus of utah: geological society of america special paper 160, p. 1–51. anderson, j.j., and rowley, p.d., 2002, th e oligocene isom formation, utah and nevada—a regional ash-fl ow tuff sheet containing fl uidal features of a lava fl ow [abs.]: geological society of america abstracts with programs, v. 34, no. 4, p. 9. andrews, g.d.m., and branney, m.j., 2005, folds, fabrics, and kinematic criteria in rheomorphic ignimbrites of the snake river plain, idaho —insights into emplacement and fl ow, in pederson, j., and dehler, c.m., editors, interior western united states: geological society of america field guide 6, p. 311–327. armstrong, r.l., ekren, e.b., mckee, e.h., and noble, d.c., 1969, space-time relations of cenozoic silicic volcanism in the great basin of the western united states: american journal of science, v. 267, p. 478–490. best, m.g., baar, d.l., christiansen, e.h., grommé, c.s., deino, a.l., and tingey, d.g., 2009, th e great basin altiplano during the middle cenozoic ignimbrite fl areup—insights from volcanic rocks: international geology review, v. 51, p. 589–633. best, m.g., christiansen, e.h., and blank, r.h., jr., 1989a, oligocene caldera complex and calc-alkaline tuff s and lavas of the indian peak volcanic fi eld, nevada and utah: geological society of america bulletin, v. 101, p. 1076–1090. best, m.g., christiansen, e.h., deino, a.l., grommé, c.s., hart, g.l., and tingey, d.g., 2013, th e 36-18 ma indian peak-caliente ignimbrite fi eld and calderas, southeastern great basin, usa— multicyclic super-eruptions: geosphere, v. 9, no. 4, p. 1–87. best, m.g., christiansen, e.h., deino, a.l., gromme, c.s., mckee, e.h., and noble, d.c., 1989b, excursion 3a—eocene through miocene volcanism in the great basin of the western united states: new mexico bureau of mines and mineral resources memoir 47, p. 91–133. biek, r.f., rowley, p.d., anderson, j.j., maldonado, f., moore, d.w., hacker, d.b., eaton, j.g., hereford, r., filkorn, h.f., and matyjasik, b., 2015, geologic map of the panguitch 30’ x 60’ quadrangle, garfi eld, iron, and kane counties, utah: utah geological survey map 270dm, 162 p., 3 plates, scale 1:62,500. cook, e.f., 1965, stratigraphy of tertiary volcanic rocks in eastern nevada: nevada bureau of mines report 11, 61 p. decelles, p.g., 2004, late jurassic to eocene evolution of the cordilleran thrust belt and foreland basin system, western u.s.a.: relay ramp hanging wall footwall markagunt plateau brian head peak cbnm overlook hurricane fault cedar valley parowan valleynormal fault tip dip figure 11. diagram of a relay ramp between parallel strands of a fault zone. th e ramp links displacement between the faults. cbnm = cedar breaks national monument. 10 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 american journal of science, v. 304, p. 105–168. dickinson, w.r., 2006, geotectonic evolution of the great basin: geosphere, v. 2, p. 353–368. eaton, j.g., hutchinson, j.h., holroyd, p.a., korth, w.w., and goldstrand, p.m., 1999, vertebrates of the turtle basin local fauna, middle eocene, sevier plateau, south-central utah, in gillette, d.d., editor, vertebrate paleontology in utah: utah geological survey miscellaneous publication 99-1, p. 463–468. feist, m., eaton, j.g., brouwers, e.m., and maldonado, f., 1997, significance of charophytes from the lower tertiary variegated and volcaniclastic units, brian head formation, casto canyon area, southern sevier plateau, southwestern utah, in maldonado, f., and nealey, l.d., editors, geologic studies in the basin and range–colorado plateau transition in southeastern nevada, southwestern utah, and northwestern arizona, 1995: u.s. geological survey bulletin 2153, p. 27–42. geissman, j.w., holm, d., harlan, s.s., and embree, g.f., 2010, rapid, high-temperature formation of large-scale rheomorphic structures in the 2.06 ma huckleberry ridge tuff, idaho, usa: geology, v. 38, p. 263–266. golder, k.b., and wizevich, m.c., 2009, wetland trace fossils in the paleogene brian head formation, southwest utah [abs.]: geological society of america abstracts with programs, v. 41, no. 3, p. 88. golder, k.b., wizevich, m.c., simpson, e.l., and storm, l.p., 2009, oligocene ichnofossils in non-marine limestone of the brian head formation, utah [abs.]: geological society of america abstracts with programs, v. 41, no. 7, p. 262. humphreys, e., 2009, relation of flat subduction to magmatism and deformation in the western united states, in kay, s.m., ramos, v.a., and dickinson, w.r., editors, backbone of the americas—shallow subduction, plateau uplift, and ridge and terrane collision: geological society of america memoir 204, p. 85–98. hurlow, h.a., 2002, the geology of cedar valley, iron county, utah, and its relation to ground-water conditions: utah geological survey special study 103, 74 p., 2 plates. korth, w.w., and eaton, j.g., 2004, rodents and a marsupial (mammalia) from the duschesnean (eocene) of the sevier plateau, utah, in dawson, m.r., and lillegraven, j.a., editors, fanfare for an uncommon paleontologist—papers on vertebrate evolution in honor of malcolm c. mckenna: carnegie museum of natural history bulletin, no. 36, p. 109–119. lund, w.r., hozik, m.j., and hatfield, s.c., 2007, paleoseismic investigation and long-term slip history of the hurricane fault in southwestern utah, paleoseismology of utah, v. 14: utah geological survey special study 119, cd, 81 p. mackin, j.h., 1960, structural significance of tertiary volcanic rocks in southwestern utah: american journal of science, v. 258, no. 2, p. 81–131. maldonado, f., 1995, decoupling of mid-tertiary rocks, red hills-western markagunt plateau, southwestern utah, in scott, r.b., and swadley, wc, editors, geologic studies in the basin and range-colorado plateau transition in southeastern nevada, southwestern utah, and northwestern arizona, 1992: u.s. geological survey bulletin 2056, p. 233–254. maldonado, f., sable, e.g., and nealey, d.l., 1997, cenozoic low-angle faults, thrust faults, and anastomosing high-angle faults, western markagunt plateau, southwestern utah, in maldonado, f., and nealey, l.d., editors, geologic studies in the basin and range-colorado plateau transition zone in southeastern nevada, southwestern utah, and northwestern arizona, 1995: u.s. geological survey bulletin 2153, p. 151–176. rowley, p.d., 1998, cenozoic transverse zones and igneous belts in the great basin, western united states—their tectonic and economic implications, in faulds, j.e., and stewart, j.h., editors, accommodation zones and transfer zones—the regional segmentation of the basin and range province: geological society of america special paper 323, p. 195–228. rowley, p.d., and barker, d.s., 1978, geology of the iron springs mining district, utah, in shawe, d.r., and rowley, p.d., editors, guidebook to mineral deposits of southwestern utah: utah geological association publication 7, p. 49–58. rowley, p.d., biek, r.f., sable, e.g., boswell, j.t., vice, g.s., hatfield, s.c., maxwell, d.j., and anderson, j.j., 2013, geologic map of the brian head quadrangle, iron county, utah: utah geological survey map 263dm, 38 p., 2 plates, scale 1:24,000. rowley, p.d., and dixon, g.l., 2001, the cenozoic evolution of the great basin area, u.s.a. —new interpretations based on regional geologic mapping, in erskine, m.c., faulds, j.e., bartley, j.m., and rowley, p.d., editors, the geologic transition, high plateaus to great basin a symposium and field guide (the mackin volume): utah geological association and pacific section of the american association of petroleum geologists, utah geological association publication 30, p. 169–188. rowley, p.d., steven, t.a., and mehnert, h.h., 1981, origin and structural implications of upper miocene rhyolites in kingston canyon, piute county, utah: geological society of america bulletin, part i, v. 92, p. 590–602. sable, e.g., and maldonado, f., 1997, the brian head formation (revised) and selected tertiary sedimentary rock units, markagunt plateau and adjacent areas, southwestern utah, in maldonado, f., and nealey, l.d., editors, geologic studies r.f. biek and p.d. rowley brian head peak 11 in the basin and range-colorado plateau transition zone in southeastern nevada, southwestern utah, and northwestern arizona, 1995: u.s. geological survey bulletin 2153, p. 7–26. schinkel, t., 2012, investigation of the origin of silicified layers within paleogene-aged volcaniclastic brian head formation, southern utah: new britain, central connecticut university, m.s. thesis, 100 p. stewart, j.h., and carlson, j.e., 1976, cenozoic rocks of nevada— four maps and a brief description of distribution, lithology, age, and centers of volcanism: nevada bureau of mines and geology map 52, scale 1:100,000. stewart, j.h., moore, w.j., and zietz, i., 1977, east-west patterns of cenozoic igneous rocks, aeromagnetic anomalies, and mineral deposits, nevada and utah: geological society of america bulletin, v. 88, p. 67–77. uga 50:11 geochemical studies of the green river formation in the piceance basin, colorado: ii. chemofacies from hierarchical cluster analysis off-canvas toggle member login donate about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle thisistheplace2.jpeg subpage1.jpeg subpage2.jpeg subpage3.jpeg uga 50:11 geochemical studies of the green river formation in the piceance basin, colorado: ii. chemofacies from hierarchical cluster analysis tengfei wu school of geosciences, university of oklahoma jeremy boak hurricane peak geosciences justin e. birdwell u.s. geological survey doi: https://doi.org/10.31711/ugap.v50i.115 abstract hierarchical cluster analysis (hca) was applied to a geochemical dataset representing the eocene green river formation in the piceance basin of colorado to identify chemofacies in core and outcrop samples from the basin margin and the basin center. the input dataset consisted of inductively coupled plasma optical emission spectroscopy and mass spectrometry and total organic carbon (toc) content analyses of 186 basin margin outcrop samples and 190 basin center core samples discussed in part 1 of this study (this volume). toc values and twenty-five major and trace elements were used as variables to define statistical clusters of samples for the overall dataset, for the two basin center cores, and for each separate core or outcrop dataset by hca applying euclidean distance and ward’s method algorithms. for each dataset, five cluster-defined chemofacies were identified. the chemofacies for each dataset show chemical affinities with five informally defined rock types– mudstone, marlstone, carbonate-rich mudstone, siliciclastic-rich mudstone/siltstone/sandstone, and na-rich (saline) mudstone, with each showing variations in toc content and abundance of redox sensitive minor and trace elements. a close relationship between enrichment of redox sensitive elements, particularly as and mo, and toc is identified in the basin center. whereas enrichment factors (relative to average shale) are relatively low for many period iv (piv) transition metals, as discussed in part 1 of this study, their consistent coherence in enrichment or depletion in hca-defined chemofacies demonstrates the expected relationship between redox state and organic richness. enrichment in piv transition metals also shows a correlation to enrichment in elements with affinity for siliciclastic sediment. enrichment/depletion among several groups of redox indicators is not everywhere consistent, with some chemofacies showing, for example, enrichment of piv transition metals and depletion of sulfur and arsenic. early timing of saline conditions in the basin margin is clearly displayed in the chemofacies log display, consistent with observations based on geochemical interpretations of concentrations and elemental ratios discussed in part 1. overall, the chemofacies are consistent with major mineralogical units and lake history stages defined in previous work, but provide more detail on the fluctuations in lake chemistry that occurred during deposition of green river oil shale in the piceance basin. order it ($5.00) cart your cart is empty login form login for discounted rates on document downloads and to access members only content username password show password remember me log in forgot your password? forgot your username? create an account utah geological association p.o. box 526356 salt lake city, ut 84152-0100  facebook  youtube © 2025 utah geological society | site by third sun member login off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf uga 50:10 geochemical studies of the green river formation in the piceance basin, colorado: i. major, minor, and trace elements off-canvas toggle member login donate about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle thisistheplace2.jpeg subpage1.jpeg subpage2.jpeg subpage3.jpeg uga 50:10 geochemical studies of the green river formation in the piceance basin, colorado: i. major, minor, and trace elements jeremy boak hurricane peak geosciences tengfei wu school of geosciences, university of oklahoma justin e. birdwell u.s. geological survey doi: https://doi.org/10.31711/ugap.v50i.114 abstract the eocene green river formation contains the largest oil shale deposits in the world and is a welldocumented example of a lacustrine depositional system. in addition, mineral resources associated with oil shale in the piceance basin nahcolite [nahco3] and dawsonite [naal(co3)(oh)2)] are of current and potential economic value, respectively. detailed geochemical analysis across the basin can aid in the understanding of the depositional environment, sedimentary processes, and water-chemistry evolution in this system. quantitative geochemical data for green river oil shale from the piceance basin of colorado were collected by inductively coupled plasma optical emission spectroscopy and mass spectrometry as part of this study. the basin margin is represented by samples from exposures at douglas pass (garfield county) and the basin center area is characterized by core samples from two drilled wells: the shell 23x-2 and john savage 24-1 (rio blanco county). major elements and groups of elements are used as proxies for clastic influx (si, al, k, ti), carbonate deposition (ca, mg), salinity (na), paleo-productivity (p), and redox state (fe, s), respectively. minor and trace elements reinforce observations based on major elements, including rb, zr, nb for clastic influx and mn, sr for carbonate. trace elements are used to characterize redox conditions (as, mo, u, v, co, ni, cu, zn) and salinity (rb/k, b/ga). chemical distinctions between the basin margin and the basin center, in terms of these components and total organic carbon concentrations, support the model of a permanently stratified lake through most of the depositional interval. a primary purpose of the study was to conduct more extensive sampling to confirm conclusions of a previous reconnaissance study. geochemical data from this study indicates elevated na around the basin margin occurring earlier than in the deeper basin. early in the history of lake uinta, the salinity may have been elevated first in the shallower marginal waters, due to increased evaporation, which then led to elevated salinity in the basin center through transport of saline density currents. other indicators of salinity (rb/k, b/ga) do not track na content in intervals where clay minerals are absent due to diagenetic alteration under hypersaline conditions but may be used to indicate the salinities at which authigenic na-bearing minerals begin to form. most na-rich samples show high proportions of clastic constituents (si, al, k, ti) compared to conventional carbonate constituents (ca, mg). redox-sensitive period iv transition metal elements (v, co, ni, cu, zn) show only local occurrence of significant enrichment relative to average shale abundances. analysis of fe/al ratios for this dataset suggests that the depletion of these elements may be related to source rocks depleted in mafic constituents, with apparent redox-related enrichments subdued by this effect. the basin margin samples reflect generally oxic bottom waters, with some intervals deposited under more reducing, possibly dysoxic to anoxic conditions. the basin center results indicate more reducing conditions, with mo and u enrichment factors suggesting operation of a particulate shuttle mechanism that scavenged mo on fe/mn-oxyhydroxides that redissolved at depth, with mo precipitating along with sulfides and/or organic matter at or near the sediment/water interface. order it ($5.00) cart your cart is empty login form login for discounted rates on document downloads and to access members only content username password show password remember me log in forgot your password? forgot your username? create an account utah geological association p.o. box 526356 salt lake city, ut 84152-0100  facebook  youtube © 2025 utah geological society | site by third sun member login off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf uga 50:13 spatiotemporal quantification of organic matter accumulation in the eocene green river formation, bridger basin, wyoming off-canvas toggle member login donate about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle thisistheplace2.jpeg subpage1.jpeg subpage2.jpeg subpage3.jpeg uga 50:13 spatiotemporal quantification of organic matter accumulation in the eocene green river formation, bridger basin, wyoming marshal tofte alan r. carroll department of geoscience, university of wisconsin-madison doi: https://doi.org/10.31711/ugap.v50i.117 abstract it has long been recognized that lakes can bury large amounts of organic carbon (corg) in their sediment, with important consequences for conventional and unconventional petroleum resources and potentially for the global carbon cycle. the detailed distribution of lacustrine organic carbon through space and time is important to understanding its commercial and climatic implications, but has seldom been documented in detail. the green river formation offers a unique opportunity to improve this understanding, due to extensive fischer assay analyses of its oil generative potential and to recently published radioisotopic age analyses of intercalat ed volcanic tuffs. fischer assay analyses reveal distinctly different patterns of organic matter enrichment that correlate with different lacustrine facies associations. histograms of oil generative potential for evaporative facies of the wilkins peak member exhibit an approximately exponential distribution. this pattern is interpret ed to result from episodic expansion and contraction of eocene lake gosiute across a low-gradient basin floor that experienced frequent desiccation. in contrast, histograms for fluctuating profundal facies of the upper rife bed of the tipton member and the lower laclede bed of the laney member exhibit an approximately normal or log normal distribution, with modes as high as 16–18 gallons per ton. this pattern is interpreted to reflect generally deeper conditions when the lake often intersected basin-bounding uplifts. within the bridger basin, burial of corg was greatest in the south during initial wilkins peak member deposition, reflecting greater rates of accommodation near the uinta uplift. the locus of corg burial shifted north during upper wilkins peak member deposition, coincident with a decrease in differential accommodation. corg burial during deposition of the upper rife and lower laclede beds was greatest in the southeast, due either to greater accommodation or localized influx of river-borne nutrients. average corg burial fluxes are consistently ~4-5 g/m2 yr for each interval, which is an order of magnitude less than fluxes reported for small holocene lakes in the northern hemisphere. maximum rates of corg burial during deposition of organic-rich mudstone beds (oil shale) were likely similar to holocene lakes however. deposition of carbonate minerals in the bridger basin resulted in additional, inorganic carbon burial. overall it appears that carbon burial by eocene lakes could have influenced the global carbon cycle, but only if synchronized across multiple lake systems. order it ($5.00) cart your cart is empty login form login for discounted rates on document downloads and to access members only content username password show password remember me log in forgot your password? forgot your username? create an account utah geological association p.o. box 526356 salt lake city, ut 84152-0100  facebook  youtube © 2025 utah geological society | site by third sun member login off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf uga 50:12 recurring lacustrine depositional successions in the wilkins peak member, green river formation: the basin-center evaporite perspective off-canvas toggle member login donate about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf off-canvas toggle thisistheplace2.jpeg subpage1.jpeg subpage2.jpeg subpage3.jpeg uga 50:12 recurring lacustrine depositional successions in the wilkins peak member, green river formation: the basin-center evaporite perspective elizabeth m. klonowski department of geological sciences and environmental studies, binghamton university tim k. lowenstein department of geological sciences and environmental studies, binghamton university alan r. carroll department of geoscience, university of wisconsin-madison m. elliot smith school of earth science and environmental sustainability, northern arizona university matteo paperini mine engineering dept., solvay chemicals, inc. jeffrey t. pietras department of geological sciences and environmental studies, binghamton university doi: https://doi.org/10.31711/ugap.v50i.116 abstract mineralogy, petrographic textures, and sedimentary structures from the world’s largest trona deposit, the wilkins peak member (wpm) of the early eocene green river formation (grf), bridger subbasin, wyoming, provide key data about depositional conditions and paleoenvironments. the 250 m-long wpm interval in the solvay s-34-1 drill core analyzed in this study contains a detailed record of sedimentation in the bridger subbasin at the deepest area of a hydrologically-closed basin during peak cenozoic atmospheric co2 concentrations. large accumulations of trona (na3(hco3)(co3)·2h2o), shortite (na2ca2(co3)3), northupite (na3mg (co3)2cl), and halite (nacl; now replaced by trona), occur in the lower half of the wpm. modern saline lake environments such as lake magadi, kenya, and the dead sea, israel-jordan, are useful analogues for interpreting paleolake conditions associated with evaporite deposition in the solvay s-34-1 core. solvay saline lake deposits are organized into meter-scale shallowing-upward successions, beginning with (1) oil shale overlain by (2) trona, in places interbedded with oil shale, followed by (3) peloidal dolomite grainstone and/or silty dolomitic mudstone, and (4) massive mudstone with disruption features or desiccation cracks, and/or siliciclastic sandstone with ripple cross-stratification. based on observations of modern hypersaline lake environments, wpm evaporite deposition at the basin depocenter is interpreted to be controlled by inflow water composition and volume, evaporative concentration, and seasonally-driven lake temperature fluctuations, resulting in recurrent patterns in evaporite mineralogies and textures. order it ($5.00) cart your cart is empty login form login for discounted rates on document downloads and to access members only content username password show password remember me log in forgot your password? forgot your username? create an account utah geological association p.o. box 526356 salt lake city, ut 84152-0100  facebook  youtube © 2025 utah geological society | site by third sun member login off-canvas toggle about publications geology of the intermountain west guidebooks uga geosites special publications uga 28 topical papers members member benefits become a member resources newsletter meeting minutes geologic road signs geology links uga constitution outreach student scholarship fund teacher of the year become a sponsor uga sponsors & affiliates lehi hintze award events contact ugf about ugf donate to ugf uga-geosite-kowallis-rock-canyon.indd 2 figure 2. looking down rock canyon to the west toward utah lake. the prominent folded rocks shown here are mississippian humbug formation, deseret limestone, and gardison formation. photo is courtesy of beau walker. figure 3. headline and part of the article from the provo daily herald, 29 july 1936, p. 1, reporting on the flood out of rock canyon the previous day. rain had begun in the mountains in the early afternoon. lewis richards, who was at his homestead in the canyon heard the roaring of the flood at about 2 p.m. and reported that, “it looked like the whole mountain had begun to move.” flooding also occurred from similar storms in other central utah communities the same day. figure 4. flood waters flowing out of rock canyon along temple view drive in provo in early june 1983. photo by bart kowallis. utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors bart j. kowallis and laura c. wald department of geological sciences, brigham young university provo, utah 84604 bkowallis@byu.edu cover image: looking down rock canyon to the west toward utah lake. rock canyon near provo, utah county: a geologic field laboratory 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e geosites refl ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: kowallis, b.j., and wald, l.c., 2019, rock canyon near provo, utah county—a geologic fi eld laboratory, in milligan, m., biek, r.f., inkenbrandt, p., and nielsen, p., editors, utah geosites: utah geological association publication 48, 15 p., https://doi.org/10.31711/geosites.v1i1.58. presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president b.j. kowallis and l.c. wald rock canyon 3 introduction rock canyon near provo, utah is an ideal outdoor laboratory. th e canyon has been known and explored for many years by scientists and students for its fascinating geology, biology, and botany. it is also a favorite location for rock climbers, hikers, and other outdoor enthusiasts. facilities near the mouth of the canyon including parking, restrooms, a lecture amphitheater, and a covered pavilion with picnic tables provide an ideal location for visitors (fi gure 1). geology is the focal point of this beautiful canyon with a history that stretches from the precambrian (about 700 million years ago) to the wasatch fault and lake bonneville, which covered much of western utah at its peak roughly 18,000 years ago (fi gure 2). excellent exposures of the rocks allow visitors to see features clearly and piece together the history of the canyon. th e oldest rocks are glacial deposits of the mineral fork tillite. th e tillite is overlain by a thick section of paleozoic rocks of cambrian to permian age, all of which have been deformed into an asymmetric, overturned fold formed during the sevier orogeny, a roughly 140 to 50 million year old mountain building event. th e upper reaches of the canyon were sculpted by glaciers during the pleistocene and deposits of the provo and bonneville levels of lake bonneville are found at the mouth of the canyon, now cut by a recent alluvial fan. also, at the mouth of the canyon are excellent exposures of features associated with the wasatch fault. visitors are asked to please respect the canyon’s natural beauty and minimize human impact by adhering to regulations and by hiking on designated paths. please be careful not to litter; remember, what you take in, you must also take out. th e mouth of rock canyon, including the pavilion, restrooms, and parking lot is maintained by the city of provo. th e amphitheater is available to the public for special programs or outdoor classroom use. th e remainder of the canyon lies within the uintah national forest and is under the jurisdiction of the pleasant grove ranger district. if you have any questions about the pavilion and park area please contact the provo city parks and recreation department. for questions regarding the canyon please contact the national forest service. although the canyon is beautiful, it also contains some hazards that visitors should be aware of. among the plants that thrive in the lower reaches of the canyon is poison ivy, a plant that causes a severe, blistering rash in most people. hikers should learn to recognize this plant before exploring the canyon. th e canyon is also home to rattlesnakes. th ese snakes will generally try to avoid humans, but if disturbed or cornered, they will strike and their bite is painful and dangerous. if bitten, you should seek immediate med1 map key major roads streets hiking trails parks n scale miles0 0.5 brigham young university university mall y mountain squaw peak y provo temple to i-15 riverside country club lavell edwards stadium rock canyon parking area & restrooms r o c k c a n y o n u nivers i ty parkw ay state street bu l ldog b lvd 9 0 0 e ast 2230 north t im p view d rive u n ive rsity a ve n u e u n ivers i ty a ven u e c an yo n r o ad no r t h te mple drive colum bia lane figure 1. index map showing the location of rock canyon and the major roads providing access to the canyon. the star shows the location of the parking area, rest rooms, and pavilion at the mouth of the canyon. figure 1. index map showing the location of rock canyon and the major roads providing access to the canyon. th e star shows the location of the parking area, rest rooms, and pavilion at the mouth the of the canyon. 2 figure 2. looking down rock canyon to the west toward utah lake. the prominent folded rocks shown here are mississippian humbug formation, deseret limestone, and gardison formation. photo is courtesy of beau walker. figure 3. headline and part of the article from the provo daily herald, 29 july 1936, p. 1, reporting on the flood out of rock canyon the previous day. rain had begun in the mountains in the early afternoon. lewis richards, who was at his homestead in the canyon heard the roaring of the flood at about 2 p.m. and reported that, “it looked like the whole mountain had begun to move.” flooding also occurred from similar storms in other central utah communities the same day. figure 4. flood waters flowing out of rock canyon along temple view drive in provo in early june 1983. photo by bart kowallis. figure 2. looking down rock canyon to the west toward utah lake. th e prominent folded rocks shown here are mississippian humbug formation, deseret limestone, and gardison formation. photo is courtesy of beau walker. 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 ical attention. finally, the canyon has many great places for rock climbing, but loose rock and improper gear or training can and has resulted in disaster. one or more serious accidents or deaths occurs each year in the canyon due to falls from rocky slopes or cliffs. early canyon exploration j. marinus jensen (1924) self-published a book entitled the early history of provo, utah. jensen reports that the first mention of rock canyon in local history is related to the difficult time the early settlers of fort utah (later named provo) had with the ute indians. the utes were understandably disturbed by the pioneers taking over their valley. the disputes peaked in 1850 when a battle broke out along the provo river between approximately 70 indians, led by big elk, and a militia of mormon men. for several days the battle continued until one evening the indians, many of whom were wounded, took flight. in their escape they split into two groups; one headed south toward spanish fork, and the other smaller group went toward rock canyon. big elk headed the rock canyon group; however, he was severely wounded and died near the mouth of the canyon. as the militia approached rock canyon in pursuit of the indians, it is rumored that big elk’s squaw attempted to climb the cliffs, but fell and was killed. squaw peak, near the mouth of the canyon, was supposedly named for the incident. frequent travel into rock canyon began after july of 1855 when “william m. wall was given a grant to build a road. on completion of the road he was to be permitted to charge fifty cents for each load of wood hauled out of the canyon. there was also a provision in the order to the effect that the grantee should allow all persons desiring to do so, to work on the road; and that they should receive, as compensation therefore, the right to haul out eight loads of wood for every day’s work performed.” no mention was made of any fee for geology excursions, rock collecting, or rock climbing; evidently they were not “popular” activities back then. due to the need for water and timber, the early settlers of utah valley exploited these resources and depleted them (bush, 1995). during the 1860’s, rock canyon was settled by a few pioneer families and rock canyon creek also provided the early settlers in provo with drinking water (jensen, 1924). some land was cleared in the canyon to use for farming including fruit orchards (bush, 1995). however, the land proved to be too hostile for the european farming methods, and was quickly abandoned. the canyon proved to be more amenable to raising cattle and sheep. livestock were taken up the canyon to graze during the summer and this quickly became so popular that overgrazing stripped the land of most of its vegetation (jensen, 1924; wilson, 1988). this in turn led to flooding where debris and water from the canyon flowed out into the valley often causing damage to farms and structures built in or near the mouth of the canyon. flood incidents occurred quite frequently between 1896 and 1984 (bush, 1995). one of the most disastrous floods occurred on 28 july 1936 (figure 3). starting in about 1925, citizen groups in the valley began to try and address the flooding problems, and with the help of the federal government, some progress was made, including the terracing of the high, steep, rocky slopes of the canyon by the civilian conservation corps (ccc) during 1933 and 1934 (bush, 1995). but despite these efforts, overgrazing in the canyon continued and it was not until after the flooding in three consecutive years in the 1950’s that the overgrazing problem was addressed (bush, 1995). the most recent serious flooding problems occurred in 1983 after a series of wetter than normal months beginning in september of the previous year. the wet cool weather persisted until late may and early june of 1983 when the temperatures warmed and flooding began. a provo daily herald headline on the front page of the may 30th edition that year read “crews sandbag temple view,” as the road out of rock canyon was temporarily turned into a river (figure 4). after the 1983 floods, the catchment basin at the mouth of the canyon was enlarged to guard against further problems. this basin now serves as a city park when not needed for flood control. 2 figure 2. looking down rock canyon to the west toward utah lake. the prominent folded rocks shown here are mississippian humbug formation, deseret limestone, and gardison formation. photo is courtesy of beau walker. figure 3. headline and part of the article from the provo daily herald, 29 july 1936, p. 1, reporting on the flood out of rock canyon the previous day. rain had begun in the mountains in the early afternoon. lewis richards, who was at his homestead in the canyon heard the roaring of the flood at about 2 p.m. and reported that, “it looked like the whole mountain had begun to move.” flooding also occurred from similar storms in other central utah communities the same day. figure 4. flood waters flowing out of rock canyon along temple view drive in provo in early june 1983. photo by bart kowallis. figure 3. headline and part of the article from the provo daily herald, 29 july 1936, p. 1, reporting on the flood out of rock canyon the previous day. rain had begun in the mountains in the early afternoon. lewis richards, who was at his homestead in the canyon heard the roaring of the flood at about 2 p.m. and reported that, “it looked like the whole mountain had begun to move.” flooding also occurred from similar storms in other central utah communities the same day. 2 figure 2. looking down rock canyon to the west toward utah lake. the prominent folded rocks shown here are mississippian humbug formation, deseret limestone, and gardison formation. photo is courtesy of beau walker. figure 3. headline and part of the article from the provo daily herald, 29 july 1936, p. 1, reporting on the flood out of rock canyon the previous day. rain had begun in the mountains in the early afternoon. lewis richards, who was at his homestead in the canyon heard the roaring of the flood at about 2 p.m. and reported that, “it looked like the whole mountain had begun to move.” flooding also occurred from similar storms in other central utah communities the same day. figure 4. flood waters flowing out of rock canyon along temple view drive in provo in early june 1983. photo by bart kowallis. figure 4. flood waters flowing out of rock canyon along temple view drive in provo in early june 1983. photo by bart kowallis. b.j. kowallis and l.c. wald rock canyon 5 3 rock canyon, provo utah 0-50 0-200 0-13,000 alluvium, colluvium, soil lake bonneville deposits buried valley fill in utah valley includes pre-bonneville alluvial & lake deposits with some interbeds of volcanic tuff (see cook et al., 1997 for bouguer gravity map) 820010,200 300 bear canyon formation (atokan-desmoinesian) shingle mill ls (lower missourian) wallsburg ridge formation (missourian-virgilian) granger mountain formation (wolfcampian) 3250 bridal veil limestone 1245 manning canyon shale 1650 upper limestone member 1800 long trail sh mbr great blue ls topliff ls mbr humbug formation deseret limestone gardison limestone fitchville dolomite maxfield limestone ophir formation tintic quartzite mineral fork tillite big cottonwood formation 300 700 540 840 550 260 600 0-250 1100 145 775 515 510 505 360 350 345 335 330 325 320 315 310 305 300 295 290 285 15 10 1 700 1350+ o qu ir rh g ro up 40007000 wasatch fault zone separates utah valley fill from paleozoic bedrock of wasatch mountains shown below 0-10,000 yrs old 10,000-30,000 yrs old gulf oil-banks #1 23-8s-2e 13,000 ft deep, bottomed in beds containing late miocene pollen 10 m.y. + 1000’ sandstone 1200’ curry peak siltstone facies mostly fine sandstn with some siltstone interbeds fine grained sandstone with heavy minerals from uncompahgre source 1400’ curry peak siltstone facies mostly fine grained sandstone with fewer limestone interbeds than same interval in oquirrh mountains equivalent to jordan & commercial lst at bingham morrowan fossils are abundant brachiopods corals, bryozoans deep water facies forms strike valley taphrognathus spathognathodus regional unconformity orange, highly fractured in rock canyon white dolostone in rock canyon 6 feet of pinyon pk fm at base olive-drab, phyllitic diamictite purple & tan argillite in slate cny only in subsurface in rock cny triticites eowaeringella fusulina profusulinella dictyoclostus cravenoceras lepidodendron apatognathus fenestella syringopora, gnathodus glossopleura schwagerina pseudoschwagerina thickness ft.age notesformations age millions years q m io -p li o pe r m ia n pe n n sy lv a n ia n m is si ss ip pi a n d pr ot c a m b figure 5. stratigraphic column for the rock canyon area modified from hintze and kowallis (2009). the big cottonwood formation shown on the column is not exposed in the canyon, but it is just a couple miles south in slate canyon. also, the mio-pliocene sediments are only in the subsurface of the valley deposits. this thick pile of sediments in the valley are evidence of the amount of uplift and erosion that has occurred along the wasatch fault zone over the last 10-12 million years. in the column of notes are listed a few of the fossils (in italics) that occur in these rocks. figure 5. stratigraphic column for the rock canyon area modifi ed from hintze and kowallis (2009). th e big cottonwood formation shown on the column is not exposed in the canyon, but it is just a couple miles south in slate canyon. also, the mio-pliocene sediments are only in the subsurface of the valley deposits. th is thick pile of sediments in the valley are evidence of the amount of uplift and erosion that has occurred along the wasatch fault zone over the last 10-12 million years. in the column of notes are listed a few of the fossils (in italics) that occur in these rocks. 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 4 qlsqa mineral fork tillite tintic quartzite ophir formation maxfield limestone fitchville dolomite gardison limeston deseret limestone humbug formation great blue limestone prebonneville deposits lake bonneville deposits stream (qal) fan (qaf) colluvium (qac) deposits landslide deposits qlbtqafmgbmhmdmgdmfcmcoctpcmf n scale miles 0 0.50.40.3 0.20.1 figure 6. geologic map of the rock canyon area showing the location of the byu campus in relation to rock canyon reproduced from hintze (1978). the current parking lot, rest rooms, and pavilion are located approximately where the words “rock canyon” are on the map. figure 7. close-up view of mineral fork tillite showing small pebbles of quartzite and dolomite floating in a fine-grained matrix that has been metamorphosed producing a secondary foliation that in places becomes a crenulation cleavage. land land stage 2 stage 1 ice ice ew ? ? sea sea figure 8. two-stage model for deposition of the mineral fork tillite proposed by matsch and ojakangas (1991). stage 1 was a deposition in a terrestrial glacial setting while stage 2 was a marginal marine glacial setting. arrows indicate input of sediment into the deposits. figure 6. geologic map of the rock canyon area showing the location of the byu campus in relation to rock canyon reproduced from hintze (1978). th e current parking lot, rest room, and pavilion are located approximately where the words “rock canyon” are on the map. geologic history of rock canyon th e rocks exposed in rock canyon range in age from neoproterozoic to recent (fi gures 5 and 6; baker, 1972; hintze, 1978; davis, 1983). th e oldest rocks are exposed near the mouth of the canyon while the youngest bedrock units are found on the high peaks (cascade mountain, y mountain, and provo peak) at the head of the canyon. neoproterozoic rocks (precambrian) th e oldest rocks found in rock canyon are diamictites of the mineral fork tillite that are well exposed on the north canyon wall near the water storage tank (fi gure 6). th e outcrops are composed of cobbles and pebbles of quartzite, dolomite, and limestone (with rare clasts of other rock types) surrounded by a fi ne-grained matrix (baker 1973), which has been weakly metamorphosed producing a secondary foliation and phyllitic character in the matrix (fi gure 7). over the years, there has been some debate as to the origin of these deposits, but the general consensus currently is that they were deposited by ancient glaciers (varney, 1976; ojakangas and matsch, 1980; christie-blick, 1983; crittenden and others, 1983; link and others, 1994; christie-blick, 1997). ojakangas and matsch (1980) and matsch and ojakangas (1991) proposed that the lower part of the formation was deposited in a terrestrial glacial environment, while the upper part was deposited in a marginal-marine glacial environment (fi gure 8), and that the entire sequence was deposited by glaciers in a more temperate environment rather than a polar environment. christie-blick (1982) questioned the idea of a terrestrial glacial environment, instead preferring a submarine, grounded-ice environment. 4 qlsqa mineral fork tillite tintic quartzite ophir formation maxfield limestone fitchville dolomite gardison limeston deseret limestone humbug formation great blue limestone prebonneville deposits lake bonneville deposits stream (qal) fan (qaf) colluvium (qac) deposits landslide deposits qlbtqafmgbmhmdmgdmfcmcoctpcmf n scale miles 0 0.50.40.3 0.20.1 figure 6. geologic map of the rock canyon area showing the location of the byu campus in relation to rock canyon reproduced from hintze (1978). the current parking lot, rest rooms, and pavilion are located approximately where the words “rock canyon” are on the map. figure 7. close-up view of mineral fork tillite showing small pebbles of quartzite and dolomite floating in a fine-grained matrix that has been metamorphosed producing a secondary foliation that in places becomes a crenulation cleavage. land land stage 2 stage 1 ice ice ew ? ? sea sea figure 8. two-stage model for deposition of the mineral fork tillite proposed by matsch and ojakangas (1991). stage 1 was a deposition in a terrestrial glacial setting while stage 2 was a marginal marine glacial setting. arrows indicate input of sediment into the deposits. figure 7. close-up view of mineral fork tillite showing small pebbles of quartzite and dolomite fl oating in a fi ne-grained matrix that has been metamorphosed producing foliation that in places becomes a crenulation cleavage. b.j. kowallis and l.c. wald rock canyon 7 cambrian to devonian rocks th e cambrian was a time of shallow seas and shorelines in what is now utah. th e oldest cambrian rock unit in the canyon is the tintic quartzite. it is an orangeto tan-weathering quartzite that is light greenish gray on unweathered surfaces. th e sandstone from which the quartzite originated was composed of fairly clean, rounded to subrounded quartz grains of fi neto medium-grain size with local coarse-grained interbeds (baker and others, 1949; lochman-balk, 1976). cross-bed sets up to 3 to 4 feet (1 m) thick can be observed in this rock unit. some layers contain abundant well-rounded pebbles of quartzite or vein quartz (fi gure 9) and at several localities, trace fossil burrows (called skolithos) occur in the outcrops (fi gure 10; see also peterson and clark, 1974). th e tintic quartzite unit outcrops as highly-fractured, rough ledges and cliff s in the lower part of the canyon (fi gure 11). at the mouth of the canyon, the tintic quartzite is gently dipping, but a short distance into the canyon the unit is vertical to slightly overturned. it is in some of these vertical quartzite beds where rock climbing has become quite popular. toward the top of the tintic quartzite the unit interfi ngers with the overlying ophir shale (baker, 1972, 1973). like the tintic quartzite, the ophir shale has been slightly metamorphosed and would be more properly classifi ed as a phyllite rather than a shale. poorly preserved fragments and pieces of trilobites may be found in the ophir shale. th e fi nal cambrian unit exposed in rock canyon is the maxfi eld limestone (fi gures 5 and 6). th e lower two-thirds of this formation consists of a mottled darkto medium-gray limestone, while the upper third is a mediumto very light-gray banded dolomite (granger and other, 1952; hintze, 1978) fossils are sparse in the maxfi eld limestone, but some trilobite, gastropod, and bryozoan fragments have been found, as well as conodont microfossils (hintze, 1978; clark and others, 2014). 5 figure 9. small pebbles of dark quartzite and white vein quartz within the tintic quartzite. weathered (orange-tan, lower right) and unweathered (light gray) surfaces of this unit are shown. figure 10. rusty-colored skolithos burrows in the tintic quartzite. figure 11. squaw peak as seen from north temple drive. the orangish tan, gently dipping tintic quartzite is distinctive at the base of the peak and overlain by gray paleozoic dolomites and limestones. figure 12. cambrian depositional system (modified from hintze and kowallis, 2009). outer shelf west east sea level carbonate platform lagoon older rocks d olomite l i m e s t o n e s h a l e s h a l e s a n d f i t c h v i l l e d o l o . f i t c h v i l l e d o l o . p i n y o n p e a k f m . p i n y o n p e a k f m . ? ? m a x f i e l d l s t . m a x f i e l d l s t . 1 m 1 m figure 13. contact between the cambrian maxfield limestone and devonian fitchville dolomite in rock canyon with a few feet of pinyon peak formation between them (modified from clark and others, 2014). figure 9. small pebbles of dark quartzite and white vein quartz withing the tintic quartzite. weathered (orange-tan, lower right) and unweathered (light gray) surfaces of this unit are shown. th e cambrian depositional system is represented in fi gure 12 with shoreline sands eastward of a broad subtidal carbonate platform and deeper water slope deposits to the west (hintze and kowallis, 2009). in the rock canyon area, the tintic quartzite was deposited as shoreline and near shore sands. th en as sea level rose, the shoreline moved farther east. first the lagoonal muds of the ophir shale and then the carbonates of the maxfi eld limestone were deposited atop the sand due to a major pulse of sea level rise. 5 figure 9. small pebbles of dark quartzite and white vein quartz within the tintic quartzite. weathered (orange-tan, lower right) and unweathered (light gray) surfaces of this unit are shown. figure 10. rusty-colored skolithos burrows in the tintic quartzite. figure 11. squaw peak as seen from north temple drive. the orangish tan, gently dipping tintic quartzite is distinctive at the base of the peak and overlain by gray paleozoic dolomites and limestones. figure 12. cambrian depositional system (modified from hintze and kowallis, 2009). outer shelf west east sea level carbonate platform lagoon older rocks dolomite l i m e s t o n e s h a l e s h a l e s a n d f i t c h v i l l e d o l o . f i t c h v i l l e d o l o . p i n y o n p e a k f m . p i n y o n p e a k f m . ? ? m a x f i e l d l s t . m a x f i e l d l s t . 1 m 1 m figure 13. contact between the cambrian maxfield limestone and devonian fitchville dolomite in rock canyon with a few feet of pinyon peak formation between them (modified from clark and others, 2014). figure 10. rusty-colored skolithos burrow in the tintic quartzite. 5 figure 9. small pebbles of dark quartzite and white vein quartz within the tintic quartzite. weathered (orange-tan, lower right) and unweathered (light gray) surfaces of this unit are shown. figure 10. rusty-colored skolithos burrows in the tintic quartzite. figure 11. squaw peak as seen from north temple drive. the orangish tan, gently dipping tintic quartzite is distinctive at the base of the peak and overlain by gray paleozoic dolomites and limestones. figure 12. cambrian depositional system (modified from hintze and kowallis, 2009). outer shelf west east sea level carbonate platform lagoon older rocks d olomite l i m e s t o n e s h a l e s h a l e s a n d f i t c h v i l l e d o l o . f i t c h v i l l e d o l o . p i n y o n p e a k f m . p i n y o n p e a k f m . ? ? m a x f i e l d l s t . m a x f i e l d l s t . 1 m 1 m figure 13. contact between the cambrian maxfield limestone and devonian fitchville dolomite in rock canyon with a few feet of pinyon peak formation between them (modified from clark and others, 2014). figure 11. squaw peak as seen from north temple drive. th e orangish tan, gently dipping tintic quartzite is distinctive at the base of the peak and overlain by gray paleozoic dolomite and limestone beds. 4 qlsqa mineral fork tillite tintic quartzite ophir formation maxfield limestone fitchville dolomite gardison limeston deseret limestone humbug formation great blue limestone prebonneville deposits lake bonneville deposits stream (qal) fan (qaf) colluvium (qac) deposits landslide deposits qlbtqafmgbmhmdmgdmfcmcoctpcmf n scale miles 0 0.50.40.3 0.20.1 figure 6. geologic map of the rock canyon area showing the location of the byu campus in relation to rock canyon reproduced from hintze (1978). the current parking lot, rest rooms, and pavilion are located approximately where the words “rock canyon” are on the map. figure 7. close-up view of mineral fork tillite showing small pebbles of quartzite and dolomite floating in a fine-grained matrix that has been metamorphosed producing a secondary foliation that in places becomes a crenulation cleavage. land land stage 2 stage 1 ice ice ew ? ? sea sea figure 8. two-stage model for deposition of the mineral fork tillite proposed by matsch and ojakangas (1991). stage 1 was a deposition in a terrestrial glacial setting while stage 2 was a marginal marine glacial setting. arrows indicate input of sediment into the deposits. figure 8. two-stage model for deposition of the mineral fork tillite proposed by matsch and ojakangas (1991). stage 1 was a deposition in a terrestrial glacial setting while stage 2 was a marginal marine glacial setting. arrows indicate input of sediment into the deposits. 4 qlsqa mineral fork tillite tintic quartzite ophir formation maxfield limestone fitchville dolomite gardison limeston deseret limestone humbug formation great blue limestone prebonneville deposits lake bonneville deposits stream (qal) fan (qaf) colluvium (qac) deposits landslide deposits qlbtqafmgbmhmdmgdmfcmcoctpcmf n scale miles 0 0.50.40.3 0.20.1 figure 6. geologic map of the rock canyon area showing the location of the byu campus in relation to rock canyon reproduced from hintze (1978). the current parking lot, rest rooms, and pavilion are located approximately where the words “rock canyon” are on the map. figure 7. close-up view of mineral fork tillite showing small pebbles of quartzite and dolomite floating in a fine-grained matrix that has been metamorphosed producing a secondary foliation that in places becomes a crenulation cleavage. land land stage 2 stage 1 ice ice ew ? ? sea sea figure 8. two-stage model for deposition of the mineral fork tillite proposed by matsch and ojakangas (1991). stage 1 was a deposition in a terrestrial glacial setting while stage 2 was a marginal marine glacial setting. arrows indicate input of sediment into the deposits. 8 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 mississippian, pennsylvanian, and permian rocks th e rock canyon area remained at or near sea level for the next 100 million years, throughout the mississippian, pennsylvanian, and permian periods. however, in the mississippian, an area of increased subsidence developed on the continental shelf margin in the area of northwestern and north-central utah, which persisted through the permian. th is basin, where thousands of feet of sediment accumulated, is called the oquirrh basin (jordan and douglass, 1980: erskine, 1997) and was part of a series of basins and uplift s that formed the ancestral rocky mountains (fi gure 14; lee, 1918; ver wiebe, 1930; kluth, 1986; ye and others, 1996). in the rock canyon area, deposits into the oquirrh basin during this time period totaled over 20,000 feet (6000 m) or approximately four miles (6.5 km) of strata (baker, 1972; hintze, 1978). th ese rocks (including the gardison limestone, deseret limestone, humbug formation, great blue limestone, manning canyon shale, and oquirrh group) make up the bulk of the bedrock located from about one mile (1.6 km) into the canyon to the tops of the highest peaks surrounding the canyon (fi gure 15). limestones and sandstones of the humbug formation form the capping rocks atop squaw peak, while the manning canyon shale forms the valley and slopes below cliff s of the oquirrh group through which the squaw peak/rock canyon picnic area road winds. as you hike through the canyon, you may fi nd some marine fossils (corals, snails, bivalves, etc.) within these rock formations as well as darker and harder lumps of chert. limestones will fi zz with the applicafollowing cambrian time, a major period of erosion and non-deposition occurred in the rock canyon area due to a drop in sea level. th is unconformity disappears to the west where deposition was still occurring. beginning about 20 to 25 miles (30-40 km) west of rock canyon, cambrian-age strata become thicker. in addition, ordovician rocks, which are absent in rock canyon, appear and also thicken to the west. th e area between these thicker deposits off to the west and the thinner deposits to the east is called the “hinge line” or “wasatch line” and marks the transition from the stable shallow continental shelf margin to a more rapidly subsiding part of the continental margin called a miogeosyncline (kay, 1951; bissell, 1974, stokes, 1976). although rocks of ordovician and silurian age are not currently found in rock canyon, thin deposits of rocks of these ages likely once existed in this area, but during the early part of the devonian period they were eroded away. later in the devonian, the ocean again covered the area and sediments were deposited atop the erosional surface. th ese devonian-age sediments are relatively thin, accumulating only to a thickness of about 250 feet (75 m) in the rock canyon area (fi gures 5 and 6) and are called the fitchville dolomite or fitchville formation (baker, 1972; hintze, 1978; hintze and kowallis, 2009). th e fitchville dolomite consists of about 20 feet (6 m) of dolomitic sandstone at the base overlain by relatively unfossiliferous lightto medium-gray, vuggy dolomite (hintze, 1978). clark and others (2014) found that the lowermost few feet of the fitchville formation in rock canyon contain fossil conodonts (microscopic fossils) that correlate this part of the formation to an older devonian formation called the pinyon peak formation (fi gure 13), which outcrops in thicker beds to the west of rock canyon. 5 figure 9. small pebbles of dark quartzite and white vein quartz within the tintic quartzite. weathered (orange-tan, lower right) and unweathered (light gray) surfaces of this unit are shown. figure 10. rusty-colored skolithos burrows in the tintic quartzite. figure 11. squaw peak as seen from north temple drive. the orangish tan, gently dipping tintic quartzite is distinctive at the base of the peak and overlain by gray paleozoic dolomites and limestones. figure 12. cambrian depositional system (modified from hintze and kowallis, 2009). outer shelf west east sea level carbonate platform lagoon older rocks dolomite l i m e s t o n e s h a l e s h a l e s a n d f i t c h v i l l e d o l o . f i t c h v i l l e d o l o . p i n y o n p e a k f m . p i n y o n p e a k f m . ? ? m a x f i e l d l s t . m a x f i e l d l s t . 1 m 1 m figure 13. contact between the cambrian maxfield limestone and devonian fitchville dolomite in rock canyon with a few feet of pinyon peak formation between them (modified from clark and others, 2014). figure 12. cambrian depositional system (modifi ed from hintze and kowallis, 2009). 5 figure 9. small pebbles of dark quartzite and white vein quartz within the tintic quartzite. weathered (orange-tan, lower right) and unweathered (light gray) surfaces of this unit are shown. figure 10. rusty-colored skolithos burrows in the tintic quartzite. figure 11. squaw peak as seen from north temple drive. the orangish tan, gently dipping tintic quartzite is distinctive at the base of the peak and overlain by gray paleozoic dolomites and limestones. figure 12. cambrian depositional system (modified from hintze and kowallis, 2009). outer shelf west east sea level carbonate platform lagoon older rocks dolomite l i m e s t o n e s h a l e s h a l e s a n d f i t c h v i l l e d o l o . f i t c h v i l l e d o l o . p i n y o n p e a k f m . p i n y o n p e a k f m . ? ? m a x f i e l d l s t . m a x f i e l d l s t . 1 m 1 m figure 13. contact between the cambrian maxfield limestone and devonian fitchville dolomite in rock canyon with a few feet of pinyon peak formation between them (modified from clark and others, 2014). figure 13. contact between the cambrian maxfi eld limestone and devonian fitchville dolomite in rock canyon with a few feet of pinyon peak formation between them (modifi ed from clark and others, 2014). 6 figure 15. fall colors grace the maples growing in the slope-forming manning canyon shale (mississippian) at the base of a cliff on cascade mountain. the cliffs of cascade mountain are formed from the lightand medium-gray beds of the bridal veil limestone with tree covered crags above the cliffs of the light-tan sandstones of the bear canyon formation, both part of the oquirrh group (pennsylvanian). figure 14. middle paleozoic basins and uplifts showing the approximate location of rock canyon in the oquirrh basin (modified from hintze and kowallis, 2009). this period of uplift and basin subsidence formed mountains that are called the ancestral rocky mountains. the dashed line shows the location of the oquirrh basin at the time it formed and deposits accumulated. later these rocks were thrust to the east during the cretaceous sevier orogeny. o q u i r r h b a s i n p a r a d o x b a s i n e a g l e c e n t r a l t a o s b a s i n f r o n t r a n g e u p l i f t rock canyon emery uplift z u n i u p l i f t h o l b r o o k b a s i n f o u n t a i n b a s i n u n c o m p a h g r e u p l i f t wyoming new mexico idaho colorado u t a h n e v a d a a r i z o n a figure 14. middle paleozoic basins and uplift s showing the approximate location of rock canyon in the oquirrh basin (modifi ed from hintze and kowallis, 2009). th is period of uplift and basin subsidence formed mountains that are called the ancestral rocky mountains. th e dashed line shows the location of the oquirrh basin at the time it formed and deposits accumulated. later these rocks were thrust to the east during the cretaceous sevier orogeny. b.j. kowallis and l.c. wald rock canyon 9 tion of weak hydrochloric acid, while dolostones will only fizz if you first scratch the surface or powder some of the rock. chert will not fizz and is much harder than limestone or dolostone. mesozoic mountain building – the sevier orogeny following the end of the paleozoic era, utah ceased to be the western margin of the north american continent. during a long span of approximately 150 million years, fragments of continental and oceanic plates collided with western north america adding much of what is now nevada, california, oregon, and washington. the collision of these fragments caused uplift and deformation in the rocks that had previously been deposited during the precambrian and paleozoic. this mountain building event, called the sevier orogeny, created a high range of mountains in central utah that we call the sevier fold and thrust belt (armstrong, 1968; decelles and coogan, 2006). erosion of these mountains left thick deposits of sandstone, conglomerate, and shale off to the east (young, 1966). none of this erosional debris is visible in rock canyon, but the deposits created by the erosion of the sevier mountains can be seen in spanish fork canyon a few miles to the south. what is visible in rock canyon are the folds, faults, and fractures created when the precambrian and paleozoic rocks, which had been deposited several tens of miles to the west of rock canyon, were folded and thrust up over younger rocks (figures 16, 17 and 18). the thrust fault and younger rocks underneath the thrust are not exposed in rock canyon but they are exposed to the south near the town of nephi, utah and to the east in spanish fork canyon. as the strata were folded and fractured, some of the rock units were deformed more plastically than others. the ophir shale, for example, is not highly fractured and has been substantially thickened in some parts of the fold and thinned or completely squeezed out of other parts (figure 18). this is similar to what happens when you squeeze or fold a tube of toothpaste—the paste gets thinned and thickened in different parts of the tube. on the other hand, units like the tintic quartzite behaved in a brittle manner, fracturing into pieces rather than flowing. figure 18 shows the major fractures and faults in rock units of rock canyon. dense fracturing and faulting occur in the tintic quartzite compared to the other stratigraphic units partly because it was in the core of the fold where deformation was most intense and partly because the quartzite is simply more brittle than the other rock units (figures 18 and 19). mountain building ended about 40 million years ago and over the next 20 million years the sevier fold and thrust belt collapsed and eroded, but left behind the structures formed during mountain building as well as sediments shed to the east off of the mountains as they eroded away. 6 figure 15. fall colors grace the maples growing in the slope-forming manning canyon shale (mississippian) at the base of a cliff on cascade mountain. the cliffs of cascade mountain are formed from the lightand medium-gray beds of the bridal veil limestone with tree covered crags above the cliffs of the light-tan sandstones of the bear canyon formation, both part of the oquirrh group (pennsylvanian). figure 14. middle paleozoic basins and uplifts showing the approximate location of rock canyon in the oquirrh basin (modified from hintze and kowallis, 2009). this period of uplift and basin subsidence formed mountains that are called the ancestral rocky mountains. the dashed line shows the location of the oquirrh basin at the time it formed and deposits accumulated. later these rocks were thrust to the east during the cretaceous sevier orogeny. o q u i r r h b a s i n p a r a d o x b a s i n e a g l e c e n t r a l t a o s b a s i n f r o n t r a n g e u p l i f t rock canyon emery uplift z u n i u p l i f t h o l b r o o k b a s i n f o u n t a i n b a s i n u n c o m p a h g r e u p l i f t wyoming new mexico idaho colorado u t a h n e v a d a a r i z o n a figure 15. fall colors grace the maples growing in the slope-forming manning canyon shale (mississippian) at the base of a cliff on cascade mountain. the cliffs of cascade mountain are formed from the lightand medium-gray beds of the bridal veil limestone with tree covered crags above the cliffs of the light-tan sandstones of the bear canyon formation, both part of the oquirrh group (pennsylvanian). 10 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 7 oquirrh formation manning canyon shale great blue limestone humbug formation gardison limestone tintic quartzite maxfield limestone 1 mi sea level -1 mi -2 mi 2 mi 3 mi -3 mi (c) post wasatch fault development (a) pre-deformation (b) during thrusting great blue limestone humbug formation gardison limestone tintic quartzite maxfield limestone great blue limestone great blue limestone humbug formation gardison limestone tintic quartzite maxfield limestone oquirrh formation oquirrh formation oquirrh formation valley fill sediments utah valley squaw peak rock canyon manning canyon shale great blue limestone humbug formation gardison limestone tintic quartzite maxfield limestone scale miles 0 543 21 1 mi 1 mi figure 16. three stages in the development of the current geology of rock canyon. a) sedimentary rock units prior to the sevier orogenic deformation. b) deformation of the rocks due to the sevier orogeny causing rock layers to be thrust eastward several 10’s of kilometers and to be folded and faulted. c) formation of the wasatch fault and infilling of utah valley with material eroded from the uplifted side of the fault. the approximate position of squaw peak is shown by the dashed red line. modified from wald (2007). figure 16. th ree stages in the development of the current geology of rock canyon. a) sedimentary rock units prior to the sevier orogenic deformation. b) deformation of the rocks due to the sevier orogeny causing rock layers to be thrust eastward several 10’s of kilometers and to be folded and faulted. c) formation of the wasatch fault and infi lling of utah valley with material eroded from the uplift ed side of the fault. th e approximate position of squaw peak is shown by the dashed red line. modifi ed from wald (2007). b.j. kowallis and l.c. wald rock canyon 11 the wasatch mountains and wasatch fault beginning about 20 million years ago western north america began to uplift , extend and pull apart due to warm rocks in the earth’s mantle rising buoyantly beneath the area. th is extension and uplift of the west continues up to the present day and is responsible for the formation of the basin and range province (hintze and kowallis, 2009). in the rock canyon area, this extension is responsible for the formation of the wasatch mountains and utah valley. a major fault zone, called the wasatch fault, separates the uplift ed mountain block from the down-dropped valley. over the last 10 to 11 million years the mountains have been going up at a rate of about 0.6 to 1.0 mm/yr relative to the valley (naeser and others, 1983; parry and bruhn, 1987; kowallis and others, 1990). th e uplift along the wasatch fault causes stress to build up within the earth and when the rocks cannot store up any more stress, it is released in the form of earthquakes. th e fault is composed of about 10 segments that break somewhat independently of each other (fi gure 20; machette and others, 1991). trenches dug across the segments of the wasatch fault, including one at the mouth of rock canyon, have shown that earthquakes of moment magnitude 7.1 to 7.5 occur on each fault segment approximately every 2000 to 4000 years (machette and others, 1991). with ten fault segments, we can expect that a large surface-rupturing earthquake will occur every 200-400 years somewhere along the wasatch fault. 8 figure 17. folded mississippian strata of the gardison, deseret, and humbug formations in rock canyon. figure 18. geologic units mapped onto the profile of squaw peak on the north wall of rock canyon. during the sevier orogeny, these rocks were thrust eastward several tens of miles and emplaced in the rock canyon area. the main faults and fractures formed during this mountain building event are shown in this sketch of the north wall of the canyon. the red box shows the area blown up in figure 19. modified from wald (2007). scale miles 0 0.50.40.3 0.20.1 figure 17. folded mississippian strata of the gardison, deseret and humbug formation in rock canyon. 9 20 ft covered slope covered slope tintic quartzite m i n e r a l f o r k t i l l i t e b r e c c i a t e d z o n e ( b e d d i n g n o t p r e s e r v e d ) figure 19. core zone of rock canyon fold. top – photo of the core zone. bottom – interpretive drawing showing the contact between the mineral fork tillite and tintic quartzite with faults (red) and bedding (gray). bedding has been completely destroyed in the central brecciated zone. modified from wald (2007). figure 20. wasatch fault segments as described by machette and others (1991). each segment is capable of producing magnitude 7.0 to 7.5 earthquakes. on average earthquakes occur at intervals of between 500-2500 years on each segment. this figure is modified from hintze (2005). rock canyon is located approximately where the word provo is on the figure. figure 19. core zone of rock canyon fold. top – photo of the core zone. bottom – interpretive drawing showing the contact between the mineral fork tillite and tintic quartzite with faults (red) and bedding (gray). bedding has been completely destroyed in the central brecciated zone. modifi ed from wald (2007). 9 20 ft covered slope covered slope tintic quartzite m i n e r a l f o r k t i l l i t e b r e c c i a t e d z o n e ( b e d d i n g n o t p r e s e r v e d ) figure 19. core zone of rock canyon fold. top – photo of the core zone. bottom – interpretive drawing showing the contact between the mineral fork tillite and tintic quartzite with faults (red) and bedding (gray). bedding has been completely destroyed in the central brecciated zone. modified from wald (2007). figure 20. wasatch fault segments as described by machette and others (1991). each segment is capable of producing magnitude 7.0 to 7.5 earthquakes. on average earthquakes occur at intervals of between 500-2500 years on each segment. this figure is modified from hintze (2005). rock canyon is located approximately where the word provo is on the figure. 8 figure 17. folded mississippian strata of the gardison, deseret, and humbug formations in rock canyon. figure 18. geologic units mapped onto the profile of squaw peak on the north wall of rock canyon. during the sevier orogeny, these rocks were thrust eastward several tens of miles and emplaced in the rock canyon area. the main faults and fractures formed during this mountain building event are shown in this sketch of the north wall of the canyon. the red box shows the area blown up in figure 19. modified from wald (2007). scale miles 0 0.50.40.3 0.20.1 west eastfigure 18. geologic units mapped onto the profi le of squaw peak on the north wall of rock canyon. during the sevier orogeny, these rocks were thrust eastward several tens of miles and emplaced in the rock canyon area. th e main faults and fractures formed during this mountain building event are shown in this sketch of the north wall of the canyon. th e red box shows the area blown up in fi gure 19. modifi ed from wald (2007). 8 figure 17. folded mississippian strata of the gardison, deseret, and humbug formations in rock canyon. figure 18. geologic units mapped onto the profile of squaw peak on the north wall of rock canyon. during the sevier orogeny, these rocks were thrust eastward several tens of miles and emplaced in the rock canyon area. the main faults and fractures formed during this mountain building event are shown in this sketch of the north wall of the canyon. the red box shows the area blown up in figure 19. modified from wald (2007). scale miles 0 0.50.40.3 0.20.1 8 figure 17. folded mississippian strata of the gardison, deseret, and humbug formations in rock canyon. figure 18. geologic units mapped onto the profile of squaw peak on the north wall of rock canyon. during the sevier orogeny, these rocks were thrust eastward several tens of miles and emplaced in the rock canyon area. the main faults and fractures formed during this mountain building event are shown in this sketch of the north wall of the canyon. the red box shows the area blown up in figure 19. modified from wald (2007). scale miles 0 0.50.40.3 0.20.1 12 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 from studies of the trench across the fault at rock canyon (located near the heliport on fi gure 6), the most recent large earthquake appears to have occurred approximately 600 years ago (lund and black, 1998; lund, 2005). th is earthquake created a fault scarp (a surface fault rupture) 6 to 10 feet (2-3 m) high, the remnants of which can still be seen along the road up to rock canyon (fi gure 21). also, if you examine the outcrops of the tintic quartzite on the north side of the canyon (ene of the heliport at the mouth of the canyon on fi gure 6), you will fi nd that the outcrop of quartzite has been broken and fractured into many small pieces by the movement along the wasatch fault. lake bonneville and the pleistocene ice age about 1.8 million years ago the world-wide climate cooled signifi cantly creating large continental glaciers in north america and europe. scientists studying this ice age have identifi ed several periods of ice advances and retreats as the climate alternately cooled and warmed (petersen and others, 1973). in utah during periods of cooler climate, mountain glaciers formed in the wasatch and uinta mountains and large, fresh-water glacial lakes, called pluvial lakes, formed in the basins to the west (atwood, 1909; blackwelder, 1931; madsen and currey, 1979). th e most recent, and one of the largest of these pluvial lakes was lake bonneville (gilbert, 1890; oviatt, 1997). reheis and others (2014) suggested that about 30,000 years ago the basin now occupied by the great salt lake began to fi ll and lake bonneville was born. th e lake grew until it covered almost one quarter of the state of utah (hintze and kowallis, 2009). th e lake reached its peak at about 18,000 years ago forming the highest prominent lake terrace. th is terrace, called the bonneville terrace or shoreline, can be seen from the mouth of rock canyon (fi gure 22). once the lake reached this level, a natural dam near the utah-idaho border was breached allowing water to fl ood and drain out to the north catastrophically onto the snake river plain (reheis and others, 2014; oviatt and jewell, 2016) causing the lake level to drop rapidly to a lower level, called the provo level, where it stabilized for several hundred years. continuing climate change caused regression of the lake to the present level of the great salt lake. th e provo-level terrace is the fl at surface on which most of the brigham young university campus is built (fi gure 6). in addition to these lake terraces near the mouth of rock canyon, visitors can also observe the change in the character of the canyon as they hike to the east. th e lower reaches of the canyon have a typical “v” shape produced by stream erosion, but farther into the canyon, just beyond the picnic and camping area on fi gure 6, it opens up into a wider “u” shape carved by glaciation (fi gure 23). recent alluvial deposits aft er the disappearance of lake bonneville, sediments carried by the stream fl owing out of rock canyon were dumped at the mouth of the canyon forming an alluvial fan (see the fan-shaped deposit on fi gure 6 outward from the mouth of the canyon). today, many homes as well as the provo temple and missionary training cenfigure 20. wasatch fault segments as described by machette and others (1991). each segment is capable of producing magnitude 7.0 to 7.5 earthquakes. on average earthquakes occur at intervals of between 500-2500 years on each segment. th is fi gure is modifi ed from hintze (2005). rock canyon is located approximately where the word provo is on the fi gure. 9 20 ft covered slope covered slope tintic quartzite m i n e r a l f o r k t i l l i t e b r e c c i a t e d z o n e ( b e d d i n g n o t p r e s e r v e d ) figure 19. core zone of rock canyon fold. top – photo of the core zone. bottom – interpretive drawing showing the contact between the mineral fork tillite and tintic quartzite with faults (red) and bedding (gray). bedding has been completely destroyed in the central brecciated zone. modified from wald (2007). figure 20. wasatch fault segments as described by machette and others (1991). each segment is capable of producing magnitude 7.0 to 7.5 earthquakes. on average earthquakes occur at intervals of between 500-2500 years on each segment. this figure is modified from hintze (2005). rock canyon is located approximately where the word provo is on the figure. figure 21. looking north toward mount timpanogos along the wasatch fault from the road atop the alluvial fan at the mouth of rock canyon. notice the abrupt change in the slope of the chain link fence along the road. th is change in the fence line occurs where it crosses the fault scarp created during the most recent large earthquake (which has been dated at about 600 years ago) on the provo segment of the wasatch fault. th ere is no break in the road here because fi ll was brought in to smooth out the slope. because the fi ll is not as fi rmly compacted as the rest of the alluvial fan sediments, the asphalt typically has cracks running across the road at this location. 10 figure 21. looking north toward mount timpanogos along the wasatch fault from the road atop the alluvial fan at the mouth of rock canyon. notice the abrupt change in the slope of the chain link fence along the road. this change in the fence line occurs where it crosses the fault scarp created during the most recent large earthquake (which has been dated at about 600 years ago) on the provo segment of the wasatch fault. there is no break in the road here because fill was brought in to smooth out the slope of the roadway when it was built. because the fill is not as firmly compacted as the rest of the alluvial fan sediments, the asphalt typically has cracks running across the road at this location. figure 22. looking south from the parking lot at the mouth of rock canyon. the power line poles are placed on the bonneville terrace, a flat surface created when lake bonneville was at its highest level. the steep slope in the middle of the photo to the left of the power poles represents the approximate location of the wasatch fault along which the mountains have been uplifted. figure 23. in the high elevations to the east, near the head of rock canyon, the canyon opens up into a wide “u” shape sculpted out by glaciers during the pleistocene ice ages. b.j. kowallis and l.c. wald rock canyon 13 ter of the church of jesus christ of latter-day saints are built on this alluvial fan. as described earlier in this paper, the alluvial fan is a place where flooding out of the canyon can still be a problem. summary rock canyon is a location where many of the events that shaped utah’s geologic history can be studied and observed. it is an excellent place for students, scout groups, families, or anyone interested in learning more about geology to spend a few hours. the hiking trails, picnic areas, rock climbing areas, rest rooms, along with easy access and parking make this one of utah’s best outdoor geologic laboratories. references cited armstrong, r.l., 1968, sevier orogenic belt in nevada and utah: geological society of america bulletin, v. 79, p. 429-458. atwood, w.w., 1909, glaciation of the uinta and wasatch mountains: u.s. geological survey professional paper 61. baker, a.a., 1972, geologic map of the bridal veil falls quadrangle, utah: u.s. geological survey map gq-998 (1:24,000). baker, a.a., 1973, geologic map of the springville quadrangle, utah: u.s. geological survey map gq-1103 (1:24,000). baker, a.a., huddle, j.w., and kinney, d.m., 1949, paleozoic geology of north and west sides of uinta basin, utah: bulletin of the american association of petroleum geologists, v. 33, p. 1161-1197. bissell, h.j., 1974, tectonic control of late paleozoic and early mesozoic sedimentation near the hinge line of the cordilleran miogeosynclinal belt: tectonics and sedimentation, the society of economic paleontologists and mineralogists, v. sp22, p. 83-97. blackwelder, e., 1931, pleistocene glaciation in the sierra nevada and basin ranges: geological society of america bulletin, v. 42, p. 865-922. bush, s., 1995, rock canyon watershed: a history and resource review: honors thesis, brigham young university, 53 p. (as 36.b752 b8812 1995). christie-blick, n., 1983, glacial-marine and subglacial sedimentation upper proterozoic mineral fork formation, utah, in molnia b.f., editor, glacial-marine sedimentation: springer, boston, ma, p. 703-776. christie-blick, n., 1997, neoproterozoic sedimentation and tectonics in west-central utah, in link, p.k, and kowallis, b.j., editors, proterozoic to recent stratigraphy, tectonics, and volcanology, utah, nevada, southern idaho and central mexico: brigham young university geology studies, v. 42, part i, p. 1-30. clark, d.l., derenthal, d., kowallis, b.j., and ritter, s.m., 2014, the major pre-mississippian unconformity in rock canyon, central wasatch range, utah: geology of the intermountain west, v. 1, p. 1-5. crittenden, m.d., jr., christie-blick, n., and link, p.k., 1983, evidence for two pulses of glaciation during the late proterozoic in northern utah and southeastern idaho: geological society of america bulletin, v. 94, p. 437-450. davis, f.d., 1983, geologic map of the southern wasatch front, utah: utah geological and mineral survey map 55-a (1:100,000). decelles, p.g., and coogan, j.c., 2006, regional structure and kinematic history of the sevier fold-and-thrust belt, central utah: geological society of america bulletin, v. 118, p. 841864. ersking, m.c., 1997, the oquirrh basin revisited: american association of petroleum geologists bulletin, v. 81, p. 624-636. gilbert, g.k., 1890, lake bonneville: u.s. geological survey monograph 1, 438 p. figure 22. looking south from the parking lot at the mouth of rock canyon. the power line poles are on the bonneville terrace, a flat surface created when lake bonneville was at its highest level. the steep slope in the middle of the photo to the left of the power poles represents the approximate location of the wasatch fault along which the mountains have been uplifted. 10 figure 21. looking north toward mount timpanogos along the wasatch fault from the road atop the alluvial fan at the mouth of rock canyon. notice the abrupt change in the slope of the chain link fence along the road. this change in the fence line occurs where it crosses the fault scarp created during the most recent large earthquake (which has been dated at about 600 years ago) on the provo segment of the wasatch fault. there is no break in the road here because fill was brought in to smooth out the slope of the roadway when it was built. because the fill is not as firmly compacted as the rest of the alluvial fan sediments, the asphalt typically has cracks running across the road at this location. figure 22. looking south from the parking lot at the mouth of rock canyon. the power line poles are placed on the bonneville terrace, a flat surface created when lake bonneville was at its highest level. the steep slope in the middle of the photo to the left of the power poles represents the approximate location of the wasatch fault along which the mountains have been uplifted. figure 23. in the high elevations to the east, near the head of rock canyon, the canyon opens up into a wide “u” shape sculpted out by glaciers during the pleistocene ice ages. figure 23. in the high elevations to the east, near the head of rock canyon, the canyon opens up into a wide “u” shape sculpted out by glaciers during the pleistocene ice ages. 10 figure 21. looking north toward mount timpanogos along the wasatch fault from the road atop the alluvial fan at the mouth of rock canyon. notice the abrupt change in the slope of the chain link fence along the road. this change in the fence line occurs where it crosses the fault scarp created during the most recent large earthquake (which has been dated at about 600 years ago) on the provo segment of the wasatch fault. there is no break in the road here because fill was brought in to smooth out the slope of the roadway when it was built. because the fill is not as firmly compacted as the rest of the alluvial fan sediments, the asphalt typically has cracks running across the road at this location. figure 22. looking south from the parking lot at the mouth of rock canyon. the power line poles are placed on the bonneville terrace, a flat surface created when lake bonneville was at its highest level. the steep slope in the middle of the photo to the left of the power poles represents the approximate location of the wasatch fault along which the mountains have been uplifted. figure 23. in the high elevations to the east, near the head of rock canyon, the canyon opens up into a wide “u” shape sculpted out by glaciers during the pleistocene ice ages. 14 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 granger, a.e., sharp, b.j., crittenden, m.d., and calkins, f.c., 1952, geology of the wasatch mountains east of salt lake city: utah geological society, geology of the central wasatch mountains, utah, guidebook to the geology of utah, no. 8, p. 1-37. hintze, l.f., 1978, geologic map of the y mountain area, east of provo, utah: brigham young university geology studies, special publication 5. hintze, l.f., and kowallis, b.j., 2009, geologic history of utah: brigham young university geology studies special publication 9, 225 p. jensen, j.m., 1924, early history of provo, utah: provo, utah, self-published, available at brigham young university harold b. lee library (f 834 .p8 j45x 1924), 182 p. jordan, t.e., and douglass, r.c., 1980, paleogeography and structural development of the late pennsylvanian to early permian oquirrh basin, northwestern utah, in fouch, t.d., and magathan, e.r., editors, paleozoic paleogeography of west central united states: society of economic paleontologists and mineralogists, rocky mountain section, west-central united states paleogeography symposium i, p. 217-238. kay, m., 1951, north american geosynclines: geological society of america memoir 48, 143 p. kluth, c.f., 1986, plate tectonics of the ancestral rocky mountains: part iii, middle rocky mountains: paleotectonics and sedimentation in the rocky mountain region, united states: american association of petroleum geologists memoir 41, p. 353-369. kowallis, b.j., and walker, b., 2009, rock canyon field guide: brigham young university department of geological sciences, 1 folded sheet. kowallis, b.j., ferguson, j., and jorgensen, g.j., 1990, uplift along the salt lake segment of the wasatch fault from apatite and zircon fission track dating in the little cottonwood stock: nuclear tracks and radiation measurements, v. 17, p. 325-329. lee, w.t., 1918, early mesozoic physiography of the southern rocky mountains: smithsonian miscellaneous collections, v. 69. link, p.k., miller, j.m.g., and christie-blick, n., 1994, glacial-marine facies in a continental rift environment: neoproterozoic rocks of the western united states cordillera, in deynoux, m., miller, j.m.g., domack, e.w., eyles, n., fairchild, i.j., and young, g.m., editors, earth’s glacial record: cambridge university press, p. 29-46. lochman-balk, c., 1976, the cambrian section of the central wasatch mountains, in hill, j.g., editor, geology of the cordilleran hingeline: rocky mountain association of geologists – 1976 symposium, p. 103-108. lund, w.r., 2005, consensus preferred recurrence-interval and vertical slip-rate estimates: review of utah paleoseismic-trenching data by the utah quaternary fault parameters working group: utah geological survey bulletin 134, 109 p. lund, w.r., and black, b.d., 1998, paleoseismic investigation at rock canyon, provo segment, wasatch fault zone, utah county, utah: utah geological survey special study 93, 21 p. machette, m.n., personius, s.f., nelson, a.r., schwartz, d.p., and lund, w.r., 1991, the wasatch fault zone, utah–segmentation and history of holocene earthquakes: journal of structural geology, v. 13, p. 137-149. madsen, d.b., and currey, d. r., 1979, late quaternary glacial and vegetation changes, little cottonwood canyon area, wasatch mountains, utah: quaternary research, v. 12, p. 254-270. matsch, c.l., and ojakangas, r.w., 1991, comparisons in depositional style of “polar” and “temperate” glacial ice, late paleozoic whiteout conglomerate (west antarctica) and late proterozoic mineral fork formation (utah), in anderson, j.b., and ashley, g.m., editors, glacial marine sedimentation: paleoclimatic significance: geological society of america, special paper 261, p. 191-206. naeser, c.w., bryant, b., crittenden, m.d., jr., and sorensen, m.l., 1983, fission-track ages of apatite in the wasatch mountains, utah: an uplift study: geological society of america memoir 157, p. 29-36. ojakangas, r.w., and matsch, c.l., 1980, upper precambrian (eocambrian) mineral fork tillite of utah: a continental glacial and glaciomarine sequence: geological society of america bulletin, part i, v. 91, p. 495-501. oviatt, c.g., 1997, lake bonneville fluctuations and global climate change: geology, v. 25, p. 155-158. parry, w.t., and bruhn, r.l., 1987, fluid inclusion evidence for minimum 11 km vertical offset on the wasatch fault, utah: geology, v. 15, p. 67-70. petersen, m.s., rigby, j.k., and hintze, l.f., 1973, historical geology of north america: dubuque, iowa, wm. c. brown co., 193 p. peterson, d.o., and clark, d.l., 1974, trace fossils plagiogmus and skolithos in the tintic quartzite (middle cambrian) of utah: journal of paleontology, v. 48, p. 766-768. reheis, m.c., adams, k.d., oviatt, c.g., and bacon, s.n., 2014, pluvial lakes in the great basin of the western united states–a view from the outcrop: quaternary science reviews, v. 97, p. 33-57. stokes, w.l., 1976, what is the wasatch line?, in hill, j.g., editor, geology of the cordilleran hingeline: rocky mountain association of geologists–1976 symposium, p. 11-25. 15 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 varney, p.j., 1976, depositional environment of the mineral fork formation (precambrian), wasatch mountains, utah, in hill, j.g., editor, geology of the cordilleran hingeline, rocky mountain association of geologists – 1976 symposium, p. 91-102. ver wiebe, w.a., 1930, ancestral rocky mountains: american association of petroleum geologists bulletin, v. 14, p. 765788. wald, l.c., 2007, structural analysis of rock canyon near provo, utah: m.s. thesis, brigham young university, 63 p. wilson, m.l., 1988, analysis of a proposed economic development: the case of heritage mountain: m.s. thesis, brigham young university, provo, utah, 123 p. (g1.02 .w55 1988). ye, h., royden, l., burchfiel, c., and schuepbach, m., 1996, late paleozoic deformation of interior north america: the greater ancestral rocky mountains: american association of petroleum geologists bulletin, v. 80, p. 1397-1432. young, r.g., 1966, stratigraphy of coal-bearing rocks of book cliffs, utah-colorado: utah geological and mineral survey bulletin 80, p. 7-22. uga-geosite-shapiro-antelope-island.indd 7 figure 7 figure 7. view of the outcrops of the lower dolostone of the kelley canyon formation at the geosite. the great salt lake is in the background, toward the west. utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors russell shapiro1 and carol dehler2 1 department of geological and environmental sciences, california state university-chico, ca 95929, rsshapiro@csuchico.edu 2 department of geology, utah state university, logan, ut 84322 geobiology of “snowball earth” deposits of antelope island cover image: view of the outcrops of the lower dolostone of the kelley canyon formation at the geosite. th e great salt lake is in the background, toward the west. 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e g eosites refl e ct the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e li st is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e u tah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: shapiro , r., and dehler , c., 2022, geobiology of "snowball earth"—deposits of antelope island: utah geological association publication, v. 1, no. 1, p. 1-8., doi: 10.31711/ ugap.v1i1.100. r. shapiro and c. dehler geobiology of “snowball earth” deposits of antelope island 3 introduction antelope island on great salt lake provides an excellent opportunity to look at one of the world’s great geobiological records—the “snowball earth.” snowball earth refers to a unique time in earth history before the dawn of skeletonized animals where there is substantial evidence to support glaciers at sea level in the equatorial regions. many researchers have proposed that the only way to achieve this unique condition is to freeze the entire planet, hence the “snowball earth” (refs). we use quotation marks around the name of this global phenomenon because the scope and details of this major climatic phenomenon are still debated. after 30 years of rigorous testing since the idea was proposed (kirschivink, 1992), this hypothesis is still holding up (hoffman and others, 2017). besides being a record of two global glaciations lasting tens of millions of years between 717 and 635 million years ago, there may be a connection between these mega-scale climate changes and the evolution of animal life. geobiology is a continually evolving field that looks at the intersection of geological context with biological and ecological constraints, using models and assessing validity through a variety of biogeochemical proxies. put another way, geobiologists recognize that classic studies of biological communities, evolution, ecology, and extinction need to be filtered through geology at scales from plate tectonics to bedrock and substrate. also, geological processes within a few miles of the surface where sediments convert to rock likely involve microbial interactions. so why should geologists care? microbe-mineral interactions are now known to be integral to the study of environmental remediation, reservoir characteristics, some economic deposits, as well as the lofty academic pursuits of early earth evolution and astrobiology. it is with this view that the following discussion of the snowball earth deposits on antelope island is framed, albeit from a largely geological focus. location and access the snowball earth geosite is located on antelope island state park (figure 1). it is legal to visit the locations without a permit, although no collecting is allowed. please take only photos! the following directions start at the visitor’s center on antelope island. from the visitor’s center, head west then south along bridger bay. turn left (east) at the end of the road, away from bridger bay campground. follow the signs to the white rock group campground and park at the far, southeastern corner. from here you will walk along the bone road trail, essentially south. after a mile (1.6 km) or so, the trail will slowly climb, eventually connecting with the junction trail off the white rocks loop trail. the last section of the trail has a steeper incline (~600 foot [180 m] gain) to the crest. the total trail is 3 miles (5 km) each way from the parking area. wearing sturdy shoes and insect repellant are recommended for the walk and there is the potential to take a dip in the lake, so wearing a bathing suit is also recommended. bathrooms, showers and a snack bar are available at the beach at bridger bay. hats, sunscreen, lip balm, gnat repellent, and several liters of water are highly recommended, especially between may and september. gps locations: 41° 1'29.75"n; 112°14'22.38"w parking area 40°59'13.37"n; 112°13'19.81"w geosite snowball earth the recognition of glacial deposits interbedded with marine sediments at low latitudes during several times in earth’s history led to the hypothesis of a “snowball earth.” originally coined by joe kirschvink (1992) for deposits of paleoproterozoic age (~2.3 billion years), the term is more commonly associated with the neoproterozoic record (~1000-540 ma; hoffman and others, 1998). recently, a new geological period, the cryogenian (717-635 ma), was established to encompass the duration of the two global glaciations of the neoproterozoic era, although the global stratotype, section, and point (gssp) for the base of the cryogenian is still being determined (halverson and others, 2018). in addition to glacial deposits, primarily diamictite (or mud-supported conglomerate), these glacial intervals also include dramatic sea-level changes and enigmatic, massive to banded carbonates with “tubestone” (see a more detailed description below) (figure 2). these carbonate units may also include large crystal fans, that were originally aragonite, which indicate rapid deposition of carbonate and high alkalinity during the deglaciations (hoffman and others, 2018). collectively, these carbonate units are called “cap carbonates” because they cap the glacial deposits of each of the two glaciations. figure 1. a) location of geosite on antelope island. upper image shows the major roads (solid) and walking trails (dashed). black lines illustrate the route in the text; additional routes are in grey. upper road is the causeway. b) a simplified geological map of northern antelope island. red dot is the geosite. xwf=farmington canyon complex; zmk=mineral fork and kelly canyon fms.; ct=tintic fm.; t=tertiary conglomerate. bp=buffalo point; eh=elephant head. based on yonkee and others (2000). 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 finally, in the neoproterozoic, there is a small resurgence of iron formation, suggesting low oxygen or high iron input conditions that is also likely related to the decoupling of the ocean from the atmosphere by ice sheets and the continued input of reduced forms of elements such as iron from sea floor spreading, subaqueous mantle input, and sedimentation (cox and others, 2013). while each piece of evidence is suggestive, it is the sum total that points to the occurrence of extreme glaciations. what makes this particular interval a global “snowball earth” is the presence of glacial deposits and related cap carbonates in paleoequatorial regions as well as higher latitudes (figure 3). furthermore, climate models show that once polar ice sheets grow to about 30 degrees north and south latitudes, there will be a runaway ice-albedo effect (ice reflecting the sun’s heat) in which ice covers the planet by growing into equatorial regions (budyko, 1969). there is also mounting evidence that conditions led to sluggish ocean currents that would have significantly impacted global nutrient fluxes. altogether, the evidence points to a dramatic time of globally extensive glaciations, hence the “snowball earth” (e.g., hoffman and others, 2017). so how does this work? the geological and climatic forcing mechanisms that led into a global glaciation and also terminated these glaciations are still debated. in part, this is due to the paucity of neoproterozoic geologic deposits and challenges in finding dateable material (such as volcanic ash), although many new age models for many cryogenian strata are being generated; all point to two global glaciations, the sturtian glaciation (717-635 ma) and the marinoan glaciation (>640-635 ma) (e.g., calver and others, 2013; macdonald and others, 2010; prave 2016). also, geologic data influence climate models and there are multiple competing models that are equally valid (e.g., bechstaedt and others, 2018). it is generally accepted that the conditions were unique, both with regards to continent location and ocean currents. also, there were a series of positive feedback loops that, once initiated, led to rapid cooling of the continents. the major feedback loop had to do with earth’s albedo. the ‘albedo’ is a measure of the reflectivity of solar radiation. white ice is reflective, so as permanent ice builds on the continents and oceans, more solar radiation is reflected, leading to cooling and the growth of more ice. additionally, by gathering the continents in the lower latitudes (as evidenced by the paleomagnetic record (e.g., sohl and others, 1999), there is also an increase in overall albedo because continental rocks reflect more energy than ocean waters and solar radiation is focused on the tropics (hoffman and schrag, 2002. the cryogenian was the time when rodinia, a major supercontinent, was in the process of breaking apart. although the orientation is still being debated, the location and orientation of rodinia continental pieces affected ocean circulation and nearly all models show the individual continental plates amassing in middle to low latitudes (li and others, 2013). once ice develops, sea level is lowered, exposing more continental shelves, thus further increasing albedo (kirshvink, 1992). one of the big challenges is covering the oceans 2 figure 2 figure 2. highly idealized composite stratigraphic column through a snowball earth sequence. in reality, no one location shows all the features perfectly. on antelope island, one sees the diamictite of the mineral fork formation and the thinly laminated and stromatolitic tubestone facies of the kelley canyon formation. figure 2. highly idealized composite stratigraphic column through a snowball earth sequence. in reality, no one location shows all the features perfectly. on antelope island, one sees the diamictite of the mineral fork formation and the thinly laminated and stromatolitic tubestone and large wave ripple facies of the kelley canyon formation. figure 3. global distribution of the three main pulses of snowball earth glaciation, plotted on modern distribution of the continents. note that in the neoproterozoic, all the landmasses were more or less joined in a supercontinent, rodinia. upper (a) is the earlier marinoan (645-635 ma) and lower (b) is the later sturtian (717-659 ma). yellow dots indicated glacial or periglacial deposits and red dots show associated iron oxide formation. from hoffman and others (2017, figure 4). r. shapiro and c. dehler geobiology of “snowball earth” deposits of antelope island 5 with ice. ocean water has a high heat capacity, making it possible to insulate against ice growth, no matter how strong the overall global albedo. in many models, regions of the tropical ocean remain ice free (e.g., baum and crowley, 2001; abbott and others, 2011). either way, the evidence thus far is overwhelming that vast regions of the planet had permanent ice cover, all the way down to the tropics. finally, all this ice would lead to an overall drying of the atmosphere, eventually leading to the shutdown of the hydrological cycle during the cyrogenian (hoffman and others, 1998). of course, the big problem is, how does the earth recover from a snowball earth? the positive feedback loops to transition the planet into a “snowball earth” state means there is a tipping point at which the earth could remain permanently frozen. the most accepted models for getting out of a frozen world rely on the buildup of volcanic co2 in the atmosphere triggering a rapid meltdown through greenhouse heating (kirschvink, 1992; caldeira and kasting, 1992; hoffman and schrag, 2002). normally, the co2 gas would be consumed by continental weathering and entrenchment in the oceans, but in a frozen world, these sinks would be blocked. but, as the co2—as well as methane and other greenhouse gases—built up in the atmosphere, they would block the reflected solar radiation and earth’s own heat emissions from escaping to space, thus leading to global warming. the next phase varies between models, but it is clear that warming would ensue. as the ice eventually thins and melts off the continents, there would be intensive continental weathering by carbonic acid, leading to a rapid buildup of carbonates in shallow water (hence ‘cap carbonates’). the δ13c values (ratio of heavy carbon to the biologically favored light carbon) through this interval is well below typical ocean values of the past 550 million years. the light values are interpreted as a result of influx of mantle co2, hence volcanic input (veizer and others, 1999), and a great reduction in biological fractionation of carbon isotopes (hoffman and others, 1998). an alternate model, called the “slushball hypothesis” allows for a dominantly ice-covered earth but with enough open water, particularly in the low latitudes, for a full hydrologic cycle (e.g., pollard and kasting, 2005) and a refugia for biota (kirschivink, 1992). this could explain some of the sedimentary features that are difficult to reconcile with a fully frozen earth and also takes other climate models into account which do not fully freeze-over the oceans. a new model relies more heavily on the break-up (rifting) of rodinia into its lithospheric progeny (li and others, 2013) and the impact of rifted margins on glacial stratigraphy, cap carbonate facies distribution and the enigmatic iron formations. specifically, rifting of rodinia would lead to rift margins that could trap oxygen-poor water, leading to basin-specific iron-formations (baldwin and others, 2012). the tectonic reorganization could also shift ocean currents, thus driving more localized ice conditions. geological setting antelope island is a classic basin-and-range uplift and, though it is small relative to the other ranges, it is the largest island (42 square miles [110 km2]) in the great salt lake and exposes a wealth of geologic information (king and willis, 2000) (figure 1b). the oldest outcropping units are early paleoproterozoic metamorphic units of the farmington canyon complex. sharply overlying the metamorphic basement are the conglomerate and finer-grained units of the mineral fork formation (glacial deposits). these strata are overlain by dolostone of the basal kelley canyon formation (cap carbonate). the remainder of the kelley canyon is composed of shale and siltstone and minor carbonate. the cambrian tintic quartzite unconformably overlies the kelley canyon and dominates the northern part of the island, where one enters via the causeway (antelope island road) (figure 1b). like most of the eastern basin-and-range, cretaceous-age sevier thrusting dictates the structural geology, resulting in a series of thrust slices (yonkee and weil, 2015). the sevier orogeny caused brittle-to-ductile deformation with outstanding exposures of shearing on antelope island. in fact, it is important to note that the neoproterozoic succession on antelope island is significantly thinner than nearby age-equivalent units to the west and north because it is on a different thrust sheet (levy and christie-blick, 1989; yonkee and others, 2000, 2014). this geosite focuses on an excellent outcrop of the upper stratigraphy of one of these thrusts that brings the neoproterozoic mineral fork and kelly canyon formations over the cambrian tintic quartzite (figure 4). neoproterozoic succession on antelope island the most accessible exposures of the neoproterozoic strata lie at the top of the ridge running east from elephant head. at this location, the mineral fork formation comprises mostly diamictite (conglomerate with a fine-grained matrix) up to 200 feet (60 m) thick though structural thinning and thickening throughout limits the integrity of this value (yonkee and others, 2000) (figure 4). some of the thickness variation is due to irregularities on the nonconformity surface atop the basement. clast sizes within the diamictite vary from granules to boulders, and clast lithology is limited to metamorphic material derived from the underlying 4 figure 4 figure 4. simplified bedrock stratigraphy of geosite on antelope island. no vertical scale is implied. units of focus are the mineral fork formation and lower dolostone member of the kelley canyon formation. based on yonkee and others (2000; 2014). figure 4. simplified bedrock stratigraphy of geosite on antelope island. no vertical scale is implied. units of focus are the mineral fork formation and lower dolostone member of the kelley canyon formation. based on yonkee and others (2000; 2014). 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 basement (figure 5). the lack of sorting, great variety of clast sizes, and fine-grained matrix suggest a glacial or mass-flow origin. the occurrence of dropstones east of elephant head indicates a definitive glacial origin (figure 6). dropstones are outsized clasts that drop from melting icebergs or sea ice and disrupt the underlying laminations. at a minimum, one can see at this outcrop the great diversity in size and composition of the clasts. this certainly points to a dynamic process! overlying the mineral fork formation is the basal dolostone of the kelley canyon formation. to find the best outcrops, walk west from the diamictite wall to the low outcrops in the saddle (figure 7). it is worth your time to poke around as there are a variety of fabrics preserved in the dolostone. while most of the pink to white dolostone is massive to thinly laminated (figure 8), some blocks contain beds of the exotic “tubestone” (cloud and others, 1974) (figure 9). both fabrics are indicative of “cap carbonates” associated with neoproterozoic glaciations globally. the current model is that once co2 reached a tipping point in the atmosphere, there was rapid global warming which then led to melting of the ice sheets. the co2, in the form of carbonic acid, was free to chemically weather the continents, leading to a massive flux of dissolved material into the oceans. furthermore, the co2 also built up in the oceans itself, leading to a rapid deposition of calcium carbonate minerals, hence the thick beds of in situ tubestone, and at other localities, in situ aragonite fans (hoffman and schrag, 2002). (the dolomite mineral was probably not original but replaced the original aragonite after burial.) the tubestone is more enigmatic, though also found globally in cap carbonate sequences. here on antelope island, the facies consists of parallel tubes, a few centimeters wide, separated by regions that are finely laminated (hayes and dehler, 2011) (figure 9). both fabrics are now dolomite. for many years, it was believed the tubes represented gas escape structures, where ascending gas bubbles disrupted the laminated sediments (cloud and others, 1974). more recently, kennedy and others (2001) specified that the gas was methane and the release of such a volume of greenhouse gas would have further intensified the exit from a snowball earth. other ideas have also been suggested, such as early vertical animal burrows, but the evidence does not support these ideas. in a critical paper on the subject, corsetti and grotzinger (2005) noted some key features overlooked by previous researchers. first, the tubes were not necessarily round in cross-section. in fact, most were of complex, multi-lobed forms. also, while they looked to be filled with solid mud (micrite) compared to the laminated facies, in truth it is quite common to find laminae extending across the tubes. if you look around carefully at the outcrop on antelope island, you will find examples of these tube-crossing laminae. 5 figure 5 figure 5. outcrop of strained diamictite from antelope island. note the variety of clasts. 2014 wasatch-uinta field camp students for scale. figure 5. outcrop of strained diamictite from antelope island. note the variety of clasts. 2014 wasatch-uinta field camp students for scale. 6 figure 6 figure 6. dropstones in the mineral fork formation on antelope island. the criteria for recognizing dropstones is that the laminations under the clasts are ‘pierced’, indicating a definitive glacial origin for the deposit (see red arrows). the clasts drop from melting icebergs and fall onto the seafloor, piercing and deforming the mudstone layers. figure 6. dropstones in the mineral fork formation on antelope island. the criteria for recognizing dropstones is that the laminations under the clasts are ‘pierced’, indicating a definitive glacial origin for the deposit (see red arrows). the clasts drop from melting icebergs and fall onto the seafloor, piercing and deforming the mudstone layers. 7 figure 7 figure 7. view of the outcrops of the lower dolostone of the kelley canyon formation at the geosite. the great salt lake is in the background, toward the west. figure 7. view of the outcrops of the lower dolostone of the kelley canyon formation at the geosite. the great salt lake is in the background, toward the west. r. shapiro and c. dehler geobiology of “snowball earth” deposits of antelope island 7 this new evidence pointed to the ‘tubes’ not being erosional but rather constructional features. the most likely hypothesis is that the laminated facies represents stromatolites and the ‘tubes’ are the spaces between the stromatolites, often filled with dolomicrite, but also sparry dolomite and iron oxides. stromatolites are a type of fossil that records the growth of microbial mats. as the mats grow and decay and repeat, they may deposit layers of cement or grains, and are thus preserved in the rock record as a sequence of thin, millimeter-scale layers. one layer grows on another, ultimately building up a vertical or domal structure. in the antelope island example, the stromatolites are mostly vertical, resembling folded curtains. occasionally, the microbial mats expanded across the mud flat to the next stromatolite; this is the preservation of laminae across the tubes. summary outcrops of the neoproterozoic mineral fork formation and basal kelley canyon formation exposed on the ridge above elephant head on antelope island state park provide an accessible and excellent window into the ‘snowball earth’ episode. that unique time of global glaciation led to glaciers at sea level near the equator while the end-glaciation is recorded by thick dolostones preserving bizarre ‘tubestone’ stromatolites. while there are many other features diagnostic of ‘snowball earth’ such as iron formation and other dolostone fabrics, antelope island preserves the two main facies: diamictite as evidence of glacial deposition and ‘tubestone’ cap carbonate. furthermore, the views of sevier-aged deformation and expansive views of the island and lake make this an excellent geosite for a short day trip from salt lake city. acknowledgements this project was an outgrowth of several field trips to antelope island by the wasatch-uinta field camp and the authors wish to thank the students and other faculty for the discussions. kathleen nicoll and kurtis burmeister provided comments on an earlier draft and robert biek and mark milligan reviewed the final manuscript. references abbot, d.s., voigt, a., and koll, d., 2011, the jörmungand global climate state and implications for neoproterozoic glaciations: journal of geophysical research: atmospheres, v. 116 (d18). baldwin, g.j., turner, e.c. and kamber, b.s., 2012, a new depositional model for glaciogenic neoproterozoic iron formation: insights from the chemostratigraphy and basin configuration of the rapitan iron formation: canadian journal of earth sciences, v. 49, no. 2, p. 455-476. baum, s.k., and crowley, t.j., 2001, gcm response to late precambrian (~590 ma) ice-covered continents: geophysical research letters, v. 28, no. 4, p. 583-586. 8 figure 8 figure 8. finely laminated lithofacies of the basal kelley canyon formation, antelope island. this is the dominant fabric in most basal cap carbonates worldwide, though it is not by itself diagnostic of the “snowball earth.” magnet is 5 mm across. figure 8. finely laminated lithofacies of the basal kelley canyon formation, antelope island. this is the dominant fabric in most basal cap carbonates worldwide, though it is not by itself diagnostic of the “snowball earth.” magnet is 5 mm across. 9 figure 9 figure 9. “tubestone” of the kelley canyon formation from antelope island. upper photo is a field view of upper surface showing ‘tube’ tops, which are sediment infilling between columnar stromatolites. lower photo is a polished vertical cut. note the stromatolitic laminations of the columns, separated by sediment and cement-filled ‘tubes.’ figure 9. “tubestone” of the kelley canyon formation from antelope island. upper photo is a field view of upper surface showing ‘tube’ tops, which are sediment infilling between columnar stromatolites. lower photo is a polished vertical cut. note the stromatolitic laminations of the columns, separated by sediment and cement-filled ‘tubes.’ lower image from hayes (2013). 8 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 bechstaedt, t., jaeger, h., rittersbacher, a., schweisfurth, b., spence, g., werner, g., and boni m., 2018, the 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origin and significance of tube structures in neoproterozoic post-glacial cap carbonates: example from noonday dolomite, death valley, united states: palaios, v. 20, p. 348-363. cox, g.m., halverson, g.p., minarik, w.g., le heron, d.p., macdonald, f.a., bellefroid, e.j., and strauss, j.v., 2013, neoproterozoic iron formation: an evaluation of its temporal, environmental and tectonic significance: chemical geology, v, 362, p. 232-249. halverson, g.p., porter, s.m., and gibson, t.m., 2018, dating the late proterozoic stratigraphic record: emerging topics in life sciences, v. 2, no. 2, p. 137-147. hayes, d. s., and dehler, c. m., 2011, a new locality of neoproterozoic tube structures in northern utah; insight into genesis and age of a cap carbonate [abs.]: geological society of america abstracts with programs, v. 43, p. 56. hoffman, p.f., kaufman, a.j., halverson, g.p., and schrag, d.p., 1998, a neoproterozoic snowball earth: science, v. 281, no. 5381, p. 1342-1346. hoffman, p.f., and schrag, d.p., 2002, the snowball earth hypothesis: testing the limits of global change: terra nova, v. 14, no. 3, p. 129-155. hoffman, p.f., and 19 others, 2017, snowball earth climate dynamics and cryogenian geology–geobiology: science advances, v. 3, no. 11. (e1600983. issn 2375-2548). kennedy, m.j., christie-blick, n., and sohl, l.e., 2001, are proterozoic cap carbonates and isotopic excursions a record of gas hydrate destabilization following earth’s coldest intervals?: geology, v. 29, no. 5, p. 443-446. king, j.k., and willis, g.c., 2000, the geology of antelope island, davis county, utah: utah geological survey, 163 p. kirschvink, j.l., 1992, late proterozoic low-latitude glaciation: the snowball earth, in schopf, j.w., and klein, c., editors, the proterozoic biosphere: cambridge, cambridge university press, p. 51-52. levy, m., and christie-blick, n., 1989, pre-mesozoic palinspastic reconstruction of the eastern great basin (western united states): science, v. 245, no. 4925, p. 1454-1462. li, z.x., evans, d.a., and halverson, g.p., 2013, neoproterozoic glaciations in a revised global palaeogeography from the breakup of rodinia to the assembly of gondwanaland: sedimentary geology, v. 294, p. 219-232. macdonald, f., schmitz, m., crowley, j., roots, c., jones, d., maloof, a., strauss, j., cohen, p., johnston, d., and schrag, d., 2010, calibrating the cryogenian: science, v. 327, no. 5970, p. 1241-1243 (doi:10.1126/science.1183325). pollard, d., and kasting, j.f., 2005, snowball earth: a thin-ice solution with flowing sea glaciers: journal of geophysical research: oceans, v. 110 (c7). prave, a.r., condon, d.j., hoffmann, k.h., tapster, s., and fallick, a.e., 2016, duration and nature of the end-cryogenian (marinoan) glaciation: geology, v. 44, no. 8, p. 631-634. sohl, l.e., christie-blick, n., and kent, d.v., 1999, paleomagnetic polarity reversals in marinoan (ca. 600 ma) glacial deposits of australia: implications for the duration of low-latitude glaciation in neoproterozoic time: geological society of america bulletin, v. 111, no. 8, p. 1120-1139. veizer, j., ala, d., azmy, k., bruckschen, p., buhl, d., bruhn, f., carden, g.a., diener, a., ebneth, s., godderis, y., and jasper, t., 1999, 87sr/86sr, δ13c and δ18o evolution of phanerozoic seawater: chemical geology, v. 161, nos. 1-3, p. 59-88. yonkee, w.a., and weil, a.b., 2015, tectonic evolution of the sevier and laramide belts within the north american cordillera orogenic system: earth-science reviews, v. 150, p. 531-593. yonkee, w. a., willis, g. c., and doelling, h. h., 2000, proterozoic and cambrian sedimentary and low-grade metasedimentary rocks on antelope island, in king, j. k., and willis, g. c., editors, the geology of antelope island, davis county, utah: utah geological survey, p. 37-49. yonkee, w.a., dehler, c.d., link, p.k., balgord, e.a., keeley, j.a., hayes, d.s., wells, m.l., fanning, c.m., and johnston, s.m., 2014, tectono-stratigraphic framework of neoproterozoic to cambrian strata, west-central us: protracted rifting, glaciation, and evolution of the north american cordilleran margin: earth-science reviews, v. 136, p. 59-95. uga-geosite-chidsey-green-river.indd utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors thomas c. chidsey, jr.1, and hellmut h. doelling2 (retired) 1utah geological survey, po box 146100, salt lake city, utah 84114-6100, tomchidsey@utah.gov 2utah geological survey, 483 e. 200 s., manti, utah 84642 cover image: view to the southwest from the green river overlook. green river overlook, island in the sky district, canyonlands national park, southeastern utah 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 t.c. chidsey, jr., and h.h. doelling green river overlook, canyonlands national park 3 introduction the stream is still quiet, and we guide along through a strange, weird, grand region. the landscape everywhere, away from the river, is of rock—cliffs of rock, tables of rock, plateaus of rock, terraces of rock, crags of rock—ten thousand strangely carved forms; rocks everywhere, and no vegetation, no soil, no sand. in long, gentle curves the river winds about these rocks. these are the words major john wesley powell (figure 1) used to describe what is called stillwater canyon along the green river on july 17, 1869, on his way to the grand canyon during his famous journey exploring the canyons of the colorado river and its tributaries (powell, 1895). this same region is spectacularly displayed from the green river overlook in the island in the sky district of canyonlands national park (figure 2). the exposed rocks consist of early permian(299 million years ago [ma]) through early jurassic-age (176 ma) rock layers that were uplifted and subjected to massive erosion. changes in the color, thickness, and composition of the rock layers and erosive work of running water and gravity (i.e., mass wasting) created the magnificent landscape seen at the overlook today. the primitive and remote maze district of canyonlands national park is the area west of the green river. prominent features viewed in the maze district from the green river overlook include ekker butte, turks head, elaterite butte, and the orange cliffs (figure 2). on a clear day, the 11,000-foot (3350 m), henry mountains, which are cored by volcanic remnants (laccoliths) can be seen to the southwest. the green river overlook is one of utah’s most iconic views, often appearing on book covers, scenic calendars, and postcards, and therefore it is fitting as a geosite for its beauty and the exposures of rocks that so much impressed john wesley powell over 150 years ago. how to get there the island in the sky district of canyonlands national park is about 250 miles (400 km) or a little less a than 4-hour drive from salt lake city, utah, via interstate 15 (i-15), u.s. highway 6, i-70, and u.s. highway 191; 30 miles (48 km) or about 45 minutes if coming from the town of moab, utah. turn west onto state highway 313 and continue to canyonlands national park (figure 3), at which point the road becomes grand view point road, proceeding through the entrance station (fee required) to the junction with upheaval dome road, a total of 28 miles (45 km). turn right onto upheaval dome road and then left after 0.3 miles (0.5 km) onto green river overlook road proceeding 1.4 miles (2.3 km) to the parking lot for the overlook (restroom facilities, camping, and picnic grounds are available). a short walk along a flat, paved sidewalk leads to the overlook, 38°22'38" n., 109°53'19" w., elevation 6000 feet (1829 m). a wooden guard fence separates the view area from dangerous sheer cliffs. stratigraphy and depositional history consolidated sedimentary rocks seen in the vistas from the green river overlook include strata of early permian to early jurassic age (299–176 ma) (figures 2, 4, 5, and 6). these strata have a cumulative thickness of 3000 feet (900 m) (hintze and kowallis, 2009). permian rocks the permian section viewed from the green river overlook is part of the cutler group (figure 5). northeast of the island in the sky district of canyonlands national park, equivalent rocks represent a thick sequence of conglomeratic, arkosic alluvial-fan sediments deposited in front of the pennsylvanian-permian-age uncompahgre highland to the northeast. these strata grade southwest into the fluvial, coastal dune, and tidal flat deposits exposed in canyonlands. two units are viewed from the green river overlook—the lower permian (299–271 ma) organ rock formation and white rim sandstone. organ rock formation the organ rock formation consists of reddish-brown finegrained to silty sandstone, sandy shale, and minor siltstone (figure 2). it is medium to thick bedded and forms steep slopes to ledgy cliffs. the organ rock represents a floodplain depositional environment that was near the edge of a large dune field (figure 7a). figure 1. major john wesley powell with tau-gu, a southern paiute guide; circa 1873. photo courtesy of the national park service. 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 the upper contact of the organ rock strata with the overlying white rim sandstone is sharp and conformable. the organ rock ranges in thickness from 200 to 400 feet (60–120 m) in the island in the sky district (hintze and kowallis, 2009). from the green river overlook, the green river cuts through the organ rock formation in stillwater canyon (figures 2 and 3). the softer units within the organ rock cause the canyon to widen as the green river meanders through it. white rim sandstone the white rim sandstone forms the prominent white rim cliff around the island in the sky district and is the surface of the popular biking trail by the same name (figure 2). the white rim is a fineto medium-grained quartzose sandstone exhibiting both planar and cross-stratified beds. the upper section is a reworked marine unit whereas the lower unit is eolian dominated largescale cross-stratification. the white rim formed as a near-shore, beach, back-beach deposit, and as coastal dunes (figure 7b). the white rim sandstone is up to 250 feet (76 m) thick, thinning to the northeast and pinching out below dead horse point (see figure 4 in the dead horse point geosite, this volume). the upper contact of the white rim sandstone strata with the overlying triassic beds is sharp, marked by local scouring and channeling and is unconformable. in most cases the plane of unconformity appears flat. the unconformity is regional and indicates a period of non-deposition lasting at least 30 to 35 million years (doelling and others, 2010). figure 2. view to the southwest from the green river overlook; unannotated (a) and annotated (b). the green river can be seen in the center distance flowing through stillwater canyon with the permian organ rock formation at river level. the deeper canyon is rimmed by the prominent permian white rim sandstone, whose name best describes the formation in outcrop. the highly eroded red-brown triassic moenkopi formation is extensively exposed in the foreground whereas the far distant cliffs display outcrops of the triassic-jurassic wingate sandstone. the faint outlines of the laccolithic henry mountains can be seen in the very far distance on clear days. photo courtesy of richard emerson. moenkopi fm. t.c. chidsey, jr., and h.h. doelling green river overlook, canyonlands national park 5 figure 3. canyonlands national park and vicinity, southeastern utah, showing the location of the green river overlook geosite and stillwater canyon as well as surrounding parks, towns, and highways. 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 figure 4. geologic map of canyonlands national park and surrounding areas. modified from hintze (1980), hintze and others (2000), and baars (2010). t.c. chidsey, jr., and h.h. doelling green river overlook, canyonlands national park 7 figure 5. stratigraphic column of exposed rocks in canyonlands national park and surrounding areas. note the section viewed from the green river overlook. modified from hintze (2005). figure 6. view to the southeast from the green river overlook; unannotated (a) and annotated (b). the fluvial jurassic kayenta formation forms the ledgy rim of the overlook followed down section by the massive, vertical cliffs of the eolian triassic-jurassic wingate sandstone overlying the slope-forming triassic chinle and moenkopi formations representing fluvial/floodplain and tidal-flat deposition, respectively. a small section of the eolian lower jurassic navajo sandstone can be seen above the kayenta in the far upper left of the photo. a b 8 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 triassic rocks the triassic period in the island in the sky district is represented by the moenkopi and chinle formations and the lower part of the wingate sandstone (figures 2, 4, 5, and 6). during moenkopi time, the open sea or ocean retreated farther to the southwest and deposition was from sluggish streams migrating over broad tidal flats that were at or near sea level (figure 7c). during chinle time, deposition was dominated by fluvial environments and the influence of the sea was gone (figure 7d). moenkopi deposits were laid down approximately 245 ma and again between 242 and 238 ma and chinle deposits were laid down from 220 to 215 ma. events occurring between moenkopi and chinle time are not recorded by rocks in this area. sediments may have been deposited but later eroded during those 18 million years (doelling and others, 2010). the top of the triassic is within the wingate sandstone of the glen canyon group and not at the base of the formation as once thought; the triassic-jurassic boundary is somewhere in the formation (see the triassic-jurassic section below). moenkopi formation the moenkopi formation represents a major change in environmental conditions from continental to marine and tidal flat. the basal section of the moenkopi consists mostly of a chocolate-brown, fine-grained, poorly sorted, micaceous to sub-arkosic sandstone. it is poorly bedded and displays irregular to wavy laminae. the formation creates a steep slope in its lower part and a cliff in its upper part (figure 6). the middle section of the moenkopi consists of chocolate-brown, thinto medium-bedded siltstone slopes with numerous widely spaced, 1to 10-foot-thick (0.3–3 m) sandstone ledges that display low-angle cross-stratification. siltstone and shale beds display a plethora of ripple mark types. ripple cross-stratification is particularly widespread. the uppermost section of the moenkopi is a slope-forming unit consisting of homogeneous, gray-red to pale-red-brown siltstone, with widely spaced, thin (less than 3 feet [1 m] thick) sandstone ledges. the siltstone displays horizontal and ripple cross-lamination. a b c d e f figure 7. paleogeographic maps of utah during the: (a) early permian, organ rock formation – 285 ma; (b) late early permian, white rim sandstone – 272 ma; (c) early triassic, moenkopi formation – 245 ma; (d) late triassic, chinle formation – 215 ma; (e) late triassic-early jurassic, wingate sandstone – 205 ma; and (f) early jurassic, kayenta formation – 200 ma. modified from blakey and ranney (2008). t.c. chidsey, jr., and h.h. doelling green river overlook, canyonlands national park 9 the contact with the overlying chinle formation is sharp and unconformable, but commonly poorly exposed. it is placed at the base of a distinctive white to mottled gritstone (figure 6), or between the grayred or gray-green mudstone and siltstone of the lower part of the chinle and the orange-red siltstones of the upper part of the moenkopi. chinle formation the chinle formation is famous for its petrified wood, uranium in red-brown lenticular channel sandstone beds, and beautiful multicolored mudstone and shale derived from altered volcanic ash. the chinle consists of complex interbedding and lensing arrangements of sandstone, pebble conglomerate, siltstone, mudstone, and rare limestone. the red-brown, tan, and gray-red sandstones are very fine to coarse grained, moderately to well sorted, quartzose, and slightly micaceous. sedimentary structures include low-angle cross-stratification, horizontal stratification, asymmetric ripples, and channeling, all of which point to deposition in a floodplain having northwest-flowing river channels, adjacent oxbow lakes, ponds, and swamps (figure 7d). pebble and intraformational conglomerates occur as lenses and in scour channels concentrated in the lower parts of sandstone beds. siltstone is interbedded with the sandstones and conglomerates, and displays low-angle cross-stratification and ripple lamination. mudstone is gray-red to gray-green, bentonitic, and poorly exposed. white to variegated gritstone locally marks the base of the chinle formation (figure 6). the grit is poorly sorted and contains rounded to angular, coarse sand to pebble-size grains of quartz. where variegated, the gritstone may represent a paleosol (ancient soil). mudstones and siltstones dominate in the lower slope-forming section. sandstones and conglomerate channels, when found in the lower slope-forming member, are locally mineralized with uranium, vanadium, and copper minerals. the lighter overall color of this lower section is caused by the reduced iron in this part of the formation (figure 6). this reduction of iron commonly extends as much as 3 feet (1 m) into the underlying moenkopi formation. the middle section (black ledge member) is dominated by redbrown sandstone and black, desert varnish-stained conglomerate (figure 6). the sandstones commonly contain scattered logs and branches of petrified wood. lowermost lenses of sandstone are locally mineralized with uranium and copper minerals. the upper contact is gradational into the upper section (church rock member). the upper section is mostly a red-brown sandstone and siltstone (figure 6), but sandstone ledges are more common in the lower part. some beds include distinctive ripple-laminated sandstone, but the bedding in much of this unit is indistinct. blocky, redbrown, fine-grained, well-sorted, thick-bedded sandstone is common in the upper 10 to 30 feet (3–9 m) of the section. the contact of the chinle formation with the overlying wingate sandstone of the glen canyon group is sharp and conformable, commonly being placed below the massive cliff of well-sorted sandstone typical of the wingate (figure 6). no regional channeling or angular unconformity is apparent, and the lowermost part of the wingate does include thin, bedding-parallel sandstone and siltstone beds that suggest continuous deposition across the chinle and wingate contact. the chinle and wingate contact was once thought to represent an unconformity, named the j-0 unconformity, at the triassic-jurassic boundary (pipiringos and o’sullivan, 1978). however, the wingate contains beds of both triassic and jurassic age (molina-garza and others, 2003; lockley and others, 2004; lucas and others, 2005). the j-0 unconformity, if it exists here, is within the triassic and likely represents a short period of time (doelling and others, 2010). upper triassic-lower jurassic rocks the upper triassic-lower jurassic wingate sandstone is the prominent cliff below the green river overlook (figure 6). the wingate forms red-brown, nearly vertical cliffs streaked and stained with desert varnish on weathered surfaces. the chinle and moenkopi slopes below the cliff are commonly littered with large blocks of the wingate sandstone (figure 6). the wingate is ordinarily described as one massive unit because partings or bedding planes are rare except near the base of the formation. the wingate consists mostly of light-orange-brown, moderate-orange-pink, or pale-red-brown, fine-grained, well-sorted, cross-stratified sandstone. the rock is usually well cemented and well indurated; weathered exposures are nearly smooth. jointing and rockslides are common due to unstable swelling clays in the underlying chinle formation. the eolian wingate was deposited in another great erg that extended from the four corners area to north-central utah (figure 7e), as indicated by the high-angle cross-stratified sandstone. however, bedding-parallel sandstone beds near the base of the formation suggest that fluvial processes were still a large part of its early depositional history (doelling and others, 2010). the wingate sandstone ranges in thickness from 250 to 400 feet (60–120 m) in the island in the sky district (hintze and kowallis, 2009). the contact of the wingate sandstone with the overlying lower jurassic kayenta formation is generally sharp and conformable as seen from a distance, but difficult to place close-up. generally, the line is placed at the horizon where the smooth cliff is replaced by thick cliffy ledges (figure 6). lower jurassic rocks with the exception of the modern (quaternary) unconsolidated sediments, lower jurassic strata form the top of the geologic column at the green river overlook (figures 5 and 6). these strata include the upper part of the wingate sandstone, kayenta forma10 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 tion, and navajo sandstone, and were deposited between 205 and 187 ma. the kayenta represents a floodplain environment (figure 7f) whereas the navajo a sand-dune desert. a few sandstone beds in the upper kayenta were also deposited by wind. kayenta formation the kayenta formation caps most of the upper benches at the green river overlook; the viewpoint is located on this unit (figures 4 through 6); it is rarely completely exposed. the kayenta consists mostly of stream-deposited sandstone lenses, with lesser amounts of eolian sandstone, intraformational conglomerate, siltstone, and shale. the unit is primarily red-brown, but individual lenses and beds vary considerably in color; some are purple, lavender, tan, orange, or white. in outcrop, the kayenta is ledgy and step-like (figure 6). sandstone in the kayenta exhibits both high-angle and low-angle cross-stratification. some lenses display channeling, current ripple marks, and rare slump features. the grain size is more variable than in the wingate and navajo, ranging mostly from fine to medium. siltstone, shale, and intraformational conglomerate appear as partings or are interlayered with the sandstone. the kayenta formation was deposited in a sandy braided river system although the environment was arid. perennial streams flowed west from the remaining ancestral rockies (figure 7f) and the appalachians far to the east (lynds and hajek, 2006; blakey and ranney, 2008). floodplains (overbank deposits) formed adjacent to active channels. the kayenta is 100 to 300 feet (30–90 m) thick in the island in the sky district (hintze, 2005). the upper contact is mostly sharp, but intertonguing between the kayenta and the overlying navajo sandstone is common. however, this contact is only observed in a few places in this area (doelling and others, 2010). navajo sandstone the navajo sandstone is mostly seen as cliffy to rounded barerock exposures (figure 6). the navajo is a mostly orange to lightgray, mostly fine grained, generally well-sorted, massive sandstone. high-angle, cross-stratificated laminae lie as much as 30 degrees from the true attitude of the unit. navajo cross-beds etch out in relief in contrast to the smooth-weathering habit of the wingate sandstone. the top of the navajo is not exposed in the island in the sky district. the eolian navajo is a classic example of a major sahara-like erg environment. structural and geologic history regional setting canyonlands national park is located in the paradox basin of southeastern utah and southwestern colorado. the basin is an elongate, northwest-southeast-trending evaporitic basin that predominantly developed during pennsylvanian (desmoinesian) time about 330 to 310 ma. during the pennsylvanian, a pattern of basin and fault-bounded uplifts developed in the area of what is now utah to oklahoma. this pattern included the uncompahgre highland of eastern utah and western colorado and the adjacent paradox basin. rapid basin subsidence, particularly during the pennsylvanian and into the permian, accommodated deposition of large volumes of deeper basin evaporite and marine sediments which intertongue with basin-margin, non-marine arkosic material shed from the mountain area to the northeast (hintze and kowallis, 2009). later, the uncompahgre highland was eroded down during the triassic and jurassic. the area was uplifted again during the late cretaceous (about 70 ma) and early tertiary (about 40 ma) laramide orogeny (a regional mountain-building event) to form the uncompahgre uplift observed today (hintze and kowallis, 2009 and references therein). the region continued to lie near sea level, but received continental deposition through early jurassic time. because no bedrock units younger than the early jurassic navajo sandstone are present in the island in the sky district and surrounding area, no record exists here of the geologic events that happened in the interval from 187 to 1.5 ma. the area likely received continental deposition during the remainder of the jurassic and the early cretaceous and shallow-marine deposition during the late cretaceous, as attested in surrounding areas. late cretaceous–early tertiary uplifts and intrusions large uplifts and basins developed in the colorado plateau during the laramide orogeny between the latest cretaceous (maastrichtian, about 70 ma) and the eocene (about 38 ma). the island in the sky district is located on the northern end of the broad laramide-age monument uplift, which is responsible for exposing the impressive stratigraphic section and the incredible view at the green river overlook (figures 2 and 8). other nearby laramide features include the san rafael swell to the west, the circle cliffs uplift and henry mountains basin to the southwest, the uinta basin to the north, and a rejuvenated uncompahgre uplift to the northeast (figure 8). the regional dip of strata in the island in the sky district is 2° to 4° northward. the henry mountains (figure 2), on the horizon to the southwest, are the type locality for laccolithic intrusions as first described by the famous geologist g.k. gilbert (1877). the henry mountains are a chain of five dome-shaped peaks that represent a complex of granitic (granodiorite porphyry) rocks intruded into the jurassic morrison formation and overlying cretaceous formations about 28 to 25 ma during the oligocene (doelling, 1975; hunt, 1980; nelson, 1998). each mountain consists of a laccolith with multiple sills (nelson, 1998). these peaks range from 7930 to 11,522 feet (2417–3512 m) above sea level and were glaciated during the pleistocene ice ages (hunt,1980). t.c. chidsey, jr., and h.h. doelling green river overlook, canyonlands national park 11 late tertiary–quaternary regional uplift and erosion the colorado plateau began rising in late cenozoic time during the miocene epoch (23 ma) (hunt, 1956; lucchitta, 1979; hintze and kowallis, 2009). this regional uplift changed the landscape from one of deposition to one of massive erosion. several thousand feet of sedimentary rocks have been removed by the erosive processes of mass wasting, wind, and running water. most of this material has been carried to the sea by the colorado river system. the green and colorado rivers probably came into existence synchronously with the uplift of the colorado plateau. as the plateau rose, the rivers and their tributaries rapidly cut into the strata. the incision rate near the center of the colorado plateau (lee’s ferry, arizona) is calculated at about 35 centimeters per thousand years (14 in./k.y.) (pederson and others, 2013). this incision produced stillwater canyon and the countless other smaller canyons visible from the green river overlook. the green river is classified as a superposed stream. it crosses structures formed long before it began to flow, such as split mountain to the north, east of the town of vernal, utah. when the colorado plateau rose, the ancestral green river and its tributaries flowed through meandering channels in wide valleys on easily eroded rocks such as the now-removed cretaceous mancos shale. once these river channels were established, they later became superimposed and entrenched into resistant rocks such as the wingate and navajo sandstones. when the green river reached the softer sedimentary rocks of the chinle and moenekopi formations, stillwater canyon widened until the river eroded down to the white rim sandstone. again, the canyon likely narrowed before the river encountered the mudstone of the organ rock formation near its current depth resulting in stillwater canyon widening, the process occurring today as observed from the overlook. acknowledgments support for this paper was provided by the utah geological survey (ugs). we thank the staff at canyonlands national park for providing access to park lands. cheryl gustin, jay hill, and lori steadman of the ugs drafted figures. this paper was carefully reviewed by michael d. vanden berg, stephanie m. carney, michael d. hylland, and bill keach of the ugs, along with the editors of this publication. their suggestions and constructive criticism greatly improved the manuscript. references baars, d.l., 2010, geology of canyonlands national park, utah, in sprinkel, d.a., chidsey, t.c., jr., and anderson, p.b., editors, geology of utah’s parks and monuments (third edition): utah geological association publication 28, p. 61–83. blakey, r., and ranney, w., 2008, ancient landscapes of the colorado plateau: grand canyon, arizona, grand canyon association, 156 p. doelling, h.h., 1975, geology and mineral resources of garfield county, utah: utah geological and mineral survey bulletin 107, 175 p., 1 plate, scale 1:250,000. doelling, h.h., chidsey, t.c., jr., and benson, b.j., 2010, geology of dead horse point state park, utah, in sprinkel, d.a., chidsey, t.c., jr., and anderson, p.b., editors, geology of utah’s parks and monuments (third edition): utah geological association publication 28, p. 409–428. gilbert, g.k., 1877, report on the geology of the henry mountains: u.s. geographical and geological survey of the rocky mountain region, government printing office, 160 p. hintze, l.f., 1980, geologic map of utah: utah geological survey map m-a-1, 2 sheets, scale 1:500,000. figure 8. major uplifts (orange) and basins (green) created during the paleocene-eocene laramide orogeny. grb = green river basin, um = uinta mountains, ub = uinta basin, uu = uncompahgre uplift, srs = san rafael swell, ku = kaibab uplift, kb = kaiparowits basin, ccu = circle cliffs uplift, hmb = henry mountains basin, and mu = monument uplift which, at the north end, includes the location of the green river overlook and the island in the sky district of canyonlands national park. 12 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 hintze, l.f., 2005, utah’s spectacular geology—how it came to be: provo, utah, brigham young university geology studies 8, 203 p. hintze, l.f., willis, g.c., laes, d.y.m., sprinkel, d.a., and brown, k.d., 2000, digital geologic map of utah: utah geological survey map 179dm, scale 1:500,000. hintze, l.f., and kowallis, b.j., 2009, geologic history of utah: provo, utah, brigham young university geology studies special publication 9, 225 p. hunt, c.b., 1956, cenozoic geology of the colorado plateau: u.s. geological survey professional paper 279, 99 p. hunt, c.b., 1980, structural and igneous geology of the henry mountains, utah, in picard, m.d., editor, henry mountain symposium: utah geological association 8, p. 25–106. lockley, m.g., lucas, s.g., hunt, a.p., and gaston, r., 2004, ichnofaunas from the triassic-jurassic boundary sequences of the gateway area, western colorado—implications for faunal composition and correlations with other areas: ichnos, v. 11, no. 1, p. 89–102. lucas, s.g., tanner, l.h., and heckert, a.b., 2005, tetrapod biostratigraphy and biochronology across the triassic-jurassic boundary in northeastern arizona, in heckert, a.b., and lucas, s.g., editors, vertebrate paleontology in arizona: new mexico museum of natural history and science bulletin 29, p. 83–92. lucchitta, i., 1979, late cenozoic uplift of the southwestern colorado plateau and adjacent colorado river region: tectonophysics, v. 61, p. 63–95. lynds, r., and hajek, e., 2006, conceptual model for predicting mudstone dimensions in sandy braided-river reservoirs: american association of petroleum geologists bulletin, v. 90, no. 8, p. 1273–1288. molina-garza, r.s., geissman, j.w., and lucas, s.g., 2003, paleomagnetism and magnetostratigraphy of the lower glen canyon and upper chinle groups, triassic-jurassic of northern arizona and northeastern utah: journal of geophyscial research, v. 108, no. b4, p. 1–24. nelson, s.t., 1998, reevaluation of the central colorado plateau laccoliths in light of new age determinations: u.s. geological survey bulletin 2158, p. 37–39. pederson, j.l., cragun, w.s., hidy, a.j., rittenour, t.m., and grosse, j.c., 2013, colorado river chronostratigraphy at lee’s ferry, arizona, and the colorado plateau bull’s-eye of incision: geology, v. 41, no. 4, p. 427–430. pipiringos, g.n., and o’sullivan, r.b., 1978, principal unconformities in triassic and jurassic rocks, western interior united states—a preliminary survey: u.s. geological survey professional paper 1035-a, 29 p. powell, j.w., 1895, canyons of the colorado: new york, argosy-antiquarian ltd., 400 p. uga-geosite-kettler-loope-wonderstone.indd utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors richard m. kettler1 and david b. loope1 1earth & atmospheric sciences, university of nebraska lincoln, ne 68588-0340; rkettler1@unl.edu cover image: multistory exposure of shinarump sandstone in quarry with thin interbeds of less resistant mudstone. the origin of shinarump wonderstone, hildale, washington county 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 r.m. kettler and d.b. loope the origin of shinarump wonderstone 3 introduction southern utah’s “wonderstone” is shinarump sandstone, variably cemented and stained with iron oxide, forming intricate patterns reminiscent of landscapes. it is cut and sold as absorbent drink coasters and decorative objects, and is seen in rock shops across the country. the wonderstone pattern comprises thick bands of iron oxide mineralization that fills pore space (referred to as iron oxide cement or ioc) and more delicate bands of iron oxide mineralization that coats sand grains but does not fill pore space (referred to as iron oxide stain or ios) (figure 1). the wonderstone pattern is of interest to geologists because it formed after the shinarump sandstone was deposited from iron that was transported in aqueous solution. the iron that now resides in the cement and stain occurs as oxidized iron (iron-iii) minerals (e.g., goethite and hematite). significant amounts of iron-iii can be transported in aqueous solution only under very unusual conditions. on the other hand, if an electron is added to iron-iii, the resultant reduced iron (iron-ii) can be transported readily in aqueous solutions. but iron-ii forms a different group of minerals, typically pyrite (fes2) and siderite (feco3) that do not have the characteristic red color of the wonderstone cement and stain. how was the iron that now resides in the wonderstone transported to its current location? what was the chemical mechanism for removing the iron from natural waters and fixing it as iron-iii minerals? the typical explanation for the wonderstone pattern is that the bands of iron oxide cement and stain are liesegang bands. liesegang bands were discovered originally by chemists and are a form of chemical self-organization that produces bands of insoluble material from the mixing of two solutions. the conventional interpretation is that when pyrite is exposed to oxygen-rich groundwater the pyrite will dissolve, producing a strongly acidic, iron-rich solution. iron-iii will migrate in solution toward the source of oxygen. this aqueous iron-iii will then precipitate as the solution is neutralized to form the liesegang bands of iron oxide cement. this conventional interpretation was developed before geologists recognized the importance of microbes to processes that occur at low temperature. our interpretation is that iron was introduced to the rock as ironii shortly after sediment deposition and formed the mineral siderite. as the colorado plateau experienced uplift more oxygen-rich groundwaters invaded the shinarump sandstone. iron-oxidizing bacteria thrive by transferring an electron from iron-ii to oxygen to make iron-iii. energy is released during this transfer that the bacteria use to survive (in the same way that humans transfer electrons from the carbon in food to oxygen and survive using the energy released in those reactions). the ioc was produced through dissolution of siderite followed by oxidation of aqueous iron-ii by microbes at a succession of oxidation-reduction interfaces. the ioc bands mark the position of interfaces where iron-oxidizing bacteria converted aqueous iron ii to iron-iii with a consequent precipitation of iron iii oxide. we consider the iron oxide staining, on the other hand, to be liesegang produced by the inter-diffusion of iron ii and oxygen after the bands of cement were produced. see kettler and others (2015) for a more complete description of the processes. the outcrops and blocks of wonderstone in this quarry provide a good summary of the evidence that falsifies the pyrite oxidation hypothesis in favor of our hypothesis. gps locations quarry location: n 37.062482, w 113.130681. intersection of 400 n. and highway 59: n 37.083304, w 113.110919. driving instructions and hiking suggestions this quarry is located 9 miles (14.5 km) northwest of hildale, in southernmost utah (figure 2). our directions begin at the intersection of 400n and utah highway 59 (state street). as of 2019 this intersection is shown on google earth as the intersection of little creek mesa road with highway 59. the sign marking the intersection reads 400n and state street. this intersection is located 1.6 highway miles (2.6 km) southeast of the intersection of apple valley way and utah state highway 59 (site of a convenience store in apple valley), whereas the distance from the utah-arizona state line is 8.7 miles (14 km). proceed west on 400n for 0.25 miles (0.4 km) where the road will turn south. proceed south on 7400e for 0.5 miles (0.80 km) to little creek mesa road and turn to the southwest. after 0.9 miles (1.4 km) on little creek mesa road, take the left fork and leave little creek mesa road. drive an additional 0.25 miles (0.4 km) to the southwest and park. figure 1. example of wonderstone mineralization. the thick reddish-brown bands are iron-oxide cement whereas the wispy reddish-brown lines are iron-oxide stain. 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 th ere are a number of wonderstone quarries located on both public and private lands in southern utah. th is particular quarry was operated by a company that sells absorbent, decorative, stone coasters. th e quarry is still private property with public access (as of december 2019) administered by the utah department of natural resources (dnr) under the walk-in-access program (http://wildlife.utah.gov/walkinaccess/). th e property is designated as midnightpropwashington2 in the walk-in-access program database. visitors should adhere to all regulations prescribed by the utah dnr. visitors are required to obtain a walk-in access authorization number and present it if challenged. th e authorization number is free of charge and available at http://wildlife.utah.gov/ walkinaccess/authorization.php. all outcrops of interest are within 500 yards (460 m) of the parking area. th e area is relatively free of hazards but visitors should be cognizant of hazards endemic to abandoned quarries (e.g., sharp edges on scrap, broken glass, loose rock, abandoned chemicals and equipment). sturdy shoes, long pants, long sleeves, head and eye protection are recommended. geologic context th e site is located in the grand staircase of the colorado plateau. th e grand staircase comprises fi ve step-like layers of sedimentary rock that are encountered as one moves north from the northern rim of the grand canyon to bryce canyon national park. state highway 59 connects hildale and hurricane, utah, traversing the lowermost tread of the grand staircase (fi gures 2 and 3). to the northeast of the road are the vermillion cliff s of the moenave and kayenta formations (fi gure 4), whereas southwest of the road the moenkopi formation of the chocolate cliff s faces south. th e resistant shinarump member of the chinle formation forms the top of the chocolate cliff s. th e shinarump was deposited in an extensive braided stream system that fl owed from southeast to northwest across the southern portion of north america (dickinson and gehrels, 2008). th e unit locally has iron-oxide cementation, contains abundant fossil wood, and has produced signifi cant amounts of uranium and copper mineralization. th e beds in the area of the quarry are subhorizontal with regional dips of 1-2° to the northeast. the shinarump sandstone and iron-oxide mineralization four diff erent types of iron oxide mineralization can be observed in this quarry. (1) iron oxide mineralization that fi lls pore space and forms bands that are cm thick. th ese bands of ioc may cross-cut or be parallel to bedding. (2) between the bands of ioc it is typical to observe iron oxide stain that coats mineral grains but does not fi ll pore space. th ese fi ner features may give the rock a wood-grain appearance or mimic cross-bedding (fi gure 1). (3) disseminated patches of iron oxide cement that locally assumes rhombic shapes. (4) hollow concretions of iron oxide-cemented sandstone that may contain small fragments of mudstone. th e accumulation of centimeter-thick masses of iron oxide cement is evidence that at least some of the iron was introduced to the rock as iron-ii while in aqueous solution. iron oxide mineralization is controlled by the presence of fractures excellent exposures in the southeastern quarry wall demonstrate that the shinarump is a multistory sandstone with little preservation of fl oodplain muds (fi gure 5). closer inspection reveals both disseminated and fracture-controlled iron-oxide mineralization. in the south wall of the quarry, a series of vertical joints fracture the shinarump sandstone. iron-oxide cement is concentrated in the sandstone immediately adjacent to these fractures but does not fi ll the fractures (fi gure 6). th e relationship between the iron oxide cementation and fracturing is strong evidence that the fractures formed before the iron oxide mineralization. we interpret the accumulations along these fractures and those sub-parallel accumulations deeper in the sandstones to have been paleo-oxidation fronts. reduced siderite dissolved and the resulting aqueous iron-ii migrated to the oxidation front where microbes oxidized iron-ii to iron-iii. th e iron-iii immediately precipitated as insoluble iron-iii oxide. in some blocks we can see a succession of these paleo-oxidation fronts that have been preserved by successive accumulations of iron oxide (fi gure 7). figure 2. location of quarry described in this report. r.m. kettler and d.b. loope the origin of shinarump wonderstone 5 figure 3. geologic map with quarry location marked. geology modified from biek and others (2010). figure 4. mesozoic and cenozoic stratigraphy. quarry is located in the shinarump member of the chinle formation ( ). r.m. kettler and d.b. loope the origin of shinarump wonderstone 6 siderite was the most likely iron oxide precursor mineral this quarry is unusual in that it is one of the few places where one can see both the wonderstone as well as evidence for the precursor iron mineralization. in the quarry walls and quarried blocks we observe disseminated patches of iron oxide (figure 8). many of the disseminated patches have a rhombohedral form that we interpret to be iron oxide pseudomorphs after poikilotopic siderite (feco3) cements. siderite is a typical byproduct of organic matter fermentation during diagenesis, a process that is typical in stream sediments (burgess and others, 2016). most groundwaters that saturate river sediments contain small amounts of sulfate compared to the amount of sulfate found in ocean water. whereas iron in oceanic sediment may be converted largely to pyrite, pyrite in river sediments tends to be ubiquitous (always present) but present in lesser abundance than siderite. sufficient sulfate was present here during diagenesis that small pyrite concretions formed and can be observed in the quarry face. it is important to note that the pyrite, which has a cubic crystal structure, is preserved whereas the precursor to the rhombohedral pseudomorphs has been oxidized. we interpret the morphologies and the continued preservation of pyrite as evidence that the rhombohedra did not have a pyrite precursor. the final bit of evidence in support of a siderite precursor rather than pyrite is the presence of the bands of iron-oxide cement. pyrite oxidation occurs when pyrite is attacked by aqueous iron iii; the aqueous iron-iii contacts the surface of the pyrite and an electron is transferred from sulfur in the pyrite to the iron. the iron-ii that forms during this process is then oxidized back to iron-iii by iron-oxidizing microbes. the rate of pyrite oxidation will therefore be slow unless the pyrite is bathed in an acidic, oxidizing fluid. the oxidizing environment is necessary to convert aqueous iron-ii to iron-iii. because iron-iii oxides are highly insoluble in neutral solutions, elevated concentrations of aqueous iron-iii can exist only in acidic waters. the presence of the bands figure 5. multistory exposure of shinarump sandstone in quarry with thin interbeds of less resistant mudstone. figure 6. shinarump sandstone exposure illustrating fracture control of iron-oxide mineralization. note iron-oxide cement and staining parallel to vertical fracture right of hammer. figure 7. wonderstone mineralization advancing (from left to right) on disseminated rhombohedral patches of iron oxide cement (small dark red dots on right half of photo) that are interpreted to be oxidized pseudomorphs after siderite. figure 8. rhombic patches of iron oxide cement (circled) interpreted to be pseudomorphs of iron oxide after siderite. r.m. kettler and d.b. loope the origin of shinarump wonderstone 7 of iron oxide indicate that solid iron-iii oxide precipitated upon oxidation. iron-iii that was trapped in mineral form as iron oxide would have been unavailable to attack pyrite. siderite, on the other hand, will undergo simple dissolution to produce aqueous iron-ii and carbonate. subsequent oxidation of iron-ii to iron-iii and precipitation of iron-iii oxides would not retard the dissolution of siderite, rather it would enhance it. the conversion of iron-ii to iron-iii removes the iron required to make siderite. although the system will be driven to reach siderite saturation by dissolving more siderite, continued depletion of iron-ii from the solution will prevent siderite saturation. siderite is, therefore, doomed to dissolve in oxidizing systems rattle stones occur with other coarse clastics if one looks carefully in the quarry—particularly in areas where pebble and cobble-sized clasts are present in the thin shinarump mudstone beds or at the base of sandstone channels—one can find ovoid concretions. upon shaking these concretions, one can typically hear a rattling sound from the concretion interior. these concretions are known as rattle stones (hollow iron oxide concretions; van der burg, 1969; loope and others, 2012; burgess and others, 2016). although very little of the floodplain mud is preserved, the presence of rattle stones is evidence that these muds contained abundant organic matter. rattle stones are the product of early diagenetic cementation of mud by siderite. these siderite concretions may then be excavated and redeposited as the fluvial system migrates across the floodplain. during later oxidation, dissolution of siderite and microbially-mediated oxidation of ironii converts the siderite concretion to a rinded iron-oxide concretion. the siderite cement in the mudstone concretions is an early displacive cement. in other words the cement forms before the mud is converted to rock and pushes the mineral grains apart as it precipitates. the siderite concretion, therefore, consists of silt and clay-sized detritus that is floating in a mass of siderite cement. dissolution of this cement and conversion to iron oxide leaves empty space that may contain fragments of mudstone. thee mudstone fragments form the rattles within the rattlestone concretions. acknowledgments our wonderstone research was supported in part by the nasa nebraska space grant program. a careful review by peter mozley (new mexico tech) improved the manuscript. references antal, t., droz, m., mangin, j., rácz, z., and zrinyi, m., 1998, derivation of the matalon-packter law for liesegang patterns: journal of chemical physics, v. 109, no. 21, p. 9479-9486 (doi: 10.1063/1.477609). biek, r.f., rowley, p.d., hayden, j.m., hacker, d.b., willis, g.c., hintze, l.f., anderson, r.e., and brown, k.d., 2010, geologic map of the st. george and east part of the clover mountains 30' x 60' quadrangles, washington and iron counties, utah: utah geological survey map 242 dm, 107 p., 3 plates, scale 1:100,000. burgess, d.t., kettler, r.m., and loope, d.b., 2016, the geologic context of wonderstone—a complex, outcrop-scaled pattern of iron-oxide cement: journal of sedimentary research, v. 86, no. 5, p. 498-511 (doi: 10.2110/jsr.2016.35). disckinson, w.r., and gehrels, g.e., 2008, u-pb ages of detrital zircons in relation to paleogeography—triassic paleodrainage networks and sediment dispersal across southwest laurentia: journal of sedimentary research, v. 78, no. 12, p. 745-764 (doi: 10.2110/jsr.2008.088). jablczynski, k., 1923, la formation rhythmique des précipités: les anneaux de liesegang, bulletin de la société chimiquede france, v. 33, p. 1592-1597. kettler, r.m., loope, d.b., weber, k.a., and niles, p.b., 2015, life and liesegang: outcrop-scale microbially-induced diagenetic structures and geochemical self-organization phenomena produced by oxidation of reduced iron: astrobiology, v. 15, no. 8, p. 616-636 (doi: 10.1089/ast.2015.1305). loope, d.b., kettler, r.m., weber, k.a., hinrichs, n.l., and burgess, d.t., 2012, rinded iron oxide concretions—hallmarks of altered siderite masses of both early and late diagenetic origin: sedimentology, v. 59, no. 6, p. 1769-1781 (doi: 10.1111/j.13653091.2012.01325.x). van der burg, w.j., 1969, the formation of rattle stones and the climatological factors which limited their distribution in the dutch pleistocene, 1. the formation of rattle stones: paleogeography, paleoclimatology, paleoecology, v.6, p.105-124 (doi: 10.1016/0031-0182(69)90007-8). uga-geosite-biek-cedar-breaks.indd utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors robert f. biek1 and peter d. rowley2 1utah geological survey, p.o. box 146100, salt lake city, ut 84114-6100, bobbiek@utah.gov 2geologic mapping inc., p.o. box 651, new harmony, ut 84757 cover image: view east to north view overlook at cedar breaks national monument showing upper claron and lower brian head strata. cedar breaks national monument, north view overlook 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e geosites refl ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: biek, r.f., and rowley, p.d., 2019, cedar breaks national monument, north view overlook, in milligan, m., biek, r.f., inkenbrandt, p., and nielsen, p., editors, utah geosites: utah geological association publication 48, 11 p., https://doi. org/10.31711/geosites.v1i1.50. presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president r.f. biek and p.d. rowley cedar breaks national monument north view overlook 3 introduction cedar breaks national monument straddles the western rim of the markagunt plateau, and, at over 10,000 feet (3050 m) in elevation, offers spectacular views westward into the adjacent great basin (figure 1). the heart of the monument consists of a series of overlooks along a beautiful, 4-mile-long (6 km), west-facing amphitheater that is eroded into the plateau’s margin. these overlooks are adorned with a fantastic variety of hoodoos and sculpted fins eroded from alternating resistant and soft, jointed and fractured layers of the claron formation. of these overlooks, the north view overlook encapsulates the entire geologic story preserved at cedar breaks—a history not only of the enigmatic claron formation, but also of the demise of its depositional basin and eventual burial by volcanic sediments of the brian head formation. location the north view overlook is near the north end of cedar breaks national monument on utah highway 143. the overlook itself is open year round, but for much of the year during the extended winter season, it is only accessible via parowan canyon or panguitch; the park and its facilities are closed during the winter. during summer months, the overlook is also accessible from cedar city via utah highways 14 and 148. the overlook is at an elevation of nearly 10,500 feet (3200 m); a short, level paved trail leads from the parking area to the overlook at latitude 37° 39' 23.6", longitude 112° 50' 0.4". figures 2, 3, and 4 provide a geologic map, stratigraphic column, and cross section of the overlook area. stratigraphy claron formation the claron formation contains some of the most visually arresting rocks in southwestern utah—they form the pink cliffs, the uppermost riser and tread of the grand staircase, that great sequence of resistant and non-resistant sedimentary strata that steps up from the grand canyon northward to bryce canyon and cedar breaks. the claron’s multi-hued pink, orange, and white sandstone, siltstone, mudstone, and minor limestone were deposited in stream, floodplain, and minor lake environments in early tertiary time. these environments were in a high-elevation basin bounded by highlands created by compressional deformation during a mountain-building event known as the laramide orogeny. the pink member was deposited on floodplains and alluvial plains with channel conglomerates along basin margins, whereas the white member is lacustrine near the center of the basin and interbedded with alluvial-plain strata along its margins (goldstrand, 1990, 1991, 1992, 1994; bown and others, 1997). the formation is divided into two members, a lower pink member and a comparatively thin upper white member (figure 5). the claron’s upper white member is readily divisible into four informal parts: two, white, cliff-forming limestone units separated by intervening colorful, variegated mudstone units much like that of the underlying pink member (rowley and others, 2013; biek and others, 2105). figure 1. view west from the north view overlook. cedar canyon is at left, beyond which is cedar valley and mountains bordering its west side. 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 figure 2. geologic and location map of the overlook area. the overlook rests on the lowermost brian head formation (pink on map), but it is the claron formation (orange and yellow on map), eroded into a vast amphitheater, that visitors come to see. note the large area of modern landslides (yellow with red triangle pattern) that resulted from failure of the brian head formation. the brian head ski area, and the southern end of the yankee meadows graben, is in the upper left part of the map. the seldom visible contact of the brian head and claron formations is well exposed in a steep, cliffy area several hundred yards southeast of the overlook. that contact marks the inception of volcanism in southwestern utah. map unit names are shown on figure 3. various young surficial deposits are shown in light yellow. from rowley and others (2013). l 3000 ft brian head peakbrian head peakbrian head peak cedar breaks north view overlook cedar breaks cedar breaks cedar breaks cedar breaks tcp qc tcwt ti qms tcwu ti tcwm tbh tql ti tcwm qms qacf tbh qac tm(ti) tcwl qc tm qmt qc tcp qac qac qms tm qmt qacf qms qacf qms ti tcwl ti qmt qmt qacfo qmsh tbh qms tbh tm(ti) qes tm(ti) qes tbh qac tcwutcwm td tm(ti) qal black fault le dge bh071808-4, -5, -6 bh071808-7 bh071808-2, -3 bh071808-1 bh062310-1 2 6 42 53 37 14 71 16 contact normal fault gravity slide fault sinkhole bedding attitude15 r.f. biek and p.d. rowley cedar breaks national monument north view overlook 5 figure 4. cross section through cedar breaks national monument from rowley and others (2013). note that line of cross section lies just south of and covers a wider area than that shown in figure 2. cretaceous strata: ksj = john henry member and ksd = drip tank member of straight cliffs formation; kw = wahweap formation, kgc and kwcs = grand castle formation and capping sandstone member of wahweap formation; km = cretaceous strata on markagunt plateau. tertiary strata: tcp = pink member and tcw = white member of claron formation; tm = markagunt gravity slide. figure 3. stratigraphic column showing rock units of the overlook area. the markagunt megabreccia is the deposit of the markagunt gravity slide. from rowley and others (2013). miocene markagunt megabreccia tm(ti) tm(tnw) tm(tbh) tm 450 (140) 300 (90) 40 (12) 150 (45) mount dutton fm. td 10 (3) tm(tqcb) leach canyon formation tql 55–100 (17–30) isom formation ti 350 (107) oligocene wah wah springs fm. tnw 3–8 (1–3) brian head formation tbh 500 (150) eocene uppermost unit middle unit upper limestone unit tcw lower limestone unit tcwm tcwl 47 (14) tcwu tcwt w hi te m em be r 440 (135) 310 (94) te r ti a r y paleocene pink memberc la ro n fo rm at io n tcp 1000 (305) 109 (33) 45–60(14–18) 23.0 33.9 ?55.8 50+ (15+) about 20–22 ma 23.8 ma 26–27 ma 29.5 ma 35 ma 36 ma ma system and stage lithology age map unit symbol map thickness feet (meters)series stratigraphic column miocene markagunt megabreccia tm(ti) tm(tnw) tm(tbh) tm 450 (140) 300 (90) 40 (12) 150 (45) mount dutton fm. td 10 (3) tm(tqcb) leach canyon formation tql 55–100 (17–30) isom formation ti 350 (107) oligocene wah wah springs fm. tnw 3–8 (1–3) brian head formation tbh 500 (150) eocene uppermost unit middle unit upper limestone unit tcw lower limestone unit tcwm tcwl 47 (14) tcwu tcwt w hi te m em be r 440 (135) 310 (94) te r ti a r y paleocene pink memberc la ro n fo rm at io n tcp 1000 (305) 109 (33) 45–60(14–18) 23.0 33.9 ?55.8 50+ (15+) about 20–22 ma 23.8 ma 26–27 ma 29.5 ma 35 ma 36 ma ma system and stage lithology age map unit symbol map thickness feet (meters)series stratigraphic column thin surficial deposits not shown 11,000 10,000 9,000 8,000 7,000 6,000 e le va tio n in fe et b west b' east 11,000 10,000 9,000 8,000 7,000 6,000 c e d a r b r e a k s n a t i o n a l m o n u m e n t cedar breaks kw ksj tcp kgc & kwcs km kw ksj cretaceous, undivided kgc & kwcs km tcp tm cretaceous, undivided sr 143 adams canyon ksdksd m a r k a g u n t p l a t e a u 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 what makes the claron so interesting is that most of its finegrained floodplain and stream deposits were extensively modified by soil-forming processes such that it now represents a stacked sequence of paleosols (ancient soils) (mullett and others, 1988a, 1988b; mullett, 1989; davis and pollock, 2010). the formation locally contains abundant crayfish burrows and trace fossils of ants, wasps, and bees that record nest activity during soil formation (bown and others, 1995a, 1995b, 1997). crayfish burrows (figure 6) in claron strata of the markagunt plateau record a relatively deep and highly fluctuating water table in the pink member, and a relatively shallow water table in alluvial parts of the white member (hasiotis and bown, 1997). but as renowned as the claron is for its spectacular scenery, it is equally known among geologists for being frustratingly unfossiliferous (apart from trace fossils such as burrows) and devoid of datable volcanic ash or age-constraining detrital zircons (a hard, resistant mineral that contains trace amounts of uranium and thorium, making it amenable to radiometric dating). the lack of fossils is disheartening in that the formation was deposited during a period of abrupt mammalian diversity that accompanied a gradual 6-million-year-long global warming trend culminating in the eocene climatic optimum (see, for example, röhl and others, 2000). oxidation, burrowing, and weathering associated with the formation of paleosols doubtless led to the destruction of what otherwise may have been highly fossiliferous beds. the lack of volcanic ash stems directly from the near cessation of volcanism that accompanied relatively rapid, “flat-slab” subduction at the western margin of north america during the paleocene to early eocene (65 to 45 million years ago), when the claron was deposited (see, for example, dickinson, 2006, and humphreys, 2009). cedar breaks national monument is justly famous for its hoodoos that adorn the western escarpment of the markagunt plateau. hoodoos form by relatively rapid erosion of alternating resistant figure 5. a. view east to north view overlook at cedar breaks national monument showing upper claron and lower brian head strata. the north view overlook is on basal strata of the gray volcaniclastic unit of the brian head formation (tbh). b. view northwest to overlook, showing variegated uppermost claron strata (tcwt) and contact with brian head formation (tbh). the top of the claron is marked by a thin, calcareous, pebbly conglomerate that lacks volcanic clasts and thus predates volcanism in the region; this conglomerate marks an unconformity, a surface of erosion and therefore a gap in geologic time where no strata are preserved that might record past environments. tcp = pink member; tcwl = lower limestone unit of the white member, tcwm = middle white unit, tcwu = upper limestone unit, tcwt = variegated, nontuffaceous mudstone, siltstone, and minor sandstone and pebble conglomerate unit of the claron formation. a b figure 6. intensely bioturbated (burrowed) lower part of the pink member of the claron formation at the west edge of the markagunt plateau in first left hand canyon, north of cedar breaks. inset shows close-up of closely packed crayfish burrows. r.f. biek and p.d. rowley cedar breaks national monument north view overlook 7 and less-resistant beds in combination with pervasive jointing, high precipitation rates, and some 200 freeze-thaw cycles per year. freeze-thaw cycles loosen rock, whereas summer thunderstorms work to wash it away. importantly however, new research shows that it is the local stress field that gives rise to the hoodoos; erosional processes act within that stress field to create these fantastic forms (bruthan and others, 2014). the vertical stress imparted by the weight of a resistant caprock holds grains together; as erosion removes some grains, the stress between the remaining grains increases, holding them more tightly together and making it harder to erode the sides of the hoodoos. erosion is thus slowed and the hoodoos tend to “grow” taller over time. in addition, much of the precipitation that falls on the hoodoos quickly evaporates, leaving behind any cement that the water started to dissolve and thus making the rind of the hoodoos more resistant (called case hardening). sinkholes are common in the pink member in the central markagunt plateau (figure 7) (moore and others, 2004; hatfield and others, 2010; biek and others, 2011; rowley and others, 2013). they form due to the region’s abundant precipitation and susceptible rocks—slightly acidic rain and snowmelt serves to dissolve carbonate cement in claron strata, creating karst topography with interconnected open fractures and caves. large sinkholes visible on 1:20,000-scale aerial photographs are plotted on the regional geologic map of biek and others (2015), and doubtless many smaller sinkholes are present. these sinkholes capture local runoff and serve to shunt shallow groundwater rapidly down dip where it emerges as springs, including the large mammoth, asay, and cascade springs (wilson and thomas, 1964; spangler, 2010; weaver, 2010). karst terrain developed in the pink member of the claron formation makes the groundwater of the markagunt plateau particularly susceptible to contamination. conglomerate at boat mesa claron deposition ended by the latest middle eocene, about 38 million years ago, with deposition of the conglomerate at boat mesa above an erosionally beveled surface. at cedar breaks and elsewhere on the markagunt plateau, this conglomerate is thin, 1 to 10 feet (0.3–1 m) thick; still, it marks a widespread unconformity in the claron basin (biek and others, 2015). farther east on the paunsaugunt plateau where the conglomerate is best developed and as much as 100 feet (30 m) thick, this unconformity cuts gently down section from north to south. this pebbly conglomerate contains rounded clasts of chert, quartzite, and claron limestone, but lacks volcanic or intrusive clasts. the lack of igneous clasts suggests that the conglomerate predates the inception of volcanism in this part of southwest utah. still, the conglomerate contains a robust suite of late eocene detrital zircons that yielded a u-pb detrital age of 37.97 +1.78/-2.70 ma (biek and others, 2015). the conglomerate at boat mesa represents deposits of braided-stream channels and minor floodplains incised into the claron formation. brian head formation the north view overlook itself rests on the lowermost brian head formation, named for poor exposures of white volcaniclastic mudstone, siltstone, sandstone, volcanic ash, micritic limestone, and minor conglomerate and multi-hued chalcedony at its type area of brian head peak, visible just 1.5 miles (2.5 km) north of the overlook (see brian head peak geosite). these strata, rich in volcanic ash, were deposited in low-relief fluvial, floodplain, and lacustrine environments (sable and maldonado, 1997). the contact of brian head and claron strata is visible and well exposed a few hundred yards east of the overlook, where a thin, 35.77 ± 0.28 ma rhyolitic air-fall tuff marks the base of the brian head formation (figure 8). several additional radiometric ages from brian head strata from across the region show that it was deposited from about 37 to 30 million years ago (biek and others, 2015). the formation is thus mostly late eocene in age, reaching into the lower oligocene to which it was previously assigned; it forms the base of the volcanic section in southwestern utah. figure 7. sinkhole on the southeast flank of blowhard mountain, immediately west of utah highway 148. the sinkhole formed by dissolution of carbonate cement in the pink member of the claron formation and here propagated upward through overlying markagunt gravity slide residuum. 8 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 the brian head formation is known for abundant trace fossils, including possible crayfish burrows and root traces (golder and wizevich, 2009; golder and others, 2009), but aside from its basal variegated interval it is surprisingly unfossiliferous. it is also known for its colorful beds of chalcedony in various shades of white, gray, yellow, red, black, and brown, all typically with a white weathering rind. the chalcedony, probably derived from remobilization of silica in glass shards from the tuff beds, forms resistant beds commonly 1 to 3 feet (0.3–1 m) thick but locally as much as 8 feet (2.5 m) thick, and may have resulted from silicification of limestone beds (maldonado, 1995; sable and maldonado, 1997; schinkel, 2012). being resistant, the chalcedony commonly litters slopes developed on brian head strata, and it was commonly used for tools and arrowheads by early native americans. above all, however, because of its abundant smectitic clay derived from weathered volcanic ash, the brian head formation is known for its swelling soils and for its susceptibility to landsliding. nearly all exposures on steep hillsides form large landslide complexes, the largest of which fills the 15-mile (24 km) length of yankee meadows graben just north of cedar breaks national monument. the non-resistant, clay-rich brian head formation was also the principal detachment surface for the early miocene (23 to 21 million-year-old) markagunt gravity slide (hacker and others, 2014; see sidney peaks geosite). structure cedar breaks national monument lies at the western edge of the markagunt plateau, on the southwestern margin of utah’s high plateau’s physiographic province. this region is a structural and stratigraphic transition zone between the highly extended great basin to the west and the colorful, mostly flat-lying strata of the colorado plateau to the east. the plateau itself is a gently east-tilted fault block bounded on the west by the active hurricane fault (on the east side of cedar city) and the parowan and paragonah faults (on the south and east sides of parowan valley) (figure 9). between cedar city and parowan, the western margin of the markagunt plateau is cut by high-angle normal faults that create a series of horsts (fault blocks that are uplifted) and grabens (fault blocks that are downthrown) that step down from the plateau to the adjacent great basin (figure 10) (maldonado and others, 1997; biek and others, 2015). the most prominent of these faults is the active hurricane fault at the base of the plateau east of cedar city, which has a down-to-the-west vertical displacement of at least 6000 feet (1800 m) (hurlow, 2002; lund and others, 2007). overall, the collection of horsts and grabens forms a highly faulted relay ramp between the paragonah fault and the hurricane fault (figure 11). cedar breaks, as with bryce canyon national park to the east, is dominated by beautiful and starkly contrasting pink and white sedimentary rocks of the claron formation eroded into picturfigure 8. basal brian head strata just southeast of the north view overlook at cedar breaks national monument (see figure 3 for location). utah geological survey paleontologist don deblieux (standing) and gary hunt (former ugs geologist) are at contact with underlying claron formation. inset shows thin rhyolitic ash (the upper part of which is a deep reddish brown) at base of the brian head formation that yielded a radiometric age on zircon of about 36 million years. west east cedar breaks national monument qt tv tc k k tc tc tv tc k k hurricane fault brian head cedar valley p a u n s a u g u n t p l a t e a u sevier fault paunsaugunt fault m a r k a g u n t p l a t e a u g r e a t b a s i n tropic h i g h p l a t e a u s t r a n s i t i o n z o n e figure 9. schematic west-to-east cross section through the central markagunt and paunsaugunt plateaus. note gentle east dip of strata in the plateaus, which are bounded by the hurricane, sevier, and paunsaugunt faults. cbnm = cedar breaks national monument. qt = quaternary-tertiary basin fill, tv = tertiary volcanic rocks, tc = claron formation, k = cretaceous strata. r.f. biek and p.d. rowley cedar breaks national monument north view overlook 9 esque hoodoos and fi ns. th e uplift by large normal faults of the markagunt plateau resulted in the erosion that is evident today. th e north view overlook provides the opportunity to examine and learn how some of the most colorful, beautiful, and famous sedimentary rocks in utah formed during the eocene, and why they subsequently weathered the way they did. acknowledgments our knowledge of southwestern utah geology comes from combined decades of geologic mapping supported largely by the utah geological survey and u.s. geological survey. as the acknowledgments of our many published geologic maps attest, we are indebted to a great many people for their help over the years. th anks to grant willis, stephanie carney, and mike hylland (all with the utah geological survey) for their insightful reviews, and to jenny erickson (ugs) for draft ing the fi gures. references biek, r.f., moore, d.w., and nealey, l.d., 2011, geologic map of the henrie knolls quadrangle, garfi eld, kane, and iron counties, utah: utah geological survey map 252dm, 2 plates, scale 1:24,000. biek, r.f., rowley, p.d., anderson, j.j., maldonado, f., moore, d.w., hacker, d.b., eaton, j.g., hereford, r., filkorn, h.f., and matyjasik, b., 2015, geologic map of the panguitch 30' x 60' quadrangle, garfi eld, iron, and kane counties, utah: figure 10. major faults of the western markagunt plateau and red hills, and named grabens (shaded) and horsts. cbnm = cedar breaks national monument. from biek and others (2015). 113°00´ 112°30´ 38°00´ 37 °3 7´ 30 ˝ 0 0 5 miles 5 kilometers 143 143 148 14 r e d en o c h g r ab en h i l l s h o r s t northern cedar valley graben hur ric an e fa u l t zo ne iro n sp rin gs th ru st f a ul t p a r o w a n v a l l e y g r a b e n r ed h ills fault zo ne fi d d le r s c an yo n g r ab en m id dl e r id ge fa u l t m ap le r id g e h o r st jo ne s h ill fa ul t su m m it m o u n ta in g r ab en su m m it m ou nt ai n fa ul t parowan graben su ga rlo af m ou nt ai n fa ul t su ga rl oa f g ra be n na va jo r id ge fa ul t na va jo ri dg e ho rs t ratt les na ke can yo n fa ult yankee mead ow s g ra be n blac k le dg e fau lt m a r k a g u n t p l a t e a uparagonah parowan summit pa ra go na h fa u l t pa ro w an can yon h o r st iro n pe ak we st fa u lt ir on peak gra be n iro n peak eas t faul t w es t be ar va lle y fa ult ea st be ar va lle y fa ult be ar va ll ey graben location map normal fault, bar and ball on down-dropped side thrust fault, teeth on upper plate parowan fault 15 salt lake city brian head peak cbnm north view overlook ¯ figure 11. diagram of a relay ramp between parallel strands of a fault zone. th e ramp links displacement between the faults. cbnm = cedar breaks national monument. relay ramp hanging wall footwall markagunt plateau brian head peak cbnm overlook hurricane fault cedar valley parowan valleynormal fault tip dip 10 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geological survey map 270dm, 162 p., 3 plates, scale 1:62,500. bown, t.m., hasiotis, s.t., and brouwers, e.m., 1995a, reassessment of sedimentary paleoenvironments, late paleocene to middle eocene claron formation, paunsaugunt plateau, southwestern utah [abs.]: geological society of america abstracts with programs, v. 27, no. 6, p. a-277. bown, t.m., hasiotis, s.t., genise, j.f., maldonado, f., and bowers, e.m., 1997, trace fossils of hymenoptera and other insects, and paleoenvironments of the claron formation (paleocene and eocene), southwestern utah, in maldonado, f., and nealey, l.d., editors, geologic studies in the basin and range–colorado plateau transition in southeastern nevada, southwestern utah, and northwestern arizona, 1995: u.s. geological survey bulletin 2153, p. 43–58. bown, t.m., hasiotis, s.t., maldonado, f., brouwers, e.m., and eaton, j.g., 1995b, trace fossils of ants, wasps, and bees (hymenoptera) from the lower tertiary claron formation of southwestern utah [abs.]: geological society of america abstracts with programs, v. 27, no. 4, p. 3. bruthans, j., soukup, j., vaculikova, j., filippi, m., schweigstillova, j., mayo, a.l., masin, d., kletetschka, g., and rihosek, j., 2014, sandstone landforms shaped by negative feedback between stress and erosion: nature geoscience, v. 7, p. 597–601, online, doi: 10.1038/ngeo2209. davis, g.h., and pollock, g.l., 2010, geology of bryce canyon national park, in sprinkel, d.a., chidesy, t.c., jr., and anderson, p.b., editors, geology of utah’s parks and monuments, third edition: utah geological association and bryce canyon natural history association, utah geological association publication 28, p. 37–60. dickinson, w.r., 2006, geotectonic evolution of the great basin: geosphere, v. 2, p. 353–368. golder, k.b., and wizevich, m.c., 2009, wetland trace fossils in the paleogene brian head formation, southwest utah [abs.]: geological society of america abstracts with programs, v. 41, no. 3, p. 88. golder, k.b., wizevich, m.c., simpson, e.l., and storm, l.p., 2009, oligocene ichnofossils in non-marine limestone of the brian head formation, utah [abs.]: geological society of america abstracts with programs, v. 41, no. 7, p. 262. goldstrand, p.m., 1990, stratigraphy and paleogeography of late cretaceous and paleogene rocks of southwest utah: utah geological survey miscellaneous publication 90-2, 58 p. goldstrand, p.m., 1991, tectonostratigraphy, petrology, and paleogeography of the upper cretaceous to eocene rocks of southwest utah: reno, university of nevada, ph.d. dissertation, 205 p. goldstrand, p.m., 1992, evolution of late cretaceous and early tertiary basins of southwest utah based on clastic petrology: journal of sedimentary petrology, v. 62, no. 3, p. 495–507. goldstrand, p.m., 1994, tectonic development of upper cretaceous to eocene strata of southwestern utah: geological society of america bulletin, v. 106, p. 145–154. hacker, d.b., biek, r.f., and rowley, p.d., 2014, catastrophic emplacement of the gigantic markagunt gravity slide, southwest utah (usa)—implications for hazards associated with sector collapse of volcanic fields: geology, v. 42, no. 11, p. 943–946. hasiotis, s.t., and bown, t.m., 1997, crayfish (decapoda, cambaridae) burrows and their paleohydrologic and paleoenvironmental significance, paleocene-eocene claron formation, markagunt plateau, southwestern utah [abs.]: geological society of america abstracts with programs, v. 29, no. 2, p. 13. hatfield, s.c., rowley, p.d., sable, e.d., maxwell, d.j., cox, b.v., mckell, m.d., and kiel, d.e., 2010, geology of cedar breaks national monument, utah, in sprinkel, d.a., chidsey, t.c., jr., and anderson, p.b., editors, geology of utah’s parks and monuments: utah geological association and bryce canyon natural history association, utah geological association publication 28, third edition, p. 145–160. humphreys, e., 2009, relation of flat subduction to magmatism and deformation in the western united states, in kay, s.m., ramos, v.a., and dickinson, w.r., editors, backbone of the americas—shallow subduction, plateau uplift, and ridge and terrane collision: geological society of america memoir 204, p. 85–98. hurlow, h.a., 2002, the geology of cedar valley, iron county, utah, and its relation to ground-water conditions: utah geological survey special study 103, 74 p., 2 plates. lund, w.r., hozik, m.j., and hatfield, s.c., 2007, paleoseismic investigation and long-term slip history of the hurricane fault in southwestern utah, paleoseismology of utah, v. 14: utah geological survey special study 119, cd, 81 p. maldonado, f., 1995, decoupling of mid-tertiary rocks, red hills-western markagunt plateau, southwestern utah, in scott, r.b., and swadley, wc, editors, geologic studies in the basin and range–colorado plateau transition in southeastern nevada, southwestern utah, and northwestern arizona, 1992: u.s. geological survey bulletin 2056, p. 233–254. maldonado, f., sable, e.g., and nealey, d.l., 1997, cenozoic low-angle faults, thrust faults, and anastomosing high-angle faults, western markagunt plateau, southwestern utah, in maldonado, f., and nealey, l.d., editors, geologic studies in the basin and range–colorado plateau transition zone in southeastern nevada, southwestern utah, and northwestern arizona, 1995: u.s. geological survey bulletin 2153, p. 151–176. r.f. biek and p.d. rowley cedar breaks national monument north view overlook 11 moore, d.w., nealey, l.d., rowley, p.d., hatfield, s.c., maxwell, d.j., and mitchell, e., 2004, geologic map of the navajo lake quadrangle, kane and iron counties, utah: utah geological survey map 199, 2 plates, scale 1:24,000. mullett, d.j., 1989, interpreting the early tertiary claron formation of southern utah [abs.]: geological society of america abstracts with programs, v. 21, no. 5, p. 120. mullett, d.j., wells, n.a., and anderson, j.j., 1988a, early cenozoic deposition in the cedar-bryce depocenter—certainties, uncertainties, and comparisons with other flagstaff-green river basins [abs.]: geological society of america abstracts with programs, v. 20, no. 3, p. 217. mullett, d.j., wells, n.a., and anderson, j.j., 1988b, unusually intense pedogenic modification of the paleocene-eocene claron formation of southwestern utah [abs.]: geological society of america abstracts with programs, v. 20, no. 5, p. 382. röhl, u., bralower, t.j., norris, r.d., and wefer, g., 2000, new chronology for the late paleocene thermal maximum and its environmental implications: geology, v. 28, no. 10, p. 927–930. rowley, p.d., biek, r.f., sable, e.g., boswell, j.t., vice, g.s., hatfield, s.c., maxwell, d.j., and anderson, j.j., 2013, geologic map of the brian head quadrangle, iron county, utah: utah geological survey map 263dm, 38 p., 2 plates, scale 1:24,000. sable, e.g., and maldonado, f., 1997, the brian head formation (revised) and selected tertiary sedimentary rock units, markagunt plateau and adjacent areas, southwestern utah, in maldonado, f., and nealey, l.d., editors, geologic studies in the basin and range-colorado plateau transition zone in southeastern nevada, southwestern utah, and northwestern arizona, 1995: u.s. geological survey bulletin 2153, p. 7–26. schinkel, t., 2012, investigation of the origin of silicified layers within paleogene-aged volcaniclastic brian head formation, southern utah: new britain, central connecticut university, m.s. thesis, 100 p. spangler, l.e., 2010, geology and hydrology of a vulcanokarstic terrain on the markagunt plateau, southwestern utah, in carney, s.m., tabet, d.e., and johnson, c.l., editors, geology and geologic resources of south-central utah: utah geological association guidebook 39, p. 93–108. weaver, l., 2010, cascade falls, kane county, utah: utah geological survey, survey notes, v. 42, no. 1, p. 12. wilson, m.t., and thomas, h.e., 1964, hydrology and hydrogeology of navajo lake, kane county, utah: u.s. geological survey professional paper 417-c, 26 p. uga-geosite-biek-silver-reef.indd utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors robert f. biek1 and j. chris rohrer2 1utah geological survey, salt lake city, ut, bobbiek@utah.gov 2utah division of oil, gas and mining, abandoned mine reclamation program, salt lake city, ut cover image: view northeast to the silver reef mining district. cottonwood creek is in the foreground, the snow-capped peaks in the distance are in the kolob canyons section of zion national park, and interstate 15 is on the right. . silver reef mining district 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e geosites refl ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: biek, r.f., 2019, silver reef mining district, in milligan, m., biek, r.f., inkenbrandt, p., and nielsen, p., editors, utah geosites: utah geological association publication 48, 14 p., https://doi.org/10.31711/ugap.v1i1.93. presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president r.f. biek silver reef mining district 3 introduction the silver reef mining district in southwestern utah is a geologic anomaly, a historical curiosity, and an ecological novelty. it is one of the few places in the world where economic disseminated silver chloride (chlorargyrite or horn silver) was produced from sandstone. the area is a little-known ghost town, now reborn as the upscale residential community of silver reef with deep ties to its history. the old wells fargo bank building, home of the silver reef museum, is listed on the national and utah state registers of historic buildings, and several other historic buildings and sites make this a fascinating area to visit (figure 1). finally, the mining district lies near the junction of the mohave and great basin ecological provinces and so contains an assemblage of plants and animals common to both regions; its mines are habitat for bats, including species considered imperiled in the state. proctor and shirts (1991) provided a fascinating account of the discovery, disbelief, re-discovery, and development of this unusual silver chloride deposit, and we previously summarized the geology, mining history, and reclamation of the district (biek and rohrer, 2006). the silver reef mining district straddles the interstate 15 corridor in central washington county, about 15 miles (24 km) northeast of st. george. the district consists of four “reefs” on the northeast-plunging nose of the virgin anticline: white, buckeye, and butte reefs are on the anticline’s northwest flank, whereas east reef is on the anticline’s east flank (a “reef ” is a mining term for a lode or vein, and was also used by early pioneers to describe figure 1. location of the silver reef (sr) and east reef (er) areas of the silver reef mining district. inset shows location of nearby pioche, nevada. rcra=red cliffs recreation area, qcsp=quail creek state park. prominent ridges that were a barrier to travel) (figure 2). the district is noted for its uncommon occurrence of ore-grade silver chloride in sandstone that is unaccompanied by obvious alteration or (with the exception of copper) substantial base-metal ores. the ore horizons are contained in the springdale sandstone member of the kayenta formation, which is repeated by thrust faults on the anticline’s northwest flank. high-grade silver chloride float was first discovered near harrisburg in 1866, and in situ mineralization was found in 1868, but it was not until 1876 that the silver rush was underway in earnest (proctor, 1953; proctor and brimhall, 1986). that it took 10 years between the discovery of silver in sandstone and the great pioche stampede, when miners in pioche, nevada, left for the new silver boom at silver reef, gives an idea of the difficulty early prospectors had in believing economic silver could be found in sandstone. chlorargyrite is at best an inconspicuous mineral and is normally associated with the weathered portions of silver-bearing sulfide deposits such as nevada’s famous comstock lode. the principal mining activity in the district lasted only through 1888, with lessee operations through 1909, after which major mining essentially ceased. prior to 1910, the district produced over 7 million ounces of silver from ore that averaged 20 to 50 ounces silver per ton, but varied from only a few ounces to about 500 ounces per ton (heikes, 1920; proctor and brimhall, 1986; eppinger and others, 1990). sporadic production between 1949 and 1968 yielded about 10 ounces of gold, 165,000 ounces of silver, 34 short tons of copper, and at least 2500 pounds of uranium oxide. in 1979, a leachpad operation was established between white and buckeye reefs to process tailings, but this venture soon closed with the collapse of silver prices. figure 2. view northeast to the silver reef mining district. cottonwood creek is in the foreground, the snow-capped peaks in the distance are in the kolob canyons section of zion national park, and interstate 15 is on the right. 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 initial reclamation at white, buckeye, and butte reefs, completed during 1996 to 1997 by the utah abandoned mines reclamation program (uamrp), involved 465 mine closures at a cost of $469,000. reclamation of 184 mine openings at the east reef area was completed in 2000 at a cost of $170,000. this initial reclamation of the silver reef mining district was restricted to immediate hazard abatement—the closure of adits and shafts for public safety. it was complicated by the fact that (1) the district was considered a “rural historic landscape” eligible for the national register of historic places; (2) it is home to one of the state’s largest corynorhinus townsendii (townsend’s big-eared bat) maternity colonies and has numerous roosts for other bat species; and (3) the adjacent town of silver reef was experiencing a second real estate and population boom. ironically, the historical status and other constraints at silver reef forced the uamrp to toss normal reclamation convention on its head. while reclamation efforts usually strive to make the mining disturbance disappear, at silver reef the goal was to preserve the mining disturbance while keeping the reclamation invisible. the uamrp used several mine closure techniques, with the majority of mines closed by backfilling. much of the backfilling was done by hand to minimize the effects of heavy equipment on the landscape. significant care was taken to preserve the appearance of nearby dumps, and in many cases, the fill was recessed slightly below grade to eliminate the fall hazard but maintain the appearance of the opening and collar features. because of both cultural and biological concerns, reclamation made extensive use of steel gates and grates. grates allow bats to come and go and they maintain ventilation in the mines. much of the historical interest and value of silver reef derives from its surviving mine openings, mine dumps, and structures, now carefully preserved by uamrp. other reclamation efforts include the u.s. environmental protection agency’s encapsulation of the 5m heap leach pad, and utah department of environmental quality’s remediation of tailings at the leeds and christy mills and fencing and other work at the big hill shaft. location the best place to learn about the history and geology of the silver reef mining district is at the silver reef museum, located in the old wells fargo building (37° 15.183' n., 113° 22.037' w.). check the museum’s informative web site (www.silverreefutah.org) for hours that they are open and for their calendar of events. an interpretative trail is open year round even when the museum is closed; the trail guide is available on their website. the museum itself is well worth the small entrance fee. from the north (cedar city/salt lake city) traveling south on i-15: take i-15 south to exit 23 for leeds/silver reef. turn right on silver reef road and travel west toward the red cliffs area, about 1.5 miles (2.5 km). ***follow directions below. from the south (st. george/las vegas) traveling north on i-15: take i-15 north to exit 22 for leeds/silver reef. drive north on main street through the town of leeds for about 1.3 miles (2 km). turn left on silver reef road at the sign for i-15 north. cross under i-15 and head west toward the red cliffs area, about 1.5 miles (2.5 km). ***follow directions below. ***you will see signs for the silver reef museum and the cosmopolitan restaurant. from silver reef road turn left at the “y” onto silver reef drive and continue for about 0.25 mile (0.4 km) to the museum on the right at the corner of silver reef drive and wells fargo drive. the street address is 1903 wells fargo rd., leeds. the mining district is also accessible from the south on hiking trails that are part of the red cliffs desert reserve administered by washington county in cooperation with several state and federal agencies. a trail map and other information is available on the reserve’s website www.redcliffsdesertreserve.com/trail. an entrance fee is required. see biek (2003b) for a geologic map of this area. for those with high-clearance vehicles, a trip to nearby east reef, also part of the red cliffs desert reserve, is well worth the effort. there, you will find some of the best exposures of late triassic to early jurassic strata in southwestern utah (including the springdale sandstone, the productive interval at silver reef), a relatively young basaltic lava flow and its inverted valley, and mine workings of the east reef part of the silver reef mining district (biek, 2003a). again, check out trail maps at the reserve’s website. to access the east reef area from leeds, proceed north on north main street (utah highway 228) for 0.75 mile (1.2 km) past silver reef road and the i-15 interchange, then turn right onto 900 north. follow this dirt road for 3.8 miles (6 km) and park at the historic babylon trailhead (37° 12' 07.3" n., 113° 20' 55.2" w.). this road may be impassible when wet and the part of the road that traverses the east reef lava flow is rough, thus a high-clearance vehicle is recommended. from the parking area, a short walk through the springdale water gap showcases a complete section of the lower kayenta, moenave, and upper chinle formations, as well as dinosaur tracks and petroglyphs (biek and hayden, 2007). finally, although several state and federal agencies have worked hard on reclamation of the area, this is an old mining district still with potential hazards. stay away from and do not enter old mine workings. collapse or caving of backfilled adits, and caving of shaft collars beyond their steel grates, is not uncommon. geology because of its unusual mineral occurrence, the silver reef mining district has captured the attention of numerous geologists, mining engineers, and historians. four “reefs,” or low ridges of resistant r.f. biek silver reef mining district 5 sandstone, comprise the silver reef mining district, with white, buckeye, and butte reefs on the northwest fl ank of the virgin anticline and east reef on the anticline’s eastern fl ank. because of its complicated structure on the folded and faulted northeast-plunging nose of the anticline, it was not until the early 1950s, through the careful work of paul proctor at brigham young university, that the structure and stratigraphy of the district were fi nally understood. several books and countless articles were written on confl icting interpretations of the structure, stratigraphy, and genesis of ore deposits of the silver reef mining district, seemingly out of proportion to the size of the district and the wealth it generated (proctor and shirts, 1991; biek and rohrer, 2006). structure and stratigraphy go hand-in-hand in interpreting the geology of an area, and nowhere is this truer than in the silver reef mining district. early geologists who studied the silver reef mining district concluded that the three western reefs were three separate stratigraphic layers and that their similarity was due to similar conditions of deposition (fi gure 3). to their credit, they noted the remarkable similarities among the three reefs, but at the time geologists did not understand the concept of thrust faulting, where older rock strata can be shoved on top of younger strata along gently dipping faults. we now know that only the section at east reef is undisturbed by subsidiary faulting or folding, and it in fact off ers one of the best sections of moenave and underlying chinle strata in all of southwestern utah. had all of this been known in the 1870s, the history of the silver reef mining district would be much less colorful and certainly less would have been written about the district. stratigraphy late triassic and early jurassic strata of the chinle, moenave, and kayenta formations form the heart of the silver reef mining district. figures 4, 5, and 6 show a simplifi ed lithologic column, geologic map, and cross sections of the silver reef area. figure 4. lithologic column showing stratigraphic units in the vicinity of silver reef, utah. modifi ed from biek and rohrer (2006). figure 3. cross sections (oriented northwest to southeast) showing historical development of fault theories for the silver reef area (modifi ed from proctor and shirts, 1991), beginning with the idea that white, buckeye and butte reefs were three similar but separate beds (a). louis janin, a mining engineer who helped organize the leeds mining company and establish the town of silver reef, was apparently the fi rst to conclude that the principal ore horizon was in fact a single bed cut by a fault between buckeye and white reefs. on march 31, 1880, the silver reef miner published a cross section (b) showing buckeye reef was truncated at relatively shallow depth by a down-tothe-east normal fault, which had the eff ect of limiting the extent of ore-bearing horizons at buckeye reef. at the time, henry lubbock, superintendent of the christy company, which operated the mines on buckeye reef, was in new york city trying to sell the christy mill and mining company; with the fault model in circulation, the value of his holdings was assumed to be diminished. he failed to fi nd a buyer and sued the paper for libel. th e district court in beaver rendered a verdict of not guilty, but the three-bed concept continued to be promoted by subsequent geologists. it was not until 1953 that paul proctor correctly showed that the ore horizons of white, buckeye, and butte reefs on the anticline’s northwest fl ank, and east reef on the anticline’s east fl ank, are all part of the same bed (then known as the silver reef sandstone member of the chinle formation) (c). 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 chinle formation the chinle formation of southwest utah has long been divided into the lower shinarump conglomerate member and the upper petrified forest member (stewart and others, 1972), but preliminary research on this interval suggests petrified forest strata may in fact contain beds better assigned to both younger and older chinle formation (martz and others, 2017). the shinarump conglomerate forms a prominent hogback around the virgin anticline, whereas the petrified forest member is both poorly and exceptionally well exposed in adjacent strike valleys. the chinle formation is late triassic in age, based principally on vertebrate and plant remains, and was deposited in a variety of fluvial and lacustrine environments of a low-relief, forested basin (stewart and others, 1972; dubiel, 1994). shinarump strata were deposited principally in braided-stream channels that flowed north and northwest. the streams were probably similar to the modern platte river that flows eastward from the rocky mountains and that consist of shallow, interconnected or braided channels and intervening gravel bars. the petrified forest fluvial systems mimicked this paleoflow, but with a much greater abundance of meandering stream deposits and floodplain mudstones (dubiel, 1994). amphibians, reptiles— including the crocodile-like phytosaur—freshwater clams, snails, ostracods, and fish made their home on this once vast, coastal lowland, and petrified conifer trees are common in chinle strata; fossil cycads, ferns, and horsetails are also present (stewart and others, 1972; blakey and others, 1993; dubiel, 1994; decourten, 1998). the shinarump conglomerate consists of cliff-forming, yellowish-brown, fineto very coarse-grained sandstone, pebbly sandstone, and minor pebbly conglomerate. small, subrounded pebbles are primarily quartz, quartzite, and chert. coarser sandstones figure 5. simplified geologic map of the silver reef mining district. see figure 4 for explanation of map unit symbols. note that jm here includes the springdale sandstone (now reassigned as the basal part of the kayenta formation). quaternary units include qa=undifferentiated alluvium, qafo=old alluvial-fan deposits, qes=eolian sand, and qb=undifferentiated basaltic lava flows. harrisburg junction, hurricane, pintura, and signal peak quadrangle boundaries also shown. cross sections a-a' and b-b' shown on figure 6. from biek and rohrer (2006). figure 6. simplified cross sections of the silver reef mining district. note that jm here includes the springdale sandstone (now reassigned as the basal part of the kayenta formation). see figure 4 for explanation of map unit symbols and figure 5 for cross section locations. from biek and rohrer (2006). r.f. biek silver reef mining district 7 and pebbly sandstones locally contain poorly preserved petrifi ed wood, commonly replaced in part by iron-manganese oxides. th e shinarump conglomerate varies widely in thickness, from 5 to 200 feet (2–60 m), due to deposition over an eroded landscape. some of the best and most complete exposures of petrifi ed forest strata in southwestern utah are at east reef, along the east side of the virgin anticline (fi gure 7). th e petrifi ed forest member consists of variably colored mudstone, claystone, siltstone, lesser sandstone and pebbly sandstone, and minor chert and nodular limestone. it contains a wider lithologic variation than might be expected given the prominent multicolored swelling mudstones that typify the member. th ese mudstones swell conspicuously when wet and give weathered surfaces a “popcorn” appearance— they are also responsible for numerous foundation problems and landslides in the region (lund and others, 2008). th e upper contact of the petrifi ed forest member is the j-0 unconformity, which represents a gap of about 10 million years during the late triassic and early jurassic (pipiringos and o’sullivan, 1978). moenave formation th e moenave formation is divided into the dinosaur canyon and whitmore point members (the former springdale sandstone member is now reassigned as the lower member of the kayenta formation). th ese distinctive strata are locally famous for the variety of fossils and dinosaur tracks discovered near st. george in february 2000 by sheldon johnson, a retired optometrist who was using a backhoe to prepare his land for development (kirkland and milner, 2006). sheldon donated the land to the city of st. george and a museum—the st. george dinosaur discovery site at johnson farm—opened in 2005. th e site contains thousands of dinosaur tracks in multiple track layers, including some of the best preserved tracks in the world, with skin and claw impressions, tail drag marks, squatting marks, and the fi rst unequivocal evidence of swimming dinosaurs anywhere in the world (milner and kirkland, 2006; milner and others, 2006). most tracks are the threetoed eubrontes, likely made by the crested, meat-eating dinosaur dilophosaurus, which reached 20 feet (6 m) long, 7 feet (2 m) high at the hips, and weighed about 1000 pounds (450 kg). many other types of tracks and huge, gar-like fi sh are also found at the site. th e dinosaur canyon member consists of slope-forming, generally thin-bedded, very fi ne grained sandstone, silty sandstone, and lesser siltstone and mudstone, all uniformly colored reddish brown. dinosaur canyon strata were deposited in river and fl oodplain environments (clemmensen and others, 1989; blakey, 1994; peterson, 1994; decourten, 1998); they span the triassic-jurassic boundary (suarez and others, 2017). if the dinosaur canyon member is noted for its uniformity— mostly reddish-brown, thin-bedded, very fi ne grained sandstone and silty sandstone—the whitmore point member must be known for its variety. it contains sandstone and siltstone similar to the dinosaur canyon, but also reddish-purple to greenish-gray mudstone and claystone and thin dolomitic limestone beds. th e limestones contain fossil fi sh scales and bones of semionotus kanabensis (schaeff er and dunkle, 1950). th e fi sh fossils were originally thought to be restricted to the triassic and so confl icted with fossil pollen from the whitmore point member that indicated the unit is early jurassic (peterson and others, 1977; imlay, 1980). we later learned that semionotus kanabensis is not age diagnostic, which fi nally resolved the long-standing debate on the age of the early jurassic moenave, kayenta, and navajo formations (olsen and padian, 1986). whitmore point strata were deposited in fl oodplain and lacustrine environments (clemmensen and others, 1989; blakey, 1994; peterson, 1994; decourten, 1998; kirkland and milner, 2006; milner and kirkland, 2006). kayenta formation th e contact between the whitmore point and springdale sandstone is a regional unconformity and corresponds to a pronounced break in slope, with the resistant springdale sandstone forming prominent cliff s and ledges above gentle whitmore point slopes (fi gure 8). th e springdale sandstone forms the imposing ridges of figure 7. view north of the petrifi ed forest strike valley at east reef, just north of where grapevine wash passes through the reef. trcs=shinarump conglomerate and trcp=petrifi ed forest members of the chinle formation; jtrmd=dinosaur canyon and jmw=whitmore point members of the moenave formation; jks=springdale sandstone member of the kayenta formation (formerly of the moenave formation); qb=east reef basaltic fl ow and cinder cone. th e white arksoic sandstone of proctor (1953) is also shown. th e southern end of black ridge near toquerville is visible in the distance. 8 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 white, butte, buckeye, and east reefs, and is host to the ore deposits of the silver reef mining district. in the district, the member was known as the silver reef sandstone, which was informally divided into the lower, white to brown leeds sandstone and the upper, lavender tecumseh sandstone (proctor, 1953). at silver reef, paul proctor’s careful geologic mapping and stratigraphic studies proved that the three reefs of the mining district were in fact one and the same bed, the springdale sandstone (since reassigned from the moenave formation to the overlying kayenta formation). th e springdale sandstone consists of thick-bedded, fi ne-grained sandstone with channel-form cross-stratifi cation. springdale sandstones are distinguished from overlying sandstones of the main body of the kayenta by their more variable pastel colors of pale red, pale pink, and yellowish orange, as opposed to reddish-brown hues that dominate kayenta beds (fi gure 9). where exposed along the fl anks of the virgin anticline, including in the silver reef mining district, the springdale sandstone is distinctly lighter in color than in exposures elsewhere, suggestive of bleaching by hydrocarbons once trapped in the core of the virgin anticline. th e springdale sandstone contains thin, discontinuous lenses of intraformational conglomerate, with mudstone rip-up clasts and poorly preserved, petrifi ed and carbonized fossil plant remains indicative of deposition in braided-stream and minor fl oodplain environments (clemmensen and others, 1989; blakey, 1994; peterson, 1994; decourten, 1998). structure th e silver reef mining district, in the transition zone between the colorado plateau and asin and range physiographic provinces, lies just 4 to 5 miles (6–8 km) west of the hurricane fault, the largest active earthquake fault in southwestern utah, capable of producing damaging earthquakes of about magnitude 7 (fi gure 1) (lund and others, 2007; see also the hurricane fault geosite). strata in the mining district are typical of the generally fl at-lying rocks of the colorado plateau. however, they were folded into the virgin anticline and subsidiary folds and cut by thrust faults during the sevier orogeny (see sidebar). in southwestern utah, the transition zone includes two major down-to-the-west fault zones that step down from the colorado plateau to the basin and range province. th e mining district lies on the intermediate structural block thus created, bounded on the east by the hurricane fault zone and on the west by the gunlock-grand wash fault (fi gure 10). figure 8. view west to white reef, just south of the old barbee mill site. trcp=petrifi ed forest member of the chinle formation; jtrmd=dinosaur canyon and jmw=whitmore point members of the moenave formation; and jks=springdale sandstone member of the kayenta formation. th e navajo sandstone forms the distant hills. figure 9. view north of east reef immediately north of where grapevine wash passes through the reef. southeast-dipping springdale sandstone (jks) and kayenta (jk) strata are beveled fl at and capped by old stream-terrace deposits (qat6). th e gradational springdale-kayenta contact corresponds to the base of the fi rst laterally continuous mudstone bed. east reef cinder cone forms the skyline on right. gunlockgrand wash fault hurricane fault washington fault virg in an tic lin e n silver reef east reef figure 10. schematic block diagram showing the relationship between the hurricane and gunlock-grand wash faults. both faults are “normal” faults that formed during regional extension, allowing rocks on the west side of the faults to slip down relative to rocks on the east side. r.f. biek silver reef mining district 9 virgin anticline the virgin anticline is a 30-mile-long (50 km), northeast-trending, symmetrical fold that trends parallel to, but is structurally distinct from, the kanarra anticline to the north (figure 1) (hurlow and biek, 2003). collectively, these two anticlines mark the eastward limit of significant sevier-age compressional deformation in southwestern utah. the flanks of the fold provide spectacular exposures of parts of the lower triassic moenkopi formation and overlying upper triassic chinle formation, including hogbacks of the shinarump conglomerate member that dramatically outline the fold’s shape (figure 12). along the northern reaches of the anticline, the carapace of shinarump stands in dramatic relief above the virgin river lowlands (figure 13). the anticline has three structural domes along its length, all located south of the mining district. from south to north these are bloomington dome, washington dome, and harrisburg dome, each of which exposes gypsum-bearing lower permian harrisburg member of the kaibab formation. the fold formed during the sevier orogeny, about 85 to 72 million years ago, above a blind thrust fault that soles into cambrian strata (davis, 1999; hurlow and biek, 2003). several west-dipping thrust faults are present along the west flank of the virgin anticline and figure heavily in the history of the mining district. proctor (1948, 1953) was the first to recognize the largest and westernmost thrust fault amid considerable controversy over structural interpretations of the district. proctor’s recognition of this and other thrust faults was the key to understanding that just one sandstone bed—the springdale sandstone—hosts the silver ore deposits of the silver reef mining district. this westernmost fault separates buckeye reef and white reef and extends at least 7 miles (11 km) to the southwest (biek, 2003b). the fault repeats the springdale sandstone in the reefs and farther south on the anticline’s western flank (figure 5). proctor and brimhall (1986) estimated that in the silver reef mining district the springdale sandstone was displaced eastward at least 2000 feet (600 m) on this fault. ore deposits ore-grade silver chloride (chlorargyrite or horn silver) accounted for more than 90% of the silver-bearing minerals recovered at the silver reef mining district (proctor, 1953). chlorargyrite (agcl) is a very soft, usually massive, inconspicuous silver mineral that is colorless or grayish white when underground, but when brought to the surface and exposed to sunlight, it develops like photographic film to a waxy grayish brown. because this and other silver minerals are nearly invisible, the ore is conspicuous only where colorful copper and uranium minerals are present (heyl, 1978). samples assaying 10 to 15 ounces (280–425 gm) per ton silver can lack visible silver minerals (james and newman, 1986). about the sevier orogeny, or mountain-building episode, began in utah by middle jurassic time, about 180 million years ago, as deformation associated with subduction of the farallon oceanic plate along the western edge of north america spread eastward into western utah (decelles, 2004). these eastward-directed compressional forces created the sevier orogenic belt and associated broad deformation zone, which consists of, from west to east, a thrust belt with wedge-top basins, foredeep basin, forebulge, and backbulge basin (decelles, 2004; willis, 1999, 2000) (figure 11). each of these four parts of the thrust system migrated eastward over time, and each created unique environments of deposition or erosion. figure 11. typical parts of a thrust system. the thickened, eastward-moving, leading-edge thrust wedge on the left overloads the earth’s crust, which flexes in response, similar to loading rock on a wooden raft floating on water. in utah, the entire thrust system migrated eastward over time during the middle mesozoic to early tertiary, but this simple pattern is commonly complicated due to variations in crustal strength and pre-existing faults. from willis (1999). figure 12. view northeast towards quail creek reservoir nestled in the eroded core of the virgin anticline, showing white (wr), buckeye (br), butte bur), and east (er) reefs on the northeast-plunging nose of the anticline. the flanks of the anticline are neatly outlined by the resistant shinarump conglomerate member of the chinle formation (trcs), below which are ledgy slopes of the upper red member of the moenkopi formation (trmu). the “bacon-striped” shnabkaib member of the moenkopi formation (trms) forms the eroded floor of the anticline. the snow-covered pine valley mountains are on the skyline at left and the kolob canyons part of zion national park is on the skyline at right; state route 9 cuts across the virgin anticline at the bottom of the photograph. photo courtesy of janice hayden. 10 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 70% of the district’s production came from the prolific buckeye reef (heikes, 1920). chlorargyrite was one of the important ore minerals at nevada’s famous comstock lode. its presence at silver reef is unusual in that chlorargyrite is usually associated with the weathered, near-surface portions of silver-bearing sulfide deposits, where silver was leached out of the rock by groundwater or hydrothermal fluids and redeposited in localized concentrations. malachite, a green copper carbonate, is the most common copper mineral in the district; it typically replaces plant material and is locally found as stains on the rock. azurite is locally present in combination with malachite. carnotite is the predominant uranium and vanadium mineral. like the chlorargyrite, uranium minerals are generally found in association with plant fragments and on bedding and fracture planes, but may occur disseminated within the sandstone. other minerals known to occur in the district include embolite (silver bromide), native silver, and argentite (silver sulfide). native silver and silver sulfide are only known from early reports of the district and apparently occur at depth within the mines, mostly near or below the water table (proctor, 1953). controls and origin of mineralization the silver reef mining district is part of a broad area relatively enriched in silver mineralization (james and newman, 1986). anomalously high silver concentrations occur in the springdale sandstone throughout southwestern utah, but were only economic at silver reef where concentrated around the nose of the anticline. the genesis of the silver reef ore deposits has been the subject of considerable debate since their discovery, but the consensus is that the metal deposits are of sedimentary origin later concentrated by reducing fluids in the northeast-plunging nose of the anticline (proctor, 1953; james and newman, 1986; houser and others, 1988). the silver and other metals were likely derived from leaching of volcanic ash beds in the chinle formation and redeposited in the springdale sandstone, the first overlying permeable bed with organic material. the silver was probably transported as sulfide-poor, chlorine-rich brine that passed upward into anticlinal traps where it encountered reducing conditions or low-salinity groundwater that caused silver to precipitate. a low-sulfide brine accounts for the lack of lead and zinc, which are soluble in such a solution. james and newman (1986) also noted evidence for petroleum residue in the springdale sandstone in the silver reef mining district, and both shinarump and springdale strata of the virgin anticline are more strongly bleached than exposures elsewhere. the virgin anticline may thus be a flushed-oil structure. subsequent enrichment created high-grade areas and petroleum escaped from the system as erosion breached the virgin anticline. ore processing silver ore of the mining district was processed principally in local pan-amalgamation mills. only the higher-grade ore, generally 20 ounces (550 gm) silver per ton or more, went to the stamp mills, where it was pulverized. the sandstone was easily crushed, freeing silver minerals in pore spaces and in associated carbonaceous debris. the pulverized ore was mixed with water in circular containers called mullers, to which was added “about a pound and a half of mercury per ton of crushed rock, 25 pounds of rock salt, and two pounds of crushed copper sulfate …” (proctor and shirts, 1991, p. 69). continuous mixing and heating of the slurry eventually brought the silver minerals in contact with mercury, salt, and copper sulfate, forming a frosty, gray, paste-like amalgam of silver and mercury. the amalgam was skimmed off the mullers and sent to retorts for silver separation and recycling of mercury. the mills typically recovered only about 80% of the assayed silver. mining history two of the best accounts of the silver mining history at silver reef are by stucki (1966) and proctor and shirts (1991). the summary below is extracted in large part from their research. there are several tales about the discovery of silver at silver reef, none more colorful than that of “metalliferous murphy.” murphy was a scottish assayer in the silver and lead mining town of pioche, nevada. some claimed that he could find rare metals in any sample sent to him. as a practical joke, some pioche miners broke up a grindstone and submitted it to him for assay. “in all soberness, figure 13. view north to the resistant shinarump conglomerate, which forms a whale-like carapace along the central part of the virgin anticline near east reef. the hurricane cliffs and kolob canyons part of zion national park are in the distance at right. r.f. biek silver reef mining district 11 after he had performed his assay work, and to the delight of all, he reported that it contained over $200 of silver per ton. some say that he was hanged on the spot. others claim he was tarred and feathered and run out of town on a rail. one happy ending of the story is that metalliferous murphy traced the origin of the grindstone to the leeds, utah area where he promptly went, located his claims, and became fabulously rich” (proctor and shirts, 1991, p. 26). proctor (1953), however, credits long-time mineral prospector john kemple with discovery of silver at silver reef. kemple spent the winter of 1866-67 at the old harrisburg site, eight years after mormon pioneers first wintered at the nearby confluence of quail creek and the virgin river. kemple doubtless recognized sandstone stained by green and blue copper carbonates; he assayed them and found them to contain silver. he sent samples to other assayers to confirm his finding and at least one refused to assay it, claiming kemple must be crazy to ask him to assay a sandstone rock. therein lies the problem—at the time, ore-grade silver was not known to occur in sandstone. even kemple doubted his own findings and moved on to the silver boomtown of white pine, nevada. still, kemple was drawn to southern utah and returned to old harrisburg the following year. after additional prospecting, he established the first mining claim in 1871. kemple’s original discovery shaft, now backfilled, is not far southwest of the orson adams home, recently restored and interpreted by the u.s. bureau of land management. in 1875, j.s. ferris and elijah thomas discovered high-grade silver ore at white reef, but it was william tecumseh barbee—agent of the walker brothers, well-known merchants and mining men of salt lake city—who set the silver rush in motion. on august 23, 1875, barbee recorded the barbee and walker claims on white reef and returned to salt lake city to report his findings. on the advice of their mining experts, the walker brothers refused further financing. undeterred, barbee and his co-workers returned and located 22 claims in october 1875; in november, they located the prolific tecumseh claim on buckeye reef, where barbee’s high-graded discovery samples assayed $500 silver per ton. barbee promoted the silver reef area in several articles in the salt lake tribune and pioche record and soon generated a fever among miners in the region, especially from the pioche area, which was playing out. the result was the pioche stampede, which began in october 1876, but interestingly the rush was not so much for mining claims (which were already staked out), but for business locations. barbee even platted out the real estate development of bonanza city, just south of silver reef, but that failed due to high property costs. most businesses and miners settled in silver reef, a real estate promotion of the san francisco-based leeds mining company. barbee eventually sold his interests at silver reef for $75,000 and continued prospecting in the west. in 1877, the leeds mining company built a mill on leeds creek (figure 14). soon thereafter, three other mining companies established mills and consolidated mines at silver reef. in 1881, rolker reported four companies at work in the silver reef mining district: the christy mining and milling company and the leeds mining company were under san francisco management, and the barbee & walker mining company and the stormont mining company were under new york control. in addition, the kinner mine at buckeye reef, and a few small claims at the south end of white reef were privately worked. when all four mills were running, the population of silver reef leveled off at about 1500, comparable to the mining towns of park city and bingham at the time. the principal mining activity at silver reef lasted only through 1888. the decline of the district is attributable to several factors. silver prices steadily declined—the average price for silver was $1.20 per ounce in 1877, $1.11 in 1883, and just $0.93 in 1888— and costs increased as mines got deeper. also important are reports that ore grade generally decreased with depth in the mines, and the host rock became more cemented and difficult to break. during the boom years, water was a significant problem only in mines of the buckeye reef, where it had to be pumped continually, but it played an important factor in closing of the mining district. stucki (1966) reported that a three-month-long strike beginning figure 14. oblique aerial view to the northeast of white, buckeye, and butte reefs. photo taken january 2, 1994, courtesy of utah division of oil, gas & mining. 12 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 on february 1, 1881, over wages that were to be cut from $4.00 to $3.50 per day, was the beginning of the end for silver reef. that may have been the first organized labor strike in utah, and many miners left for tombstone, arizona, and other western mining camps. those who remained eventually settled for the $3.50 daily wage. while the district’s heyday ended in 1888, individual lessors continued to work the district from 1889 to 1909. stucki (1966) and proctor and shirts (1991) also described the more recent history of the district, including sinking of a 540-footdeep (165 m) shaft at white reef in 1929 by american smelting and refining company (asarco); a u.s. atomic energy commission drilling program to evaluate uranium mineralization in the early 1950s; and, in 1979, the construction of a leach-pad operation between white and buckeye reefs to process tailings by hurricane-based 5m, inc. (which later collapsed with the hunt brothers scheme to control silver prices). mining interest in silver reef continued through the 1990s, but given the residential growth around silver reef, historical and ecological concerns, and the district’s small size and potential, silver reef ’s mining days are over. acknowledgments this geosite summary was extracted from geology, mining history, and reclamation of the silver reef mining district, washington county, utah by biek and rohrer (2006). as the lengthy acknowledgments of that paper attest, we benefited greatly from the research of many geologists and historians who studied the silver reef mining district. we thank grant willis, don clark, and mike hylland (ugs), and kimm harty (ugs retired), for helpful reviews. references biek, r.f., 2003a, geologic map of the hurricane quadrangle, washington county, utah: utah geological survey map 187, 61 p., 2 plates, scale 1:24,000. biek, r.f., 2003b, geologic map of the harrisburg junction quadrangle, washington county, utah: utah geological survey map 191, 42 p., 2 plates, scale 1:24,000. biek, r.f., and hayden, j.m., 2007, off the beaten path—spectacular geology of utah’s dixie, in lund, w.r., editor, field guide to geologic excursions in southern utah: utah geological association publication 35, 72 p. biek, r.f., and rohrer, j.c., 2006, geology, mining history, and reclamation of the silver reef mining district, washington county, utah, in bon, r.l., gloyn, r.w., and park, g.m., editors, mining districts of utah: utah geological association publication 32, p. 479–512. blakey, r.c., 1994, paleogeographic and tectonic controls on some lower and middle jurassic erg deposits, colorado plateau, in caputo, m.v., peterson, j.a., and franczyk, k.j., editors, mesozoic systems of the rocky mountain region, usa: denver, colorado, rocky mountain section of the society of economic paleontologists and mineralogists, p. 273–298. blakey, r.c., bashem, e.l., and cook, m.j., 1993, early and middle triassic paleogeography, colorado plateau and vicinity, in morales, m., editor, aspects of mesozoic geology and paleontology of the colorado plateau: museum of northern arizona bulletin 59, p. 13–26. clemmensen, l.b., olsen, henrik, and blakey, r.c., 1989, erg-margin deposits in the lower jurassic moenave formation and wingate sandstone, southern utah: geological society of america bulletin, v. 101, p. 759–773. davis, g.h., 1999, structural geology of the colorado plateau region of southwestern utah, with special emphasis on deformation bands: geological society of america special paper 342, 168 p. decelles, p.g., 2004, late jurassic to eocene evolution of the cordilleran thrust belt and foreland basin system, western u.s.a.: american journal of science, v. 304, p. 105–168. decourten, f., 1998, dinosaurs of utah: salt lake city, university of utah press, 300 p. dubiel, r.f., 1994, triassic deposystems, paleogeography, and paleoclimate of the western interior, in caputo, m.v., peterson, j.a., and franczyk, k.j., editors, mesozoic systems of the rocky mountain region, usa: denver, colorado, rocky mountain section of society of economic paleontologists and mineralogists, p. 133–168. eppinger, r.g., winkler, g.r., cookro, t.m., shubat, m.a., blank, h.r., crowley, j.k., kucks, r.p., and jones, j.l., 1990, preliminary assessment of the mineral resources of the cedar city 1° x 2° quadrangle, utah: u.s. geological survey open-file report 90-34, 142 p., scale 1:250,000. heikes, v.c., 1920, silver reef (harrisburg, leeds) district—history and production, in butler, b.s., loughlin, g.f., heikes, v.c., and others: u.s. geological survey professional paper 111, p. 585–586. heyl, a.v., 1978, silver reef, utah, ores and the possibilities of unoxidized ores at greater depths, in shawe, d.r., and rowley, p.d., editors, field excursion c-2, guidebook to mineral deposits of southwestern utah: utah geological association publication 7, p. 65. houser, b.b., jones, j.l., kilburn, j.e., blank, h.r., jr., wood, r.h., ii, and cook, k.l., 1988, mineral resources of the cottonwood canyon wilderness study area, washington county, utah: u.s. geological survey bulletin 1746-c, 14 p., scale 1:24,000. r.f. biek silver reef mining district 13 hurlow, h.a., and biek, r.f., 2003, geologic map of the pintura quadrangle, washington county, utah: utah geological survey map 196, 20 p., scale 1:24,000. imlay, r.w., 1980, jurassic paleobiogeography of the conterminous united states in its continental setting: u.s. geological survey professional paper 1062, 134 p. james, l.p., and newman, e.w., 1986, subsurface character of mineralization at silver reef, utah, and a possible model for ore genesis, in griffen, d.t., and phillips, w.r., editors, thrusting and extensional structures and mineralization in the beaver dam mountains, southwestern utah: utah geological association publication 15, p. 149–158. kirkland, j.i., and milner, a.r.c., 2006, the moenave formation at the st. george dinosaur discovery site at johnson farm, in harris, j.d., lucas, s.g., spielmann, j.a., lockley, m.g., milner, a.r.c., and kirkland, j.i., editors, tracking dinosaur origins—the triassic/jurassic terrestrial transition: new mexico museum of natural history and science bulletin 37, p. 289–309. lund, w.r., knudsen, t.r., vice, g.s., and shaw, l.m., 2008, geologic hazards and adverse construction conditions, st. george-hurricane metropolitan area, washington county, utah: utah geological survey special study 127, variously paginated, 14 plates, scale 1:24,000, dvd-rom. lund, w.r., hozik, m.j., and hatfield, s.c., 2007, paleoseismic investigation and long-term slip history of the hurricane fault in southwestern utah: utah geological survey special study 119, paleoseismology of utah, volume 14, 81 p., cd. martz, j.w., kirkland, j.i., milner, a.r.r., parker, w.g., and santucci, v.l., 2017, upper triassic lithostratigraphy, depositional systems, and vertebrate paleontology across southern utah—a field guide prepared for the society of vertebrate paleontology annual meeting, october 26–29, 2016, salt lake city, utah: utah geological association, geology of the intermountain west, v. 4, p. 99–180. milner, a.r.c., and kirkland, j.i., 2006, preliminary review of the early jurassic (hettangian) freshwater lake dixie fish fauna in the whitmore point member, moenave formation in southwest utah, in harris, j.d., lucas, s.g., spielmann, j.a., lockley, m.g., milner, a.r.c., and kirkland, j.i., editors, tracking dinosaur origins—the triassic/jurassic terrestrial transition: new mexico museum of natural history and science bulletin 37, p. 510–521. milner, a.r.c., lockley, m.g., and kirkland, j.i., 2006, a large collection of well-preserved theropod dinosaur swim tracks from the lower jurassic moenave formation, st. george, utah, in harris, j.d., lucas, s.g., spielmann, j.a., lockley, m.g., milner, a.r.c., and kirkland, j.i., editors, tracking dinosaur origins—the triassic/jurassic terrestrial transition: new mexico museum of natural history and science bulletin 37, p. 315–328. olsen, p.e., and padian, k., 1986, earliest records of batrachopus from the southwestern united states, and a revision of some early mesozoic crocodylomorph ichnogenera, in padian, k., editor, the beginning of the age of dinosaurs—faunal changes across the triassic-jurassic boundary: cambridge university press, p. 260–273. peterson, f., 1994, sand dunes, sabkhas, streams, and shallow seas—jurassic paleogeography in the southern part of the western interior basin, in caputo, m.v., peterson, j.a., and franczyk, k.j., editors, mesozoic systems of the rocky mountain region, usa: denver, colorado, rocky mountain section of the society for sedimentary geology, p. 233–272. peterson, f., cornet, b., and turner-peterson, c.e., 1977, new data on the stratigraphy and age of the glen canyon group (triassic and jurassic) in southern utah and northern arizona [abs.]: geological society of america abstracts with programs, v. 9, no. 6, p. 755. pipiringos, g.n., and o’sullivan, r.b., 1978, principal unconformities in triassic and jurassic rocks, western interior united states—a preliminary survey: u.s. geological survey professional paper 1035-a, 29 p. proctor, p.d., 1948, geologic map and sections of the silver reef mining area, in stugard, f., jr., 1951, uranium resources in the silver reef (harrisburg) district, washington county, utah: u.s. geological survey trace elements memorandum report tem-214, scales 1:31,360 and 1:4560. proctor, p.d., 1953, geology of the silver reef (harrisburg) mining district, washington county, utah: utah geological and mineral survey bulletin 44, 169 p. proctor, p.d., and brimhall, w.h., 1986, silver reef mining district, revisited, washington county, utah, in griffen, d.t., and phillips, w.r., editors, thrusting and extensional structures and mineralization in the beaver dam mountains, southwestern utah: utah geological association publication 15, p. 159–177. proctor, p.d., and shirts, m.a., 1991, silver, sinners and saints—a history of old silver reef, utah: provo, utah, paulmar, inc., 224 p. schaeffer, b., and dunkle, d.h., 1950, a semionotid fish from the chinle formation, with consideration of its relationships: american museum novitates, no. 1457, p. 1–29. stewart, j.h., poole, f.g., and wilson, r.f., 1972, stratigraphy and origin of the chinle formation and related upper triassic strata in the colorado plateau region: u.s. geological survey professional paper 690, 336 p. 14 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 stucki, a.b., 1966, a historical study of silver reef, southern utah mining town: provo, utah, brigham young university m.a. thesis, 128 p. suarez, c.a., knobbe, t.k., crowley, j.l., kirkland, j.i., and milner, a.r.c., 2017, a chronostratigraphic assessment of the moenave formation, usa using c-isotope chemostratigraphy and detrital zircon geochronology—implications for the terrestrial end triassic extinction: earth and planetary science letters, v. 475, p. 83-93. willis, g.c., 1999, the utah thrust system—an overview, in spangler, l.e., and allen, c.j., editors, geology of northern utah and vicinity: utah geological association publication 27, p. 1–9. willis, g.c., 2000, utah’s sevier thrust system: utah geological survey, survey notes, v. 32, no. 1, p. 1–4. uga-geosite-biek-hurricane-fault.indd utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors robert f. biek utah geological survey, p.o. box 146100, salt lake city, ut 84114-6100 bobbiek@utah.gov cover image: view north-northwest along hurricane fault zone just west of state route 9. hurricane fault 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e geosites refl ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: biek, r.f., 2019, hurricane fault, in milligan, m., biek, r.f., inkenbrandt, p., and nielsen, p., editors, utah geosites: utah geological association publication 48, 12 p., https://doi. org/10.31711/geosites.v1i1.51. presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president r.f. biek hurricane fault 3 introduction the hurricane fault is the big earthquake fault in southwestern utah. it stretches at least 155 miles (250 km) from south of the grand canyon northward to cedar city and is capable of producing damaging earthquakes of about magnitude 7.0. the hurricane fault is a “normal” fault, a type of fault that forms during extension of the earth’s crust, where one side of the fault moves down relative to the other side. in this case, the down-dropped side (the hanging wall) is west of the fault; the upthrown side (the footwall) lies to the east. like most long normal faults, the hurricane fault is composed of discrete segments that tend to rupture independently (figure 1). the fault lies at or near the base of the hurricane cliffs, which form an impressive, little-eroded fault scarp several hundred feet high. conspicuous, west-tilted, faulted slivers of mostly triassic and jurassic red beds are locally exposed at the base of the cliffs, and contrast strongly with gray permian carbonates exposed in the cliffs themselves. several pleistocene basaltic lava flows flowed across and are now offset by the fault zone, dramatically recording long-term slip rates. should you make the mistake of pronouncing the name “hurricane” as one would when describing a mighty storm on the east coast, you should stand to be corrected, for locals pronounce it as “hurricun” even though pioneers named the town after ferocious winds common to the local area. locations the hurricane fault is one of the most prominent geologic structures in southwestern utah. the eastern, upthrown side of the fault forms the hurricane cliffs, offering dozens of instructive places to view the fault. three readily accessible sites—the hurricane and la verkin overlooks and a polished, striated exposure of the hurricane fault—reveal the variety of deformation seen along the fault (figures 2, 3, 4, and 5). the la verkin overlook (37° 11' 55" n., 113° 15' 44" w.) is accessible by car on an unpaved 1.5 mile (2.5 km) dirt road that leads southwest off state route 9. the hurricane overlook (37° 11' 00" n., 113° 16' 44" w.) is at a trailhead off state route 59. the fault exposure showing slickenlines and a polished fault surface is partway up the hurricane cliffs on a relay ramp between strands of the hurricane fault. park at the wide pullout on the south side of route 9, at 37° 13' 15" n., 113° 15' 31." w. carefully cross the highway and hike 0.3 mile (0.5 km) north of the parking area to the fault exposure immediately west of and below route 9 (at 37° 13' 30" n., 113° 15' 31.7" w.). this same fault surface is exposed in a road cut on route 9, but is located at a dangerous curve with no shoulder; that site is best viewed driving slowly past as you head east. as a bonus, a small wash 450 feet (150 m) northeast of the parking area provides good exposures of the permian-triassic unconformity. hurricane fault the hurricane fault marks the western boundary of the colorado plateau in southwestern utah, forming the eastern boundary of a structural and stratigraphic transition zone with the nearby basin and range province. displacement increases northward along the hurricane fault, from 800 to 1300 feet (250–400 m) at the colofigure 1. the six segments of the hurricane fault zone, each of which has a different rupture history and rate of long-term slip; arrows indicate segment boundaries. the three geosites described here lie at the northern end of the anderson junction segment. the fault continues north of cedar city but that section was not part of the study of lund and others (2007) from which this figure is taken. 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 figure 3. geologic map of the hurricane-la verkin area. th e northeast-plunging nose of the virgin anticline is at upper left (see quail creek reservoir geosite for more information on the anticline). th e virgin river crosses the hurricane fault between hurricane and la verkin; the north wall of the canyon provides good exposures of the fault zone just north of the hurricane overlook. state route 9 takes advantage of a break in the hurricane cliff s created by a relay ramp between parallel strands of the hurricane fault zone; the relay ramp exposes a complete, though faulted, section of the moenkopi formation. rock unit names are shown in fi gure 4; various surfi cial deposits are shown in yellow and light brown. cross sections a-a' and b-b' are shown on fi gure 5. modifi ed from biek and others (2009). figure 2. locations of the hurricane and la verkin overlooks and state route 9 geosites. park here v ir g in river hurricane la verkin la verkin overlook timpoweap canyonhurricane overlook sr 9 site 0.5 mile 1 km r.f. biek hurricane fault 5 figure 4. stratigraphic column of rock units exposed in the greater hurricane area. from biek (2003). jk jms jmw jmd trcp trcs jt r u tr u tr m t r m u trms tr mtrmu trmv trml tr m trmt trmr p kt pkh pkf ptw ptb q qb kis jcco jn pts jts k ju 0-160+ (0-50+) 0-170+ (0-50+) 100-160 (30-50) lithologic column bentonitic beds pentacrinus sp. gypsum high-angle cross beds transition zone silver reef mining district semionotus kanabensis swelling, brightly colored clays "picture stone" "purgatory sandstone" gypsum "bacon striped" 3 limestone ledges oil seeps chert-clast conglomerate medial limestone brachiopods "black banded" collapse structures gypsum system and series formation member sy m b o l th ic k n es s fe et (m et er s) lithology quaternary c r et a c eo u s u pp er ju r a ss ic kayenta formation m id d le lo w er u pp er m . lo w er lo w er tr ia ss ic pe r m ia n navajo sandstone carmel formation iron springs formation surficial deposits basalt flows and associated deposits co-op creek limestone member temple cap formation sinawava member moenave formation springdale ss mbr whitmore point mbr dinosaur canyon member petrified forest memberchinle formation shinarump conglomerate mbr upper red member shnabkaib member middle red member moenkopi formation virgin limestone mbr lower red member timpoweap member rock canyon cgl harrisburg member fossil mountain member woods ranch member brady canyon member kaibab formation toroweap formation seligman member 600+ (180+) 130-200 (40-60) 220 (67) 2,0002,300 (600-700) 935 (285) 95 (29) 61 (18) 200 (60) 408-500 (124-150) 115-162 (35-49)) 400 (120) 400-600 (120-180) 400-500 (120-150) 100 (30) 200-250 (60-75) 30-130 (9-40) 208-286 (63-87) 130-200 (40-60) 160-230 (49-70) 30-50 (9-15) 0-80+ (0-24+) rado river in the grand canyon (karlstrom and others, 2007) to about 7500 feet (2300 m) along the north part of the fault near cedar city (hurlow, 2002). anderson and christenson (1989) found displacements of about 3600 feet (1100 m) and 4900 feet (1500 m) at the latitudes of st. george and toquerville, respectively. this is the largest active normal fault in southwestern utah. like most long normal faults, the hurricane fault is composed of discrete segments that tend to rupture independently. the fault zone comprises six segments, three of which—the cedar city, ash creek, and anderson junction—are in utah (figure 1). the anderson junction and ash creek segments are linked by a structurally complicated segment boundary near toquerville (stewart and taylor, 1996; biek, 2003; hurlow and biek, 2003; lund and others, 2007). this segment boundary is at the southern end of the kanarra anticline, a late cretaceous frontal fold of the sevier orogenic belt, and at the northern end of a relay ramp, which exposes a much-faulted yet complete panel of northwest-dipping moenkopi formation (figure 6). in the hurricane area, the fault juxtaposes upper cretaceous sandstone and mudstone of the iron springs formation on the hanging-wall block against lower permian carbonates in the footwall block (biek, 2003). triassic red beds, tilted moderately to the west, are caught up in splays of the fault at the entrance to timpoweap canyon (figure 7). in 2007, utah geological survey geologist bill lund and his colleagues showed that the most recent surface-faulting event on the fault probably occurred in the holocene (<10,000 years ago), at the north end of the fault near cedar city (lund and others, 2007). they further noted that multiple surface faulting earthquakes have occurred in the late quaternary (about the past 125,000 years) along most, if not all, of the utah portion of the fault. based on dated basaltic lava flows that crossed and are now offset by the fault, they calculated an average slip rate of about 8 inches per 1000 years (0.21 mm/yr) for the anderson junction segment of the hurricane fault since about 350,000 years ago. near ash creek reservoir, on the ash creek segment of the hurricane fault, they documented an average slip rate of about 22 inches per 1000 years (0.57 mm/yr) since about 850,000 years ago for that segment of the fault. the fault is considered capable of generating damaging earthquakes of about magnitude 7.0 (lund and others, 2007); the 1992 magnitude 5.8 st. george earthquake likely occurred on the hurricane fault (pechmann and others, 1995). faulting on the southern segments of the hurricane fault began in the pliocene, much later than for northern segments (rowley and others, 1978; hurlow, 2002). billingsley and workman (2000) described offset relationships of late tertiary and quaternary basaltic lava flows in arizona and showed that, based on equal offset 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 figure 5. cross sections through the hurricane cliff s between the la verkin and hurricane overlooks. see fi gure 3 for section locations and fi gure 4 for rock unit names. note that map (1:62,500) and cross sections (1:24,000) show diff erent levels of detail and so do not match exactly. from biek (2003). -2,000 -1,000 0 1,000 2,000 3,000 4,000 feet feet -2,000 -1,000 0 1,000 2,000 3,000 4,000 trmv kis jcco trcp jm jk jn jm trcp trmu trcs trms trmm trml jn jn jk jk trcp trmu jts jcco jts trml trmr + trmt trmr + trmt trml pk pt pk pt pk pt pk pk pt pt pq pq pq pq pq pt pk trml pzu pzu pzu qat5 + qat6 qbp qafo pzu jm trcs trmr + trmt a west east a ' ash creek la verkin creek h u r r i c a n e f a u l t z o n e complex fault zone -2,000 -1,000 0 1,000 2,000 3,000 4,000 feet jm rt m trcs jk trcp jn jm trcp(?) jk jts jcco kis jn jn(?) jk(?) jm(?) trcs(?) trm(?) jts jcco kis qb + qat6 pk ? ? ? ? ? pzu pq pk pttrm southeast b' feet -1,000 0 1,000 2,000 3,000 4,000 -2,000 b northwest hurricane fault zone virgin river faults projected and inferred ash creek la verkin creek complex fault zones thin surficial deposits not shown. thin surficial deposits not shown. -2,000 -1,000 0 1,000 2,000 3,000 4,000 feet feet -2,000 -1,000 0 1,000 2,000 3,000 4,000 trmv kis jcco trcp jm jk jn jm trcp trmu trcs trms trmm trml jn jn jk jk trcp trmu jts jcco jts trml trmr + trmt trmr + trmt trml pk pt pk pt pk pt pk pk pt pt pq pq pq pq pq pt pk trml pzu pzu pzu qat5 + qat6 qbp qafo pzu jm trcs trmr + trmt a west east a ' ash creek la verkin creek h u r r i c a n e f a u l t z o n e complex fault zone -2,000 -1,000 0 1,000 2,000 3,000 4,000 feet jm rt m trcs jk trcp jn jm trcp(?) jk jts jcco kis jn jn(?) jk(?) jm(?) trcs(?) trm(?) jts jcco kis qb + qat6 pk ? ? ? ? ? pzu pq pk pttrm southeast b' feet -1,000 0 1,000 2,000 3,000 4,000 -2,000 b northwest hurricane fault zone virgin river faults projected and inferred ash creek la verkin creek complex fault zones thin surficial deposits not shown. thin surficial deposits not shown. relay ramp hanging wall footwall n hurricane overlook la verkin overlook hurricane fault normal fault tip dip figure 6. a. oblique aerial photograph (view east, courtesy of janice hayden) of the relay ramp between two strands of the hurricane fault. state route 9 loops in and out of the photo on the right, and a gravel pit in old alluvial-fan deposits (qafo) is in the foreground. hurricane mesa, with its fl at-lying moenkopi strata capped by the shinarump conglomerate member of the chinle formation (trcs), is in the distance. members of the moenkopi formation are: trmu, upper red; trms, shnabkaib; trmm, middle red; trmv, virgin limestone; trml, lower red; and trmt, timpoweap. b. inset block diagram shows the main features of a relay ramp. r.f. biek hurricane fault 7 of lava flows and underlying mesozoic strata, most normal faults on the shivwits and uinkaret plateaus became active after 3.6 and possibly after 2.6 million years ago. geosites la verkin overlook the la verkin overlook offers expansive views in all directions— west across the st. george basin, north to black ridge, east towards zion national park, and south into arizona (figure 8). the overlook is on brownish-gray limestone, cherty limestone, fine-grained sandstone, siltstone, and mudstone of the timpoweap member of the moenkopi formation. these lower triassic shallow-marine strata form a gently undulating surface on top of the permian-triassic unconformity (see state route 9 geosite below) (nielson and johnson, 1979; dubiel, 1994; lucas and others, 2007); here, the intervening basal rock canyon member, also described under the state route 9 geosite, is thin or absent. the lower triassic moenkopi formation of southwestern utah, with its alternating reddish-brown, white, and gray layers, documents shallow-marine sedimentation along the western margin of the supercontinent pangea, when what is now southwest utah lay just north of the equator. the moenkopi consists of three transgressive members (the timpoweap, virgin limestone, and shnabkaib members that each record an interval of sea-level rise), each of which is overlain by an informally named regressive red bed member (the lower, middle, and upper red members, respectively, which record sea-level fall) (figure 9); the rock canyon conglomerate member locally forms the base of the moenkopi formation (reeside and bassler, 1921; stewart and others, 1972; dubiel, 1994). these members record a series of incursions and retreats of a shallow ocean across a gently sloping continental shelf, where sea-level changes of several feet translated into shoreline changes of many miles (blakey and others, 1993; dubiel, 1994). timpoweap strata unconformably overlie limestone, cherty limestone, and gypsum of the lower permian harrisburg member of the kaibab formation. these older, lower permian, strata were deposited when the flat landscape of the pangean coast was alternately inundated by warm, shallow seas and exposed as broad coastal beaches and sabkhas (broad, flat surfaces near sea level where high evaporation rates commonly led to the accumulation of salts in sediments, such as today in some parts of the arabian peninsula) (mckee, 1938; rawson and turner-peterson, 1979; nielson, 1986). this is also a good vantage point from which to contemplate longterm erosion rates for the st. george-zion national park area. such rates can be calculated using known ages of basaltic lava flows and their height above major drainages. what we find is that long-term incision rates vary systematically west-to-east across the region (hamblin and others, 1981; willis and biek, 2001). at st. george, west of the washington fault, rates average 0.2 feet (0.06 m) per thousand years; between the washington fault and hurricane fault, rates average 0.4 feet (0.11 m) per thousand years; and east of the hurricane fault, rates average 1.25 feet (0.35 m) per thousand years. using these long-term incision rates, we can show, for example, that most of zion canyon was carved within figure 7. view northeast to the entrance of timpoweap canyon, where the virgin river exits the hurricane cliffs. three main strands of the hurricane fault zone are shown. the fault zone is characterized by a west-dipping panel of triassic strata caught between splays of the fault; these red beds act as a barrier to groundwater flow, causing warm water to rise through permeable permian strata in the footwall and emerge as pah tempe hot springs. here, the 350,000-year-old volcano mountain lava flow (qbv2) is offset about 240 feet (73 m), yielding an average slip rate of about 8 inches per 1000 years (0.21 mm/yr). trcp, petrified forest member of the chinle formation; trms, shnabkaib member of the moenkopi formation; ptw and ptb, woods ranch and brady canyon members of the toroweap formation, respectively; pkf, fossil mountain member of the kaibab formation. 8 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 the past 2 million years (biek and others, 2010). variation in the amount and rate of long-term incision along the virgin river is a function of relative base-level lowering along faults at and near the boundary of the basin and range. the rates are fastest east of the hurricane fault, on its upthrown side, because it is the most active fault in the region. hurricane overlook beginning in miocene time in southwest utah and continuing nearly to the present day, dozens of widely scattered basaltic lava flows poured out onto this landscape. today, they constitute part of the western grand canyon basaltic field, which extends across the southwest part of the colorado plateau and adjacent transition zone in southwest utah, northern arizona, and easternmost nevada (hamblin, 1963, 1970; best and brimhall, 1970, 1974; best and others, 1980; smith and others, 1999). the hurricane overlook is on one of these lava flows (the volcano mountain lava flow), which erupted about 350,000 years ago from a vent at volcano mountain (also known as sullivan knoll) just southwest of hurricane. this lava flow is displaced about 240 feet (73 m) by the hurricane fault, which is at the base of the cliffs. the lava flow is eroded to form the deep gorge of the virgin river north of hurricane. the cliffs reveal a single lava flow about 170 feet (50) thick; at its base is a 20to 40-foot-thick (6-12 m) rubbly zone with pillow basalts (figures 10 and 11). the pillow basalt, which only forms when hot, fluid lava flows into water, shows that the flow temporarily blocked the virgin river. figure 9. view north to hurricane mesa from state route 9 showing the lower red member (trml), virgin limestone member (trmv), middle red member (trmm), the “bacon-striped” shnabkaib member (trms), and upper red member (trmu) of the moenkopi formation. hurricane mesa, capped by the resistant shinarump conglomerate member of the chinle formation (trcs), has a world war ii-era test track once used by the air force for testing ejection seats. note landslide (area of disrupted bedding, qms) in the down-dropped shnabkaib and middle red strata at the right center of the photo. fi gu re 8 . p an or am ic vi ew fr om th e l a ve rk in o ve rlo ok . b eg in ni ng a t t he ri gh t ( no rt he as t o f t he o ve rlo ok ), a co m pl et e s ec tio n of th e m oe nk op i f or m at io n is ex po se d be lo w th e s hi na ru m p co ng lo m er at e, w hi ch ca ps h ur ric an e m es a. b la ck r id ge is ca rv ed fr om th e e as t l im b of th e k an ar ra a nt icl in e a nd is lo ca lly ca pp ed b y th e 8 50 ,0 00 -y ea rol d pi nt ur a la va fl ow . th is la va fl ow h ad it s p rin cip al so ur ce w es t o f t he h ur ric an e c liff s an d ha s s in ce b ee n di sp la ce d ov er 1 00 0 fe et (3 00 m ) b y do w nto -th ew es t m ov em en t o n th e h ur ric an e f au lt (w hi ch is m os tly o ut o f v ie w a t t he w es te rn b as e o f b la ck r id ge ). th e p in e v al ley m ou nt ai ns , a n ea rly m io ce ne la cc ol ith in tr ud ed in to th e c la ro n fo rm at io n, fo rm th e s ky lin e t o th e n or th w es t. to th e w es t a nd so ut hw es t, th e v irg in r iv er cu ts th ro ug h se ve ra l b as al tic la va fl ow s b et w ee n h ur ric an e a nd l a ve rk in . th e b as al t flo w ex po se d in th e f oo tw al l o f t he h ur ric an e f au lt zo ne , j us t s ou th o f t he v irg in r iv er , i s 3 50 ,0 00 y ea rs o ld a nd er up te d fro m v ol ca no k no ll; it h as b ee n di sp la ce d ab ou t 2 40 fe et (7 3 m ) b y th e h ur ric an e f au lt. m ol lie s n ip pl e, on th e s ky lin e a t l eft , i s c ap pe d by a b as al tic la va fl ow th at er up te d at iv an s k no ll, im m ed ia te ly so ut h of v ol ca no k no ll. th e i va ns k no ll flo w is a bo ut 1 m ill io n ye ar s o ld a nd h as b ee n di sp la ce d ab ou t 1 30 0 fe et (4 00 m ) b y th e h ur ric an e f au lt. r.f. biek hurricane fault 9 three strands of the hurricane fault are exposed at the entrance to timpoweap canyon, a short walk north of the hurricane overlook (figure 7). pah tempe hot springs (also known as la verkin or dixie hot springs) issues from the canyon bottom just east of the fault zone. the west-dipping panel of triassic red beds caught in the hurricane fault zone acts as a barrier to groundwater flow, whereas the fractured and cavernous permian carbonates east of the fault act as conduits to groundwater flow. according to research by brigham young university geologists, as groundwater encounters the hurricane fault, probably at depths of at least 2000 to 4000 feet (700-1300 m), it hits a moenkopi seal and is forced laterally and upward, eventually discharging along the virgin river (dutson, 2005). the source of the hot water is thus structurally controlled and results from deeply circulating groundwater and the earth’s natural geothermal gradient (here probably about 130°f/ mile depth); it is not related to eruption of relatively young basaltic lava flows. basaltic lavas in the region originated at far greater depth, rising to the surface though narrow conduits, thus precluding the basaltic magmas from being a significant heat source. pah tempe hot springs has the highest recorded spring-water temperature in the st. george basin (budding and sommer, 1986). the springs average about 104°f (40°c), have a ph of 6.3, and discharge sodium-chloride-type water with a very high average of 9600 to 9900 mg/l total dissolved solids (tds) concentration. they issue from and immediately above the bed of the virgin river for a distance of nearly 1500 feet (500 m) upstream of the hurricane fault. the average annual dissolved-solids discharge at the springs is about triple that of the virgin river just upstream of the springs (mundorff, 1970; yelken, 1996). the high tds concentration of the spring water contributes to poor downstream water quality, and is the reason virgin river water is collected above the springs and piped to quail creek and sand hollow reservoirs. in april 1985, sinkholes appeared in the bed of the virgin river— adjacent to an abandoned diversion structure and just downstream from the newly completed quail creek diversion structure—about 2 miles (3.2 km) straight-line distance upstream from pah tempe hot springs. the sinkholes captured an estimated 7200 acre-feet (about 9,000,000 m3) of water over a period of several months. this stream-capture event acted as a natural aquifer test that was evaluated by ben everitt of the utah division of water resources and his colleague martin einert. discharge at the hot springs surged from 11 to about 20 cubic feet per second (0.3-0.6 m3/s), the temperature dropped about 9°f (5°c), and the tds concentration dropped by about 2000 mg/l (everitt and einert, 1994). within a year all parameters began to recover and returned to normal after about 2.5 years. figure 11. a. close-up of pillow basalt near the state route 9 bridge; hammer for scale. b. note altered, glassy rind on pillow. a b figure 10. volcano mountain lava flow and state route 9 bridge over the virgin river. this cliff reveals a single lava flow about 170 feet (50 m) thick; at its base is a rubbly zone with pillow basalts, indicating that the flow once blocked the ancestral virgin river, creating a small lake into which lava flowed. this rubbly zone is overlain by about 10 feet (3 m) of dense, fine-grained basalt with prominent, widely spaced columnar joints (the lower colonnade), which in turn is overlain by about 100 feet (30 m) of similar basalt that is prominently and chaotically jointed (the entablature). the upper roughly 20 feet (6 m) of this exposure consists of vesicular basalt with few columnar joints (the upper colonnade). this same sequence is present in exposures on the upthrown side of the hurricane fault. 10 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 highway 9 road cut in this area, the hurricane fault zone forms a relay ramp between two en echelon faults (figure 6) (biek, 2003). the ramp exposes a complete section of moenkopi strata, cut by numerous down-to-theeast and down-to-the-west normal faults. this relay ramp also forms an easy way up the hurricane cliffs, a fact exploited by state route 9. beautiful slickenlines on a polished fault surface, described below, are found 0.1 mile (0.15 km) up the road from the suggested parking area, but are located at a dangerous curve of state route 9 (at 37° 13' 23.3", 113° 15' 29.6"). however, even more dramatic slickenlines are present to the north below the highway; to access this area, carefully cross the highway and walk about 1300 feet (400 m) north, to the north end of the highway fill (to 37° 13' 29.7", 113° 15' 34.0"), where the lower red member is faulted down-to-thewest against the rock canyon conglomerate (figure 12). the fault plane dips 85 degrees west and slickenlines developed on rock canyon strata are nearly vertical (figure 13). the permian-triassic boundary is exposed in the cliffs of a small box canyon just ahead of the parking area on the east side of the road (figure 14). in southwest utah, this boundary is a major unconformity that spans 10 to 20 million years (nielson, 1981, 1991; sorauf and billingsley, 1991). this is the tr-1 unconformity of pipiringos and o’sullivan (1978), the first regional unconformity of the triassic period in the western u.s. it also represents a period of dramatic, world wide sea-level drop and the largest global extinction event in the history of the world (see, for example, ward, 2004). erosion during this time produced an irregular surface, locally with several hundred feet of relief, upon which conglomerate and breccia of the rock canyon conglomerate member of the moenkopi formation was deposited in paleocanyons, karst depressions, and as regolith (nielson, 1991; higgins, 1997). limestone of the overlying timpoweap member of the moenkopi formation was deposited in broader paleovalleys (nielson and johnson, 1979). the cliffs of this small box canyon are capped by timpoweap strata that overlie channel-form conglomerates of the rock canyon conglomerate. cherty limestone of the harrisburg member of the kaibab formation (late early permian) is exposed at the base of the wash. acknowledgments my knowledge of southwestern utah geology comes from years of geologic mapping supported largely by the utah geological survey and u.s. geological survey. as the acknowledgments of the resultant geologic maps attest, i am indebted to a great many people for their help over the years. thanks to grant willis and mike hylland (ugs) and kimm harty (ugs retired) for their insightful reviews and to jenny erickson (ugs) for drafting the figures. figure 12. view north-northwest along hurricane fault zone just west of state route 9. the lower red member of the moenkopi formation (trml) is down-tothe-west against the rock canyon conglomerate member (trmr). slickenlines are shown by white line. black ridge, capped by the 850,000-year-old pinture lava flow, is on the skyline in middle of photograph. figure 13. slickenlines and polished fault surface developed on the rock canyon conglomerate member. here, the fault plane and slickenlines are nearly vertical. figure 14. here, a strand of the hurricane fault exposes the permian-triassic unconformity. the fault places lower red (trml) strata down on the west against the rock canyon conglomerate member (trmr) overlain by the timpoweap member (trmt). the early permian harrisburg member of the kaibab formation is exposed at the base of the wash, just out of view in this photo. view northeast. r.f. biek hurricane fault 11 references anderson, r.e., and christenson, g.e., 1989, quaternary faults, folds, and selected volcanic features in the cedar city 1○ x 2○ quadrangle, utah: utah geological and mineral survey miscellaneous publication 89-6, 29 p., scale 1:250,000. best, m.g., and brimhall, w.h., 1970, late cenozoic basalt types in the western grand canyon region, in hamblin, w.k., and 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geological society guidebook, 9th field conference, p. 87–99. reeside, j.b., jr., and bassler, h., 1921, stratigraphic sections in southwestern utah and northwestern arizona: u.s. geological survey professional paper 129-d, p. 53–77. rowley, p.d., and barker, d.s., 1978, geology of the iron springs mining district, utah, in shawe, d.r., and rowley, p.d., editors, guidebook to mineral deposits of southwestern utah: utah geological association publication 7, p. 49–58. smith, e.i., sanchez, a., walker, j.d., and wang, k., 1999, geochemistry of mafic magmas in the hurricane volcanic field, utah—implications for smalland large-scale chemical variability of the lithospheric mantle: the journal of geology, v. 107, p. 433–448. sorauf, j.e., and billingsley, g.h., 1991, members of the toroweap and kaibab formations, lower permian, northern arizona and southwestern utah: the mountain geologist, v. 28, no. 1, p. 9–24. stewart, j.h., poole, f.g., and wilson, r.f., 1972, stratigraphy and origin of the triassic moenkopi formation and related strata in the colorado plateau region, with a section on sedimentary petrology by r.a. cadigan: u.s. geological survey professional paper 691, 195 p. stewart, m.e., and taylor, w.j., 1996, structural analysis and fault segment boundary identification along the hurricane fault in southwestern utah: journal of structural geology, v. 18, p. 1017–1029. ward, p.d., 2004, gorgon—paleontology, obsession, and the greatest catastrophy in earth’s history: new york, viking, 257 p. willis, g.c., and biek, r.f., 2001, quaternary incision rates of the colorado river and major tributaries in the colorado plateau, utah, in young, r.a., and spamer, e.e., editors, colorado river origin and evolution—proceedings of the symposium held at grand canyon national park in june 2000: grand canyon association monograph 12, p. 119–123. yelken, m.a., 1996, trace element analysis of selected springs in the virgin river basin: las vegas, university of nevada, m.s. thesis, 156 p. uga-geosite-chidsey-breccia-pipe.indd 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors thomas c. chidsey, jr.1, david e. eby2, and douglas a. sprinkel1 (retired) 1utah geological survey, po box 146100, salt lake city, utah 84114-6100, tomchidsey@utah.gov 2 eby petrography & consulting, inc., 2830 w. 9 th ave., denver, colorado 80204 cover image: large breccia pipe penetrating the deseret limestone—the unique feature that defi nes the geosite. a breccia pipe in the deseret limestone, south flank of the uinta mountains, northern utah 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e geosites refl ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: chidsey, t.j., eby, d.e., and sprinkel, d.a., 2019, a breccia pipe in the deseret limestone, south fl ank of the uinta mountains, northern utah, in milligan, m., biek, r.f., inkenbrandt, p., and nielsen, p., editors, utah geosites: utah geological association publication 48, 10 p., https://doi.org/10.31711/ geosites.v1i1.55. presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president t.c. chidsey, jr., d.e. eby, and d.a. sprinkel breccia pipe in the deseret limestone 3 introduction a breccia pipe is a cylindricalor irregular-shaped mass of brecciated rock. a breccia consists of broken, angular fragments of rock cemented together by a fine-grained matrix. hydrothermal breccia pipes form when hydrothermal solutions force their way towards the surface through zones of weakness or fracture zones and naturally break up the rocks in the process, i.e., hydrofracturing (figure 1); breccia pipes can also form by collapse. hydrothermal breccia pipes can contain ore deposits and, as will be discussed later, are associated with some large oil and gas accumulations in southeastern utah. mississippian (359 to 318 million years ago [ma]) rocks outcrop along the flanks of the east-west-trending uinta mountains in northern utah (figure 2). uplift of this range occurred during the laramide orogeny, a regional mountain-building event, between the latest cretaceous (maastrichtian, about 70 ma) and the eocene (about 34 ma). the mississippian rock section along the south flank of the uinta mountains is over 1600 feet (490 m) thick (sprinkel, 2018) (figures 2 and 3). units include the fitchville (upper devonian-lower mississippian [385–340 ma]), gardison, deseret, humbug, and doughnut formations (figure 3). these units collectively were mapped as the madison limestone by bryant (1990). these formations generally have the same characteristics as the oiland gas-productive mississippian leadville limestone in fields of the paradox fold and fault belt of the northern paradox basin (southeastern utah and southwestern colorado (figure 4). the deseret limestone (osagean [346 ma] through middle meramecian [334 ma]) contains local zones of breccia due to either natural hydrofracturing or collapse. a hydrothermal breccia pipe in the deseret is well-exposed on the western end of the south flank of the uinta mountains and is the main focus of this geosite (figures 2 and 5). breccia associated with sediment-filled collapsed cavities is also present in nearby outcrops. brecciation caused by explosive natural hydrofracturing, described below, created similar shattered-looking, pulverized rock in well cores from the paradox basin. the best examples are identified from lisbon field (figure 4), the largest leadville oil field in utah (about 51.5 million barrels of oil produced through 2018 [utah division of oil, gas and mining, 2019]). breccia pipes and high-temperature dolomitization (the process by which the mineral dolomite is formed when magnesium ions replace calcium ions in the mineral calcite, 2caco3(calcite) + mg2+ ↔ camg(co3)2(dolomite) + ca2+) may be related to past hydrothermal activity (eby and others, 2005). figure 1. large breccia pipe penetrating the deseret limestone—the unique feature that defines the geosite. note pulverized nature of the material that comprises the pipe, the sharp contact with the country rock and parallel, calcite-filled vertical fractures. figure 2. generalized geologic map of the uinta mountains, northeastern utah, showing the location of the hydrothermal breccia pipe geosite in the mississippian section of the western part of the south flank (red square and arrow) and the diamond plateau breccia site (red circle) in the eastern part. modified from hintze and others (2000). 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 thus, the hydrothermal breccia pipe and nearby features were selected as a geosite because they serve as a visible example of an oil and gas reservoir model for the lisbon and other oil fields in the paradox basin (eby and others, 2009). combined with the details and characteristics of nearby outcrops, the breccia pipe can be used as a “template” for evaluating data from conventional oil well core, well logs (geophysical and petrophysical), and possibly subsurface imagery created with seismic surveys in the basin. how to get there the deseret limestone hydrothermal breccia pipe geosite is located along the south fork provo river on the western end of the south flank of the uinta mountains, wasatch county (figures 2 and 5). the breccia pipe geosite is about 50 miles (83 km) or a little less than an hour drive from salt lake city, utah, via interstate 80 and u.s. highway 40 to state highway 32. proceed eastbound on state highway 35 past the town of francis for 11 miles (18.4 km) to the a large roadcut on the east side which exposes the breccia pipe (40°33'03" n., 111°06'11" w., elevation 7400 feet; 2242 m) (figures 1 and 5); the geosite is about 25 miles (41.8 km) northwest of the town of hanna, utah, to the south on state highway 35. upon approaching the breccia pipe geosite, carefully slow down with flashers on and cross the westbound lane to park on the east side just below the roadcut, thus avoiding having to walk across the highway. to examine the outcrops up close requires scaling a fairly steep slope with loose rocks; caution is urged (figure 1)! general geologic characteristics of the mississippian section, uinta mountains the deseret limestone and other carbonate formations in the mississippian section around the uinta mountains were deposited in a shallow, warm-water, variable energy, epicontinental sea that extensively covered a large part of the craton (figure 6). the deseret is mostly lightto dark-gray, fineto coarse-crystalline, cherty limestone (figure 7). dolomitic units are gray to tan, sucrosic to crystalline, and medium bedded with occasional silty partings; both limestone and dolomite would be prime reservoir lithologies as found in the leadville limestone. chert is typically light gray, forming lenses and nodules. the most common carbonate fabrics of the mississippian rocks in northeastern utah include peloidal, skeletal, and oolitic grainstones, packstones, and wackestones; skeletal and intraclast rudstones and floatstones are also present. cross-bedded grainstones of crinoid debris are referred to as encrinites. mudstone appears as microcrystalline and cryptocrystalline limestone and dolomite. the mississippian section is generally thick to massive and unevenly bedded, forming vertical cliffs and dip slopes. marine fauna in the mississippian section are represented by corals, brachiopods, pelecypods, bryozoans, and crinoids; however, fossils are relatively rare in some areas. other common biota include ostracods, benthic forams, and gastropods. microbial-dominated rocks are present but uncommon. depositional environments include muddy tidal flats; burrowed peloidal muds in subtidal settings; high-energy oolitic shoals; storm-dominated, outer shelf open-marine, crinoid shoals with muddy intershoals; and offshore low-energy, open-marine settings below wave base. figure 3. lithologic column of a part of the paleozoic section along the western end of the south flank of the uinta mountains. modified from hintze and kowallis (2009), and sprinkel (2018). figure 4. location of fields (lisbon highlighted) that produce from the mississippian leadville limestone, arizona, utah, and colorado. thickness of the leadville is shown; contour interval is 100 feet (30.3 m) (modified from parker and roberts, 1963). the leadville limestone paradox basin play area is colored dark tan. modified from morgan (1993). t.c. chidsey, jr., d.e. eby, and d.a. sprinkel breccia pipe in the deseret limestone 5 figure 5. a – location map for the breccia pipe geosite (yellow dot), northern utah, and major highways, cities, and towns. b – google earth image (© 2018 google) showing the location of the hydrothermal breccia pipe geosite along the western end of the south flank of the uinta mountains in wasatch county, utah. the mississippian deseret limestone outcrops are best exposed in roadcuts within the elongated yellow oval; the red square is the location of the breccia pipe and red arrow is the location of possible paleokarst features. a b 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 the contacts between the fitchville, gardison, deseret, humbug, and doughnut formations are mostly conformable (sadlick, 1955, 1957; carey, 1973; hintze and kowallis, 2009), whereas the mississippian leadville limestone and overlying pennsylvanian molas formation are separated by a major unconformity in southeastern utah. this same unconformity is found at the top of the stratigraphically equivalent mississippian redwall limestone in the grand canyon (mckee, 1969), where subaerial exposure resulted in development of karst topography with carbonate breccia-filled collapse features (paleo-sinkholes) and terra rosa (cave fills) near the top of the formation. fractures are common in the mississippian section. they are best expressed as closely spaced, vertical fractures throughout thinto medium-thick beds or as swarms associated with large and small faults and collapse features. stylolites and jointing are also present. geology of the hydrothermal breccia pipe description the region that includes the deseret limestone hydrothermal breccia pipe geosite was mapped by eskelsen (1953), mcdougald (1953), and bryant (1990). sprinkel (2018) mapped the adjacent region along the south flank of the uinta mountains and western uinta basin. dockal (1980) published ten nearly complete measured sections of the mississippian units from the canyons around the core of the uinta mountains and they serve as an excellent reference set for further comparison. the deseret limestone at the geosite is typically darkto light-gray limestone consisting of skeletal grainstone to packstone (figure 7a). skeletal grains are composed of disarticulated crinoids and rugose coral fragments representing a high-energy, open-marine environment. some units contain in-place syringopora corals and common burrows (figure 7b), indicating a low-energy environment. other units are dolomitized and include chert nodules (figure 7c). vertical fractures are also common (figure 7d). the contact with the overlying humbug formation is difficult to recognize due to the poor nature of the outcrops and extensive slope cover. eskelsen (1953), mcdougald (1953), and bryant (1990) mapped them together in the area as deseret-humbug undifferentiated. the large breccia pipe is the most striking feature at the geosite (figure 1). the subvertical pipe is about 17 feet (~5 m) wide at the base of the outcrop and cuts vertically through about 30 feet (9 m) of deseret limestone. the interior of the pipe contains a poorly sorted breccia with small to large clasts surrounded by pulverized rock (figure 8a). calcite veins, dolomitized zones, and vugs are widespread (figure 8b). the contacts of the pipe with the unaltered limestone country rock are sharp. vertical, commonly calcite-filled fractures are prevalent on both sides of the breccia pipe. thin sections reveal the presence of mini-herkimer (i.e., doubly terminated) quartz crystals (figure 8c), which were also found in leadville cores from lisbon field. there, the mini-herkimer crystals had high-temperature fluid inclusions (joseph n. moore, energy & geoscience institute, written communication, 2009). the presence of mini-herkimer crystals suggests that a high-temperature event occurred at the geosite, presumably emplacing the breccia pipe. in deseret outcrops immediately southeast of the breccia pipe geosite, we recognized some of the same characteristics found in leadville cores and the redwall limestone outcrops in the grand canyon. stratiform brecciation (collapse) is extensive (figures 9a and 9b) but without the calcite veins, dolomitization, and explosive characteristic of a breccia pipe. red staining at the top of the deseret limestone, mapped as the madison limestone by bryant (1990), looks like possible terra rosa weathering (i.e., a reddish-colored clayey layer containing hematite that ranges in thickness from a few inches to several feet and can be found coating limestone in ancient and modern karst areas) but could be coming from the overlying humbug formation (figure 9c). early karsting and formation of collapse breccias may have begun shortly after deposition of the deseret, which would argue for an unconformity between the deseret and humbug; however, this has not been identified elsewhere in the region. several sinkholes are mapped in undivided deseret and humbug outcrops, as well as the gardison limestone and weber sandstone (pennsylvanian-permian), in the area by eskelsen (1953). these karst-related features likely postdate the mississippian beginning as early as late eocene to oligocene (godfrey, 1985; spangler, 2005). mayo and others (2010) also suggested the cave system formed relatively recently (pleistocene). figure 6. paleogeography of utah and eastern nevada during mississippian time. a warm shallow sea covered much of utah during the early meramecian (340 million years ago). western utah was the site of a deep starved basin. modified from blakey and ranney (2008). t.c. chidsey, jr., d.e. eby, and d.a. sprinkel breccia pipe in the deseret limestone 7 b d figure 7. typical characteristics of the deseret limestone just northwest of the breccia pipe (within the yellow oval on figure 5b). a – skeletal (crinoid and rugose coral) grainstone and packstone. b – in-place syringopora coral. c – chert nodules in dolomitized packstone. d – vertical fractures. figure 8. characteristics of the explosive nature of the hydrothermal breccia pipe. a – brecciated rock in shattered-looking, pulverized groundmass. b – close-up of sharp contact with unaltered limestone country rock. note vuggy dolomite and white calcite veins. c – photomicrograph (plane light) of dolomite containing a mini-herkimer quartz crystal (center) suggesting a high-temperature event. a c a b c 8 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 discussion all deseret limestone depositional environments described from outcrops above are also observed in leadville cores from lisbon field (figure 3) (chidsey, 2004; chidsey and others, 2004). carbonate buildups having good porosity/permeability represent the best reservoir analog units, whereas low-porosity/permeability, open-marine packstones and wackestones represent less attractive reservoir analog units, unless they have experienced dolomitization in the subsurface that results in increased reservoir quality. breccia pipes, paleokarst features, and fractures can also enhance reservoir quality. the post-burial breccias associated with hydrothermal events, fracturing, and dissolution in the leadville limestone yield the best reservoirs at lisbon field (eby and others, 2009). the breccia pipe geosite and extensive breccia zones discovered on the diamond plateau northeast of vernal, utah (figure 1) (eby and others, 2009) are likely the result of hydrothermal activity in the geologic past. the presence of the basal cambrian tintic quartzite in the southern flank of the western uinta mountains (figure 2) and the stratigraphically equivalent lodore formation in the eastern part of the range, may serve as hydrothermal recharge aquifers and thus are important contributors to the hydrothermal story. the tintic is a very coarse to granular, or pebble, sandstone with moderately sorted, subrounded to spherical, monocrystalline and polycrystalline quartz grains. it has thin to thick cross-bedding, is moderately indurated, and contains a few shaley partings, which have small amounts of mica and some biogenic feeding trails (dockal, 1980). the tintic’s contact with overlying mississippian strata is fairly sharp. the lodore is a very fine to medium-grained, well-sorted, very thinly bedded to cross-bedded (with somewhat undulatory surfaces) sandstone marked by argillaceous partings. quartz grains are subrounded to spherical. the lodore can be calcareous and slightly ferruginous. the top appears to be eroded (dockal, 1980). both the lodore and tintic can have porous and permeable units. as aquifers, they likely supplied hot water to the former hydrothermal system. three-dimensional numerical models of seafloor hydrothermal convection by coumou and others (2008) demonstrated that convection cells organize themselves into pipe-like upflow zones surrounded by narrow zones of warm downflow. recharge can occur over an extensive area or along faults as water migration pathways. the tintic quartzite is mapped on the western end of the uinta mountains whereas the lodore formation is present on the eastern end. through the central part of the south flank of the uinta mountains, porous cambrian sandstone is absent and the mississippian lies unconformably on precambrian (middle neoproterozoic) red pine shale and older formations of the uinta mountain group (figure 3). no hydrothermal breccia zones or pipes are found in the central part of the south flank, lending a b c figure 9. possible paleokarst features just southeast of the breccia pipe (see red arrow within the yellow oval on figure 5b). a – extensive collapse polymictic breccia. b – close-up of limestone and chert breccia clasts. note the lack of calcite veins and dolomite. c – red staining, possibly terra rosa weathering, at the top of the deseret limestone near the contact with the overlying humbug formation(?). t.c. chidsey, jr., d.e. eby, and d.a. sprinkel breccia pipe in the deseret limestone 9 credence to the concept that aquifers in the tintic and lodore were a required condition for past hydrothermal activity to have occurred. after pressure builds up to a certain point, the hydrothermal fluids within the aquifer can very rapidly and explosively break through zones of weakness at the intersections of fractures or along faults; this can occur as a single event or over several stages. thus, when targeting the leadville limestone in the paradox basin for potential hydrothermal dolomite and enhanced reservoir quality due to natural hydrofracturing, the presence of an underlying aquifer, and fracture zones or faults may be necessary ingredients, supported in part from what can be observed at the breccia pipe geosite. acknowledgments support for this paper was provided by the utah geological survey (ugs). cheryl gustin, jay hill, and lori steadman of the ugs drafted figures. this paper was carefully reviewed by michael d. vanden berg, stephanie m. carney, michael d. hylland, and bill keach of the ugs, along with the editors of this publication. their suggestions and constructive criticism greatly improved the manuscript. references blakey, r., and ranney, w., 2008, ancient landscapes of the colorado plateau: grand canyon, arizona, grand canyon association, 156 p. bryant, b., 1990, geologic map of the salt city 30' x 60' quadrangle, north-central utah and uinta county, wyoming: u.s. geological survey miscellaneous investigations map i-1944, 2 plates, scale 1:100,000. carey, m.a., 1973, chesterian-morrowan conodont biostratigraphy from northeastern utah: salt lake city, university of utah, m.s. thesis, 83 p. chidsey, t.c., jr., 2004, the mississippian leadville limestone exploration play, utah and colorado: rocky mountain association of geologists, the outcrop, v. 53, no. 10, p. 1 and 6. chidsey, t.c., jr., morgan, c.d., mcclure, k., and eby, d.e., 2004, the mississippian leadville limestone exploration play, utah and colorado [abs.]: american association of petroleum geologists, rocky mountain section meeting official program book, p. 94. coumou, d., driesner, t., and heinrich, a., 2008, the structure and dynamics of mid-ocean ridge hydrothermal system: science magazine, v. 321, p. 1825–1828. dockal, j.a., 1980, petrology and sedimentary facies of redwall limestone (mississippian) of uinta mountains, utah and colorado: iowa city, iowa state university, ph.d. dissertation, 423 p. eby, d.e., chidsey, t.c., jr., morgan, c.d., mcclure, k., humphrey, j.d., moore, j.n., taylor, l.h., and weyland, v.h., 2005, dolomitization of the mississippian leadville reservoir at lisbon field, utah [abs.]: american association of petroleum geologists annual convention, official program with abstracts, v. 14, p. a40. eby, d.e., chidsey, t.c., jr., sprinkel, d.a., and laine, m.d., 2009, a tale of two breccia types in the mississippian leadville limestone, lisbon field, paradox basin, southeastern utah [abs.]: american association of petroleum geologists annual convention abstracts, v. 18, p. 61. eskelsen, q.m., 1953, geology of the soapstone basin and vicinity, wasatch, summit, and duchesne counties, utah: salt lake city, university of utah, m.s. thesis, 56 p., 2 plates, scale 1:31,680. godfrey, a.e., 1985, karst hydrology of the south slope of the uinta mountains, utah, in picard, m.d., editor, geology and energy resources, uinta basin of utah: utah geological association publication 12, p. 277–293. hintze, l.f., and kowallis, b.j., 2009, geologic history of utah: provo, utah, brigham young university geology studies special publication 9, 225 p. hintze, l.f., willis, g.c., laes, d.y.m., sprinkel, d.a., and brown, k.d., 2000, digital geologic map of utah: utah geological survey map 179dm, scale 1:500,000. mayo, a.l., herron, d., nelson, s.t., tingey, d.g., and tranel, m.j., 2010, geology and hydrology of timpanogos cave national monument, utah, in sprinkel, d.a., chidsey, t.c., jr., and anderson, p.b., editors, geology of utah’s parks and monuments: utah geological association publication 28 (third edition), p. 269–283. mcdougald, w.d., 1953, geology of beaver creek and adjacent areas, utah: salt lake city, university of utah, m.s. thesis, 54 p., 2 separate plates, scale 1:31,680. mckee, e.d., 1969, paleozoic rocks of the grand canyon, in baars, d.l., editor, geology and natural history of the grand canyon region: four corners geological society, 5th field conference, powell centennial river expedition, p. 78–90. morgan, c.d., 1993, mississippian leadville limestone, in hjellming, c.a., editor, atlas of major rocky mountain gas reservoirs: new mexico bureau of mines and mineral resources, p. 94. parker, j.w., and roberts, j.w., 1963, devonian and mississippian stratigraphy of the central part of the colorado plateau: four corners geological society, 4th field conference guidebook, p. 31–60. sadlick, w., 1955, the mississippian-pennsylvanian boundary in northeastern utah: salt lake city, university of utah, m.s. thesis, 77 p. 10 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 sadlick, w., 1957, regional relations of carboniferous rocks of northeastern utah, in seal, o.g., editor, guidebook to the geology of the uinta basin: intermountain association of petroleum geologists 8th annual field conference, p. 57–77. spangler, l.e., 2005, geology and karst hydrology of the eastern uinta mountains—an overview, in dehler, c.m., pederson, j.l., sprinkel, d.a., and kowallis, b.j., editors, uinta mountain geology: utah geological association publication 33, p. 201–214. sprinkel, d.a., 2018, interim geologic map of the duchesne 30' x 60' quadrangle, duchesne and wasatch counties, utah: utah geological survey open-file report 689, 37 p., 2 plates, scale 1:62,500. utah division of oil, gas and mining, 2019, oil and gas production report by field, december 2018: online, https://oilgas. ogm.utah.gov/oilgasweb/publications/monthly-rpts-by-fld. xhtml?rpttype=fld, accessed april 2019. uga-geosite-chidsey-dead-horse-point.indd 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors thomas c. chidsey, jr.1, and hellmut h. doelling2 (retired) 1utah geological survey, po box 146100, salt lake city, utah 84114-6100, tomchidsey@utah.gov 2utah geological survey, 483 e. 200 s., manti, utah 84642 cover image: panorama to the east-northeast from dead horse point. dead horse point, southeastern utah 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e geosites refl ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: chidsey, t.c., and doelling, h.h., 2019, dead horse point, southeastern utah, in milligan, m., biek, r.f., inkenbrandt, p., and nielsen, p., editors, utah geosites: utah geological association publication 48, 17 p., https://doi.org/10.31711/ geosites.v1i1.56. presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president t.c. chidsey, jr., and h.h. doelling dead horse point 3 introduction the dead horse point geosite, within the state park by the same name, is located in the heart of the canyonlands region of utah between canyonlands and arches national parks (figure 1). the views are spectacular, sublime, awe-inspiring, and majestic, and hard to surpass anywhere on the colorado plateau. the mood of the vistas changes by season and time of day. here, one of nature’s engineers, in this instance the colorado river and its tributaries, has carved and exposed strata of late pennsylvanian (307 million years ago [ma]) to early jurassic (200 ma) age within just the past 5 million years (figures 2 and 3). dead horse point state park is located in a geologic area known as the paradox fold and fault belt in the northern part of the paradox basin, southeastern utah and southwestern colorado. the regional structural setting was created by the (1) movement of subsurface salt layers in the paradox formation (intermittently active from the pennsylvanian to the present day), (2) late cretaceous-early tertiary (about 70 to 30 ma) laramide orogeny (a regional mountain-building event), and (3) late tertiary-quaternary (23 ma to the present) regional uplift. these structural events have locally folded and fractured the rock layers in a manner that favored the deposition or accumulation of economic natural resource deposits (oil and gas, uranium, and potash). the signs of human activity to extract these deposits are apparent as one approaches the park or looks from the canyon edge at dead horse point. the question is often asked “why the name dead horse point?” according to one legend, around the late 1800s, dead horse point was used as a corral for wild horses roaming around the mesa that now includes an area called big flat to the north and the island in the sky district of canyonlands national park to the west. cowboys rounded up these horses and herded them across the narrow neck of land that separates the mesa from the point (figure 2). the neck, which is only 90 feet (27 m) wide, was then fenced off with juniper branches and brush. this created a natural corral surrounded by precipitous cliffs which afforded no escape. sadly, one time for some unknown reason, horses were carelessly left corralled on the waterless point where they died of thirst within view of the colorado river, 2000 feet (600 m) below (utah division of state parks, 2018). dead horse point is one of utah’s most visited scenic sites, and like delicate arch and the green river overlook (see these geosites elsewhere in this volume), the views are used on book covers, scenic calendars, and postcards. along with the incredible layered rocks, prominent features viewed from dead horse point include large up-ward folds called anticlines, mountains cored by volcanic remnants (laccoliths), immense solar evaporation ponds, and entrenched meanders of the colorado river— all revealing a geologic story. no compilation of classic utah geosites would be complete without dead horse point. figure 1. location map for the dead horse point geosite (red dot), southeastern utah, and surrounding parks, towns, and highways. 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 figure 2. geologic map of the dead horse point area, grand and san juan counties, utah. modified from hinrichs and others (1967), huntoon and others (1982), and doelling (2001). t.c. chidsey, jr., and h.h. doelling dead horse point 5 how to get there dead horse point state park is about 250 miles (400 km) or a little less than a 4-hour drive from salt lake city, utah, via interstate 15 (i-15), u.s. highway 6, i-70, and u.s. highway 191; 33 miles (53 km) or about 40 minutes if coming from the town of moab, utah. from highway 191 turn west onto state highway 313 and proceed 30 miles (48 km) to dead horse point state park (figure 1), bearing left at the junction with grand view point road, passing through the entrance station (fee required) to the parking lot for the overlook (restroom facilities, camping, and picnic grounds are available). a short walk along a flat, paved sidewalk leads to the overlook, 38°28'10" n., 109°44'21" w., elevation 5967 feet (1819 m). a stone wall separates the view areas from dangerous sheer cliffs. the visitor should walk west to east (or east to west) to obtain various panoramic views of the canyonlands area described in detail below. stratigraphy and depositional history consolidated sedimentary rocks seen in the vistas from dead horse point include strata of late pennsylvanian to early jurassic age. these strata have a cumulative thickness of 2700 to 3600 feet (820–1100 m) depending on which direction one looks (doelling and others, 2010) (figures 3 through 6). data in this section is largely taken from doelling and others (2010). figure 3. lithologic column showing formations and members exposed below dead horse point, including environments of deposition, thickness, age, weathering habits, and lithology. after doelling and others (2010). figure 4. panorama to the southwest from dead horse point; unannotated (a) and annotated (b). the gooseneck of the colorado river is on the left (south) and the shafer trail in the center; permian cutler formation at river level to jurassic navajo sandstone on the skyline in the distance in canyonlands national park (left of center). dead horse point is in the transition area between the areas where cutler facies change. note that the white rim sandstone pinches out below dead horse point (right of center). modified from doelling and others (2010). a b 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 figure 6. panorama to the east-northeast from dead horse point; unannotated (a) and annotated (b). the view shows the cane creek anticline, solar evaporation ponds, and jointed navajo sandstone outcrops of the behind the rocks area in front of the la sal mountains in the distance. the jurassic kayenta formation forms the rim of the overlook at the point followed down section by wingate, chinle, moenkopi, and cutler formations at the base. modified from doelling and others (2010). figure 5. panorama to the southeast from dead horse point; unannotated (a) and annotated (b). the view shows north-dipping permian cutler formation of the shafer anticline, pyramid butte, and the colorado river; abajo mountains in the far distance. the triassic through lower jurassic section is exposed in the distant cliffs and slopes below hatch point in distance (upper right). modified from doelling and others (2010). a b a b t.c. chidsey, jr., and h.h. doelling dead horse point 7 rocks of the pennsylvanian period pennsylvanian period rocks were deposited in a subsiding paradox basin beneath a restricted sea or embayment of an ocean that extended to the west of the park. to the northeast of this embayment, a land mass, known as the uncompahgre highland, emerged and rose (figures 7a and 7b). the boundary between the embayment and mountainous uplift was a sharp northwest-trending fault that lay to the northeast of dead horse point. rivers and streams carried the erosional debris from the mountains to the embayment where it was deposited. the uplift did not occur at a uniform rate. when the mountains rose rapidly, erosional debris and sediment were dumped into the embayment. when the uplift slowed, marine precipitates, such as gypsum, limestone, and salt were deposited. in middle pennsylvanian time, when the paradox formation was laid down, the embayment was intermittently cut off from open-ocean circulation and thick layers of salt were deposited. in late pennsylvanian time, the honaker trail formation was laid down (early to middle virgilian), and ocean waters freely circulated into the subsiding embayment, which continued to receive the erosional debris shed from the uncompahgre highland. paradox formation although not exposed in the area observed from dead horse point, the paradox formation played an important role in shaping the surrounding geology. it was deposited beginning about 308 ma as a marine deposit consisting of layered salt, anhydrite, shale, siltstone, limestone, and dolomite. these rocks were laid down within the subsiding basin in cycles that alternated from restricted-marine conditions with limited communication to the ocean (figure 7a) to open-marine conditions where the connection with the ocean was unimpeded and water circulated freely into the embayment. more often than not, physical barriers were established that cut off circulation with the ocean. the area had a hot, dry climate with high rates of evaporation. as the water evaporated, the dissolved salt was precipitated in the bottom of the bay. most of this salt is halite (nacl) (common table salt). however, at times the magnesium and potassium content of the seawater increased and sylvite (kcl) and carnallite (kclmgcl2.6h2o) were deposited. a typical paradox cycle began when connection with the open ocean was hindered. limestone and dark shale beds were first deposited; ensuing evaporation led to deposition of dolomite (primary), gypsum, and finally salt. thereafter free circulation with the open ocean was re-established only to gradually be hindered again. as many as 35 such depositional cycles are recognized in the paradox formation (rasmussen, 2010). these cycles have been correlated to glacio-eustatic sea-level fluctuations (goldhammer and others, 1991). the paradox formation is probably 3000 to 5500 feet (900–1700 m) thick beneath the area around dead horse point. of this amount, 75% consists of soluble salts. a single salt bed may be as much as 300 feet (90 m) thick. honaker trail formation like the paradox formation, most of the honaker trail formation is in the subsurface. however, the uppermost strata of the formation are exposed on the tops of the cane creek and shafer anticlines (figures 2, 5, and 6). the outcrops consist of interbedded sandstone, limestone, and siltstone. the sandstone is very fine to fine grained, well to moderately sorted, micaceous, and calcareous; some beds are cross-bedded. bedding is thick to massive. limestone interbeds are gray to very light gray, variably argillaceous (clayey), and 1 to 10 feet (0.3–3 m) thick. many limestone beds are fossiliferous, containing a marine fauna of crinoid debris, brachiopods, bryozoans, gastropods, foraminifera, and rare trilobites. the siltstone is micaceous, locally bioturbated, and cross-stratified. the honaker trail formation was deposited in the same embayment as the paradox formation (figure 7b). however, the circulation of seawater with the open ocean was no longer impeded. deposition was in very shallow water and the deposits were laid down in shallow marine, beach, lagoonal, and deltaic depositional environments. marine fusulinids indicate a latest pennsylvanian age (virgilian) for the honaker trail below dead horse point (doelling and others, 1994). the formation was deposited between 305 and 295 ma. from subsurface data we learn that the total honaker trail formation is 1600 to 2100 feet (490–640 m) thick (figure 3). the contact between the pennsylvanian honaker trail formation and the overlying pennsylvanian-permian rocks of the cutler formation is a paraconformity—rocks above and below the contact show little or no erosional relief, with beds of both formations parallel to one another but separated by a 10-million-year time gap. rocks of the pennsylvanian-permian period during latest pennsylvanian (late virgilian) time and throughout the permian, the uncompahgre highland to the northeast continued to rise and erosion eventually exposed proterozoic granitic and mafic rocks; radiometric ages of these rocks from near grand junction, colorado, range between about 1740 to 1400 ma (scott and others, 2001). the clastic erosional debris shed to the southwest into the dead horse point area was deposited as a series of alluvial fans at the foot of the highlands; the embayment was filled and the ocean shorelines retreated farther to the southwest (figure 7c). the iron content of the source rocks imparted a red coloration to the rocks deposited in this area. granites, made up of quartz, feldspar, mica, and small percentages of dark iron-bearing 8 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 figure 7. paleogeographic maps of utah during the: (a) middle pennsylvanian, paradox formation – 308 ma; (b) late pennsylvanian, honaker trail formation – 300 ma; (c) late early permian, cutler formation – 272 ma; (d) early triassic, moenkopi formation – 245 ma; (e) late triassic, chinle formation – 215 ma; (f) late triassic-early jurassic, wingate sandstone – 205 ma; and (g) early jurassic, kayenta formation – 200 ma. modified from blakey and ranney (2008). a b c d g e f t.c. chidsey, jr., and h.h. doelling dead horse point 9 minerals, were eroded and provided source material for sandstones rich in feldspar, or arkoses. the permian rocks below dead horse point were deposited between 285 and 270 ma. pennsylvanian-permian rocks in the dead horse point area are part of the cutler formation/group. northeast of dead horse point, these rocks are part of a thick sequence of conglomeratic, arkosic alluvial-fan sediments deposited in front of the uncompahgre highland and are given formation status (figure 3). these strata grade southwest into fluvial, coastal dune, and tidal flat deposits that are exposed in canyonlands national park southwest of dead horse point, where they are divided into five units—the pennsylvanian halgaito and elephant canyon formations, and the permian cedar mesa sandstone, organ rock formation, and white rim sandstone (see green river overlook geosite, island in the sky district, canyonlands national park, this volume). in that area, the cutler is elevated to group status and the five units are considered formations. dead horse point is in the transition between these areas. the eolian white rim sandstone (the upper formation) is the only formation of the cutler group that extends as far as dead horse point. the point where the white rim pinches out is visible below dead horse point (figure 4). in this area, beds below the white rim are more like the northeast facies than the other four formations of the cutler group. thus, these arkosic beds are still considered cutler formation, but are divided into informal upper and lower members (condon, 1997). the lower member is late pennsylvanian (late virgilian) to early permian (wolfcampian) in age based on regional correlations and fusilinid zones (sanderson and verville, 1990; condon, 1997). the upper and lower members of the cutler formation near dead horse point are alike and consist of interbedded arkosic sandstone, subarkosic sandstone, sandstone, conglomerate, micaceous siltstone, and limestone. many of the arkosic and subarkosic sandstone beds display trough cross-bedding, and cut-and-fill structures indicating fluvial deposition. conglomeratic lenses locally contain pieces of granite as much as 15 centimeters (6 in.) in diameter. quartzose sandstones are mostly fine grained, well sorted, and micaceous. many were deposited by the wind. the eolian rocks are mostly orange and red-orange, whereas the fluvially deposited arkoses, subarkoses, and conglomerates are dark red to purple. the gray limestones are marine deposits laid down during short intervals when the sea was able to push shorelines northeastward. these beds are more prevalent in the lower member of the cutler formation and form the basis for dividing the lower from the upper member (the popular shafer trail is located on this boundary below dead horse point). the number of limestone beds decreases northeastward. these are locally fossiliferous, containing brachiopods, bryozoans, gastropods, crinoid debris, and rare cephalopods and trilobites. the lower permian white rim sandstone is only present southwest of dead horse point and forms a white rim around the island in the sky in canyonlands national park (figure 4). the white rim is a fineto medium-grained quartzose sandstone exhibiting both planar and cross-stratified beds. the white rim formed as a near-shore, beach, and back-beach deposit. the white rim sandstone is 0 to 60 feet (0–18 m) thick. the upper contact of the permian strata with the overlying triassic beds is sharp, marked by local scouring and channeling and is unconformable. in most cases the plane of unconformity appears flat. the unconformity is regional and indicates a period of non-deposition lasting at least 30 to 35 million years. the lower member of the cutler formation is 180 to 220 feet (55–67 m) and the upper member is 700 to 1000 feet (200–300 m) thick in the dead horse point area (figure 3). rocks of the triassic period the triassic period in the dead horse point area is represented by the moenkopi and chinle formations and part of the wingate sandstone (figure 3). during deposition of the moenkopi, the open sea or ocean retreated farther to the southwest and deposition was from sluggish streams migrating over broad tidal flats that were at or near sea level (figure 7d). the uncompahgre highlands to the northeast were now much lower and the sediments derived from them contain more mica. fluvial environments dominated during deposition of the chinle (figure 7e) and the influence of the sea was gone. toward the end of chinle deposition, the uncompahgre highlands were nearly leveled or reduced to very low relief near dead horse point. moenkopi deposits were laid down approximately 244 ma and again between 242 and 238 ma and chinle deposits were laid down from 220 to 215 ma. events occurring between deposition of the moenkopi and chinle formations are not recorded by rocks in this area. sediments may have been deposited, but were later eroded during those 18 million years. the top of the triassic is within the wingate sandstone of the glen canyon group. moenkopi formation the moenkopi formation is divided into four members in the dead horse point area (figures 3 and 6), which in ascending order are: hoskinnini, tenderfoot, ali baba, and sewemup (sew-em-up). all four members total about 300 to 520 feet (90–160 m) thick at dead horse point. thicknesses of the four members vary independently. the hoskinnini member consists mostly of a chocolate-brown, fine-grained, poorly sorted, micaceous to sub-arkosic sandstone. it is poorly bedded and displays irregular to wavy laminae. locally, it has a thin bed of gypsum near the base. it forms a steep slope in its 10 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 lower part and a cliff in its upper part (figure 6). the hoskinnini member here is generally 60 to 110 feet (18–34 m) thick. its upper contact with the tenderfoot member is mostly sharp, an unconformity leaving as much as 2 million years with no record. the tenderfoot member forms a 100to 160-foot-thick (30–49 m) slope of light chocolate-brown, thin-bedded siltstone with widely spaced, 1to 3-foot-thick (0.3–1 m) sandstone ledges (figure 6). ripple cross-stratification is particularly widespread. the ali baba member forms a 100to 160-foot-thick (30–49 m) series of chocolate-brown ledges and slopes (figure 6). it consists of thinto medium-bedded siltstone with numerous sandstone beds. the sandstone beds are 1 to 10 feet (0.3–3 m) thick and display low-angle cross-stratification. the ali baba member locally contains a thin, yellow-gray limestone and calcareous siltstone at its base, which may correlate with the sinbad member of the moenkopi formation in the san rafael swell and circle cliffs areas of utah. the sewemup member is a 60to 120-foot(18–37-m) thick slope-forming unit consisting of homogeneous, gray-red to palered-brown siltstone, with widely spaced, thin (less than 3 feet [1 m] thick) sandstone ledges. the siltstone displays horizontal and ripple cross-lamination and is crisscrossed by thin gypsum veinlets. the contact with the overlying chinle formation is sharp and unconformable, but commonly poorly exposed. it is placed at the base of a distinctive white to mottled gritstone (figure 6), or between the gray-red or gray-green mudstone and siltstone of the lower part of the chinle and the orange-red siltstones of the upper part of the moenkopi. chinle formation the chinle formation is divided into three members in the cliffs below dead horse point: an unnamed lower slope-forming member, the black ledge member, and the upper church rock member (figures 3 and 6). the formation consists of complex interbedding and lensing arrangements of sandstone, pebble conglomerate, siltstone, mudstone, and rare limestone. the red-brown, tan, and gray-red sandstones are very fine to coarse grained, moderately to well sorted, quartzose, and slightly micaceous. primary sedimentary structures include low-angle cross-stratification, horizontal stratification, asymmetric ripples, and channeling, all of which point to deposition in a floodplain with northwest-flowing river channels, adjacent oxbow lakes, ponds, and swamps (figure 7e) (blakey and ranney, 2008). soft-sediment deformation, including disharmonic folds and low-angle detachments, is common, especially below thick sandstone ledges. pebble and intraformational conglomerates occur as lenses and in scour channels concentrated in the lower parts of sandstone beds. quartzose and micaceous siltstone is interbedded with the sandstones and conglomerates and displays low-angle cross-stratification and ripple lamination. mudstone is gray red to gray green, bentonitic, and poorly exposed. a few gray red nodular limestone beds are locally present. white to variegated gritstone locally marks the base of the chinle formation. the grit is poorly sorted and contains rounded to angular, coarse to pebble-size grains of quartz. where variegated, the gritstone may have existed as a paleosol. mudstone and siltstone dominate in the lower slope-forming member. sandstone and conglomerate channels, when found in the lower slope-forming member, are locally mineralized with uranium, vanadium, and copper minerals. fine carbonaceous plant debris is abundant in these channels. the lighter overall color of this lower member is caused by reduced iron in this part of the formation. this reduction of iron commonly extends as much as 3 feet (1 m) into the underlying moenkopi formation. the lower slope-forming member is 70 to 90 feet (20–30 m) thick, and the upper contact is abrupt. the black ledge is dominated by red-brown sandstone and black, desert varnish-stained conglomerate. the sandstones commonly contain scattered logs and branches of petrified wood. lowermost lenses of sandstone are locally mineralized with uranium and copper minerals. the black ledge is about 90 to 110 feet (30–34 m) thick. the upper contact is gradational into the church rock member. the church rock member is mostly a red-brown sandstone and siltstone, but sandstone ledges are more common in the lower part. some beds include distinctive ripple-laminated sandstone, but the bedding in much of this unit is indistinct, and the rock breaks into equidimensional fragments. blocky, red-brown, finegrained, well-sorted, thick-bedded sandstone, that mimics the overlying wingate sandstone, is common in the upper 10 to 30 feet (30–49 m) of the member. these upper sandstones have informally been referred to as the hite beds. the church rock member is 160 to 200 feet (49–60 m) thick. the contact of the church rock member with the overlying wingate sandstone is sharp but conformable, commonly being placed below the massive cliff of well-sorted sandstone typical of the wingate. no regional channeling or angular unconformity is apparent, and the lowermost part of the wingate does include thin, bedding-parallel sandstone and siltstone beds that suggest continuous deposition across the chinle and wingate contact. the chinle and wingate contact was once thought to represent an unconformity, named the j-0 unconformity, at the triassic-jurassic boundary (pipiringos and o’sullivan, 1978). however, the wingate contains beds of both triassic and jurassic age (molina-garza and others, 2003; lockley and others, 2004; lucas and others, 2005). the j-0 unconformity, if it exists here, is within the triassic and likely represents a short period of time. t.c. chidsey, jr., and h.h. doelling dead horse point 11 rocks of the triassic-jurassic period the wingate sandstone is a prominent cliff former below the promenade at dead horse point and below the island in the sky bench to the southwest (figures 4 through 6). the wingate forms red-brown, nearly vertical cliffs streaked and stained with “desert varnish.” the chinle and moenkopi slopes below the cliff are commonly littered with large blocks of the wingate. the wingate sandstone is ordinarily described as one massive unit because partings or bedding planes are rare except near the base of the formation. the wingate consists mostly of light-orange-brown, moderate-orange-pink, or pale-red-brown, fine-grained, well-sorted, cross-bedded sandstone. the rock is usually well cemented and well indurated; weathered exposures are nearly smooth. the eolian wingate was deposited in a great erg that extended from the four corners area to north-central utah (figure 7f) (blakey and ranney, 2008), as indicated by the high-angle cross-bedded sandstone. however, the bedding-parallel sandstone beds near the base of the formation suggest that fluvial processes were still a large part of its early depositional history. the formation is 300 to 400 feet (90–120 m) thick in the dead horse point area. the contact of the wingate with the overlying kayenta formation is generally sharp and conformable as seen from a distance, but difficult to place close-up (figures 4 through 6). generally, the line is placed at the horizon where the smooth cliff is replaced by thick cliffy ledges. rocks of the lower jurassic period with the exception of the quaternary unconsolidated sediments, lower jurassic strata form the top of the geologic column at dead horse point (figure 3). these strata include the upper part of the wingate sandstone, kayenta formation, and navajo sandstone, and were deposited between 205 and 187 ma. the kayenta was deposited in a fluvial environment (figure 7g), whereas the navajo formed in a sand-dune desert. a few sandstone beds in the upper kayenta were also deposited by the wind. kayenta formation the kayenta formation caps most of the upper benches at dead horse point; the view points are located on this unit (figures 3 through 6). the kayenta consists mostly of stream-deposited sandstone lenses, with lesser amounts of eolian sandstone, intraformational conglomerate, siltstone, and shale. the unit is primarily red brown, but individual lenses and beds vary considerably in color; some are purple, lavender, tan, orange, or white. in outcrop, the kayenta is ledgy and step-like. sandstone in the kayenta exhibits both high-angle and low-angle cross-bedding. some lenses display channeling, current ripple marks, and rare slump features. the grain size is more variable than in the wingate and navajo, ranging mostly from fine to medium. siltstone, shale, and intraformational conglomerate appear as partings or are interlayered with the sandstone. these softer constituents are rare in the lower half of the formation and become common in the upper part. the kayenta formation was deposited in a sandy braided river system although the environment was arid (figure 7g). perennial streams flowed west from the remaining ancestral rockies and the appalachians far to the east (lynds and hajek, 2006; blakey and ranney, 2008). floodplains (overbank deposits) formed adjacent to active channel belts where the channel threads were only about 3 feet (1 m) deep. tracks of dinosaurs, including eubrontes, a tridactyl (three-toed) theropod (bipedal) dinosaur, are commonly found in the kayenta in dead horse point state park and elsewhere in southern utah. the kayenta formation is 220 to 300 feet (67–90 m) thick in the dead horse point area. it forms the bench tops and is rarely completely exposed. the upper contact is mostly sharp, but intertonguing between the kayenta and the overlying navajo sandstone is common. remnants of the navajo sandstone rest on the benches, and at least at dead horse point, are far from the cliff edge. navajo sandstone the navajo sandstone is not exposed at dead horse point but can be viewed to the east in an area known as behind the rocks (figure 6). the navajo is mostly exposed as cliffy to rounded bare sandstone. the formation is an orange to light-gray, massive cross-bedded sandstone. the eolian navajo is a classic example of a major sahara-like erg (dune) environment. structural and geologic history regional setting the fold and fault belt of the paradox basin, where dead horse point is located, extends southeast from here into colorado. this structural belt is characterized by a series of northwest-trending anticlines and faults that developed in response to movement caused by crustal forces or by shallow deposits of pennsylvanian salt. the most obvious structural feature observed from dead horse point is the spectacular cane creek anticline that fills the panoramic view to the east (figure 6). the events that caused this and other structural features to form in the region began over a billion years ago during the precambrian (proterozoic) when movement began on high-angle basement faults and fractures 1700 to 1600 ma (stevenson and baars, 1987). the paradox basin of southeastern utah and southwestern colorado, like the fold and fault belt within it, is an elongate, northwest-southeast-trending evaporitic basin. the paradox basin predominantly developed during pennsylvanian time about 318 to 300 ma as part of a pattern of basins and fault-bounded uplifts from what is now utah to oklahoma. one result of this tectonic 12 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 event was the uplift of the ancestral rockies in the western united states. the uncompahgre highland (uplift) of eastern utah and western colorado formed the westernmost range of the ancestral rockies during this ancient mountain-building period. the uncompahgre highland is bounded along its southwestern flank by a large basement-involved, high-angle reverse fault (frahme and vaughn, 1983; kluth and duchene, 2009). this fault has been identified from geophysical seismic surveys and exploration drilling. as the mountains rose, movement along this fault caused an accompanying depression, a foreland basin, to form to the southwest—the paradox basin. rapid basin subsidence, particularly during the pennsylvanian and into the permian, accommodated deposition of large volumes of deeper basin evaporitic and marine sediments which intertongued with basin-margin non-marine arkosic material shed from the mountain area to the northeast (hintze and kowallis, 2009). later, the uncompahgre highland was eroded down during the triassic and jurassic. the area was uplifted again during the late cretaceous and early tertiary laramide orogeny to form the uncompahgre uplift observed today (hintze and kowallis, 2009). salt movement salt, which was deposited as part of the paradox formation, has a low specific gravity, is ductile and less dense than the surrounding rocks, and behaves plastically. the salt moves due to high confining pressure from the weight of the overlying column of rocks or sediments. salt movement can push up and raise overlying strata in some places, while the salt is squeezed and thinned in others. this movement (often piercing the strata above [diapiric movement]) progresses along zones of weakness or areas of low confining pressure, forming large folds such as the cane creek and shafer anticlines (figures 2, 5, and 6). the weak zones likely developed above and along the northwest-trending basement faults in the region which experienced continued movement (shoemaker and others, 1958). salt-cored anticline development has been intermittently active from the pennsylvanian to the present day (shoemaker and others, 1958; cater, 1970; case and joesting, 1973; baars and doelling, 1987; doelling, 1988; oviatt, 1988). jurassic and cretaceous sedimentation the region continued to lie near sea level, but received continental deposition through early jurassic time. because no bedrock units younger than the early jurassic navajo sandstone are present near the park, no record exists here of the geologic events that happened in the interval from 187 to 1.5 ma. the area likely received continental deposition during the remainder of the jurassic and the early cretaceous and shallow-marine deposition during the late cretaceous, as attested in surrounding areas. late cretaceous–early tertiary folding, faulting, and intrusions large uplifts and basins developed in the colorado plateau during the laramide orogeny between the latest cretaceous (maastrichtian, about 70 ma) and the eocene (about 38 ma). the northern end of the laramide monument upwarp can be viewed to the south from dead horse point where permian-age rocks are extensively exposed (figures 2, 4, and 5). other nearby laramide features include the san rafael swell to the west, the circle cliffs uplift and henry mountains basin to the southwest, the uinta basin to the north, and a rejuvenated uncompahgre uplift to the northeast. the regional dip of strata in the area is 2 to 4 degrees northward. the gentle northwest-trending anticlinal and synclinal folds in the dead horse point area disrupt this regional slope. the axes of the cane creek and shafer anticlines (figures 5 and 6) are aligned directly over local bulges of paradox salt. the overlying rocks were fractured and locally extended by minor faults just off the crest of the anticlines (morgan and others, 1991). other anticlines in the paradox fold and fault belt that are aligned with salt walls have major normal faults such as the moab fault to the northeast. total displacement across the moab fault is estimated at 1500 to 2400 feet (460–730 m) of displacement (doelling and others, 1994; mauch and pederson, in press). the timing of these events is unclear, except that they occurred between 80 and 15 ma. the la sal mountains (figure 6), on the horizon to the east above the cane creek anticline, are a classic example of a laccolithic intrusion. the la sal mountains are a complex of granitic (diorite porphyry) rocks intruded into a salt wall about 27.9 to 25.1 ma during the oligocene (nelson, 1998). these peaks are more than 12,000 feet (3600 m) above sea level and were glaciated during the pleistocene ice ages (hintze and kowallis, 2009). in the distance to the southeast, another laccolithic intrusion, the abajo mountains, can be seen. this intrusion is dated between 28 and 25 ma (nelson, 1998). jointing and fractures the sandstones in and around dead horse point are brittle and when folded or bent produce joints and fractures. fracture patterns and joint sets are related to regional tectonics and salt movement. largescale, northeast-trending fractures are found along the cane creek anticline (morgan, 1992). jointing is prevalent in the cutler, wingate, and kayenta formations. a dominant set of joints strikes northwest to north-northwest (doelling and others, 1994). joints are closely spaced in many areas, but vary with lithology and bed thickness. late tertiary–quaternary regional uplift and erosion the colorado plateau began rising in late cenozoic time during the miocene (23 ma) (hunt, 1956; lucchitta, 1979; hintze and kowallis, 2009). this regional uplift changed the landscape from t.c. chidsey, jr., and h.h. doelling dead horse point 13 one of deposition to one of massive erosion. several thousand feet of sedimentary rocks have been removed by the erosive processes of mass wasting, wind, and running water. most of this material has been carried to the sea by the colorado river system. as the colorado plateau rose, the colorado river and its tributaries rapidly cut into the strata. the incision rate near the center of the colorado plateau (lee’s ferry, arizona) is calculated at about 35 centimeters per thousand years (14 in./k.y.) (pederson and others, 2013). the results of this action are the countless canyons and entrenched meanders visible from dead horse point. most of the eroded material from the colorado plateau has been carried to the sea by the colorado river system. economic geology dead horse point is surrounded by activity associated with the extraction of three principal minerals—potash, petroleum, and uranium. all three saw periods of “boom and bust” during the 20th century. they have been, and remain, an important part of the local economy, as well as the subject of debate in how the lands around the dead horse point state park should be managed. only the activities of potash production and targets for oil drilling can be seen from dead horse point. potash this region of the paradox basin is underlain by the salt-bearing paradox formation. seventeen of the depositional cycles in the paradox formation contain halite (common table salt [nacl]) and potash salts which are mined near dead horse point (doelling and others, 1994). the most common question asked at dead horse point state park is “what are those big blue ponds off to the east?” the “big blue ponds” are solar evaporation ponds (figure 6) designed to extract potash from brines pumped from wells penetrating the salts of the paradox formation. potash is an important economic resource used to manufacture a variety of products including fertilizer (93%), chemicals, pharmaceuticals, soap, glass, synthetic rubber, and explosives (tripp, 2010; carney, 2019). the potash minerals are carnallite and sylvite (ritzma, 1969). sylvite is less abundant but is the primary commercial mineral. chemically, it is composed of 52.4% potassium and 47.6% chlorine. in 1964, texas gulf sulfur inc. (now intrepid potash, inc.) began underground room-and-pillar mining of potash from the paradox formation cycle number 5 along the cane creek anticline at depths from 2500 to 3000 feet (760–900 m). underground mining operations were difficult due to intensely folded and faulted salt beds, salt movement, pockets of natural gas, high temperatures, and mine-roof maintenance problems that made it uneconomic to mine underground (phillips, 1975). before mine operations began, 18 miners were killed in a gas explosion while constructing lateral shafts (huntoon, 1986). in 1970, the operations were converted to solution mining. colorado river water is now pumped into the old mine workings to dissolve the potash and salt (figure 8). the resulting brine is pumped out and piped to the solar evaporation ponds where the minerals are eventually harvested. potash production is roughly 100,000 tons annually (andrew rupke, utah geological survey, verbal communication, 2019). halite is a by-product and 40,000 to 60,000 tons are produced annually. the halite is used in water softeners, animal feed, for road salt, and oil field drilling fluids. between 1994 and 2003, potash prices remained stable ranging from $150 to $200 per ton. potash prices started rising in 2005 and peaked during 2009 at nearly $800 per ton, but by 2019 had declined to over $250 per ton (andrew rupke, utah geological survey, verbal communication, 2019). petroleum since the early 1920s, the immediate area around dead horse point has been the site of oil and gas exploration drilling (chidsey and others, 2016; chidsey and eby, 2017). oil and gas is produced from the cane creek shale zone of the paradox formation. the cane creek shale zone consists of thinly interbedded, black, organic-rich marine shale, dolomitic siltstone, dolomite, and anhydrite (morgan, 1992; grove and others, 1993; chidsey and others, 2016). petroleum is trapped in fractured reservoirs usually on the crest of anticlinal closures. the cane creek anticline (figure 6) was the most obvious structural drilling target and was tested sporadically in the cane creek oil field near the colorado river eight times since 1924. however, only 1887 barrels of oil and 25 million cubic feet of gas were produced from the structure (stowe, 1972). in 1960, lands on the crest of the cane creek anticline were withdrawn from oil and gas leasing to prevent interference with potash mining (smith, 1978). figure 8. schematic cross section showing the solution-mining process on the cane creek anticline where salt and potash in the sylvite number 5 depositional cycle of the pennsylvanian paradox formation are dissolved by injecting river water. the resulting brine is withdrawn by an extraction well and pumped to the solar evaporation ponds where the minerals are eventually harvested. from doelling and others (2010). 14 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 the oil field area has since been abandoned as a target for oil. horizontal drilling was never used at the cane creek anticline. the 1990s saw dramatic improvements in the ability to successfully complete wells in fractured reservoirs using horizontal drilling technology and the cane creek shale zone again became an important target. several new fields were discovered in the area near dead horse point as seen on the drive to the park (morgan, 1992; grove and others, 1993; chidsey and others, 2016) and exploration is ongoing; however, none of these fields can be seen from dead horse point. panoramic viewpoints dead horse point is well known for its remarkable panoramic views (figures 4 through 6). although there are some features of interest on dead horse point itself, most of the incredible sites associated with the park are not within its boundaries—the exciting panoramic views looking over the canyonlands area of utah. changes in the strata’s color, rock types, thickness, composition, weathering, and erosion enhance the interest of these magnificent sights. if at all possible, the visitor should take advantage of road logs found in doelling and chidsey (2012). dead horse point should be visited before and after enjoying the sites along these roads. panoramic view #1 (figures 4 and 9) the colorado river meandering around the gooseneck is a favorite view (figure 4); note the shafer trail. during the laramide orogeny (latest cretaceous through eocene time [about 70 to 30 ma]), the colorado plateau region developed basins and uplifts. these basins and uplifts were all elevated during the general uplift of the colorado plateau beginning during the miocene epoch (23 ma). rocks in the area of dead horse point dip gently northward from the monument uplift. the axis of the monument uplift is northsouth. the view from dead horse point is generally along the axis, which plunges northward. strata dip gently eastward and westward on each side of the axis. the colorado river has therefore eroded deeper into the earth’s crust as it crossed the uplift axis, revealing the oldest formations exposed in this part of the colorado plateau region and creating the canyonlands country. dead horse point owes its beautiful panoramic vista to the fact that the river has cut into these older spectacularly colored rock formations. the stratigraphic section begins with the permian cutler formation at river level to jurassic navajo sandstone on the skyline in the distance in canyonlands national park (figures 2 and 4). dead horse point is in the transition area where cutler facies change. the permian white rim sandstone that was so well-displayed to the southwest pinches out just below dead horse point (figure 4). without the white rim, the cutler formation is subdivided into informal upper and lower members. the upper member is dominated by fluvial and eolian arkosic sandstone, representing material shed from the uncompahgre highlands of the ancestral rockies to the northeast. the lower member forms a prominent bench (see shafer trail below) and is characterized by numerous, distinct, thin limestone ledges containing marine fossils. panoramic view #2 (figures 5 and 9) from panoramic view #2 we see pyramid butte, the colorado river, and the north-dipping cutler formation on the flank of the shafer anticline; the needles district of canyonlands national park and abajo mountains are in the distance to the south (figure 5). the abajos are another laccolithic intrusion, dated between 25 and 28 ma (nelson, 1998). the light band in the distant cliffs is in the lower part of the chinle formation with the vertical cliffs of the wingate sandstone above. during the middle to late cenozoic, the ancestral colorado river and its tributaries flowed through meandering channels in wide valleys on easily eroded rocks such as the now-removed cretaceous mancos shale. once these river channels were established, they later became superimposed and entrenched into buried structures, such as the shafer and cane creek anticlines, and into resistant rocks such as the wingate and navajo sandstones. the deeply entrenched gooseneck of the colorado is a classic example of this geomorphologic process (figure 2). this magnificent entrenched meander of the colorado river is easily seen just below dead horse point to the southwest. at the narrowest neck of the peninsula, the distance across bedrock between the river channel on either side of the peninsula is less than a quarter mile (0.4 km). here, it is entrenched into about 400 feet (120 m) of the lower cutler formation. the river level is about 2000 feet (600 m) below the elevation of the pavilion at the dead horse point overlook. for a distance of about 4 river miles (6 km) along the gooseneck, the river proceeds less than 0.3 land miles (0.5 km)! the river is capable of cutting through the 400-foot (120-m) wall of lower cutler rocks and has cut through similar “necks” in other areas, but it will probably take thousands of years to do so. the shafer anticline (sometimes referred to as shafer dome) can be seen just south of dead horse point (figure 5). the axis curves essentially around the point, trending northwest on the west, eastwest to the south, and northeast to the east (figure 2). although less obvious than the cane creek anticline to the east, the shafer anticline is a major structural feature. beds dip 5 to 10 degrees from the crest, where rocks as old as the pennsylvanian honaker trail formation are exposed along the colorado river on the eastern axis of the structure. a series of minor, high-angle extension faults generally parallels the northwest-trending axis along the southwest limb (figure 2). the shafer anticline most likely developed from salt bulging in the paradox formation, and movement on deep, older faults. the structure was drilled in a few localities, but no petroleum was found. t.c. chidsey, jr., and h.h. doelling dead horse point 15 panoramic view #3 (figures 6 and 9) looking east we see an incredible view of the cane creek anticline (middle ground), solar evaporation ponds, and jointed navajo outcrops of the behind the rocks area in front of the la sal mountains in the distance (figure 6). the northwest-trending cane creek anticline is the largest structural feature in the area (figure 2). beds dip gently, 5 to 8 degrees in both directions away from the crest. a series of minor, high-angle extension faults parallel the axis of the anticline along the northeast limb (figure 2). the cane creek anticline likely developed from a combination of salt bulging and renewed movement on deep, older structures. the kayenta formation forms the rim of the overlook followed down section by wingate, chinle, moenkopi, and upper cutler at the base. the various members of the chinle and moenkopi are well displayed (figure 6). the orange-red upper member of the cutler formation is very arkosic and forms steep, ledgy slopes, whereas the lower member of the cutler capped by thin limestone overlying arkosic clastic rocks forms the relatively flat and extensive terrain observed. the pennsylvanian honaker trail formation is exposed in the core of the cane creek anticline where the colorado river cuts across the axis of the structure. the upper 240 feet (73 m) of the honaker trail crops out here as yellow-gray to gray ledgy cliffs. it consists of interbedded sandstone, limestone, and siltstone. some of the limestone beds are fossiliferous containing crinoid debris, brachiopods, bryozoa, corals, gastropods, and rare trilobites. these rocks are the oldest exposed in the area. the brine from the solution mine (described earlier, figure 8) is pumped to the large solar evaporation ponds observed from dead horse point (figure 6); it is saturated with respect to salt and close to saturation with respect to sylvite. halite is a contaminant and is partially removed in a carefully controlled evaporative process. brine, upon evaporating to a suitable concentration, is drained into a new pond, each time leaving some of the halite behind. the process is repeated several times before the final liquid becomes suitably concentrated in potassium salts. the slurry is then pumped to a processing facility where the potash is separated from the remaining salt by flotation (bryce tripp, utah geological survey, verbal communication, 2009). the potash is then bagged and sold. the potash solar evaporation ponds so prominently displayed east of dead horse point actually consist of 23 ponds in four areas. the ponds cover a total of 400 acres (160 ha) and are 2 to 3 feet (0.6–0.9 m) deep. the solar ponds are lined with heavy vinyl to prevent leakage of the brine into the groundwater aquifers and colorado river. the beautiful blue color of the water is the result of a dye (similar to food coloring) which is introduced to enhance evaporation. after evaporation, the potash and salt crystals are harvested by 25-ton scraper-loaders that can also be seen from dead horse point. if the product cannot be harvested within two years the operation becomes uneconomical. the operators do not like wet years and cloudy skies. figure 9. google earth image (© 2018 google) of dead horse point (red dot) and surrounding area showing the direction (red arrows) of the views shown on figures 4 through 6. 16 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 acknowledgments support for this paper was provided by the utah geological survey (ugs). we thank the staff at dead horse point state park for providing access to park lands. cheryl gustin, lori steadman, jay hill, and jim parker (deceased) of the ugs drafted figures. this paper was carefully reviewed by michael d. vanden berg, stephanie m. carney, michael d. hylland, and bill keach of the ugs, along with the editors of this publication. their suggestions and constructive criticism greatly improved the manuscript. references baars, d.l., and doelling, h.h., 1987, moab 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r.s., geissman, j.w., and lucas, s.g., 2003, paleomagnetism and magnetostratigraphy of the lower glen canyon and upper chinle groups, triassic-jurassic of northern arizona and northeastern utah: journal of geophyscial research, v. 108, no. b4, p. 1–24. morgan, c.d., 1992, horizontal drilling potential of the cane creek shale, paradox formation, utah, in schmoker, j.w., coalson, e.b., and brown, c.a., editors, geologic studies relevant to horizontal drilling—examples from western north america: rocky mountain association of geologists guidebook, p. 257–265. morgan, c.d., yonkee, w.a., and tripp, b.t., 1991, geological considerations for oil and gas drilling on state potash leases at cane creek anticline, grand and san juan counties, utah: utah geological survey circular 84, 24 p. nelson, s.t., 1998, reevaluation of the central colorado plateau laccoliths in light of new age determinations: u.s. geological survey bulletin 2158, p. 37–39. oviatt, c.g., 1988, evidence for quaternary deformation in the salt valley anticline, southeastern utah, in salt deformation in the paradox region: utah geological and mineral survey bulletin 122, p. 61–76. pederson, j.l., cragun, w.s., hidy, a.j., rittenour, t.m., and grosse, j.c., 2013, colorado river chronostratigraphy at lee’s ferry, arizona, and the colorado plateau bull’s-eye of incision: geology, v. 41, no. 4, p. 427–430. phillips, m., 1975, cane creek mine solution mining project, moab potash operations, texasgulf inc., in fassett, j.e., editor, canyonlands country: four corners geological society, 8th annual field conference, p. 261. pipiringos, g.n., and o’sullivan, r.b., 1978, principal unconformities in triassic and jurassic rocks, western interior united states—a preliminary survey: u.s. geological survey professional paper 1035-a, 29 p. rasmussen, d.l., 2010, halokinesis features related to flowage and dissolution of the pennsylvanian hermosa salt in the paradox basin, colorado and utah [abs.]: american association of petroleum geologists, rocky mountain section meeting program with abstracts, p. 59. ritzma, h.r., 1969, potash in the san juan project area, utah and colorado, in mineral resources, san juan county, utah, and adjacent areas: utah geological and mineral survey special studies 24, p. 17–33. sanderson, g.a., and verville, g.j., 1990, fusilinid zonation of the general petroleum no. 45-5-g core, emery county, utah: mountain geologist, v. 27, no. 4, p. 131–136. shoemaker, e.m., case, j.e., and elston, d.p., 1958, salt anticlines of the paradox basin: intermountain association of petroleum geologists 9th annual field conference, p. 39–59. scott, r.b., harding, a.e., hood, w.c., cole, r.d., liviccari, r.f., johnson, j.b., shroba, r.r., and dickerson, r.p., 2001, geologic map of colorado national monument and adjacent areas, mesa county, colorado: u.s. geological survey geologic investigations series i-2740, 40 p., 1 plate, scale 1:24,000. smith, k.t., 1978, cane creek, in fassett, j.e., editor, oil and gas fields of the four corners area: four corners geological society, v. 2, p. 624–626. stevenson, g.m., and baars, d.l., 1987, the paradox—a pull-apart basin of pennsylvanian age, in campbell, j.a., editor, geology of cataract canyon and vicinity: four corners geological society, 10th field conference, p. 31–55. stowe, carlton, 1972, oil and gas production in utah to 1970: utah geological and mineral survey bulletin 94, 179 p. tripp, b.t., 2010, utah potash—resources, production, and exploration: utah geological survey, survey notes, v. 42, no. 1, p. 1–3. utah division of state parks, 2018, dead horse point state park: https://stateparks.utah.gov/parks/dead-horse/discover/, accessed january 2019. uga-geosite-chidsey-ferron.indd 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors thomas c. chidsey, jr.1, and paul b. anderson2 1utah geological survey, po box 146100, salt lake city, utah 84114-6100, tomchidsey@utah.gov 2geologic consultant, po box 101, emery, utah 84522 cover image: clinoform medial facies in the kf-1-iv[a] bedset. ancient delta deposits in the ivie creek area, ferron sandstone member of the mancos shale, western san rafael swell, east-central utah 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 t.c. chidsey, jr., and p.b. anderson ancient delta deposits in ivie creek area, ferron sandstone 3 introduction in contrast to the beautiful array of colorful layers and spectacular cliffs of the triassic and jurassic (251 to 148 million years ago [ma]) sections in the san rafael swell of east-central utah, most of the upper cretaceous (96 to 86 ma) mancos shale produces a drab, barren landscape. however, lying within the mancos, the ferron sandstone, is the most studied unit in the san rafael swell. the ferron has world-class outcrops of rock layers deposited near the shorelines of a sinking, fluvial(stream) dominated delta system. along the west flank of the san rafael swell (figure 1), the 80-mile-long (130 km) ferron outcrop belt of cliffs and side canyons (e.g., the coal cliffs, molen reef, and limestone cliffs [not actually limestone, just misnamed]) provides a three-dimensional view of vertical and lateral changes in the ferron’s rock layers (facies and sequence stratigraphy), and, as such, is an excellent model for fluvial-deltaic oil and gas reservoirs worldwide (e.g., chidsey and others, 2004). the ferron sandstone consists of a stacked series of sandstone-dominated rock layers deposited as local sea level rose and fell (depositional transgressive-regressive cycles). collectively these sandstone layers form an eastward-thinning wedge into the mancos shale. the ivie creek area along the north side of interstate 70 (i-70) (figure 2), displays spectacular and abrupt changes in two of these regional-scale depositional cycles (“parasequence sets” referred to as kf-1 and kf-2 [figure 3]). these cyclic deposits represent the those typically found in a deltaic oil and gas reservoir. the kf-1 parasequence set displays spectacular clinoforms. the term “clinoform” is used to identify a group of beds which are inclined seaward in an en echelon pattern and generally separated from one another by a distinctive bounding surface. as a reservoir model, the ferron in the ivie creek area displays variations that influence the reservoir (both its compartmentalization and permeability). these outcrops are often a standard stop for geology figure 1. the san rafael swell and vicinity, east-central utah, showing the location of the ferron sandstone geosite as well as major physiographic features, surrounding towns, and highways. 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 field trips. thus, the ferron sandstone in the ivie creek area was selected as a geosite (also see anderson’s ferron geosite in emery county, this volume) providing a more complete picture of this classic outcrop. how to get there the ferron sandstone ivie creek area is about 170 miles (270 km) or a 2 hour and 45-minute drive from salt lake city, utah, via i-15 and u.s. highway 6 to state highway 10 to i-70. proceed eastbound on i-70 for 6 miles (10 km) to the junction with county road 912 (miller canyon road), cross the overpass and re-enter i-70 heading westbound for about 2.3 miles (3.7 km) to a well-used but difficult to spot dirt pull-off area for the best view (to the north) of the ferron clinoforms, parasequence sets, and fluvial-deltaic facies in the “ivie creek amphitheater,” an informal name applied to a broad, curving area of cliffs north of i-70 (38°48'35" n., 111°15'10" w., elevation 5813 feet [1772 m]) (figures 1 and 3). the ferron ivie creek geosite is about 61 miles (98 km) west of the junction of u.s. highway 6 and i-70 near the town of green river, utah (figure 1). upon approaching the ivie creek stop carefully slow down with flashers on and pull off the interstate down a slight incline to a well-used parking area along the right-of-way fence. there are no well-marked trails to examine the ferron outcrops up close; the terrain is rugged with vertical cliffs and requires first negotiating a fence designed to keep livestock and wildlife from the interstate. anderson and others (1997) provided a geologic field guide to several key locations within the ferron section in the ivie creek amphitheater (figures 2 and 3). geologic overview san rafael swell the ferron sandstone is exposed in an outcrop belt along the west flank of the san rafael swell, a broad, asymmetric, north-southto southwest-northeast-trending anticline about 75 miles (120 km) long and 35 miles (56 km) wide. the structure formed in response to compressional forces of the laramide orogeny between latest cretaceous time (about 70 ma) and the eocene (about 40 ma) (hintze and kowallis, 2009 and references therein) (figure 1). uplift and erosion has made it a showcase of colorado plateau geology with a colorful array of sedimentary rocks over 7000 feet (2100 m) thick, ranging in age from permian (276 ma) to cretaceous (86 ma), exposed in spectacular cliffs along cuestas, mesas, and deep canyons. the rocks in the san rafael swell are folded, faulted, jointed, fractured, and uplifted, with deformation likely controlled by a large, blind, basement-involved reverse fault (up on the west side) bounding the east flank of the structure. beds figure 2. geologic map of the ferron sandstone geosite along ivie creek (red box) and surrounding area, westernmost flank of the san rafael swell. the blue box represents the area of paleogeographic maps shown on figure 15. modified from doelling and kuehne (2016). t.c. chidsey, jr., and p.b. anderson ancient delta deposits in ivie creek area, ferron sandstone 5 in the ferron outcrop belt dip west from 2° to 12° with only a few minor faults present (gloyn and others, 2003; quick and others, 2004). small to large subsidiary anticlines and synclines are found north to south along the uplift. ferron field produces natural gas from the ferron sandstone in one of these subsidiary structures northeast of the ivie creek area; farther north coalbed methane is produced from ferron coal seams in a series of major fields (wood and chidsey, 2015). ferron sandstone stratigraphy the mancos shale is divided into four members along the west flank of the san rafael swell, which in ascending order are: tununk, ferron sandstone, lower blue gate, and emery sandstone (figure 4). the ferron sandstone and other upper cretaceous units pinch out to the east; the ferron is stratigraphically equivalent to the juana lopez member of the mancos (figure 5). the mancos is about 2000 to 2500 feet (600–760 m) thick in the san rafael swell (hintze and kowallis, 2009 and references therein). it epitomizes deposits from the cretaceous interior seaway on its western margin and the influence of the sevier orogenic belt in western utah. the ferron sandstone ranges in thickness from 150 to 400 feet (45–230 m) on the west flank of the san rafael swell (witkind, 1979, 1980, 1988; weiss and others, 1990; doelling and others, 2009, 2015; hintze and kowallis, 2009 and references therein; doelling and kuehne, 2012, 2016). the ferron is informally divided into lower and upper sections. the lower ferron consists of two thin-bedded silty intervals called, in ascending order, the clawson figure 3. the upper cretaceous ferron sandstone member of the mancos shale in the ivie creek area north of i-70 showing the area well known for the contrasting delta-front architectural styles displayed in the kf-1 and kf-2 ferron parasequence sets (see figure 6); view to the northeast. the kf-1 has distinctive steeply rightto left-dipping clinoforms representing fluvial-dominated deposition (white box represents location of close-up shown on figure 9a); note that kf-1 bedsets are also annotated (see figure 9c). the kf-2 represents a significant change to wave-modified deposition (subtle left-to-right dips) as seen by shoreface, distributary channel, bay, and coastal plain/swamp facies. modified from anderson and others (2004). photograph by michael chidsey, sqwak productions inc. figure 4. stratigraphic column of exposed cretaceous rocks along the west flank of the san rafael swell, including age, thickness, lithology, weathering profile, and depositional environment. modified from hintze and kowallis (2009). 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 figure 5. stratigraphic relations of the mancos shale, its subunits, and adjacent formations in the uinta basin. after birgenheier and others (2015). and washboard units, that are within the tununk member (figure 6). the upper ferron is a series of major cliff-forming sandstone units described by ryer (1981, 1991, 2004), gardner (1993), barton (1994), barton and others (2004), garrison and van den bergh (2004), anderson and ryer (2004), and many other workers. they have been referred to as “delta-front units” (ryer, 1981), “genetic units” (gardner, 1993), and “stratigraphic cycles” (barton, 1994). anderson and ryer (2004) defined these units as parasequence sets based on genetically related parasequences within each set, bounded by major flooding surfaces, and they paired the parasequence sets to major coal zones (note: a parasequence is a small-scale, genetically related succession of bedsets bounded by marine flooding surfaces or their correlative surfaces on top and at the bottom [van wagoner and others, 1988]). the ferron parasequence sets create stacks of seaward-stepping, vertically stacked, and landward-stepping packages of rocks, each generally composed of sandstone, siltstone, mudstone, and coal intervals (gardner and others, 2004). there are nine parasequence sets in the ferron, which anderson and ryer (2004) referred to as kflast chance through kf-8: “kf ” for cretaceous ferron sandstone, followed by the first, second, third, etc. (kf-lc, kf-1, kf-2…) (figures 6 and 7); we use this nomenclature. the parasequences in each set include an abbreviation for the location followed by the letter “a” for the lowermost unit (e.g., kf-2-iv-a, kf-2-iv-b, and kf-2-iv-c parasequences, where iv stands for ivie creek [figure 7]) (anderson and ryer, 2004). bedsets are recognizable on outcrop as distinct and mappable genetic units but unrelated to flooding surfaces (van wagoner and others, 1990; anderson and ryer, 2004). the nomenclature used for the bedsets in the ivie creek area are the kf-1-iv[a] and kf-1-iv[c] (anderson and ryer, 2004). lithology the ferron sandstone consists of yellow-gray, light-brown, to white sandstone and siltstone, gray sandy to black carbonaceous shale, and coal. sandstone (quartz arenites to quartzo-feldspathic arenites) is very fine to coarse grained, poor to moderately sorted, subrounded to angular, and cemented with calcite, dolomite, or iron oxide (jarrard and others, 2004). sedimentary structures in the sandstone that determine facies include ripples, channel scours, soft-sediment deformation, cross-stratification (trough, swaley, and hummocky), and planar beds. many beds contain rooted zones, plant fossils, and a large variety of trace fossils from burrows to dinosaur tracks. fauna such as bivalves and gastropods are common; some are so enriched with shells that the rocks are considered coquinas. bedding in sandstone is thin or lenticular to massive, forming vertical cliffs, whereas siltstone, shale, and coal create recesses and slopes. coal beds are persistent but lenticular and commonly burned (due to spontaneous combustion or lightning strikes), creating “clinker” zones with red oxidation staining. paleogeography and facies the ferron sandstone was deposited in fluvial-dominated and wave-modified deltaic environments that prograded into the cretaceous interior seaway (figure 8). sediment was transported by eastto northeast-flowing rivers and streams that originated in the nearby sevier orogenic belt. the alluvial to lower delta or coastal plain of the ferron contained meandering streams (creating singleand multi-storied complexes), swamps and peat bogs, distributary channels, levees, crevasse spays/overbanks, and bays. deltaic facies varied and were primarily controlled by sediment input and accommodation space. wave energy at the coastline influenced the redistribution of these sediments and fluctuations in sea level had profound effects on the accumulating sediment package. sediment supply was high during early ferron deposition resulting in seaward-prograding fluvial-dominated conditions that t.c. chidsey, jr., and p.b. anderson ancient delta deposits in ivie creek area, ferron sandstone 7 figure 6. ferron sandstone parasequence sets. there are nine parasequence sets in the ferron referred to as kf-last chance through kf-8: “kf” for cretaceous ferron sandstone, followed by the first, second, third, etc. (kf-lc, kf-1, kf-2…). the system of letter designations for major coal beds of lupton (1916) is still followed. the green area represents coastaland alluvial-plain deposits. after ryer (1991), anderson and others (1997), and anderson and ryer (2004). figure 7. ivie creek area ferron stratigraphy including parasequence sets, parasequences, and named bedsets. gray = marine shale, blue and gray-blue = marine siltstone, light orange and yellow = shoreface sandstone, green = bay-fill mudstone and siltstone, black = coal, grass green = coastal and alluvial plain siltstone and mudstone, tan = fluvial sandstone. after anderson and others (2004). figure 8. paleogeographic map of utah about 85 ma. modified from utah geological survey (1998). 8 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 created lobate deltas consisting of delta-front deposits, distributary channels, splay complexes, and interdistributary bays. later, conditions changed at ivie creek to wave-dominated or wave-modified deltaic deposition forming cuspate deltas that grade laterally into strandplains and barrier islands. facies associated with wave-dominated or wave-modified deltas consist of prodelta deposits; lower, middle, and upper shoreface; foreshore; distributary channels; and distributary-mouth bars. the strandplains and barrier island facies include washover fans, lagoons, bays, and tidal inlets and associated flood-tidal deltas (ryer, 1991; ryer and anderson, 2004). elongate silty sand bodies, such as the clawson and washboard units (figure 6), represent offshore bars produced by longshore drift or sand plumes/turbidites that accumulated on the shallow-marine shelf to the east. landslides and slumps landslides and slumps can occur when erosion oversteepens slopes and undercuts resistant bedrock. the ferron sandstone is highly jointed and underlain by shale beds of the tununk member of the mancos shale and thus very susceptible to rockfalls (figures 2 and 3). the tununk contains significant amounts of swelling clay, particularly bentonite/smectite in volcanic ash beds. these unstable soft units cause joints in the overlying ferron to further enlarge. large pleistocene-(?)-age (1,800,000 to 12,000 years ago) landslides can be observed along the south-facing slopes of the ferron escarpment, east of quitchupah creek along the i-70 route (figure 2). emery coalfield the emery coalfield is located along the west flank of the san rafael swell from the limestone and coal cliffs continuing into the subsurface under castle valley and wasatch plateau (figure 1). the coalfield consists of 13 coal beds contained in the ferron sandstone that are given letter designations from a to m in ascending order based on lupton’s (1916) ferron description system (figure 6); this system is still used by geologists working the ferron over 100 years later. ferron coal beds are lenticular, which has limited their commercial development. seven seams exceed 4 feet (1.2 m) in thickness; the maximum coal seam thickness is 30 feet (9 m) (doelling, 1972). however, several seams coalesce to form a single, thicker coal with at least one at 25 feet (8 m) thick (doelling, 1972, 1976). coal beds a (exposed in the ivie creek area), c, i, and m are the most important (figure 6), and several are stratigraphically close to each other. analyses of these beds indicate an overall high-volatile c bituminous coal (doelling, 1972, 1976). estimated available, inplace coal resources for the southern emery coalfield are 2.2 billion short tons, based on coal beds greater than 4 feet (1.2 m) thick and less than 3000 feet (900 m) deep, of which about 500 million short tons are recoverable (quick and others, 2004). presently, one coal mine is extracting coal from the coalfield from the i seam. sequence stratigraphy, facies, and paleogeography of the ferron sandstone in the ivie creek area kf-1 parasequence set the kf-1 parasequence set consists of river-dominated deltaic deposits that prograded from the south-southeast to the north-northwest, proximal to distal, as delta lobes across the ivie creek area. progradation was parallel or onshore to the northwest-southeast regional shoreline trend. the ivie creek area is best known for its distinctive, steeply inclined bedsets or clinoforms dipping west-northwest 10° to 15° (figures 3 and 9). these unique sedimentary features have been the subject of major academic and industry studies and are incorporated in hydrocarbon reservoir simulation models (anderson and others, 1997, 2004; mattson, 1997; forster and others, 2004; jarrard and others, 2004; mattson and chan, 2004; enge and howell, 2010; deveugle and others, 2011; graham and others, 2013, 2015a, 2015b). the kf-1 parasequence set lies conformably upon the tununk member. in the ivie creek area the kf-1 parasequence set consists of only one parasequence, kf-1-ivie creek (iv), with two bedsets referred to here as kf-1-ivie creek[a] and kf-1-ivie creek[c] (figures 3 and 7), that represent the two delta lobes sourced from a common point (anderson and ryer, 2004). kf-1-iv[a] bedset the kf-1-iv[a] bedset can be traced on outcrop from landward to seaward pinchout for about 1.5 miles (2.4 km). delta-front sandstones of a modified gilbert delta make up these deposits, which change from proximal to distal as they pass from east to west across the ivie creek area. this bedset is a fan-like deposit that thickens to the east and was deposited into an area having minimal wave influence; therefore, the primary beds are preserved as clinoforms (figures 3 and 9). the clinoforms change characteristics from the shallower water, more proximal locations, to the deeper water, more distal locations. they thin rapidly to the west, as well as thin to the more typical seaward direction of north. this is anomalous for ferron delta lobes. the clinoforms are classified into four facies: proximal, medial, distal, and cap (figures 10 through 14) (anderson and others, 1997, 2004; mattson, 1997; mattson and chan, 2004). clinoform facies are based on grain size, sedimentary structures, bedding thickness, inclination angle, and stratigraphic position. • clinoform proximal facies is sandstone, mostly fine to medium grained. the chief sedimentary structure is low-angle cross-stratification with minor horizontal and trough cross-stratification and rare hummocky bedding (figure 11). this facies is dominantly thickto medium-bedded, well to moderately indurated. t.c. chidsey, jr., and p.b. anderson ancient delta deposits in ivie creek area, ferron sandstone 9 clinoform proximal facies is interpreted to be the highest energy and most proximal to the sediment supply. the steep inclinations are interpreted to represent deposition into a relatively localized deep portion of an open bay environment. • clinoform medial facies is dominantly sandstone with about 5% shale (figure 12). the sandstone is primarily fine grained with slightly more fineto very fine grains than fine to medium grains. horizontal beds dominate with some rippled, trough, and low-angle cross-stratified beds (figure 13). bed thickness ranges from laminated to very thick, but most beds are medium. clinoform medial facies is generally transitional between end members clinoform proximal and distal. • clinoform distal facies is sandstone (sometimes silty) and about 10% shale (figure 14). the sandstone grain size is dominantly fine to very fine grained, with considerable variation. sedimentary structures in clinoform distal facies are chiefly horizontal laminations and ripples in medium to thin beds (figure 13). clinoform distal facies is gradational with clinoform medial facies and represents the deepest water and lowest energy deposits within the clinoform. it can be traced distally into prodelta to offshore facies. • clinoform cap facies is sandstone, generally very fine to fine grained. the beds are horizontal, with some trough and low-angle cross-stratification in thick to medium beds (figure 11). the clinoform cap facies is interpreted to represent an eroded and reworked delta top. the deposits of the kf-1-iv[a] bedset accumulated on an arcuate delta lobe that prograded into a deeper-water, fully marine bay. the main delta, located to the east and northeast, created a protected embayment in the northwest part of the ivie creek area (anderson and others, 1997). the kf-1 clinoforms represent deposition into the embayment fed by river channels from the southeast (figure 15a) (anderson and others, 2004). the distributary complexes/delta front, shallow marine, and deep marine environments produced the clinoform proximal, medial, and distal facies, respectively. kf-1-iv[c] bedset the kf-1-iv[c] bedset is sand rich and varies in thickness within the ivie creek area. in contrast to the kf-1-iv[a], delta-front, subtle clinoforms of the lower part of the bedset dip less than 5°. this sand-rich bedset thins in both up-depositional-dip and down-depositional-dip directions due to lateral facies changes. a major flooding surface is at the top of the sub-a coal zone (figure 3 and 16) and corresponds to the boundary with the overlying kf-2 parasequence set (figures 6 and 7). the lower section of the kf-1-iv[c] laps onto the more distal parts of the kf-1-iv[a] in the western part of the area and represents the distal part of another delta lobe, probably sourced from the southwest. this lower section may be completely absent in some locations as a result of erosion and/or non-deposition. it also includes a distributary channel sandstone body, lenticular in cross section, deposited by a northwesterly flowing stream (figures 9b and 9c). figure 9. close-up photos of depositional features and stratigraphy of the ferron sandstone in the ivie creek area. a – clinoforms in the kf-1defined by pronounced bounding surfaces (see white box on figure 3 for location of figure 9a). b – distributary channel cut into clinoforms within the kf-1 parasequence set. the kf-2 parasequence set above shows relatively thin bedding. c – same photo as b annotated with kf-1 bedsets and kf-2 parasequences (see figure 7). modified from anderson and others (2004). photographs by michael chidsey, sqwak productions inc. a b c 10 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 figure 10. scaled cross section, oriented west to east, of the kf-1-iv[a] bedset from the ivie creek amphitheater based on a portion of interpreted photomosaics showing clinoform facies distribution and bounding layer measurement locations (modified from mattson, 1997; for data and detailed descriptions see anderson and others, 2004). figure 3, kf-1 captures approximately the same area as the cross section. the artificially abrupt end of the medial facies is representative of most of the contacts portrayed between proximal, medial, and distal clinoform facies in this study because polygonal packages of rock are necessary for reservoir modeling and simulation. in reality, the transition from medial to distal (as well as medial to proximal) is gradational. figure 11. kf-1-iv[a] bedset in the ivie creek amphitheater. a – clinoform cap facies with trough cross-bedding exposed at the top of the photo. the head of the hammer is within the cap facies above a hard, reddish bed within the underlying clinoform proximal facies. b – low-angle cross-beds are the primary sedimentary structure in the clinoform proximal facies. the general direction of the depositional dip is from right to left. note the gentle up depositional dip inclinations of many laminae in the bed below the hammer. holes in the rock are from outcrop permeability plugs taken in the unit. c – close-up of low-angle trough cross-bedding in the clinoform proximal facies. figure 12. clinoform medial facies in the kf-1-iv[a] bedset. the facies is dominated by horizontal bedding as found in the thicker bedded units. the recessive parts of the outcrop are typically ripple laminated and finer grained than the resistant sandstone beds. figure 13. clinoform medial/distal facies showing close-up of (a) thin horizontal bedding and (b) ripple laminations. a b c a b t.c. chidsey, jr., and p.b. anderson ancient delta deposits in ivie creek area, ferron sandstone 11 the uppermost section of the kf-1-iv[c] is continuous across the entire ivie creek area and has wave and fluvial influences (figure 3). it is capped by unidirectional, trough cross-bedded sandstone. the kf-1-iv[c] contains slump or rotated block features near the mouth of ivie creek canyon. an excellent study of these features is presented in braathen and others (2018). in the ivie creek amphitheater, above the basal cross-bedded sandstone, are 10 to 15 feet (3–5 m) of coarseningand bed-thickening-upward, brackish-water/bay-fill deposits consisting of carbonaceous mudstone; thin, rippled to horizontally laminated, sparsely to intensely burrowed sandstone and siltstone; fossiliferous mudstone to sandstone; oyster coquina; and ash-rich coal (figures 16 and 17). the presence of a brackish-water fauna (crassostrea, corbula securis, lucinid, caryo corbulais, aff. varia, and serrthid) is the distinguishing characteristic of this facies. this fauna is present within “dirty” sandstone, siltstone, mudstone, and carbonaceous mudstone. oyster coquinas commonly are found within the brackish-water/bay facies (figure 17b). typically, the rocks are rich in carbonaceous debris. the uppermost, carbonaceous mudstone or ash-rich coal is the sub-a coal zone (figures 3 and 16). the coal zone (swamp [paludal] facies) shows considerable variation in thickness (0 to 1 foot [0–0.3 m]) and intertongues with underlying bay deposits. the paludal facies is commonly underlain and/or overlain by coastal plain facies (the top is frequently rooted) and may grade into shoreface deposits below. at the mouth of ivie creek canyon, the kf-1-iv[c] bedset is anomalously thick due to large slump features or rotated blocks that formed shortly after deposition. failure of the rotated blocks is consistently toward the north to northwest, the direction that the delta lobe appears to have prograded. kf-2 parasequence set the ferron sediment source switched from the east-southeast to the west going from kf-1 to kf-2 deposition. the kf-2 parasequence set represents wave-modified deltaic deposits that generally coarsen east to west, consisting of shoreface and distributary complex facies gently inclined (< 3°) to the east (figure 3). these relatively clean, sand-rich deposits accumulated along a local north-south shoreline trend defined by a landward pinchout of marine shoreface facies (observed in the i-70 roadcut just to the west of the ivie creek area and farther east in ivie creek), as opposed to the more common regional northwest-southeast shoreline trend recognized in other ferron cycles above and below kf-2. the kf-2 parasequence set contains three parasequences: the kf-2-iv-a, kf-2-iv-b, and kf-2-iv-c (figures 7, 9b, and 9c). these parasequences show less lateral variation in lithofacies than the kf-1 bedsets due to greater wave influence (reworking). within the ivie creek area, there is also little lateral variation in thickness of sand-rich kf-2 facies, even when lateral change occurs from one depositional subfacies to another. kf-2 cycles accumulated in sheet-like bodies that pinchout laterally, forming a wedge due to wave-action along the delta front. the contact between kf-1 and kf-2 is generally drawn at the top of the sub-a coal zone in the ivie creek area (figure 3). the top of kf-2 lies near the top of the c coal and includes all of the a coal zone and delta-plain strata which separate the a and c coal zones (figure 6). in general, the western part of the ivie creek area, eastto northeast-flowing distributary channels deposited large amounts of sand in north-south-trending distributary-mouth bars. shallowto moderate-depth marine conditions existed in the eastern part of the area. an uncommon transition from shoreface, to bay, to coastal plain/ swamp occurred during the late stage of kf-2 deposition (anderson and others, 2004). above the kf-2 parasequence set are coastal-plain and alluvial-plain facies which represent the landward equivalents of the marine portions of kf-3 through kf-7 parasequence sets. the wave-modified parasequences of the kf-2 parasequence set consist predominantly of shoreface, distributary complex, and coastal plain/swamp facies. shoreface the shoreface facies consists of a relatively steeply dipping zone (compared to the shelf/slope) from the subaerial beach to a poorly defined point where the slope flattens on the sea floor. wave energy is sufficient to move sand-size grains in this zone. at the seaward end of the shoreface is the prodelta facies. this mud-dominated facies represents the area just seaward of the dominant influence of wave energy and typically interfingers with the lower shoreface defined below. figure 14. clinoform distal facies in the kf-1-iv[a] bedset. the base of the ferron sandstone is exposed at the base of the interbedded sandstone, siltstone, and mudstone. ripple and planar bedding are common; burrowing is sparse. 12 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 figure 15. paleogeographic interpretations of the ferron sandstone in the ivie creek area. a – kf-1-iv[a] bedset showing river channels flowing from the south and southeast that deposited sands into a protected embayment in the northwest part of the area. b – kf-2-iv-a parasequence showing a large distributary-mouth bar indicating the shoreline was probably not too far to the west and the shoreline was now primarily oriented north-south. c – kf-2-iv-b parasequence showing the continuation of the north-south-oriented shoreline where the landward area is dominated by distributary-mouth bar and distributary channels associated with a wave-modified delta which gives way seaward to shallow marine facies. d – kf-2-iv-c parasequence showing the north-south-oriented shoreline now farther east and a large bay with associated distributary channels and bay-head deltas. late stages of this parasequence (not shown) include the development of large meander belt and channel system cut into coastal-plain deposits at the top of the kf-2-iv-c. see blue box on figure 2 for location of these maps. modified from anderson and others (2004). a b c d t.c. chidsey, jr., and p.b. anderson ancient delta deposits in ivie creek area, ferron sandstone 13 the lower shoreface facies consists of thinly interbedded shale to siltstone and very fine to fine-grained sandstone having wave ripples to horizontal laminations; hummocky stratification is common (figure 16). burrowing is generally found on the top of thin sandstone beds and the shale is often bioturbated. middle shoreface facies consists of very fine to fine-grained sandstone composed of hummocky, swaley, and planar laminations with minor ripple laminations (figure 18). this facies is generally thick, representing 50% or more of the shoreface sequence. the most common burrow types are thalassinoides and ophiomorpha with numerous other shallow marine traces and the amount of burrowing varies from moderately to intensely bioturbated. upper shoreface facies is characterized by fineto medium-grained, multidirectional to bimodal, cross-stratified sandstone, in sets that are occasionally separated by planar laminations, and that is generally moderately to well sorted (figure 19). this facies is about 10 feet (3 m) thick, but greater in the vicinities of the landward pinchouts of the parasequences where the upper shoreface may reach 20 feet (6 m) in thickness. the upper shoreface facies is slightly to moderately burrowed and ophiomorpha is the most common trace fossil. the foreshore facies consists of fineto medium-grained sandstone with planar to inclined bedding, slightly to intensely burrowing, and is sometimes rooted (figure 19). this facies is sometimes absent at the top of the shoreface sequence due to erosion, but when present ranges up to a few feet in thickness. distributary complex the distributary complex facies is subdivided into distributary channel and distributary mouth-bar in the ivie creek area (figures 9b and 9c shows an example in the kf-1). this facies is commonly characterized by the complicated geometry of cross-bedding and large-scale bounding surfaces that are related to cut-and-fill processes, in contrast to the flat to very gently inclined surfaces of the lower delta-front. figure 16. lower shoreface facies found above the kf-1/kf-2 parasequence set boundary in the ivie creek area. typically, the facies consists of interbedded very fine grained sandstone, siltstone, and shale, and grain size coarsens and beds thicken upward. this is stratigraphically the lowest facies in the wave-modified delta deposits in the area. figure 17. brackish water/bay fill facies in the kf-2-iv-c in the ivie creek amphitheater. a – carbonaceous mudstone, siltstone, and sandstone, intensely burrowed. b – oyster coquina containing brackish-water fauna within carbonaceous mudstone. core from ugs drill hole ivie creek no. 11 (see figure 15 for drill-hole location). a b 14 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 the distributary channel facies is common in river-dominated delta-front deposits and is also found within the wave-modified shoreline deposits of the ivie creek area (figures 9b and 9c). it is characterized by channels with high height-to-width ratios and unidirectional trough cross-stratified and current ripple cross-laminated deposits. channel fills are sandstone-dominated, but heterolithic channel fills are also found. in sand-filled channels, the grain size is commonly coarser than the surrounding delta-front or shoreface and fines upward. troughs in the channel base generally contain mud rip-up clasts, woody fragments, and rare shark teeth. this facies grades seaward into the distributary-mouth bar facies. the distributary-mouth bar facies is found in the upper parts of delta-front sequences of kf-2 and is associated with distributary channel deposits. this facies is characterized by fine-grained, or coarser, trough-cross-stratified sandstone, with moderate to intense burrowing in areas that had lower flow velocities and decreased sedimentation rate; some intervals between trough sets are completely bioturbated; ophiomorpha is common. paleoflow directions show a strong offshore component with the amount of scatter increasing with increased wave influence and distance from the distributary channel. traced laterally and seaward, the facies commonly grades into middle shoreface or lower delta-front facies (fluvial-dominated shoreline). coastal plain/swamp the coastal plain/swamp facies is represented by a sequence of non-marine rocks dominated by mudstone, carbonaceous mudstone, and siltstone with minor sandstone. coal is commonly interstratified with the other rock types. these rocks are interpreted as inter-fluvial environments along a low-gradient coast. kf-2-iv-a parasequence the oldest parasequence of the kf-2 parasequence set in the ivie creek area is the kf-2-iv-a. the kf-2-iv-a thickens to 50 feet (15 m) westward across the area. the base of the kf-2-iv-a parasequence consists of interbedded sand and minor shale deposited in prodelta to lower-shoreface environments (figures 9c and 16). these deposits are thin, typically less than 10 feet (3 m) thick. the figure 18. middle shoreface facies in parasequence kf-2-iv-a usually forms a vertical cliff and consists of very fine to fine-grained sandstone that is horizontally bedded and often moderately burrowed to bioturbated. note the horizontal bedding indicated by the light horizontal lineations in the photograph. figure 19. upper shoreface and foreshore facies in parasequence kf-2-iv-c, exposed in ivie creek canyon. note the cross-beds in the base of the upper shoreface, typical of this facies. ophiomorpha is common in both facies. the foreshore bedding is horizontal to sometimes slightly inclined. inset shows vertical rootlets penetrating the facies and originating from the overlying a coal. t.c. chidsey, jr., and p.b. anderson ancient delta deposits in ivie creek area, ferron sandstone 15 prodelta and lower-shoreface deposits are overlain by a 0.5to 1-foot-thick (0.2–0.3 m) zone of highly carbonaceous to coaly sandstone which grades into 20 to 30 feet (6–9 m) of very fine grained, silty, and slightly carbonaceous sandstone representing a middle shoreface environment (figure 18). the middle-shoreface unit is moderately to intensely bioturbated. near the mouth of ivie creek canyon, the kf-2-iv-a parasequence consists of a thin sequence of lower shoreface heterolithics overlain by about 28 feet (8 m) of middle-shoreface deposits (figure 18). the westernmost exposure of kf-2-iv-a is in ivie creek canyon where the facies is a distributary-mouth bar. this indicates that the shoreline was probably not too far to the west. utah geological survey (ugs) drill hole ivie creek no. 11 contained distributary-mouth bar facies of the unit (figure 15b), but interpretation of drill holes farther west indicate the shoreline has been crossed and only non-marine deposits are present. the shoreline orientation of the kf-2-iv-a parasequence was primarily northsouth (figure 15b). possible foreshore deposits are near the last outcrop of the kf-2-iv-a in the bottom of ivie creek. kf-2-iv-b parasequence in the ivie creek amphitheater, the kf-2-iv-b parasequence was deposited in a middleshoreface environment. kf-2-iv-b exhibits gently seaward-inclined beds that are very conspicuous when viewed from west to east along the outcrop. the kf-2-iv-b parasequence consists of horizontally bedded, silty sandstone at the base (middle shoreface) (figure 9c) and unidirectional, trough cross-bedded sandstone at the top (distributary channel to mouthbar deposits). these units are moderately to intensely bioturbated. along ivie creek canyon, kf-2-iv-b is dominantly unidirectional, trough cross-bedded sandstone of a mouth-bar complex which continues westward up the canyon and is present in the subsurface at ugs drill hole ivie creek no. 10 to the northwest (figure 15c). the unit thickens slightly from west to east, in contrast to the underlying parasequence. like the preceding unit, the shoreline orientation during deposition of the kf-2-iv-b parasequence was generally north-south. in the landward exposures at ivie creek canyon, the unit is dominated by distributary-mouth bar and distributary channels associated with wave-modified delta deposits (figure 15c). farther seaward, the distributary-mouth bar deposits give way to middle-shoreface deposits. seaward (east) of the ivie creek area, a well-developed lower shoreface is present at the base of the unit. paleoflow of several of the distributary channels indicates a general northerly trend in the local area. at the i-70 road cut, a highway drainage channel was cut through solid rock to accommodate storm runoff. this cut provides superb views of kf-2-iv-b mouth-bar deposits. from the downstream end of the drainage channel, cross-bedded, mouth-bar deposits can be visually followed into distal bar or middle shoreface deposits to the east. kf-2-iv-c parasequence the kf-2-iv-c parasequence is separated from the underlying kf2-iv-b parasequence by a siltstone to shale interval that varies in thickness across the east part of the ivie creek area. generally, the entire parasequence fines from west to east. in the ivie creek amphitheater, kf-2-iv-c is interpreted as a bay-fill deposit (although it is devoid of body fossils). there is evidence for bay-head deltas and tidal channels feeding the bay to the northeast in quitchupah canyon (figures 2 and 15d). at the top of this sequence is a thin, medium-grained carbonaceous sandstone, which may represent the migration of a low-energy beach (foreshore deposits) across the bay fill prior to capping by coastal-plain deposits and deposition of the overlying a-coal zone (which is locally burned). in ivie creek canyon, kf-2-iv-c forms a 10-foot (3 m) cliff having excellent upper shoreface facies exposed (figure 19). the top of the unit is rooted by the overlying coastal-plain vegetation. the landward pinchout of the marine facies of kf-2-iv-c trends just slightly east of south toward i-70. at the mouth of ivie creek canyon, shoreline sandstone changes over a distance of about 300 feet (90 m) from a strongly wave-modified shoreface unit to a much lower energy unit that contains mud interbeds and finer sand and has a silvery gray color in outcrop, but is carbonaceous on a fresh surface. this suggests a change from a coast directly facing the sea to one that was sheltered from wave energy. the environment of this shoreline unit is a wave-modified coast, probably shoreface facies in the proximal part, transforming laterally to low-wave-energy bay facies. a large meanderbelt channel system cuts into the top of the kf2-iv-c parasequence just south of the immediate ivie creek area. the channel system flowed to the northeast, but much of this ferron channel system has been removed by erosion. a late-stage episode of lateral channel migration across the top of the kf-2-iv-c delta-plain deposits is recorded by scours into the meanderbelt deposits. the best example of this channel system is exposed on the south side of i-70 across from the ivie creek amphitheater, just east of the road cut through the ferron. this same channel is exposed in the road cut on the north side of i-70, but the channel orientation and the nature of its exposure in the cut face do not present the classic channel shape exhibited in the southern exposure. the meanderbelt and younger channel systems fed the continued progradation of kf-2-iv-c and stratigraphic equivalents to the east and northeast. 16 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 acknowledgments support for this paper was provided by the utah geological survey (ugs). cheryl gustin, jay hill, and martha jensen of the ugs drafted figures. this paper was carefully reviewed by michael d. vanden berg, stephanie m. carney, michael d. hylland, and bill keach of the ugs, along with the editors of this publication. their suggestions and constructive criticism greatly improved the manuscript. references anderson, p.b., chidsey, t.c., jr., and ryer, t.a., 1997, fluvial-deltaic sedimentation and stratigraphy of the ferron sandstone, in link, p.k., and kowallis, b.j., editors, mesozoic 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p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 wagoner, j.c., ross, c.a., and kendall, c.g., editors, sea-level changes—an integrated approach: society for sedimentary geology (sepm) special publication 42, p. 39–46. van wagoner, j.c., mitchum, r.m., campion, k.m., and rahmanian, v.d., 1990, siliciclastic sequence stratigraphy in welllogs, cores, and outcrop: american association of petroleum geologists methods in exploration series, v. 7, 55 p. weiss, m.p., witkind, i.j., and cashion, w.b., 1990, geologic map of the price 30' x 60' quadrangle, carbon, duchesne, uintah, utah, and wasatch counties, utah: u.s. geological survey miscellaneous investigations series map i-1981, scale 1:250,000. witkind, i.j., 1979, reconnaissance geologic map of the wellington quadrangle, carbon county, utah: u.s. geological survey miscellaneous investigations series map i-1178, scale 1:24,000. witkind, i.j., 1980, reconnaissance geologic map of the mounds quadrangle, carbon and emery counties, utah: u.s. geological survey miscellaneous investigations series map i-1202, scale 1:24,000. witkind, i.j., 1988, geologic map of the huntington 30' x 60' quadrangle, carbon, emery, grand, and uintah counties, utah: u.s. geological survey miscellaneous investigations series map i-1764, scale 1:100,000. wood, r.e., and chidsey, t.c., jr., 2015, oil and gas fields map of utah: utah geological survey circular 119, scale 1:700,000. uga-geosite-biek-sevier-fault.indd utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors robert f. biek utah geological survey, p.o. box 146100, salt lake city, ut 84114-6100 bobbiek@utah.gov cover image: view north-northeast across utah highway 12 to the sevier fault and black mountain at the entrance to red canyon. sevier fault at red canyon 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e geosites refl ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: biek, r.f., 2019, sevier fault at red canyon, in milligan, m., biek, r.f., inkenbrandt, p., and nielsen, p., editors, utah geosites: utah geological association publication 48, 6 p., https://doi.org/10.31711/geosites.v1i1.53. presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president r.f. biek sevier fault at red canyon 3 introduction the sevier fault is spectacularly displayed on the north side of utah highway 12 at the entrance to red canyon, where it offsets a 500,000-year-old basaltic lava flow. the fault is one of several active, major faults that break apart the western margin of the colorado plateau in southwestern utah. the sevier fault is a “normal” fault, a type of fault that forms during extension of the earth’s crust, where one side of the fault moves down relative to the other side. in this case, the down-dropped side (the hanging wall) is west of the fault; the upthrown side (the footwall) lies to the east. the contrasting colors of rocks across the fault make the fault stand out in vivid detail (figure 1). immediately south of red canyon, the 5-million-year-old rock canyon lava flow, which erupted on the eastern slope of the markagunt plateau, flowed eastward and crossed the fault (which at the time juxtaposed non-resistant fan alluvium against coarse-grained volcaniclastic deposits) (figure 2) (biek and others, 2015). the flow is now offset 775 to 1130 feet (235-345 m) along the main strand of the fault, yielding an anomalously low vertical slip rate of about 0.05 mm/yr (lund and others, 2008). however, this eastern branch of the sevier fault accounts for only part of the total displacement on the fault zone. a concealed, down-to-the-west fault is present west of coarse-grained volcaniclastic strata at the base of the claron cliffs. seismic reflection data indicate that the total displacement on the fault zone in this area is about 3000 feet (900 m) (lundin, 1987, 1989; davis, 1999). location park at the scenic overlook and rest stop 2.6 miles (4.2 km) east of the junction of utah highway 12 and u.s. highway 89. coordinates are 37° 44' 38", 112° 19' 46". picnic tables and pit toilets are available. figures 3 and 4 show a geologic map and cross section of this area. several scenic pullouts and a paved bicycle trail make red canyon a popular spot for visitors on their way to nearby bryce canyon national park. a u.s. forest service visitor center with restroom facilities is about one mile (1.6 km) east up red canyon. figure 1. view north-northeast across utah highway 12 to the sevier fault and black mountain at the entrance to red canyon. the 500,000-year-old red canyon lava flow (qbrc) on the north side of the canyon erupted from vents on the down-dropped side of the sevier fault (red line, d = down, u = up) and flowed eastward across the fault, which then did not form a significant topographic barrier as it does today (at the time the lava flow erupted, the fault juxtaposed tilted, non-resistant late tertiary basin-fill deposits in the hanging wall against non-resistant, basal brian head strata in the footwall). the lava flow has since been displaced about 650 feet (200 m), yielding a vertical slip rate of 0.4 mm/yr (lund and others, 2008). the trace of the fault is largely concealed by talus and landslides. tvf, late tertiary basin-fill deposits; tbhv, basal brian head strata; tbm, conglomerate at boat mesa; tcwl, lower unit of the white member of the claron formation; tcp, pink member of the claron formation. figure 2. view south along the sevier fault zone just south of red canyon. pinyon-juniper-covered hill between the two fault strands (u = upthrown side, d = down-dropped side) is mostly coarse-grained volcaniclastic strata, locally capped by the 5-million-year-old rock canyon lava flow, a remnant of which is preserved just out of view to left in this photo. wilson peak is at upper left. 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association 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qms? qhd qms qac qacfo qc kwcg qms(ti) qafo2 kw qmsc tbhqam qc kt ksjs qgtou qac kwcs qms qblf tcp tql qc qacf qms qacf qhd tbhm qms kwcs km ti tcwm qaly qat qgtu qacf qgtou tcwl qafo kw tql tcw tcp tcwm kw qms tbh? qacfo qac qc tcwt tcwm ti qafy tcwl tcwttcwu ksd tbh qblf kwcs qms tbh qacfo tcwm jcc qmtc kwcs ti qac ktu ksjc kt tcwl jcx qms? qbpl3 tcwm tbrc qaf1 tcwm qms?qmsh tcwm tcwu qacfo qaf1 qac tcwu qms ti qaly qafo tcwl qms kwcs kw qc qc tbh tcwm jcp qbmk3 qafo tbh qblfc kk kw qaly qms? tcwl kt qal qmtc ksj tm(ti) qac jcx qc qmtc ksd qms tm(ti) qc qafy jcw qalo qac tcp kwcg qafy qmt qbpl3 qaly qaly ksj tcp qmt qmsh ktu qc ksjs qms qc kwcs qc ksst qms tbbm taf tcwl qacf qafy tbh tcp qms ksd tcp qac qmt ksst kwcs qafo kwcs qmt tcwu ksjc ksj qms qafy qmsc qaly qmtc qmsqhd qafo2 qlao/qbmk3 tcw kkl ti qmtc qms qafy qac qac ti tcw? qalo qms kw tbh tbh qtbx tvg qafo qap qmtco tbh qmtc tm(ti) qbpl2 qap kkltm(tnw) qmt qacfqgop qbck qap ksd kwcs ksj qmsc kw qmt qac qacfo qc ksst kwcg qaly kkl qms? qat qmtc ti qafy qmtc kw qam kkl qc kwcs qmtc qms qtap ti qmtc qms tql qms tbrc qac qms qacf qc/tbh qafo2 tbh kk tql qms ksj tql tcwm qc qap qc qgtuqms qcqc qac qc tcwu ksd qlao ksjs qmtc qaco qacf tc taf qat qms qc ksd tbsp qc qms(ti) tcp qac tm(tbh) qmt qc qmt qac ksst qhd km qms qac tql qms qtbx ti tm(ti) qms kwcs qc qmtc ksjc tbh tm(ti) qacfo ktd qaly tql qafo qc qc kw kwcs tbhtcwu qaly qms(ti) kw qaly qco qmt qaly tcw qaly qlatm(ti) qtap qc taf qmt qc qc qc kkl tbrc qbpl1 tcwt qacf qms ksjc qms(ti) tm(ti) qc qafotbbm qms qac qtap qms?qacfo qhd qms tbh ksj qc taf qmt ti ksj qaly qms qaly tcwt qms qap tvfqmtc qafy ksd tcwu qmt qc qac taf qafy qac qms(tql) qmt qac qalo tbh tcwu qactvgqacf qgtp tcwt qac tbh qc qc qafo2 qaly qafo tcwt qmtc qmt qc qbpl3 qalo ktd qc tm(tnw) qacqms(ti) qao kw qac tm(ti) kw qc qaly tm(ti) tafqafo qacf tvg qc tm qtap kst qms qms qmsc kk qms? qaco qmt tm(tdm) qbckc qmtc qms qc qafo2qacf tbsp qacf qao tbh qbck qaly qms? kwcs qaly kw qmtc qafo qmtcqaly qms qc qacf qac qac ksjs? qc tbh tcwt tm(ti) tcwl? qc tcp taf qc tbh qacf qafy tcp qafy qhd tm(ti) tql qaly qc qac qc ksj tm(tdm) tbh qms qc qao qacf qafy tm(tbh) qac jcp qafo2 tbh qms qac qc qms tbsp qms(tm) tbrc ti qafo tm ksjs qms? tbh tql qmt tm(ti)qaco ti qmsc qacfo tm qc jcw qafy tbrc qat qc kst qmtc qms tm(tdm) qac tm(tnw) qc qmtc qmt qms(ti) tm(ti) tcp qafy qap taf qmtc qacfo ksd qaf2 tm(tbh) tbsp tcwl qgtp tbh qacfo qc qbtp qms qafo qaly qms tbh tbrcc qmt ksd qac qafy ksd qcqms qms qc kk qms qafo qc qalo qmtc ta qalo tbrc tbh tvg qacf tbh qalo qac qac ksdtbh tm(tbh) tbh qac kw td qaly qmtc kwtaf tm qacf qmtc tql qafo qms qaly qmtc tcp qmscqms qaly tm(tnw) qmtc tql qms qc tm(ti) qmt qmtc qacf qaly tm(ti) qms qacfo tbh tcp tm(ti) qms(ti) qac tql? ktd qafo qms? ti? qc kwcs tm tbh? qmtc tbrc qms tm(ti) qc qc ksd qc kk kst qms kw qgtp qms qacf qaly qalo tbrc qaotm qc kgc qac tql qacqap qms qmstbh qacfo qms tcp kw qafo qaly qacf qc qmtc qms qacf qca qmt qms ksj tql qmsc qms tcw? tm qc qalo qc qac tm(tdm) qc kw tcp qap ti? qtap qacf tm(ti) qaf2tm(tbh) thma qms kk tbhm qafy qmtc qat tm(ti) kw qc tbh kwcs qaf2 qalyqmtc qacf qc/tcp qms qac qafo tcp? qac thm ksd qafy ksd qtap tbh qacf qacfo tbh tbh qap tm(ti) qmtc qalo qms qacf tm qaly qms qc tm(td) qaly qacf tm(ta) qmtc qac tm(td) tcp qafo tql qc tbh qafo tm(ta) tbrc qacf tm(td) tm(ti) tm? qalo qac qaly tbh tvg qafoqms qacf qaly qaly tm qacf qacf qmtc qms taf qaly qc qafo qaly qafotm qc qtap taf qat thm qaly qms qac qaf2 tbh tbrc?qc qmtc tvf tcp taf qmsc qmtc qafy ti qacf qmtcqafy qms tcp kw qac qacf qtap tm(tbv) tcp qap tm(ti) qaf2 qc/tcp qmtc ti qacf qms ksj qc tm(td) qaco kw kgc qacfo tm(tbh) qc tcwl qc qac qc ti qat qafo qmt thm qms qc qms tbh ksj qaco qafo tm(ta) tm(ti) tm(tbh) qafy qmt qc tm qmtc qafytbh ksd qacf qco qacf qms? qms? qacfo ti qaly qalo qc qc qaly qac qafy qaly qms? tcwl qtap qms qac qap tcp qms? taf tcwm tm(tqcb) kw qco tbh qmtc tm qaly qacf qc qmtc qac qtap qap tm(tbv) qac qafo qac qc qaf2 tm tm tcw tm(ti) ksd qbrc qc qc qmtc qafy qc qtap qc qc tcw tvf qafo qmt tcp tcw qap qafy tm qap qc kwcs qalo tcwl qafo qms qap qms qms qaly qafy kk qaf2 qaf2 kgc tm qc qce kw qms(ti) qafo qmtc kw tcwl qbrc qafy tm(ti) qaf2 qafo tcw qaco qaly qac qac ti tmf qtap qms qac kwcs qmtc tm(td) qms qtapksj qac tbhv tm qaly qmtc ksd qc tm qc ti qco qafo qafo qc tbrc? qco qtap kw tm(ti) qafo qalytbmqaf1 qmtc tm(ti) tm ksjs tcw tm(td) ksj taf qc qafo qac tm tcw qc tm tcp qacf qms? tcwl tql tm? taf qms?qtap tbdh qms? thm qac tm(tbv) kst tcwl tbm kw ksjs qaly tm qac tm(ti) tm qtap qaf2 qaly qtap qms tm(td) tcw tm(ti) qcqap tbrc? qafy ti tm qafo qmtc qat tm(tbv) qms qacf tcw tm ta qac qms qap qc tcp tcw qac qmtc tm(ti) qafy qaco qmtc qalo taf qhdqacqms qmsc qmtc tbdh tcp? qacqat ksd qacf ti qbrc tm(ti) qac qac qacf qc ksj ti tm tcwm qaly kw kst qac qaly qac qacf qafo qctbm tm qms kw kwcs qms qc kgc tm(td) qacfo tcp tbm tbrc? qtap qafo tqcb ti qbe td qacf qaf1 qac tcpqacf tm(ti) qmtc qms qac qacf qtaptbh qalo qc tm(ti) qtap tqh tbhv qc tcp taf tql tcwm tbm qtap tql kgc qaf2 tcwu qc qc qmt qc qms qacfo tm(tbv) qacf td tbm qmt qc qc tm qms qms tbhv qaly ksdqms qaf2 taf qaly qac qat qms ti tbhv qc qmtc qap kwcs qmt tcw qafy tm(ti) qacf qc qacf qac tm(ti) ti tbrc? ksd tm(tbv) tql tm(tdbv) qmt tvf qmtc tcp qaf2 qaf2 tqcb kgc qaf2 tm(ti) ksd qc qafy tmf(td) tcw qtap tbh qms kwcs tbh qms qafy qbe qc ti qactcp qmtc thm qc taf qc qc qafy ksj qacf qaly qacf qc tm(td) qac tql ti td qbc qbs tbhv qc qaly qmtc ksd ti qac qms tbm tm qafo qap qmtc tql? qtlf qacf tqh qc tm(tbv) qafy qac qbsc tm(ti) tbhv qc qmt qac qafo tbh? qaly qmtc qacfo tm ksd qacqac qac qac qmtcqat qms? qacfo ksd tm(ti) qaly qc tcp qac qms? taf qmtc qaf2 qc ksj qacf tm qaf2 tbh1 qtaf qalo tm qap qtap kwcs thm tqcb tbhv qms qms qbs thm qat tbfr qaf2 qat qacf tcwm qac qac ksj tm qafo tcp qms qc qc qmtc qc qal qafo2 ti tql tbhvtvf qc qc qc qms qac kgc qaf2 qafy qac tcwu tbfr taf qac thma qms tcw tbh tm(td)qc tbhv qc qmtc tm kw qac qafo qms? qap qc tcwm qaly qap qac qc/tbh qc qms tm qc tbdhctm(ti) tcwl qap qmtc kwcs qap ksj qac qacfo tbhv qc ksd tbm qc ksj ksd ti tm(td) tm(ti) ksjs qap ksj qaf1 tcwm qms qc tcwl qacf kwcsqaf2 qac qafo qaf2 tbh tm(tbv) ksd qc ti qat kgc tm(tbv)qaly qmtc qacf qafy tm qc qms qco qmtc ti qafy tcwm qmtc tbh qafo tqcb qmtc tm(ti) td qaly qmtco qac tcwqap tm tcwu qap tbhv ksj tcw tbm qat qaly tbm? qacf tcp tbhv tm qaf1 qms qacqc qmtc tm(tbv) ti tbdh tcp qacf ti qbs qc tm(tbv) tbhv qafy qc qacf qaly tbfr tcwu kwcs qc qms tql tm(ti) qac qc qc qc qtap qms? qaly tbh qms tm qafy qmtco tm(ti) qc qafo kwcs qc qms taf tcp qc qco tcp qalo qacf tcp qc qac tbh qc tbm tm qafo tbdh qms? qc tm(tdbv) tm(td) tbh qc taf tcwqtap qms tm(tbh)tbdh qca qaf2 ti kgc tbh tm tql qaf2 kwcs tm(td) qap tcw qafy tbm tbh tm(ti) ksd qbs qacf qaly qaly ti? tm(tbv) tm(td) qac qmtc qac ti qms qac tm(tbv) qacf taf qaly qmtc qc qmtco qco qmt qmtco tbvkwcs qaly tbhv ksjs qap qaly tbhv tcp tm(tdbv) qafy tm(td) qmtc qaf1 ti qca qmt ti qafy tm(tbv) taf tcp tbh qms qafo? qacf qc tbm ksj kgc tcwu? qmtc qtaf tcp qacf ti qmtco qac tcwm? qac kwqc tcw tm(tql) tm(tbv) tm(td) tcp tvf qmt qac tm(ti) tbh qc/tbhkwcs qap qco tbhv tcwt? ksd qmt qafy qaco tm(td) tcwl qac qtap qc qaf2 ti? qbsqmt qaf2 qc/tbh qafo thm qms ksd qc ksd tm(tbv) qc tbh qms? ksd qc qmtc qap tm(td) qc qac tm(tbv) qafo kw qac qacf qmtc td tbh1 qat qaly qc/tbh ti qms? qmtc ksd kwcs qmtc tm qaco tcp taf tm(td) thma tbh1 kgc qmtc qc qacf qmtc tbh2 ti kwcs qap qap qac qms qafo tcwu? tbh qafo qmtc ksj taf qat qafo tm qafy qap qtaf tm(tbv) ti qaly qafo tcw tbhv taf tbh qms qco tcwl qap qacf qc tcwm? ti thm qac qac qms tm td qc qaf2 tcwm qtap ksj kgc qat tcwl tbh tcp qac tbh qacf tcp qac qc/tbh qafoqaly qc/tcpqc/tcp qla qc/tbh tcwl? qc tbh qms qafy qafy ti qac ti tcw qaf2 tm(tbh2) tcw qtaf tm(tbv) tm(td) qaf2 qc qafo tbh1 qms qac qc/tbh td ti tm(tdbv) tcwu qms? qmtc qc taftm tm(ti) tm(tbv) qms tbfr tbh qaly tcw qmtc ti qac qmtc qac qaco qc tm(td) tm(ti) qaf2 qap tm(tbh3)ti qac qafo tbh qafo tvfqmtc ti qac qms qms tbhv qms ksj qaco qc qc qc kgc tbh1 qafo qms tm(ti) qmt tm(td) qms qac tbh qacf qafy taf tm(tnw) qafo ksd taf qal tm(td) tm(tdbv) ti qbrh tbml qc qafo tcp thm tbm qac qactm(td) qafo taf td qaly tcwm qc qms? qc/kws tbh qaly qaco qacf qmtc qaly qaf1 qmtc tm(ti) qco qce qafo qac qaly tcp kwcs qacf tcp qac tcw qmtc tcwm tm(ti) qacf tm(tnw) tbfr qmtc tm(td) qaco qacf tcwl tbh3 tbh1 tm(ta) qc qmtc qmtc qms tcwl qafy qms? tcwl ti qafo tm(ti) qmtc qbwc qc qc qc qac qap ksj qac qc/tbh qap qafo kil qacf tbhtaf qmtc tbh tbm qaly tm(td) tbh2 qmtctcw qco qat qafy tm(tdbv) qc tbhb ti tcp qmtc qafo qap qal tbfr qat qacf tm(ti) qafo tm(tbv) tm(ta) qms ti kgc qacf qafo ti qmtc tm(ti) qms qc qmtc qaly tm(td) qaf1qac tbh qc qac taf qc tbh tm(ti)qms(ti) qms tbm tdk? qtap qmtc tm(td) kiu ti qac qac tm(td) tbh qacf qc qms? qaly qmtc qaly thm qms qmtc tm(td) tm(tbv) qmtc qmtc qafo qacotm(tbv) tbh tbhv qmtc qc qaf2 qacf qtap qms qc qac tbh kgc tm(tbv) qac tm(ti) qms?ti qafy qea qmsc qmtc qaly qms tcw qac qafo tcp ksj qc ksd tm(tbv) tm tm ksj qafy ksd tm(tbv) qap qms qc qaly taf qacf qms? tm(tbh) qms tbh1 tbh2 qafy tm(tdbv) tm tm(ti) tm qms? qac qac qacftm(tn) qap qafo2 tbm tcp qc tcp tm(td) qc qafo kgc qmstm(tnw) qc qafy tm(tnw) tbh3 qac qaco qap qc qms? qms tm(ti) qafo qaf2 qacf qac qms kwcs tcw qc tbv tbm tm(tbv) kiu qalo tm(td) qc tm(ti) qc tbh tm(td) qcqc tbh? tbh qmscqaf2 tm(ti) qacf tcwm tcwl qms qafo qap qc qafy qat qafo tm(td) qac tm(ti) qc qbw qbrhc tbh tm tm(td) tm qc tm(tnw) qaly qap qms? qms tm qc qacf tbml qc/tbh? tbml qac tm(td) tm(tbv) qaf2 qafo tbh3 tm(ti) tmtcp qat kwcs tbh2 tcw ksd kgc qafo qms qms tm(tbv) qafo qacf kwcs qalo qms tbml qac qaf1 ksj qac tm qafo tm(ti) tbm qaly qafo tm(tbv) qmsc tbm tm qac qac tm(tbv) qafy qafy qac qco qcqms qc tm(tbh1) qmsh qmsco qacf tbh3 qtaf tbh1 tbh qms qms tbh1 tm qc qaf2 tbm qaly qap qaly qafy qmtc qaly qc qafo tm tm qaly qbw tm(td) qafo kil tcw tcw tm(tbv) qafo tbh qafy tbml qafo2 tbm tbh1 tbh2 tm qap qae tm(td) tbh tm(td) qmt tbh tm(ti)kgc tm(td) qctbh1 qafo qms? qaf2 kiu tbh1 tbh qmsc tbh3 tbh2 qafc tm(ti) qap kil qms qafo qc tm(td) qc tm qac tm(tbb) tbh1 tbh qafy qafo tbh2 qac tm(tbv) tbh1 qafo qco qmtc qc tm(ti) qlm qafo qafo qat qc tm(ti) tbh1 qaly kwcsd tm(ti) qaly qc qafo qc tm(td) qms qaly tm(tn) qafo qce qmsc qaly qafy tbh qap tm(tbv) kgc tm(tbv) tm(ti) qaly qafo tm(td) qmsc tm(tbv) qbrh qmsc tbh3 qms tbh qms qms tnw kgc td qmt tcp qacf qc qafy tm(tn) taf qth qc tbh qaly kiu qat qaf2 tm(ti) tm(tbb) tm(tbv) qms qac tbh qc qafo qafy tm(tn) kgc qc qat qms? qafo qms qafo qms kgc tbhtm tm(tnw) qaf2 qms? tbh tcp tm(td) qmtc qacfo kiu tm(tbb) ku tbh3 tm(ti) td qac qc qlm qc kil ksj qc qafo qac qaly qac qafy qaf2ku qafc qafy qacf tm(td) qc qc ksj ku tm(tbvv) tbh3 tbh tdk tm(tbv) tm(ti) tm(tn) qc ksj qaly qmsco qacf qaf2 tbh3 qc tbh3 qap tm(td) qmsc kiu qaly qaly qes tm(tbv) qafo jcn tm(td) qafo qafo qal qc tbh3qap qafo tm(tbv) qc qc tm(td) jn qmt qaf2 qafo tbh3 qac qap tm(td) jc tbh qafo tbh qaf2 qaf2 tm(ti) tm(td) qafo tm(ti) qac qc qaf2 qaf2 tm(tbv) qc tm(td) kgc qaf2 qc tm qms qaf2 qafo qmtc tbh3 qcqc qmsc tm(td) qap qms tm(ti) qms qafo qc/tbh qmsc qacf tm(tbv) qes qac qms kgc qms qafo qacf qmsc tm(tn) tm(td) tbh tbh3 qafo qlm qac tbh ksj qms tbh3 qmtc qap jn? qaf2 tm(tbh) tm(ti) tm(tnw) tbh qafo qms tdk qafo ti tip kwcsd qms qafo qes qafo qmtc tm(td)qac qafy qac tm(tbv?) tm(ti) tm(tbv?) qafy qc qmsc tdk qms qc tbh qaly qac qmtc qac tbh tm(tbh) qms qmsc tm(tnw) qafy qms? qafo qc qafo qms qms? tbh qap qtaf qaf2 tdk tbh qc tm(td) qat qafo qth tbh3 qc qms tm(tn) tnw tbt qmsc qaf2 qap qc/tbh qc/tbh ti tm(tbvb)tnl tnl tdk tbh qafo tbhu1 qafy qafo tbt qaf2 tm(tbb) qms tnw qaf2 qc qmt qac qac tdk qms qaf1 qms qafo qms tm(tbh) ti tnw tbh tnl qaf2 qap tm(td) tm(ti) tm(td)tm(td) qafo tbh3 qat tm(tn) qc tnw qaly qalyqc qac tnw qac qaly qaly ti qafo qaf2 qes qac qlm tbh tbh3 qacf qafo tm(tbb) qaly qafy qaly qafy qes qafy qaly tnw qac tm(td) qaco qms qaco qmt tdk qaly qaf2 taf qat qacfo qc qacf qafo tbh qlm qtaf tm(tn) ti qafotm(td) tnw qlp tn qafo tbh tbh tn qac ti tbh qms tbh tnw qctbh qth tnw tm(tbv) ti qc tm(tbv) tql tm(ti) tbt qc qapqc tm(ti) tcpl tm(tnw) qat ti qmsc tbh tm(td) qac ti qc tcptm(tbh) qes tm(tnw) tcp qms? tn qc qafo tdk ti qafy qafy ti tm(tbv) tnl qms qafo qms qafo qaly qap tbh tm(td) tn tm(tbh)qaf2 tnl tbh tm(ti) tbt qc tm(ti) tm(tn) tm(tbv) ti tm(tbhu1) tm(tnw) tm(ti) tm(tnw) tm(ti) tm(tn) tm(tbhu2) tbh td tnl qaf2 qac tm(tbb) qmtc tnw tbh qlm qafo qac tm(tbhu2) ti tn tm(td) tm(tbv) qaly qes qafy ti qafo qap qms tm(tn) qat tm(tbhu2) ti tm(ti) tm(ti) tm(ti) qac qaf2 tm(tbv) qap qaly tm(tbb) tm(tbv) qat qc qafo tm(tnw) tnw qtaf qafo tdk tnl qat qafo qac qmtc tm(td) tcp qafy qaf2 tm(ti) tm(tqcb) tm(tbv)tm(td) tdk qafy qac tnw tm(td) tnw tm(tbb) qaly tm(ti) tql qafo ti td tm(tdli) tbh tm(tbv) qms qafo qms tm(tbvb) tbt tm(tql) tnl tdk tdbrp ti qac tnw qat qc qc tdk qafo tm(tbv) qmtc tm(td) qafo tm(td) qac qaf2 tqltcp tbrp qms qafo qafo qc tm(td) qaly qafy qms? tbt td taf qac tdk qaf1 qafo qafo tdbrp tdk qaco tm(tbb) tbrp tm(td) qap qtaf qaly qc tm(ti) qafo tm(ti) tql tm(tbv) tm(tql) qaf2 qafo qap tm(tbv) qaly qc tbh qat tm(td) qms? tm(td) qafy tm(tbv) qac qafo tnl tiqafy qtaf qafo tdk tm(tbv) tm(tbv) qac tm(tbv) tcp qaly ti qafotm(tn) qafy tm(td) tbv qafo tbv tm(tcp) tm(tbb) qmtc qms td qaly qap tm(tbb) tm(tn) tbv ti tm(tbb) tm(tn) tm(td) qaly tm(tbv) tbht tql qafo qafo tm(td) tm(ti) td qacf tql td qms qafo qafotm(td) tm(td) tm(tbrp) tm(td)tm(tbrp) qafy qap tm(tn) qat tbh taf tm(td) qafy qap tm(td) tm(tbb) tm(ti) tm(td) qmtc tm(tds) tm(td) tm(tds) qafy tm(td) tm(td) tm(tbb) qafo qaly qms qafo tm(td) qafo tm(tbv) tm(tbb) qap tm(td) qac qap qap qaly qap qaly qap qat tnw qafy tm(tbv) qafy qafy tbv qafy qafc tm(td) tnl qafy qafc qafo2 tql qafo qaly qafy tql ti tbv ti tqin qafo tm qafy tm(ti) qms? tm(td) qth qaly qafo qaly qms? kil tm qms(tql) qaf1 qaly qc qal qal qafy qaf1 tm(tbv) tm tm(tbv) tm(tcp) tm(tbv) tm(tql) kgc qafo tnl tm(ti) kil kgc kgc tcp tcp jct tbh jn tm(tn) qafo qaf2 ti tcp tm(tbvb) tm(tbvb) tbh tm(td) tbh tm(td) qms qaly tm(tnw) tm(tbv) tm(ti) tbh tcp tm(td) tm(td) tm(td) tm(tnw) tm(tbv) tm(tbvb) tm(tbb) tm(td) tm(td) tm(td) tm(td) tm(tbv) tm(tbv) tm(td) tm(td) tm(tbv) tm(tbv) tm(tbv) tm(tbv) tm(ti) tm tm(tbv) tm(tbv) tm(tbv) tm(tbv) tm(tbv) tm(tbv) tm(ti) tm(tbb) tm(td) tm(tbb) tm(tbv) tm(ti) tm(tn) tm(tbv) tm(ti) tm(ti) tm(tbv) tm(tn) tm(tbv) qac tm(td) tm(td) tm(tbv) tm(tbv) tm(tbv) tm(tbv) tm(tbv) tm(td) tm(tbv) qafy tm(tbhu1) tm(tbh) tm(tbhu2) tm(tbb) qafy tm(td) tm(tbh) qafo tcw td td tbhu1 td qac qac kw tdk tbt tbt tdk tnw tcw tmf(tbh)tcw qap tcp tcp tcp kw qacf qaly tcptcwtbm tbh qac qac tcw tbm tcp tcp ksj qaly ksj qap qap qap qap tbm tbhv tbh qap qap ksj tbh1 qbrc tbh tbm tbh tm(td) tm(td) taf tcwu tvf qal tbrc qacf tvg tbrc tbrc kwcg qafo2 qafo qal qafotbdh tbrc?taf tbdh tbrc?tbdh taf qat qafy tbrc qal qtap tbh qal taf tbfr taf tm(td) thm tm(ti) qac thm tbh taf tm(ti) tm(ta) tm(tbh) ta ta tm(ti) tm(ti) tm(td) tm(tbv) tbh qacf tbh qtbx ti qms ti tm tm(tdbv) tm tm(tdbv) tm(tdbv) tm(tdbv) tm(tdbv) thm tm(tbv) tm(tdbv) tm tm(tbvf) tm(tbv) ti tm tm tm ti ti? tbh tm(ti) qms(ti) qms tm tcp tcp tm(ti) qacfo km tm(tbh) kw tcp qms tm tm qbe tm qbs qafc qbs qafo ti tql ti qafy tcp tqcb tqh tm(ti) ksj tcp tbh tcp ti kcm ksd qbwk qacfo kw ksd qalo kw ksd ksjs kw ksd qacfo ksd qat ksd kgc qafo calico bed? qbal calico bed? kgc tcwl tcp kgc qbef qtbx tcp qalo qacf qbmk2 qtap taf tbh qc tbh qbhk qac tcwu tcwl tbm tbm tbm tcwu tcwu tbm tcwu qap kw qafo2 qal tcp kkl kwcs tcwl kkl tcp qc tcw tcwl tcwl tbm kw ksst kw kw ksd jcw tbt qaly tm(ti) kil tm(tql) tql tm(tn) tm(tbv) tm(ta) tbh tm(ti) tm ti ti ti td kgc kgc ti tbh tbh ta tbh qtbx tbm tcwu tl kw qafy qms(ktd) tcwl tcwm kgc kgc kgc 25 11 15 3 8 4 20 9 22 20 18 18 15 6 45 4 11 3 3 1 2 2 2 2 25 7 5 5 8 70 12 9 4 6 2 27 3 5 3 3 2 2 12 5 4 3 3 2 2 2 2 3 2 2 30 35 7 6 8 6 30 7 4 12 13 8 22 17 2 3 3 45 25 12 45 7 9 20 18 42 34 22 5 15 3 5 5 10 9 12 8 14 30 10 20 10 8 11 10 4 45 40 40 2 40 35 21 5 30 40 30 50 35 35 30 18 10 35 10 30 15 10 55 12 10 8 10 15 30 27 15 22 25 65 35 35 35 10 32 30 5 4 30 45 7 25 80 2 14 4 28 33 15 7 30 2025 25 5 5 7 5 10 5 14 15 45 25 15 20 35 3 2 35 22 12 8 2 1 7 6 6 5 10 5 10 15 5 55 5 25 25 20 5 9 9 5 20 5 18 8 22 8 17 2218 15 32 5 30 25 32 20 15 25 10 5 10 8 12 6 4 12 4 11 20 35 10 35 70 15 15 80 35 10 8 17 8 14 20 12 5 11 6 24 15 7 10 18 22 15 15 25 25 15 20 15 13 9 10 15 85 40 40 20 35 12 55 8 4 4 3 4 6 5 65 52 80 34 65 10 4 12 3 7 12 6 10 5 16 30 5 12 7 20 2 3 5 11 14 14 25 12 4 5 8 11 30 25 2 5 10 6 2 8 6 10 5 1 2 15 2 1 4 1 1 3 5 2 35 10 40 40 15 26 30 20 15 13 15 25 20 5 4 5 2 10 20 50 75 10 25 10 70 3 5 45 40 30 12 46 15 35 14 20 3 2 3 4 2 2 1 10 2 1 2 3 2 3 3 2 2 2 2 2 2 10 10 5 7 3 20 60 63 55 58 37 78 80 72 85 10 2 2 2 2 7 5 2 8 12 10 40 14 50 30 25 15 33 15 4 12 4 26 25 12 70 60 75 41 6 2 30 80 2217 15 85 65 45 16 17 15 14 13 12 11 10 9 7 8 5 6 4 3 2 1 18 1 2 3 5 6 7 8 9 10 11 12 13 14 15 16 18 20 1 2 3 4 5 6 7 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28 30 31 32 33 34 35 36 37 38 39 40 41 42 43 44 45 46 47 48 49 50 51 52 53 54 55 56 57 58 59 60 61 62 63 64 65 66 67 68 69 70 71 72 73 74 75 76 77 78 79 80 81 82 83 84 85 86 87 88 89 90 91 92 93 94 95 96 97 98 15 100 101 102 103 104 105 106 107 110 114 115 116 118 119 120 121 122 123 126 127 128 129 130 131 132 133 134 135 136 137 138 139 140 141 142 143 144 145 146 147 148 149 150 151 152 153 154 155 156 157 158 159 160 161 162 163 164 165 166 167 168 169 170 171 172 173 174 175 176 177 178 179 180 181 182 183 184 185 186 189 190 191 194 195 196 197 198 199 200 201 202 203 204 205 206 207 208 209 210 211 212 213 214 215 216 218 219 220 221 223 224 225 226 227 228 229 230 231 232 233 234 235 236 237 238 239 240 241 242 243 245 1 2 3 4 5 6 7 9 10 11 12 13 14 99 16 17 18 19 20 21 22 23 24 26 27 29 30 31 32 33 3435 38 40 41 4243 44 222 46 48 49 50 246 247 248 249 250 251 252 253 254 255 256 4 17 29 108 109 111 112 113 117 124 125 187 188 192 193 217 244 8 25 28 37 39 47 15 syncline panguitch an ti cli ne c an yo n br yc e sy n c li n e ca ny o n bryce anticline bench johnson zo n e fa u lt paunsaugunt zo ne fa ult pa u n sa u g u n t fault thrustinn rubys fault thrust hills pine fa ult thrust valley johns fa ult th ru st creek huntzo n e fa u lt se v ie r zo ne fa u lt se vi er fa u lt pa s s sa n d zo n e fa u lt s e v ie r fa ul t pa n g u it c h fa ult valley be ar ea st fault va ll ey bear w est fa ul t ea st peak ir o n fa u lt w e st pe ak ir on fa ult ledge black fa u lt kn o ll w o o d fa ul t canyon ra tt le sn ak e fa ul t r id g e na va jo fault zo ne fault parowan fa u lt th r u st sp r in g s mounta in su g ar lo af fa u lt m o un ta in su m m it fa u lt h il l jo n es fa u lt r id g e m id dl e zo n e fa ul t pa r ag o n ah zo n e fault h il ls re d ir o n zo n e fa u lt h u r r ic an e f' f e' e d' d a' a b' b c' c 2 1 2 14 25 53 58 14 86 55 5947 20 3636 45 19 21 ? ? 1 0 1 2 3 40.5 miles 1 0 1 2 3 4 5 60.5 kilometers 1:62,500scale contour interval 50 meters disclaimer although this product represents the work of professional scientists, the utah department of natural resources, utah geological survey, makes no warranty, expressed or implied, regarding its suitability for a particular use, and does not guarantee accuracy or completeness of the data. the utah department of natural resources, utah geological survey, shall not be liable under any circumstances for any direct, indirect, special, incidental, or consequential damages with respect to claims by users of this product. for use at 1:62,500 scale only. this geologic map was funded by the utah geological survey and the u.s. geological survey, national cooperative geologic mapping program through usgs statemap award numbers 08hqag0096, g09ac00152, g10ac00386, and g11ac20249. the views and conclusions contained in this document are those of the authors and should not be interpreted as necessarily representing the official policies, either expressed or implied, of the u.s. government. plate 1 utah geological survey map 270dm geologic map of the panguitch 30' x 60' quadrangle geologic map of the panguitch 30' x 60' quadrangle, garfield, iron, and kane counties, utah 2015 by robert f. biek1, peter d. rowley2, john j. anderson3, florian maldonado4, david w. moore4, david b. hacker5, jeffrey g. eaton6, richard hereford4, harry filkorn7, and basia matyjasik1 1 utah geological survey, p.o. box 146100, salt lake city, ut 84114-6100 2 geologic mapping inc., p.o. box 651, new harmony, ut 84757 3 kent state university, retired 4 u.s. geological survey, retired 5department of geology, kent state university trumbull campus, warren, oh 44483-1998 6department of geosciences, weber state university, ogden, ut 84408-2507 7department of physics and planetary science, pierce college, woodland hills, ca 91371 approximate mean declination, 2014 tr u e n o r th m ag n et ic n o r th 111/2° base from usgs panguitch 30' x 60’ quadrangle (1980) projection: utm zone 12 datum: nad 1927 spheroid: clarke 1866 project manager: douglas a. sprinkel gis and cartography: jay c. hill, and basia matyjasik utah geological survey 1594 west north temple, suite 3110 p.o. box 146100, salt lake city, ut 84114-6100 (801) 537-3300 geology.utah.gov quadrangle location u t a h utah geological survey a division of utah department of natural resources in cooperation with u.s. geological survey sevier fault sevier fault black mountain a a’ parking area 12 89 utah ¹ 1 mile 1 kilometer 0 0 tcp taf figure 3. geologic map of the red canyon area, showing parking area to view the sevier fault zone. note numerous vent areas (stars) on rock canyon lava flow. q = various surficial deposits; qbrc = red canyon lava flow; tbrc = rock canyon lava flow; taf = old alluvial-fan deposits; tvg = coarse-grained volcaniclastic deposits; tvf = fine-grained volcanic deposits; tbh = brian head formation; tcp = pink member of the claron formation. modified from biek and others (2015). figure 4. schematic cross section showing displacement across the sevier fault near red canyon. basin-fill sediments (qta) were stripped off the upthrown side of the fault prior to deposition of the red canyon lava flow (qbrc). see figure 3 for cross section location and rock unit names. a a’ sevier valley west east black mountain paunsaugunt plateau qbrc qtaqta qta tbh tbhtbh tbh tc ku kuku tc tc tctctc tbhqbrc 89 r.f. biek sevier fault at red canyon 5 sevier fault th e sevier fault extends from south of the grand canyon northward to the southern part of the marysvale volcanic fi eld, a distance of at least 150 miles (250 km); it is known as the toroweap fault in arizona (fi gure 5) (lund and others, 2008). th e fault lies at the western base of the southern sevier plateau and paunsaugunt plateau. northward, fault-bounded basins in the heart of the marysvale volcanic fi eld, including the circleville, marysvale, and monroe areas, are on structural trend with the sevier fault zone, indicating that the fault bisects the volcanic fi eld and links up with down-to-the-west normal faults in southern sevier valley south of richfi eld (hintze and others, 2003; rowley and others, 2005). like most long normal faults, the sevier fault is composed of discrete parts that tend to rupture independently (lund and others, 2008). although lund and others (2008) suggested that hillsdale canyon, about 3 miles (5 km) south of red canyon, may mark a segment boundary on the fault zone, biek and others (2015) suggested a diff erent boundary farther north, east of panguitch. th ere, a complexly faulted bedrock high and associated graben in the hanging wall separate the hatch part of the valley on the south from the panguitch part of the valley on the north. in this area, numerous scarps are present on old alluvial fans where they collectively form a northeast-trending graben. th is area also corresponds to the south end of the sevier plateau, which is of higher elevation and relief than the paunsaugunt plateau. th is northern, “panguitch” section of the fault zone appears to terminate several miles north of the junction of utah highway 20 and u.s. highway 89, where it devolves into an anastomosing series of down-to-the-west and down-to-theeast normal faults before linking up with the eastern bounding fault of circleville valley (fi gure 5) (rowley and others, 2005). displacement on the sevier fault generally increases south to north along its length; it is about 3000 feet (900 m) at red canyon (lundin, 1987, 1989; davis, 1999; lund and others, 2008). apart from scarps in the hanging wall near panguitch, no scarps on unconsolidated surfi cial deposits are present along the main trace of the sevier fault in utah. th e inception of faulting on the sevier fault zone is poorly known but widely cited as middle miocene, 12 to 15 million years ago (see, for example, davis, 1999). claron formation many people get their fi rst glimpse of the spectacularly colored, hoodoo-forming claron formation not at bryce canyon national park or cedar breaks national monument, but here in red canyon. wide pullouts along the road attest to the area’s popularity, a stark contrast from the sage-covered, high-elevation basin and tree-covered plateau to the west. apart from the bright pink, orange, and red rocks themselves (fi gure 6), what attracts most people are the fantastic variety of fi ns, spires, and hoodoos that the claron formation naturally erodes into. hoodoos form by relatively rapid erosion of alternating resistant and less-resistant beds in combination with pervasive jointing, high precipitation rates, and some 200 freeze-thaw cycles per year. freeze-thaw cycles loosen rock, whereas summer thunderstorms work to wash it away. importantly however, new research shows that the local stress fi eld is a key factor that gives rise to the hoodoos; erosional processes act within that stress fi eld to create these fantastic forms (bruthans and others, 2014). th e vertical stress imparted by the weight of a resistant caprock holds grains together; as erosion removes some grains, the stress between the remaining grains increases, holding them more tightly together red canyon geosite red canyon geosite circleville marysvale panguitch section figure 5. th e sevier-toroweap, hurricane, and paunsaugunt faults in the basin and range-colorado plateau transition zone. arrows indicate fault section boundaries. each section likely has a diff erent rupture history and rate of longterm slip. modifi ed from lund and others (2008). 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 and making it harder to erode the sides of the hoodoos. erosion is thus slowed and the hoodoos tend to grow taller over time as intervening slopes erode. in addition, in the dry climate of southern utah, much of the precipitation that falls on the hoodoos quickly evaporates, leaving behind any cement that the water started to dissolve and thus making the rind of the hoodoos more resistant. acknowledgments my knowledge of southwestern utah geology comes from years of geologic mapping supported largely by the utah geological survey and u.s. geological survey. as the acknowledgments of those many published geologic maps attest, i am indebted to a great many people for their help over the years. th anks to grant willis and mike hylland (ugs) and kimm harty (ugs retired) for their insightful reviews and to jon good (ugs) for draft ing the fi gures. references biek, r.f., rowley, p.d., anderson, j.j., maldonado, f., moore, d.w., hacker, d.b., eaton, j.g., hereford, r., filkorn, h.f., and matyjasik, b., 2015, geologic map of the panguitch 30' x 60' quadrangle, garfi eld, iron, and kane counties, utah: utah geological survey map 270dm, 162 p., 3 plates, scale 1:62,500. bruthans, j., soukup, j., vaculikova, j., filippi, m., schweigstillova, j., mayo, a.l., masin, d., kletetschka, g., and rihosek, j., 2014, sandstone landforms shaped by negative feedback between stress and erosion: nature geoscience, v. 7, p. 597–601, published online, 20 july 2014 | doi: 10.1038/ngeo2209. davis, g.h., 1999, structural geology of the colorado plateau region of southern utah, with special emphasis on deformation bands: geological society of america special paper 342, 157 p. hintze, l.f., davis, f.d., rowley, p.d., cunningham, c.g., steven, t.a., and willis, g.c., 2003, geologic map of the richfi eld 30' x 60' quadrangle, southeast millard county and parts of beaver, piute, and sevier counties, utah: utah geological survey map 195, 2 plates, scale 1:62,500. lund, w.r., knudsen, t.r., and vice, g.s., 2008, paleoseismic reconnaissance of the sevier fault, kane and garfi eld counties, utah: utah geological survey special study 122, paleoseismology of utah, volume 16, 31 p. lundin, e.r., 1987, th rusting of the claron formation, the bryce canyon region, utah: tucson, university of arizona, m.s. thesis, 51 p. lundin, e.r., 1989, th rusting of the claron formation, the bryce canyon region, utah: geological society of america bulletin, v. 101, p. 1038–1050. rowley, p.d., vice, g.s., mcdonald, r.e., anderson, j.j., machette, m.n., maxwell, d.j., ekren, e.b., cunningham, c.g., steven, t.a., and wardlaw, b.r., 2005, interim geologic map of the beaver 30' x 60' quadrangle, beaver, piute, iron, and garfi eld counties, utah: utah geological survey open-file report 454, 27 p., 1 plate, scale 1:100,000. figure 6. brightly colored claron formation at the entrance to red canyon. yellowish-brown caprocks are pebbly conglomerate, evidence that these sedimentary rocks were deposited in stream and fl oodplain settings. utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors douglas a. sprinkel utah geological survey, p.o. box 146100, salt lake city, utah 84116 douglassprinkel@utah.gov the palisades at sheep creek canyon geological area—a geosite in the uinta mountains, daggett county, utah cover image: panoramic view of the palisades as seen from the sheep creek canyon geological area overlook. 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. the geosites reflect the interests of the many volunteers who wrote to share some of their favorite geologic sites. the list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. the utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. this is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and figures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: presidents message i have had the pleasure of working with many different geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several different geological provinces. we have the basin and range to the west and the central utah hingeline and thrust belt down the middle. the uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava flows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. the general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. the “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. the articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. this guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, thank you to the american association of petroleum geologists, rocky mountain section foundation for their financial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fill the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president sprinkel , d., 2022, the palisades at sheep creek canyon geological area: a geosite in the uinta mountains, daggett county, utah: utah geological association publication 481, 10 p., https://10.31711/ugap.v1i1.95. d.a. sprinkel the palisades at sheep creek canyon geological area—a geosite in the uinta mountains, daggett county, utah 3 abstract the palisades is an impressive ridge within the sheep creek canyon geological area—an area nestled on the north flank of the eastern uinta mountains not far from flaming gorge national recreation area. sheep creek cuts through the palisades, as well as the heart of the geological area, to reveal about 800 million years of geology, from ancient environments to the rise and ultimate erosion of the uinta mountains. the oldest rocks exposed at the palisades comprise the upper part of the neoproterozoic (about 770 million years ago) uinta mountain group, which have been thrusted upon the mississippian (about 350 million years ago) deseret limestone (equivalent to the upper madison limestone). that thrust fault and others exposed along the north and south sides of the palisades are part of the uinta thrust fault zone, which is responsible for intense folding of both formations. although the uplift of the uinta mountains and related deformation along the uinta fault zone set the stage for development of the palisades, it was erosion that revealed and shaped this spectacular feature. introduction the palisades is an impressive natural rampart of contorted and contrasting reddishand grayish-colored rocks through which sheep creek has cut a narrow gash. rising at least 1200 feet (370 m) above the creek bottom, the palisades reveal an important part of the geologic history of the uinta mountains, from ancient landscapes to the upheaval that formed the mountain range. arguably, it is the showcase geologic feature within the sheep creek canyon geological area. the palisades location coordinates are: n 40° 54' 24.07" latitude, w 109° 47' 49.45" longitude, wgs84; 40.906686°, -109.797070°. the u.s. forest service designated nearly 3600 acres (1460 hectares) of land as the sheep creek canyon geological area on may 13, 1962, to preserve the spectacular geology of the canyon for future generations (schell, 1969). this remarkable area is located on the north flank of the uinta mountains in west daggett county, utah, and is part of the sheep creek drainage west of and discharging into flaming gorge reservoir (figure 1). the elevation within the sheep creek canyon geological area drops from about 9000 feet (2700 m) along a ridge in the southwest part of the geological area to 6400 feet (1950 m) in the northeast part of the area where sheep creek enters the broader, alluviated valley in the lower part of sheep creek canyon. the peaks in the eastern uintas are generally below 10,000 feet (3000 m) in elevation, 3000 to 4000 feet (900–1200 m) lower than the high peaks of the west part of the range, due to extensional faulting that occurred during the past 25 million years (sprinkel, 2014), but are still impressive. one figure 1. color-relief map of northeast utah and southwest wyoming showing the two routes to the palisades geosite within the sheep creek canyon geological area. the shortest route from salt lake city, utah, in blue, is by traveling through southwest wyoming. the longer route, in red, is by traveling through northeast utah. the inset map shows the location of the sheep creek canyon geological area and main roads. the p on the inset map shows the approximate center of the palisades. 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 paved road, with small sections being graded gravel, goes through the geological area to form a loop that connects utah state road 44 (sr 44) near dowd spring on the south with sheep creek gap on the north. when and by whom the name “palisades” was first applied to this impressive feature is unclear. however, the name “palisades” was used in a 1969 publication by schell (1969), but it is not recognized by the united states board of geographic names. however, palisades campground, located near the foot of the palisades, is formally recognized by the board of geographic names and is on the 1963 version of the jessen butte 7.5’ quadrangle topographic map. palisades campground was later renamed palisades memorial park in honor of the seven lives that were lost in a flash flood (debris flow) that swept through the campground on the night of wednesday, june 9, 1965 (dietrich, 1979; sprinkel and others, 2010). the name “palisades memorial park” is shown on the 1996 version of the jessen butte quadrangle and is also recognized by the board of geographic names. the southeast-trending ridgeline south of sheep creek is formally named windy ridge—that name is shown on the 1963 and subsequent versions of the jessen butte quadrangle and is recognized by the board of geographic names. how to get there from salt lake city, utah, the palisades (within the sheep creek canyon geological area) can be reached by traveling east on interstate 80 (i-80) through southwest wyoming to manila, utah, and then south on sr 44. this route is about 190 miles (306 km), or about three hours, from salt lake city. the palisades can also be reached from salt lake city by traveling east to vernal, utah, on i-80 and u.s. highway 40 (u.s. 40), north on u.s. 191 along the flaming gorge-uintas national scenic byway, and then west on sr 44. this route is about 230 miles (370 km) and will take about 4.5 hours. driving the shorter route to the palisades from salt lake city (figure 1)—travel east on i-80 to fort bridger, wyoming (exit 34), then to mountain view, wyoming, on wyoming state road 414, and finally to manila, utah. note that wyoming 414 changes to utah sr 43 at the state line. from manila, travel south on sr 44 for about 6 miles (10 km) through sheep creek gap and turn west (right) onto paved u.s. forest road 218; small section of this road is graded gravel. a sign in sheep creek gap alerts drivers of this intersection. the palisades is about another 6 miles (10 km) on u.s. forest road 218. a good overview of the palisades is at the sheep creek canyon geological area overlook that is past palisades memorial park and at the top of the switchbacks. driving the longer route to the palisades from salt lake city (figure 1)—travel east on i-80 to the u.s. 40 exit at park city, utah, and travel u.s. 40 east to vernal, utah. from vernal, travel north about 35 miles (56 km) on u.s. 191 to the intersection with sr 44. continue straight ahead on sr 44 for about 15 miles (24 km) to the intersection of the spirit lake-browne lake road (u.s. forest road 218). continue about 7.5 miles (12 km) on u.s. forest road 218 to the sheep creek canyon geological area overlook. geology of the sheep creek canyon geological area sheep creek canyon geological area is on the north flank of the eastern uinta mountains (figure 1). about 800 million years of geologic history are showcased within the geological area, from the deposition of lower neoproterozoic sedimentary rocks about 770 to 750 million years ago (dehler and others, 2007; rybczynski, 2009; sprinkel, 2016) to the classic faults and folds of the laramide orogeny that uplifted the uinta mountains about 70 to 34 million years ago (hansen, 1965; sprinkel, 2010, 2018a). ten bedrock formations having a total thickness of about 6600 feet (2000 m) (figure 2) are exposed in northwest-trending bands that dip northeastward; the oldest rocks are exposed in the south part of the geological area and the bands young to the north (figure 3). the oldest rocks in sheep creek canyon compose the lower neoproterozoic formation of hades pass of the uinta mountain group. regionally, this formation is overlain by the red pine shale of the uinta mountain group, which is exposed about 4 miles (6 km) to the west of the sheep creek canyon geological area. at the palisades, however, the red pine shale is faulted out by a fault within the uinta fault zone, which is described in detail below. the rocks of the uinta mountain group were deposited by marine and non-marine processes from about 770 to 750 million years ago in a faulted basin (dehler and others, 2005; dehler and others, 2007; rybczynski, 2009; sprinkel, 2010). the basin that accumulated the thick sediments of the uinta mountain group was inverted by uplift before the end of the late neoproterozoic causing subtle folding, tilting, and erosion of the neoproterozoic section (stone, 1993; sprinkel, 2014, 2018a). the uplift and subsequent erosion formed an irregular topographic surface that affected deposition and distribution of cambrian tintic quartzite and equivalent lodore sandstone. afterwards, the late cambrian to devonian was a prolonged period—about 140 million years—of interspersed deposition and erosion. erosion prevailed—any sediments deposited during this time were later eroded. it was not until the mississippian, a quiet time during which much of the western united states was covered by a warm epicontinental sea (blakey and ranney, 2008), that sediments were laid down on the unconformable surface and preserved. marine conditions dominated the rest of the paleozoic, with a hiatus during late pennsylvanian to early permian time during which the eolian weber sandstone was deposited (figure 2). the youngest formation in the sheep creek canyon geological area is the lower triassic dinwoody formation. d.a. sprinkel the palisades at sheep creek canyon geological area—a geosite in the uinta mountains, daggett county, utah 5 figure 2. stratigraphic column of bedrock formations and surficial map units shown in figure 3. the palisades is a ridge of nearly vertical beds of the mississippian deseret limestone (upper madison limestone). the underlying mississippian gardison limestone and neoproterozoic red pine shale are faulted out at the palisades by the south splay of the uinta fault zone (fz). figure 3. geologic map and cross section of the sheep creek canyon geological area. the palisades is a structurally thinned sliver of mississippian deseret limestone that is sandwiched between splays of the uinta fault zone. a set of faults that are part of the hanging wall of the uinta fault zone displaces the formation of hades pass down to the south. 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 the uinta mountains rose along several segmented fault zones along the north and south flanks of the range during latest cretaceous to early tertiary time (70 to 34 ma). precambrian, paleozoic, and mesozoic rocks were uplifted, tilted steeply northward, folded into monoclines, and faulted (hansen, 1965; stone, 1993; sprinkel, 2010; sprinkel and others, 2010). about 18,000 to 25,000 feet (5500–7600 m) of rock was eroded from the rising highlands, eventually exposing the precambrian core of the uinta mountains. the landscape and drainage system of the eastern uinta mountains, which includes sheep creek canyon, continued to change as the gilbert peak erosion surface formed early in the oligocene during a time of relatively regional stability and planation (hansen, 1986). the erosion surface and the generally coextensive oligocene bishop conglomerate was later faulted and tilted during miocene extension, beginning as early as about 25 million years ago (aslan and others, 2017). this extensional faulting resulted in the structural collapse of the eastern uinta mountains that caused the reorganization of streams and rivers (hansen, 1986; aslan and others, 2017). renewed uplift of the colorado plateau during basin and range extension that began about 20 million years ago, and continues today, rejuvenated the upper colorado river basin and caused deep canyon cutting and active headward erosion of many creek and rivers in the uinta mountains. geologic description of the palisades the ridge of brownish-gray rocks that form the palisades consists of mississippian-age limestone and dolomite beds called the madison limestone, with only the upper part of the formation present because of faulting (schell, 1969; sprinkel, 2006; sprinkel and others, in revision). about 4 miles (6 km) west of the palisades at sheep creek, schell (1969) indicated that the madison is thicker and can be subdivided into a lower, thin-bedded unit of mostly limestone and an upper, thick-bedded to massive brecciated limestone and dolomite with brownish-gray chert. caves and sinkholes are common karst features in the upper madison caused by the dissolution of the limestone and dolomite beds that result in a complex groundwater network. water from big spring gushes from caves located near the base of the palisades into sheep creek (see discussion of big spring near the end of this article). continuing stratigraphic work and geologic mapping in the uinta mountains indicate that the lower madison correlates to the gardison limestone and the upper madison correlates to the deseret limestone of the wasatch range to the west on the basis of similar lithologies and stratigraphy (sprinkel and others, in revision). the identification of the basal delle phosphatic member of the deseret limestone of sandberg and gutschick (1984) is key to this correlation. a slope-forming interval that is the delle phosphatic member can be traced from the west part of the uintas (sprinkel, 2018b) into the central part of the range (kinney, 1955; poduska, 2015); however, this unit thins and becomes predominantly a slope-forming, thin-bedded cherty limestone and dolomite without phosphatic shale. i have chosen to use the terms gardison and deseret interchangeably for the madison limestone in this article. figure 4. panoramic view of the palisades as seen from the sheep creek canyon geological area overlook. view is to the northwest. the palisades is an impressive natural rampart of contorted and contrasting reddishand grayish-colored rocks through which sheep creek has cut a narrow gash. rising at least 1200 feet (366 m) above the creek bottom, the palisades reveal an important part of the geologic history of the uinta mountains, from ancient landscapes to the upheaval that formed the mountain range. lines labeled suf (south thrust splay of uinta fault zone) and nuf (north thrust splays of uinta fault zone) are thrust faults with the barbs on the upthrown side. the thick line labeled d (down) and u (up) along the left edge of view is a fault that generally parallels mahogany creek. the zone of light-colored rocks at the top of the lower formation of hades pass is likely the result of “bleaching” by post-depositional fluids. zhl – lower formation of hades pass, zhu – upper formation of hades pass, md – deseret limestone, iprv-ipm – round valley limestone and morgan formation, pipw – weber sandstone. d.a. sprinkel the palisades at sheep creek canyon geological area—a geosite in the uinta mountains, daggett county, utah 7 the deseret limestone forms the palisades in the sheep creek canyon geological area (figure 4). as elsewhere in the uinta mountains and the wasatch range, the deseret commonly displays karst features and springs similar to those found at the palisades. the deseret bedding attitudes in the palisades are upright, steeply dipping to the north, to locally vertical and overturned with steep dips to the south (figure 3). the reddish-colored rocks juxtaposed against the deseret limestone on the south side of the palisades are the lower neoproterozoic formation of hades pass of the uinta mountain group (figure 3). the formation of hades pass in this area is informally subdivided into lower and upper parts. the lower hades pass is mostly reddish-brown, mediumto coarse-grained, quartz sandstone with lesser amounts of feldspathic sandstone, and typically forms cliffs. the unit also includes some interbeds of reddish-brown siltstone and greenish-gray shale. sedimentary features preserved in the sandstone beds indicate a fluvial origin (dehler and others, 2005; kingsbury-stewart and others, 2013). an interval of light-colored sandstone beds is conspicuously exposed along mahogany draw (figure 4), which may be related to reducing fluids (“bleaching”) associated with the west-trending fault that generally parallels mahogany draw (figure 3). however, this interval of light-colored sandstone beds is also found near the top of the lower hades pass on the downthrown block above the west-trending segment of sheep creek south of mahogany draw, which can be traced westward. thus, the “bleached” interval may be related to laramide or older post-deposition fluid flow that is unrelated to faulting. the upper hades pass is lithologically similar to the lower hades pass with reddish-brown quartz and feldspathic sandstone, reddish-brown siltstone, and greenish-gray shale, except (1) the percentage of feldspathic sandstone beds increases relative to quartz sandstone, (2) the amount of interbedded fine-grained beds (siltstone and shale) also increases, and (3) it is thinner bedded forming ledges instead of cliffs. structurally, outcrops of the formation of hades pass are on the hanging wall (upper plate) of the south thrust splay of the uinta fault zone, and is up relative to the deseret limestone (figure 3). the formation of hades pass is also displaced down-to-the-south by faults along mahogany draw and the west-trending segment of sheep creek. beds of the formation of hades pass dip north 12° to 15° at the south end of the geological area but increase in dip to 58° northeast approaching the palisades (figure 3), and ultimately become nearly vertical to locally overturned (figure 5). the formations immediately north of the palisades represent the normal stratigraphic succession of the overlying mississippian humbug formation to the pennsylvanian-permian weber sandstone; however, the humbug and doughnut formations and, to some degree, the round valley limestone, are structurally complicated because of thrust faults in the north part of the uinta fault zone (figure 6). the deseret limestone appears to be faulted against the humbug formation, which steeply tilted the beds 74° northeast to vertical and overturned. another thrust splay is in the doughnut formation, but bedding attitudes are uncertain because exposures are poor. all formations are structurally thinned. bedding is upright and attitudes become gentler northward and away from the uinta fault zone, ranging from 37° northeast nearest the fault zone to 18° north-northeast near the north boundary of the geological area (figure 3). figure 5. the south thrust splay of the uinta fault zone (barbs on the upthrown side of thrust fault) on the south side of the palisades placed the upper formation of hades pass (zhu) up against the deseret limestone (md) or upper madison limestone. beds of the formation of hades pass and the deseret limestone are steeply dipping upright to vertical and overturned. the arrowhead on the white lines (representing strike) indicates the direction of bed tops. view is to the west. figure 6. the two thrust splays of the uinta fault zone (barbs on the upthrown side of thrust fault) on the north side of the palisades steeply tilts and structurally thins the formations involved in the faulting. bedding in the footwall is not as steep, becoming gentler away from (to the right) the thrust fault in the drainage (also see the cross section in figure 3). the arrowheads on the white lines (representing strike) indicate the direction of bed tops. thin solid lines indicate formation contacts; dashed lines indicate where contact is approximately located. view is to the northwest. md – deseret limestone, mh – humbug formation, mdo – doughnut formation, iprv – round valley limestone, ipm – morgan formation, pipw – weber sandstone. 8 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 the uinta fault propagates at depth and resurfaces west of the sheep creek canyon geological area where much of the red pine shale is faulted out. stratigraphic throw on the uinta fault increases eastward through the geological area. schell (1969) estimated about 1000 feet (300 m) of stratigraphic throw along the palisades, but my estimate is about 2600 feet (800 m) because the red pine shale, the gardison limestone (lower madison), and lowermost part of the deseret limestone (upper madison) are faulted out. stratigraphic throw continues to increase eastward as the paleoproterozoic red creek quartzite is thrusted upon the paleocene-eocene wasatch formation (hansen, 1965; sprinkel, 2006). hansen (1965) estimated as much as 34,000 feet (10,300 m) of stratigraphic throw along the uinta fault zone in flaming gorge national recreation area; however, it may be only about 26,000 to 28,000 feet (8000–8600 m) because the uinta mountain group is likely only 16,000 to 18,000 feet (4900–5500 m) thick (sprinkel and waanders, 2005), less than the estimated thickness of 24,000 feet (7300 m) by hansen (1965). how and when did the palisades form? the palisades’ spectacular scenery results from the combination of movement on the uinta fault zone and differential erosion of the deseret limestone. the deformation along the uinta fault zone is associated with uplift of the uinta mountains, which occurred during a laramide mountain-building event that began about 70 million years ago and ended about 34 million years ago (sprinkel, 2018a, 2018b). upward movement along the south thrust splay of the uinta fault zone placed the reddish-colored formation of hades pass up against the deseret limestone and folded rocks on both sides of the fault as reflected in the steeply dipping to overturned beds. additional upward movement on the two north thrust splays structurally thinned involved formations and increased the dips on the entire hanging wall of the uinta fault zone. as a result, the faults sandwiched the resistant deseret limestone between less-resistant formations. erosion of the stratigraphic section by a network of streams began as the mountain range was uplifted. erosion continued to remove formations as streams cut deeper into older rocks even after uplift stopped. however, rocks are not removed evenly. less-resistant rocks tend to erode faster, leaving more-resistant rocks standing relatively higher. the palisades is an example of differential erosion where the less-resistant formation of hades pass on the south and the humbug-doughnut formations on the north eroded more easily than the resistant deseret limestone. the steeply dipping beds also played a role in forming a knife-edge ridge of deseret limestone. the contorted and contrasting colored beds of the formations in the palisades area creates an impressive view. big spring as noted previously, big spring is located at the base of the palisades along the sheep creek canyon loop road (figure 1). the spring is part of a karst system developed within the deseret limestone (upper madison limestone) and is one of many similar springs that occur in the uinta mountains. like other karst springs, the source of the water for big spring is snowmelt and rain that feeds streams which disappear underground through fractures and sinkholes at the surface. the groundwater generally flows through a network of solution-enlarged fractures and caves to finally emerge at the surface as springs. water from big spring discharges from sheep creek caves and flows down a small talus slope into sheep creek (figure 7). the flow rate of big spring is not regularly measured but varies depending on snowmelt and the amount of rain received during precipitation events in the drainage basin during the spring. mundorff (1971) estimated a flow that ranged from less than 5 cubic feet/second (cfs) to at least 36 cfs (0.14–1.0 cms); that equates to about 2244 to 16,158 gallons/minute. figure 7. water from big spring (discharge area in the trees at the base of the palisades) discharges from sheep creek cave, which is in the deseret limestone, and flows down a talus slope into sheep creek. the cave openings visible in the cliff face likely represent a former springhead (spangler, 2005). water from the spring originates, in part, from lost creek sink, more than 14 miles (22.5 km) to the west. view is to the west. d.a. sprinkel the palisades at sheep creek canyon geological area—a geosite in the uinta mountains, daggett county, utah 9 a major source of water for big spring is from a sinkhole found at the base of a ridge of madison limestone (deseret limestone) along lost creek about 14 miles (23 km) west of the spring. dye tracing in 1979 and again in 2001 confirms that one of the sources of water for big spring is from lost creek, which disappears into lost creek sinkhole. this underground system is one of the longest in utah documented by dye tracing (spangler, 2005). the lost creek sinkhole is at an elevation of about 9100 feet (2800 m) and big spring is at an elevation of about 6945 feet (2100 m). notably, lost creek is part of the burnt fork drainage and big spring is part of the sheep creek drainage with a surface hydrologic divide between the two drainages (see figure 3 of spangler, 2005). the water disappearing into lost creek sinkhole flows through a network of fractures and caves within the steeply dipping deseret limestone and drops more than 2100 feet (640 m) in elevation before discharging at big spring, despite crossing a surface hydrologic divide. results of the dye-tracing test also indicate that the travel time of water from lost creek to the spring is no more than two weeks (spangler, 2005). acknowledgments i am grateful to mary beth bennis (utah field house museum of natural history state park), larry spangler (u.s. geological survey), and grant willis (utah geological survey) for reviewing this manuscript. their reviews improved the technical aspects, as well as the readability and clarity of the manuscript. they have my thanks. i also thank utah geological survey technical editors, stephanie carney and michael hylland, for their keen review and editing. finally, i thank ashley national forest for their support of my geologic mapping over the past 20 years. references aslan, a., boraas-connors, m., sprinkel, d.a., karlstrom, k.e., heizler, m., lynds, r., and becker, t.p., 2017, cenozoic collapse of the eastern uinta mountains and 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stratigraphy of north flank of eastern uinta mountains and southern green river basin: utah geological survey open-file report 650, 1 chart, unscaled. sprinkel, d.a., 2018a, mysteries of the uinta mountains: utah geological survey, survey notes, v. 50, no. 3, p. 1–3. sprinkel, d.a., 2018b, interim geologic map of the duchesne 30' x 60' quadrangle, duchesne and wasatch counties, utah: utah geological survey open-file report 689, 38 p., 2 plates, scale 1:62,500. sprinkel, d.a., 2010, geology of flaming gorge national recreation area, utah-wyoming, in sprinkel, d.a., chidsey, t.c., jr., and anderson, p.b., editors, geology of utah’s parks and monuments: utah geological association publication 28 (third edition), p. 285–308. sprinkel, d.a., park, b., and stevens, m.d., 2010, geology of sheep creek canyon geological area, northeastern utah, in sprinkel, d.a., chidsey, t.c., jr., and anderson, p.b., editors, geology of utah’s parks and monuments: utah geological association publication 28 (third edition), p. 539–551. sprinkel, d.a., rybczynski, d.j., kowallis, b.j., dehler, c.m., and pederson, j.l., in revision, geologic map of the dutch john 30' x 60' quadrangle, daggett and uintah counties, utah, moffat county, colorado, and sweetwater county, wyoming: utah geological survey map xxdm, gis data, 3 plates, scale 1:62,500. sprinkel, d.a., and waanders, g., 2005, stratigraphy, organic microfossils, and thermal maturation of the neoproterozoic uinta mountain group in the eastern uinta mountains, northeastern utah, in dehler, c.m., pederson, j.l., sprinkel, d.a., and kowallis, b.j., editors, uinta mountain geology: utah geological association publication 33, p. 63–73. stone, d.s., 1993, tectonic evolution of the uinta mountains— palinspastic restoration of a structural cross section along longitude 109o15', utah: utah geological survey miscellaneous publication 93-8, 19 p. 1 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 abstract the modern (holocene-age) great salt lake (gsl) and pleistocene lake bonneville of the bonneville basin (bb) together make a geosite (gsl-bb system) of exceptional scientific, cultural, aesthetic, and societal value. gsl is the largest saline lake in the western hemisphere and a sensitive recorder of climate. for millennia, this distinctive salty water body has been a dynamic and complex natural ecosystem, including an important waterway for birds and other wildlife and an archive of environmental change and history. lake bonneville is a seminal part of the history of science in the united states through the work of g.k. gilbert, who in the 1870s and 1880s developed both critical scientific concepts (e.g., isostasy) and methods (e.g., multiple working hypotheses), which are still employed today. gsl is a major tourist attraction, an economic driver, and a place of scientific exploration. yet today, the gsl is in grave danger of near total desiccation due to a combination of factors: human removal of waters that would normally replenish the lake, climate change, and other environmental pressures. over the past few decades there has been a growing international movement to recognize and respect our geoheritage, by raising visibility and protection of high-priority geosites. the gslbb system is a geoheritage site that urgently needs our protection. the holocene great salt lake and pleistocene lake bonneville system: conserving our geoheritage for future generations marjorie a. chan1, charles g. oviatt2, bonnie k. baxter3, basil tikoff4, and genevieve atwood5 1department of geology and geophysics, university of utah, salt lake city, utah, marjorie.chan@utah.edu 2department of geology, kansas state university, manhattan, kansas 3great salt lake institute, westminster university, salt lake city, utah 4department of geoscience, university of wisconsin, madison, wisconsin 5earth science education, salt lake city, salt lake city, utah geoheritage concept an international movement over several decades, a growing international geoconservation movement recognizes that exceptional geological sites need to be protected and managed as part of our geoheritage. the geological society of america position statement (geological society of america, 2022) defines geoheritage sites as areas with geologic features of significant scientific, educational, cultural, and/or aesthetic value. these sites are key to advancing knowledge and support the broad understanding of the environment, its geodiversity and biodiversity, and the factors that influence climate change (see america’s geoheritage ii workshop proceedings, 2021 https://nap.edu/26316). although biodiversity is notably visible to the public, the geologic setting – its geodiversity and the convergence of geographic to environmental conditions – commonly form the underpinnings and context for biodiversity. the extensive and rapidly expanding body of literature on geoheritage is too extensive to detail here (e.g., see summaries of brilha, 2015, 2018; reynard and brilha, 2018; brilha and others, 2018). the united states is endowed with many sites that embody a rich geoheritage. the u.s. and state park systems have had an important impact on the conservation movement, but there has been growing recognition for more coordinated global recognition of natural sites. thus, geoheritage calls for global communication and cooperation, and provides the context that covers much of the science and education related to important geosites, while also embracing ethics, outreach, inclusivity, protection, and management. geoheritage also relies on modern technology to understand and model how natural systems are impacted. as earth scientists, we understand earth systems, with their change and interrelationships, and feedbacks in time and space. we must be caretakers and advocates for gsl, as we have both the knowledge and responsibility to help balance nature and societal needs. geosite locality in the basin and range province, ancient lake bonneville (figure 1a) covered much of western utah during the last glacial maximum. the modern gsl (figs. 1b, c) is the recent version of the closedbasin gsl-bb system, which, during the past few 10.31711/ugap.v51i.132 https://nap.edu/26316 2 m.a. chan, c.g. oviatt, b.k. baxter, b. tikoff, and g. atwood the holocene great salt lake and pleistocene lake bonneville system million years, has been dominated by various saline to hypersaline lakes similar to holocene gsl. between 30,000 and 13,000 yr bp the lake system was deeper and more extensive (lake bonneville) and was dominated by freshwater (figure 1d, currey and others, 1984). there is much interest in the gsl as shown by the considerable literature covering more than a century, including this volume (also see gilbert, 1886, 1890; oviatt and shroder, 2016a). the significant runoff that resulted from the wet winter of 2022-2023, does not significantly ameliorate the long -term decline in water level of the gsl. for centuries, gsl has been the largest saline lake in the western hemisphere, recording a history of change (madsen, b.d., 1989; gwynn, 2002a). but now in the anthropocene, drying of gsl (figs. 1b, c) and the probability of it disappearing, leaving behind a bowl of toxic dust with a few pools of salty water, has understandably raised alarm (e.g., flavelle, 2022). with growing pressures of urbanization in utah, the geologic features in antelope island state park in davis county, utah, provides one of the few sites left to easily access the gsl and see the context of its history over millennia, including the cyclic rises and falls of gsl and lake bonneville. this paper focuses on the broad spectrum and overview of geoheritage values and why it is important to protect the gsl. geoheritage values cultural and historical value the gsl-bb system has significant cultural as well as historical value because of the role that the landscape played for indigenous peoples as well as in the subsequent exploration of the west by european americans. humans have occupied the great basin for thousands of years. native american tribes that have lived in the gsl region, and that are still an important presence, include the western shoshone, goshute, ute, paiute, and washoe peoples (national park service, 2015). the landscape was a vital resource where native people hunted and gathered for sustenance, and the gsl watershed provided an exceptional bounty (e.g., madsen, d.b, 1989). today many tribal descendants feel an important connection to the land, particularly where open spaces retain much of their original, natural expression. in the 19th century with expansion and exploration of the west by european americans (e.g., stegner, 1954), early scientific studies included the documentation of lake bonneville, based on studies of its shorelines, deltas, and sediments by renowned american geologist g. k. gilbert (1886, 1890). his careful studies on foot and horseback allowed him to deduce that valley floors were previously covered by water and the isolated mountain ranges had been islands and peninsulas in a pleistocene water body he named “lake bonneville.” gilbert used the bonneville basin to investigate the idea of isostasy (equilibrium adjustments of earth’s crust to changing distributions of weight at the surface, in this case the growth and eventual loss of the water load of lake bonneville). individual shorelines of lake bonneville vary in elevation with the highest elevations occurring where the lake was deepest (the weight of the water in the lake depressed the underlying crust, and when the water evaporated, the crust rebounded). this work was the case example for gilbert to illustrate the methodology of multiple working hypotheses to overcome bias in human reasoning (gilbert, 1886). his recognition of the dynamic equilibrium of landforms and his correlation of shoreline elevations was seminal to understanding the complex interplay of isostasy and basin tectonics. gilbert identified and quantified evidence of shoreline superelevation and effects of fetch on shoreline elevations of gsl and lake bonneville. because of gilbert’s work, the gsl-bb system represents a seminal part of the history of science in america. gilbert used his experiences in this basin to understand distinctive shoreline barriers, terraces, and spits, and he chronicled the causal changes in hydrology based on rises and falls of lake bonneville and the highstands of gsl during the 1870s compared to falling levels of gsl during the 1880s. remarkably, gilbert’s seminal work has been an inspiration to people all over the world who have studied the history of closed-basin lakes. to be able to retrace gilbert’s thoughts and walk in his footsteps has deep meaning for those who value historical significance. many bonneville shorelines are now being rapidly lost or covered due to urbanization, but antelope island state park preserves near-pristine records of these ancient shorelines. scientific and educational value the gsl-bb system encompasses a rich geoheritage (figure 2) and contains many classic textbook geologic features and landscapes, that are significant to both education and research. much of the specific science is detailed in other papers of this volume. geomorphology and ice age history the geoheritage value of the bonneville basin’s prominent ice-age landforms is explained in more detail in other publications (chan and currey 2001; 3 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 1. the gsl-bb system. a. location of gsl and lake bonneville in western utah. b. overview map of gsl showing the historic average elevation, and the new 2022 historic low (figure from clark and baxter, 2023.) c. corresponding landsat satellite imagery of gsl elevations showing the record high of gsl in 1986 at left vs. historic low in 2022. ai = antelope island. images (images are public domain.) d. known bonneville basin lake cycles. the blue line labeled b in the main graph marks the bonneville deep-lake cycle. vertical black bars represent older deep-lake cycles. the base of the main graph is the elevation of modern gsl. inset shows the shoreline history of lake bonneville (blue) and gsl (red) with named shorelines (also see figure 3). (inset figure from oviatt and shroder, 2016a). 4 m.a. chan, c.g. oviatt, b.k. baxter, b. tikoff, and g. atwood the holocene great salt lake and pleistocene lake bonneville system figure 2. the gsl-bb system has many geoheritage values including historical, scientific, educational, aesthetic, economic, and societal. a. polygonal cracks south of gunnison island. b. colorful imagery at antelope island. c. gsl lies at the intersection of urban and natural settings (antelope island looking east, herd of antelope in the foreground). gsl is a major attraction that draws tourists. gsl enhances the quality of life in the salt lake valley. images: j. long. 5 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 chan and others, 2003; chan and godsey, 2004, 2016). since lake bonneville was the largest pluvial lake in the western hemisphere (that is, it was caused by climate change and an increase in effective moisture in the basin and was not fed by glacial meltwater), it is a natural laboratory for study (figure 3), bolstered by the well-dated shorelines that provide a precise lake hydrograph linked to pleistocene climate change. study of the gsl-bb has unparalleled analog value for many other large lake systems. the varied character of the lake is a result of climate change in the basin, causing the lake to range from small and hypersaline to large and nearly fresh. because the basin is so big and deep relative to the amount of water that enters the system, the lake has remained hydrographically closed for most of its history. connecting lake bonneville and gsl to lakes farther back into the pleistocene, subsurface cores like the burmester core (eardley and others, 1973; oviatt and others, 1999), tell the story of only four deep-lake cycles during the past 800,000 years. lake bonneville was the most recent of those deep-lake cycles, and the deepest because it had the benefit of input from the upper bear river and rivers in cache valley, which were diverted into the basin after 50,000 yr bp. all together those four deep-lake cycles took up less than 10% of the past 800,000 yr — the rest of the time the lake was shallow, similar to the historic gsl (oviatt and shroder, 2016b). prior to 800,000 yr bp, the lake system stayed at low levels back to about 3 million years ago. thus, other than the four deep-lake cycles, the history of the gsl-bb system indicates that our modern view of gsl is typical of the past few millions of years — a shallow hypersaline lake in a desert environment. the combination of geomorphic and the sediment records are valuable analogs for other large lake studies, in part because the record in the gsl-bb system is so intact, with distinctive markers of change over documentable spatial and temporal scales. the landscape expressions are also analogs to understanding geologic processes and applying them to regions of mars (e.g., chan and others, 2016). gsl ooids gsl is known as the world’s largest lacustrine carbonate depositional system (baskin and others, 2022). distinctive carbonate ooids (figure 4) — coated grains formed where waves agitate the lake bottom sediment — of gsl are long-standing world class examples. these sand-size features form when finegrained particles, such as brine-shrimp pellets or tiny sand grains, become coated with successive thin, concentric layers of calcium-carbonate crystals (crystals of the mineral aragonite arranged radially outward from the center of the ooid; e.g., sandberg, 1975; figure 4a). recent work of lincoln and others (2022) suggests that the radial pattern is derived from recrystallization. the gsl ooids contrast with other classic examples, such as bahamian ooids that have calcium carbonate crystals arranged parallel to the grain coatings rather than radially. oolitic sand is commonly cemented into beachrock (figure 4b), which is an indicator of lithification along older shorelines, with cementation aided by microbial activity (lincoln and others, 2022). microbialites microbialites are organo-sedimentary mounds formed by the actions of complex microbial mats (burne and moore, 1987; lindsay and others, 2017), and gsl has an extensive distribution in the highsalinity water (baskin and others, 2022; carney and vanden berg, 2022; pedone and others, 2023). photosynthesis by cyanobacteria and sulfate metabolism by other microorganisms create conditions that precipitate calcium carbonate (burne and moore 1987). in addition, the extra-polymeric substance (eps; a term commonly used by people who study microbialites) secreted by the cyanobacteria trap carbonate sediment, which creates a substrate on which new mats grow toward sunlight, hence the mound shape. some microbialites follow older polygonal crack patterns (figure 4), possibly because they are texturally different sites that might enhance biomediated growth, but microbialites also occur as individual mound buildups (figure 5) up to 1.5 m high that cover as much as a quarter of the lake floor (chidsey and others, 2015; vanden berg, 2019; baskin and others, 2022; wilcock and others, 2024). microbialite growth is sensitive to water chemistry and depth (light), wave energy, substrate, and other environmental factors (kanik and others, 2020). cyanobacteria-based mats represent the earliest fossilized life form on earth; layered and mounded accumulations of microbialites are well-preserved in carbonate rocks in the geologic record. the longevity and adaptability of microbialites accounts for their distribution on our planet in modern extreme environments, such as gsl. the study of gsl microbialites has implications for the search for biosignatures on mars (noffke, 2015; chan and others, 2019; gill and others, 2023). mineralogy and mirabilite evaporite minerals such as halite (nacl) have a long history of being extracted from gsl waters 6 m.a. chan, c.g. oviatt, b.k. baxter, b. tikoff, and g. atwood the holocene great salt lake and pleistocene lake bonneville system (gwynn, 2002b). additionally, unusual cold-water, saline-lake minerals, such as mirabilite (hydrated sodium sulfate, na2so4•10h2o, also known as glauber’s salt), occur in spring mounds that are visible during winter months (figure 6). groundwater seems to be partially dissolving a subsurface mirabilite layer, and then the mirabilite minerals are reprecipitated at the surface where spring water emerges. once the sodiumsulfate-rich spring water hits the cold winter air, mirabilite crystals form and build up a collection of small, mounded terraces, with beautiful crystals (figure 6) that are stable only in sub-freezing dry environments. some of the mirabilite-rich springs have colorful pools that are being studied for the associated microbial life (e.g., jagniecki and others, 2021; gill and others, 2023). these unusual mineralogies have implications for astrobiology and understanding life in extreme environments. figure 3. shorelines of lake bonneville in antelope island state park preserved at white rock bay and superimposed on the mountain bedrock. the landforms are a valuable record of geologic history and climate change. a. prominent shorelines (photo taken in 2014): s = stansbury shoreline, b = bonneville shoreline, p = provo shoreline. b. many shorelines formed during the rising and falling phases of lake bonneville, here showing wellpreserved examples between the stansbury and provo shorelines on this hillside; gsl at far right (barely in sight; photo taken in 2012). images: m. chan. 7 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 ecosystem significance gsl is a delicately balanced ecosystem (figure 7). the extreme conditions of gsl gives rise to a rich biodiversity and a special set of lifeforms, including brine shrimp and brine flies and the microorganisms that feed them, which have implications for understanding life adaptations in extreme environments (baxter and butler, 2020). the gsl provides important food and shelter to over 10 million migrating birds (sorenson and others, 2020; gslep, 2022), in addition to generating billions of dollars in revenue from tourism and the brine-shrimp industry (bioeconomics, 2012). life on earth needs water, yet water in the gsl watershed has been extracted and diverted for many purposes, such as for growing alfalfa and building housing subdivisions and supporting infrastructure. this has significantly impacted the inflow and replenishment of the lake, which has been drying and could figure 4. distinctive gsl features from the north side of antelope island that have scientific value. a. loose spheroidal oolitic sand (mostly ~ 0.2 to 0.5 mm diameter) with scattered weathered siliciclastic granules derived from nearby bedrock exposures; diagrammatic inset shows gsl radial structure of gsl ooids vs. the common marine tangential structure exemplified in bahama ooids. b. cemented beachrock composed of oolitic sand. images from bridger bay, antelope island, m. chan. 8 m.a. chan, c.g. oviatt, b.k. baxter, b. tikoff, and g. atwood the holocene great salt lake and pleistocene lake bonneville system potentially leave a basin of toxic dust that could impact regional communities (flavelle, 2022). declining gsl water levels threaten economic activity, public health in adjacent communities and ecosystems of gsl (larsen, 2022; great salt lake strike team report, 2023). it is clear that strategies to improve water management and increase deliveries to the lake are critical. gsl is an extreme ecosystem of biodiversity and geodiversity that is too important to lose. societal value there is no doubt that the gsl and the bonneville basin comprise an aesthetical geoheritage landscape that is visually appealing and that inspires a sense of awe and wonder (figure 8). the landscape of gsl, enhanced by open space and the natural setting of flora and fauna, has cultural and historical roots, and impacts economic development and tourism as well as quality of life. shrinking water levels of gsl have put this ecosystem into a state of crisis. diminishment of the gsl will threaten wildlife and further degrade utah’s air quality. society needs geoheritage sites like gsl because these sites are critical to advancing knowledge about water, climate and environmental changes, evolution of life, minerals and resources, and other aspects of the nature and history of earth (geological society of america, 2022). numerous studies show that nature and the outdoors provide positive impacts on mental health and cognition (e.g., bratman and others, 2019; weir, 2020). gsl is an outdoor classroom that enhances public understanding and engagement with science (figure 8), while providing recreational areas that improve quality of life, as well as economic support to local and regional communities as tourist destinations and as vital mineral and water resources. conclusions drying of pleistocene lake bonneville, which ended about 13,000 years ago, left both ancient shorefigure 5. microbial mounds of gsl at bridger bay, antelope island state park have important implications for understanding early life, with applications to astrobiology. a. and b. buildups; c. and d. cyanobacteria growth holding together oolitic sand grains, with some elongate brine-fly pupae cases. images a, b: b. baxter. images c, d: m. chan. 9 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 lines and the modern gsl, the largest saline lake in the western hemisphere. this gsl-bb system, as a whole, is a unique and valuable geoheritage archive of climate change and an extreme ecosystem that is often underappreciated and is now under threat of being lost. the gsl-bb hosts world class examples of landforms related to climate history, ooids, microbialite mounds, and evaporite minerals (e.g., halite and mirabilite). the microbialite and mirabilite features have implications for astrobiology and understanding life in extreme environments. specifically, geoheritage sites like gsl are critical to the geoscience profession, to conserve sites of geoscience importance related to earth processes, earth history, and history of geologic thought. these sites are the training ground for the next generation of environmental scientists who will grapple with global societal issues and the complexities and balance of nature. the biodiversity and geodiversity of gsl and the bonneville basin make this a remarkable geoheritage jewel of utah’s west desert. acknowledgments we gratefully acknowledge professional photographer joel long for use of his great salt lake images, and reviewers mark milligan and thomas casadevall. figure 6. mirabilite mounds and terrace structures (a, b), with large, cm+ scale crystals growing in cold colorful pools (c, d colored green by cyanobacteria) at white rock bay, antelope island state park. these mineralogies have important implications for life in extreme environments. winter images (2020): a-c: m. chan. image d: d. eby. 10 m.a. chan, c.g. oviatt, b.k. baxter, b. tikoff, and g. atwood the holocene great salt lake and pleistocene lake bonneville system figure 7. gsl is a delicately balanced ecosystem. waterfowl at farmington bay wildlife refuge on the eastern edge of gsl include migratory populations of tundra swans (a) and phalaropes (b). c: a male gsl brine shrimp with impressive claspers; brine shrimp produce eggs/cysts that are harvested from the lake and sold in aquaculture shops (e.g., fish food), and they provide food for migratory birds. d: brine-fly detritus, including pupae cases along the beach of bridger bay, antelope island state park. images a, b: j. long. c: bridget dopp. d: m. chan. 11 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 references 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c.m., and vanden berg, m.d., 2024, use of remote imagery to map microbialite distribution at great salt lake, utah: implications for microbialite exposure, in vanden berg, m.d., ford, r., frantz, c. hurlow, h., gunderson, k., and atwood, g. (eds.), great salt lake and the bonneville basin: geologic history and anthropocene issues: utah geological association publication 51. https://wildlife.utah.gov/gslep/wildlife/birds.html https://www.sltrib.com/news/environment/2022/11/08/great-salt-lakes-ecological/ https://www.sltrib.com/news/environment/2022/11/08/great-salt-lakes-ecological/ https://doi.org/10.2110/001c.56183 https://doi.org/10.1111/j.1365-3091.1975.tb00244.x rowland record low water surface elevations.pub 1 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2023 utah geological association publication 51 abstract the united states geological survey (usgs) operates two long-term water-surface elevation (wse) gages on great salt lake, utah, one north of the union pacific railroad causeway in the historic little valley boat harbor (saline gage), and one south of the causeway in the harbor at great salt lake state park (saltair gage). from september 28 to december 15, 2022, lake levels were too low in the harbor for the saltair gage to operate and wse data was measured at the south causeway gage, a relatively new gaging station (installed in 2020) located immediately south of the causeway. data collected at the south causeway gage were used to estimate the daily mean wse record for the saltair gage for the period it was shut down, preserving the continuity of the 175-year wse record that is associated with this gage. the long-standing historic low daily mean wse measured at the saltair gage on october 15, 1963 (4,191.35 feet, relative to the national geodetic vertical datum of 1929 (ngvd29)) was broken on july 21, 2021. seasonal lake-level declines from july 2021 to october 2021 and april 2022 to early november 2022 resulted in a new historic low daily mean wse of 4,188.5 feet ngvd29, measured during several days during november 2022 at the south causeway gage. the same value is also the new historic low daily mean wse for the saline gage and was measured during several days in november and december 2022 (the previous historic low of 4,188.98 feet ngvd29 was measured in september and october 2016 and was related to closure of two railroad causeway culverts). usgs also operates streamgages on major surface-water inflows including the bear river, weber river, jordan river, and surplus canal. the combined annual discharge measured at these gages in water years 2021 and 2022 was 0.704 and 0.743 million acre-feet, respectively, which is less than half of the combined median annual discharge (1.57 million acre-feet) based on the period of record for each gage. record low water-surface elevations at great salt lake, utah, 2021–2022 ryan c. rowland and mike l. freeman u.s. geological survey, utah water science center, west valley city, utah, rrowland@usgs.gov introduction the united states geological survey (usgs), in cooperation with the utah department of natural resources, operates two long-term water-surface elevation (wse) gages on great salt lake (gsl), utah (figure 1). usgs station 10010000 great salt lake at saltair boat harbor, ut (saltair gage) (u.s. geological survey, 2023), is located about 35 miles south of the union pacific railroad causeway (referred to as the causeway in the rest of this document) in the harbor at gsl state park. this gage is associated with a wse data record dating back to 1847. the record from 1847 to 1874 was compiled by grove karl gilbert, first chief geologist of the usgs, and is based on oral reports from stockmen who had ridden horses across sandbars to reach antelope and stansbury islands (gilbert, 1890; arnow and stephens, 1990). the accuracy of the early measurements does not compare to those made with modern methods (for example, arnow and stephens (1990) state that water levels from 1847 to 1874 should be considered accurate only to within 1 foot (ft)); however, this does not detract from the scientific value of those early observations. systematic lake level measurements at gsl began in 1875 as described by gilbert (1890): “in the year 1875, dr. john r. park, of salt lake city, at the suggestion of prof. joseph henry of the smithsonian institution and with the cooperation of other citizens, instituted a series of observations. there was erected at the water’s edge at black rock a granite block cut in the form of an obelisk and engraved on one side with a scale of feet and inches; and mr. john t. mitchell was engaged to observe the water-height at intervals of a few days.” from 1875 to 1938, the lake level was measured at staff gages by many different individuals and organizations at variable intervals ranging from weekly to monthly. since 1939, lake levels associated with the saltair gage have been measured continuously with various recorder devices operated by the usgs. the saltair gage has been moved several times within gsl state park because of storm damage, rebuilding of the harbor dikes, high lake levels, and low lake levels. from september 28 to december 15, 2022, lake levels were too low in the harbor for the saltair gage to operate. during this period, wses south of the causeway were obtained from usgs station 10.31711/ugap.v51i.135 2 r.c. rowland and m.l. freeman record low water-surface elevations at great salt lake, utah, 2021–2022 figure 1. locations of selected united states geological survey gaging stations (u.s. geological survey, 2023) at and near great salt lake, utah. base from maxar imagery digital data, various scales, 2019-2022. universal transverse mercator projection, zone 12 n, north american datum of 1983. 3 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 10010024 gsl south side of causeway, 6 miles east of lakeside, utah (referred to as the south causeway gage in the rest of this document) (u.s. geological survey, 2023), a relatively new gage installed in 2020. the second long-term wse gage is usgs station 10010100 great salt lake near saline, ut (saline gage) (u.s. geological survey, 2023), is in the historic little valley boat harbor on the west side of promontory point, about 2.7 miles north of the causeway. this gage was installed in 1966, about 7 years after the causeway was completed (figure 1). watersurface elevations at this gage have been measured continuously with various recorder devices operated by the usgs. it has only been moved once within the harbor (in may 1996) so that the pier it was mounted to could be removed by the owner. the usgs also operates four long-term streamgages on major surface-water inflow sources to the south part of gsl (collectively referred to as inflow gages in this document). while these gages do not measure all surface-water inflows to gsl (there are several unmeasured surface-water inflows), and also unmeasured losses or gains of water between the gages and gsl, they provide important insight into gsl wse changes over time. these inflow gages are usgs stations 10126000 bear river near corinne, ut (bear river gage); 10141000 weber river near plain city, ut (weber river gage); 10171000 jordan river at 1700 south at salt lake city, ut (jordan river gage); and 10170500 surplus canal at salt lake city, ut (surplus canal gage) (u.s. geological survey. 2023) (figure 1). the weber river gage is among the 10 oldest streamgages in utah and has been active since october 1907, with some discretedischarge measurements starting in 1904. the other inflow gages were installed in the 1940s. the primary objective of this document is to summarize selected data from the locations listed above through november 2022, which includes the lowest daily mean wse measured at gsl. this document 1) reports how record low wses at gsl were measured and validated; 2) summarizes extended periods of wse and inflow gage data; 3) compares wses to inflow gage data; and 4) compares wse and inflow gage data to a standardized measure of drought severity in utah. the locations of usgs monitoring stations discussed in this document are shown in figure 1 and a summary of parameters used in this document, including the period of record associated with each parameter, are summarized in table 1. data summarized in table 1 are available via the usgs national water information system (nwis) (u.s. geological survey, 2023). station name station number parameter and units available period of record great salt lake at saltair boat harbor, utah 10010000 daily and annual mean watersurface elevation above national geodetic vertical datum of 1929, in feet 10/15/1847 to current great salt lake near saline, utah 10010100 daily and annual mean watersurface elevation above national geodetic vertical datum of 1929, in feet 4/15/1966 to current great salt lake south side of causeway, 6 miles east of lakeside, utah 10010024 daily mean watersurface elevation above national geodetic vertical datum of 1929, in feet 2/25/2020 to current bear river near corinne, utah 10126000 daily and monthly mean discharge, in cubic feet per second 10/1/1949 to 9/29/1957 and 10/1/1963 to current (no data 9/30/1957 to 9/30/1963) weber river near plain city, utah 10141000 daily and monthly mean discharge, in cubic feet per second 10/1/1907 to current jordan river at 1700 south at salt lake city, utah 10171000 daily and monthly mean discharge, in cubic feet per second 12/1/1942 to current surplus canal at salt lake city, utah 10170500 daily and monthly mean discharge, in cubic feet per second 12/1/1942 to current table 1. u.s. geological survey (usgs) station name, number, and available period of record for data used in this report. data are available via the usgs national water information system (u.s. geological survey, 2023). 4 r.c. rowland and m.l. freeman record low water-surface elevations at great salt lake, utah, 2021–2022 description of study area great salt lake is a closed-basin lake bordered on the west by desert and on the east by the wasatch range. its abundant food and wetlands attract nearly 2 million shorebirds, including over 1.5 million grebes (podiscipedidae) and several million migrating waterfowl (wurstbaugh and others, 2017). construction of a rock-fill causeway across gsl in 1959 created two separate but connected parts of the lake with different wses, salinities, and densities resulting from more than 95 percent of all freshwater surface inflow entering the lake south of the causeway (loving and others, 2000). the differences between the wses and densities of the south and north parts of gsl provide the potential for water (gsl water is technically brine because it contains more than 35,000 milligrams per liter of dissolved solids; however, for simplicity, the term water is used in this document) to flow in both directions through the causeway conveyances. generally, the less-dense water from the south part flows northward through the upper part of the causeway conveyances (breaches and causeway fill) and the more-dense water from the north part flows southward through the lower part of the causeway conveyances (loving and others, 2000). currently, the means of conveyance include the following: a 290 ft wide breach (often referred to as the lakeside breach) near the west end of the causeway that was completed in 1984 with a bottom elevation of 4,200 ft that was lowered to 4,193 ft relative to the national geodetic vertical datum of 1929 (ngvd29) in 2000; a relatively new 150 ft wide breach about 4.5 miles from the west end of the causeway that was opened on december 1, 2016, with an adjustable berm that has a current top elevation of 4,192 ft ngvd29; and the permeability of the rockfill material used to construct the causeway. the 150 ft wide breach completed in 2016 replaced two culverts, referred to as the east and the west culverts, that were in service from causeway completion in 1959 until closure in november 2012 (east culvert) and december 2013 (west culvert). methods measurement of water-surface elevation gaging stations located on gsl are used to measure wse. the measurement of wse at gsl follows usgs protocols outlined in sauer and turnipseed (2010) which details the measurement of stage. in summary, each gaging station has a network of reference points and reference marks. these reference locations are surveyed, using a variety of techniques, to establish an elevation relative to an assigned datum. once an elevation is assigned to these reference locations, a nonrecording reference gage can be established at the gaging station. the reference gage is used to physically measure the wse of gsl. once the wse is measured using the reference gage, a recording water-level instrument can be set up to measure the wse at a set interval relative to the reference gage wse reading. currently, wse recorders for gsl are set up to measure every 15 minutes. gaging stations are visited every 1–2 months to read the wse from the reference gage and compare those readings with the wse recorder. if a difference is observed between the reference and recorded values because of instrument drift, a correction is applied to the recorded data so that the wse is accurately reported. daily and annual mean wses discussed in this report are available via nwis (u.s. geological survey, 2023). water-surface elevation reported datum reference marks, reference gages and recording gages are all referenced to ngvd29. ngvd29 is similar in elevation to dynamic heights reported by the national geodetic survey (ngs). dynamic height values are defined by an equipotential surface allowing for accurate representation of hydrologic gradient when measuring wses over a large geographic area (meyer and others, 2006). the reporting of vertical datums using dynamic heights to accurately measure water gradients is best documented in the establishment of the international great lakes datum of 1985 (meyer and others, 2006). the equipotential surface applied to dynamic heights provides a wse that flows downhill as expected. dynamic heights for the gsl region are most accurately reported when referencing wse to ngvd29. in contrast, the more commonly used north american vertical datum 1988 (navd88) is influenced by gravitational models that can cause wses that suggest water flowing in an upstream direction when a downstream gradient is known and expected. because of the causeway and the dividing of gsl, it is important to accurately represent hydraulic gradient across the causeway and, therefore, elevation should always be reported with respect to an equipotential surface so that hydraulic gradient can accurately be measured. to accurately report wses of gsl, the stability of the gaging station’s reference to ngvd29 is verified using a variety of survey methods. the survey method used to verify vertical datum is determined by the location of the three lake gaging stations. the saltair and south causeway gaging stations are located on earthen-fill material and have shown vertical 5 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 movement in previous years related to the rising and falling wse of the gsl. as wse of gsl increases, the earthen material rises with increasing water level, whereas when the wse of the gsl decreases the earthen material subsides within the lake substrate. to maintain accurate reporting of wse at these locations, the two sites are surveyed at the peak and the trough of the annual hydrograph. by surveying at the peak and the trough, wse data are corrected based on the annual fluctuations of gsl. in contrast, the saline gage has demonstrated vertical stability and is surveyed annually to ensure the gage is reporting wse accurately. because of the complex surveying techniques required to verify vertical datum at gsl gaging stations, wse reported by the usgs gsl gages are considered to be accurate to within +/-0.10 ft of the datum in use (loving, 2002). usgs station 10010000 great salt lake at saltair boat harbor, ut to verify that this gaging station is reporting accurately to ngvd29, trigonometric and differential leveling techniques are used to carry vertical datum from ngs vertical control point c-174 (table 2) to the gaging station. starting from ngs reference mark c174, a double-run spur traverse (drst) using trigonometric leveling techniques documented in noll and rydlund (2020) is used to carry datum approximately 0.5 miles from the reference mark near kennecott smelter to the harbor at great salt lake state park. the level line from the drst establishes an elevation relative to ngvd29 to a reference mark closer to the saltair gaging station where differential leveling techniques can be used to verify the datum of the reference gage (kenney, 2010). if the reference gage has moved (+/-0.05 ft) a datum correction is applied to the wse record of the gage to correct for movement of the reference gage. usgs station 10010024 gsl south side of causeway, 6 miles east of lakeside, ut  to verify that this gaging station is reporting accurately to ngvd29, survey-grade global navigational satellite systems (gnss), trigonometric, and differential leveling techniques are used to verify vertical datum. to begin datum verification, gnss static survey techniques outlined in rydlund and others (2012) are performed on four independent benchmarks near the causeway. the four independent benchmarks, documented in table 2, are occupied for a minimum of 2 hours with all four static surveys overlapping in time for a minimum of 1 hour. once the static survey is complete, a ngs online positioning user service (opus) project is performed to verify that the four independent benchmarks are stable and to determine the elevation of the reference mark (rm4) at the causeway bridge. once the elevation of rm4 is verified, the navd88 elevation from the opus project is converted to ngvd29 using ngs vertcon (national oceanic and atmospheric administration, 2023a). a drst is then performed to carry the datum approximately 0.3 miles to an established reference mark near the south causeway gaging station. the level line from the drst establishes an elevation relative to ngvd29 to a reference mark closer to the south causeway gaging station where differential leveling techniques can be used to verify the datum of the reference gage. if the reference gage has moved (+/-0.05 ft) a datum correction is applied to the wse record of the gage to correct for movement of the reference gage. usgs station 10010100 great salt lake near saline, ut  in 2009, differential leveling techniques were used to carry ngvd29 vertical datum from ngs benchmark fmk-77 (table 2) to this gaging station. the level loop was approximately 1.0 mile long and predated the trigonometric leveling techniques used at other gsl gaging stations. differential levels carried ngvd29 vertical datum to three independent benchmarks near the saline gaging station which have remained stable as referenced in kenney (2010). differential levels are run annually to verify the reference gage at the saline gaging station. if the reference gage has moved (+/-0.05 ft) a datum correction is applied to the wse record of the gage to correct for movement of the reference gage. benchmark name latitude longitude ngvd29 (ft) navd88 (ft) usgs gage c-174 40° 43' 34.00" 112° 12' 19.00" 4230.63 4233.87 10010000 77-fmk 41° 14' 33.21" 112° 29' 28.95" 4231.16 4234.20 10010100 and 10010024 moore 41° 14' 50.06" 112° 15' 33.03" 4237.63 4240.60 10010024 120-fmk 41° 13' 10.04" 112° 51' 07.35" 4223.31 4226.23 10010024 rm4 41° 13' 15.49" 112° 45' 56.49" 4216.02 4218.95 10010024 table 2. benchmarks used to maintain vertical datum at u.s. geological survey great salt lake gaging stations (u.s. geological survey, 2023). 6 r.c. rowland and m.l. freeman record low water-surface elevations at great salt lake, utah, 2021–2022 estimating water-surface elevations for the saltair gage from september 2022 to december 2022 the harbor at great salt lake state park was mostly dry and the usgs saltair gaging station could not measure wse of gsl. as a result, the elevation record was estimated for the saltair gaging station by comparing hydrographs with the south causeway gage. the two gaging stations have a nearly identical hydrograph during calm conditions. because of the location of the gages in the south part of gsl, the two gages can exhibit inverse hydrographs during wind-driven lake seiches. with the south causeway gage located on the north end of the south part, and the saltair gage being located on the southern tip of the south part, when a lake seiche occurs, one gage will have an elevated wse whereas the opposing gage will have a suppressed wse. figure 2 provides a time-series comparison of the two gages and the inverse wse observed during higher lake levels in may 2021. considering the inverted relationship, when estimating the wse for the saltair gage, the wses associated with seiche events were estimated to account for the high and low water levels that most likely occurred during the storm events. measurement of discharge a streamgage is a structure that contains equipment that measures and records the water level of a stream. the water level of a stream is often referred to as gage height or stage, reported in feet, and is measured using methods outlined in sauer and turnipseed (2010). stage is typically recorded by an instrument at a set interval ranging from 5 to 15 minutes. the continuous record of stage is then used as a surrogate to compute discharge in cubic feet per second (cfs). to compute and report discharge at a given stage, discharge measurements are made at a variety of stages to cover low, medium, and high flow conditions. discharge measurements at all stages follow methods outlined in turnipseed and sauer (2010). once a range of stage and discharge measurements have been made, a stage-discharge rating curve can be developed. a rating curve is a graphical representation of the relationship between stage and discharge, with the assumption that for every stage, there is a unique discharge. once a stage-discharge rating curve is established for a streamgage station, the continuously recorded stage at the streamgage can be used to compute a continuous discharge record. the stage-discharge method for computing discharge is applicable to gaging stations 10141000 (weber river gage) and 10126000 (bear river gage) (u.s. geological survey, 2023). the stage-discharge relationship becomes inaccurate when backwater conditions occur. backwater conditions cause the stage-discharge relationship to fail because the same discharge can occur at a range in stage values due to the backwater conditions. if backwater conditions exist at a streamgage, discharge can be computed using an index velocity method (levesque and oberg, 2012). index velocity methods require that, in addition to continuously measured figure 2. example of seiche event impact on 15-minute interval water-surface elevations measured at usgs station 10010000 great salt lake at saltair boat harbor, utah (saltair gage), and usgs station 10010024 gsl south side of causeway, 6 miles east of lakeside, utah (south causeway gage) (u.s. geological survey, 2023). 7 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 stage, a velocity sensor is installed at the stream to continuously measure water velocity at the same measurement interval as the stage sensor (5 to 15 minutes). discharge measurements are made over a range in stage and velocity to develop a mathematical relationship between measured indexed velocity and the mean channel velocity at the streamgage. once this relationship is established, the measured index velocity is used to compute a mean velocity for the channel. the velocity is multiplied by a known crosssectional area (computed from the stage value and documented channel geometry) to compute a continuous discharge at the streamgage. the index velocity method for continuous discharge is applicable to gaging stations 10171000 (jordan river gage) and 10170500 (surplus canal gage) (u.s. geological survey, 2023). most streamgage stations are located on natural channels which are subject to changes over time. these changes can be seasonally influenced or occur over several years. streamgage stations are visited routinely throughout the year to verify accurate stage data and to maintain an accurate stage-discharge or index velocity relationship. annual discharge values discussed in this report are in units millions of acre-feet (maf). these values were computed for each inflow gage by downloading daily mean discharge values, in cfs, from nwis and converting these values to daily discharge, in acrefeet per day, followed by summing these values for each water year of interest. monthly mean discharge values, in cfs, also are discussed in this report and are available via nwis (u.s. geological survey, 2023). data presentation and discussion water-surface elevations figures 3 shows the complete period of record of daily mean wse for the saltair gage (a daily mean value is the average of the recorder values logged each day; periodic wse observations made by individuals at gsl prior to installation of recorders are considered daily mean values). for detailed descriptions of the early record before june 1986, see arnow (1984) and arnow and stephens (1990); it is briefly summarized below. the high stands in the 1870s and 1980s are prominent features of the early saltair gage record along with a succession of low stands in the early 1900s, 1930s, and early 1960s. seasonal variation, where lake levels increase from approximately late autumn to late spring and decrease from approximately early summer to mid-autumn, becomes more apparent in the record after systematic measurements began in 1875. seasonal variation is driven by the balance between inflows and evaporation where lake levels increase when inflow exceeds evaporation and decrease when evaporation exceeds inflow (arnow and stephens, 1990). until july 2021, the record low mean daily wse was 4,191.35 ft, measured at the saltair gage on october 15, 1963. at the time of the record low in 1963, many people thought the lake was going to become dry and roads, railroads, wildfowl management areas, recreational facilities, and industrial installations were established on the exposed figure 3. daily mean water-surface elevation at usgs station 10010000 great salt lake at saltair boat harbor, utah (saltair gage), 1847-2022 (u.s. geological survey, 2023). 8 r.c. rowland and m.l. freeman record low water-surface elevations at great salt lake, utah, 2021–2022 lakebed (arnow, 1984). from the low in 1963 to 1976 lake levels increased about 11 ft leading to discussions about pumping water from the lake into the undeveloped desert west of gsl, but in 1977 lake levels began to decline ending concerns about high water (arnow, 1984). from september 1982 to its historic peak on june 3, 1986 (4,211.60 ft ngvd29), lake levels had a net rise of about 12 ft. this period of rapid rise culminated in $240 million in flood damages and prompted completion of the lakeside breach in the causeway in august 1984 to help decrease the approximate +3.5 ft wse difference between the south and north parts of the lake. it also prompted completion of the west desert pumping project in june 1987. the pumps associated with that project were shut down on june 30, 1989, after pumping 2.2 million acre-feet of water from gsl into the west desert pond, which reduced gsl’s wse by about 2.2 feet (austin, 2002). a plot of daily mean wses for the saltair and saline gages from june 1986 to november 2022 is shown in figure 4. the difference between the wses for the two gages depends on factors such as inflows, densities of the south and north parts of gsl (which provides the potential for gsl water to flow in both directions through the causeway conveyances), evaporation, and modifications to causeway conveyances. it is beyond the scope of this paper to discuss each of these factors in detail; however, modifications were made to the causeway conveyances during the periods that are associated with observed wse differences (figure 4). because almost all surface-water inflow is to the south part of the lake, wses at saltair are usually higher than saline (median value for period shown in plot is +0.7 ft). the increased difference in wse between saltair and saline from november 1991 through january 1998 occurred during an extended period when the culverts were frequently plugged with debris (loving, 2002). the effective depth of the lakeside breach was deepened from about 4,200 ft to 4,198 ft ngvd29 in august 1996 (loving, 2002), which likely contributed to the subsequent reduction in wse difference between saltair and saline from 1996 to 1998. the increased wse difference from september 2014 to february 2017 is associated with closure of the east (november 2012) and west (december 2013) culverts. the rapid decrease in wse difference from december 2016 to june 2017 is associated with the opening of the new breach on december 1, 2016. this breach has an adjustable berm on the north side of the causeway. to help manage the salinity in the southern half of the lake the top of the berm was raised from 4,183 to 4,187 ft ngvd29 (completed july 27, 2022) and from 4,187 to 4,192 ft ngvd29 (completed february 9, 2023). the latter modification raised the top of the berm above the wse of the south part of the lake at the time and contributed to the increased wse difference after february 9, 2023. the magnitude of seasonal fluctuations in the daily mean wse record from 1986 to 2022 are shown in figure 5. at saltair, the average seasonal increase is 1.8 ft and the average seasonal decrease is 2.4 ft. the largest seasonal lake level increase (5.1 ft) occurred from autumn 2010 to late spring 2011. the largest seasonal lake level decrease (3.2 ft) occurred during figure 4. daily mean water-surface elevation at usgs station 10010000 great salt lake at saltair boat harbor, utah (saltair gage), and usgs station 10010100 great salt lake near saline, utah (saline gage), 1985-2022 (u.s. geological survey, 2023). 9 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 spring and summer 1988 and 2018. lake level decreases exceeded increases in 25 of the 36 seasonal cycles shown. at the saline gage, the average seasonal increase is 1.5 ft and the average seasonal decrease is 2.1 ft. the largest seasonal lake level increase (5.5 ft) occurred from december 2016 to may 2017, after the new 150 ft wide breach was opened restoring open channel connection between the south and north parts of the lake that was previously associated with the east and west culverts. the largest seasonal lake level decrease (3.3 ft) occurred during spring and summer 1988. lake level decreases exceeded increases in 22 of the 36 seasonal cycles shown in figure 5. with seasonal decreases exceeding increases for most years following the record high wse in june 1986, both the north and south parts of the lake had net wse drops that resulted in record low wses in november 2022 (figure 4). the long-standing historic figure 5. seasonal water-surface elevation increase and decrease at usgs station 10010000 great salt lake at saltair boat harbor, utah (saltair gage, top) and usgs station 10010100 great salt lake near saline, utah (saline gage, bottom), 1986 to 2022 (u.s. geological survey, 2023). seasonal increases generally occur from late fall/early winter through the following spring/early summer. seasonal decreases generally occur from late spring/early summer through mid fall/early winter. 10 r.c. rowland and m.l. freeman record low water-surface elevations at great salt lake, utah, 2021–2022 low daily mean wse measured at the saltair gage on october 16, 1963 (4,191.35 ft ngvd29) was broken on july 21, 2021. the wse continued to decrease until october 18, 2021, when it reached a short-lived historic low of 4,190.2 ft ngvd29. the seasonal lake level increase from october 18, 2021, to early april 2022, was relatively low at 1.2 ft, and by july 3, 2022, the wse dropped to 4,190.1 ft ngvd29, breaking the short-lived historic low set less than 9 months prior. by september 28, 2022, continued seasonal decrease resulted in too little water in the harbor at great salt lake state park for the saltair gage to operate and it was shutdown. water-surface elevation data for the south part of the lake continued to be measured at the relatively new (installed august 2020) south causeway gage, maintaining continuity of the 175-year wse record that is associated with the south part of gsl. seasonal decreases continued until november 3, 2022, when the south causeway gage recorded the new record low daily mean wse for the south part of the lake of 4,188.5 ft ngvd29. two days prior, on november 1, 2022, the saline gage recorded 4,188.5 ft ngvd29, which also is the new historic low daily mean wse for the north part of the lake (the previous historic low of 4,188.98 ft ngvd29 was measured in september and october 2016 and was related to closure of the two railroad causeway culverts). net wse decreases from june 1986 to november 2022 for the south and north parts of the lake were 23.1 ft and 22.7 ft, respectively. by late november 2022, the south part of the lake began its seasonal increase (figure 4). the saltair gage was restarted on december 15, 2022, and, as of may 4, 2023, the south part had risen to 4,192.6 ft ngvd29, a 4.1 ft increase. the north part of the lake did not start increasing until late december 2022 and, as of may 4, 2023, it had risen to 4,189.3 ft ngvd29, a 0.8 ft increase. the wse of the south part of the lake reached 4,192.1 ft ngvd29 on april 17, 2023, exceeding the top of the berm at the new breach allowing for south to north flow (on may 3, 2023, usgs measured south to north discharge of 129 cfs). with a significant snowpack remaining in the bear, weber, and jordan river basins, south to north flows were expected to continue until evaporation exceeds inflows and lake levels begin their seasonal decrease. streamflow and great salt lake water surface elevation annual discharge for water years 1985 to 2022 for each inflow gage are shown in figure 6. also included in figure 6 are the median annual discharge values for each inflow gage. the median annual discharge values are based on the period or record associated with each gage (table 2). the bear river gage has the highest median annual discharge (0.958 maf), followed by the weber river gage (0.343 maf), surplus canal gage (0.170 maf), and jordan river gage (0.101 maf). relatively high flow years in the mid1980s, late 1990s, 2011 and 2017; and relatively low flow years in the late 1980s, early 1990s, mid-2010s, and early 2020s are apparent in the data for the bear river, weber river, and surplus canal gages (note that a significant amount of flow in the jordan river is diverted to the surplus canal). of the 36 years of annual discharge record shown for each gage, the bear river gage had 12 years where annual discharge exceeded its median annual discharge, the weber river gage had 10 years, the surplus canal gage had 21 years, and the jordan river gage had 12 years. annual discharge measured at the inflow gages and annual mean wse measured at the saltair gage for water years 1985 to 2022 are shown in figure 7. annual mean wse is related to annual discharge as consecutive years of relatively high flows from 19851987, 1995-1999, 2005–2006, and 2011–2012 contributed to annual mean wse increases. consecutive years of relatively low flows from 1988–1994, 2001– 2004, 2013–2016, and 2021–2022 contributed to annual mean wse decreases. during water years 2021 and 2022, when new record low wses were measured at gsl, combined annual discharge values were 0.704 and 0.743 maf, which are less than half of the median combined annual discharge of 1.57 maf (indicated by the horizontal dashed line in figure 7). to put the wse and discharge records into context with broader climatological conditions, monthly palmer drought severity index (pdsi) values for utah (national oceanic and atmospheric administration, 2023b), monthly mean wse at saltair, and combined monthly mean discharge for inflow gages, in cfs, are plotted in figure 8. the pdsi uses precipitation and temperature data to evaluate moisture supply and demand using a simple water balance model. a pdsi value of greater than 4 represents very wet conditions, while a pdsi less than -4 represents an extreme drought. extended periods of negative pdsi values from november 1988 to november 1992, october 1999 to august 2004, november 2006 to november 2009, november 2011 to august 2016, and august 2019 to october 2022 correspond to net wse declines and lower monthly mean discharge. these periods of extended negative pdsi are offset by relatively few periods of extended positive pdsi values and increased monthly mean wse and higher combined monthly mean discharge. 11 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 6. annual discharge measured at usgs inflow gages, water years 1985-2022 (u.s. geological survey, 2023). the median annual discharge for the period of record associated with each gage is indicated with a dashed line. the yaxis scale, discharge in millions of acre-feet, is customized for each gage. 12 r.c. rowland and m.l. freeman record low water-surface elevations at great salt lake, utah, 2021–2022 conclusions a new historic low daily mean wse of 4,188.5 ft ngvd29 was measured during november 2022 at gages north and south of the gsl causeway. from september 28 to december 15, 2022, there was too little water in the harbor at great salt lake state park for the saltair gage to operate and the new historic low daily mean wse for the south part of the lake was measured at a relatively new gage located just south of the causeway (south causeway gage). data collected at the south causeway gage were used to estimate the daily mean wse record for the saltair gage for the period it was shut down, preserving the continuity of the 175-year wse record that is associated with this gage. many factors, including direct precipitation, groundwater inflow, west desert pumping project withdrawals (1987–1989), evaporation, and surfacewater inflow contribute to the water balance and thus wse of gsl. in this document, data were presented only for a portion of the surface-water inflow budget as measured by four long-term streamgages. for water years 1985 to 2022, trends in saltair gage wses correspond to trends in combined annual discharge measured at the four streamgages, and both wse and combined monthly mean discharge correspond to trends in pdsi values for utah. this basic observation is true when the data records are examined back to 1950 when concurrent monitoring began at all sites (cordova and angeroth, 2012). for detailed analyses of gsl wse variation and climate, see wang and others (2010) and mann and others (1995). the impact of upstream diversions from surface water inflow sources to gsl is beyond the scope of this document; however, wurtsbaugh and others (2017) estimated that 11 ft of gsl wse decrease from 1847 to 2016 can be attributed to consumptive use and related changes to evaporation. detailed monitoring of gsl’s water budget may support quantification of the complex interplay between drought cycles, consumptive use, and wses. figure 7. annual discharge for usgs inflow gages and annual mean water surface elevation at usgs station 10010000 great salt lake at saltair boat harbor, utah (saltair gage), water years 1985-2022 (u.s. geological survey, 2023). the combined median annual discharge for all four gages, based on the period of record for each gage, is indicated by the horizontal dashed line. 13 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 any use of trade, firm, or product names is for descriptive purposes only and does not imply endorsement by the u.s. government. references austin, l.h., 2002, problems and management alternatives related to the selection and construction of the west desert pumping project, in gwynn, j.w., editor, great salt lake an overview of change, special publication of the utah department of natural resources, p. 303-312. arnow, t., 1984, water-level and water-quality changes in great salt lake, utah, 1847-1983: u.s. geological survey circular 913, 22p. arnow, t. and stephens, d., 1990, hydrologic characteristics of the great salt lake, utah, 18471986: united states geological survey watersupply paper 2332, 32 p. cordova, j.t. 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https://www.ncei.noaa.gov/access/monitoring/climate-at-a-glance/statewide/time-series https://www.ncei.noaa.gov/access/monitoring/climate-at-a-glance/statewide/time-series https://www.ncei.noaa.gov/access/monitoring/climate-at-a-glance/statewide/time-series http://pubs.usgs.gov/tm/tm3-a7/ http://pubs.usgs.gov/tm/tm3-a7/ http://pubs.usgs.gov/tm/tm3-a8/ http://pubs.usgs.gov/tm/tm3-a8/ https://waterdata.usgs.gov/nwis https://www.nature.com/articles/ngeo3052 << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /none /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (srgb iec61966-2.1) /calcmykprofile (u.s. web coated \050swop\051 v2) /srgbprofile (srgb iec61966-2.1) /cannotembedfontpolicy /error /compatibilitylevel 1.4 /compressobjects /tags /compresspages true /convertimagestoindexed true /passthroughjpegimages true /createjobticket false /defaultrenderingintent /default /detectblends true /detectcurves 0.0000 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/hrv (za stvaranje adobe pdf dokumenata najpogodnijih za visokokvalitetni ispis prije tiskanja koristite ove postavke. stvoreni pdf dokumenti mogu se otvoriti acrobat i adobe reader 5.0 i kasnijim verzijama.) /hun 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can be opened with acrobat and adobe reader 5.0 and later.) >> /namespace [ (adobe) (common) (1.0) ] /othernamespaces [ << /asreaderspreads false /cropimagestoframes true /errorcontrol /warnandcontinue /flattenerignorespreadoverrides false /includeguidesgrids false /includenonprinting false /includeslug false /namespace [ (adobe) (indesign) (4.0) ] /omitplacedbitmaps false /omitplacedeps false /omitplacedpdf false /simulateoverprint /legacy >> << /addbleedmarks false /addcolorbars false /addcropmarks false /addpageinfo false /addregmarks false /convertcolors /converttocmyk /destinationprofilename () /destinationprofileselector /documentcmyk /downsample16bitimages true /flattenerpreset << /presetselector /mediumresolution >> /formelements false /generatestructure false /includebookmarks false /includehyperlinks false /includeinteractive false /includelayers false /includeprofiles false /multimediahandling /useobjectsettings /namespace [ (adobe) (creativesuite) (2.0) ] /pdfxoutputintentprofileselector /documentcmyk /preserveediting true /untaggedcmykhandling /leaveuntagged /untaggedrgbhandling /usedocumentprofile /usedocumentbleed false >> ] >> setdistillerparams << /hwresolution [2400 2400] /pagesize [612.000 792.000] >> setpagedevice atwoodshorelinesuperelevation.pub 1 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2023 utah geological association publication 51 abstract coastal processes create the shoreline evidence of great salt lake. shoreline superelevation is the difference in elevation between still water lake level and the shoreline evidence produced by the lake at that level. processes of formation include effects of wind strength, fetch, beach attributes, coastline aspect, and coast morphology. a series of field studies from 1986 through 2000 concluded strong storm winds from the northwest contribute to the patterns and magnitude of shoreline superelevation. weather data for 2020-2023 for gunnison island and hat island document strong storm winds from the north and northwest for gunnison bay and with more complexity for gilbert bay. the strongest wind patterns are consistent with the geologic evidence of shoreline superelevation produced by the high lake stands of 1986-1987. wind strength, fetch, and storm duration cause great salt lake wave regimes. the wave-regimes of great salt lake are fetch-limited due to the size and morphology of the water body. in contrast, the long fetch of large lakes such as lake bonneville (the enlarged manifestation of the great salt lake lacustrine system), determines the magnitude and patterns of their shoreline superelevation. geologic evidence of shoreline superelevation of modernand paleofetch-limited lakes similar to great salt lake may be durable evidence of storm wind direction. shoreline superelevation, clues to coastal processes of great salt lake genevieve atwood1, tamara j. wambeam2, and charles g. oviatt3 1earth science education, salt lake city, utah, genevieveatwood@comcast.net 2salt lake city, utah 3department of geology, kansas state university, manhattan, kansas introduction great salt lake (gsl) is a closed-basin lake located in the lowest region of the gsl drainage basin, and it has no surface outlet (figure 1). its shorelines and lake bottom sediments record lake conditions. gsl shoreline elevation fluctuates as the lake’s volume fluctuates in response to the balance of water entering the lake by direct precipitation and runoff, and water leaving the lake by evapotranspiration. therefore, patterns of shoreline elevations are interpreted as patterns of climate. understanding the chronology of lake fluctuations underpins interpretations of changed climate over time. however, the details of the history of climate changes have not yet been deciphered for post-lake bonneville time from the geomorphic and stratigraphic records (oviatt and others, 2021). shoreline materials also contribute to the understanding of lake processes (gilbert, 1890). terrigenous materials deposited by waves become the geologic record. if a paleoshoreline defines a horizontal plane, it can be used to distinguish post-depositional change. examples of the use of this assumption include studies of isostatic rebound, tectonic displacements, and effects of wind and waves (gilbert, 1890; tackman, 1993; adams and wesnousky, 1998; tackman and others, 1998; adams and others, 1999; adams and bills, 2016; and chen and maloof, 2017). however, should initial shoreline conditions not define a horizontal plane, the original non-horizontality introduces uncertainty to interpretations (gilbert, 1890, currey, 1982). this paper summarizes a series of field studies documenting the shoreline left by utah’s 1980s wet cycle (1982-1987). in 1986 and again in 1987, gsl reached its historic highstand elevation, 4212.15 ft (arnow and stephens, 1990). it left pristine, undisturbed, continuous evidence around the perimeter of antelope island as lines of organic and inorganic debris. this paper explores the coastal processes that caused the original non-horizontality of the 19861987 shoreline. wind waves that are higher and more energetic in some places than others cause patterns of shoreline superelevation. the following definitions contribute to understanding “superelevation” (figure 2). lake setup is “elevated lake surface caused by any process whether or not storm-related.” wind setup is “the component of lake setup caused by wind” and is accompanied by lake setdown, a lowered lake level. lake seiche is “the oscillation of the lake’s surface initiated by lake setup.” wave runup is “the rush of water with entrained sediment landward and upward to 10.31711/ugap.v51i.138   g. atwood, t.j. wambeam, and c.g. oviatt shoreline superelevation, clues to coastal processes of great salt lake 2 figure 1. location maps for great salt lake and antelope island. adapted from atwood (2006) (a). great salt lake: place names include bays of great salt lake and major islands: antelope island (ai), carrington island (ci), fremont island (fi), gunnison island (gi), hat island (hi), and stansbury island (si). lake-level monitoring gages are saltair marina boat harbor (bh), promontory (pr), and saline (sa). names of communities are shown in italics. the dark line indicates the extent of 1986-1987 highstand flooding. (b). antelope island: formal and informal names (in italics) for locations of shoreline superelevation surveyed during 1997-1998. 3 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 its highest shoreline expression.” wave runup is the highest elevation reached by waves, and the entrained sediment deposited by the waves provides a record of shoreline superelevation. shoreline superelevation is the difference in elevation between the shoreline evidence and the independently monitored still water elevation. (atwood 2006). earlier field studies documented patterns of shoreline superelevation and suggested that the patterns were the effects of wind strength and direction as well as of fetch (distance across open water). this paper reports how present-day meteorological data supplement the findings of field surveys of 1986-2000 (atwood, 2006), which did not have the advantage of 2020-2023 records from weather stations on hat and gunnison islands. the strongest winds across gsl blow from the north and west, corroborating the geomorphic evidence. patterns of shoreline superelevation document the effects of wind strength and direction because gsl is fetch-limited. “fetch-limited” refers to water bodies where the size of the wave generation area limits wave height and energy. methods and data purpose and methods of the field surveys, 1986-2000 several surveys conducted between 1986 and 2000 by d.r. currey, d.r. mabey and g. atwood provide field-based data for the present paper. a summary of methods, data, and results is given below and is set out fully in atwood (2006). shoreline features were observed, described, and their elevations were measured directly in the field during, immediately after, and in the decades following the 1986-1987 gsl highstand. unmistakable floated debris (e.g., wood, plastic, and windrows of organic matter), as well as fresh gravel ridges, identified 1986-1987 shoreline evidence that persisted for over a decade (figure 3). surveyed shoreline debris defined shoreline superelevation patterns. over the past four decades, some of that evidence has degraded, but gravel ridges remain in many places where they can be spotted by their vegetation (sunflowers). 1986 survey currey and mabey on the eastern shore of antelope island the purpose of the 1986 survey was to repeat g.k. gilbert’s survey in 1877 of the evidence of the 1870s highstand shoreline (gilbert, 1890). mabey and currey (mabey, 1986), concerned that the rising lake would rework and destroy the 1870s evidence, repeated gilbert’s survey on the east shore of antelope island using hand-held equipment similar to gilbert’s era. they identified three places on aerial photographs and surveyed them on the ground using the united states geological survey (usgs)-monitored still water level for vertical control. in the century between gilbert’s survey in 1877, and the work of currey and mabey in 1986, the shoreline evidence had become difficult to recognize, except as patterns on aerial photographs and patches of gravel. according to mabey (1986), “in the spring of 1986 when the lake was at a level of 4211.85 ft, a storm line was formed on the east side of antelope isfigure 2. shoreline superelevation, evidence of interactions of earth systems. the schematic simplifies and summarizes diverse conditions and processes that result in shoreline superelevation. under strong winds or as storms progress, waves develop, and lake water is pushed up against windward shores. winds blow across the surface causing waves, and the waves deposit the terrigenous debris that becomes the durable geologic evidence of lake elevation. patterns of shoreline superelevation include interactions among the atmosphere, the hydrosphere, the geosphere, and the biosphere. 4 g. atwood, t.j. wambeam, and c.g. oviatt shoreline superelevation, clues to coastal processes of great salt lake figure 3. evidence of shoreline superelevation. adapted from atwood (2006). (a). the sketch illustrates the shorezone features relative to shoreline superelevation. the difference between the 1986 and 1987 usgsmonitored still water elevation (4212 ft) and the 1986-1987 highstand debris lines on antelope island is shoreline superelevation. shore features include lagoons, killed vegetation, and higher and older shorelines. (b). the photograph taken in 1998 looks east along the northern exposure of ladyfinger east. the 1986-1987 shoreline expression, foreground, includes terrigenous debris of cobbles, gravel, and sand. contrasts in vegetation patterns, the upper center of the photograph, and lumber and timber in the beach zone are evident a decade after the 1986-1987 flooding. (c). the photograph taken in 1998 looks northeast toward the intersection of the northern and southern expressions of the spit at unicorn point. the man with the rod stands on the northern, northeast-facing, lower expression, and the younger man stands on a southeast-facing expression. (d). the photograph taken in 1998 looks north from timely gull bay toward curlew bay along the west side of antelope island. note the stacked timber and lumber at the south of the bay indicating transport by wind waves from the northwest. (e). the photograph taken in 1986, during the highstand years, looks north along the eastern exposure of tin lambing shed, south of harbor bay. note the terrigenous wash-over deposits of sand and 20th-century evidence of lumber. 5 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 land at 4213.5 ft, the same elevation measured by gilbert for the storm line formed in the 1870s.” the shoreline evidence of both surveys was superelevated compared to the usgs-monitored still water elevations (gilbert, 1890; mabey, 1986). based on records of the elevation of the highstand taken along the south shore of gsl, gilbert estimated shoreline superelevation of one foot on the eastern side of antelope island. 1986 survey by atwood and mabey in 1986 g. atwood and d.r. mabey conducted a survey to compute the frequency of holocene flooding of gsl (atwood and mabey, 2000). the idea was to survey the 1986 shoreline evidence and to count the shorelines between the historic highstands (1870s and 1980s) and murchison’s (1989) “holocene high” of approximately 4217 ft. for the 1986 survey, we used an electronic measuring device (edm) to measure elevations. initially, we expected that the 19861987 highstand evidence would provide the horizontal datum from which to survey the higher, older shoreline elevations. however, the 1986-1987 shoreline evidence did not define a horizontal plane. therefore, we used only the usgs-monitored still water lake level as vertical control for the survey. the evidence of the 1986-1987 highstand was unmistakable and included debris lines of floated debris, gravel ridges and beaches, erosional steps, and vegetation lines (salt-kill zones). floated debris included 20th-century wood and anthropogenic material (such as plastic). terrigenous evidence included well-sorted cobbles, coarse and fine gravel, and sand. shoreline evidence of 1986-1987 had no observable surface staining in contrast to older shorelines. in places, erosional steps had been cut into poorly consolidated, sandy sediments. the evidence of murchison’s (1989) “holocene high” at an elevation between 4217 and 4222 ft was discontinuous, subtle, and subject to interpretation, consisting of widely scattered gravel and cobble patches and subtle breaks in slope. for detailed discussion of holocene lake fluctuations, see oviatt and others (2021). a summary of the findings of the 1986 survey on antelope island is as follows (atwood and mabey, 2000): (a) the plot of shoreline elevations (figure 4) indicated at least three highstand shorelines between the 1986-1987 highstand and 4226 ft. counting the two historic excursions to 4212 ft in 1986 and 1987, gsl had risen a minimum of five times to elevations equal to or higher than 4212 ft. (b) the 1986-1987 shoreline debris did not define a horizontal plane from which to measure relative elevations of holocene shorelines. evidence of the 1986-1987 gsl highstand was consistently superelevated, wellabove the usgs-monitored still water lake level. the elevations of the higher, older shorelines appeared to have trends of superelevation resembling those of 1986-1987. 1997-1998 atwood and mabey survey of 1986-1987 highstand evidence on antelope island the 1997-98 survey aimed to document the character of the 1986-1987 highstand shoreline in detail before ephemeral evidence became unrecognizable (atwood, 2006). we documented elevations of the 1986-1987 highstand evidence and recorded shorezone characteristics around the island's perimeter using a sokia total station to survey elevations and global positioning system (gps) data loggers to record horizontal positions of observations of shorezone characteristics. the usgs-monitored still water lake level provided vertical control. throughout the day, we used a survey staff to measure water-level changes for the purpose of maintaining accurate vertical control. we also corroborated our elevations with davis county public works’ road elevations. elevations were surveyed on the upper surface of the mostinland terrigenous deposits at 1,228 locations along the 64 km shoreline of antelope island. the data were downloaded into a geographic information system database, projected to a single route using esri arc/info linear referencing, and analyzed with simple spatial statistics (atwood, 2006). a decade after the 1986-1987 gsl highstand, much of the 1986-1987 flotsam (windrows of brinefly carapaces, vegetative evidence, and automobile tires) was lost to disintegration, fire, and trash collection. large debris, which included lumber and timber, became reliable evidence and it persisted long after deposition. some smaller debris persisted included plastic and other 20th-century debris. gravel ridges and sand beaches were intact. erosional steps were evident but no longer had the angular shape of 19861987. vegetation, specifically sunflowers, grew on the 1986-1987 gravels. patterns of shoreline superelevation along the 1986-1987 shoreline were consistent with observations of the 1986 survey. patterns of shoreline superelevation were not random. they did not define a horizontal plane from which post-depositional change could be measured with confidence. 6 g. atwood, t.j. wambeam, and c.g. oviatt shoreline superelevation, clues to coastal processes of great salt lake patterns of shoreline superelevation were compared with patterns of shorezone characteristics (figures 5 and 6). variations of shoreline superelevation from place-to-place record relative wave energy modified by diverse factors. shoreline evidence was found consistently above still water lake level because, when there is no wind, there are no waves to rework materials and deposit evidence. wind waves are the most significant agents of coastal processes that affect lake shorelines (komar, 1998). waves erode and deposit the shoreline evidence. wave height and wave energy largely determine shoreline superelevation. however, other factors affect coastal dynamics. wind setup and wind setdown due to atmospheric conditions lead to lake seiche (wang, 1978). seiche alone has little effect on shoreline erosion and deposition but may affect the magnitude of shoreline superelevation due to wind setup. interference and harmonics of normal “gravity waves” create widely spaced infra-gravity waves (bertin and others, 2020). offshore and on-shore currents affect wind wave processes and wave heights. these factors make the initial, generally higher lake levels from which waves run up the shore. shoreline superelevation records the net effect of wave energy and shorezone conditions, including aspect, fetch, steepness, and materials. aspect (the direction that the beach faces) was used as a proxy for wind direction. figures 7, 8, and 9 show contrasts of antelope island shores. high shoreline superelevation correlates with long fetch and with north and northwest aspect. low shoreline superelevation correlates with short fetch and geomorphic shielding. this observation implied that wind might be a recognizable contributing factor to shoreline superelevation of gsl in addition to the effects of fetch (figure 10). figure 4. great salt lake shoreline elevation data of the 1987-1988 survey on antelope island. plot of elevation (ft) at a progression of surveyed locations clockwise around antelope island beginning at white rock bay (wrb), continuing to lady finger point (lfw, lfe), to seagull point (sgn, sge), to unicorn point (unc, unp), to dry canyon (drc) (refer to fig. 1(b) for locations. blue dots show surveyed elevations of shoreline debris of the 1986-1987 highstand. black dots show surveyed elevations of older, higher shoreline evidence around antelope island. lines between locations indicate lateral correlation of shorelines. the blue dots, surveyed locations of 1986-1987 debris, consistently lie above 4212 ft, the elevation of the usgs-monitored still water level and documented flooding hazards above still water lake level. the lettered points (b, c, d, h) were grouped based on trends and position relative to the 1986-1987 shoreline evidence at each location. adapted from atwood and mabey, 2000. 7 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 5. superelevation of shoreline evidence of antelope island, surveyed in 1998-1999. adapted from atwood (2006). two maps of surveyed locations on antelope island showing shoreline superelevation in equal increments versus shoreline superelevation classes. (a). shoreline superelevation displayed in equal 1-ft increments above the 4200 ft datum of the field study. shoreline evidence ranged from 4211 ft to 4223 ft (11 to 23 on the key). shoreline superelevation elevations ranged from at or slightly below usgs-monitored still water level along vegetated shore stretches to the highest levels, 11 ft above still water lake level, on bedrock outcrops bordering pocket beaches. (b). shoreline superelevation classified in approximately equal populations. high superelevation is superelevation equal to or greater than 3.4 feet. intermediate superelevation is superelevation between 2.2 and 3.4 feet. low superelevation is superelevation less than 2.2 feet. each of the three classes consists of approximately 400 surveyed elevations. g. atwood, t.j. wambeam, and c.g. oviatt shoreline superelevation, clues to coastal processes of great salt lake figure 6. associations of shoreline superelevation with fetch and aspect. the two sets of maps show associations of maximum fetch, aspect in 15-degree increments, and shoreline superelevation. visual inspection indicated correlations among shoreline superelevation, fetch, and aspect. adapted from atwood (2006). 8 9 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 7. lady finger point and south point patterns: superelevation, maximum fetch, and shore aspect. the orthophotos show great salt lake near its highstand at the two red-circled locations on the map. lady finger point juts into gilbert bay as a bedrock headland. unicorn point immediately to the east of south point is named for the “unicorn” described by its 1986-1987 lagoons and spits. the table relates high shoreline superelevation, at a detailed scale with maximum fetch and shore aspect. the dots of the table entries indicate locations surveyed in the 1997-1998 antelope island survey. shoreline superelevation was surveyed, whereas maximum fetch and aspect were interpreted from maps. the patterns show west-east contrasts. at lady finger point, high shoreline superelevation correlates visually with medium fetch and western aspect. at south point, high shoreline superelevation correlates visually with maximum fetch and shore aspects facing west and southwest. 10 g. atwood, t.j. wambeam, and c.g. oviatt shoreline superelevation, clues to coastal processes of great salt lake figure 8. white rock bay and harbor bay patterns: superelevation, maximum fetch, and shore aspect. the orthophotos show great salt lake near its highstand at the two red-circled locations on the map. white rock bay, a broad, shallow bay opens to the west. harbor bay, a complex bay opens to the north and east. the dots of the table entries indicate locations surveyed in the 1997-1998 antelope island survey. shoreline superelevation was surveyed, whereas maximum fetch and aspect were interpreted from maps. the patterns show west-east contrasts. at white rock bay, high shoreline superelevation correlates visually with medium fetch and western aspect, while low shoreline superelevation does not appear to correlate with fetch or aspect and may result from sheltering by geomorphic features. harbor bay has limited high shoreline superelevation with low shoreline superelevation correlating visually with low fetch and aspects facing southeast. 11 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 9. buffalo scaffold curlew bays and ranch house south patterns: superelevation, maximum fetch, and shore aspect. the orthophotos show great salt lake near its highstand at the two red-circled locations on the map. the bays and headlands of buffalo scaffold curlew bay on the southwestern shore of the island contrast with the straight shore of ranch house south. the dots of the table entries indicate locations surveyed in the 1997-1998 antelope island survey. shoreline superelevation was surveyed, whereas maximum fetch and aspect were interpreted from maps. the patterns show west-east contrasts. at buffalo scaffold and curlew bays, high shoreline superelevation correlates visually with maximum fetch and with western aspects. at ranch house south, low shoreline superelevation correlates visually with low fetch and eastern aspect. 12 g. atwood, t.j. wambeam, and c.g. oviatt shoreline superelevation, clues to coastal processes of great salt lake as with previous surveys, elevations of shoreline evidence were not at the usgs 1986-1987 monitored still water lake level. patterns of shoreline superelevation were not random and could be quantified. for example, highest shoreline superelevation was associated with fetch greater than 55 km and on shores facing north, northwest, and west. shoreline superelevation ranged from as low as 4211.1 ft to 4223.4 ft. with a mean of 4214.5 ft. the shoreline superelevation of the west side of the island was generally higher and more variable than of the east side of the island. the patterns of shoreline superelevation of the 1986-1987 shorelines on antelope island provide evidence of the geomorphic effects of wind waves. but because both the longest fetch and the strongest winds were from the northwest, patterns of shoreline superelevation on antelope island could not clarify the relative contributions of fetch and aspect to wave energy. g.k. gilbert observed shoreline superelevation on the southern shores of lake bonneville and cautioned that fetch, not wind strength or wind direction, caused the superelevation of lake bonneville shores. gilbert (1890, p.107) expressed his recognition of the effects of long fetch in the following quote, which conveys his surprise, humility, and acceptance that long fetch, regardless of wind strength and direction, accounted for the high shoreline superelevation of lake bonneville’s shores. at an early stage of the investigation, the writer thought that the coasts facing in certain directions gave evidence of exceptional amounts of wave work, and imagined that he had discovered therein the record of prevalent westerly winds or westerly storms in ancient times. this belief was dissipated by further study; and he discovered, as students of modern shores long ago discovered, that there is a close sympathy between the magnitude of the shore features and the "fetch" of the efficient waves. the greater the distance through which waves travel to reach a given coast, the greater figure 10. associations of high superelevation, long fetch, and aspects facing north and west. series of three maps showing aspect, superelevation, and fetch. patterns of (a) west-facing shores; (b) high shoreline superelevation; and (c) longest fetch are similar. of the 400 surveyed locations with high superelevation, 86 percent have fetch ≥ 50 km; 55 percent have fetch ≥ 55 km; and 50 percent have aspect 240-290°. of the 200 surveyed locations with both high superelevation and aspect 240-290°, 199 have fetch ≥ 50 km and 140 have fetch ≥ 55 km. on antelope island, because the patterns of fetch and aspect so closely resemble each other, it is difficult, if not impossible, to distinguish the relative importance of wind from fetch on shoreline superelevation. adapted from atwood (2006). 13 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 the work accomplished by them. the highest cliffs, the broadest terraces, and the largest embankments are those wrought by the unobstructed waves of the main body; and opposite coasts appear to have been equally affected. might processes of a fetch-limited lake such as gsl at its highstand level leave long-lasting evidence of wind strength and direction and therefore lasting geomorphic clues to storm conditions and weather patterns? atwood and mabey 1999-2000 survey at places around gilbert and gunnison bays the 1999-2000 survey (atwood 2006) aimed to confirm whether patterns of shoreline evidence along the shores of gilbert and gunnison bays resembled those along the shores of antelope island. we explored relationships among fetch, aspect, and shoreline superelevation. disturbance of 1986-1987 shoreline evidence, accessibility, and inadequate vertical survey control limited the choice of locations with which to compare diverse conditions of fetch and aspect (figure 11). for the 1999-2000 survey, we followed the same procedures as for antelope island in 1997-1998. we used the same equipment, including the sokia total station and the gps data loggers. vertical control was carried from first-order survey markers and/or from usgs-monitored still water lake level. we interpreted factors of fetch such as length and direction of the longest fetch, length of fetch north and northwest, the distance from the bay axis, shorezone aspect, shorezone slope, elevation of the lake shore bed, and bedrock outcrops between 4200 and 4220 ft a.s.l. from maps. much of the non-terrigenous evidence of the 1986 -1987 shoreline had been lost to natural disintegration, land cultivation, and development onto the lakebed as gsl retreated. orthophotos documented shore features of 1986-1987. that evidence and 20thcentury debris, such as large logs and railroad ties, confirmed field identification of the 1986-1987 highstand in contrast to higher, older shorelines. the survey data were plotted on orthophotos and checked against geomorphic features. figure 12 shows contrasts of patterns of shoreline superelevation at strongs knob near the southwestern shore of gunnison bay with those of rozel point along gunnison bay’s eastern shore. the classifications of high, medium and low superelevation are those of the antelope island survey. patterns of shoreline superelevation of gilbert and gunnison bays shores resembled those of antelope island. they confirmed that shoreline superelevation was a lake-wide phenomenon. as with the findings on antelope island, patterns of shoreline superelevation were not random and were quantifiable. differences in elevation from place to place were easily detected. the 1986-1987 shoreline around gilbert and gunnison bays, just as around antelope island, did not define a horizontal plane. coastal processes of gsl cause shoreline superelevation. we used a series of steps to explore whether fetch alone caused spatial variations in shoreline superelevation of gunnison and gilbert bays (figures 13 and 14). we assumed that equal fetch causes patterns of equal shoreline superelevation. fetch lengths between surveyed locations on opposite sides of the lake were plotted on a diagram with midpoints placed on a center point. if fetch alone, as gilbert noted for lake bonneville (a fetch-dominated, much-larger version of the lake system), controlled the magnitude of shoreline superelevation for gsl, then the magnitude of shoreline superelevation would be similarly high at both ends of gsl in the direction of the longest fetch. in the direction of the shortest fetch, superelevation would be low at both ends of gsl. in addition, if fetch were the dominant control on superelevation, the midpoint patterns would resemble a bullseye. however, a pattern of the midpoint diagrams that showed trends of low to high superelevation could indicate that wind strength, in addition to fetch, caused the differences in shoreline superelevation. the patterns shown in figure 14 indicate strong storm winds from the northwest (atwood, 2006). wind data for gunnison and gilbert bays were not available in 2006 to corroborate or refute these conclusions. 2023 – analysis of patterns of shoreline superelevation and 2020-2023 wind data in 2023, wind data from weather stations on hat island in gilbert bay and gunnison island in gunnison bay was analyzed to corroborate or refute interpretations of the earlier studies. atwood (2006) suggested that patterns of shoreline superelevation were influenced by wind direction and strength, not simply by fetch. in 1990-2000, regional meteorological data, other than for salt lake international airport, were unavailable, and the salt lake international airport records were considered possibly non-representative of the open-lake conditions of gilbert and gunnison bays. instead, atwood (2006) used estimates of wind parameters by w. alder, utah state meteorologist, as 14 g. atwood, t.j. wambeam, and c.g. oviatt shoreline superelevation, clues to coastal processes of great salt lake figure 11. maps of the great salt lake perimeter surveys 1999-2000. adapted from atwood (2006). (a). the map shows the names of the ten places along the perimeter of great salt lake selected to test the findings of the 1997-1998 antelope island field surveys and explore relationships among aspect, fetch, and shoreline superelevation. (b). the numbers identify surveyed stretches at places along the perimeter of gilbert and gunnison bays. they indicate the 20 contrasting shores of the field survey, classified as generally high (red), intermediate (orange), or low (green) shoreline superelevation using the criteria of the 1997-1998 antelope island survey. 15 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 the empirical basis for definitions of wave environments for 1986-1987 (personal communication, reported in atwood, 2006; alder, 1986, 1987). university of utah mesowest weather stations on gunnison and gilbert bays (www.mesowest.utah.edu) now provide real-time wind data for gsl. j.d. horel (atmospheric sciences, university of utah, personal communication, 2023) provided the wind roses shown in figure 15. figure 16 displays the wind rose patterns combined with patterns of shoreline superelevation. downwind patterns explain patterns of gsl’s physical evidence of shoreline superelevation. they corroborate interpretations that the strongest winds that form the waves that cause shoreline superelevation come from the north, northeast and northwest. the cartoon sketches of figure 17 show the progression of a low-pressure system from offshore the pacific northwest, across california and nevada to utah and gsl (shafer and steenburgh, 2008). south winds precede the front's arrival, followed by strong northerly winds during and after the front’s passage. this substantiates the field surveys’ findings that the durable geologic evidence of shoreline superelevation in gsl documents strong storm winds from the northwest. figure 12. patterns of shoreline superelevation at strongs knob and at rozel point. two location maps with two orthophotos showing survey locations in contrasting areas of gunnison bay. (a). plot of surveyed places on the three shorezone stretches (8, 9, 10) of strongs knob. strongs knob, an island during 1986-1987, is located in southwestern gunnison bay immediately north of the railroad causeway. during the 1986-1987 highstand, location #8 with southeast aspect and long fetch, had low shoreline superelevation. location #9, a bay, had two northerly aspects. the northeast aspect had long fetch and high shoreline superelevation. the north-facing shore had long fetch and low shoreline superelevation. location #10, two separate shores with east-facing aspect had long fetch and high shoreline superelevation. (b). plot of surveyed places on the two shorezone stretches (18, 19) of rozel point. rozel point is located mid-bay on the east shore of gunnison bay. the rise of great salt lake flooded the spiral jetty immediately offshore. location 18 had a southwest aspect, intermediate fetch, and intermediate shoreline superelevation. location 19 had south and southeast aspects, short fetch, and low shoreline superelevation. 16 g. atwood, t.j. wambeam, and c.g. oviatt shoreline superelevation, clues to coastal processes of great salt lake figure 13. pairings of surveyed locations. adapted from atwood (2006). the four figures represent four steps to display fetch vectors at a center point. (a). step 1. locate places of surveyed shoreline superelevation. the numbers represent the surveyed places in gunnison bay on figure 12 with the number’s color indicating the generalized superelevation. draw lines representing fetch digitally between all pairs of places. obviously, the distance, for example, from place 15 to place 20 (82 km) is the same as from place 20 to place 15. the lines represent distance and direction both ways. (b). step 2. locate the midpoint of each line. (c). step 3. copy each line digitally. snap the lines across each other at their midpoints. (d). step 4. create the diagram that compares the effects of wind direction for places of fetch of equal length. the snapped diagram shows equal fetch to both ends. the color dot of the endpoints indicates relative shoreline superelevation, high (red), intermediate (yellow), and low (green). 17 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 discussion wave theory and wave dynamics, including the interaction of waves with coastlines and beaches, have generated extensive literature. scientific aspects have been discussed, for instance, by munk (1951), wright and short (1984), komar (1998), and wmo (2018). bertin and others (2020) recently reviewed infra-gravity waves. coastal landforms, morphodynamics, and processes of fetch-limited shorelines have been documented and discussed by cooper and others (2007) and freire and others (2009). fiedler and others (2020) provide a numerical modeling approach to beach erosion, wave overtopping, and street flooding from storm wave runup and superelevation where historical data are scarce or lacking. in contrast to the shoreline features of gsl, theuerkauf and others (2021) present the patterns and processes of geomorphic change caused by coastal storms on the shorelines of longer-fetch lake michigan. applequist (2013) presents a framework for assessing hazards in coastal environments linked to climate change, increasingly recognized as a factor in the evolution of weather patterns and storm intensity. from gilbert (1890) to schofield and others (2004) and jewell (2007) fetch has been a subject of shore processes of lake bonneville, a lake with wave environments not limited by fetch. wind transfers energy from the atmosphere into the water, creating wind waves (fontaine, 2013). the stronger and longer the wind blows, the higher and more energetic the wind waves. the transfer generates a chaos of wave heights and wave trajectories in a storm zone. the waves interact. as waves travel from a storm zone across a large open lake, the lake surface becomes progressively organized into a “fully arisen sea” of swell. swell transfers energy with negligible energy loss toward shore. the “sea” becomes more organized with longer fetch. wave energies and wind waves do not become fully organized if a lake is not big enough. wave development may be cut off during regime growth by lack of fetch, and this defines fetch-limited conditions (komar, 1998). under strong winds and as storms progress, waves develop, and lake water is pushed up against windward shores (wind setup). waves lose energy as they encounter the shoreface and then break. more energetic waves run farther up the shore, depositing their entrained and floated debris above the still water level. the entrained and floated debris becomes the sufigure 14. visual analysis of the fetch vector diagram. adapted from atwood (2006). (a). if fetch alone accounted for shoreline superelevation, the pattern of the colored endpoints would resemble a bull’s eye. the green dots representing low superelevation would cluster closer to the center and red dots representing high superelevation away from the center. the color dots on the bullseye diagram do not have a bullseye pattern. (b). the pattern of the dots indicates generally lower shoreline superelevation for northwestern, upwind locations and higher shoreline superelevation for southeastern, downwind locations. this trend implies that strong storm winds from the northwest contribute to patterns of shoreline superelevation of great salt lake, and both wind strength and fetch contribute to shoreline superelevation. 18 g. atwood, t.j. wambeam, and c.g. oviatt shoreline superelevation, clues to coastal processes of great salt lake perelevated shoreline evidence of still water lake level. winds directed straight at the shoreline inevitably produce greater superelevation than those of oblique incidence. storm duration, wind strength, and fetch determine the energy input into wind waves. complicating factors that affect wave runup and therefore shoreline superelevation, include wave setup, wind setup, wind-driven currents, the slope of the shore, shoreline morphology, lakebed and shoreline materials, and geographic features such as headlands that create sheltered zones. for example, antelope island shelters its eastern shore from most winds coming from the northwest. shoreline superelevation at any one location, although dominated by the triad of wave energy, wind setup, and wave setup, is the cumulative effect of all contributing factors. the us army shore protection manual (cerc 1984) treats the subject of coastal protection comprehensively and provides the empirical smb-84 nomograph developed by sverdrup and munk (1947) and modified by bretschneider (1952). the smb-84 chart is a simple graphical method to identify fetch-limited wave regimes such as gsl. figure 18, and figure 19 its key present a nomogram for gsl modified from cerc (1984). it indicates that wave regimes of gsl are fetch-limited. lo re and others (2016) found that simple empirical wind-wave models, such as smb84, give reliable results, and they remain popular among coastal engineers. the dark green line of the nomogram of figure 18 indicates that neither gunnison nor gilbert bay has sufficient fetch to develop a fully developed wave regime in response to storm winds. strong winds across bays of gsl transfer energy into the waves that cause shoreline superelevation under fetch-limited conditions. although maximum fetch across any direction of either gunnison or gilbert bay (figure 6) is too short for the wave environment to become fully developed during storm conditions, gentle winds over a long period of time can produce a fully developed refigure 15. wind data for gunnison and gilbert bays. the wind roses show wind direction and wind strength from over 300,000 total observations per location by mesowest for 2020 to 2023 (j.d. horel, atmospheric sciences, university of utah, personal communication, 2023). each wedge represents one of 16 cardinal directions. the length of the wedge represents the percent of the total observations for that site. the colors of the wedge represent the observations that fall in each of the speed classifications. (a) the wind rose for gunnison island in gunnison bay shows about 14% of the winds come from the northwest. most of the strong winds come from the north and west and not from the south and east. (b) the wind rose for hat island in gilbert bay shows about 13% of the winds come from the east, about 8% from the north and 9% from the southwest. most of the strong winds come from the north and southwest with fewer from the east. 19 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 16. gunnison and gilbert bays: wind patterns and patterns of shoreline superelevation. map of great salt lake overlaid with the data from figure 12 and wind roses from figure 16. the dark line indicates the extent of the 1986-1987 highstand. the wind rose diagrams of wind direction and strength appear to explain some of the patterns of shoreline superelevation of gunnison bay. patterns are more complex for gilbert bay. 20 g. atwood, t.j. wambeam, and c.g. oviatt shoreline superelevation, clues to coastal processes of great salt lake gime (wind energy at equilibrium with wave energy that arrives on gsl shores), as indicated by the small yellow triangle in figure 18. rowers and sailors commonly observe swell less than 1-2 ft on gsl (g. atwood and t. wambeam, personal observations). those conditions may affect currents and sedimentation patterns but are not the wave environments that leave evidence of storm wind direction. wind speeds recorded for hat island and gunnison island, from 2020 through 2023 (j.d. horel, atmospheric sciences, university of utah, personal communication, 2023), together with limits of fetch, provide constraints on wave regimes represented by the green rectangle on figure 18. a possible path of wave regime development over time under storm winds is plotted on the chart as a succession of three green stars (1, 2, 3). the blue star representing the empirical evidence of 1986-1987 conditions lies on the trajectory. the values given by the blue polygon on figure 18 for gilbert bay wave environments, for the lake at its 1986-1987 highest historic level, were based on infigure 17. cartoon of the progress of a low-pressure storm system. source: figure a), b), c) adapted from shafer and steenburgh (2008). key added. figure d), from j.d. horel, personal communication (2023). the low-pressure system progresses from offshore the pacific northwest coast to great salt lake, where its winds create wind waves that leave evidence of shoreline superelevation. a). a cyclonic system arrives at the pacific northwest coast. b). the system digs in and progresses across the great basin. c). the cold front arrives and crosses great salt lake. d). the two maps show wind direction and strength before and after a cold front crosses great salt lake. strong winds from the south precede the front's passage. strong winds from the north and northwest follow. strong winds transfer energy into the lake surface and create the wave regime that results in the superelevation of shoreline evidence. 21 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 18. wave regime chart for gilbert bay. the wave regime chart (see key next page) has four variables: a) wind stress in miles per hour and knots; b) fetch in miles and kilometers; c) duration of strongest winds in hours; and d) significant wave height in feet (the average height of the tallest one-third of waves). this chart (adapted from the smb-84 nomograph in cerc, 1984) summarizes the complex wave environment for given conditions. it shows a plot (blue polygon) for the conditions for gilbert bay reported in 1986-1987 (atwood, 2006). the wind speeds of figure 15 together with fetch suggest constraints on the limits of wave regime development in gilbert bay at that time (green rectangle). see the key and text for an explanation of the stars. 22 g. atwood, t.j. wambeam, and c.g. oviatt shoreline superelevation, clues to coastal processes of great salt lake terviews with w. alder, utah state meteorologist, who estimated the duration of strongest storms and wind speeds and david shearer, harbormaster of gsl saltair boat harbor marina, who estimated significant wave height and wind stress (atwood, 2006). fetch was measured from maps. the values for these four parameters define the blue polygon of figure 18, lying well within fetch-limited conditions (the white region of the chart). summary of gsl lake processes shoreline superelevation is evidence of the lake processes of gsl. wind develops waves and transfers energy into them. under strong winds, lake water stacks up against windward shorelines (wind setup). waves dissipate energy as they encounter the shoreface, run up, break, and deposit their entrained materials well above the static still water level monitored by the usgs. storm duration, wind strength, and fetch determine the energy input for the waves that leave the superelevated shoreline evidence. factors affecting wave run-up on shorelines include wave setup, wind setup, wind-driven currents, the slope of the shore, shoreline morphology, from convex to straight to concave, lakebed and shoreline materials, and geographic features that block winds or create sheltered zones. the cumulative effect of these diverse contributing factors is that shoreline superelevation may at any one location, although dominated by wave energy, includes wind set up and wave set up. the patterns of figure 19. key to figure 18 23 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 shoreline superelevation of antelope island 19861987 shoreline were caused by differences in the energy of wind waves arriving on shore. those energy differences, although primarily due to differences in fetch, were noticeably affected by wind strength. insights from documentation of shoreline superelevation on antelope island, corroborated by the recent analysis of winds across gunnison bay, suggest that geomorphic patterns of shoreline superelevation of fetch-limited paleolakes can provide evidence of strongest wind direction and clues to regional paleoclimate and weather. conclusions previous papers by the authors defined and presented evidence to quantify shoreline superelevation of the 1986-1987 highstand on gsl, documenting that geomorphic shoreline evidence is not at the still water level of the lake and does not define a horizontal plane from which to measure post-depositional change with confidence. this paper extends the findings of earlier work with empirical evidence of wind patterns across gilbert and gunnison bays from wind records from weather stations at hat and gunnison islands. we further explore the processes of shoreline superelevation. because gsl is fetch-limited, its geomorphic evidence at the highstand has a signal of wind direction and strength. wind records of weather stations on gsl indicate the strongest winds across gsl are from the north and northwest and correlate with geomorphic evidence. patterns of shoreline superelevation of gravel ridges and other geomorphic features along the shores of gsl are durable evidence of the direction of the strongest storm winds as well as effects of fetch. examination of shoreline superelevation of additional modernand paleofetch-limited lakes will lead to better understanding of their regional wind direction and strength and perhaps regional climate. acknowledgments we gratefully acknowledge robert baskin for his knowledge of gsl and his assistance with figures. we appreciate how william alder, david shearer, and john horel shared knowledge via personal communications. this paper was completed with the knowledge and encouragement of peter w. homewood and monique mettraux, geosolutions, trd. cameron scharrer assisted with figures 13 and 14. we appreciate the 1986-2000 collaboration of the utah geological survey, the university of utah, and u.s. geological survey colleagues who worked together to better understand gsl 1986-1987 shorelines as evidence of lake processes. we are grateful to ian schofield and daren nelson, whose reviews improved this paper. references adams, k.d. and wesnousky, g.g., 1998, shoreline processes and the age of the lake lahontan highstand in the jessup embayment, nevada: geological society of america bulletin, v.110 (10), p 1318-1332. adams, k.d., wesnousky, g.g., and bills, b.g., 1999, isostatic rebound, active faulting, and potential geomorphic effects in the lake lahontan basin, nevada and california: geological society of america bulletin, v. 111, issue 12, p. 17391756. adams, k.d., and bills, b.g., 2016, isostatic rebound and palinspastic restoration of the bonneville and provo shorelines in the bonneville basin, ut, nv, and id, in oviatt, c.g., shroder, j.f., jr., editors, lake bonneville: a scientific update. developments in earth surface processes 20: elsevier, p. 145-164. alder w., 1986, 1987, daily journal entries: unpublished: salt lake city ut. applequist l.r., 2013, generic framework for mesoscale assessment of climate change hazards in coastal environments: journal of coastal conservation, v. 17, p. 59–74. arnow, t., and stephens, d., 1990, hydrologic characteristics of the great salt lake, utah: 18471986: u.s. geological survey water-supply paper 2332, 32 p., map scale 1:125,000 atwood, g., 2006, shoreline superelevation, evidence of coastal processes of great salt lake, utah geological survey, miscellaneous publication 06-9, 231 pp. atwood, g., and mabey, d.r., 2000, shorelines of antelope island as evidence of fluctuations of the level of great salt lake, in king, j.k., and willis, g.c., eds., geology of antelope island, davis county, utah: utah geological survey miscellaneous publication 00-1, p. 85-97. bertin, x., martins, k., de bakker, a., chataigner, t., guérin, t., coulombier, t. and de viron, o., 2020, energy transfers and reflection of infragravity waves at a dissipative beach under storm waves: journal of geophysical research: oceans, v. 125, no. 5, e2019jc015714. bretschneider, c.l., 1952, the generation and decay of wind waves in deep water: transactions american geophysical union, vol 33, p 381-389. cerc (u.s. army coastal engineering research 24 g. atwood, t.j. wambeam, and c.g. oviatt shoreline superelevation, clues to coastal processes of great salt lake center), 1984, shore protection manual, 4th edition: u.s. government. chen, c.y., and maloof, a.c., 2017, revisiting the deformed high shoreline of lake bonneville: quaternary science reviews, v. 159, p. 169-189. cooper j.a.g., lewis d.a., and pilkey o., 2007, fetch-limited barrier islands: overlooked coastal landforms: geological society of america geology, geology today, v. 17, no. 3, p. 4-9. currey, d.r., 1982, lake bonneville: selected features of relevance to neotectonic analysis of great salt lake and vicinity, in gwynn, j.w., ed., great salt lake: a scientific, historical and economic overview; bulletin 116: salt lake city, utah geological and mineral survey, p 69-82. fiedler, j.w., young, a.p., ludka, b.c., o’reilly, w.c., henderson, c. merrifield, m.a. and guza, r.t. 2020, predicting site-specific storm wave run -up: natural hazards, v. 104, p. 493–517. fontaine, e., 2013, a theoretical explanation of the fetchand duration-limited laws: journal of physical oceanography, v.43, p. 233-247. freire, p., ferreira, ó., taborda, r., oliveira, f.s.b.f., carrasco, a.r., silva, a., vargas, c., capitão, r., fortes, c.j., coli, a.b., and santos, j.a., 2009, morphodynamics of fetch-limited beaches in contrasting environments: journal of coastal research si 56, proceedings of the 10th international coastal symposium, p. 183-187. gilbert, g.k., 1890, lake bonneville: u.s. geological survey monograph 1, 438 p. jewell, p.w., 2007, morphology and paleoclimatic significance of pleistocene lake bonneville spits, quaternary research v. 68, p. 421-30. komar, p.d., 1998, beach processes and sedimentation (2nd edition): prentice hall, upper saddle river, new jersey, 544 p. lo re, c., cannarozzo, m., and ferreri, g.b., 2016, present-day use of an empirical wave prediction method: maritime engineering, v. 169, issue ma1, 14 p. mabey, d.r., 1986, notes on the historic high level of great salt lake: utah geological and mineral survey, survey notes, v. 20, no. 2, p. 13-15. mesowest, www.mesowest.utah.edu munk, w.h., 1951, origin and generation of waves: proceedings of the 1st coastal engineering conference, long beach, california, p. 1–4, doi: 10 .9753/ icce, v1 .1. murchison, s.b., 1989, fluctuation history of great salt lake, utah, during the last 13,000 years: salt lake city, university of utah, ph.d. dissertation, 137 p. oviatt, c.g., atwood, g., and thompson, r.s., 2021, history of great salt lake, utah, usa, since the termination of lake bonneville, in rosen, m.r., finkelstein, d., park-boush, l., and pla-pueyo, s., editors, limnogeology: progress, challenges and opportunities: a tribute to beth gierlowskikordesch: syntheses in limnogeology no. 2, new york, springer nature ag, p. 233-271. schofield, i., jewell, p.w., chan, m., currey, d.r., and gregory, m., 2004, shoreline development, longshore transport and surface wave dynamics, pleistocene lake bonneville, utah: earth surface processes and landforms, v. 29, p. 1675-1690. shafer, j.c., and steenburgh, w.j., 2008, climatology of strong intermountain cold fronts: monthly weather review, v 136, 784-807. sverdrup, h.u. and munk, w.h., 1947, wind, sea, and swell theory of relationships in forecasting: us dept. of the navy hydrographic office publication, no 601. tackman, g.e., 1993, middle and late pleistocene hydrologic history of diamond valley, eureka and elko counties, nevada, with climatic and isostatic implications: salt lake city, university of utah, masters of science thesis, 192 p. tackman, g.e., currey, d.r., bills, b.b., and james, t.s., 1998, paleoshoreline evidence for postglacial tilting in southern manitoba: journal of paleolimnology, v. 19, 443-463. theuerkauf, e. j., mattheus, c.r., braun, k.n. and bueno, j., 2021, patterns and processes of beach and foredune geomorphic change along a great lakes shoreline: insights from a year-long drone mapping study along lake michigan: shore & beach, v. 89, no. 2, p. 46-55. wmo (world meteorological organization), 2018, guide to wave analysis and forecasting. world meteorological organization report wmo no 702, 189 p. wang, p., 1978, seiches in the great salt lake: salt lake city, university of utah, ph.d. dissertation, 119 pp. wright, l.d., and short, a.d., 1984, morphodynamic variability of surf zones and beaches: a synthesis: marine geology, v. 56, p. 93–118. utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors william w. little1,2 and david a. little1 1w.w. little geological consulting, llc, 185 west 200 south, wellington, utah 84542 2department of geology, brigham young university – idaho, 525 south center street, rexburg, idaho 83460-0510 corresponding email: wwlittle@gmail.com cover image: outcrop of coastal-plain or delta-top deposits within the upper aberdeen member equivalent (kba-equiv) of the undifferentiated blackhawk formation. the outcrop is capped by the undifferentiated blackhawk formation (kb undif). field localities in the book cliffs to understand sequence stratigraphic concepts mailto:wwlittle%40gmail.com?subject=uga%20geosites 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. the geosites reflect the interests of the many volunteers who wrote to share some of their favorite geologic sites. the list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. the utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. this is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and figures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. presidents message i have had the pleasure of working with many different geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several different geological provinces. we have the basin and range to the west and the central utah hingeline and thrust belt down the middle. the uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava flows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. the general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. the “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. the articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. this guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, thank you to the american association of petroleum geologists, rocky mountain section foundation for their financial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fill the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president little, w.w. and little, d.a., 2019, field localities in the book cliffs to understand sequence stratigraphic concepts, in milligan, m., biek, r.f., inkenbrandt, p., and nielsen, p., editors, utah geosites: utah geological association publication 48, 5 p., https://doi.org/10.31711/ugap.v1i1.90. https://geology.utah.gov/apps/geosights/index.htm https://dutchiedesign.com https://geology.utah.gov/ https://creativecommons.org/licenses/by/4.0/ w. w. little and d. a. little field localities in the book cliffs to understand sequence stratigraphic concepts 3 introduction the book cliffs of utah and colorado have become the premier location globally to study and teach principles of sequence stratigraphy. continuous, well-exposed, and easily-accessible outcrops along both depositional dip (mountain to sea) and depositional strike (along the shore) make it possible for detailed three-dimensional reconstruction and analysis of sedimentary successions. most types of clastic sedimentary systems are found in the book cliffs. these include river (braided, meandering, and anastomosed), delta (riverand wave-dominated), estuary, beaches, and seafloor, demonstrating a complete sediment delivery system from near the mountains to the ocean basin. these characteristics make the book cliffs an excellent classroom to study the interrelationship between large-scale changes in sea level, tectonic processes that build mountains and form sedimentary basins, and the sedimentary deposits that fill those basins. this article differs somewhat from others in this volume in two ways: 1) it does not focus on a single location but includes a series of 25 related localities found along an approximately 100-mile stretch of the book cliffs, extending from north of the city of helper southeastward to west of the town of thompson springs (figure 1). localities are arranged first by topic then by location, with driving directions and gps coordinates given at the end of the main text. 2) the focus of this article, teaching concepts of sequence stratigraphy, makes it difficult to write meaningfully in a jargon-free manner; therefore, a brief generalized introduction is given to describe the study of sequence stratigraphy for all audiences, and the remaining text is written in a more technical fashion. simple description of sequence stratigraphy stratigraphy is the study of layered rock, in most cases sedimentary rock. in a sedimentary succession, each layer is like a chapter in a book, in that it records the history for a specific interval of time. in the case of sedimentary layers, that history includes the manner in which sediment accumulated within a particular setting; for instance, adjacent to the channel of a river or along a shoreline. the boundaries between layers are surfaces that denote episodes during which sediment did not collect, such as between flooding events for a stream, or where sediment collected but was later removed by erosion. these surfaces represent a “gap” in time for which there is no record. if these gaps are of significant duration, they are called unconformities. unconformities are commonly the consequence of changes in sea level. when sea level rises, space is created above the previous layer into which additional sediment can be deposited, forming a new layer on top of the old. if sea level falls, previously deposited sediment is eroded and carried to a new location farther seaward. sequence stratigraphy, therefore, is a study of the history of sea-level changes as preserved in sedimentary layers and the unconformities between them. the fundamental unit of sequence stratigraphy is a sequence, which refers to a predictable vertical pattern of layering developed during a single cycle of sea-level rise and fall. it is this pattern (sequence) which gives the study its name. for the book cliffs, these layers tend to be thinner and more muddy at the base, becoming thicker and more sandy toward the top. additionally, the bottom layers mostly record environments from deeper water or from a more seaward position than those above, which represent environments more landward. in other words, in a highly simplistic manner, lower layers are likely to represent the ocean bottom, middle layers beaches, and capping layers rivers. a sequence is typically bound both on bottom and on top by erosional unconformities that developed during times when base-level fell and are called sequence boundaries. the individual layers within a sequence form mostly during a stepwise rise in sea level and are known as parasequences. a parasequence forms when the rising sea level creates new space and then pauses long enough for that space to be filled with sediment before the sea rises again. sequence stratigraphic concepts are well described in a number of sources (e.g., coe, a., 2002, is an excellent resource for those new to sequence stratigraphy; mitchum and others, 1977; van wagoner and others, 1988, 1990; van wagoner, 1995a; neal and abreu, 2009; catuneanu and others, 2011; posamentier and allen, 1999; and abreu and others, 2014 present more technical perspectives). technical description of sequence stratigraphy in a little more technical sense, sequence stratigraphy refers to the depositional pattern developed during a single cycle of base-level fluctuation (figure 2a). more precisely, it predicts the order (sequence) in which progradational depositional events (parasequences) stack in response to changes in the balance between the rate at which accommodation is produced or destroyed and the rate at which that space is either filled by sediment, bypassed, or reduced through erosion. this relationship causes parasequences to assemble in progradational (forward stepping), aggradational (vertically stacked), and retrogradational (backstepping) sets separated by discontinuity surfaces (figure 2b through 2d). these surfaces reflect either an increase in accommodation exceeding the capacity of sediment to fill that space (flooding surface), which is marked by an abrupt landward shift in sedimentary facies; an increase in space that is less than the rate of filling (type 2 sequence boundary), shown by a basinward shift in sedimentary facies in the form of a normal progradational pattern; or an actual loss of accommodation associated with a drop in base level and erosion of earlier deposits (type 1 sequence boundary), generated by a forced regression. it is important to note that depending on scale and tectonic setting, processes other than a simple fluctuation of base level, such as a variation in sediment supply, delta-lobe switching, or channel development and abandonment, can produce stratiw. w. little and d. a. little field localities in the book cliffs to understand sequence stratigraphic concepts 4 figure 1. a) conceptual model of an idealized thrust belt/foreland basin system. (from decelles and giles, 1996) b) dip-oriented cross section showing generalized stratigraphy of the cretaceous western interior foreland basin. (from armstrong, 1968) c) index map showing the region covered by field localities in this report. (from willis, 1999) d) expanded view of c. e, f, and g) detailed index maps showing localities (red dots) discussed in this report. driving directions and gps coordinates are given at the end of the report. w. w. little and d. a. little field localities in the book cliffs to understand sequence stratigraphic concepts 5 figure 2. a) conceptual sequence-stratigraphic diagram illustrating relationship between parasequence stacking patterns and associated surfaces referred to throughout this report. abbreviations are as follows: sb (sequence boundary), ts (transgressive surface), mfs (maximum-flooding surface), fsst (falling-stage systems tract), lst (lowstand systems tract), tst (transgressive systems tract), hst (highstand systems tract). parasequences can stack in b) a forward-stepping pattern, c) a vertically-stacked pattern, or d) a backstepping pattern, depending on the relative balance between sediment supply (s) and accommodation production (a). (b, c, and d modified from van wagoner and others, 1990). figure 3. dip-oriented cross section of the composite sequence made up of the star point sandstone, blackhawk formation, and lower castlegate sandstone. names printed within the gray marine mudstone region are formation members. castlegate sandstone has been abbreviated c.ss. note the overall coarsening upward trend (3rd-order sequence) with smaller-scale coarsening-upward (4th-order) successions contained within. the sawtooth pattern reflects parasequence stacking arrangements within the smaller-scale sequences. (from pattison, 2018; after young, 1955; balsley and horne, 1980; cole and others, 1997; and pattison and others, 2007a) w. w. little and d. a. little field localities in the book cliffs to understand sequence stratigraphic concepts 6 graphic patterns similar to those discussed in this report; however, a discussion of variations to the basic sequence stratigraphic model are beyond the scope of this paper. this paper concentrates entirely on showing field examples that can be used to illustrate general concepts of sequence stratigraphy in a simplified manner. geological setting the third-order composite sequence comprised of the early through middle campanian star point sandstone, blackhawk formation, and lower castlegate sandstone of the mesaverde group is the focus of this report (figure 3). these formations were deposited as a series of successive progradational sheets into the foredeep portion (figure 1a, b; decelles and giles, 1996) of the western interior foreland basin in response to eastward movement along the canyon range and adjoining segments of the sevier fold/thrust belt to the west (figure 1c; villien and kligfield, 1986; schwans, 1995). the western interior basin is a foreland basin that subsided and was filled in response to thrust loading in the sevier orogenic belt and dynamic loading associated with subduction of the underlying farallon plate (figure 1b; bally and others, 1966; armstrong, 1968; kauffman and caldwell, 1993; decelles and coogan, 2006; liu and others, 2011). emphasis in this report is placed on regional sheet sandstone bodies and their capping coal deposits, as they represent easily identifiable progradational depositional events that can be traced from landward pinchouts against coastal-plain deposits to down-dip terminations into marine mudstone (figure 3). parasequences in each of the members of the blackhawk formation are dominated by prograding shoreface successions associated with wave-dominated deltas, in which prodelta mudstone at the base abruptly overlies shoreface sandstones and coastal plain coals of the previous cycle (figure 4). upward through the succession, prodelta mudstone becomes interbedded with thin beds of hummocky to parallel-laminated sandstone (transition zone), which in turn is overlain by amalgamated intervals of hummocky sandstone figure 4. idealized facies succession for shoreface systems in the blackhawk formation. (modified from sepm strata; after van wagoner and others, 1990). photographs of a “typical” shoreface succession from the spring canyon member (sowbelly parasequence) in gentile wash. abbreviations are ss for sandstone and hcs for hummocky stratification. w. w. little and d. a. little field localities in the book cliffs to understand sequence stratigraphic concepts 7 (lower shoreface), followed by trough cross-laminated sandstone (upper shoreface), then parallel-laminated sandstone (foreshore), and capped by coal (coastal plain). the contact between each parasequence is sharp and at least somewhat erosional. report organization this report focuses on the utah portion of the book cliffs stratigraphic succession and progresses from onshore to offshore and from stratigraphically oldest to youngest through the star point/ blackhawk/lower castlegate interval, beginning with shoreline and coastal-plain deposits near the town of helper and ending with small-scale, incised, channel-fill units at thompson canyon (figure 2). the report is divided into four sections: section i illustrates general principles of sequence stratigraphy as applied to a single vertical profile through the entire third-order sequence, beginning with a hike into gentile wash and finishing with stops along u.s. highway 6; section ii focuses on the concept of a systems tract and examines lateral facies changes, both along depositional strike and down depositional dip, using the same stratigraphic units observed in section i; section 3 investigates characteristics of incised channel fill deposits; and section 4 allows observation of a small-scale (fourth-order) sequence in single outcrops. driving directions and gps coordinates for each locality are given at the end of the report. significance cretaceous deposits of the book cliffs region are often cited as an analog for subsurface exploration, particularly in foreland basins, and sequence stratigraphy has become the primary method for correlating and mapping depositional successions, leading to important discoveries of petroleum in fields that had been abandoned, as well as new discoveries. in the search for and development of both coal and hydrocarbons, recognizing facies associations and the manner in which they change laterally and stack vertically aids prediction of the geometry, spatial distribution, continuity, and resource potential of associated rock bodies. the book cliffs of east-central utah include some of the best exposed ancient examples of fluvial, coastal, and marine deposition on earth, especially with respect to foreland basins. consequently, these strata are visited each year by hundreds of geologists from across the world. continuous, undeformed exposures dissected by a large network of canyons with little vegetal cover make it possible to clearly document lateral and vertical facies changes over broad distances, as well as key surfaces that bound stratigraphic successions, both of which are critical aspects to understanding the geometry, spatial distribution, continuity, and resource quality of clastic sedimentary rock bodies. much of the initial interest in these stratigraphic units was related to the presence of commercial coal seams. subsequently, these deposits have increasingly become cited as analogs for hydrocarbon exploration, giving a clearer perspective when making interpretations of similar stratigraphic units from cores, logs, and seismic data. the primary purpose of this report is to denote exceptional outcrops that can be used as analogs for clastic depositional facies models and toward understanding and teaching principles of sequence stratigraphy. field localities section i: general characteristics of a sequence locality i-a: north helper escarpment (lithostratigraphic and sequence stratigraphic overview of the mesaverde group) the purpose of this overlook is to gain a general perspective of the star point/blackhawk/lower castlegate sequence. descriptions of individual stratigraphic units will be provided in later sections. a complete succession, including all lithostratigraphic units and sequence stratigraphic elements can be seen along the escarpment bordering the north side of the town of helper, stretching eastward from u.s. highway 6 for just under 2 miles (3 km) to kenilworth wash (also called helper canyon; figure 5). figure 5. the north helper escarpment showing the mancos shale (km), panther tongue (ksp) and storrs (kss) members of the star point sandstone, and spring canyon (kbsc) and aberdeen (kba) members of the blackhawk formation. kb is blackhawk formation undifferentiated. also labelled on the photo are corresponding sequence stratigraphic elements as identified in figure 4. 8 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 the basal part of the escarpment comprises several hundred feet of offshore marine mudstone belonging to the mancos shale, with significantly more in the subsurface. about a third of the way up the escarpment (half way up the photo of figure 5), the first resistant ledge is the panther tongue member of the star point sandstone, interpreted as having been deposited by a river-dominated delta under conditions of forced regression (howard, 1966; newman and chan, 1991; enge and others, 2010). the lower contact is conformable with the underlying mancos shale, forming the lower sequence boundary, which, being conformable, is expressed in this area as a correlative conformity. above the panther tongue is an interval of interbedded sandstone and mudstone that first forms a relatively thin fining-upward succession, followed by a much thicker accumulation that coarsens upward. the mudstones are tongues of mancos shale that thin and pinchout to the west (landward) and sandstones above the panther were deposited as shoreface successions (spieker and reeside, 1925; clark, 1928; young, 1955; balsley and horne, 1980), each of which represents an individual progradational event (parasequence). within the first muddy interval are three thin and poorly-developed benches, with each being thinner and less distinct than the lower. these are distal expressions of shoreface parasequences within the storrs member of the star point sandstone. each is silty with a thin cap of sandstone and together form part of a back-stepping parasequence pattern associated with the transgressive systems tract. the upper contact of the panther tongue is the transgressive surface, as it marks the first major flooding event following deposition of the lowstand panther tongue. the prominent sandstone ledges in the upper part of the escarpment belong to the spring canyon and aberdeen members of the blackhawk formation. in this part of the section, each sandstone is composed of a more complete shoreface parasequence than that below, and together they represent progressively more proximal shorefaces upward through the section, reflective of a progradational stacking pattern, which is diagnostic of a highstand systems tract. the surface forming the contact between the upper two benches in the star point sandstone is where retrogradation transitions to brief aggradation, followed by progradation and marking the maximum-flooding surface. in this section, aggradation in the lower part of the highstand tract is minor and is expressed by an abrupt turnaround from transgression to regression. overlying the shoreface sandstones are coastal-plain deposits of the undifferentiated blackhawk formation (spieker and reeside, 1925; clark, 1928; young, 1955; balsley and horne, 1980), which correspond to the kenilworth through desert members, each of which becomes a prominent shoreface succession to the east that is similar in characteristics to the spring canyon and aberdeen members of this locality. in this report, we include the lower castlegate sandstone (distant cliffs out of view on figure 5) as the capping part of the highstand systems tract, though many workers prefer to place a sequence boundary at the base of the castlegate (van wagoner, 1995b; yoshida, 2000). locality i-b: gentile wash (hike through a sequence – panther tongue through spring canyon members) gentile wash provides an opportunity to hike through and observe closely all parts of the sequence, from the underlying mancos shale, across the sequence boundary near the mouth of the wash, into the forced-regressive lowstand deposits of the panther tongue, across the transgressive surface, through the retrogradational parasequence set of the transgressive systems tract in the storrs member, and across the maximum-flooding surface into the progradational parasequence stacking pattern of the lower highstand systems tract found in the spring canyon member (figure 6). the hike begins at the mouth of gentile wash and continues for about 1.25 miles (2 km) up the wash. sequence boundary (mancos shale/panther tongue member contact): a sequence boundary is defined as a subaerial unconformity and its correlative marine conformity (mitchum and others, 1977). a subaerial unconformity develops as base level drops during a forced regression. simultaneously, basinward, deposition is continuous, producing a time-equivalent conformable succession of strata, the lower boundary of which is considered the correlative conformity. the panther tongue in the subsurface, northward of gentile wash, shows a downstepping pattern, consistent with a forced regression (posamentier and morris, 2000; hwang and heller, 2002). exposures at gentile wash are of the final downward and most basinward step; therefore, at this locality the sequence boundary shows a simple shallowing-upward, conformable succession from the mancos shale into the panther tongue (figure 7). figure 6. stratigraphic section in gentile wash. the bottom of the photo is approximately the top of the panther tongue. named parasequences within the spring canyon member are identified with “ps” to assist with correlation to other localities. w. w. little and d. a. little field localities in the book cliffs to understand sequence stratigraphic concepts 9 falling-stage to lowstand systems tracts – river-dominated delta-front and channel-mouth bar deposits (panther tongue member): the panther tongue member was deposited as a river-dominated delta during the falling-stage to lowstand systems tracts. it is bound by the sequence boundary below and transgressive surface above. interpretation as a river-dominated deltaic deposit is based on a coarsening-upward pattern, a lobate shape, and clinoforms that dip basinward. three subdivisions, based on facies associations, are evident at the mouth of gentile wash: 1) a basal interval of thinly interbedded sandstones dominated by bouma sequences and bioturbated mudstones, 2) a middle section of thin sandstones characterized by parallel laminations, and 3) a capping interval of lens-shaped sandstones that can be structureless, parallel laminated, trough cross laminated, or a mixture of the three (figure 8). sandstones in the capping portion can be amalgamated or separated by thin mudstones. the succession represents an upward progression from lower delta-front turbidites and storm beds at the base through upper delta front amalgamated turbidite deposits into reworked channel-mouth bar and interdistributary bay deposits. transgressive surface (panther tongue member/mancos shale contact): the primary evidence for a transgressive surface at the top of the panther tongue is an overlying lens of mancos shale in sharp contact, indicating the first significant rise in base level and subsequent drowning of the shoreline following the forced regression (figure 8b). other evidences include an undulatory and erosional surface, oscillatory ripples, mud-clast and wood-fragment molds, intense bioturbation, and shell fragments, demonstrating high-energy reworking of the former subaerial surface by waves, followed by colonization by burrowing organisms (figure 9). figure 7. conformable contact (correlative conformity) between the mancos shale (km) and the panther tongue member (ksp), marking the basal sequence boundary (sb) for the star point/blackhawk/lower castlegate sequence. figure 8. a) facies and major divisions of the panther tongue in strike-oriented view at the mouth of gentile wash. b) expression of clinoforms in depositional dip orientation, located across u.s. highway 6 at the mouth to panther canyon. sb is the sequence boundary at the base of the panther tongue member (ksp) and marks the contact with the underlying mancos shale (km). the panther tongue member is capped by the transgressive surface (ts) and overlain by a thin wedge of the mancos shale. figure 9. upper surface of the panther tongue member at gentile wash and sowbelly gulch showing characteristics common to a transgressive surface of erosion. 10 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 transgressive systems tract (storrs member): in gentile wash, the storrs member consists of three distal shoreface successions, the lower two of which are progressively less well developed than the previous moving upward from the panther tongue, suggesting a smaller amount of progradation for each succeeding parasequence, leading to a retrogradational stacking pattern characteristic of a transgressive systems tract (figure 10). the lowermost parasequence consists of interbedded mudstone and ripple-bedded sandstone at the base (distal transition zone), grading upward into interbedded mudstone and hummocky sandstone (proximal transition zone), and capped by a thin interval of amalgamated hummocky sandstone beds (lower shoreface). the middle and upper parasequences are similar but lack the amalgamated sandstone part of the section. maximum-flooding surface (storrs member): by definition, the maximum flooding surface is placed where parasequence stacking changes from a retrogradational to an aggradational or progradational pattern (van wagoner and others, 1988). figure 10 shows two parasequences within the storrs member that exhibit nearly equal degrees of development, suggesting the shoreline had stabilized and ceased its landward translation. parasequences both above and below these are more completely developed. distally, maximum flooding surfaces are characterized by slow deposition of fine-grained sediment under quiet, deep-water conditions. often, they are exemplified more by a “condensed interval,” in which a thin sedimentary deposit represents an extended period of time (galloway, 1989). in a landward direction, particularly in a foreland basin setting, the maximum flooding surface is likely to grade into a thick succession of fluvial strata. early highstand systems tract (spring canyon member): though the highstand systems tract is significantly thicker than the others and encompasses the remainder of the blackhawk formation and possibly the lower part of the castlegate sandstone, this hike ends near the top of the spring canyon member, in the lower part of the tract (figure 10). stratigraphically higher units, namely the aberdeen member and undifferentiated blackhawk formation, along with the castlegate sandstone, will be observed from outcrops along u.s. highway 6. above the maximum flooding surface, each parasequence within the spring canyon member becomes progressively coarser grained and more completely developed, until one of the upper parasequences, known as the sowbelly parasequence, consists of a complete shoreline succession, from interbedded mudstones and hummocky sandstones of the transition zone to amalgamated hummocky and contorted sandstone beds in the lower shoreface, followed by trough cross-laminated sandstone of the upper shoreface, then parallel-laminated sandstone in the foreshore (figure 3; kamola, 1987; kamola and van wagoner, 1995). this parasequence is capped by coal of the coastal plain. the upper contact of the coal bed is the flooding surface beneath the next (hardscrabble) parasequence. parasequences above the sowbelly show less of the lower (distal) parts of the succession and more of the upper (proximal) facies. locality i-c: road-maintenance shed (early highstand systems tract continued spring canyon and aberdeen members): the spring canyon and aberdeen members of the blackhawk formation are alike in character (figure 11). each consists of stacked parasequences similar to those observed in gentile wash. the primary difference between this locality and that of gentile wash is parasequences in the aberdeen member preferentially develop only the upper parts of shoreface successions, because deposits represent progressively more proximal environments upward through the section. therefore, there is less of the hummocky sandstone of the lower shoreface, virtually none of the interbedded mudstone characteristic of the transition zone, and more of the trough cross-bedded and parallel laminated sandstones of the upper shoreface and foreshore, respectively, as well as a greater number of coal beds capping the parasequences. locality i-d: monument turnout (late highstand systems tract – blackhawk formation, upper aberdeen member equivalent): at the base of the road cut on the west side of u.s. highway 6 (figure 12), wave-dominated shoreface parasequences in the aberdeen member are similar to those of the spring canyon member at gentile wash. these are overlain by coastal-plain or delta-top deposits of the undifferentiated blackhawk formation that have been correlated to incised channel-fill deposits in the upper part of the aberdeen member farther to the east (pattison and others, 2007b; kamola, pers. comm. 2014). shallow fluvial environments are indicated by channel-form sandstone bodies that show strong figure 10. backstepping (retrogradational) to vertically-stacked (aggradational) parasequence (ps) pattern in the storrs member (kss), followed by a strongly forward stepping (progradational) pattern in the spring canyon member. the change from retrogradation to a very brief interval of aggradation in the upper storrs member marks the maximum-flooding surface (mfs). km refers to thin interbedded intervals of the mancos shale. parasequences (ps) are numbered in the storrs member and named in the spring canyon member. 11 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 figure 11. a) in this roadcut, the capping (helper) parasequence of the spring canyon member (kbsc) and the aberdeen (kba) member of the blackhawk formation are alike in character. b) a relatively complete parasequence at the top of the spring canyon member (helper parasequence). lithofacies are labeled hcs for hummocky-stratified sandstone, st for trough cross-stratified sandstone, and sh for horizontally-laminated sandstone. c) close up of a coal seam, of which the upper contact forms the boundary between the spring canyon and aberdeen members. d through f) details of parasequences in the aberdeen member. note the predominance of trough cross stratification and parallel laminations and absence of hummocky laminations. f shows a small channel, possibly the result of a small outlet channel across the foreshore. figure 12. a) outcrop of coastal-plain or delta-top deposits within the upper aberdeen member equivalent (kba-equiv) of the undifferentiated blackhawk formation. the outcrop is capped by the undifferentiated blackhawk formation (kb undif). b) contact with the underlying aberdeen member (kba). c and d) filled distributary channel deposits within delta-front sandstones. note heterolithic bedding and lateral accretion surfaces. lateral accretion throughout the exposure. current and climbing ripples are common in likely splay deposits. tidal influence is evidenced by heterolithic bedding along the lateral accretion surfaces within the channel forms, and possible wave influence is suggested by trough, cross, and parallel laminations within thin sheet-like bodies. locality i-e: monument turnout (late highstand systems tract continued – blackhawk formation undifferentiated and castlegate sandstone): on the hillside across the price river to the east of the turnout, the blackhawk formation above the aberdeen member has not been subdivided into members but is referred to as blackhawk formation undifferentiated (figure 13). these strata were deposited as an overall coarsening-upward succession of coastal-plain then fluvial deposits, culminating in the lower castlegate sandstone. the blackhawk undifferentiated section correlates eastward to the kenilworth, sunnyside, grassy, and desert members of the blackhawk formation (young, 1955), each of which is stratigraphically higher than the member beneath, extends farther basinward before terminating into the mancos shale than the member below, and, in general lithological character, resembles the spring canyon and aberdeen members of gentile wash and the u.s. highway 6 road maintenance shed localities. on this hillside, the castlegate sandstone is approximately 260 feet (80 m) thick and consists of discontinuous, overlapping sheets of sandstone up to 23 feet (7 m) thick (e.g., adams and bhattacharya, 2005; mclaurin and steel, 2007). similar to members in the blackhawk formation, the lower castlegate sandstone, which in this locality was deposited by braided streams, also transitions eastward into coastal plain, then shoreface, and finally, near the colorado border, pinches out into the mancos shale. most workers interpret the base of the lower castlegate sandstone as a sequence boundary based on a regional-scale erosional surface that further down dip is overlain by a basinward shift in facies, possible evidence for onlap (e.g., van wagoner and others, 1990; van wagoner, 1995a), and a change in sandstone composition across the boundary from lithic to quartz arenite (horton and others, 2004). however, lack of a significant transition in alluvial architecture across the boundary (adams and bhattacharya, 2005), correlation with syntectonic sheetflood deposits of the indianola group landward (lawton, 1986a, b), and an overall coarsening upward succession similar to that of tectonically-driven w. w. little and d. a. little field localities in the book cliffs to understand sequence stratigraphic concepts 12 fluvial cycles in foreland deposits of the kaiparowits basin (little, 1995, 1997) suggest this section might be far enough inland to have been unaffected by eustatic base-level fluctuation and, rather, represents continued coarsening-upward and basinal progradation as the capping part of the sequence. locality i-f: rolapp mine road (late highstand systems tract concluded castlegate sandstone): at this locality, fluvial characteristics of the castlegate sandstone are plainly visible (figure 14). overall, the castlegate is thought to have been formed by shallow braided streams (van de graaff, 1972), as evidenced by sand-rich deposits that are thin and laterally discontinuous; however, many of the channel forms also show signs of lateral migration, more indicative of meandering channels, suggesting at this locality the castlegate might be somewhat of a hybrid between the two, showing a transition from braiding upstream to more of a meandering pattern downstream. section ii: lateral facies changes (systems tracts) the concept of a systems tract is that depositional systems are physically and contemporaneously linked along the same basin margin (galloway, 1989). within a systems tract, changes in facies and facies associations occur in both along-strike and down-dip directions. in order to illustrate lateral variations in sedimentation, the same units described in part i of this report are traced and observed at other localities. lateral facies changes in the panther tongue thoughts regarding the character of river-dominated deltas in the subsurface are evolving from the traditional digitate model exemplified by a few prominent distributary channels separated by vast interdistributary bays, such as the modern mississippi river delta, to a more lobate pattern in which potentially numerous channels fan from a trunk stream at a wide range of scales, mostly shallow and with highly variable orientations. these latter deposits are dominated by channel-mouth bars that show evidence for simultaneous accretion in down-dip, along-strike, and up-dip directions as they grow subaerially in all directions, making for a complex pattern of variably-oriented clinoforms (van heerden, 1983; olariu and bhattacharya, 2006). modern examples that have been used as analogs for the panther tongue are the atchafalaya and wax lake river deltas of the gulf coast (olariu and bhattacharya, 2006); in that distributary channel fill deposits of the panther tongue show a complex overlapping relationship at multiple scales, being essentially channel fills within channel fills with a wide range of orientations. in spring canyon, three scales of channel forms have been identified within the panther tongue (figure 15). the outcrop itself is a channel-form body of sandstone about 80 feet (25 m) thick in the center. within the overall channel form is a relatively thick sandstone lens that pinches into channel-mouth bars both to the east and to the west from the mouth of sowbelly gulch. a third, smaller-scale, set of channel forms showing a wide dispersal range is scattered throughout the outcrop from gilson gulch on the west to a short distance east of sowbelly gulch. figure 13. above shoreface parasequences of the aberdeen member (kba) are coastal-plain deposits of the upper (undifferentiated) blackhawk formation (kb undif), which are overlain by braided fluvial deposits of the castlegate sandstone (kc). these units form the final progradation and capping portion of the star point/ blackhawk/lower castlegate sequence. figure 14. a) shows the nature of the contact between the blackhawk formation (kb) below and lower castlegate sandstone (kcl) above. complex, discontinuous and lenticular to sheet-like nature of sandstone bodies and presence of lateral accretion surfaces are both well expressed. kcm refers to the middle castlegate sandstone, which is part of the overlying sequence and not addressed in this paper. b) close up of the lower castlegate sandstone, taken immediately to the left of a. 13 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 figure 15. a) panther tongue distributary channel fill pinch out into a channel-mouth bar succession toward the east (right) side of the photo on the east side of sowbelly gulch. the base of the panther tongue is a sequence boundary and the upper contact is a transgressive surface overlain by the mancos shale. b) the distributary channel fill consists of multiple generations of intrachannel bars and smaller-scale channel fills. the left photo is from just inside the mouth to sowbelly gulch and shows channels with an approximate east-west orientation. the right photo is from the north side of spring canyon road between sowbelly gulch and gilson gulch, with channels oriented north-south. c) facies and bedding characteristics of the panther tongue member at the mouth to sowbelly gulch. 14 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 locality ii-a: mouth of sowbelly gulch (distributary channel fill deposits): the mouth of sowbelly gulch in spring canyon (figure 15) is approximately 3 miles (4.5 km) west of the mouth to gentile wash along depositional strike. over that distance, the character of the panther tongue member changes significantly in that the ratio of reworked channel-mouth bars (top of the unit) to delta-front gravity flows (base of unit) is much higher in the spring canyon area than at gentile wash, suggesting a more proximal setting at spring canyon. this can be seen clearly along the north side of the canyon about 0.25 mile (0.4 km) to the east of sowbelly gulch, where the section is dominated by channel-mouth bars and has only a thin interval of delta-front deposits at the base. across the mouth of the gulch and extending both to the east and to the west, channel-mouth bars have been incised by a series of distributary channels, which were subsequently filled with intrachannel bar deposits. sedimentary structures within the channel fill include clay clast molds at the base of some beds and troughor horizontally-laminated sandstone within the beds, suggestive of unidirectional channelized flow. other characteristics are indicative of tidal influence, such as thin mudstone interbeds, double mud drapes, flaser bedding, and possibly the trace fossil teredolites. some intervals appear structureless, but this may be due to uniformity of grain size or weathering that has obscured structures, rather than a true absence. a prominent characteristic of the distributary channel fills is that they show repeated channelization and filling at multiple scales with highly variable orientations. similar to gentile wash, the upper surface of the panther tongue, with respect to both distributary channel and channel-mouth bar deposits, is slightly eroded, highly burrowed, and contains small fragments of wood and bivalve shells, supporting the interpretation of a marine flooding surface. it is overlain by a thin interval of mancos shale, which again demonstrates the first major marine encroachment following deposition of the panther tongue, reinforcing interpretation as the transgressive surface for the sequence. locality ii-b: gilson gulch (clinoforms): northward-dipping clinoforms in the panther tongue at gilson gulch have traditionally been considered as lateral accretion surfaces of a deltaic distributary channel (figure 16; balsley and horne, 1980; carroll, 1987; forzoni et al., 2015). this interpretation has been challenged, suggesting the clinoforms might, instead, have been formed by the upstream accretion component of channel-mouth bars (olariu and bhattacharya, 2006). the reinterpretation is based on an overall southward progradation of the shoreline during deposition of the panther tongue and an abundance of structureless to horizontally-laminated sandstone in these deposits. on the other hand, there is a significant component of trough cross-laminated sandstone, and the 80-foot-thick (24 m) channel fill that is prominently exposed on the north side of spring canyon would need to abruptly thin to zero within a distance of less than 350 feet (107 m) in the direction of transport, as, other than a very small exposure, it is not found on the south side of the canyon. this suggests the clinoforms might actually be lateral accretion surfaces associated with an east-west oriented bend in the main distributary channel. a detailed study of paleocurrent directions needs to be conducted to resolve this discrepancy. locality ii-c: north helper escarpment from helper city cemetary (scale of down-dip facies change): the north helper escarpment can be viewed from a vacant lot on the north side of the helper city cemetery to gain perspective of the rate of down-dip facies change within a systems tract (figure 17). this is significant in order to understand how quickly changes might occur in the subsurface at reservoir or aquifer scale. the down-dip distance in the exposure is deceiving, as the outcrop has a strongly oblique orientation that is just off from depositional strike. from the 80-foot-thick (24 m) section of distributary channel sandstone at the mouth of sowbelly gulch to muddy prodelta deposits on the triangular hillside just east of the escarpment is a distance of approximately 4.7 miles (7.6 km), but represents only about 0.6 mile (1 km) in down-dip direction. the length of the escarpment is close to 1.75 miles (2.8 km), but involves around only 0.25 mile figure 16. clinoforms in the panther tongue member on the east side of the mouth to gilson gulch. w. w. little and d. a. little field localities in the book cliffs to understand sequence stratigraphic concepts 15 (0.4 km) of dip direction; therefore, facies changes are even more rapid than they appear from this outcrop. despite the minor amount of down-dip distance exhibited along the escarpment, facies changes in the panther tongue are quite evident, as they are still dominated by sandstone on the west end adjacent to u.s. highway 6, becoming almost entirely mudstone by the triangular hill at the eastern end. lateral facies changes in the storrs member locality ii-d: mouth of sowbelly gulch (distal to proximal comparison): on the west side of the mouth of sowbelly gulch, above the panther tongue, are three thin ledges containing coarsening-upward facies associations, each of which begins with mudstone at the base and is capped by sandstone (figure 18). they belong to the storrs member and are the same parasequences observed in the storrs member at gentile wash. similar to the gentile wash section, each of these coarsening-upward intervals is more poorly developed than the one beneath, in that each progressively lacks more of the upper part of the succession. however, being a more proximal location at sowbelly gulch, these deposits overall are coarser grained and better developed than their more distal exposures at gentile wash, indicating an updip change in facies. both localities show a back-stepping parasequence pattern representative of a transgressive systems tract, in which accommodation formed more quickly than it was filled. the top of the storrs member is placed where parasequences begin to demonstrate increasingly more complete formation, marking a change from a retrogradational to a strongly progradational parasequence stacking pattern. the uppermost parasequence in the storrs is only slightly less-well developed than the preceding, making this interval nearly aggradational, placing the maximum-flooding surface at the boundary between them. figure 17. north helper escarpment showing landward to basinward facies changes for the panther tongue member in a strongly oblique, close to strike-parallel, orientation. figure 18. retrogradational to aggradational parasequences within the storrs member (kss) on the west side of the mouth to sowbelly gulch. ksp and km refer to the panther tongue member and mancos shale, respectively. parasequences are marked by p1, p2, and p3. the transgressive (ts) and maximum-flooding (mfs) surfaces are shown by a blue and a green line, respectively. 16 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 lateral facies changes in the spring canyon member locality ii-e: spring canyon (landward shoreface pinchout): ascending a small hill immediately on the north side of spring canyon road and looking toward the distant high ridge to the north, two white-capped sandstones outcrop on the eastern end of the ridge (figure 19). the higher of the two sand bodies (hardscrabble parasequence) disappears along the hillside a short distance to the west. using binoculars, the landward pinchout of the upper sandstone can be seen where it terminates into an equally-thick succession of thinly-interbedded sandstones and mudstones, in which sandstone beds dip westward (landward). this terminal end of the sand body is interpreted as a barrier bar deposit (van wagoner and others, 1990). to the west of the bar are back-barrier deposits formed by washover fans and quiet-water lagoonal deposition. to the east, the parasequence transitions into a highly progradational shoreface succession. the lower sandstone is the sowbelly parasequence observed at the end of the hike in gentile wash. barrier islands form as base level rises, the longshore sediment source is reduced, and waves rework former shoreface and foreshore deposits into barrier islands that migrate landward. upon a slowing of base-level rise (or an increase in sediment supply), the back-barrier lagoon fills with sediment, and the barrier island transforms to a progradational beach system. the back-barrier lagoon fills on its landward (proximal) side by bayhead deltas, from the seaward (distal) side by washover fans, and throughout by suspension settling of mud. the spring canyon view point affords observation of the distal side of the hardscrabble lagoon. the more proximal portion can be seen at sowbelly gulch. locality ii-f: sowbelly gulch alcove north of “magazine canyon” (back-barrier deposits): at the top of the old tramline in the alcove on the east side of the road just north of magazine canyon are well-exposed back-barrier deposits associated with the hardscrabble parasequence (figure 20). this is the same parasequence observed approximately 1 mile (1.6 km) more distally with binoculars on the north side of spring canyon. the lower parasequence boundary here is formed between coal underneath and lagoonal mudstone above. sandstones in this interval are quite different from those seen up close in parasequences at gentile wash, in that they are thinly bedded, beds are lenticular in geometry, and the facies association is comprised of trough cross laminations and climbing, symmetrical, flat-topped, flaser, and interference ripple patterns. these features, along with the thin interbeds of mudstone, indicate highly variable energy conditions and tidal influence, probably associated with bayhead delta progradation into a lagoon. figure 19. landward pinch out of the hardscrabble parasequence (ps) within the spring canyon member, labeled with associated environments. figure 20. a) back-barrier lagoon deposits of the hardscrabble parasequence within the spring canyon member. b and c are enlargements of the upper and lower portions of the unit, respectively. in c, the parasequence boundary, between coal below and lagoonal mudstone above, can be seen. d and e show the thin, lenticular nature of bedding for the sandstone portion of the unit. f through i are close-up views of common sedimentary structures found within the sandstones, including flaserand troughcross lamination (f), climbing ripples (g), symmetrical ripples (h), flat-topped ripples (i), and oscillatory ripples (j). these features, along with thin beds of mudstone, indicate highly variable energy conditions and probable tidal influence. k is a generalized model (kamola and van wagoner, 1995) for blackhawk barrier and back-barrier deposits, in which sl1, sl2, and sl3 indicate sequential positions of sea level. a bayhead delta could be added to the proximal margin. w. w. little and d. a. little field localities in the book cliffs to understand sequence stratigraphic concepts 17 lateral facies changes in the aberdeen and kenilworth members locality ii-g: inside the mouth of coal canyon (progradational parasequence stacking pattern): coal canyon gives an opportunity to examine the upward and basinward shift of parasequences typical of the progradational stacking pattern in a highstand systems tract. because this is private property with a locked gate, we will focus on the overall pattern as can be seen from outside the gate. from a distance, the stratigraphy of this locality appears nearly identical to that in gentile wash, however, here, the spring canyon member parasequences that formed prevalent sandstone-dominated cliffs in gentile wash consist mostly of thinly-bedded mudstone with a thin cap of hummocky-laminated sandstone, similar to the storrs member at gentile wash (figure 21; compare with figure 6). the storrs member equivalent at this locality is in the subsurface and has completely graded laterally into the mancos shale. the pronounced shoreface successions visible in the ledges surrounding the mouth to coal creek canyon belong to the aberdeen and kenilworth members. at gentile wash, the aberdeen member formed thinner shoreface successions dominated by the presence of only the upper, more proximal parts. here, parasequences in the aberdeen show complete facies successions from offshore mudstone at the base, through shoreface and foreshore sandstones, to interbedded mudstone and sandstone or coal of coastal-plain environments, similar to the spring canyon member at gentile wash. the lower part of the kenilworth member also includes parasequences similar to those of the spring canyon at gentile wash, with upper parasequences continuing to be composed of coastal-plain deposits, similar to the entire kenilworth succession viewed from the monument turnout along u.s. highway 6. as with underlying units, farther eastward, the aberdeen and kenilworth members progressively grade down dip into the mancos shale, and the overlying sunnyside through desert members, which here are primarily coastal plain in nature, become the well-developed shoreline successions, though the sunnyside member does include a small scale, fourth-order marine incursion (figure 1). locality ii-h: beckwith plateau (progradational parasequence stacking pattern): the purpose of this locality is to further reinforce the concept of a progradational parasequence stacking pattern in the highstand systems tract, as we progress upward through the section and move farther into the basin. in the more proximal helper area, the lowest prominent sandstone was the panther tongue member (figure 5). the spring canyon and aberdeen members also exhibited prominent cliff-forming shoreline parasequences in that area (figure 6). higher units consisted of coastal-plain fluvial deposits, overlain by a thick braided river system in the castlegate sandstone (figure 13). the panther tongue disappears into the mancos shale near the town of kenilworth, and the spring canyon member does the same near the mouth of coal canyon, where the prominent sandstone ledges are the stratigraphically higher aberdeen and kenilworth members (figure 21). by the beckwith plateau, the aberdeen and lower kenilworth members have also transitioned eastward into mancos shale (figure 1), and the lowest cliff face is in the upper part of the kenilworth member, which, along with the overlying sunnyside, grassy, and desert members consists of a series of shoreface parasequences (figure 22). the thin ledge at the top of the plateau is the castlegate sandstone, which has thinned significantly since the helper region, has become finer grained, and is more characteristic of a coarse-grained meandering system. from the helper region to this locality, we have traveled obliquely to depositional dip; therefore, facies changes are actually much more rapid than they might have appeared from the few observation points. at this stop, both dipand strike -oriented exposures can be viewed simultaneously (figure 22a and 22b). depositional units in the eastern cliff face show little variation for a long distance along outcrop in terms of either lithology or thickness; however, the ledges forming the north side of the beckwith plateau show a significant thinning of sandstone from west to east. this is particularly evident in the kenilworth member (basal sandstone) that thins to the east and separates into three thinner sandstone ledges separated by tongues of the mancos shale. locality ii-i: woodside canyon – north side of curve (kenilworth member – wave-modified delta front): the uppermost two parasequences of the kenilworth member are exposed on the north side of woodside canyon road (figure 23) and together superficially appear similar to the panther tongue at gentile wash (figure 8), in that there is an interval at the base consisting of thinly interbedded sandstones and mudstones, and the succession coarsens upward into thicker lenticular bodies of sandstone capped by a flat surface and overlain by a tongue of the mancos figure 21: aberdeen (kba) and kenilworth (kbk) members of the blackhawk formation (kb) near the mouth of coal creek canyon. notice similarity to the spring canyon and aberdeen members in gentile wash, as shown in figure 6. w. w. little and d. a. little field localities in the book cliffs to understand sequence stratigraphic concepts 18 figure 22. a) depositional-dip oriented photo of the upper blackhawk formation along the northern escarpment of the beckwith plateau. the kenilworth member (lowest ledge-forming sandstone, kbk) thins eastward (basinward) and splits from one thick ledge to three thinner resistant intervals interbedded with tongues of the mancos shale (km), of which only the uppermost is still dominated by sandstone. b) depositional-strike oriented view of the same units along the escarpment on the north side of woodside canyon. depositional units show little variation in thickness or lithology. other abbreviations refer to the prairie canyon (kmp), sunnyside (kbsu), grassy (kbg), and desert (kbd) members of the blackhawk formation and to the castlegate sandstone (kc). shale. however, close inspection shows facies within the kenilworth member to differ significantly from those of the panther tongue. at the base of this outcrop, the thin sandstones interbedded with mudstones are hummockyto horizontally-laminated, suggesting storm deposition, rather than the bouma successions that dominated the basal panther tongue. these are overlain by amalgamated hummockyto parallel-laminated sandstone beds and followed by a thick interval of trough cross-laminated sandstone. these characteristics indicate wave energy was the prevailing agent, rather than sediment gravity flows. the lenticular geometry of beds in the upper part of the unit are the result of wave-reworked channel-mouth bars. near the top of the outcrop, rather than coal, this parasequence is topped by a thin interval of mudstone and sandstone interbeds, with significant burrowing of both lithologies and indications of tidal currents, such as double mud or carbonaceous material drapes. sandstones are ripple figure 23. a) complete wave-dominated deltaic succession in the upper part of the kenilworth member in woodside canyon. b) close up of hummocky sandstone interbedded with mudstone in the lower part of the outcrop. c) upper shoreface trough cross-laminated sandstone from the upper portion of the exposure. d) fine-grained back-barrier deposits. e) parallel-laminated foreshore deposits of the uppermost parasequence. 19 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 laminated and can have a pinch-and-swell geometry, indicative of rapid deposition into an otherwise quiet-water, mud-dominated environment. in addition to heavy bioturbation, the upper surface of the muddy deposits contains an abundance of small wood fragment casts. this muddy interval was deposited in a back barrier setting, similar to that of the hardscrabble parasequence observed in sowbelly gulch. following either a minor flooding event or simple delta-lobe switching and compaction, the uppermost parasequence formed a relatively thin deposit almost entirely of horizontally-laminated sandstone, most likely produced within a foreshore environment. this was followed by another, much more significant base-level rise, placing a thick tongue of the mancos shale between the kenilworth and sunnyside members. locality ii-j – east side of curve: woodside canyon (kenilworth member – distributary channel): approximately 0.25 mile (0.4 km) to the southeast of the kenilworth delta-front exposure, at the same stratigraphic level as the lower, thicker parasequence, is a thick lens-shaped sandstone body encompassed on both sides by thinner lens-shaped sandstones (figure 24). the thicker sandstone appears to be an abandoned distributary-channel fill, whereas, the smaller sandstones are channel-mouth bar deposits similar to those in the delta front. the same fine-grained capping interval and overlying foreshore sandstone of the uppermost parasequence are also present at this locality. a similar matching deposit to the southwest, across the price river, suggests the paleoflow was to the northeast. locality ii-k: valley floor north side of beckwith plateau (prairie canyon member – “detached lowstand deposits”): relatively thin, isolated bodies of sandstone, forming low benches near the base of the beckwith plateau and totally encased by mancos shale (figure 25a), have been interpreted by some as detached lowstand deposits (hampson and others, 1999, 2001). it is thought they might have formed during forced regressions, being fed through bypass channels cut into marine mudstone. others consider them to have formed as isolated highstand shelf deposits produced by turbidites (pattison, 2005a, b). dominant facies include horizontally-laminated and hummocky sandstone, with a considerable amount of contorted bedding. they are often burrowed on the upper surface and show no signs of subaerial exposure. some of these have collectively been referred to as the prairie canyon member of the mancos shale and have been correlated to the aberdeen and kenilworth members of the blackhawk formation (newman, 1985; swift and others, 1987; pattison, 2005a). on the north side of the road at this locality are small lens-shaped sandstones within the mancos shale, which are possibly examples of backfilled bypass channels (figure 25b). locality ii-l: middle mountain to gunnison butte (kenilworth member down-dip parasequence pinchout): the cliff face to the north is the south side of a ridge that extends eastward from middle mountain to gunnison butte (figure 26). this is nearly along depositional strike from locality ii-h on the north side of the beckwith plateau; therefore, this view shows the same downdip relationships but can be seen a little more closely, particularly with the aid of binoculars. the down-dip pinchout of the three upper parasequences of the kenilworth member can be easily traced from complete shoreface successions for the upper two near the base of middle mountain on the west to mostly interbedded bioturbated mudstone and thin silty intervals of the distal transition zone at gunnison butte to the east. the view from this locality gives perspective on the scale over which these facies asfigure 24. lens-shaped distributary channel-fill sandstone bounded laterally by channel-mouth bar sandstone lenses and overlain by a finer-grained back-barrier succession within the upper kenilworth member. figure 25. a) the prairie canyon member of the mancos shale is expressed as low benches associated with the aberdeen and kenilworth members of the blackhawk formation. b) channel forms within the mancos shale that are possibly backfilled lateral equivalent feeder channels to the benches in a. w. w. little and d. a. little field localities in the book cliffs to understand sequence stratigraphic concepts 20 sociations can change in a reservoir. the outcrop orientation here is close to true depositional dip, so perceptions will be much more accurate than earlier observations. the muddy resistant intervals at the base of the plateau, as well as those on at the base of battleship butte on the south side of the road, are distal remnants of the lower parasequences of the kenilworth member. locality ii-m: hatch mesa (prairie canyon member – “detached lowstand deposits” continued): thin, up to about 3-foot-thick (1 m), sandstones with sharp lower and upper contacts form broad, lobate bodies contained completely within the mancos shale (figure 27). the most common facies association is structureless sandstone overlain by parallel-laminated sandstone, though some exposures consist entirely of stacked thin beds of parallel-laminated sandstone. these suggest high energy and rapid deposition. contorted bedding and water-escape structures are abundant, indicating water saturation. below the sand bodies are thinly-bedded, coarsening-upward successions of fine-grained sediment reminiscent of distal shoreface deposits. these characteristics have led to an interpretation of subaqueous “fans” deposited on a marine “shelf ” during either forced regressive or highstand episodes of base-level change (hampson and others, 1999, 2001; pattison, 2005a, b). the sand bodies are abruptly overlain by marine mudstone of the mancos shale. figure 26. down dip cross-section of the kenilworth member through castlegate sandstone (a). the focus here is on the lowest resistant bed (b), forming the top two parasequences of the kenilworth member, which thins and eventually transitions basinward into the mancos shale (c). yellow lines bracket this interval at middle mountain on the landward and gunnison butte on the basinward ends of the outcrop. figure 27. a) the prairie canyon member (low bench) at hatch mesa. b) contact with underlying distal shoreface deposits. c) typical facies association of structureless sandstone overlain by parallel-laminated sandstone. d) secondary, but also common facies association of stacked beds of parallel-laminated sandstone. e) soft-sediment deformation, including water-escape structures. 21 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 section iii: incised channel fill deposits within the overall third-order sequence represented by the star point/blackhawk/lower castlegate succession are scattered and complex lens-shaped outcrops up to several miles wide that fill depressions cut into underlying and adjacent shoreface sandstones. these deposits consist of several stacked parasequences, each ranging from a few feet to a few tens of feet in thickness. fill environments vary from poorly developed shoreface units to multistory fluvial successions. both are heavily influenced by tidal processes, as shown by abundant reactivation surfaces, double mud or carbonaceous material drapes, flaser to wavy and lenticular bedding, heterolithic deposition, sigmoidal cross bedding, small-scale channel forms, and extensive burrowing. they are commonly capped by coal. these successions accumulated in estuarine settings, and successions can either fine or coarsen upward, depending upon openness of the shoreline to marine conditions on the basinward side and the relative contribution by fluvial systems at the landward margin. lower contacts are sharp, erosional, and considered to be fourth-order sequence boundaries (miall, 1993; o’byrne and flint, 1996; davies and others, 2006; gani and others, 2015). laterally, between incised channels, the basal contact is often parallel to underlying shoreface deposits above an interfluve surface. interfluve deposits tend to show a stronger terrestrial component than those found within the channels. an example of an interfluve succession is the earlier-described upper aberdeen member on the west side of u.s. highway 6 at the monument turnout (locality id). incised channel fill at top of the aberdeen member locality iii-a: outside the mouth of coal canyon (distant observation of incised channel geometry – aberdeen member): on the north side of coal canyon at the top of the aberdeen member, the uppermost shoreface parasequence and part of the next lower parasequence are truncated, and the incision is filled with tidally-influenced estuarine deposits (figure 28). using binoculars, above the incised channel surface can be seen inclined heterolithic strata, suggestive of a tidally-influenced fluvial succession. near the top of the channel fill, on the south end, is a thicker, lens-shaped channel-fill sandstone. a similar succession can be observed closely in soldier canyon (localities iii-b and iii-c), as described below. locality iii-b: soldier canyon – both sides of highway south of coal mine entrance (tidally-dominated shoreline deposits – upper aberdeen member): the basal part of the same incised channel fill seen at the top of the aberdeen member on the hillside at the mouth of coal canyon (locality iii-a) is exposed at road and creek level in soldier canyon (figure 29). tidal influence on shoreline systems is demonstrated by parasequences that fine upward from subtidal hummockyand sigmoidally cross-laminatfigure 28. distant (a) and closer (b, c) views of an incised channel and fill at the top of the aberdeen member. fluvial lateral accretion surfaces are evident in the closest (c) view. figure 29. a) roadcut exposure of tidally-influenced shoreline deposits, including horizontally laminated (b), hummocky (c, k), and sigmoidally-laminated (d, l) sandstone of the subtidal region; heterolithic bedding (e), flaser-laminated sandstone (f, g), and ripple marks (h, j) from the intertidal zone, and coastal plain coal (i). photos j through l are from the creek bed on the south side of the road. w. w. little and d. a. little field localities in the book cliffs to understand sequence stratigraphic concepts 22 ed sandstone to thinly interbedded mudstones and sandstones of the intertidal zone. these thin sandstones are mostly flaserto lenticularly-laminated, with lesser trough cross lamination, features indicative of alternating higher and lower energy conditions. each parasequence is capped by a carbonaceous shale or coal. locality iii-c: soldier canyon – east side of highway at former coal mine entrance (tidally-dominated fluvial deposits – upper aberdeen member): the upper part of the incised channel fill is well exposed at the former mine entrance, across the highway from the now closed soldier canyon coal mine facility (figure 30). these deposits are similar to those observed by binoculars at the mouth to coal canyon (locality iii-a), dominated by inclined heterolithic stratification, having accumulated, in part, through lateral accretion, suggesting tidally-influenced fluvial deposition. thin lenses of sandstone are flaser laminated and contain water escape structures. the succession is capped by mudstone that accumulated in overbank and swamp environments. above the channel-fill succession is the same shoreface parasequence seen at the base of the kenilworth member in coal canyon; therefore, the contact between the aberdeen and kenilworth members is a flooding surface. section iv: identification of a “complete” small-scale sequence in single outcrops locality iv-a: tusher canyon (desert member and lower castlegate sandstone): on the northwest side of a major bend in the tusher canyon streambed/road is a complete fourth-order sequence within the upper part of the desert member (figure 31). the lower contact (sequence boundary) is well expressed at the southwestern edge of the channel fill and clearly shows erosion along the base and margin. beneath and lateral to the erosional surface, are stacked shoreface deposits of the lower desert member, identified by hummocky-laminated sandstone. the fill above the sequence boundary consists of multistoried lens-shaped sand bodies. channel fill sandstone is coarser grained than underlying shoreface deposits, is characterized by low-angle trough cross stratification, some thin interbeds of mudstone, asymmetrical (current) ripples and thin veneers of carbonaceous material. the fill deposits include molds of fossilized tree trunks with teredolites burrows, indicating tidal influence. laterally (northeastward) from the sandstone lenses, the succession becomes dominated by inclined, thinly interbedded sandstones and mudstones. sandstones in this interval are structureless to trough cross stratified, figure 30. a) hillside exposure of an incised-valley fill (ivf) capping the aberdeen member (kba). the light-colored bed above is the basal shoreface parasequence of the kenilworth member (kbk). photo b shows a close up of heterolithic bedding located at the base on the west (right) side of the outcrop. photos c (pinch-and-swell bedding), d (flaser lamination), and f (water-escape structure) are from near the right margin of photo b. figure 31. incised-channel fill in tusher canyon. a) major sequence stratigraphic surfaces. abbreviations: kbd (desert member), kc (castlegate sandstone), sf (shoreface), ivf (incised channel fill), f (fluvial), tf (tidally-influenced fluvial), sb (sequence boundary), ts (transgressive surface), and mfs (maximum flooding surface). b) bedding characteristics within the incised channel fill. c) double mud and carbonaceous material drapes, d) bioturbation and wood fragment molds. e) heterolithic bedding of tidally-influenced fluvial deposits. f) coal bed found below the maximum flooding surface. 23 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 with common ripple bedding on their surfaces and load casts at the base. these are likely splay systems that developed onto a very wet floodplain or into a shallow interdistributary bay. locality iv-b: thompson canyon (desert member and lower castlegate sandstone): observed from a bedrock bench located directly on top of the yellowish petroglyphs on the west side of thompson creek, a complete fourth order sequence similar to that studied at tusher canyon can be seen in the alcove to the east of thompson creek, with a few differences from that in tusher canyon (figure 32). along the base of the alcove are stacked shoreface sandstones of the lower desert member, characterized by hummocky and trough cross laminations that are separated by continuous, planar contacts. these deposits are the capping part of a fourth-order highstand systems tract within the lower desert member. at the north end of the alcove, the lower desert member shoreface deposits are cut by an angular erosional surface and overlain by a thick lenticular sandstone body that contains lateral-accretion clinoforms, is coarser grained than the sand units beneath the surface, and contains few visible sedimentary structures. these are the lower sequence boundary and lowstand systems tract, respectively, both part of the upper desert member. in the upper part of this sand lens is a second inclined erosional surface, above which sand beds are thinner, show more distinct lateral accretion, contain mud drapes on some lateral accretion surfaces, are distinctly cross bedded, and include a significant degree of contorted bedding. laterally (south) of the upper part of the sand lens are inclined heterolithic strata very much like those described at tusher canyon. here, they form a clear coarsening-upward succession that includes abundant water-escape structures, climbing ripples, well-developed load casts, and ball and pillow structures, indicating rapid deposition under very wet and variable energy conditions. together, the upper sandstone and adjacent heterolithic deposits appear to have formed as tidally-influenced channel fill and splays onto a ponded floodplain or into an interdistributary bay. the upper part of the sand lens and associated heterolithic strata together form the transgressive systems tract, also within the upper desert member. the boundary at the facies change within the thick sandstone lens, continuing along the base of the heterolithic strata is the transgressive surface. incision by the sequence boundary through shoreface deposits is also apparent at the south side of the alcove; however, in the southern cut, the incised channel is filled entirely by inclined heterolithic strata of the transgressive systems tract. between the two ends of the alcove is an interfluve surface, where the sequence boundary and transgressive surface merge. the upper sandstone-heterolithic unit are overlain by a thin interval of carbonaceous mudstone and coal, which in turn is covered by thin sandstone sheets dominated by trough cross bedding, some of which has been highly contorted. these sandstone sheets above the coal indicate a return to shoreface conditions, and the contact between them and the underlying coal represents the most significant marine transgression for this sequence and, therefore, the maximum flooding surface. this is also the contact between the desert member below and castlegate sandstone above. a second, stratigraphically higher, interval of heterolithic strata, similar to that already described, but with thinner sandstone beds and scattered small sandstone channel fills, overlies the shoreface units and is again capped by coal and carbonaceous mudstone. teredolites borings and brackish water bivalves within this unit indicate a significant tidal influence. terrestrial strata overlying shoreface units without intervening foreshore deposits mark the upper boundary for this fourth-order sequence. note that both sequence boundaries in this outcrop are examples of where lithostratigraphic and sequence stratigraphic bounding surfaces do not coincide. the entire outcrop is capped by a coarser-grained, trough cross-bedded sandstone with distinct lateral accretion development and an erosional basal contact belonging to a purely fluvial interval of the lower castlegate sandstone. summary the star point/blackhawk/lower castlegate sequence in the book cliffs region of east-central utah provides an excellent field laboratory to view features and teach concepts associated with sequence stratigraphy. all major aspects of a stratigraphic sequence are well expressed and easily accessible. this report progresses through a complete third-order sequence from its base to the upper boundary and simultaneosly moves from proximal to distal localities within the basin. particular emphasis is placed on the nature of sequence stratigraphic surfaces, the character of stacking patterns, lateral changes in facies association within a sequence, the special case of incised channel fill deposits above a sequence boundary, and recognition of small-scale (fourth-order) sequences within the larger-scale third-order sequence. together, these features provide figure 32. incised-channel fill in thompson canyon showing sequence stratigraphic surfaces and systems tracts. abbreviations as indicated in earlier figures. w. w. little and d. a. little field localities in the book cliffs to understand sequence stratigraphic concepts 24 a detailed three-dimensional stratigraphic reconstruction in terms of accommodation production and filling in a foreland basin setting. driving directions and gps coordinates field localities are arranged in the body of the report according to topic, rather than order in which they would be encountered traversing from a single starting to ending point. in this log, they are listed first according to the report section under which they are initially encountered. second, to avoid repetition of driving directions, closely spaced localities are grouped by general area (e.g. spring canyon), followed by locality number (e.g., ii-a); therefore, localities from more than one section of the report might be listed under the same general area. within a general area, directions to localities are mostly described in the order in which they appear in the report body, not by proximity to a reference point. driving directions for each locality start at u.s. highway 6, except those in vicinity of the town of green river, which begin at main street. distances to each locality are estimated from google earth and tied to an easily identified point along the route. these are followed by gps coordinates. section i localities helper city cell towers from u.s. highway 6, near the north end of the town of helper, take exit 232 (north main street). turn west onto north main street. drive one block and turn south (left) onto uintah street. turn west (right) where uintah street dead ends into hill street. at the end of hill street turn south (left) onto duchesne street. at the base of the hill turn east (left) onto reservoir street, which will curve to the south as it climbs the hill. shortly after the curve, turn west (right) at gun club road (second unmarked street, identified by a red fire hydrant). climb the road to the hilltop. after cresting the hill, take a small utility road to the north, just past the second water tank. see gps coordinates under locality i-a. locality i-a (observation point for the north helper escarpment): park at the end of the loop next to the cell towers. walk to the edge of the hill to observe the escarpment across the valley that runs along the north side of the town of helper. gps coordinates are 39°41'04.7"n 110°51'44.8"w. gentile wash the entrance to gentile wash is located on the west side of u.s. highway 6 at mile post 231. locality i-b (gentile wash hike): park just inside the mouth to gentile wash. hike the road into the wash. localities are described in the order they are encountered during the hike and can be identified by accompanying photographs. gps coordinates for the canyon mouth are 39°42'41.73"n 110°52'4.77"w. u.s. highway 6 north of helper: locality i-c (udot road maintenance facility): park in the lot on the east side of the highway at mile 229.8. gps coordinates are 39°43'40.17"n 110°52'1.82"w. locality i-d (“monument turnout”): park in the large turnout on the east side of the highway at mile 229.4. gps coordinates are 39°43'55.73"n 110°52'14.71"w. discussion is of the roadcut on the west side of the highway. locality i-e: same as id, but discussion is of the hillside across price river on the east side of the highway. locality i-f (rolapp mine road): park at the turnout on the west side of the highway at mile 227.8. gps coordinates are 39°45'5.70"n 110°53'10.72"w. sections ii and iii localities spring canyon from u.s. highway 6 turn south onto a small paved road located about 0.25 miles (0.4 km) north of the official north main street exit. this is also north main street but is not marked at its intersection with the highway. drive between the pick & rail market on the north (right) side of north main street and the gift/rock shop on the south (left) side to where north main street curves to the east (left), and make an immediate right turn onto bryner street, heading south. turn west (right) onto canyon street and continue into spring canyon. distances are approximated from the intersection of canyon and bryner streets to gps coordinates for each locality: locality ii-a (mouth of sowbelly gulch): distance from bryner street is 3.85 miles (6.20 km). gps coordinates are 39°41'59.96"n 110°55'17.23"w. park at the spring canyon rest stop. locality ii-b (standardville/mouth of gilson gulch): distance from bryner street is 4.55 miles (7.32 km). gps coordinates are 39°41'52.70"n 110°55'55.59"w. park at the entrance to the eastern standardville road. locality ii-c: this stop is not in spring canyon. see directions to helper city cemetery below. locality ii-d: same as locality ii-a. locality ii-e (observation point for unnamed ridge east of peerless mine): distance from bryner street is 2.00 miles (3.22 km). gps coordinates are 39°41'20.72"n 110°53'47.01"w. park along spring canyon road, and hike 0.1 miles (0.16 km) to near hilltop on the north side of the road. 25 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 locality ii-f (alcove in sowbelly gulch north of “magazine canyon”): distance from spring canyon road (not bryner street) is 0.41 (0.66 km). gps coordinates are 39°42'20.76"n 110°55'12.04"w. park to side of sowbelly gulch road. helper city cemetery from u.s. highway 6, take exit 232 (north main street), and turn east onto north main street. north main street curves to the south (right). continue to janet street and turn east (left). after crossing the railroad, turn south (right) onto 2nd east/webley street. second east curves to the east and becomes spruce street. past the city park, spruce street turns into cemetery road and ends at the cemetery. locality ii-c (observation point for the north helper escarpment): park at the east end of the large open space on the north side of the cemetery. gps coordinates are 39°41'15.82"n 110°50'44.33"w. coal canyon from u.s. highway 6, at about mile 249.1, turn north onto paved east coal creek road. after about 3.2 miles (5.1 km), the paved road becomes north coal creek road, and east coal creek road is a right turn, where it continues to the north as a long, straight, well-maintained gravel road. this turn is easy to miss and is about 0.60 miles (0.97 km) after crossing a small bridge with metal side and overhanging girders. distances are approximated from where east coal creek road becomes gravel. locality ii-g (locked gate inside mouth of coal canyon): distance from paved road is 8.80 miles (14.16 km). gps coordinates are 39°42'11.26"n 110°40'42.90"w. park at the small turnout just outside the gate. locality iii-a (outside mouth of coal canyon): the incised channel fill at the top of the aberdeen member on the north side of the mouth to coal canyon can be seen clearly from coal creek road over a distance of about 2 miles (3.2 km) centered at 7 miles (11 km) from the paved road or gps coordinates 39°41'13.07"n 110°40'54.92"w. woodside canyon from u.s. highway 6, at about mile 278.1, turn east onto woodside-lower price river road, just north of the old woodside city townsite. distances to field locality stops are estimated from the intersection with highway 6. locality ii-h (beckwith plateau observation point): distance from highway 6 is 1.65 miles (2.66 km). gps coordinates are 39°15'24.63"n 110°19'19.73"w. park at intersection with minor side road that connects from the south (right) side. locality ii-i (north side of curve in road): distance from highway 6 is 4.65 miles (7.48 km). gps coordinates are 39°15'4.13"n 110°16'18.87"w. park next to a small cliff face of coarsening-upward sandstone on north (left) side of road, shortly before a curve in the road to the south (right). locality ii-j (east side of curve in road): distance from highway 6 is 5.05 miles (8.13 km). gps coordinates are 39°14'56.40"n 110°15'56.17"w. park inside fenced area on the east (left) side of the road, a short distance south of the curve. locality ii-k (valley floor north side of beckwith plateau): distance from highway 6 is 2.1 miles (3.4 km). gps coordinates are 39°15'20.85"n 110°18'49.72"w. park next to a hill of mancos shale on the north (left) side of the road. the cove/blue castle road from main street in the town of green river, turn north onto long street. drive 6.75 miles (10.86 km) and turn west (left) onto a dirt road just north of a low bench on the west side of long street. continue to the top of the ridge and along the dirt road toward the cove and blue castle. locality ii-l (middle mountain to gunnison butte observation point): distance from long street is 0.68 miles (1.09 km). gps coordinates are 39° 4'58.83"n 110° 9'32.09"w. park where a smaller road intersects and heads to the southwest toward battleship butte. hatch mesa take i-70 exit 175. turn north onto ruby ranch road, which shortly curves to the west and becomes old u.s. highway 6. about 0.25 miles (0.40 km) from the i-70 offramp, turn north (first right) onto blm road 225. locality ii-m (hatch mesa): distance from beginning of blm road 225 is about 1.50 miles (2.41 km). gps coordinates are 38°56'12.45"n 109°55'30.19"w. outcrops extend for some distance along a low bench on the west (left) side of the road. soldier canyon near mile 249.2 at the east end of the town of wellington, from u.s. highway 6, turn north onto north 2200 east. this becomes soldier creek/nine-mile canyon road. locality iii-b (roadcut on west and creek bed on east sides of the road south of mine entrance): distance from u.s. highway 6 to a roadcut on the north side of the road is 12.31 miles (19.81 km), with gps coordinates of 39°41'48.83"n 110°36'53.80"w. park across the road from the outcrop. for creek bed exposures, follow a path on the south (right) side of the road that begins at a chainw. w. little and d. a. little field localities in the book cliffs to understand sequence stratigraphic concepts 26 link fence just east of the roadcut down the hillside to the creek at gps coordinates 39°41'50.94n"110°36'49.80"w. locality iii-c (former mine entrance): distance from u.s. highway 6 is 12.65 miles (20.36 km). gps coordinates are 39°42'1.66"n 110°36'40.54"w. park in the old roadway at the mine’s former entrance. section iv localities tusher canyon from east main street near the eastern end of the town of green river, turn north onto hastings road, which changes to beach road. drive 7.50 miles (12.07 km) from the intersection with east main street and turn east (right) onto a well-maintained gravel road. there is no sign, but this is blm road 156. in 1.53 mi (2.46 km), the road enters tusher creek bed, continue to the north (left). locality iv-a (northwest side of a major bend in the creek bed/ road): distance from where road 156 enters tusher creek bed is 5.05 miles (8.13 km). gps coordinates are 39° 6'1.70"n 110° 1'50.89"w. thompson canyon from interstate i-70, take exit 187 (thompson springs). after exiting the freeway, turn north onto sego canyon/thompson canyon road and continue into the canyon. locality iv-b (thompson canyon petroglyph parking lot): drive 4.40 miles (7.08 km) from i-70 to a large pullout on the west (left) side of the road. it is just past a cattle guard to the north of a parking/rest stop for petroglyphs and a fenced coral for loading cattle. walk northward across the creek bed to a trailhead at the base of cliffs on the west side of the creek, near where the hillside curves to the east. follow the trail southward to a ledge directly over a petroglyph panel. the hike is about 0.25 miles (0.40 km). gps coordinates for the parking spot are 39° 1'8.04"n 109°42'37.14"w. acknowledgments a thank you to debora little for many years of assistance in the field and for initial editing. we thank paul anderson, thomas fisher, lisa fisher, ali jaffri, and diane kamola for constructive reviews; paul anderson and diane kamola for introducing us to some of the outcrops used in this report; blackhawk coal company and dorrell barker for access to private property; peter nielsen and tom dempster for assistance at the utah core research center; and daniel w. little for drafting original figures. references cited abreu, v., pederson, k., neal, j., and bohacs, k., 2014, a simplified guide for sequence stratigraphy—nomenclature, definitions, and method: — london, u.k., the geological society, william smith meeting, the future of sequence stratigraphy—evolution or revolution: adams, m.m., and bhattacharya, j.p., 2005, no change in fluvial style across a sequence boundary, cretaceous blackhawk and castlegate formations of central utah, u.s.a.: journal of sedimentary research, v. 75, p. 1038-1051. armstrong, r.l., 1968, sevier orogenic belt in nevada and utah: geological society of america bulletin, v. 79, p. 429-458. bally, a.w., gordy, p.l., and stewart, g.a., 1966, structure, seismic data, and orogenic evolution of southern canadian rocky mountains: bulletin of canadian petroleum geology, v. 14, p. 337-381. balsley, j.k. and horne, j.c., 1980, cretaceous wave-dominated delta systems, book cliffs, east central utah: unpublished field guide, 163 p. catuneanu, o., galloway, w.e., kendall, g.s.c., miall, a.d., posamentier, h.w., strasser, a., and tucker, m.e., 2011, sequence stratigraphy—methodology and nomenclature: newsletters on stratigraphy, v. 44, no. 3, p. 173-245. clark, f.r., 1928, castlegate, wellington, and sunnyside quadrangles, carbon county, utah: u.s. geological survey, bulletin 793, p. 165 p. coe. a.l. 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(editor), paleotectonics and sedimentation in the rocky mountain region, united states: american association of petroleum geologists memoir 41, p. 423-442. little, w.w, 1995, the influence of tectonics and eustasy on alluvial architecture, middle coniacian through campanian stratigraphy of the kaiparowits basin, utah: boulder, university of colorado, 328 p. little, w.w., 1997, tectonic and eustatic controls on cyclical fluvial patterns, upper cretaceous strata of the kaiparowits basin, utah, in hill, l.m. and koselak, j.j. (editors), learning from the land—scientific inquiry for planning and managing the grand staircase-escalante national monument: washington, d.c., u.s. department of the interior, bureau of land management, p. 489-504. liu, s., nummedal, d., and liu, l., 2011, migration of dynamic subsidence across the late cretaceous united states western interior basin in response to farallon plate subduction: geology, v. 39, p. 555-558. miall, a.d., 1993, the architecture of fluvial-deltaic sequences in the upper mesaverde group (upper cretaceous), book cliffs, utah, in best, j.l., and bristow, c.s. (editors), braided rivers: geological society special publication no. 75, p. 305-332. mclaurin, b.t., and steel, r.j., 2007, architecture and origin of an amalgamated fluvial sheet sand, lower castlegate formation, book cliffs, utah: sedimentary geology, v. 197, p. 291-311. mitchum, r.m., vail, p.r., and thompson, s., 1977, seismic stratigraphy and global changes of sea level, part 2—the depositional sequence as a basic unit for stratigraphic analysis: american association of petroleum geologists memoir 26, seismic stratigraphy – applications to hydrocarbon research, p. 53-62. neal, j., and abreu, v., 2009, sequence stratigraphy hierarchy and the accommodation succession method: geology, v. 37, p. 779-782. newman, k.f., and chan, m.a., 1991, depositional facies and sequences in the upper cretaceous panther tongue member of the star point formation, wasatch plateau, utah, in chidsey, t.c. (editor), geology of east-central utah: utah geological association publication 19, p. 65-76. newman, s.l., 1985, facies interpretations and lateral relationships of the blackhawk formation and mancos shale, east-central utah: sepm mid-year meeting field guidebook 10, p. 69-113. o’byrne, c.j., and flint, s., 1996, interfluve sequence boundaries in the grassy member, book cliffs, utah—criteria for recognition and implications for subsurface correlation, in howell, j.a. and aitken, j.f. (editors), high resolution sequence stratigraphy—innovations and applications: geological society special publication no. 104, p. 208-220. 28 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 olariu, c., and bhattacharya, j.p., 2006, terminal distributary channels and delta front architecture of river-dominated delta systems: journal of sedimentary research, v. 76, p. 212-233. pattison, s.a.j., 2005a, recognition and interpretation of isolated shelf turbidite bodies in the cretaceous western interior, book cliffs, utah, in pederson, j. and dehler, c.m. (editors), interior western united states: geological society of america field guide 6, p. 479-504. pattison, s.a.j., 2005b, storm-influenced prodelta turbidite complex in the lower kenilworth member at hatch mesa, book cliffs, utah, u.s.a.—implications for shallow marine facies models: journal of sedimentary geology, v. 75, p. 420-439. pattison, s.a.j., 2018, rethinking the incised-valley fill paradigm for campanian book cliffs strata, utah-colorado, u.s.a.—evidence for discrete parasequence-scale, shoreface-incised channel fills: journal of sedimentary research, v. 88, p. 1381-1412. pattison, s.a.j., ainsworth, r.b., and hoffman, t.a., 2007a, evidence of across-shelf transport of fine-grained sediments—turbidite-filled shelf channels in the campanian aberdeen member, book cliffs, utah, usa: sedimentology, v. 54, p. 1033-1063. pattison, s.a.j., williams, h., and davies, p., 2007b, clastic sedimentology, sedimentary architecture, and sequence stratigraphy of fluvio-deltaic, shoreface, and shelf deposits, upper cretaceous, book cliffs, eastern utah and western colorado, in raynolds, r.g. (editor), roaming the rocky mountains and environs—geological field trips: geological society of america field guide 10, p. 17-43. posamentier, h.w., and allen, g.p., 1999, siliciclastic sequence stratigraphy—concepts and applications: sepm concepts in sedimentology and paleontology, v. 9, 210 p. posamentier, h.w., and morris, w.r., 2000, aspects of the stratal architecture of forced regressive deposits, in hunt, d. and gawthorpe, r.l. (editors), sedimentary responses to forced regressions: geological society of london special publications 172, p. 19-46. schwans, p., 1995, controls on sequence stacking and fluvial to shallow marine architecture in a foreland basin, in van wagoner, j.c. and bertam, g. (editors), sequence stratigraphy of foreland basin deposits, examples from the cretaceous of north america: american association of petroleum geologists memoir 64, p. 55-102. sepm strata, accessed 2011: http://www.sepmstrata.org/page.aspx?&pageid=1&1. spieker, e.m., and reeside, j.b., 1925, cretaceous and tertiary formations of the wasatch plateau, utah: geological society of america bulletin, v. 36, p. 435-454. swift, d.j.p., hudelson, p.m., brenner, r.l., and thompson, p., 1987, shelf construction in a foreland basin—storm beds, shelf sandbodies, and slope-shelf depositional sequences in the upper cretaceous mesaverde group, book cliffs, utah: sedimentology, v. 34, p. 423-457. van de graaff, f.r., 1972, fluvial-deltaic facies of the castlegate sandstone (cretaceous), east-central, utah: journal of sedimentary petrology, v. 42, p. 558-571. van heerden, i.l., 1983, deltaic sedimentation in eastern atchafalaya bay, louisiana: center for water resources, louisiana state university, 117 p. van wagoner, j.c. 1995a, overview of sequence stratigraphy of foreland basin deposits: terminology, summary of papers, and glossary of sequence stratigraphy, in van wagoner, j.c., and bertam, g. (editors), sequence stratigraphy of foreland basin deposits, examples from the cretaceous of north america: american association of petroleum geologists memoir 64, p. ix-xxi. van wagoner, j.c. 1995b, sequence stratigraphy and marine to nonmarine facies architecture of foreland basin strata, book cliffs, utah, u.s.a., in van wagoner, j.c., and bertam, g. (editors), sequence stratigraphy of foreland basin deposits, examples from the cretaceous of north america: american association of petroleum geologists memoir 64, p. ix-xxi. van wagoner, j.c., posamentier, h.w., mitchum, r.m., vail, p.r., sarg, j.f., loutit, t.s., and hardenbol, j., 1988, an overview of the fundamentals of sequence stratigraphy and key definitions, in wilgus, c.k., hastings, b.s., kendal, c.g.s.c., posamentier, h.w., ross, c.a., and van wagoner, j.c., (editors), sea-level changes—an integrated approach: society for sedimentary geology (sepm) special publication, no. 42, p. 39-46. van wagoner, j.c., mitchum, r.m., campion, k.m., and rahmanian, v.d., 1990, siliciclastic sequence stratigraphy in well logs, cores, and outcrops—concepts for high-resolution correlation of time and facies: american association of petroleum geologists methods in exploration, no. 7, 55 p. villien, a., and kligfield, r.m., 1986, thrusting and synorogenic sedimentation in central utah: american association of petroleum geologists memoir, no. 41, p. 281-307. willis, g.c., 1999, the utah thrust system—an overview, in spangler, l.w., and allen, c.j. (editors), geology of northern utah and vicinity: utah geological association publication 27, p. 1-9. yoshida, s., 2000, sequence and facies architecture of the upper blackhawk formation and the lower castlegate sandstone (upper cretaceous), book cliffs, utah, usa: sedimentary geology, v. 136, p. 239-276. young, r.g., 1955, sedimentary facies and intertonguing in the upper cretaceous of the book cliffs, utah-colorado: geological society of america bulletin, v. 66, p. 177-201. http://www.sepmstrata.org/page.aspx?&pageid=1&1 http://www.sepmstrata.org/page.aspx?&pageid=1&1 uga-geosite-hayden-inverted-topography.indd utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors janice m. hayden dixie state university, 225 south university ave, st. george, utah 84770 hayden@dixie.edu inverted topography in st. george, washington county, utah cover image: photo looking west-northwest at the higher west black ridge, capped by the 2.3 million year old twin peaks lava fl ow, behind the lower “old airport ridge,” capped by the 1.2 million-year-old cedar bench lava fl ow. 2 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 utah geosites 2019 utah geological association publication 48 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors utah geosites showcases some of utah’s spectacular geology, both little-known localities and sites seen by visitors to utah’s many national and state parks and monuments. th e geosites refl ect the interests of the many volunteers who wrote to share some of their favorite geologic sites. th e list is eclectic and far from complete, and we hope that additional geosites will be added in the coming years. th e utah geological survey also maintains a list of geosites https://geology.utah.gov/apps/geosights/index.htm. we thank the many authors for their geosite contributions, utah geological association members who make annual uga publications possible, and the american association of petroleum geologists—rocky mountain section foundation for a generous grant for desktop publishing of these geosite papers. design and desktop publishing by jenny erickson, graphic designer, dutchiedesign.com, salt lake city, utah. th is is an open-access article in which the utah geological association permits unrestricted use, distribution, and reproduction of text and fi gures that are not noted as copyrighted, provided the original author and source are credited. see the utah geological association website, www.utahgeology.org, and creative commons https://creativecommons.org/licenses/by/4.0/ for details. suggested citation for this geosite: presidents message i have had the pleasure of working with many diff erent geologists from all around the world. as i have traveled around utah for work and pleasure, many times i have observed vehicles parked alongside the road with many people climbing around an outcrop or walking up a trail in a canyon. whether these people are from utah or from another state or country, they all are quick to mention to me how wonderful our geology is here in utah. utah is at the junction of several diff erent geological provinces. we have the basin and range to the west and the central utah hingeline and th rust belt down the middle. th e uinta mountains have outcrops of some of the oldest sedimentary rock in utah. utah also has its share of young cinder cones and basaltic lava fl ows, and ancient laccoliths, stratovolcanoes, and plutonic rocks. th e general public comes to utah to experience our wonderful scenic geology throughout our state and national parks. driving between our national and state parks is a breathtaking experience. th e “utah geosites” has been a great undertaking by many people. i wanted to involve as many people as we could in preparing this guidebook. we have had great response from authors that visit or work here in the state. several authors have more than one site that they consider unique and want to share with the rest of us. i wanted to make the guidebook usable by geologists wanting to see outcrops and to the informed general public. th e articles are well written and the editorial work on this guidebook has been top quality. i would like to personally thank mark milligan, bob biek, and paul inkenbrandt for their editorial work on this guidebook. th is guidebook could not have happened without their support. i would like to thank jenny erickson for doing the great desktop publishing and the many authors and reviewers that helped prepare the articles. your work has been outstanding and will certainly showcase the many great places and geology of utah. last, but not least, th ank you to the american association of petroleum geologists, rocky mountain section foundation for their fi nancial support for this publication. guidebook 48 will hopefully be a dynamic document with the potential to add additional “geosites” in the future. i hope more authors will volunteer articles on their favorite sites. i would like to fi ll the map with locations so that a person or family looking at the map or articles will see a great location to read about and visit. enjoy guidebook 48 and enjoy the geology of utah. peter j. nielsen 2019 uga president hayden , j., 2022, inverted topography in st. george, washington county, utah: utah geological association publication, v. 1, p. 1-11., doi: 10.31711/ugap.v1i1.99. j.m. hayden inverted topography in st. george, washington county, utah 3 introduction washington county, utah has several classic examples of inverted topography, where now topographically high ridges are capped by basalt that once flowed as lava down low stream drainages. this paper focuses on the ridges that trend north-south on either side of downtown st. george (figures 1 and 2). the city of st. george boasts three of these ridges. west black ridge capped by the twin peaks lava flow, and “old airport ridge” capped by the cedar bench lava flow are both located to the west of downtown. middleton black ridge capped by the lava ridge lava flow is located to the east of downtown. the two lower elevation ridges are now being covered with homes, some of which have spectacular views. these ridges also remain a favorite place from which to view firework displays during city celebrations and events. visiting the water tank on the red hills, north of downtown offers an excellent perspective from which to view these ridges. figure 2. geologic map of the downtown st. george area (from hayden and willis, 2011). note the north-south trending ridges both east and west of downtown that are capped by lava that flowed down stream valleys but, because of subsequent downcutting and erosion, now cap ridges, forming classic examples of inverted topography or “inverted valleys.” the sedimentary bedrock that comprise the ridges underneath the lava flows strikes generally east to west-northwest, just opposite that of the ridges, with northeastward tilting rock layers getting progressively younger from south to north. a significant portion of the map area is covered by recent quaternary sediments. see figure 4 for a-a' cross section and figure 8 for a partial stratigraphic column with unit names. 1 0 10.5 mile 1000 0 1000 2000 3000 4000 5000 6000 7000 feet 1 0 10.5 kilometer 1:24,000scale ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! 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qat5 qes qes jkm qms qag ^cp ^cp qms jn jkm jn qat3 qat4 ^cp qat2 qmt qag qat2 jks hist. prehist. up pe r m id dl e lo w er p le is to ce ne q u a t e r n a r y . ol o h qal1 qat2 qat3 qat4 qat5 qato qag qes ? qmt ? qac qae qms ? ? qaeo qca qea qeca qbcb qbl qbt ? ? qeao jn jkm jks jk jmw j^md j^m ^cp ^cs m e s o z o i c tr ia s s ic ju r a s s ic lo w er u pp er correlation of map units hist. prehist. up pe r m id dl e lo w er p le is to ce ne q u a t e r n a r y . ol o h qal1 qat2 qat3 qat4 qat5 qato qag qes ? qmt ? qac qae qms ? ? qaeo qca qea qeca qbcb qbl qbt ? ? qeao jn jkm jks jk jmw j^md j^m ^cp ^cs m e s o z o i c tr ia s s ic ju r a s s ic lo w er u pp er correlation of map units 1 0 10.5 mile 1000 0 1000 2000 3000 4000 5000 6000 7000 feet 1 0 10.5 kilometer 1:24,000scale ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! 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jkm jmw qat3 jks qeca jkm qat4 qeca jksqmt qac jkmqac jkm ^cp jkm qmt j^md qal1 qmt qea qac jkm jks qat4 jkm qat2 jkm qat3 qat4 jmw jkm jkm qac qat2 qac qacj^md qag jmw qea qag qmt qat3 qea qag qac j^md jkm qac qaeoqbl qacqbt jmw qag qag jkm qatojkm jmw qmt jks qbcb qat3 jkm qac qaeoqeca qmt jkm qmt jn qat3 qea qmt qag jkm qeca jkmqato qbl jkm qmtqac qbl jkm qag jkm qat4 qag qaeo qaeo qaeo qat3 qat5 qes qes jkm qms qag ^cp ^cp qms jn jkm jn qat3 qat4 ^cp qat2 qmt qag qat2 jks u t a h map location figure 1. map showing inverted valleys in st. george, washington county, utah. 4 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 location traveling on i-15, use exit 8 to access st. george boulevard. travel west to 1000 east then turn right to go north up the hill for one block. turn left onto red hills parkway and proceed 0.5 miles (0.8 km) before turning left into a parking lot. just to the west of the parking lot you’ll see a large water tank built in 1948 with stairs and a railing. community dances were held atop the tank on friday and saturday nights for many years (washington county historical society, website). should you miss pulling in to this first parking lot that is on the side of the stairs, another parking lot is located just to the west of the water tank. you can see the view from the parking lots, but it is more fun to climb to the top of the water tank, which is perched atop the red hills with a view down onto the city center. look left and right for a great view of the inverted topography ridges that flank downtown st. george. gps location: 37°06'46.07"n 113°34'22.70"w. elevation: 3020 feet (921 meters) above mean sea level on top of the tank. inverted valleys and basaltic lava flows the concept of inverted topography in the st. george area was first described in detail by hamblin (1963, 1970, 1987) and hamblin and others (1981). typically, the lava flowed down the bottom of stream valleys and cooled, forming a hard surface (figure 3). streams commonly re-established on top of the flows, as evidenced by thin gravel deposits, before slipping to the sides of the flows to preferentially erode the softer sedimentary bedrock. continued downcutting then left the resistant lava flows isolated as elevated, sinuous ridges called inverted valleys; thus, flows that used to form the valley floors now cap the ridges. because most small basaltic volcanoes are monocyclic, meaning that each vent produces only one eruptive cycle that may last less than a year to a few tens of years, the resistant flows document the local drainage pattern as it existed when the flow erupted (in contrast, flows from a single eruptive cycle may consist of several pulses of lava, called cooling units, that can be confused as separate flows). ages for these lava flows and their heights above major drainages provide a means for calculating long-term incision rates for major rivers and streams in the st. george area (willis and biek, 2001). the calculations reconfirm and expand on many of the findings of hamblin and others (1981), who similarly documented incision rates in the st. george basin. however, the old axiom that “the higher the lava flow is above the current drainage the older it is” is only valid when comparing flows on the same side of active faults (figure 4). the eastern middleton black ridge (lava ridge lava flow) has comparably greater inversion that do those flows to the west. this anomaly is because the ridges are on the opposite sides of the st. george fault, a late cenozoic extensional, down-to-the-west fault figure 3. diagrams showing the sequential development of inverted topography (created by jerry d. harris, dixie state university). a) water flows down a stream valley. b) a cinder cone erupts upstream, filling the stream valley with lava. c) the lava cools and hardens into a basaltic rock that is more resistant to weathering and erosion than the sedimentary bedrock it rests on. d) streams downcut and erode along the edges of the flow. e-f) what used to be the valley floor now caps a ridge, creating a classic example of inverted topography. a. c. b. d. e. f. j.m. hayden inverted topography in st. george, washington county, utah 5 that offsets strata about 400 feet (120 m) but does not offset exposed surficial deposits (hayden and willis, 2011). the comparably greater topographic inversion of this middle-aged flow (of the three flows) is directly attributable to its position on the footwall (upthrown part) of a separate, relatively more elevated structural block. thus, position on structural blocks is important when estimating relative ages of lava flows based on the amount of “topographic inversion” (“stage” designations of hamblin, 1963, 1970, 1987). hamblin (1963, 1970, 1987), best and others (1966, 1980), lowder (1973), leeman, (1974), best and brimhall (1970, 1974), hamblin and others (1981), nelson and tingey (1997), nusbaum and others (1997), smith and others (1999), downing (2000), and biek and others (2009) all described lava flows in the greater st. george area, their tectonic setting, and their petrogenesis, and proposed that the geochemical variability between individual lava flows could be explained by their derivation from the partial melting of compositionally heterogeneous lithospheric mantle, and by fractional crystallization. twin peaks lava flow of west black ridge the oldest of the three lava flows is the lower pleistocene twin peaks lava flow (qbt) which caps west black ridge above the dixie state university “d.” it is dark-gray basaltic trachyandesite with large plagioclase and quartz, and small olivine and clinopyroxene phenocrysts (hayden and willis, 2011). it has strong columnar jointing and weathers to large, angular, blocky rubble. there are two cooling units that are well exposed. geochemistry suggests that this flow, previously called west black ridge lava flow (willis and biek, 2001), erupted from vents at extensively eroded cinder cones at twin peaks, about 8 miles (13 km) to the north, and it is now considered the southernmost part of the twin peaks lava flow (biek and others, 2009). the flow yielded radiometric k-ar ages of 2.3 + 0.1 million years (ma) (best and others, 1980) and 2.24 + 0.11 ma (hamblin and figure 4. cross-section showing the three, north-south trending inverted topography ridges that bracket downtown st. george. generally, the higher the ridge is above the current drainage, the older the capping lava flow. however, note that when comparing ridges on different structural blocks, the higher elevation block commonly weathers and erodes at a faster rate than the block that has been dropped down by normal faulting. the st. george fault separates middleton black ridge (lava ridge lava flow, east of downtown) from west black ridge (twin peaks lava flow) and “old airport ridge” (cedar bench lava flow) on the west side of downtown. older rock units that are not exposed at the surface in the map area are shown on the cross-section. (from hayden and willis, 2011.) q j^m jk ^cp ^cs ^mu ^ms ^mm ^mv *^ml & ^mt pk pt pq qq qblqms qbt q j^m jk ^cp ^cs ^mu ^ms ^mm ^mv ^mt & ^ml* pk pt pq pp qat 5 santa clara river west black ridge middleton black ridgeold airport ridge 3000 2000 1000 sea level -1000 -2000 -3000 qbcb southwest elevation (feet) a 3000 2000 1000 sea level -1000 -2000 -3000 northeast elevation (feet) a'st. george fault some quaternary deposits not shown * ^mr may be present in scattered channels at base of ^mt city of st. george downtown others, 1981), and an 40ar/39ar age of 2.34 + 0.02 ma (biek and others, 2009). it is 20 to 80 feet (6-24 m) thick. this flow, accessible to pioneers because of a landslide at the southern tip of west black ridge, was quarried and pounded into the ground with a pile driver fashioned from an old cannon to make the foundation for the church of jesus christ of latter-day saints st. george temple, a process taking two years. the remnants of the road and quarry can be visited by taking an easy 2 mile (3.2 km) round-trip hike that begins at the city park located at the top of “old airport ridge” after continuing west on st. george boulevard (figure 5). (trailhead: 37°06'10.2"n 113°35'45.2"w) figure 5. photo looking west-northwest at the higher west black ridge, capped by the 2.3 million year old twin peaks lava flow, behind the lower “old airport ridge,” capped by the 1.2 million-year-old cedar bench lava flow. these flows are quite young. comparing the 4.5 billion-year-old age of earth to the length of a football field, where one inch on the field represents 1.25 million years, these two flows occurred about twoand then one-inch shy of the endzone. note the “d” on the slope between the two flows. both lava flows are partially covered by sediment deposited by eolian, colluvial and alluvial processes (qeca, figure 2). much of the ridge slope is covered by rockfall blocks as talus deposits (qmt, figure 2), but the triassic-jurassic sedimentary layers are exposed by roadcuts along bluff street and in drainages, and get progressively younger going left to right in the photo. note that petrified forest member of the chile formation underlies the south (left) end of both ridges, creating landslides (qms, figure 2) that include the lava flows. photo taken aug, 24, 2018. 6 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 lava ridge lava flow of middleton black ridge the middle-aged of the three flows is the lower pleistocene lava ridge flow (qbl) that caps middleton black ridge east of downtown (figure 6). it is a moderately jointed, dark-gray basaltic trachyandesite with prominent euhedral plagioclase phenocrysts up to 0.4 inch (1 cm) wide, common quartz and pyroxene phenocrysts, and small olivine phenocrysts (hayden and willis, 2011). it was previously called middleton lava flow (willis and biek, 2001), but petrographic and limited geochemical data suggest it is the southern extension of the lava ridge flow (biek and others, 2009). it consists of three flows in a road cut on middleton drive near the intersection with red rock road (37°07'16.26"n 113°33'02.96"w) (hamblin and best, 1970), where the more mafic oldest flow, about 5 feet (1.5 m) thick, overlies alluvial gravel deposited on bedrock. it is overlain by another well-developed alluvial gravel, a lava flow about 20 feet (6 m) thick, another gravel, and then an upper lava flow about 15 feet (4.5 m) thick. a nearby roadcut on interstate 15 reveals that only the upper flow continues south, capping middleton black ridge and forming a two-milelong (3.2 km), straight, narrow inverted valley where the flow was confined in a narrow channel, and a broad “foot” where it entered the more open channel of the ancestral virgin river (figure 7). it erupted from a group of heavily weathered cinder cones on lava ridge, about 8 miles (13 km) north of st. george. samples taken from the upper flow on middleton black ridge yielded a radiometric k-ar age of 1.5 + 0.1 ma (best and others, 1980) and an 40ar/39ar age of 1.41 + 0.01 ma (biek and others, 2009). the lower two flows are probably about the same age. it is generally 20 to 40 feet (6-12 m) thick. cedar bench lava flow of “old airport ridge” the youngest of the three lava flows included here is the lower pleistocene cedar bench lava flow (qbcb) that caps “old airport ridge” below the dixie state university “d” to the west of downtown. it is a brownish-black trachybasalt with small phenocrysts of clinopyroxene and olivine. it has prevalent columnar jointing and is strongly weathered along joints, forming a mottled texture (hayden and willis, 2011). previously called the airport lava flow (willis and biek, 2001), it is now considered the southern extension of the cedar bench lava flow because of similar geochemistry (biek and others, 2009). two cooling units are well exposed along the southeast edge of the flow, which erupted from vents at two overlapping cinder cones about 10 miles (16 km) north of st. george. the rock yielded a radiometric 40ar/39ar plateau age of 1.23 + 0.01 ma (biek and others, 2009), which fits well with regional downcutting rates (willis and biek, 2001). the flow is typically 10 to 30 feet (3-9 m) thick. these three lava flows sit atop lower pleistocene, poorly to moderately sorted, clayto boulder-size, stream-deposited gravel (qag) with mostly well-rounded cobbles and small boulders that are exotic to the quadrangle, including igneous rocks derived from the pine valley mountains. these alluvial gravels are best exposed in road cuts. figure 6. photo looking north-northwest at the northern part of middleton black ridge, which is capped by the 1.4 million-year-old lava ridge flow. like the other two lava flows, this one is partially covered by sediment (qeca, figure 2). talus (qmt, figure 2) covers much of the slope of the ridge and the jurassic-triassic section is best exposed along roadcuts and in drainages. i-15 cuts through the flow near middle of photo. photo taken mar. 1, 2007 by jerry d. harris. figure 7. photo looking north at the i-15 roadcut showing the red beds of the early jurassic-age main body of the kayenta formation (jkm) below, separated from the black lava ridge lava flow (qbl) above, by a tan layer of stream gravel (qag), indicating that the lava flowed down an ancestral stream valley. photo taken aug. 26, 2006 by jerry d. harris. j.m. hayden inverted topography in st. george, washington county, utah 7 stratigraphy these three lava flows unconformably overlie triassic to jurassic age sedimentary rocks that were tilted to the northeast as part of the broadly folded st. george syncline, a down-arched fold of rock layers (figure 8). the fold axis runs under downtown, through the south end of “old airport ridge,” and continues over the ridge to the south-southwest. there is a profound change in the direction of strike of these sedimentary rocks from west-southwest on the east side of town, to northwest on the west side of town, this change is shown nicely by the bend in the ridge of the shinarump conglomerate member of the chinle formation (trcs). triassic chinle formation, shinarump conglomerate member the upper triassic chinle formation (trc) covers the southern portion of the geologic map (figure 2). the shinarump conglomerate member (trcs), shown in the southeast corner, is a yellowish-brown, mediumto coarse-grained sandstone with locally well-developed limonite bands (“picture stone” or “landscape rock”), grading to brown pebbly conglomerate with subrounded clasts of quartz, quartzite, and chert (hayden and willis, 2011). it is mostly thick to very thick bedded with both planar bedding and low-angle cross-stratification, although thin, platy beds with ripple cross-stratification occur locally. strongly jointed with common slickensides on multiple surfaces, it contains poorly preserved petrified wood, commonly replaced in part by iron-manganese oxides. regionally it caps the chocolate cliffs step of the grand staircase (gregory, 1950) and is variable in composition and thickness because it represents braided stream-channel deposition over late triassic paleotopography (stewart and others, 1972b; dubiel, 1994). it ranges from 5 to 200 feet (1.5-60 m) thick. triassic chinle formation, petrified forest member the petrified forest member (trcp) of the chinle formation consists of highly variegated, light-brownish-gray, pale-greenish-gray, to grayish-red-purple, smectitic (swelling) shale, mudstone, siltstone, and claystone, with several lenticular interbeds of yellowish-brown, cross-bedded, resistant sandstone up to 10 feet (3 m) thick. there is a pebble to small-cobble conglomerate near the base with primarily chert and quartzite clasts. it also contains minor chert, nodular limestone, very thin coal seams and lenses as much as 0.5 inch (1 cm) thick, and locally abundant, brightly colored fossilized wood (hayden and willis, 2011). shale and mudstone layers of the petrified forest member weather to a “popcorn” surface with abundant mudcracks due to expansive clays that cause road and building foundation problems. it weathers to badland topography and is prone to landsliding along steep hillsides. it is also the local primary source of radon gas (solomon, 1992a, 1992b). it forms the well-developed strike valley of the santa clara and virgin rivers. it is well exposed only where protected from erosion by stream-terrace deposits and in road cuts along the south edges of middleton and west black ridges. where underlain by petrified forest beds, the lava flow capped ridges commonly exhibit a series of steps created by landslides. the petrified forest member was deposited in lacustrine, floodplain, and fluvial environments of a back-arc basin formed inland of a magmatic arc associated with a subduction zone along the west coast of north america. a significant portion of its sediment was supplied by volcanic ash (stewart and others, 1972b; dickinson and others, 1983; blakey and others, 1993; lucas, 1993; dubiel, 1994; decourten, 1998; lucas and tanner, 2007). it is 700 feet (215 m) thick. moenave formation, dinosaur canyon member the lower jurassic moenave formation (jtrm) trends across the middle of the map area. the dinosaur canyon member (jtrmd) is interbedded, generally thin-bedded, reddish-brown, very fine to fine-grained sandstone, very fine grained silty sandstone, and lesser siltstone and mudstone with laminated cross-beds with common ripple marks and mud cracks that forms a ledgy slope (hayden and willis, 2011). regionally it forms the base of vermilion cliffs step of the grand staircase (gregory, 1950). it is locally exposed in excavations below basalt talus near the south end of middleton and west black ridges, in stream drainages on either side of the ridges, and where protected from erosion by overlying stream-terrace deposits. several outcrops exposed by construction have revealed plant fossils (tidwell and ash, 2006). it was deposited on a broad, low floodplain that was locally shallowly flooded (fluvial mud flat) (clemmensen and others, 1989; blakey, 1994; peterson, 1994; decourten, 1998). it is 185 feet (55 m) thick (kirkland and milner, 2006). qb unconsolidated surficial deposits basaltic lava flows jn kayenta formation navajo sandstone 0–100 (0–30) 10–80 (3–24) 200+ (60+) main body jkm 810 (247) jks 115 (35)springdale sandstone mbr. moenave formation whitmore point member jmw 55 (17) dinosaur canyon member j^md 185 (55) chinle formation petrified forest member ^cp 700 (215) conglomerate mbr. ^cs 5–200 (1.5–60) e ra th er m s ys te m s er ie s formation member s ym bo l thickness feet (meters) lithology c e n . q ua t. p le is tr ec en t ju ra ss ic lo w er m e s o z o i c u pp er tr ia ss ic cross-beds cedar bench lava flows lava ridge lava flows twin peaks lava flows dinosaur tracks semionotus kanabensis swelling clays landslides petrified wood “picture stone” dinosaur tracks stratigraphic column figure 8: stratigraphic column showing the exposed rock layers within the map area. (from hayden and willis, 2011.) 8 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 moenave formation, whitmore point member the whitmore point member (jmw) is an interbedded, pale-reddish-brown, greenish-gray, and grayish-red mudstone and claystone, with thin-bedded, reddish-brown, very fine to fine-grained sandstone and siltstone. it also includes yellowish-gray, dolomitic limestone that includes several 2to 6-inch-thick (5-15 cm), bioturbated, cherty, dolomitic limestone beds with algal structures, some altered to jasper, and fossil fish scales (hayden and willis, 2011), likely of semionotid fish (milner and kirkland, 2006). it is nonresistant and poorly exposed in excavations along bluff street, in drainages next to middleton black ridge, and beneath a few protective stream terraces now largely removed by construction along riverside drive. the nearby st. george dinosaur discovery site at johnson farm (harris and milner, 2016) revealed exceptionally well-preserved theropod tracks (three-toed dinosaur tracks called eubrontes and grallator) near the base of the member. the site also includes swim tracks (kirkland and milner, 2006; milner and others, 2006), other trace fossils (lucas and others, 2006), and a variety of invertebrate fossils (lucas and milner, 2006). note the dinosaur footprint symbol near the east edge of the geologic map (fig. 2) for the location of the museum along riverside drive, which is definitely worth a visit. the member was deposited in low-energy lacustrine and fluvial environments (clemmensen and others, 1989; blakey, 1994; peterson, 1994; decourten, 1998; and milner and kirkland, 2006) and is 55 feet (17 m) thick. kayenta formation, springdale sandstone member the lower jurassic kayenta formation (jk) fills much of the northern portion of the map area beneath the lava flows and is well exposed at the base of the red hills. the lower, springdale sandstone member (jks) is mostly grayish-yellow, moderately sorted, fineto medium-grained, mediumto very thick bedded, ledgeto small-cliff-forming sandstone, with minor, thin, discontinuous lenses of intraformational conglomerate and thin interbeds of reddish-brown or greenish-gray mudstone and siltstone (hayden and willis, 2011). it contains locally abundant petrified and carbonized fossil plant remains. theropod dinosaur tracks are common in upper horizon, known as the springdale megatracksite (lucas and others, 2005; hamblin and others, 2006). the sandstone produced silver at the silver reef mining district 15 miles (24 km) to the northeast (james and newman, 1986; proctor and shirts, 1991; biek and rohrer, 2006), and has local copper and uranium mineralization (james and newman, 1986). it is resistant to erosion and forms isolated outcrops that protrude from beneath basalt talus along the slopes of the basalt-capped ridges and is completely exposed in washes east and west of middleton black ridge. it was deposited in braided-stream and minor floodplain environments (clemmensen and others, 1989; blakley, 1994; peterson, 1994; decourten, 1998; lucas and tanner, 2006). it is 115 feet (35 m) thick. kayenta formation, main body the main body of the kayenta formation (jkm) is reddish-brown, thin-bedded siltstone and mudstone interbedded with very fine to fine-grained, planer to lenticular, mottled sandstone with climbing ripple marks (hayden and willis, 2011). the upper surface of sandstone ledges is commonly bioturbated. it forms a steep, ledgy slope to ledgy cliff that is mostly covered by talus but is best exposed by construction and roadcuts along bluff street, in the drainage on the east side of middleton black ridge, and at the base of red hills along the northern edge of the map. the top of the formation is close to red hills parkway, the road you took to reach the view from the water tank. the main body of the kayenta was deposited in distal river, playa, and minor lacustrine environments (tuesink, 1989; blakey, 1994; peterson, 1994). it is 810 feet (247 m) thick. this member was quarried from the red hills by early settlers for the stone used to build the st. george temple, tabernacle, historic courthouse, washington cottonmill and many other buildings. a 0.6 mile (1 km) round-trip trail to the quarry begins at the north end of 700 west and goes along the edge of red hills golf course. navajo sandstone the lower jurassic navajo sandstone (jn) is a reddish-orange, massively cross-bedded, moderately well-cemented sandstone with well-rounded, fineto medium-grained, frosted quartz sand grains and locally common ironstone bands and concretions (hayden and willis, 2011). it forms the cliffs and slopes of the upper portion of the red hills. it is strongly jointed in two main joints sets: generally north-northeast-trending joints that are parallel, high-angle, and typically open, and northwest-trending joints that are commonly brecciated and strongly cemented. these joints are prominent in the outcrops on the north side of red hills parkway in the area of “dixie” rock. a local favorite just up the one-way road from the red hills desert garden parking lot is known as “the crack.” regionally this sandstone forms the white cliffs step of the grand staircase (gregory, 1950), although it does not happen to be white locally. the formation is also the principal aquifer in the area (clyde, 1987; hurlow, 1998; heilweil and others, 2000, 2002; rowley and dixon, 2004) and springs are common at the lower contact with the kayenta formation. the sand was deposited in a vast coastal and inland dune field with prevailing winds principally from the north, and in rare interdunal ephemeral lakes and playas (blakey, 1994; peterson, 1994). only the lower 200 feet (60 m) is present in the area shown on figure 2, but the navajo’s total thickness is 1800 to 2000 feet (550-600 m) in snow canyon state park. quaternary age sediments much of the map area is covered by quaternary-age sediment deposited by several agents of gradation: streams, wind, debris flows, and landslides. landslides (qms) and broken blocks of rock called talus j.m. hayden inverted topography in st. george, washington county, utah 9 (qmt) that cover many bedrock slopes are deposited as gravity pulls them down the hillsides. piles of unconsolidated sand are deposited by wind (eolian) processes (qes). stream systems deposit alluvial sediment in the active channel of the virgin and santa clara rivers (qal1), and old stream deposits are preserved as terraces (qat2-5) at higher elevations than current drainages. some much older terraces (qato) are not connected to present drainages. there are also units deposited by a combination of processes, such as both alluvial and colluvial (qac) in minor drainage dissected areas, and both alluvial and eolian (qea and qae) in mostly broad, flat areas. the valley underneath downtown is an older alluvial and eolian surface (qaeo), while the sediment covering part of the top of the lava flows is deposited by a combination of eolian, colluvial, and alluvial processes (qeca). summary when trying to understand landscape development, it is helpful to know that whatever landscape you are seeing is due for a change. due stands for deposition, uplift, and erosion. although these processes are not mutually exclusive, with one ending before another one can begin, it is meaningful to consider one at a time. rock layers must be made through deposition before they can experience either uplift or erosion. said in a different way, once the rock layers are deposited, they are subjected to (1) earth’s internal energy, which drives plate tectonics that roughens up the surface of the earth (uplift), and (2) external energy, which powers the hydrologic cycle that smooths the surface (erosion). the modern landscape you see is a result of that interaction. usually, to shift regimes from deposition to erosion, base level must lower in some way. in southwest utah this is commonly achieved by uplifting of rock layers, by, for example, movement on major faults. once there is any topographic relief, erosion begins. the inverted topography of these three basalt-capped ridges can be viewed as experiencing this simplified sequence twice. first, the triassic-jurassic sedimentary rock was deposited (along with younger sedimentary rock that was subsequently removed), then the area experienced compression that tilted the rock (st. george syncline); subsequent relative uplift allowed stream channels to be eroded into the layers. much later, the area experienced deposition, but this time lava flows filled the stream valleys, and the area elevations were adjusted by tension (extension) that created normal faults (st. george fault). now, with this most recent period of predominantly downcutting and erosion, the lava flows that filled the stream valleys instead cap the ridges, thus inverting the topography and giving the ridges the landform name of “inverted valleys” or “inverted topography.” as you look out on the ridges and the valley of downtown st. george from the vantage point of the water tank on the red hills, it can be quite a mind-expanding experience to think of how different the topography looked not so very long ago. but then again, many concepts in geology are mind-expanding. now go take a closer look. references best, m.g., and brimhall, w.h., 1970, late cenozoic basalt types in the western grand canyon region, in hamblin, w.k., and best, m.g., editors, the western grand canyon district: utah geological society guidebook to the geology of utah, no. 23, p. 57-74. best, m.g., and brimhall, w.h., 1974, late cenozoic alkalic basaltic magmas in the western colorado plateaus and the basin and range transition zone, u.s.a., and their bearing on mantle dynamics: geological society of america bulletin, v. 85, no. 11, p. 1677-1690. best, m.g., hamblin, w.k., and brimhall, w.h., 1966, preliminary petrology and chemistry of late cenozoic basalts in the western grand canyon region: brigham young university geology studies, v. 13, p. 109-123. best, m.g., mckee, e.h., and damon, p.e., 1980, space-time-composition patterns of late cenozoic mafic volcanism, southwestern utah and adjoining areas: american journal of science, v. 280, p. 1035-1050. biek, r.f., and rohrer, j.c., 2006, geology, mining history, and reclamation of the silver reef mining district, washington county, utah, in bon, r.l., gloyn, r.w., and park, g.m., editors, mining districts of utah: utah geological association publication 32, p. 479-512. biek, r.f., rowley, p.d., hayden, j.m., hacker, d.b., willis, g.c., hintze, l.f., anderson, r.e., and brown, k.d., 2009, geologic map of the st. george and east part of the clover mountains 30' x 60' quadrangles, washington and iron counties, utah: utah geological survey map 242, 2 pl., 101 p., scale 1:100,000. blakey, r.c., 1994, paleogeographic and tectonic controls on some lower and middle jurassic erg deposits, colorado plateau, in caputo, m.v., peterson, j.a., and franczyk, k.j., editors, mesozoic systems of the rocky mountain region, usa: denver, colorado, rocky mountain section of the society for sedimentary geology, p. 273-298. blakey, r.c., bashem, e.l., and cook, m.j., 1993, early and middle triassic paleogeography, colorado plateau and vicinity, in morales, m., editor, aspects of mesozoic geology and paleontology of the colorado plateau: museum of northern arizona bulletin 59, p. 13-26. clemmensen, l.b., olsen, h., and blakey, r.c., 1989, erg-margin deposits in the lower jurassic moenave formation and wingate sandstone, southern utah: geological society of america bulletin, v. 101, p. 759-773. clyde, c.g., 1987, groundwater resources of the virgin river basin in utah: logan, utah state university, utah water research laboratory, 104 p. 10 m. milligan, r.f. biek, p. inkenbrandt, and p. nielsen, editors 2019 utah geological association publication 48 decourten, f., 1998, dinosaurs of utah: salt lake city, university of utah press, 300 p. dickinson, w.r., beard, s.l., brakenridge, g.r., erjavec, j.l., ferguson, r.c., inman, k.f., knepp, r.a., lindberg, f.a., and ryberg, p.t., 1983, provenance of north american phanerozoic sandstones in relation to tectonic setting: geological society of america bulletin, v. 94, p. 222-235. downing, r.f., 2000, imaging the mantle in southwestern utah using geochemistry and geographic information systems: las vegas, university of nevada, m.s. thesis, 128 p. dubiel, r.f., 1994, triassic deposystems, paleogeography and paleoclimate of the western interior, in caputo, m.v., peterson, j.a., and franczyk, k.j., editors, mesozoic systems of the rocky mountain region, usa: rocky mountain section of society of economic paleontologists and mineralogists, p. 132-168. gregory, h.e., 1950, geology and geography of the zion park region, utah and arizona: u.s. geological survey professional paper 220, 200 p. hamblin, a.h., lockley, m.g., and milner, a.r.c., 2006, more reports of theropod dinosaur tracksites from the kayenta formation (lower jurassic), washington county, utah—implications for describing the springdale sandstone, in harris, j.d., lucas, s.g., spielmann, j.a., lockley, m.g., milner, a.r.c., and kirkland, j.i., editors, tracking dinosaur origins—the triassic/jurassic terrestrial transition: new mexico museum of natural history and science bulletin 37, p. 276-281. hamblin, w.k., 1963, late cenozoic basalts of the st. george basin, utah, in heylmun, e.b., editor, guidebook to the geology of southwestern utah: intermountain association of petroleum geologists 12th annual field conference, p. 84-89. hamblin, w.k., 1970, late cenozoic basalt flows of the western grand canyon, in hamblin, w.k., and best, m.g., editors, the western grand canyon district: utah geological society guidebook to the geology of utah, no. 23, p. 21-38. hamblin, w.k., 1987, late cenozoic volcanism in the st. george basin, utah: geological society of america centennial field guide—rocky mountain section, p. 291-294. hamblin, w.k., and best, m.g., 1970, road log, in hamblin, w.k., and best, m.g., editors, the western grand canyon district: utah geological society guidebook to the geology of utah, no. 23, p. 93-154. hamblin, w.k., damon, p.e., and bull, w.b., 1981, estimates of vertical crustal strain rates along the western margins of the colorado plateau: geology, v. 9, p. 293-298. harris, j. d., and milner, a. r. c., 2016, tracks in deep time: the st. george dinosaur discovery site at johnson farm; university of utah press, 112 p. hayden, j.m. and willis, g.c., 2011, geologic map of the st. george 7.5' quadrangle, washington county, utah: utah geological survey map 251dm, 20 p., 2 plates, scale 1:24,000. heilweil, v.m., freethey, g.w., stolp, b.j., wilkowske, c.d., and wilberg, d.e., 2000, geohydrology and numerical simulation of ground-water flow in the central virgin river basin of iron and washington counties, utah: utah department of natural resources technical publication 116, 139 p. heilweil, v.m., watt, d.e., solomon, d.k., and goddard, k.e., 2002, the navajo aquifer system of southwestern utah, in lund, w.r., editor, field guide to geologic excursions in southwestern utah and adjacent areas of arizona and nevada: u.s. geological survey open-file report 02-172, p. 105-130. hurlow, h.a., 1998, the geology of the central virgin river basin, southwestern utah, and its relation to ground-water conditions: utah geological survey water-resources bulletin 26, 53 p., 6 plates. james, l.p., and newman, e.w., 1986, subsurface character of mineralization at silver reef, utah, and a possible model for ore genesis, in griffen, d.t., and phillips, w.r., editors, thrusting and extensional structures and mineralization in the beaver dam mountains, southwestern utah: utah geological association publication 15, p. 149-158. kirkland, j.i., and milner, a.r.c., 2006, the moenave formation at the st. george dinosaur discovery site at johnson farm, in harris, j.d., lucas, s.g., spielmann, j.a., lockley, m.g., milner, a.r.c., and kirkland, j.i., editors, tracking dinosaur origins—the triassic/jurassic terrestrial transition: new mexico museum of natural history and science bulletin 37, p. 289-309. leeman, w.p., 1974, late cenozoic alkali-rich basalt from the western grand canyon area, utah and arizona—isotopic composition of strontium: geological society of america bulletin, v. 85, p. 1691-1696. lowder, g.g., 1973, late cenozoic transitional alkali olivine tholeiitic basalt and andesite from the margin of the great basin, southwest utah: geological society of america bulletin, v. 84, no. 9, p. 2993-3012. lucas, s.g., 1993, the chinle group—revised stratigraphy and biochronology of upper triassic nonmarine strata in the western united states, in morales, m., editor, aspects of mesozoic geology and paleontology of the colorado plateau: museum of northern arizona bulletin 59, p. 27-50. lucas, s.g., lerner, a.j., milner, a.r.c., and lockley, m.g., 2006, lower jurassic invertebrate ichnofossils from a clastic lake margin, johnson farm, southwest utah, in harris, j.d., lucas, s.g., spielmann, j.a., lockley, m.g., milner, a.r.c., and kirkland, j.i., editors, tracking dinosaur origins—the trij.m. hayden inverted topography in st. george, washington county, utah 11 assic/jurassic terrestrial transition: new mexico museum of natural history and science bulletin 37, p. 128-136. lucas, s.g., and milner, a.r.c., 2006, conchostraca from the lower jurassic whitmore point member of the moenave formation, johnson farm, southwestern utah, in harris, j.d., lucas, s.g., spielmann, j.a., lockley, m.g., milner, a.r.c., and kirkland, j.i., editors, tracking dinosaur origins—the triassic/jurassic terrestrial transition: new mexico museum of natural history and science bulletin 37, p. 421-423. lucas, s.g., and tanner, l.h., 2007, tetrapod biostratigraphy and biochronology of the triassic-jurassic transition on the southern colorado plateau, usa: palaeogeography, palaeoclimatology, palaeoecology, v. 244, p. 242-256. lucas, s.g., tanner, l.h., and heckert, a.b., 2005, tetrapod biostratigraphy and biochronology across the triassic-jurassic boundary in northeastern arizona, in heckert, a.b., and lucas, s.g., editors, vertebrate paleontology in arizona: new mexico museum of natural history and science bulletin 29, p. 84-94. milner, a.r.c., and kirkland, j.i., 2006, preliminary review of the early jurassic (hettangian) freshwater lake dixie fish fauna in the whitmore point member, moenave formation in southwest utah, in harris, j.d., lucas, s.g., spielmann, j.a., lockley, m.g., milner, a.r.c., and kirkland, j.i., editors, tracking dinosaur origins–the triassic/jurassic terrestrial transition: new mexico museum of natural history and science bulletin 37, p. 510-521. milner, a.r.c., lockley, m.g., and kirkland, j.i., 2006, a large collection of well-preserved theropod dinosaur swim tracks from the lower jurassic moenave formation, st. george, utah, in harris, j.d., lucas, s.g., spielmann, j.a., lockley, m.g., milner, a.r.c., and kirkland, j.i., editors, tracking dinosaur origins—the triassic/jurassic terrestrial transition: new mexico museum of natural history and science bulletin 37, p. 315-328. peterson, f., 1994, sand dunes, sabkhas, streams, and shallow seas—jurassic paleogeography in the southern part of the western interior basin, in caputo, m.v., peterson, j.a., and franczyk, k.j., editors, mesozoic systems of the rocky mountain region, usa: denver, colorado, rocky mountain section of the society for sedimentary geology, p. 233-272. proctor, p.d., and shirts, m.a., 1991, silver, sinners and saints—a history of old silver reef, utah: provo, utah, paulmar, inc., 224 p. rowley, p.d., and dixon, g.l., 2004, the role of geology in increasing utah’s ground-water resources from faulted terranes—lessons from the navajo sandstone, utah, and the death valley flow system, nevada-california, in spangler, l.e., editor, ground water in utah—source, protection, and remediation: utah geological association publication 31, p. 27-41. solomon, b.j., 1992a, geology and the indoor-radon hazard in southwestern utah, in harty, k.m., editor, engineering and environmental geology of southwestern utah: utah geological association publication 21, p. 164-172. solomon, b.j., 1992b, environmental geophysical survey of radon-hazard areas in the southern st. george basin, washington county, utah, in harty, k.m., editor, engineering and environmental geology of southwestern utah: utah geological association publication 21, p. 173-191. stewart, j.h., poole, f.g., and wilson, r.f., 1972b, stratigraphy and origin of the chinle formation and related upper triassic strata in the colorado plateau region, with a section on sedimentary petrology by r.a. cadigan and on conglomerate studies by w. thordarson, h.f. albee, and j.h. stewart: u.s. geological survey professional paper 690, 336 p. tidwell, w.d., and ash, s.r., 2006, preliminary report on the early jurassic flora from the st. george dinosaur discovery site, utah, in harris, j.d., lucas, s.g., spielmann, j.a., lockley, m.g., milner, a.r.c., and kirkland, j.i., editors, tracking dinosaur origins—the triassic/jurassic terrestrial transition: new mexico museum of natural history and science bulletin 37, p. 414-420. tuesink, m.f., 1989, depositional analysis of an eolian-fluvial environment—the intertonguing of the kayenta formation and navajo sandstone (jurassic) in southwestern utah: flagstaff, northern arizona university, m.s. thesis, 189 p. washington county historical society website, https://wchsutah. org/churches/st-george-temple.php, accessed jan. 30, 2019. washington county historical society website, https://wchsutah. org/water/stgeorge-old-water-tank.php accessed jan. 30, 2019. willis, g.c., and biek, r.f., 2001, quaternary incision rates of the colorado plateau and major tributaries in the colorado plateau, utah, in young, r.a., and spamer, e.e., editors, colorado river origin and evolution—proceedings of the symposium held at grand canyon national park in june 2000: grand canyon association monograph 12, p. 119-123. oviattgslbasin.pub 1 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2023 utah geological association publication 51 abstract the great salt lake-bonneville basin has contained lakes for many millions of years and has been hydrographically closed for most of its history. lakes in the lacustrine system have ranged from saline to fresh, and from shallow to deep. tectonics, specifically crustal extension, which began roughly 20 million years ago as part of the formation of the basin and range province, is the cause of lake-basin formation. much of the rock record of lakes from miocene time is faulted and has been eroded and/or buried. pliocene and quaternary lakes are better known. for much of the past ~5 ma the basin has probably appeared similar to today, with a shallow saline terminal lake in a dry desert surrounded by mountains. freshwater marshes and fluvial systems existed on the basin floor during part of the past ~5 ma, probably were caused by the lack of inflow from the upper bear river during the neogene period and most of the pleistocene epoch (that river was diverted into the basin during the late pleistocene), combined with a warm and dry climate. the largest deep-lake cycles were caused by changes to a cold and wet climate, which affected the water budget of the lake system and were correlated with periods of global glaciation. based on limited data, the total length of time deep lakes existed in the basin is thought to be less than 10% of the past ~773 ka. lake bonneville, the most-recent of the deep-lake cycles, was probably the deepest and largest manifestation of the lake system in the history of the basin. named deep-lake cycles during the past ~773 ka, are lava creek (~620 ka), pokes point (~430 ka), little valley (~150 ka), cutler dam (~60 ka), and bonneville (~30 -13 ka). of the quaternary deep-lake cycles, only lake bonneville is represented by lacustrine landforms, outcrops, and cores of offshore deposits; no landforms from older deep-lake cycles exist (some may be buried under lake bonneville deposits but are not visible at the surface), and pre-bonneville lakes are represented by sediments in limited outcrops and drill holes (including a set of cores taken by a.j. eardley in the mid 20th century). during the past ~773 ka, deep-lake cycles were correlated with changes in the total volume of global glacial ice; the available evidence indicates that prior to ~773 ka deep-lake cycles were rare. late neogene and quaternary lacustrine history of the great salt lake-bonneville basin charles g. oviatt 3department of geology, kansas state university, manhattan, kansas, joviatt@ksu.edu introduction this paper discusses lakes of pliocene through quaternary age (figure 1) that have occupied the great salt lake-bonneville basin (gsl-bb). the gsl-bb is located in the eastern basin and range province and is part of the great basin (figure 2). all lakes in the gsl-bb during its long history, which includes the past 15 or 20 million years (ma, mega annum; figures 1 and 2), should be thought of as parts of a single lacustrine system — this concept is extrapolated from that of atwood and others (2016), who applied it to lake bonneville (lb) and post-lb great salt lake (gsl). lake size varied over time in response to tectonic and climatic changes; sometimes the lake was shallow and saline to hypersaline, and uncommonly it grew in depth, volume, and surface area to become brackish to fresh. an important observation emphasized in this paper is that during the pliocene and quaternary epochs the gsl-bb lacustrine system spent more time as a shallow lake than as a deep lake; deep-lake versions of the system have been relatively short lived and uncommon. a more quantitative approach to this observation is discussed below. it is not possible to give precise definitions of “deep lake,” and “shallow lake,” but for this paper, “deep” lakes are regarded as being much bigger than modern gsl. in this general sense, “deep” lakes might range from a lake roughly the size of the cutler dam (cd) lake (see below for discussions of named lakes in the gsl-bb), roughly 60 m higher than the average elevation of modern gsl (1280 m), to the size of lb, almost 350 m higher than modern gsl in the middle of the basin. “shallow” lakes would look similar to modern gsl, with average maximum depth near 10 m, but might be shallower than that or several tens of meters higher. with lake level constantly changing in the closed basin (on time scales longer than a few weeks), lake size is difficult to precisely define if shorelines are not available. 10.31711/ugap.v51i.133 2 c.g. oviatt late neogene and quaternary lacustrine history of the great salt lake-bonneville basin for this paper, the gsl-bb includes the subbasins that collectively comprise the bonneville basin of late pleistocene to modern age. the subbasins are: (1) the great salt lake (gsl) basin, (2) the great salt lake desert (gsld) basin (separated from gsl by low divides), and (3) the sevier basin (figure 2). major streams entering the system are the sevier and beaver rivers in the sevier basin, and the provo/ jordan, weber, and bear rivers in the gsl basin (figure 2). all these rivers head in the high mountains and plateaus along the eastern margin of the basin. no major rivers flow into the gsld basin, although a few rivers that are ephemeral today, were probably perennial during deep-lake episodes (streams such as thousand springs creek, grouse creek, and deep creek [the deep creek that heads in eastern nevada] built impressive deltas into lb). an upward component of groundwater flow (stephens, 1974; fitzmayer and others, 2004), and the observation that the mud of the mudflats is moist everywhere (except maybe for a few centimeters at the surface where the wind has dried it), indicates that the modern gsld is a gigantic groundwater-discharge, or evapotranspiration area (in springs flow is concentrated). within the subbasins are smaller closed basins, such as puddle valley and tule valley in the gsld basin, and cedar valley and rush valley in the gsl basin. all these hydrographically closed basins and subbasins exist because of neogene and quaternary faulting. the wasatch fault bounds the eastern margin of the gsl-bb and the great basin (and basin and range province), and has the greatest total offset of any fault system in the gsl-bb. the wasatch fault accounts for the major mountain front of the wasatch range. the maximum thicknesses of neogene and quaternary sediment in the gsl-bb vary from place to place, and the sediments may be ~4 km thick, or more, in some places (hintze and kowallis, 2021). details of the faulting history are beyond the scope of this paper, but faulting is an important long-term control on the lacustrine history. this paper summarizes what is currently known about the lacustrine history of the gsl-bb for the past ~5 ma. as is typical of geologic information, more is known about relatively recent events than about older events. the shapes and sizes of the older lacustrine basins within the gsl-bb are poorly known because of continued tectonic deformation. miocene tectonics and deposition extension associated with the neogene and quaternary tectonics of the basin and range province, including the gsl-bb in the eastern part of the provfigure 1. approximate ages (in ma) for subdivisions of the cenozoic era (after walker and others, 2018). in recent interpretations, the tertiary period (as it was called for many years) is now regarded as consisting of two geologic periods, the paleogene and neogene. the events discussed in this paper occurred during the neogene and quaternary periods (the miocene, pliocene, pleistocene, and holocene epochs). 3 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 2. map showing approximate modern drainage divides for the subbasins within the gsl-bb. pleistocene drainage divides were probably similar, but drainage divides for neogene basins are not known. the gsld basin has less than a meter of closure and it is separated from the gsl basin by two low thresholds, which are nearly imperceptible on the mudflats at identical elevations (1285 m); when first discussed by eardley and others (1957) only one threshold, the southern one, was recognized and was called the “desert threshold” (in some cases it is now called the “eardley threshold”). the approximate outline of lb (the bonneville shoreline) is shown for reference, as are major rivers that entered the basins from the east side. modern lakes are labeled. approximate locations of the eardley cores and the sevier-basin cores are shown with red dots (s28 = s28; s = saltair; b = burmester; k = knolls; w = wendover; pod = pit of death; br = black rock). the low point on the divide between the sevier basin and the gsld, is the old river bed threshold (orbt); flow from the sevier basin entered the gsld basin during the late pleistocene. major rivers are shown schematically with dashed lines. l = lakeside; slc = salt lake city. 4 c.g. oviatt late neogene and quaternary lacustrine history of the great salt lake-bonneville basin ince, began roughly 20 ma (hintze and kowallis, 2021). by at least 15 ma, lake basins had begun to form in the eastern basin and range province (patton and lent, 1980; taylor and bright, 1987; oaks and others, 1999; bortz, 2002; janecke and others, 2003; long and others, 2006; mcclellan and smith, 2020). despite ongoing tectonism and many details of the topography that have changed between late neogene time and the present, the general configurations of mountains and basins is probably similar now to what it was 5 ma ago (hintze and kowallis, 2021). some significant regional-scale changes have occurred in the slc-bb during the time period in question, such as river diversions that have changed the water budgets of lakes (discussed below). the basin and range province is still tectonically extending today (wguep, 2016; utah geological survey, 2023). thick accumulations of lacustrine and associated deposits of miocene age are exposed in such areas as cache valley, utah and idaho (oaks and others, 1999; janecke and others, 2003; mcclellan and smith, 2020), and goose creek, id and nv (perkins and others, 1995), and in many other places within or near the modern gsl-bb. janecke and others (2003) present good evidence that neogene lake basins developed in the area now called northeastern utah and southeastern idaho, many of which were associated with the evolving bannock detachment fault system. it is likely that multiple individual basins were integrated into one large gsl-bb by pleistocene times in response to continuing tectonism during the several-million-year period, but the details of the lacustrine history are still being discovered. although there is no question that miocene lakes existed in the gsl-bb, the outlines of individual basins and the shorelines of those old lakes are not preserved or are covered. the gsl-bb is large (figure 2), but it’s not an ocean basin — because of the huge spatial variability in geology, biology, topography, etc., within the basin, a core taken from one point, or one outcrop, are unlikely to contain sediments that look similar to those in cores or outcrops several kilometers away. one core or outcrop, although it may contain valuable information, is not likely to record the geologic history of the entire basin. to construct a complete geologic history of the basin, information from multiple sources throughout the lake basin needs to be integrated, a process that takes a long time and efforts by multiple generations of scientists. water budget the water budget of lakes in the gsl-bb is a fundamental consideration. although precise measurements for many of the variables in water-budget equations for modern lakes are available, the values of important variables for older lakes can only be generally estimated. water budget (or balance) can be expressed in many ways, but a simple equation shows water inflows equal to water outflows, plus-or-minus changes in storage of water in the lake (hutchinson, 1957). in the case of a hydrographically closed lake, water does not exit the system except by evaporation (there is no river or groundwater outflow). gsl is a closed-basin (or terminal, or endorheic) lake, so it has no surface outflow, and groundwater outflow is assumed to be zero (arnow and stephens, 1990). the relationship between volume and surface area (and elevation) in the modern gsl-bb is nearly linear (wambeam, 2001). for most of its history the gslbb has been hydrographically closed and short-term changes in lake level have been correlated with changes in climate. tectonics and paleoclimate in the bonneville basin the rate of tectonic deformation and sediment infilling compared to the water balance should be considered in tectonic basins (bohacs and others, 2000). if climate in a basin favors a positive water balance, where inflows exceed outflows, a basin might appear to be open, but if tectonic subsidence of the basin floor is relatively rapid and the rate of sediment infill is low the basin might remain hydrographically closed even if inflows exceed outflows. the gsl-bb would be classified as “underfilled” by bohacs and others (2000, their figure 7; bernau, 2022). in an underfilled basin plenty of space is available for water and sediment to accumulate, and that large volume of unfilled space keeps the basin from overflowing. in hydrographically closed basins, the water that remains in the basin after most of it has evaporated becomes increasingly salty over time (hardie and eugster, 1970). over its many-million-year history, the gsl-bb has remained underfilled with respect to sediment, and hydrographically closed most of the time. the rate of tectonic deformation in the gsl-bb is great enough that only one period is known where the basin was hydrographically open while remaining sedimentologically closed. this occurred when late pleistocene lb was overflowing at red rock pass into the snake river drainage basin as the provo shoreline formed (gilbert, 1890). during that period (possibly about 1000 to 3000 years in duration) climate was cooler and wetter than today and the lake was deep. 5 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 neogene climate of the gsl-bb was probably similar to that of today, although the mean annual precipitation may have been generally lower and temperature somewhat higher (moutoux, 1995; moutoux and davis, 1995, their figures 3 and 4; davis and moutoux, 1998; davis, 2002). these paleoclimate interpretations were based on pollen from samples of cuttings from drill holes in gsl (table 1); the dating was not precise, but the pollen allowed for interpretations of generalized climatic conditions during the pliocene and early pleistocene time. quaternary climate in the gsl-bb has been widely variable (rhode, 2016). when deep-lake cycles occurred, climate was relatively cool and wet and during times when the lake system was shallow, climate was relatively warm and dry (davis and moutoux, 1988; rhode, 2016). freshto brackish-water marshes on the basin floor kowalewska and cohen (1998), in an analysis of ostracodes (small crustaceans, typically about 1 mm in size) from cuttings taken from the same gsl drill holes that yielded the pollen samples mentioned above, found evidence of freshwater wetlands (marshes) and fluvial environments at various locations on the floor of the basin at different poorly dated times during the past 5 ma. during the holocene, the water of gsl has been hypersaline and has not supported ostracodes (thompson and others, 2016), but at the locations of the drill holes studied by kowalewska and cohen (1998), freshwater conditions existed at times, and at other times the same places were occupied by shallow lakes, some of which were saline. just the presence of freshwater ostracodes on the floor of the gsl-bb, which are not part of deep-lake faunas (delorme, 1969; forester, 1987), indicates hydrologic conditions much different than those of today. kowalewska and cohen (1998) compared their ostracode results with pollen results described by moutoux and davis (1995), and they were not able to find meaningful correlations between the ostracode interpretations and pollen interpretations of the paleoclimate in the gsl-bb. one possibility to help explain why marshes and/or freshwater fluvial systems might appear low in the basin if it was hydrographically closed, is that, because of local tectonic activity, the basin floor was probably not smooth and uniform, but instead consisted of multiple shallow depressions separated by low ridges and hills. fresh river water could flow into some depressions (and feed freshwater marshes and/or streams), but not into others, which might contain shallow saline lakes. the number and distribution of drill holes from which cutfigure 3. the eardley cores. this figure was assembled using data from published (eardley and gvosdetsky, 1960; eardley and others, 1973; williams, 1994; oviatt and others, 1999) and unpublished sources (shuey, 1971; thompson and oviatt, 1995, notes from core examinations; j. bright, d.s. kaufman, and r.m. forester-late ‘90s data on ostracode faunas and amino acid results for samples collected by thompson and oviatt). 6 c.g. oviatt late neogene and quaternary lacustrine history of the great salt lake-bonneville basin tings were obtained is not sufficient to determine if this explanation is viable; also the available geochronological control is not good enough to make reliable correlations between cores. in a different data set (core gsl00-4; balch and others, 2005, their figure 6), the youngest ostracode fauna from a freshwater marsh on the basin floor is on the order of ~45 ka (kilo [1000] annum; presumably prior to the diversion of the upper bear river and cache valley tributaries into gsl (see discussion below). an important contributing cause of the appearance of marshes and/or freshwater fluvial systems on the basin floor involves the diversion into the gslbb of the upper bear river plus the rivers that drain cache valley. these rivers contribute water and dissolved solids to modern gsl. the precise ages of incisions of canyons along the path of the bear river have not been totally resolved, but it’s likely that the incisions occurred during the late pleistocene. according to pederson and others (2016, their table 2.1) oneida narrows (figure 2; on the topographic divide of cache valley) was fully incised, allowing the upper bear river to enter cache valley, based on optically stimulated luminescence ages, after 55.0 ± 5.6 ka and before 48.9 ± 6.9 ka (a round number near the middle of that overlapping range is 50 ka). prior to the incision of oneida narrows, the upper bear river had a complicated history involving flow into the portneuf river (a tributary of the snake river) and ponding upstream from oneida narrows to form lake thatcher (gilbert, 1890; bright, 1963; pederson and others, 2016). another canyon through which the upper bear river now flows into the gsl-bb is the cutler narrows (“gate of bear river,” gilbert, 1890, his plate xxx), where the upper bear river plus its cache valley tributaries exit cache valley. the exact timing of the incision of cutler narrows, and the mechanism of the incision, has not been determined, but all the incision (it’s possible the incision occurred in stages?) probably was not completed until sometime after the cd lake cycle (that is, after ~60 ka; oviatt and others, 1987; kaufman and others, 2001; oaks and others, 2024). the incision of cutler narrows was traditionally interpreted to be the result of superposition probably combined with antecedence (williams, 1958; maw, 1968). the word “anteposition” was coined by hunt (1982) to describe situations where incision began with superposition and continued because of tectonic uplift across the path of the river. williams (1958), maw (1968), and hunt (1982) did not give specific ages or directly discuss which river was superposed to ultimately create cutler narrows. movement on the wasatch and west cache valley fault zones would easily account for tectonic uplift of the junction hills bedrock block across a superposed river. if the anteposition interpretation were correct, however, the river that was superimposed across the cutler divide could not have been the bear river if the upper bear did not incise oneida narrows and enter cache valley until about 50 ka. more work is needed on the geologic history of the cutler narrows. oaks and others (2018; 2024) suggested the presence of lakes in cache valley separate from lakes in the gsl-bb, but the precise ages and characteristics of those cache valley lakes have not been determined. this study adopts the relative age of incision of cutler narrows as younger than the cd lake cycle and older than the lb lake cycle (possibly close to 30 ka, but this has not been scientifically tested). the upper bear river, plus the total discharge of rivers that enter cache valley from the nearby mountains, plus discharge from the malad river, accounts for about a third of the modern annual inflow to gsl (oviatt and others, 1987; arnow and stephens, 1990). without input from the upper bear river plus the cache-valley rivers, the river inflow to the gsl-bb lake system would have been significantly reduced. figure 4. known lakes in the gsl-bb larger than modern gsl during the past 3 ma. chronologic data are from the burmester core for the bonneville, little valley, pokes point, lava creek, and unnamed lake cycles (oviatt and others, 1999; unpublished information); the age of the cd lake cycle is from kaufman and others (2001). the x axis of the graph marks the approximate elevation of modern gsl (~1280 m), and the vertical scale, which represents the relative maximum elevations of lakes, is not shown on the figure because insufficient information is available for most lake cycles. approximations of the upper elevation limits of the cd and lv lake cycles are based on outcrops of lacustrine sediment. the upper elevation limits of the pp and lc lake cycles are interpreted as being similar to that of the lv lake cycle. the elevation of the unnamed lake cycle at about 3 ma is unknown, but based on the ostracode fauna in sediments of that age from the burmester core, the lake probably did not rise higher than the cd lake cycle. b = bonneville, cd = cutler dam, pp = pokes point, lc = lava creek, u = unnamed lake cycle. drill hole id  collec on year  la tude (°n)  longitude (°w)  eleva on (m)  depth of hole (m)  age at bo om of hole (ma)  core or cu ngs  reference  gsl96‐6  1996  41.0  112.4  1272  9  0.044  core  thompson and ovia , unpublished, 1995‐2022; thompson and others, 2016  gsl96‐4  1996  41.0  112.5  1272  5.5  0.04  core  thompson and ovia , unpublished, 1995‐2022  gsl00‐4  2000  41.1  112.6  1271  120  0.280  core  schnurrenberger and others, 2001; balch and others, 2005  c  ~1980  41.0  112.4  1272  5.5  0.035  core  spencer and others, 1984; thompson and others, 1990  amoco 1  ?  41.5  112.8  ?  ?  ?  cu ngs  moutoux, 1995  amoco 2  ?  41.4  112.8  ?  ?  ?  cu ngs  moutoux, 1995  amoco 3  ?  41.4  112.8  ?  ?  ?  cu ngs  moutoux, 1995  amoco 4  ?  41.4  112.7  ?  ?  ?  cu ngs  moutoux, 1995  amoco 5  ?  41.4  112.7  ?  ?  ?  cu ngs  moutoux, 1995  amoco 6  ?  41.4  112.7  ?  ?  ?  cu ngs  moutoux, 1995  amoco 7  ?  41.4  112.6  ?  ?  ?  cu ngs  moutoux, 1995  amoco 8  ?  41.1  112.7  ?  ?  ?  cu ngs  moutoux, 1995  amoco 9  ?  40.9  112.3  ?  ?  ?  cu ngs  moutoux, 1995  amoco 10  ?  40.8  112.3  ?  ?  ?  cu ngs  moutoux, 1995  south rozel (j)  ?  41.4  112.6  1272  ?  ~5  cu ngs  moutoux, 1995; kowalewska and cohen, 1998; davis, 2002  gunnison (p)  ?  41.3  112.7  1270  ?  ~5  cu ngs  moutoux, 1995; kowalewska and cohen, 1998; davis, 2002  indian cove (i)  ?  41.3  112.6  1271  ?  ~5  cu ngs  moutoux, 1995; kowalewska and cohen, 1998; davis, 2002  bridge  ?  41.2  112.5  ?  ?  ?  cu ngs  moutoux, 1995; davis, 2002  carrington island (h)  ?  41.0  112.5  2171  ?  ~5  cu ngs  moutoux, 1995; kowalewska and cohen, 1998; davis, 2002  sandbar (n)  ?  40.7  112.4  ?  ?  ~2.3  cu ngs  kowalewska and cohen, 1998  s28  1960  40.9  112.2  1286  224  ~0.9  core  shuey, 1971; eardley and gvosdetsky, 1960; williams, 1994; thompson and ovia ,  unpublished, 1995  saltair  1956  40.8  112.1  1282  198  ~0.8  core  shuey, 1971; eardley and gvosdetsky, 1960; williams, 1994; thompson and ovia ,  unpublished, 1995  burmester  1970  40.7  112.5  1285  307  ~3.4  core  shuey, 1971; eardley and others, 1970; williams, 1994; ovia  and others, 1999;  thompson and ovia , unpublished, 1995  knolls  1960  40.7  113.3  1289  152  ~0.9  core  shuey, 1971; williams, 1994; thompson and ovia , unpublished, 1995  wendover  1960  40.7  113.9  1285  171  ~1.7  core  shuey, 1971; williams, 1994; thompson and ovia , unpublished, 1995; bright and  others, 2022  clive  2019  40.7  113.1  1307  187  ?  cu ngs  stantec, unpublished, 2019; ovia , unpublished, 2019  black rock  1993  38.7  112.9  1503  273  ~3  core  thompson and others, 1995  pit of death  1993  39.0  113.2  1383  140  ~3.1*  core  thompson and others, 1995  table 1. drill holes in the great salt lake and sevier basins that contain sediments of pre-lb age. 7 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 *this core contains an unconformity @~140 m, below which is a ~6 ma tephra. 8 c.g. oviatt late neogene and quaternary lacustrine history of the great salt lake-bonneville basin if the climate in the gsl-bb basin were dryer than today during the late neogene and pleistocene (except during deep-lake cycles), it is likely that climatically induced river inflow to the gsl would have been reduced at that time (following the logic of bekker and others, 2014, who studied tree-ring reconstructions of late holocene streamflow in the weber river and the connections with climate). climatically reduced inflow, combined with the lack of inflow from the bear and cache-valley rivers, would have caused lakes in the gsl-bb to be smaller compared to holocene gsl, and that reduced input would likely increase the probability of streams feeding marsh systems in isolated depressions on the basin floor. the information reported by balch and others (2005) suggests that the hydrologic budget of gsl about at 45 ka was different than it is today. the difference in budget could have been that the bear and cache valley rivers were not entering gsl 45 ka, and/or that climate was dryer at that time, during marine oxygen isotope stage (mis) 3. mis 3 was an interglacial period. it is interesting and seemingly paradoxical that a hypersaline condition for the lake system in the gslbb (such as modern gsl) probably requires the inflow volume to be relatively high compared to that required for freshwater marshes to appear on the basin floor. it’s clear that a decrease in water inflow to the lake causes lake level to decline; if inflow were to decrease sufficiently a hypersaline lake would cease to exist. in 2023, the upper bear river and cache valley rivers are contributing water to gsl, and the lake is dropping to alarmingly low levels, partly because of the very warm and dry climate we are now experiencing, but mostly because of water diversions by humans from the inflowing rivers before the water gets to gsl (abbott and others, 2023). if the upper bear river and cache-valley rivers were not presently entering gsl, what would be the condition of the lake in 2023? pliocene to late pleistocene deposition sevier basin cores the sevier basin (figure 2) has been part of the larger gsl-bb for at least the past ~3 ma. two sediment cores from the sevier basin record sedimentation during the period from ~3 ma to a few thousand years younger than the brunhes/matuyama paleomagnetic boundary (thompson and others, 1995), currently dated at 773 ka (channell and others, 2010). these two cores, the black rock and pit of death cores (table 1; figure 2), contain sediments of shallow lakes and muddy (playa) depositional systems. no deposits of deep lakes were encountered in those cores, an observation that is consistent with observations from the gsl basin farther north and reinforces the interpretation that lakes in the gsl-bb were low or did not exist during the period from ~3 ma to 773 ka. the deep-lake cycle at about 3 ka in the gsl-bb probably did not get high enough to flood into the sevier basin; the elevation of the topographic divide between the gsl basin and the sevier basin (orbt, figure 2) was probably on the order of 1400 m. eardley cores during the 1950s and 1960s, armand j. eardley, who was a professor of geology at the university of utah, oversaw the drilling of four deep holes and the acquisition of sediment cores from those drill holes. the cores were called s28, saltair, burmester, knolls, and wendover (figure 2; table 2). eardley and his colleague, vasyl gvosdetsky (university of utah), published a description and interpretation of one of the cores (the saltair core; they also commented on the s28 core; eardley and gvosdetsky, 1960). r.t. shuey, a colleague of eardley’s at the university of utah, obtained funding from the national science foundation (nsf) to study the paleomagnetism of the sediments in the eardley cores and wrote an unpublished report for nsf (shuey, 1971). in 1973, eardley and a group of colleagues, published a description and interpretation of part of the burmester core (eardley and others, 1973). lister (1975) described ostracodes from the saltair and s28 cores. s.k. williams, a ph.d. student of b.p. nash (also at the university of utah and a coauthor on the eardley and others, 1973, paper), studied the volcanic ashes from the cores and published important information about the eardley cores (williams, 1994). in 1995, r.s. thompson (usgs) and c.g. oviatt (kansas state university) examined the five eardley cores looking for evidence of deep-lake cycles based on the presence of carbonate marl deposited in deep lakes and deep-lake ostracode faunas. in 1999 oviatt and colleagues published a brief description and reinterpretation of the upper ~110 m of the burmester core (younger than the brunhes/matuyama geomagnetic boundary; oviatt and others, 1999). as part of that work, j. bright and d.s. kaufman (northern arizona university), and r.m. forester (usgs), studied ostracode faunas and ostracode amino acid racemization in most of the eardley cores, and some of that information was published in oviatt and others (1999). more recently j. bright and colleagues studied amino acid racemization in ostracodes from the wendover core id  plss1  la tude2  longitude2  eleva on (m)  depth (m)  sed. rate (m/ma)3  approx. age at bo om (ma)  year of drilling  recovery  references  s28  sw1/4,  se1/4, sec.  28, t1n, r2w  40.79  112.07  1286  223  230  ~0.9  1960  0% in some secons, up to 40%  in others  shuey (1971); williams (1994)  saltair  se1/4 sec. 25, t1n, r3w  40.79  112.20  1282  198  260  ~0.8  1956  50%  eardley and others (1963);  shuey (1971); williams (1994)  burmester  se1/4, sec. 7, t2s, r5w  40.65  112.45  1286  306  3.4-2.6 ma:  90 m/ma;  2.6-0 ma:  120 m/ma  3.4  1970  90%  shuey (1971); eardley and  others (1973); williams  (1994); ovia  and others  (1999)  knolls  sw1/4, sec.  15, t1s,  r13w  40.72  113.30  1289  152  170  0.9  1960  30%  shuey (1971); williams (1994)  wendover  se1/4, sec.  15, t1s,  r18w  40.74  113.87  1285  171  130  1.7  1960  50% < ~120 m;  15% > ~120 m  shuey (1971); williams (1994)  1plss = public land survey system  2datum for la tude/longitude coordinates is wgs84.  3data from williams (1994); approximate sedimenta on rates  table 2. information about the eardley cores m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 9 10 c.g. oviatt late neogene and quaternary lacustrine history of the great salt lake-bonneville basin core (bright and others, 2022), the only core not studied for that purpose in the 1990s. davis (2002) published pollen diagrams that had been constructed from data from the wendover and knolls cores in the 1960s, but which had not been previously published. the eardley cores are now completely dried out. they have been stored in cardboard boxes and sampled multiple times by different people for different purposes. observations about the geologic history of the core sites, which would have been possible when the cores were fresh, are now difficult. the eardley cores are now archived at the utah geological survey core research center. the usefulness of the eardley cores is limited because some of the core sections have crumbled. although drilling technology has been vastly improved since the 1960s, the cost of drilling and the acquisition of even one new core that might build on what has been learned from the eardley cores, would be huge. however, the scientific information (geologic, biologic, paleoclimatic, etc.) that could be obtained from a new core would be invaluable. the following sections give summaries of published and unpublished information and interpretations concerning the eardley cores (figures 3 and 4; tables 1 and 2). eardley did not publish anything related to two of the cores (knolls and wendover). no independent studies of the sediments or changing depositional environments represented in the knolls and wendover cores have been published. when bob thompson and i examined all the eardley cores in 1995, we found that the core sections had not been split and the surviving sections of the cores were covered with dried mud from the drilling operations. in order to examine the sediments, we had to look at the ends or break apart dried core sections or scrape off the mud from the surfaces. we found this to be true for all the cores, including the saltair and burmester, so it was unclear to us how eardley and his colleagues had observed any of the sediments in the cores. s28 and saltair cores these saltair and s28 cores were taken near each other (figure 2; tables 1 and 2). although some important information about pre-lb lake cycles is preserved in these cores (eardley and gvosdetsky, 1960; recognizing that interpretations of global quaternary history have changed considerably since the 1950s), the amount and quality of information about the lacustrine history of the gsl-bb the cores can provide is not great. both the s28 and saltair cores were drilled at locations dominated by the jordan river and its precursors and were not suitable as complete records of sedimentation in gsl-bb lakes. deposits of the lb cycle are not present in either the s28 or saltair cores, and it would now be difficult to determine whether lb sediments were not preserved at the coring sites or if lb sediments simply were not recovered during the drilling operations. deposits of some older deep-lake cycles are present in the cores and deposits of some deep-lake cycles are missing. although lister (1975) defined some new ostracode species based on samples from the s28 and saltair cores, and his descriptions of ostracodes are excellent and useful, he did not indicate the depths of the samples or say anything about the depositional environments of the samples he examined. burmester core the burmester core is the longest eardley core at 306 m and covers the greatest amount of time (the age at the base of the core is ~3.4 ma; williams, 1994). in our examination of the core in 1995 we found many buried calcic soils, some with enough soil carbonate to whiten the core for many meters. eardley and others (1973; their figure 1) showed 17 deep-lake cycles during brunhes time based on their work on the burmester core, whereas oviatt and others (1999) found evidence in the burmester core for only four deep-lake cycles during the same time period (an age of 750 ka for the brunhes/matuyama geomagnetic boundary was estimated by eardley and others, 1973; in 2023 the age of that geomagnetic boundary is considered to be ~773 ka [channell and others, 2010]). in the upper ~3 m of the burmester core eardley and others (1973; their figure 1) interpreted the sediments as representative of shallow to dry lakes, overprinted by a soil, but oviatt and others (1999; their figure 1) found deposits of lb in that interval, including the hansel valley basaltic ash (miller and others, 2008), which was erupted during the early transgressive phase of lb. recovery was good in the burmester core (90%; table 2) and that core has provided ages for middle and late pleistocene deep-lake cycles in the basin (figure 4). the approximate drilling site of the burmester core is low in the basin, but it is on land, not in the gsl, and no deposits of shallow lakes are preserved in the burmester core. knolls core the lb marl is present in the knolls core. however, in a shallow pit about 4 km west of the approximate location of the knolls core, only about 80 cm — approximately the lower half — of the lb marl (gilbert’s, 1890, white marl) are present, and the upper half has been deflated (oviatt and others, 2020). it 11 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 is unknown how much of the lb marl is present in the knolls core and the section may not be complete. sediments of pre-lb deep-lake cycles are present lower in the knolls core, although it is not known if those deep-lake stratigraphic units are truncated or complete. most of the core is dominated by sediments of shallow lakes (similar to the wendover core, described below). wendover core in the wendover core the lb marl is completely absent, as are deposits of pre-lb deep lakes (unpublished observations by thompson and oviatt, 1995, and by oviatt and d.l. clark, 2019-2022; bright and others, 2022; clark and others, 2023; bernau and others, 2024). drilling recovery was not good (table 2), but no non-lacustrine deposits have been observed; deposits of shallow lakes dominate the core. the wendover core helps demonstrate the importance of deflation in the gsld (bernau, 2022; bernau and others, 2023, this volume), but does not help with determining when deep-lake cycles occurred. the sediments in the wendover and knolls cores reveal important information about the pre-lb history of the gsl-bb. in both cores, the most common sediment types are carbonate mud (grain sizes of clay, silt, some fine sand) and oolitic sands, where most of the oolitic grains are rod shaped. also present are irregularly shaped carbonate lumps and gypsum grains (both primary and secondary precipitates). some carbonate mud units (not the ones dominated by rodshaped ooids) contain the ostracode limnocythere staplini, but no other ostracode species are present. l. staplini lives in brackish water with relatively low alkalinity. in this basin this means the lake was less than a few tens of meters deep — if it rose higher the water would have become diluted and other ostracode species would appear. the rod-shaped ooids probably indicate the presence of brine shrimp (eardley, 1938); spherical ooids probably formed abiotically in the wave-agitation zone of a shallow saline lake (eardley, 1938). these sediments indicate that in pre-lb times, lakes in the gsld were shallow and varied in dissolved-solid content from being salineenough to support brine shrimp at times (too saline for ostracodes), to being brackish and supporting ostracodes at other times (but no brine shrimp). taking into account the poor recovery of the wendover and knolls cores (table 2), the observations suggest that deposition in shallow lakes dominated in the gsld for thousands or millions of years. although dated shorelines of pre-lb lakes in the gsld have not been found (and may not exist), fluctuating lakes with an average elevation of roughly 1300 ± 10 m would be suitable candidates for producing this kind of sedimentary record. a rise of gsl to about 1300 m today would cause widespread flooding and destruction of human infrastructure in the gsl part of the basin, but from a geologic perspective 1300 m is close to the average level of gsl. while “1300 m” is an arbitrarily chosen elevation, it’s within the possible range of elevations of closed-basin lakes that periodically flooded the gsld during pre-b time. this range of elevations is close to the maximum elevation of the latest-pleistocene gilbert-episode lake (~1297 m). the gilbert-episode lake (about 12,000 years ago) formed after lb had evaporated, and was part of gsl. in the gsld the gilbertepisode lake was strongly influenced by fresh, cold water that flowed into the gsld from the sevier basin along the old river bed (palacios-fest and others, 2021; they referred to the gilbert-episode lake as the "old river bed delta lake"), but in the gsl part of the system the same shallow lake was brackish (thompson and others, 2016). similar pre-lb lakes in the gsl-bb with elevations in the range of 1300 ± 10 m should be considered part of ancestral gsl, but it is unknown whether freshwater from the sevier basin entered the gsld in pre-lb times. discussion figure 4 shows an estimate of the ages of known deep-lake cycles in the gsl-bb based primarily on data from the eardley cores (oviatt and others, 1999). figure 5 shows correlations between deep-lake cycles in the gsl-bb and miss (ages summarized by lisiecki and raymo, 2005). deep lakes, other than the ones that have been documented so far, may have risen and fallen during additional even-numbered miss during brunhes time (even-numbered stages were glaciations, odd-numbered stages were interglacials), but further investigations are needed to decipher details. if samples of vein-fill calcite and aragonite from outcrops at lakeside were deposited during deep-lake cycles, they may suggest deep-lake cycles during mis 8 and mis 10 (d. mcgee, mit, personal communication, 2019)( figure 5). it is possible to estimate the proportion of time that deep-lake cycles occupied the gsl-bb during the brunhes geomagnetic chron (730-0 ka). if each of the four largest deep-lake cycles lasted the same length of time as the bonneville cycle, about 17 ka, the total proportion of time that deep-lake cycles occupied the gsl-bb during the past 773 ka was roughly 9%. only one deep-lake cycle is poorly known from the period between 3 ma and 773 ka (based on limited information from the burmester core), and, based on its ostracode fauna, the lake 12 c.g. oviatt late neogene and quaternary lacustrine history of the great salt lake-bonneville basin probably did not rise higher than the cd lake cycle; deep-lake cycles account for less than 1% of that period. therefore, for over 90% of the past ~3 ma lakes in the gsl-bb were shallow. of course, if further evidence is found for deep lakes other than the ones that have so-far been described for the past 3 ma, the percentage of time during which deep lakes occupied the gsl-bb would be greater than 9 %. however, environmental conditions like what we see now (not including human influences) apparently were the rule rather than the exception for at least the past 3 ma, and probably for a longer period (based on the mis record of lisiecki and raymo [2005, their figure 4], which extends back beyond 5 ma). the domination of shallow lakes in the gsl-bb is not surprising considering that the upper bear river and the cache valley rivers did not enter the gsl-bb until just a few tens of thousands of years ago (figure 5). as shown in figure 5, very deep lakes in the gsl -bb were uncommon prior to the middle pleistocene transition (mpt; between about 1.2 ma and 700 ka), which marked a change in the magnitude and frequency of pleistocene glaciations (clark and others, 2006; clark, 2012). after the mpt, global climate varied with high-amplitude 100-ka cyclicity (as seen in mis curves; figure 5), and prior to the mpt, global climate varied with lower amplitude 41-ka cyclicity. after the mpt large northern hemisphere ice sheets began to attain great elevations and had larger volumes than earlier ice sheets (clark, 2012). very thick northern hemisphere ice sheets probably affected global atmospheric circulation patterns and may have been important in the growth of large lakes in the great basin (antevs, 1948), although it’s likely that the influence of ice sheets on global circulation was more complicated than that portrayed by antevs (oster and others, 2015). the cd lake cycle and the post-lb gilbertepisode lake are not represented by deposits in the burmester core (or in any of the eardley cores, except possibly in the knolls core — figure 3), but independently those lakes are known to have covered the burmester core site and all other eardley-core sites. perhaps those lake cycles were quick (fast up, fast down), and little sediment was available at the core sites; or perhaps sediment from those lakes was present immediately after the lake cycles but was not preserved. if sediments of those lake cycles do not exist in the eardley cores, maybe other major lake cycles occurred in the basin but have not yet been detected. balch and others (2005) in a study of core gsl00-4 from gsl, including lake sediments that ranged in age from the present to as old as ~280 ka (table 1), did not report evidence of those short-lived lake cycles. however, the spacing of the samples they figure 5. marine oxygen-isotope (mis) record, which can be interpreted as representing the relative volume of global glacial ice (simplified from lisiecki and raymo, 2005), and known deep-lake cycles in the gsl-bb during the past 3 ma. the red line is a stacked record of δ18o in foraminifera fossils from 57 sites around the world where deep-sea cores have been taken (lisiecki and raymo, 2005; values of δ18o in ocean water were relatively high at times when glacial ice attained large volumes on earth’s surface, and relatively low when ice sheets melted and the water flowed back to ocean basins; values of δ18o are also correlated with water-temperature changes). deep-lake cycles in the gsl-bb are shown in blue with their presumed correlative mis stage numbers—b (bonneville) ≈ mis 2; cd (cutler dam) ≈ mis 4; pp (pokes point) ≈ mis 12; lc (lava creek) ≈ mis 16 (oviatt and others, 1999); the mis stage number possibly correlative with the unnamed lake cycle (“u”) about 3 ma is unknown. three other even-numbered stages are marked on the figure that are likely to have been correlative with deep lakes in the gsl-bb, but deposits of those hypothetical lakes have not been found. a possible age (very approximately 30 ka) of the diversion of the upper bear river and cache valley rivers into the gsl-bb, is plotted. for reference, the green line is plotted at the level on the isotope curve approximately coincident with mis 1 (the holocene); the pale blue line is plotted at the level on the isotope curve approximately coincident with the cd lake cycle (mis 4); the darker blue line is plotted at the level on the mis curve approximately coincident with the lb cycle (mis 2). the approximate duration of the middle pleistocene transition (mpt; clark and others, 2006) is shown. 13 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 examined averaged about 1 m (this represents an average of about 2400 years in that core). even if sedimentation was continuous in some depressions on the floor of gsl (as at the site of gsl00-4), sampling at ~2400 years spacing may not have been close enough to intercept lake cycles that may have lasted only centuries or less. clearly much remains to be learned about pre-lb lakes in the gsl-bb. based on what we know now, it is safe to say that the long-term appearance of the gsl-bb has been close to what we see today, with a shallow saline lake on the floor of the basin. lb was an anomaly, as were other deep-lake cycles in the basin. our historic view of gsl (the past 170+ years) is occurring during a drop in the ocean of geologic time. from a perspective grounded in geologic time, gsl should be viewed as typical rather than as a “remnant” of lb. acknowledgments i am grateful to many people who have helped me with thoughts and principles, data, and various kinds of inspiration. andy roberts (australian national university, canberra) consulted with me about the age of the brunhes/matuyama geomagnetic boundary. genevieve atwood (earth science education, salt lake city), read an earlier draft of this paper and made good suggestions for improvement. don clark (utah geological survey) has been an able and thoughtful colleague in many phases of the work summarized here. i am grateful to the reviewers for helpful comments, and to rick ford and mike vanden berg for encouragement. references abbott, b.w., baxter, b.k., busche, k., de freitas, l., frei, r., gormez, t., karren, m.a., buck, r.l., price, j., frutos, s., sowby, r.b., brahney, j., hopkins, b.g., bekker, m.f., bekker, j.s., rader, r., brown, b., proteau, m., carling, g.t., conner, l., cox, p.a., mcquhae, e., oscarson, d., nelson, d.t., davis, r.j., horns, d., dove, h., bishop, t., johnson, a., nelson, k., bennion, j., and belmmont, p., 2023, 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bonneville-a scientific update: developments in earth surface processes 20, elsevier, p. 28-59. pederson, j.l., rittenour, t.m., jänecke, s.u., and oaks, r.q., jr., 2018, the bear river's diversion and the cutting of oneida narrows at ~55-50 ka and relations to the lake bonneville record-pro16 c.g. oviatt late neogene and quaternary lacustrine history of the great salt lake-bonneville basin ceedings volume: 2018 lake bonneville geologic conference and short course, utah geological survey mp-170, p. 74-89. perkins, m.e., nash, w.p., brown, f.h., and fleck, r.j., 1995, fallout tuffs of trapper creek, idaho a record of miocene explosive volcanism in the snake river plain volcanic province: geological society of america bulletin, v. 107, no. 12, p. 1484-1506. rhode, d., 2016, quaternary vegetation changes in the bonneville basin, in oviatt, c.g., and shroder, j.f., jr., editors, lake bonneville — a scientific update: developments in earth surface processes 20. elsevier. p. 420-441. shuey, r. t., 1971, paleomagnetic chronology and correlation of great salt lake basin sediments: national science foundation, final technical report for grant ga-16134, 15 p. schnurrenberger, d., and haskell, b., editors, 2001, initial reports of the global lakes drilling program. volume 1. glad 1—great salt lake, utah, and bear lake, utah/idaho. covering expedition glad1 of the drilling barge kerry kelts and the glad800 drilling system, 10 august, 2000 to 4 september, 2000: limnological research center, minneapolis, university of minnesota, 315 p. spencer, r.j., baedecker, m.j., eugster, h.p., forester, r.m., goldhaber, m.b., jones, b.f., kelts, k., mckenzie, j., madsen, d.b., rettig, s.l., rubin, m., and bowser, c.j., 1984, great salt lake and precursors, utah-the last 30,000 years: contributions to mineralogy and petrology, v. 86, p. 321 -334. stephens, j.c., 1974, hydrologic reconnaissance of the northern great salt lake desert and summary hydrologic reconnaissance of northwestern utah: utah department of natural resources technical publication 42, 55 p. taylor, d.w., and bright, r.d., 1987, drainage history of the bonneville basin: in cenozoic geology of western utah -sites for precious metals and hydrocarbon accumulation-utah geological association publication 16, p. 239-256. thompson, r.s., toolin, l.j., forester, r.m., and spencer, r.j., 1990, accelerator-mass spectrometer (ams) radiocarbon dating of pleistocene lake sediments in the great basin: palaeogeography, palaeoclimatology, palaeoecology, v. 78, p. 301313. thompson, r.s., oviatt, c.g., honke, j.s., mcgeehin, j.p., 2016. late quaternary changes in lakes, vegetation, and climate in the bonneville basin reconstructed from sediment cores from great salt lake. in oviatt, c.g., and shroder, j.f., jr., editors, lake bonneville-a scientific update. developments in earth surface processes 20: elsevier, p. 221-291. thompson, r.s., oviatt, c.g., roberts, a.p., buchner, j., kelsey, r., bracht, c., forester, r.m., and bradbury, j.p., 1995, stratigraphy, sedimentology, paleontology, and paleomagnetism of pliocene-early pleistocene lacustrine deposits in two cores from western utah: u.s. geological survey open-file report 95-1. utah geological survey, 2023, utah geologic hazards portal: https://geology.utah.gov/apps/hazards/ walker, j.d., geissman, j.w., bowring, s.a., and babcock, l.e., compilers, 2018, geologic time scale v. 5.0: geological society of america. wambeam, t.j., 2001, modeling lake bonneville basin morphometry using digital elevation models: salt lake city, university of utah, m.s. thesis, 60 p. williams, j.s., 1958, geologic atlas of utah, cache county: utah geological and mineral survey bulletin 64, 104 p. williams, s.k., 1994, late cenozoic tephrochronology of deep sediment cores from the bonneville basin, northwest utah: geological society of america bulletin, v. 105, p. 1517-1530. wguep [working group on utah earthquake probabilities], 2016, earthquake probabilities for the wasatch front region in utah, idaho, and wyoming: utah geological survey miscellaneous publication 16-3. paradisradiocarbonchronology.pub 1 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2023 utah geological association publication 51 abstract ooids (calcium carbonate coated grains) are common in carbonate environments throughout geologic time, but the mechanism by which they form remains unclear. in particular, the rate of ooid growth remains elusive in all but a few modern marine environments. in order to investigate the rate of ooid growth in a nonmarine setting, we used 14c to date ooids from great salt lake, utah, a well-known site of aragonitic ooids. bulk ooids obtained from the northern shore of antelope island and the northeast shore of great salt lake near spiral jetty were sieved into different size fractions and produced mean ages ranging between 2728±15 and 4373±20 14c yr bp. larger ooids were older than smaller ooids, implying that larger ooids grew in the environment for a longer duration, with the caveat that bulk age dating integrates the growth history of an ooid. to better resolve growth history, ooids from the coarse fraction were sequentially dissolved, and 14c ages were obtained for each dissolution step to create a time series of ooid growth. the results of the sequential dating indicate that the coarse great salt lake ooid growth began between 5800-6600 ± 60 14c yr bp while their outer cortices are nearly modern. sequentially dated ooids from the south arm of great salt lake at antelope island record a nearly linear growth history (~ 10-15 µm/kyr), whereas ooids from spiral jetty record somewhat faster growth between ~6000 and 4000 years ago (0.03 – 0.06 µm/yr) followed by a 10x slower growth history for the remainder of their lifespan (0.003 – 0.008 µm/yr). the lifespan of great salt lake aragonitic ooids is two to six times longer than those from modern marine environments, and thus provides a unique end member for understanding the mechanisms behind ooid formation. the ooid age range indicates that geochemical parameters measured from bulk ooid dissolution integrates over ~6000 years and thus does not represent a geochemical snapshot in time, as some previous studies have suggested. radiocarbon chronology/growth rates of ooids from great salt lake, utah olivia p. paradis1, frank a. corsetti1, audra bardsley2, douglas e. hammond1, william berelson1, xiaomei xu3, jennifer walker3, and aaron celestian4 1department of earth sciences, university of southern california, los angeles, calirforia; opiazza@usc.edu 2environmental studies program, university of southern california, los angeles, california 3department of earth system science, university of california, irvine, california 4department of mineral sciences, natural history museum of los angeles county, los angeles, california introduction ooids are small (generally <2 mm) laminated, coated grains, with a calcium carbonate cortex surrounding a nucleus. ooids are ubiquitous in the geologic record in marine and lacustrine settings, and as accretionary structures, may serve as repositories of high resolution aqueous evolution, preserving both biogeochemical (diaz and others, 2015, 2013; summons and others, 2013) and isotopic (duguid and others, 2010) information. despite their ubiquity, ooid formation remains enigmatic. both abiogenic and biogenic modes of formation have been proposed (diaz and others, 2013, 2015, 2017; o’reilly and others, 2017; pacton and others, 2012; summons and others, 2013), and the rate of ooid accretion remains elusive for the majority of ooid occurrences. without a better understanding of how rapidly ooids form, their utility as paleoenvironmental indicators is hindered and the question of biogenicity remains unclear. radiocarbon (14c; half-life = 5730 ± 40 yr) has been successfully used for the step-wise dating of marine ooids from the bahamas (beaupré and others, 2015; duguid and others, 2010), australia (beaupré and others, 2015; james and others, 2004), and hawaii (hearty and others, 2010). regardless of location, 14c ages decrease from the ooid nuclei toward their outer surfaces with the exception of a 14c anomaly of unknown origin in ooids from highborne cay, bahamas (beaupré and others, 2015). using the radiocarbon chronology, beaupré and others (2015) argued “modern” marine ooid net growth rates were slow and relatively constant, with mean lifespans ranging from 800 ± 135 to 1470 ± 280 14c years and growth rates ranging from 0.36 ± 0.03 to 2.2 ± 0.3 ng c-caco3/ooid-year. however, calculated net growth rates from these radiocarbon dating experiments on ooids are likely underestimating gross carbonate precipitation due to abrasion, as lab experiments have shown growth can be four orders of magnitude faster than radiocarbon net growth rates (trower and others, 2017). the great salt lake (gsl) in utah provides a 10.31711/ugap.v51i.137 2 o.p. paradis, f.a. corsetti, a. bardsley, d.e. hammond, w. berelson, x. xu, j. walker, a. celestian radiocarbon chronology rates of ooids unique opportunity to assess the net growth rate of radial aragonitic ooids that texturally resemble many ancient ooids, both marine and lacustrine (figure 1). in addition to their utility in understanding radial ooid formation, gsl ooids may be targets for understanding the history of gsl, which as a terminal lake with no outflow, is particularly sensitive to climatic shifts. the gsl has also been subjected to environmental alteration by human activity. especially notable is the partitioning of the lake by a railroad causeway constructed in 1959, which created a northern and southern salinity contrast. however, like marine ooids, the utility of lacustrine ooids in reconstructing paleoenvironmental changes is dependent on their placement within a proper temporal framework. the aim of this study is to use 14c as a chronometer to sequentially date ooids from great salt lake, and thus constrain modern ooid formation in this setting and provide necessary chronological context so that their potential as paleoenvironmental indicators may be explored. great salt lake great salt lake environmental setting great salt lake (gsl) is a terminal lake in northern utah with circumneutral ph. gsl represents the present phase (since 11.5 ka bp) that resulted from figure 1. examples of ancient and modern ooid microfabrics. a) neoproterozoic tangential ooids from the beck springs formation. b) radial ooids from the neoproterozoic johnnie formation. c) modern ooids from joulter’s cay, bahamas display tangential concentric laminae that are characteristic of many modern marine ooids. d) modern ooids from great salt lake, utah have a primary radial crystal orientation. 3 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 the transition of the larger and deeper lake bonneville (30-11.5 ka bp) to a shallow, hypersaline lake (oviatt and others, 1992; oviatt and others, 2015). the holocene shallow lake interval of great salt lake is accompanied by a shift to aragonite precipitation, which is in agreement with the mg2+/ca2+ ratios of the lake water (spencer 1985). the north arm of gsl is currently separated from the south arm by a rock-fill railroad causeway that was constructed in 1959. because the three rivers that feed the lake (bear, jordan, and weber rivers) enter the south arm, the north arm water is more saline (28%, at or above saturation for halite) than the south arm, which has a salinity of approximately 15% (rupke and macdonald, 2012; stephens, 1990; usgs, 2023). the causeway was breached in 2016 to restore the flow between the north and south arms, but the berm in the bottom of the breach was raised by 4 feet in july 2022 help reduce north-to-south water flow but still allow water to flow from south to north. (utah dnr, 2022). great salt lake ooids ooids are found as shoreline deposits around the entirety of gsl (baskin, 2014; eardley, 1938). eardley (1938) described the predominantly radial texture of the gsl ooid cortices, inferred ooid cortices were calcitic, and suggested that their radial texture was the result of recrystallization. the assumption of calcitic mineralogy in gsl ooids prevailed until kahle (1974) demonstrated that gsl ooids are in fact aragonite and their cortical fabric is depositional. however, kahle (1974) concluded aragonite-aragonite inversion had taken place. sandberg (1975) confirmed the aragonitic mineralogy of gsl ooids, demonstrated that the radial aragonite fabric is depositional, and found no evidence that aragonite-aragonite inversion had taken place. subsequently, reitner (1999) suggested that organic matrices on the surface of gsl ooids could be important in the mineralization of the aragonite, and lincoln et al. (2022) implicated sulfate reducing bacteria in the precipitation of mg-silicates associated with some gsl ooids, and hypothesized some of the aragonite could be secondary vs. primary. trower and others (2020) developed an approach to understand the unique cortical history of great salt lake ooids, noting that the grains within the same deposit likely record similar histories, but found differences between populations of ooids across various localities in gsl. with respect to the age of the gsl ooids, (mcguire, 2014) attempted serial dissolution of unsorted ooids from 15 cm water depth in the modern south arm of gsl that resulted in 14c ages from 2024±36 yr bp (outermost composite sample) to 8144±29 yr bp (innermost composite sample), indicating that the ooids were quite old relative to modern marine examples, but the coarse sampling resolution could not discern whether modern precipitation took place. as part of a large-scale survey of the tufa-like carbonate mounds that many refer to as “microbialites”, bouton and others (2016) measured the bulk 14c age of unsorted gsl ooids from the shoreline of the south arm of the lake. their results yielded a composite ooid 14c age of 3300 yr bp. thus, while some constraints regarding the age of the ooids exist, many questions remain. methods sample collection ooids were collected at the sediment-water interface in less than 10 cm water depth from bridger bay on antelope island and near spiral jetty in march 2014 (figure 2). samples were rinsed with deionized water, dried in an oven at 50°c, and sieved to partition the ooids into discrete size fractions (125-250 µm, 250-355 µm, 355-500 µm). ooid masses from each size fraction were normalized by mass to establish a grain size distribution (figure 3). unfiltered lake water was sampled from the shore of the northern tip of antelope island in the south arm of gsl and the beach at spiral jetty in the north arm of gsl in september 2016 for dissolved inorganic carbon 14c analysis. unfiltered river and well water were sampled in may 2017 from bear, jordan, and weber rivers as well as a well in ogden, utah (weber state university). at each site, one liter of water was collected in 1000ml size glass bottles (fisher #06-414-8) which had been previously rinsed three times with deionized water, soaked in 10% hcl, and rinsed three more times with deionized water. the bottles were field rinsed three times before water was sampled with no head space and immediately poisoned with 100 µl of saturated hgcl2 in the field to preclude later biologic activity. raman spectroscopy raman spectra of gsl ooids were obtained using a horiba xplora+ micro-raman spectrometer. specimens were measured using an incident wavelength of 532 nm, laser slits of 200 µm, 1800 gr/mm diffraction grating, a 100x (0.9 na) objective. laser spot size was approximately 2 micrometers in diameter, and the laser power measured at the sample was approximately 87 (+/3) µw. data were collected on individual grain mount ooids that were polished and thinsectioned. hyperspectral mapping was collected with 4 o.p. paradis, f.a. corsetti, a. bardsley, d.e. hammond, w. berelson, x. xu, j. walker, a. celestian radiocarbon chronology rates of ooids figure 2. map of great salt lake, modified from currey and others, 1984. ooid samples were collected from the sediment water interface from: spiral jetty in the north arm of gsl, and bridger bay on antelope island in the south arm of gsl. scale bar equals 8km. figure 3. presence of organic matter (blue) within an ooid from the north arm (a) and south arm (b) of great salt lake acquired from raman spectroscopy. only the central grain was scanned in each image. a survey of 30 ooids was carried out to confirm the distribution of organic matter within ooid cortices (si table 2). scale bars equal 100 µm. organic matter is distributed within peloidal nucleus (a) and throughout carbonate cortex (a and b). 5 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 0.1-sec exposure averaged over three acquisitions using an 8 µm x 8 µm mapping grid. these spectral acquisition parameters were determined by trial and error to maximize signal-to-noise and keep acquisition time to less than 24 hours for most maps. after collecting hyperspectral maps, principal component analysis was performed to find those unique spectra representing the total variability within each ooid. the total number of components found for all ooids was aragonite, organic material, burned organic material (burned by the laser), pigment (carotenoid), quartz, k -feldspar, and epoxy (si figure 1). microplastics were found but were exceedingly rare within the ooids. once components were identified for each ooid, heat maps were generated by least squares fitting to every spectrum in the map (in some cases, these were >80,000 spectra per map). the least squares fitting does provide an approximate percentage of each component in the spectra; however, these values often have very high errors. therefore, the heat maps were treated as the presence or absence of each component and were not used for absolute abundances. massive organic material has a characteristic raman signal and strong luminescence with the 532 nm laser. these patterns were compared to known organic materials from pigmented crustacea after digestion by red-ear slider turtles (clause 2021). the result is intense luminescence from the organic-rich parts of the sample, and therefore no individual organic molecule could be identified except for the carotenoids. the carotenoids were all found in spectra with organic luminescence. attempts to collect data with a 785 nm laser resulted in very poor signal-to-noise ratio, and it was determined that large maps could not be collected in a reasonable time frame (< 24 hrs), even though the background luminescence intensity was lower. burned organic material has characteristic d and g bands common for soot, char, and organic materials with high thermal maturity. dissolved inorganic carbon 14c age lake, river, and well water samples were prepared using the headspace-extraction method (gao and others, 2014). all radiocarbon results have been corrected for isotopic fractionation according to the conventions of stuiver and polach (1977), with δ13c values measured on prepared graphite using the ams spectrometer. these may differ from δ13c of the original material, if fractionation occurred during sample graphitization of the ams measurement, and thus, the d13c values reported herein were measured on water dissolved inorganic carbon (dic) directly using gas bench coupled with irms (finnigan delta plus). bulk inorganic and organic carbon ooid 14c ages total organic and inorganic c was extracted from each sieved ooid sample (125-250 µm, 250-355 µm, 355-500 µm) and an unsorted ooid sample. the extracted organic and inorganic carbon was analyzed for 14c at the keck carbon cycle accelerator mass spectrometer (kccams) facility at the university of california, irvine (beverly and others, 2010; southon and others, 2004). details regarding methodology for the bulk organic and inorganic carbon extractions may be found in the supplementary information (si table 1). sequential ooid acidification to assemble an ooid chronology, we measured the 14c ages of fractions of co2 collected during sequential acid addition to sieved ooids (355-500 µm) from spiral jetty and antelope island. ooids (~50g) and 150ml of deionized water (diw) were placed in a reaction vessel constructed from a 500ml graduated round media storage bottle (vwr cat. # 89000-238) and a suspended magnetic stir rod (si, figure 2). the reaction vessel was purged with n2 that was scrubbed with ascarite-ii while a stir bar spun at 700 rpm to drive off any dissolved co2 in the water for a total of 30 minutes. the sample was acidified by injecting 60 ml of 3.3m hcl at a flow rate (acid) of 10ml/min. gas was shunted for the first 5 seconds of acidification to off-gas any residual n2 before collecting the sample in tedlar bags which had been rinsed with ultra-high purity (uhp) helium scrubbed with ascariteii. gas was collected in 3 tedlar bags per each acidiwater source la tude longitude δ13c(‰) ± frac on modern ± δ14c(‰) ± 14c age (yr bp) ± dic (mm) well  41.192175  ‐111.93894  ‐15.5  0.1  0.8377  0.0014  ‐169.1  1.4  1425  15  6.6  weber river   41.218295  ‐111.987708  ‐10.2  0.1  0.9332  0.0015  ‐74.3  1.5  555  15  3.2  bear river  41.545895  ‐112.095349  ‐8.4  0.1  0.8348  0.0016  ‐171.9  1.6  1450  20  4.1   jordan river  40.771568  ‐111.975878  ‐9.7  0.1  0.8416  0.0014  ‐165.2  1.4  1385  15  4.1  table 1. 14c and δ13c composition of weber, bear, and jordan rivers as well as well water sampled in ogden, utah. dissolved inorganic carbon (dic) is reported in millimolar (mm). water was treated with hgcl2. 6 o.p. paradis, f.a. corsetti, a. bardsley, d.e. hammond, w. berelson, x. xu, j. walker, a. celestian radiocarbon chronology rates of ooids fication step. the first tedlar bag collected gas for the first 30 ml of acid added, the second tedlar bag collected gas during the second 30 ml acid addition, while the third tedlar bag collected remaining co2 that evolved after all 60 ml acid had been added and was left to sit for 3 minutes before pulling it off the vessel. four discrete acidification steps were performed, with a subsample of 5-10 ooids removed from the acidification vessel between each. the subsample of ooids was examined using a hitachi tm1000 environmental scanning electron microscope (sem) to confirm dissolution was occurring from the outside to the inside (si, figure 3). between acidifications, ooids were rinsed three times with deionized water (diw) and dried overnight. the reaction vessel and its components were rinsed in 10% hcl and dried between each acidification. the diw in the reaction vessel was replaced, and the reaction vessel was purged for 30 minutes with ascarite-scrubbed n2 to remove any atmospheric carbon. the acidification procedure, using 60 ml of 3.3 m hcl, diw rinse, acid wash, and 30-minute purge, was repeated for each acidification (four times total). following the final acidification, the remaining nuclei were rinsed three times with diw, dried overnight, and reserved for 14c analysis of the organic carbon fraction. some calcium carbonate remained on the oolitic nuclei at the end of the experiment to ensure ancient carbonate nuclei were not dissolved which might skew the oldest inorganic carbon age. 14c analysis for 14c analysis, gas samples from the sequential leach were cryogenically purified through a dry ice/ ethanol trap and collected in a liquid nitrogen trap. residuals from ooid dissolution of bulk ooids and from the sequential leach were combusted at 900°c for 3 hours to obtain co2. all purified co2 samples were graphitized using a sealed-tube zinc reduction method (xu and others, 2007). graphite was pressed into aluminum target holders and analyzed for 14c at the keck carbon cycle accelerator mass spectrometer (kccams) facility at the university of california, irvine (beverly and others, 2010; southon and others, 2004). data were normalized to oxalic acid standard ox1 and background corrected using radiocarbon-dead reference carbonates acidified in the same reaction vessel. a modern cstd standard (an in-house coral standard from ellen druffel, with a fraction modern value of 0.9445 ± 0.0018 (1σ stdev, n=262) was also processed by the acidification and measured for quality control (gao and others, 2014). 14c data are presented according to the conventions presented in stuiver and polach, 1977. non-calibrated ages (given in year before present, or ybp) are presented in this paper to be consistent with ages reported in the literature. results 14c analyses in the great salt lake the radiocarbon ages of the surface water that enters great salt lake (bear, jordan, and weber rivers) and water from a groundwater well in ogden, utah were measured in may 2017. the 14c ages of the weber river, bear river, and jordan river are 555±15 yr bp, 1450±20 yr bp, and 1385±15 yr bp, respectively (table 1). water sampled from a groundwater well on weber state university campus in ogden, ut produced a 14c age of 1425±15 yr bp. bulk ooid 14c results bulk unsorted and sieved ooids from each site yield inorganic and organic 14c ages that represent the average of a mix of older and younger carbon in the samples, and thus do not represent a unique age for the ooids. however, bulk ages can help bracket the general age of the ooids and provide some indication of their antiquity. in general, bulk ooid carbonate analyses produced ages that ranged from 2728±15 (spiral jetty) to 4373±20 yr bp (spiral jetty), whereas bulk organics produced slightly younger ages, between 1935±15 (antelope island) and 4200±15 yr bp (spiral jetty) (table 2). smaller ooids have younger average 14c ages, which is reflected in both inorganic and organic carbon. total organic carbon of bulk ooids from both sites varies from 0.43% to 1.34%; however, ooids from the north arm of gsl have more than double the organic carbon of south arm ooids (table 2). raman spectroscopy of ooid cross sections locality grain size (μm) inorganic c age (14c yr bp) organic c age (14c yr bp) % total organic carbon spiral je y ‐ north arm    unsorted  3872±15  3490±15  1.34 355‐500  4373±20  4200±15  1.3   250‐355  3759±15  3520±20  0.97   125‐250  2728±15  2335±15  1.19    unsorted  3556±15  2175±20  0.46 355‐500  3947±15  2680±20  0.43   250‐355  3834±15  2250±15  0.48   125‐250  3158±15  1935±15  0.55  antelope isl ‐‐ south arm   table 2. inorganic and organic 14c ages from bulk ooids. 7 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 shows that organic matter is distributed both in the micritic nuclei of ooids (when the nuclei are peloids) and incorporated throughout the carbonate cortices (figure 3, si figure 1). grain size analysis reveals both north and south arm ooids are skewed toward finer grain sizes which contain more organic c in the nuclei proportionally, between 63 and 355 µm. north arm ooids are less dominated by the 63 – 355 µm size class (45%), than southern arm ooids (73%) (figure 4). serial 14c ooid record 14c ages of co2 that was released during acidification of 355 – 500 µm diameter ooids from spiral jetty increased in a non-linear manner from an age of 660±15 yr bp in the first layer dissolved to 5830±60 yr bp in the last layer dissolved. the 14c ages of ooids from antelope island increased linearly from 460±20 yr to 6600±60 yr bp (si table 1) (uncorrected for reservoir effect). 14c ages of organic matter combusted from the nuclei remaining at the end of the experiment were 5975±15 and 6210±20 yr bp for spiral jetty and antelope island ooids, respectively. as dissolution progresses and ooids become smaller, each successive sample taken represents a thicker width of ooid cortex dissolved (assuming the mass dissolved from each acid addition is constant), thus homogenizing the 14c over a larger range of radial cortex depths. to account for this, ages were integrated over ooid cortices ranging from 355 – 500 µm in diameter assuming they had a spherical geometry and constant net growth (precipitation abrasion) (figure 5a, 5b) using v = 4/3 * π r3, where v is volume and r is radius in micrometers. to summarize, our experiments indicate that 355-500 µm ooids from gsl began precipitating around 5830-6600 ± 60 14c years bp with a continuous chronology to near modern ages (when corrected for reservoir effect). organic carbon extracted from the nuclei material left at the figure 4. grain size distribution of great salt lake ooids sampled from the north and south arms of the lake. ooids from both arms of the lake are dominated by finer sized ooids (63 – 355 µm), though south arm ooids are more heavily skewed toward fine grain sizes. 8 o.p. paradis, f.a. corsetti, a. bardsley, d.e. hammond, w. berelson, x. xu, j. walker, a. celestian radiocarbon chronology rates of ooids figure 5. inorganic 14c chronologies from north arm ooids (a) and south arm ooids (b) were integrated over ooid cortical ranges of 355-500 µm in diameter (represented by shaded region), assuming spherical geometry and constant net growth. 14c ages from organic carbon (dotted line) from remaining ooid nuclei were 5975±15 and 6210±20 yr for north arm (a) and south arm (b) ooids respectively. 9 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 end of the experiment yield nearly contemporaneous ages with the oldest inorganic carbon samples, lending credence to the presumed onset of ooid formation by our methods. the average ooid age extracted from the sequential dissolution yields mean ages of 3737 yr bp and 3277 yr bp for north and south arm ooids (355 – 500 µm diameter), respectively. discussion it is important to consider the behavior of 14c in the gsl, as the 14c age of lacustrine carbonates may be subject to a “reservoir effect”, whereby lakes can accumulate old dissolved inorganic carbon over time. lakes acquire some of this carbon from inflowing water that travels over ancient limestones that reside in their catchment, causing their dissolved inorganic carbon pool to have an apparent age that would be older than the atmospheric value. any calcium carbonate that precipitates from that lake water would record an apparently older 14c age than coeval atmospheric 14c. our analyses reveal that reservoir effects represent the largest source of uncertainty in our data. surface water enters the lake via three rivers: bear, jordan, and weber, all of which enter the south arm of gsl. our results show that these river waters and water from a well in ogden, utah (representing groundwater) deliver ancient inorganic carbon to the lake. the continuous exchange of co2 between the lake water and the atmosphere reduces the age of the lake water reservoir, and thus the reservoir age at any given time is a reflection of the water balance of the inputs of ancient dic, lake surface area (exchange of co2), dic removal, and the existing reservoir age. two anthropogenic changes may have influenced the lake reservoir age in contrasting ways. the causeway has reduced the north arm surface area by a factor of two, reducing the rate of atmospheric exchange in this region proportionally. however, bomb testing has increased 14c/12c in the atmosphere by an average of 50% during the past 50 years. because those two effects may largely negate one another, we assume the modern south arm reservoir effect of 295± 20 yr bp is likely more representative of pre-causeway homogeneous lake conditions and therefore more applicable to this dataset. there remains uncertainty in how the reservoir age may have varied through the past 6000 years. paired u-th and 14c ages from lacustrine cave carbonates suggest the reservoir effect for lake bonneville (from 25 to 13 ka) was 200 years or less (mcgee and others, 2012), which agrees with previous estimates of lake bonneville’s reservoir effect (oviatt and others, 1992). however, bowen and others (2019) suggest reservoir ages for much of the holocene may exceed 1000 years, and they estimate that the reservoir age decreased substantially from >1200 years to <500 years during the late holocene. bulk inorganic 14c ages from sieved ooids reveal that smaller ooids are younger than larger ooids (table 2). the younger average bulk 14c ages of finer sized ooids may be attributed to a more recent onset in formation and implies that an ooid factory has been active in great salt lake since ~6000 ybp. therefore, ooid size in the gsl would appear to scale with age rather than some later physical sorting mechanism indicating that ooids have been growing in gsl over at least the past several thousand years at these localities. the grain size distribution is skewed toward finer grain sizes in both arms of the lake, which is also suggestive of an active ooid factory when combined with bulk ooid age data (table 2, figure 4). bulk ooid ages also indicate that the great salt lake ooids are significantly older than the modern marine ooids from carbla beach, australia and the bahamas (beaupré and others, 2015; duguid and others, 2010). however, some significant caveats require exploration while interpreting bulk 14c ages. bulk ooid ages do not allow for the differentiation between relic ooids that formed thousands of years ago versus modern ooid formation if the population of ooids in a size class is a mixture of material of different ages. in addition, raman spectroscopy demonstrates that organic carbon is not exclusively found in peloidal nuclei but is also incorporated throughout the ooid cortex. thus, bulk ooid organic carbon ages represent a mixture of organic carbon from artemia pellet nuclei and younger organic carbon incorporated at various points in the growth of the aragonitic cortex. ooids from the northern arm of gsl have older bulk organic carbon ages (table 2) for each size fraction, including unsorted ooids. because the total organic carbon content in north arm ooids is twice that of south arm ooids and the bulk organic carbon ages are older, we expect this age disparity is attributed to a higher occurrence of ooids with organic-rich brine shrimp pellet nuclei in the north arm of the lake. petrographic investigation of 100 ooids in thin section from the northern and southern arm of gsl confirm this hypothesis, with 83% pellet nuclei in the north compared to 56% pellet nuclei in the south (si figure 4). the distribution of organic matter throughout ooid cortices coupled with the need to resolve a chronology from the carbonate fraction, highlight both the problems with interpreting bulk age data from ooids and the need for serial dissolution. our serial dissolution experiments present a chronology from modern lacustrine ooids that demonstrate the ancient onset of ooid formation over ~6,000 years ago. once corrected for reservoir effect, the youngest inorganic carbon ages suggest ooids continued to form, apparently up to the present. we hypothesize 10 o.p. paradis, f.a. corsetti, a. bardsley, d.e. hammond, w. berelson, x. xu, j. walker, a. celestian radiocarbon chronology rates of ooids ooid formation may still be occurring, as any modern 14c would be homogenized with slightly older 14c in our youngest sample. the 14corg from antelope island ooids is slightly younger than the oldest inorganic carbon sample, and this may be attributed to 1) residual organics of younger origin, 2) partial leaching of ancient carbonate material in the center of the ooids, and/or 3) a reservoir effect yielding inorganic carbon which is apparently older by hundreds of years. the age of the onset of ooid growth from the north and south arm of gsl is similar, as indicated by the oldest inorganic and organic carbon ages of ooids, but the growth curve of their chronologies varies. for example, the south arm ooids appear to have a nearconstant growth rate (between ~0.01 – 0.015 µm/yr) within the resolution of the data and assumptions. in contrast, the growth of the north arm ooids appears to have been initially more rapid (~0.03 – 0.06 µm/yr) and then slowed somewhat throughout their growth history (0.003 – 0.008 µm/yr). the differences in slope (figure 5) may be attributed to local sitespecific variations affecting carbonate precipitation or abrasion in each part of the lake, or assumptions made when calculating dissolution depth (i.e., constant net growth rate, spherical geometry). to determine whether the assumption of spherical geometry in age integration is responsible for the difference in the slope of the ooid growth curves, we integrated the ages over assumed ellipsoidal ooid geometries. the resulting slope differences were exacerbated when we assumed 100% ellipsoidal geometry (si figure 5), suggesting that there are likely other effects (environmental, geochemical, or physical) during the ooid growth history causing their differences in slope. it is intriguing that the north arm ooids fit the prediction that ooid growth should be rapid at first and then slow as they reach hydrologic equilibrium and spend more time as bedload versus suspended load (trower and others, 2017) but coevally, the south arm ooids display a linear growth trend. the prevailing wind direction at gsl is from the se (western regional climate center, 2023). because of the locations of the samples (figure 1), the north arm site should receive stronger wave action than the south arm site, resulting in greater abrasion and slower growth as the grain grows in size. the grain size data (fig. 4), seem to support this, as coarse grained ooids may have been selectively concentrated by stronger wave energy at the northern site. comparison to marine ooid chronologies the gsl ooid growth histories raise some unexpected questions with respect to how ooids form in the gsl and thus how ooids grow in general. the lifespan of 355-500 µm radial ooids from great salt lake is between two and six times longer than most modern marine ooids from the bahamas archipelago and australia (beaupré and others, 2015; duguid and others, 2010). the ooids are very old compared to modern marine examples, yet sequential dating reveals they experienced continuous net growth for over 6000 years while existing within the gsl environment. trower and others (2017) note that the balance of precipitation versus abrasion are key components in the formation of ooids. on the one hand, the gsl has a very different chemical environment versus the marine settings. for example, in marine settings where ooids grow, the seasonal water temperature variations are low, whereas the great salt lake experiences comparatively large temperature fluctuations (si figure 6). paradis 2019 showed that the solubility of aragonite decreases as temperature increases and co2 escapes to the atmosphere, thus the favorable window for aragonite precipitation in the gsl may only exist over a short window in the summer when the lake water is significantly warmer, whereas marine settings are likely to be supersaturated with respect to aragonite year-round. additionally, we expect abrasion is less intense in the gsl than in marine systems as gsl is a significantly lower energy environment than marine examples. finally, the much lower ca+2/mg+2 in gsl (0.03, jones and others, 2009) compared to the ocean (0.2) may slow growth rates. it is unclear how the balance between abrasion and precipitation should be reconciled given how slowly net ooid growth appears to be in the gsl system. do they experience rapid growth then significant abrasion on a yearly basis, thus accounting for such a slow net growth rate, or do they simply grow very slowly? growth could also be episodic in response to variations in salinity driven by rainfall variations on decadal (or longer) time scales. petrographic investigation reveals what appear to be relatively delicate aragonite crystals that we speculate would not survive intense abrasion, supporting the premise that that perhaps the gsl ooids simply grow very slowly. furthermore, how might the radial fabric affect or indicate growth rate versus the tangential fabric in modern marine ooids? could the low ca+2/ mg+2 in gsl facilitate growth at the tips of crystals extending into the solution? interestingly, lincoln et al. (2022) hypothesized that the large ray-like aragonite crystals common in gsl ooid cortices may represent a replacement of a precursor mg-silicate (that is, not a primary phase, but one formed later vs. adjacent, subjacent, or superjacent aragonite). we note that the 14c chronologies of all sampled ooids are coherent (inside/older-outside/younger from initiation of growth to termination for the sequentially dated ooids, with the 14c organic dates of the nuclei corroborating the initiation of ooid growth, as well as larg11 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 older-smaller/younger for the bulk dated ooid). while our work cannot comment on the paragenetic sequence of the ooid fabrics, the coherence of 14c ages is unexpected if secondary replacement of aragonite was widespread. future work, including finer scale sequential dating, may help resolve the unanswered questions surrounding the gsl ooids. ooids and the history of the great salt lake during ooid growth, the north and south arms of great salt lake would have been in communication with one another as part of one large body of water (rather than two arms separated by a railway), thus the generally similar chronologies for onset of ooid growth from each arm of the lake ~6000 years ago agrees with the lake’s history. furthermore, 10,000 14c yr bp marks the end of the gilbert episode of great salt lake, where the lake experienced a brief 15m transgression during which the lake had freshened enough to support ostracods and possibly fish (broughton and others, 2000; oviatt and others, 2015). after the gilbert episode, gsl regressed to average historic gsl levels (near 1280 m) and brine shrimp cysts and pellets appeared in lake sediment cores (oviatt and others, 2015). it is thought that great salt lake did not transgress higher than modern lake levels during early parts of the holocene (11.510.2 cal ka bp; 10-9 14c ka bp), but little is known about the remainder of holocene lake level because holocene sediments on the floor of gsl have been largely reworked (oviatt and others, 2015). on one hand, a bulk chemical analysis of ooids would represent an homogenized signal over ~6000 years and provides one outlook for the duration of aquatic history that ooids may represent, with relevance to other systems where ooids are analyzed as paleoenvironmental indicators. on the other hand, sequential dissolution of the ooids preserved in gsl ooids has the potential to resolve some of the finer scale lake level variations in gsl during the last ~6,000 years and potentially longer given that δ13c and δ18o are coupled in this closed-basin system, and δ18o can correlate with lake level (e.g., talbot 1990). conclusions the high-resolution 14c chronology of gsl ooids demonstrates that: 14c is a robust tool for dating ooids in gsl, and gsl ooids have a lifespan between two and six times longer than modern marine ooids. the long ooid lifespan confirms the need to temporally resolve accretionary structures like ooids before interpreting bulk geochemical data. the 14c ages obtained from organics in ooid nuclei corroborate the timeframe of onset of aragonite precipitation. additionally, raman spectroscopy coupled with 14c ages from bulk unsorted and sieved ooids shed light on the importance of sequentially derived chronologies due to the fact that bulk ages underestimate the maximum age of ooids by thousands of 14c years. this study highlights the disparity in net growth rate, lifespan, and seasonality in precipitation between radial ooids from great salt lake and modern marine ooids. ooids from different parts of the lake show differing growth histories, perhaps reflecting localized variations in wave energy due to prevailing wind direction or other local environmental conditions. lastly, the >6000 year chronology captured in gsl ooids highlights the caution needed in utilizing these accretionary sediments in a 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sedimentary grains. journal of geophysical research: earth surface, 125, e2019jf005452. https:// doi.org/10.1029/2019jf005452 xu, x., trumbore, s. e., zheng, s., southon, j. r., mcduffee, k. e., luttgen, m., & liu, j. c. 2007. modifying a sealed tube zinc reduction method for preparation of ams graphite targets: reducing background and attaining high precision. nuclear instruments and methods in physics research section b: beam interactions with materials and atoms, 259(1), 320–329. https:// doi.org/10.1016/j.nimb.2007.01.175 https://doi.org/10.1016/j.yqres.2015.05.001 https://doi.org/10.1016/j.yqres.2015.05.001 https://doi.org/10.1130/g32846.1 https://doi.org/10.1130/g32846.1 http://wrcc.dri.edu/climate/comp_table_show.php?stype=wind_dir_avg http://wrcc.dri.edu/climate/comp_table_show.php?stype=wind_dir_avg http://wrcc.dri.edu/climate/comp_table_show.php?stype=wind_dir_avg http://ugspub.nr.utah.gov/publications/open_file_reports/ofr-596/appa.xls http://ugspub.nr.utah.gov/publications/open_file_reports/ofr-596/appa.xls https://doi.org/10.1111/j.1365-3091.1975.tb00244.x https://doi.org/10.1111/j.1365-3091.1975.tb00244.x https://escholarship.org/uc/item/4522n2x6 https://escholarship.org/uc/item/4522n2x6 https://doi.org/10.1016/0016-7037(85)90168-1 https://doi.org/10.1016/0016-7037(85)90168-1 https://doi.org/10.1007/bf00026946 https://doi.org/10.1021/ac100494m https://doi.org/10.1111/gbi.12047 https://doi.org/10.1111/gbi.12047 https://doi.org/10.1016/0168-9622(90)90009-2 https://doi.org/10.1016/0168-9622(90)90009-2 https://doi.org/10.1016/j.epsl.2017.04.004 https://doi.org/10.1016/j.epsl.2017.04.004 https://doi.org/10.1029/2019jf005452 https://doi.org/10.1029/2019jf005452 https://doi.org/10.1016/j.nimb.2007.01.175 https://doi.org/10.1016/j.nimb.2007.01.175 14 o.p. paradis, f.a. corsetti, a. bardsley, d.e. hammond, w. berelson, x. xu, j. walker, a. celestian radiocarbon chronology rates of ooids supplementary information . methods for bulk organic and inorganic carbon extractions on sieved ooids to obtain bulk inorganic carbon ages on sieved and unsieved ooids, ooids were dissolved in 3 extractions, and the results of those extractions were weighted and reported herein. the first co2 extraction occurred after 0.8 ml h2po4 acid was added at 70°c and left to react for 1-2 hours. the second co2 extraction occurred after an additional 0.8 ml h2po4 was added and left to react for 2 more hours at 70°c. the third co2 extraction occurred after 2 hours at 70°c after the second extraction, but no extra acid was added to the samples. all purified co2 samples were graphitized using a sealed-tube zinc reduction method (xu and others, 2007). graphite was pressed into aluminum target holders and analyzed for 14c at the keck carbon cycle accelerator mass spectrometer (kccams) facility at the university of california, irvine (southon and others, 2004; beverly and others, 2010). data were normalized and background corrected using both modern coral and radiocarbon-dead reference carbonates acidified in the same reaction vessel. resulting fractions of modern (fm) carbon were weighted according to yield to calculate bulk inorganic carbon age, and 14c data are presented according to the conventions presented in stuiver and polach (1977). to measure bulk organic carbon content and radiocarbon age from the acid insoluble fraction, 1m hcl was added to 10g of ooids at 70°c for 24 hours until ph maintained at 1 for 2 hours. during acidification, the solution containing the sample was centrifuged and the solution was decanted, then a new aliquot of 1m hcl was added. this process was repeated until the ph maintained at 1 for 2 hours. the residuals were then rinsed with milli-q until the ph became neutral. figure 1. example of spectra obtained from raman spectroscopy on one ooid demonstrating the materials identified. 15 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 2. ooid dissolution reaction vessel. 50g of size-sieved ooids were added to the reaction vessel with 150cc of deionized water. ascarite-scrubbed n2 flowed through while the stir bar spun at 600 rpm for 30 min. this step ensured no atmospheric carbon remains in the reaction vessel or water. next, the outflow and gas inflow stopcocks were closed. for each dissolution step, 60cc of 3.3m hcl were added to the reaction vessel at a rate of 10cc/min. tedlar bags were filled with the resulting gas every ~3min. the gas was then moved into he-rinsed and evacuated 25cc serum vials that were submitted to uc irvine for radiocarbon analysis. 16 o.p. paradis, f.a. corsetti, a. bardsley, d.e. hammond, w. berelson, x. xu, j. walker, a. celestian radiocarbon chronology rates of ooids figure 3. scanning electron microscope images of individual ooids with close-up inset after 60ml (a – b), 120ml (c – d), and 180ml (e – f) of 3.3m hcl was added to reaction vessel. ooids maintain general shape post-acidification confirming dissolution occurred fairly uniformly from exterior to interior. 17 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 4. nuclei composition in north vs south arm gsl ooids as observed in thin section. north arm ooids have 83% brine shrimp pellet nuclei, 15% detrital grains such as quartz and feldspar, and 2% other ooids as nuclei. southern arm ooids have 56% brine shrimp pellet nuclei and 44% detrital grain nuclei. 18 o.p. paradis, f.a. corsetti, a. bardsley, d.e. hammond, w. berelson, x. xu, j. walker, a. celestian radiocarbon chronology rates of ooids figure 5. comparison of ooid chronologies from the northern arm of gsl (7b, 7d) to the southern arm of gsl (7a, 7c). figures 7a and 7b use the radiocarbon chronology and integrate the ages assuming spherical geometry of ooids (using mean ooid radius of 213 µm). the resulting slope of the north arm ooid growth curve (b) is steeper (more rapid growth) during the first several thousand years of ooid growth and slows down during the last several thousand years, while the southern arm ooid growth curve (a) is more linear. because north arm ooids have a larger occurrence of ooids with peloidal nuclei (roughly cylindrical geometry), the radiocarbon chronology was also integrated over an assumed cylindrical geometry using v = l * π r2, where v is volume, l is length (6 * radius), and r is radius (213 µm) in micrometers. the slope of the north arm ooid growth curve remains highly non-linear even after assuming cylindrical geometry, suggesting there are other factors (environmental or otherwise) accounting for the difference in slope. the error bars represent the relations for ooids at 355 and 500 µm. 19 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 6. monthly average water temperature in north (circle) and south (square) arms of great salt lake as measured by usgs from 2010-2016. (u.s. geological survey, 2016). 20 o.p. paradis, f.a. corsetti, a. bardsley, d.e. hammond, w. berelson, x. xu, j. walker, a. celestian radiocarbon chronology rates of ooids grain size (μm) extrac on yield (mgc, inorg) inorg c frac on fm bulk fm bulk 14c age (yr bp)   unsorted  1  2.29  0.6844  0.7097 ± 0.0012  0.6423 ± 0.0012  3556 ± 15    2  0.97  0.2904  0.4952 ± 0.0010    3  0.08  0.0252  0.5086 ± 0.0021    355‐500  1  2.36  0.6706  0.6676 ± 0.0012  0.6118 ± 0.0011  3947 ± 15    2  0.98  0.2792  0.4977 ± 0.0009    3  0.18  0.0501  0.5004 ± 0.0011    250‐355  1  2.22  0.6584  0.6838 ±0.0012  0.6204 ± 0.0011   3834 ± 15    2  1.04  0.3092  0.5024 ± 0.0009    3  0.11  0.0324  0.4599 ± 0.0018    125‐250  1  2.42  0.7559  0.7222 ± 0.0013  0.6749 ± 0.0012  3158 ± 15    2  0.72  0.2257  0.5254 ± 0.0009    3  0.06  0.0184  0.5659 ± 0.0027    unsorted  1  3.29  0.7682  0.6482 ± 0.0011  0.6175 ± 0.0011  3872 ± 15    2  0.84  0.1965  0.5202 ± 0.0010    3  0.15  0.0354  0.4924 ± 0.0015    355‐500  1  2.99  0.7542  0.6030 ± 0.0011  0.5802 ± 0.0011  4373 ± 20    2  0.82  0.2055  0.5217 ± 0.0010    3  0.16  0.0403  0.5020 ± 0.0012    250‐355  1  2.89  0.8339  0.6491 ± 0.0011  0.6263 ± 0.0011  3759 ± 15    2  0.52  0.1503  0.5133 ± 0.0011    3  0.05  0.0158  0.5010 ± 0.0031    125‐250  1  3.00  0.8602  0.7359 ± 0.0013  0.7120 ± 0.0013   2728 ± 15    2  0.44  0.1253  0.5646 ± 0.0011    3  0.05  0.0145  0.5685 ± 0.0032      table 1. bulk inorganic carbon extractions for sieved and unsorted ooids. the 14c ages from each extraction were pooled to calculate a bulk age for each sample. 21 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 table 2. thin sections of individual ooids (a) analyzed with raman spectroscopy (b) to map the presence of organic matter (blue) within the ooid cortex. a raman spectrum with a strong noise signal suggesting the presence of organic material was selected from a gsl ooid and established as the “organic material reference spectrum”. each point on the thin section was analyzed using raman spectroscopy and compared to the reference spectrum. the similarity of the measured spectra to the reference organic matter spectrum was mapped on thin sections with a blue overlay, with blue indicating presence of organic material. p. thin sec on label site grain size (µm) ooid number 1  ai‐oc  antelope island – s. arm  355 – 500  1  2  ai‐oc  antelope island – s. arm  355 – 500  2  3  ai‐oc  antelope island – s. arm  355 – 500  3  4  ai‐oc  antelope island – s. arm  355 – 500  4  5  ai‐oc  antelope island – s. arm  355 – 500  5  6  ai‐om  antelope island – s. arm  250 – 355  1  7  ai‐om  antelope island – s. arm  250 – 355  2  8  ai‐om  antelope island – s. arm  250 – 355  3  9  ai‐om  antelope island – s. arm  250 – 355  4  10  ai‐om  antelope island – s. arm  250 – 355  5  11  ai‐of  antelope island – s. arm  125 – 250  1  12  ai‐of  antelope island – s. arm  125 – 250  2  13  ai‐of  antelope island – s. arm  125 – 250  3  14  ai‐of  antelope island – s. arm  125 – 250  4  15  ai‐of  antelope island – s. arm  125 – 250  5  16  sj‐oc  spiral je y – n. arm  355 – 500  1  17  sj‐oc  spiral je y – n. arm  355 – 500  2  18  sj‐oc  spiral je y – n. arm  355 – 500  3  19  sj‐oc  spiral je y – n. arm  355 – 500  4  20  sj‐oc  spiral je y – n. arm  355 – 500  5  21  sj‐om  spiral je y – n. arm  250 – 355  1  22  sj‐om  spiral je y – n. arm  250 – 355  2  23  sj‐om  spiral je y – n. arm  250 – 355  3  24  sj‐om  spiral je y – n. arm  250 – 355  4  25  sj‐om  spiral je y – n. arm  250 – 355  5  26  sj‐of  spiral je y – n. arm  125 – 250  1  27  sj‐of  spiral je y – n. arm  125 – 250  2  28  sj‐of  spiral je y – n. arm  125 – 250  3  29  sj‐of  spiral je y – n. arm  125 – 250  4  << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /none /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (srgb 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/hrv (za stvaranje adobe pdf dokumenata najpogodnijih za visokokvalitetni ispis prije tiskanja koristite ove postavke. stvoreni pdf dokumenti mogu se otvoriti acrobat i adobe reader 5.0 i kasnijim verzijama.) /hun 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can be opened with acrobat and adobe reader 5.0 and later.) >> /namespace [ (adobe) (common) (1.0) ] /othernamespaces [ << /asreaderspreads false /cropimagestoframes true /errorcontrol /warnandcontinue /flattenerignorespreadoverrides false /includeguidesgrids false /includenonprinting false /includeslug false /namespace [ (adobe) (indesign) (4.0) ] /omitplacedbitmaps false /omitplacedeps false /omitplacedpdf false /simulateoverprint /legacy >> << /addbleedmarks false /addcolorbars false /addcropmarks false /addpageinfo false /addregmarks false /convertcolors /converttocmyk /destinationprofilename () /destinationprofileselector /documentcmyk /downsample16bitimages true /flattenerpreset << /presetselector /mediumresolution >> /formelements false /generatestructure false /includebookmarks false /includehyperlinks false /includeinteractive false /includelayers false /includeprofiles false /multimediahandling /useobjectsettings /namespace [ (adobe) (creativesuite) (2.0) ] /pdfxoutputintentprofileselector /documentcmyk /preserveediting true /untaggedcmykhandling /leaveuntagged /untaggedrgbhandling /usedocumentprofile /usedocumentbleed false >> ] >> setdistillerparams << /hwresolution [2400 2400] /pagesize [612.000 792.000] >> setpagedevice gsl desert landscape.pub 1 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2023 utah geological association publication 51 trip overview this one-day (~260-mile) field trip guide provides an overview of the late pleistocene to holocene history of the great salt lake desert. stops include knolls sand dunes and areas on or surrounding the bonneville salt flats, such as juke box trench, the bonneville salt flats international speedway, and the saline pan center and edge (figure 1). we cover the post-lake bonneville geomorphic evolution of the great salt lake desert including changes in land cover over the past century. the great salt lake desert area provides unique access to saline landscape features including gypsum dunes and a perennial saline pan. we discuss the origin of these features and how they fit within the area’s broader geologic context. the accessibility of sites discussed here depends on surface conditions. in general, late summer to early fall is the most opportune time to visit this area. vehicular travel to any of the off-road sites is discouraged when there is standing water or high nearsurface moisture (wet mud with little traction). surface conditions can change rapidly, and we recommend researching current conditions before initiating this trip. this desert is hot and dry during the summer and there is no shade and limited access to water; please plan accordingly. past and current great salt lake desert depositional changes provide an analog for the modern great salt lake with changing water availability, potential dust production, competing priorities, and rapidly changing land cover. the information presented here impacts understanding natural and geologic heritage, changing management strategies, and landscape dynamism over multiple spatial and temporal scales. great salt lake desert landscape change over multiple temporal scales—a field trip guide covering the bonneville salt flats and knolls sand dunes   jeremiah a. bernau1,3, brenda b. bowen1, charles g. oviatt2, and donald l. clark3 1department of geology and geophysics, university of utah, salt lake city, utah, jeremiahbernau@gmail.com 2department of geology, kansas state university, manhattan, kansas 3utah geological survey, salt lake city, utah depositional and erosional history this trip features a landscape in the heart of the great basin, the bonneville basin in northwestern utah, which includes classic examples of basin and range topography. while traveling between salt lake city and the great salt lake desert, we will cross several mountain ranges that have excellent fault-block tilting and lake bonneville shoreline exposures. the bonneville salt flats and knolls dune field areas are bounded by grabens with <300 ft of deposition in the past 600,000 to 800,000 years (shuey, 1971). the bonneville salt flats is located within the wendover graben; this area has >1000 ft of laminated carbonate muds and gypsum beds that are underlain by conglomerates (stephens, 1974; bernau and others, 2023b). the wendover and other nearby grabens began forming in the miocene (miller and others, 2021). although faulting and seismic activity are thought to have largely ceased in this area, we discuss evidence for late pleistocene to holocene fault movement and soft sediment deformation features (the cause of these features, whether seismicity, decompression dewatering, or compaction, is unknown). in addition to this field trip guide, several maps provide more insights into the geologic history of this area (cook and others, 1964; doelling, 1964; stifel, 1964; doelling and others, 1994; clark and others, 2020; bernau and others, 2023; clark and others, in progress). late pleistocene to holocene geological record late pleistocene lake bonneville provides the geologic backdrop for this trip. based on radiocarbon dating, lake bonneville was persistent between 30,000 and 13,000 calibrated radiocarbon years before present (cal yr b.p.) (figure 2) (oviatt, 2015). at its peak, it was almost as big as lake michigan and extended over one-third of the state of utah. lake bonneville extended from the wasatch range to the utah-nevada state line area and from soda springs in 10.31711/ugap.v51i.145 2 j.a. bernau, b.b. bowen, c.g. oviatt, and d.l. clark great salt lake desert landscape change over multiple temporal scales figure 1. field trip area overview. yellow stars denote field trip stops and points of interest. other symbols denote sites used to interpret the great salt lake desert depositional record in figure 3. figure modified from clark and others (in progress). basemap imagery from earthstar geographics. figure 2. lake bonneville hydrograph with depositional history from bonneville salt flats cores added. modified from oviatt (2015). 3 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 southern idaho south to parowan, utah (making it nearly twice as long as it was wide) (gilbert, 1890). in addition to its shorelines, lake bonneville left a regular stratigraphic succession of marl across its basin (~1–6 ft thick in most areas). erosion of the great salt lake desert the great salt lake and bonneville salt flats have long been considered saline remnants of lake bonneville (eardley, 1962). analysis of great salt lake desert shallow cores, pits, trenches, and other exposures indicates that story is more complicated (oviatt and others, 2020; bernau, 2022; bernau and others, 2023b; clark and others, in progress). the bonneville salt flats, rather than forming from lake bonneville’s remnant waters, began forming at ~8,000 cal yr b.p., about 5,000 years after lake bonneville’s final retreat to modern great salt lake levels. after lake bonneville desiccated, an estimated 3–6 ft of lake bonneville sediments were deflated (eroded by wind) from the area surrounding the site of the modern bonneville salt flats. this past deflation provides an analog for potential great salt lake sediment deflation that may occur if its water levels continue to decline. the great salt lake desert depositional record along interstate highway 80 (i-80) is summarized in figure 3 (louderback and rhode, 2009; oviatt and others, 2018; oviatt and others, 2020; bernau, 2022). surprisingly, because of the deflation of lake bonneville sediments, the bonneville salt flats—a great salt lake desert depositional area with up to 5 ft of holocene deposition—has a less complete geologic record than the adjoining mudflat, which has had little, if any holocene deposition but has retained all to some of the lake bonneville marl deposits. this depositional difference highlights the potential for topographic lows in arid climates to have less complete geologic records, making this an important consideration when planning and interpreting investigations of paleoenvironmental records. research on modern deflation at owens lake, california, provides a model that explains observed great salt lake desert deflationary patterns (figure 4) (reynolds and others, 2007). deflation in arid settings is strongly influenced by groundwater level and salinity. playas become deflationary surfaces when groundwater levels fall far below the surface (rosen, 1994). this fact alone, however, does not explain observed deflationary patterns. the bonneville salt flats, as a regional topographic low, should have higher groundwater levels than adjoining great salt lake desert basin floor areas, which do not have the same degree of deflation. salinity explains this apparent contradiction. the bonneville salt flats area would have had higher groundwater salinity as water flowed in and evapoconcentrated from surface waters figure 3. chronostratigraphic cross section of shallow great salt lake desert deposits. figure is based on information from great salt lake desert core and pit sites shown in figure 1 (louderback and rhode, 2009; oviatt and others, 2018; oviatt and others, 2020; bernau, 2022) and the information supporting this figure is described in bernau and others, 2023b. 4 j.a. bernau, b.b. bowen, c.g. oviatt, and d.l. clark great salt lake desert landscape change over multiple temporal scales or groundwaters. surficial salts, depending on their thickness and composition, greatly alter a surface’s erodibility. thick salt layers crystallized from standing water are very resistant to erosion. ephemeral crusts created from groundwater evaporation, however, are highly unstable; they easily break up and act as abrasives. similarly, displacive evaporite growth, commonly gypsum, may alter the properties of nearsurface mud. areas with minor salinity (groundwater table so low that the capillary fringe is below the surface, or they have continuous freshwater input) can form relatively stable surfaces that limit deflation. using the model in figure 4, it is likely that before saline pan formation at the bonneville salt flats that this area had an ephemeral crust that accelerated deflation, enabling lake bonneville sediments to be locally removed. as climate shifted at ~8,000 cal yr b.p., gypsum sands began accumulating at the saline pan, limiting deflation. finally, saline pan deposition with bedded halite deposits began to accumulate around 5,500 cal yr b.p. as the climate became cooler and wetter (bernau, 2022). these cooler and wetter conditions enabled more surface water, and potentially, groundwater, to flow into the saline pan and for deeper and longer-lived surface ponding to occur. thicker, bedded halite deposits, like the halite layers seen at the modern surface of the bonneville salt flats were then able to form. field trip road log begin by driving west along i-80 from salt lake city for ~80 miles to exit 41 (knolls). you will pass through several mountain ranges and basins of the eastern basin and range physiographic province. the great salt lake desert, because of its remoteness, hosts hazardous waste facilities and military testing. much of the great salt lake desert hass limited to public access due to military testing and training activities. near aragonite, as you enter the great salt lake desert (near exit 49), smokestacks from a hazardous waste incineration plant to the south become visible. a low-level nuclear and mixed waste landfill is west of this facility near clive (exit 49). a hazardous waste landfill is also located northwest of the clive exit. knolls sand dunes take exit 41, drive to the south, and follow the road as it bends west. after ~1 mile, you will be in the dunes (40.7244° n, 113.2821° w; all coordinates in wgs84 datum). several places on the side of the road provide some distance from the road and are safer than the road for parking. this road can be busy with atvs and utvs, particularly on weekends. site description the knolls dune field and other great salt lake desert gypsum dunes have been investigated since the 1950s (jones, 1953; eardley, 1962; dean, 1978; jewell and nicoll, 2011; boden, 2016; fitzgerald, 2019). these dunes and salt pans are considered excellent analogs for aspects of the martian landscape and may help us better understand mars’ surface evolution and past potential for the existence and preservation of life (benison and karmanocky, 2014). gypsum dunes have low preservation potential. most documented gypsum dunes are less than a few tens of thousands of years old (warren, 2006). gypsum dunes on the eastern side of the great salt lake desert are concentrated along a change in slope (doelling, 1964). dunes consist of predominantly mediumto very fine sand (figure 5) and may figure 4. deflation model for the great salt lake desert, modified from reynolds and others (2007). (a) the saline pan and surrounding area. (b) cross section a a’ shows deflation is highest in the area immediately adjacent to the persistent halite crust (figure from bernau and others, 2023b). 5 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 be >30 ft high (boden, 2016). most grains are gypsum (up to 60%), with oolites being the second most common grain. trace shell fragments (predominantly ostracodes), algal fragments, and quartz sand also occur (jones, 1953). quartz and oolitic dunes occur in other great salt lake desert areas (dean, 1978). in addition to these eastern gypsum dunes, there are several smaller dune areas to the west, approaching the bonneville salt flats (boden, 2016). gypsum dunes are evidence of a drying saline landscape. gypsum forms in arid environments 1) as bottom growth crystals at the bottom of a shallow saline lake, or 2) displacively in subsurface sediments from the evaporation of groundwaters. great salt lake desert gypsum dunes are thought to mostly originate from displacive growths. drying conditions and falling groundwater levels enable gypsum crystals to be transported (figure 5). deflation of finegrained sediments exposes displacive crystals which then are redistributed by wind and accumulate downwind along changes in slope. in the modern great salt lake desert, predominant winds are from the west/ southwest (jewell and nicoll, 2011). dunes are stabilized by vadose zone moisture (wetter conditions), when available, or vegetation, or they stabilize where prevailing winds meet. analyses of cores reveals that lake bonneville sediments were partially to fully deflated at and near knolls (eardley, 1962; oviatt and others, 2020). similarly, lake bonneville marl has been truncated in the area to the west of knolls, indicating deflation was concentrated here (relative to the mudflats to the west) before gypsum deposition. optically stimulated luminescence dating of gypsum crystals (an atypical material for this technique) suggested that gypsum dune formation has been constant since >2,300 yr b.p. and is ongoing (fitzgerald, 2019). gypsum was deposited at the bonneville salt flats between 3,500 and 1,700 cal yr b.p. (bernau, 2022). analyses of aerial photography from the years 1953, 1972, and 2015 indicate that many dunes in the knolls area are still active, with some dunes moving by several miles in that period (fitzgerald, 2019). similarly, our observations of sediment caught by a snow fence running parallel to old highway 40 (south of i-80) indicate great salt lake desert deflation is still actively occurring. sediments stopped by the snow fence consisted of mud, gypsum crystals, and carbonate lumps. transit to juke box trench return to i-80 and continue west for 37 miles to exit 4. the interstate mile markers below note areas of interest along this route. mile marker 25 the unimproved road leading to the northwest from here connects to floating island. floating island’s name stems from the mirage that occurs on hot days, creating an illusion that the small mountain is suspended or floating in the air. note: traversing the floating island dike road is only advisable with a heavy-duty high-clearance vehicle. this elevated road figure 5. schematic model for gypsum dune origin (a to c) and grain size composition of dunes at knolls (d). (a to c) is based on bowler (1986) and (d) is modified from jones (1953). 6 j.a. bernau, b.b. bowen, c.g. oviatt, and d.l. clark great salt lake desert landscape change over multiple temporal scales on an earthen dike was constructed to limit the extent of the west pond, a large lake created in the great salt lake desert by the west desert pumping project in the 1980s (wold and waddell, 1994; kohler, 2002). during this wet period, rising great sale lake levels had the potential to flood infrastructure, commercial facilities, and homes. to address this, large pumps ~11 miles west of lakeside, utah, were constructed and used to pump great sale lake water into the great salt lake desert, specifically the newfoundland evaporation basin (the area directly to the north and east of mile marker 25) (figure 1). when pumped waters evaporated, they created a saline pan that exceeds the modern bonneville salt flats in extent. over time the saline pan has decreased in size (radwin and bowen, 2021). due to the absence of a saline pan in the newfoundland basin before the pumping project, it is unlikely that the hydrological conditions in the area would naturally support the long-term persistence of a saline pan. mile marker 15 analyses of aerial imagery and past reports of the bonneville salt flats’ extent indicate that the bonneville salt flats' surface halite once extended to this location (nolan, 1927). the past thickness of halite between mile markers 15 and 13, however, was thin (<1 inch). mile marker 13.5 the ditch (which has adjoining tailings berms) stretching to the north is used to collect briny groundwater for potash production. the ditch extends under the highway and to the south where it connects to large evaporation ponds used to concentrate brine at the potash mine. these evaporation ponds are a local source of gypsum sand in the great salt lake desert. there is another set of berms between this and the next stop; they are from a brine collection ditch that has been inactive since the 1960s (the salduro loop). note: the ditches are located on private property and a fence along the road limits the pull-over area, do not stop here or enter the ditches. mile marker 10 many people know the bonneville salt flats from the i-80 rest stop. this stop has public restrooms and is the best place to explore the bonneville salt flats’ surface morphology for salt polygons. more discussion of salt polygons is available under the bonneville salt flats surface morphology section of this guide. exit 4 take exit 4. near exit 4, there is a gas station with public restrooms, the remaining stops do not have any facilities. this stop is also an option to refuel before continuing or returning to salt lake city. continue north on leppy pass road for ~0.2 miles and take a left onto the paved i-80 frontage road. continue on the i-80 frontage road for 1.4 miles until the road reaches a t intersection. take a right. continue for 0.8 miles towards the alluvial fan until there is another t in the road, take a left and continue for ~200 ft, and park. walk to juke box trench (40.7549° n, 114.0102° w) (elevation ~4255 ft) ~150 ft southeast of here. please be aware that the roads beyond the i-80 frontage road are not regularly maintained. in the event of recent precipitation or insufficient evaporation to dry the surface, these roads can become impassable. exercise caution and consider the weather conditions before venturing onto these roads. under sustained dry conditions, all sites described in the rest of this guide are accessible in 2-wheel-drive vehicles with standard clearance. accessibility is markedly reduced under wet conditions. proceed with caution. juke box trench juke box trench is located at the site of a past spring. because of the archeological significance of this area (see discussion of juke box cave and danger cave), a trench was excavated and investigated by archeologist david madsen and colleagues in the 1980s. it was enlarged in 2009 and revisited for paleoenvironmental interpretation (oviatt and others, 2018). this stop has an excellent example of prebonneville, lake bonneville, and post-bonneville deposits. the depositional section is (1) base: prebonneville oolitic sand and carbonate-cemented gravel and sand; (2) lake bonneville offshore finegrained sediments (marl); (3) an unconformity that cuts the bonneville section; (4) a gravel lens at the base of the post-bonneville sequence (possibly deposited during the gilbert episode); and (5) holocene wetland deposits. depositional record three main strata in lake bonneville marl correspond to different stages in the lake’s levels. during the lake’s rising (transgressive) stage, it left laminated marls (figure 6). the laminae are interpreted as evi7 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 dence of little bioturbation and more-rapid deposition. as lake bonneville approached its maximum size and formed the bonneville level shoreline, a more massive (layerless) marl was deposited. this layer is interpreted as originating from slower depositional rates and more bioturbation. the top of this interval is denoted by a sharp change in lithology associated with the bonneville flood. as the lake decreased in size, the color of the marl changed, reflecting changes in mineralogy associated with the evapoconcentration of lake waters (provo level and post-provo). across these stages, ostracode species also change, creating a regular sequence that can be used to aid stratigraphic interpretation (figure 2) (oviatt, 2015, 2017). based on radiocarbon dating, deflation (wind erosion) of lake bonneville sediments occurred between the terminal desiccation of lake bonneville (~13,000 cal yr b.p.) and the gilbert episode (~11,600 cal yr b.p.) (oviatt and others, 2018). radiocarbon dating and tephra indicate sustained wetland deposition since ~10,500 cal yr b.p. a significant stratigraphic marker in these deposits is the mazama ash (7,600 yr b.p.), which was deposited during the last significant eruption of the volcano at crater lake national park in oregon. juke box cave and danger cave (optional stop) this optional stop is located to the west of juke box trench. juke box cave and nearby danger cave are important archeological sites. furthermore, the vista from juke box cave’s entrance provides an excellent overview of the area. continue along the road from where you parked for ~150 ft, take the road to the right and head up the alluvial fan until you reach a large turnaround area, figure 6. overview of juke box trench. (a) 1946 aerial imagery overview of the area. danger and juke box caves, former military base, and past extent of spring noted. (b) map overview of juke box trench. (c) juke box trench sediments. note lateral change in deposition locally with little to no holocene deposition and lake bonneville (provo marl, massive marl, and laminated marl) sediments being preserved to the northwest (coincident with a notable change in surface morphology), and deflated area covered by holocene wetland sediments. this wetland was active as recently as 1946. (b and c) modified from oviatt and others (2018). 8 j.a. bernau, b.b. bowen, c.g. oviatt, and d.l. clark great salt lake desert landscape change over multiple temporal scales and park there. from here the road extends up to juke box cave (40.7570° n, 114.0130° w) (elevation ~4446 ft). this is a short, but steep, hike. to reach danger cave (40.7490° n, 114.0182° w) continue ~0.6 mi southwest along the dirt road from where you parked. a short path leads to the gated cave entrance. archeological significance juke box cave is believed to have acquired its name because during world war ii soldiers from the nearby military barracks (figure 6a) used the cave to socialize and even went so far as to construct concrete dance floor in its confines. danger cave holds the distinction of being utah's first state monument. it also holds a place on the national register of historic places and is recognized as a national historic landmark. major excavations were conducted at juke box and danger caves in the 1940s and 1950s by university of utah researchers (jennings and others, 1956). since then, intermittent work has been conducted on the caves, taking advantage of new archaeological techniques as they have become available (madsen, 2014). the initial excavation of danger cave was important in helping establish the utility of radiocarbon dating as a valid chronological tool when it proved to be one of the oldest archeological sites in north america known at the time. these caves were occupied repeatedly by indigenous people between about 12,500 cal yr b.p. to historic times (jennings and others, 1956; madsen and rhode 1990; rhode and madsen, 1998; rhode and others, 2005; rhode and others, 2006; goebbel and others, 2007). early people were present here intermittently from ~12,500 to 8,000 cal yr b.p. this area was later used by desert archaic people (~1,500 cal yr b.p.), the fremont (~1,500 to 700 cal yr b.p.), and proto-historic shoshonean groups (~700 cal yr b.p. – present). the extremely good preservation in the dry caves, coupled with their detailed stratigraphy, has provided some of the best evidence of prehistoric lifeways of great basin peoples, as well as records of ecosystem change and paleoclimate. textiles, baskets, pottery, animal bones, plant remains, weapons, chipped stones, coprolites, quids (chewed bits of fibrous food), arrowheads, and leather scraps have all been found in the caves. paleoenvironmental record juke box cave is located near the stansbury shoreline and stansbury shoreline tufas are visible near the cave’s entrance. plant and animal remains left by the cave’s inhabitants record changes in the surrounding environments, including the juke box spring marshland that existed below the cave and nearby desert and mountain ecosystems. additional paleoenvironmental information from pollen analyzed from cores taken in the marsh and from woodrat nests found in nearby caves containing well-preserved plants, insects, and vertebrate remains supplement the cave records (rhode and madsen, 1998; madsen and others, 2001). these woodrat “middens” can be preserved for tens of thousands of years, providing ecological snapshots of the past, making them invaluable paleoenvironmental tools. bonneville salt flats return to the i-80 frontage road and continue for 1.4 miles until you reach leppy pass road. take a left. continue for 5 miles (at the bend in the road, turn right/east towards the bonneville salt flats). at the end of the pavement, there is a large turn-around area. park here (40.7625° n, 113.8958° w). depending on events and surface conditions you may be able to access the salt crust (stops in these areas are described in the geomorphology section below). access to bonneville salt flats crust is limited seasonally by surface flooding. in general, if there is surface moisture at the end of the access road, stay off the salt flats. ignoring this guideline may rip up and damage the crust for years to come (figure 7). in addition, the salt can be thin. it is easy to get stuck in the underlying mud and it is expensive and damaging to be towed out. only drive on the salt when it is dry and when your tires do not leave a track. furthermore, access is limited during events such as speed week. the bonneville salt flats is on public land managed by the bureau of land management. a schedule of bonneville salt flats events is available at https://www.blm.gov/visit/bonneville-salt-flats. when on the salt flats, be careful to watch out for cross-traffic. fatal crashes have happened here before. also, be aware of the state of the crust. if your vehicle is leaving tracks, keep momentum, turn around, and return to the stable crust! the vast expansive landscape of the bonneville salt flats is treasured for different uses by many groups of people. the brines underlying the saline pan are enriched in potassium and have been mined continuously since 1939 (bingham, 1980). the landscape is valued by tourists and artists for its sharp contrasts and stark beauty (zajchowski and others, 2020; bowen and wischer, 2023). the hard flat surface is treasured by the vehicular land-speed racing community for its flatness, mechanical properties, and length (francisco, 1965). social and physical scientists also value this landscape, which provides an example of saline processes influenced by human action 9 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 7. aerial imagery of the bonneville salt flats. (a) racetracks and view of silver island mountains. (b) water input from south of interstate 80. (c) car tracks left near the end of the access road, view of potash mine to the south. note that tracks are more evident to the west, towards the edge of the saline pan where the crust is thinner. north arrows added to show orientation. images taken with pilot dr. gabe bowen on (a) august 17, 2019, (b) february 19, 2023, and (c) october 30, 2022. these and other bonneville salt flats aerial imagery available at: https://geodata.geology.utah.gov/pages/search.php?search=%21collection129324. 10 j.a. bernau, b.b. bowen, c.g. oviatt, and d.l. clark great salt lake desert landscape change over multiple temporal scales (e.g., christiansen, 1963; zajchowski and others, 2020). here, we describe a century of changes at the bonneville salt flats and provide context for these changes through the window of the bonneville salt flats’ depositional history. we then describe the surface expression of the crust and how it varies spatially and temporally. recent change the bonneville salt flats' recent history is characterized by changes in salt crust area and thickness (figure 8). furthermore, the geochemistry of bonneville salt flats brines has changed over time in response to changing management (bernau and others, 2023a, this volume). anthropogenic activities during this period are strong contributors to this change. the saline pan has been dissected by industrial activity and an interstate highway, and its waters have been collected for mineral production. these changes have upset and limited the ability of multiple stakeholders to use the site, spurring several salt crust thickness studies, research, and lobbying for action (francisco, 1965; kipnis and bowen, 2018). for a quarter century, restoration efforts at the bonneville salt flats have focused on a brine “laydown” program. to the north of juke box trench, there are alluvial-fan aquifer production wells. these brackish water wells are used to provide the potash mine with water for its operations. they have also been used to supply water for the laydown program since 1997. the laydown uses alluvial-fan water to dissolve waste halite from mine operations, the resulting brine then floods the bonneville salt flats’ surface. this project has not had anticipated results, and the crust has continued to decrease in area and thickness (figure 8b and d) (white, 2004; bowen and others, 2017; kipnis and bowen, 2018). alluvial-fan aquifer extraction may be exacerbating long-term crust declines (bernau and others, 2023a, this volume). since the onset of the laydown, groundwater levels in the alluvial fan have steadily fallen, leading hydraulic gradients to reverse; instead of water flowing towards the saline pan as it used to, it now flows away from the saline pan. this is evidenced by the salinity of the waters the wells now produce–they used to be fresh, but several wells now produce brine that is saltier than the ocean (bernau and others, 2023a, this volume). a portion of this water likely comes from groundwaters underlying the saline pan. dewatering from lowering alluvial-fan water levels (from surface to ~50 ft below the surface in 2021) has created >3-ft-wide desiccation fractures in areas near the mountain front (mason and kipp, 1998). depositional history the depositional history of the bonneville salt flats’ site provides insights and perspective on modern change. before lake bonneville, there were intermittent shallow saline lakes similar to the modern great sale lake (>45,000 to >28,000 cal yr b.p.) (figure 2; bernau, 2022). these lake deposits have small faults and soft sediment deformation features, suggesting past seismic activity; an alternative interpretation of these features is that they are dewatering structures that developed as lake levels fell, and water was released from sediments as the overlying pressure of lake waters was removed. a fault with ~1.5 ft of holocene offset along the southeastern silver island mountains in a former spring area suggests that seismic activity may be ongoing (hecker, 1993; madsen, d., personal communication, 2022) (further investigation and interpretation of these sediments is needed). from the bonneville salt flats you can see shorelines left by lake bonneville on the silver island mountains and the leppy hills (figure 7a). the bonneville salt flats salt crust consists of layers of gypsum sand and halite crystals. the gypsum sand becomes coarser with increasing depth, indicating the displacive growth of crystals after deposition (bowen and others, 2018; bernau and bowen, 2021). bonneville salt flats’ gypsum deposition began at ~8,000 cal yr b.p. (bernau, 2022). the origin of the gypsum sand is likely in-situ growth, which is seen in some modern sediments. some grains may originate from displacive crystals that were later reworked with deflation. radiocarbon dating of pollen, and similarly sized material from bedded bonneville salt flats evaporites, indicates that the bonneville salt flats is much younger than previously thought. the bonneville salt flats likely resembled today’s saline pan by 5,500 cal yr b.p., not immediately after lake bonneville (13,000 to 11,000 cal yr b.p.), as was previously thought. this new evidence indicates the saline pan may be a much more ephemeral feature than assumed. similarly, the bonneville salt flats’ depositional history with respect to regional changes in climate indicates that halite is deposited under wetter conditions whereas gypsum is deposited under drier conditions. recent records suggest this region is becoming drier (williams and others, 2022), making the bonneville salt flats likely to shift towards more gypsum accumulation even in the absence of direct anthropogenic alteration. crust surface morphology you can view a timeline of surface conditions from this location in photos collected by citizen scien11 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 8. overview of change at the bonneville salt flats. (a) photomosaic of 1953 aerial imagery giving an overview of the bonneville salt flats. dashed yellow lines show the extent of surface halite from aerial imagery and black outline shows the oldest mapped extent of surface halite (nolan, 1927). aerial imagery is available in non-photomosaic form at https://imagery.geology.utah.gov/pages/home.php and in photomosaic form at https://geodata.geology.utah.gov/pages/search.php?search=%21collection129324. (b) a consistent long-term trend of declining saline extent is evident in analyzed landsat data mapping the areal extent of end-desiccation surface halite over the bonneville salt flats area north of i-80 (bowen and others, 2017). (c) recalculated (to adjust for differences in methodology) crust volume across salt crust thickness studies (1960 to 2016) (kipnis and bowen, 2018). (d) change in the area of 3 ft crust thickness contour at the bonneville salt flats across studies (modified from kipnis and bowen, 2018). https://geodata.geology.utah.gov/pages/search.php?search=%21collection129324 https://geodata.geology.utah.gov/pages/search.php?search=%21collection129324 12 j.a. bernau, b.b. bowen, c.g. oviatt, and d.l. clark great salt lake desert landscape change over multiple temporal scales tists here: https://www.chronolog.io/site/bsf101. these images highlight how variable the surface of the crust can be seasonally, or from week to week. less than 1 cm of precipitation has led to >20 cm of flooding at this location. heavy summer precipitation can rapidly alter conditions and event plans. storms and deteriorating crust conditions led to the cancellation of racing events in 1993, 1994, 2014, 2015, and 2022 (kipnis and bowen, 2018). the surface expression of evaporites at the bonneville salt flats is dynamic and changes in response to cycles of flooding, evapoconcentration, and desiccation (figure 9) (lowenstein and hardie, 1985; bernau and bowen, 2021). during the flooding stage, rainfall contributes to the full or partial dissolution of halite (figure 9). this is apparent in dissolution pits that expose darker, gypsum and microbial-rich mud that underlies surface halite layers (figure 10a1 and a2). during the evaporation stage, salt crystals begin to crystallize on the surface of the brine as rafts or at the sediment-water interface as bottom-growth crystals (figure 10b1 and b2). one unique feature at the bonneville salt flats is salt blisters (figure 10b3), these may form from the remobilization of trapped air under a crust after flooding. most people know the bonneville salt flats from its appearance during the desiccation stage. the crust is in this stage when surface water is completely removed by evaporation. many surface morphologies form during this period, with the most rapid growth occurring immediately after the surface enters the desiccation stage when nearsurface pores are larger (not filled by crystal growth) and contain water. one diagnostic feature of this period is efflorescent (or popcorn) halite (figure 10c). these aptly named crystals effloresce, or bloom, from the ground. during the desiccation stage, efflorescent growth causes the crust’s morphology to change. on the salt flats, we will cover the transition from the thincrusted, pressure-buckled crust at the western edge of the bonneville salt flats to the smooth-crusted transitional zone near the raceway and weather station (figure 11). finally, we will move towards the salduro loop, an area covered by polygonal crust. the surface expression of the crust is influenced by its thickness and history (figures 11e, 12, and 13). for more information about the surface expression of halite crusts and other surface features in similar settings see christiansen (1963), lines (1979), goodall and others (2000), wang and others (2014), el-maarry and others (2015), nield and others (2015), milewski and others (2017), lasser and others (2020), bernau and bowen (2021), and zhang and others (2021). buckled crust our first stop on the bonneville salt flats’ crust occurs on its western edge. drive onto the saline pan and follow the main traffic area (salt crust is smoother) to the northeast for 4.5 miles (40.8054° n, 113.8309° w). then turn west towards the silver island mountains and continue until you see buckled crust (figure 11a). note: if you approach this area in the midsummer or early fall and there has not been a recent flooding event you will see pressure-buckled crust. you may opt to stop and walk to this area if the surface is becoming less stable and you are beginning to create tracks. avoid getting stuck – turn around if the surface is unstable! the buckled crust is underlain by a thin layer of gypsum sand over carbonate mud. it is located near the salt-flat to mudflat transition, so wind-blown sediment may easily accumulate on these buckles. the buckled morphology forms as the crust bends to accommodate increases in crust volume. as the crust buckles it may transport sediment on its underside as well as sediments deposited by wind on its surface tofigure 9. flooding, evapoconcentration, and desiccation periods at the bonneville salt flats (modified from bernau and bowen, 2021). 13 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 10. surficial expression of the bonneville salt flats across flooding, evaporation, and desiccation periods (modified from bernau and bowen, 2021). (a) flooding features include halite-undersaturated brine and partial (dissolution pits [a2]) to full dissolution of the surface crust. (b) evaporation stage, where halite crystalizes on the water surface as rafts (white arrow) (b1 and b2). (b1) view to west of prominent lake bonneville shorelines looking west. (b3) halite blister feature where halite has bulged up after flooding. this feature is surrounded by insects, which accumulated at the water level line and bottom-growth halite. halite blisters often occur in the southern racetrack area in the autumn after the surface has desiccated after flooding. blisters are shown forming in figure 11. (c) as the surface shifts from the flooding to the desiccation period efflorescent (or popcorn) halite (black arrow) (c2) forms. 14 j.a. bernau, b.b. bowen, c.g. oviatt, and d.l. clark great salt lake desert landscape change over multiple temporal scales ward the area of buckling. this may lead to the formation of a small detrital sediment ridge; this ridge creates a preferred area for buckling to occur in the future because crusts deposited over the ridge will be thinner and easier to break, creating a feedback loop for buckling to regularly occur in the same area and for more sediment to accumulate at the same spot (figure 12 left) (lokier and steuber, 2009; lokier, 2012). observations of other areas of the saline pan indicate that the surface expression of buckling can be highly variable in height, spacing, and shape (polygonal vs. orthogonal), potentially depending on the thickness of surface halite, mineralogy, and other factors (for example, buckles often form where car tracks have created a preferential break point). flat crust our next stop is near the center of the bonneville salt flats at the bflat weather station. drive south for ~1.5 miles (40.7846° n, 113.8297° w). the weather station should become visible as you near it; it has a chain link fence surrounding it. note: watch for cross traffic, this route cuts across several racetracks and high traffic areas. you may notice as you travel towards the saline pan center that the crust’s morphology changes and is generally flatter. in areas where the racetrack has been prepped even efflorescent crystals become subdued. the racetrack is prepared by dragging heavy steel beams behind a vehicle to crush and homogenize the crust (morgan, 1985). if the crust is too thin or if conditions are too moist, preparing the racetrack can rip the crust, degrading its quality and limiting the ability to safely race. in addition to racetrack preparation, regular flooding of this area (a seasonal pond is concentrated on the bonneville salt flats’ western edge, a topographic low point) and salt crust thickness likely contribute to its flatness. seasonal flooding at the bonneville salt flats is one of the features that makes it so ideal for land-speed racing. flooding removes any buckles in the crust and redistributes sediment, and when the pond desiccates it leaves a new flat crust. areas that figure 11. surface and subsurface expression of halite crust (modified from bernau and bowen, 2021). surficial halite can vary widely, but generally is (a) buckled, (b) flat, or (c) polygonal. (d) cross section across a polygon similar to that shown in (c). (e) surface expression of halite crust in relation to surface halite thickness. pressure buckles consistently occur near saline pan’s edges. flat areas coincide with southern and central racetrack areas where seasonal ponding is persistent (bowen and others, 2017; craft and horel, 2019). the polygonal crust is concentrated within and to the northeast of the salduro loop. 15 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 flood less frequently, such as death valley and salar de atacama, can have rough surfaces that develop as the crust deforms during the desiccation stage (bobst and others, 2001). the crust in this flat area is generally thicker than the buckled crust area. a thicker crust may be harder to buckle because of its mechanical properties. similarly, the thicker crust has more pore space that efflorescent crystals could develop within, vertically distributing crystal growth and reducing lateral deformational pressures (figures 11e and 12 center). the site of the bflat weather station (operated by the utah geological survey after 2021 and by the university of utah from 2016 to 2021) highlights some of the long-term research performed on the bonneville salt flats. the weather station measures precipitation and evaporation, enabling researchers to understand how water is moving in and out of the crust. it also collects time-lapse images and logs data every 5 minutes, enabling anyone to see current surface conditions on the saline pan (https:// meso1.chpc.utah.edu/station_cameras/bflat_cam/ bflat_cam_current.jpg). similarly, several groundwater monitoring wells with multiple depths are present here. these wells are used to understand how groundwater levels change in response to climate and human actions. they also enable researchers to determine if shallow groundwater at the saline pan is moving up to figure 12. schematic of processes leading to the formation of different surface morphologies. the upper (first) section shows the surface at beginning of the desiccation stage. (left) formation of buckled crust, which typically occurs in thin salt at saline pan edge. note the accumulation of sediment at the buckle location as halite crust expands with efflorescent growth (modified from lokier, 2012; lokier and steuber, 2009). (center) formation of flat crust and blisters. the flat crust is generally thicker, limiting its ability to buckle or reach heights seen in thinner crust areas. blisters are thought to form after flat areas flood, the lack of surface buckles and breaks in these areas limit off-gassing, enabling trapped gases to move laterally, accumulate, and bubble up in an area, deforming the crust in the process (based on the description in bernau and bowen [2021] and similar to microbial mat gas domes in other settings [goodall and others, 2000; noffke and others, 2002]). (right) polygonal crust development occurs in areas where halite crust is persistent across flooding events; although a contractional origin of polygons has been proposed, the features at the bonneville salt flats may be explained by repeated cycles of flooding with preferred dissolution occurring along polygonal edges (either pre-existing or formed by buckling). the gaps then become preferred areas of dissolution (see figure 11d and bernau and bowen [2021] for further reference). note that under extended flooding periods, the surface halite can completely dissolve near the bonneville salt flats’ center; during these periods, remnant gypsum becomes rippled from wave action. 16 j.a. bernau, b.b. bowen, c.g. oviatt, and d.l. clark great salt lake desert landscape change over multiple temporal scales moving up to feed the crust or if it is moving down and removing salts from the saline pan. polygonal crust our final stop on the bonneville salt flats’ crust visits the polygonal crust. head 1 mile south from bflat towards the berm of the salduro loop (40.7700° n, 113.8283° w). the salduro loop was a brine collection ditch. the crust next to it can be soft, in some areas inside the loop you can see where the ditch filled in with halite by identifying deep brinefilled holes in the salt. to minimize the chances of getting stuck, do not drive within 100 ft of the loop. furthermore, the salt where the ditch has filled in with halite may be fragile and could collapse beneath you, step with care! this site has some of the thickest surface crust at the bonneville salt flats. it also has a distinctive polygonal fracture system. the polygons are highlighted by vertical ridges of efflorescent halite. the efflorescent halite is very porous and dissolves during flooding, inverting the local topography (leaving a crack where a ridge once was). this crust is persistent across multiple flooding events. under extended flooding periods, the surface halite can completely dissolve near the saline pan’s center and remnant gypsum becomes rippled. although a contractional origin of salt polygons has been proposed (tucker, 1981), the features at the bonneville salt flats can be explained by repeated cycles of flooding with preferred dissolution occurring along polygonal edges (either pre-existing or formed by buckling) and the growth of efflorescent salt at the surface (bernau and bowen, 2021) (figures 11d and 12 right). the polygonal geometries at the bonneville salt flats occur at multiple scales, ranging from less than a meter to over 300 meters across. the multiple scales of polygons can be seen in person, in aerial imagery (figure 13), and, at the largest scale, in multispectral satellite (resolution up to ~100 ft/pixel) spectral index images. using different methods and examining polygons at a much smaller scale, lasser and others (2020) present evidence for convection occurring beneath polygonal crusts, indicating a strong relationship between the surface expression of saline pans and groundwater movement beneath them. you may now enjoy exploring the rest of the bonneville salt flats’ crust or return to salt lake city. to return to the access road head west and follow the salduro loop berm (~4 miles), then return west towards the access road which will become be visible (40.7625° n, 113.8958° w). silver island mountains access road (optional extension) the silver island mountains access road provides another great perspective on this area, specifically on the contact between deflated and non-deflated surfaces. to access it, go west from the end of the access road for 3.8 miles. at the t in the road, go right. continue for 0.8 miles and take a slight right onto silver island road (unmarked) and continue north for ~14.5 miles (40.8926° n, 113.7978° w). note: this road is periodically maintained and may have heavily rutted or muddy areas that require a high clearance vehicle. the exposed gravel bar here is enhanced by erosion; it is known as the lozenge section and highlights the sharp contact between preserved and deflated areas. the lozenge section is capped by lake bonneville's late-regressive-phase well-rounded gravels. lake bonneville sediments are likely preserved here because the gravels limited deflation (figure 14a). below the gravels are reddish, silty, sandy beds that figure 13. aerial imagery of the bonneville salt flats on june 9, 2022. (a) view looking north. (b) large polygonal features (black arrow). images taken with pilot dr. gabe bowen. 17 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 overlie lake bonneville and underlying pre-lake bonneville sediments. the eroded mudflat is to the east and the alluvial-fan deposits, which have limited erosion, are to the west (figure 14b). if you continue northeast on this road, you will see similar erosional features and will reach a road that connects with floating island (40.9295° n, 113.7189° w). turn around and retrace your route. the final stop is an alluvial fan that demonstrates the erosional contact between deflated and preserved lake bonneville sediments. this stop coincides with the furthest northern extent of groundwater extraction, a now-dormant spring, and desiccation fissures. return south along silver island road for 11 miles. stop just after the promontory (40.7960° n, 113.9395° w). to the east of this area is a former spring (40.7947° n, 113.9325° w). the eastern linear feature is a now inactive freshwater collection ditch for the northernmost extent of now inactive brackish water production wells (figure 12c). the mudflat area between the former collection ditch and the alluvial fan has many large, deep (several feet) desiccation fissures (mason and kipp, 1998). there are no trails here so choose your steps carefully. figure 14. edge of deflation along the silver island mountains’ alluvial-fan edge. (a) the lozenge section (image modified from munroe and others [2015] and site further described in oviatt and others [2020]). (b and c) preservation and deflation areas are delineated by changes in elevation and land cover. note that the preservation of lake bonneville sediments delineated by surface cover is also evident at juke box trench (figure 6). (c) is a usgs national aerial photography program (napp) photo from july 1997. dashed yellow line delineates between the areas of preservation (alluvium and gravel bars) and deflation. 18 j.a. bernau, b.b. bowen, c.g. oviatt, and d.l. clark great salt lake desert landscape change over multiple temporal scales directions back to salt lake city return to i-80 and head east (~120 miles). if you drove on the salt flats you may wish to wash your vehicle in wendover or back in salt lake city. salt accumulation depends on surface conditions; for example, salt may more easily accumulate on your vehicle if the bonneville salt flats has flooded recently or if groundwater levels are near the surface, as they are in the mid-summer (bernau, 2022). the crust is driest in the fall to winter, when groundwater levels decline if there has been no precipitation (salt accumulation on your vehicle will be lower under these conditions). salt will cake onto the surface and get onto ledges and crannies beneath a vehicle. we recommend using a self-service carwash (with hot water if possible) to ensure salt has been removed and to limit potential corrosion. acknowledgments we acknowledge that this study was conducted on traditionally newe/western shoshone and goshute lands. we thank craig peterson and russ draper with intrepid potash; bureau of land management west desert district's current and former office staff, including kevin oliver, matt preston, mike nelson, cheryl johnson, steve allen, roxanne tea, and todd marks; university of utah researchers jory lerback, evan kipnis, and mark radwin; and utah geological survey geologists paul inkenbrandt, hugh hurlow, bill keach, and elliot jagniecki. isaac hart and andrea brunelle (university of utah), and stephanie carnie and mike hylland (utah geological survey) are appreciated for their helpful reviews and comments, and genevieve atwood is thanked for acting as the editor of this paper. we also thank steve bowman for help scanning past bonneville salt flats’ files and setting up the geodata archive and david madsen for enriching our discussion of danger and juke box caves. we also thank dr. gabe bowen for piloting several flights over the great salt lake desert, enabling us to collect aerial imagery. funding for this work was provided by an nsf coupled natural human systems award #1617473, the utah geological survey, the utah state legislature, the bureau of land management, an american association of petroleum geologists grant-in-aid, and university of utah global change and sustainability center graduate research grants. funding for geologic mapping in the area is by the statemap and fedmap components of the national cooperative geologic mapping program. references benison, k.c., and karmanocky, f.j., 2014, could microorganisms be preserved in mars gypsum? insights from terrestrial examples: geology, v. 42, no. 7, p. 615–617. bernau, j.a., 2022, spatial and temporal scales of water and salt movement at the 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cook, b.i., and smerdon, j.e., 2022, rapid intensification of the emerging southwestern north american megadrought in 2020–2021: nature climate change, v. 12, no. 3, p. 232–234. wold, s.r., and waddell, k.m., 1994, salt budget for west pond, utah, april 1987 to june 1989: water-resources investigations report 93-4028, no. april 1987, p. 1–20. zajchowski, c.a.b., brownlee, m.t.j., blacketer, m.p., peterson, b.a., craft, k., and bowen, b.b., 2020, rapid resource change and visitor-use management—social–ecological connections at the bonneville salt flats: environmental management, v. 66, no. 2, p. 263–277. zhang, g., xiao, y., xiang, m., hong, c., zhang, b.t., liu, l., shi, p., and liu, j., 2021, structure and morphological characteristics of polygonal salt crust, the west juyan lake, china: geosciences journal, p. 1-12. bernaugeochem.pub 1 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2023 utah geological association publication 51 abstract over the past century, the bonneville salt flats, which lies on the western edge of the great salt lake watershed, has experienced changing environmental conditions and a unique history of land use, including resource extraction and recreation. the perennial halite salt crust has decreased in thickness since at least 1960. an experimental restoration project to return mined solutes began in 1997, but it has not resulted in anticipated salt crust growth. here, primary observations of the bonneville salt flats surface and subsurface brine chemistry and water levels collected from 2013 to 2023 are reported. spatial and temporal patterns in chemistry, focused on density and water stable isotopes, are evaluated and compared with observations across seven periods of research spanning from 1925 to 2023. declining salinity in the areas to the east of extraction ditches and south of interstate 80 were observed. brine extracted for potash production decreased in salinity as extraction rates increased. between the years 1964 and 1997, brine in the shallow aquifer located beneath and to the east of the crust experienced a decrease in salinity. however, following this period, the salinity stabilized and subsequently increased. salinity recovery was concurrent with declines in brine extraction and the salt restoration project, with the largest decrease in brine extraction being concurrent with the largest recovery in salinity. the specific impact of the restoration project on the brine salinity increase remains unclear. to the west, the shallow aquifer in the area between the silver island mountains and the salt crust has increased in salinity. this increase is accompanied by a decline in groundwater levels, which enables the underground movement of solutes from east to west, following a salinity gradient away from the saline pan. over the past 25 years, water levels in the alluvial-fan aquifer along the silver island mountains have markedly declined, leading to the extraction of increasingly more saline and isotopically heavier basinal waters are intriguing landscapes for industrial use. this change is concurrent with the onset of the salt restoration project, which relies on alluvial-fan aquifer waters. this article’s compilation of changes in groundwater chemistry provides an important resource for stakeholders working to understand and manage this dynamic and ephemeral evaporite system. it also offers an example of decadal-scale change in a highly managed great salt lake watershed saline system. observations of decadal-scale brine chemistry change at the bonneville salt flats, utah   jeremiah a. bernau1,2, brenda b. bowen1, evan l. kipnis1, and jory c. lerback1,3 1department of geology and geophysics, university of utah, salt lake city, utah, jeremiahbernau@gmail.com 2utah geological survey, salt lake city, utah 3department of earth, planetary and space sciences, university of california, los angeles, california introduction saline pans, shallow depressions encrusted by evaporites where waters accumulate, provide an intriguing example where groundwater level and chemistry, climate, and anthropogenic activities converge. here, several decades of chemical measurements, with a focus on density, are used to examine how the bonneville salt flats (bsf) groundwater system changed in response to: 1) brine extraction for potash production, 2) alluvial-fan groundwater extraction for industrial uses, and 3) 25 years of an experimental brine “laydown” program to restore the saline pan (figures 1 to 3) (kipnis and bowen, 2018). the laydown program uses alluvial-fan aquifer groundwater to dissolve the potash mine’s halite (nacl) byproduct and transport it to the saline pan in hopes of restoring saline pan thickness and extent. multidecadal analyses of satellite imagery and reoccurring measurements of salt crust thickness show long-term declines in crust thickness and extent (bowen and others, 2017; bowen and others, 2018; radwin and bowen, 2021). one-third of the crust consists of halite (nacl) and two-thirds of it is gypsum (caso4·2h2o). the multi-decadal nature of research on bsf and the uniquely involved mix of stakeholders including racing enthusiasts, recreational visitors, the potash industry, researchers, and governmental managers make this site well-suited for examining the evolution of brine chemistry. this landscape is dynamic with seasonal to decadal-scale changes in flooding and saline pan volume (figure 2e and f) (bowen and others, 2017). here, this examination of long-term changes in brine chemistry, with groundwater levels as a secondary dataset, provides context for the relative impact of extraction and restoration 10.31711/ugap.v51i.143 2 j.a. bernau, b.b. bowen, e.l. kipnis, and j.c. lerback observations of decadal-scale brine chemistry change at the bonneville salt flats figure 1. site overview. a. off-angle international space station image of bonneville salt flats, great salt lake desert, and great salt lake (https://earthobservatory.nasa.gov/images/91765/bonneville-saltflats). b) areas of investigation. sample areas (divided as zones moving east from west) and depths of different sampling wells noted in figure explanation. inset c) shows locations of brackish water (bw) alluvial-fan aquifer production wells (a-a’) and nested observational wells (ow) (b-b’) along transects. primary production wells outlined by box in the middle of a-a’. d) schematic of investigated aquifer intervals (not to scale). basemap imagery from earthstar geographics. 3 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 2. features influencing brine density and surface variation at bsf. a). aerial photograph of potash mine looking north to bsf, letters denote relative locations of b to d. b) alluvial-fan aquifer wells collect brackish water that is used in mine operations and to create salt laydown. c) brine collection ditch (6 m deep) (bingham, 1980) east of bsf (looking north). d) laydown brine being introduced to bsf’s southwest corner, person on right for scale. e and f) time-lapse photos from bsf weather station on e) may 28, 2018, when the surface was flooded, and f) july 15, 2018, when the surface was desiccated. more field and aerial imagery of bsf is available with the utah geological survey data archive system at https://geodata.geology.utah.gov/pages/search.php?search=%21collection129324. 4 j.a. bernau, b.b. bowen, e.l. kipnis, and j.c. lerback observations of decadal-scale brine chemistry change at the bonneville salt flats figure 3. human activities (mining and laydown) that impact brine and solute mass through time. changes in brine mass balance (a and b) in millions of tonnes (mt) of nacl (not brine), the density of extracted brines (c and d), and alluvial-fan aquifer (af) groundwater extraction and groundwater levels (e). a) annual values nacl mass extracted from bsf’s western ditches, added to the saline pan through the laydown, and net annual nacl balance. b) cumulative net movement of nacl onto bsf from the laydown. c) changes in extracted brine salinity over time. note seasonality in density; summer values used for long-term trend. numbered red arrows added to b) and c) highlight notable inflection points in data. in 2006 (point 1) the salinity of extracted brine began a long-term decline. net cumulative brine contributions were neutral between 2005 and 2010 (point 2), and increased afterward; despite this increase, extracted brines did not begin to recover in salinity until 2015 (point 3). after 2017 (point 4) extracted brine salinity remains relatively high, while extraction is low. d) correlation between pumping and extracted brine salinity (each dot represents the average of 8 months of pumping and brine density values) from aug. 2018 to dec. 2020 (these values are reported as hours of pumping, ~0.85 acre-ft/hour). e) alluvial-fan annual groundwater extraction and groundwater elevation (meters above sea level) change over time (well located at 2.3 km in figure 12) (103 acre-ft ≈ 1.23*106 m3). the vertical dashed red line shows when laydown began, and the gray area on (a) and (d) emphasizes the period with elevated laydown volumes and lower than anticipated (light blue line in lower e) groundwater levels. alluvial-fan extraction data from https://www.waterrights.utah.gov/cgi-bin/wuseview.exe?modinfo=wruseage&wrnum=16-25. 5 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 activities on changing saline pan extent and volume (figure 3). this improved knowledge will help guide plans to optimize the sustained use, including both mining and racing, of this landscape. methods brine chemistry was characterized and groundwater levels were measured at bsf between 2015 and 2023. measurements prior to 2015 were compiled. this study involved extensive density measurement quality control, analyzing trends in individual wells, and analyzing data aggregated by area and aquifer. here, groundwater levels are reported to provide context for observed changes; however, they are not the primary focus of this work. data sources over time groundwater level and density data were divided over seven study periods that vary in length, intervals between periods, reported data, areas of investigation, and researchers (figures 4 and 5). nolan (1927) investigated the composition of great salt lake desert (gsld) brines. in 1925, nolan made shallow borings across the gsld and reported sample total dissolved solids (tds), major ions, and groundwater levels. nolan made major-ion measurements in the field through titration (for chloride) and by measuring the volume of precipitate formed after adding chemicals to the solution (for potassium and sulfate). several of these sites occur on the southern edge of bsf and the area between the silver island mountains and the saline pan crust (figure 4b). between 1946 and 1949, 23 alluvial-fan production wells (named bw for brackish water) were drilled to depths ranging between 32 and 111 m (bernau and others, 2023a) (figure 1c). bw wells flowed freely (1.1 to 9.5 m3/second) when they were drilled, with reported potentiometric surfaces between 1.5 and 6 m above surface level. bw well water chemistry was reported in aggregate. two researchers investigated bsf chemistry between 1964 and 1972 (figure 4c). in 1964 brine samples from shallow gsld borings were analyzed for major ions, lithium, and tds (lindenburg, 1974). between 1965 and 1967, turk (1973) installed shallow (<9 m depth) wells across and adjoining bsf’s crust, and measured groundwater levels and brine chemistry. in 1972, two brine chemistry samples from the alluvial-fan aquifer wells were collected (reported in the water quality portal; read and others, 2017). between 1975 and 1981 the u.s. geological survey performed two bsf studies (figure 4d). groundwater levels and chemistry were examined between 1975 and 1978 (lines, 1978, 1979). in 1981 the u.s. geological survey conservation division measured brine chemistry from borings and wells (reported in white, 2002). the next study period occurred leading up to the onset of the laydown in the autumn of 1997. two groups collected measurements between 1991 and 1997 (figure 4e). the u.s. geological survey measured bsf between 1991 and 1993 (mason and others, 1995; mason and kipp, 1998). the bureau of land management made annual measurements from a subset of wells between 1994 and 1997 (white, 2002). between 1998 and 2006 the bureau of land management monitored brine chemistry and water levels in several wells to evaluate the laydown’s impact (figure 4f) (white, 2002). some unpublished bureau of land management measurements collected between 2003 and 2006 are compiled here (white, field notes and files including laboratory results, 1998 to 2014). between 2003 and 2018, shaw environmental, inc. conducted biannual measurements of groundwater level and chemistry from various locations. these measurements, amounting to over 900 measurement sets, were gathered on behalf of the potash mine as part of its mine reclamation plan (shaw environmental, 2020). while most sites were within the potash mine, many samples were from the bsf study area. these include samples from the alluvial fan aquifer, east of the extraction ditch, and south of interstate 80 (i-80) areas. the samples for the period spanning from 2007 to 2012 exclusively originate from this report. between 2015 and 2022, researchers from the university of utah collected brine chemistry and groundwater level measurements, and between 2022 and 2023 the utah geological survey collected similar measurements (figure 4g) (penrod, 2016; bowen and others, 2018; kipnis and bowen, 2018; lerback and others, 2019; kipnis and others, 2020; bernau and bowen, 2021; bernau and others, 2023a). the bureau of land management made groundwater level and density measurements independently (white, field notes and files including laboratory results, 1998 to 2014) and in collaboration with the university of utah in 2015. measurements from may 2016 to may 2020, a precipitation sample collector with internal electrical heating was installed at the potash mine. mine staff monitored and collected precipitation samples regularly, offering storm-event-level resolution for collected precipita6 j.a. bernau, b.b. bowen, e.l. kipnis, and j.c. lerback observations of decadal-scale brine chemistry change at the bonneville salt flats figure 4. spatial distribution of shallow brine aquifer density measurements from the bonneville salt flats and surrounding area. a) compilation from density measurements between 1925 and 2023. b to g) locations of shallow brine aquifer density measurements over time. kriging done with empirical bayesian kriging (geostatistical analyst) in arcgis. 7 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 5. brine density measurements over time (~2000 data points). measurements are separated by area and depth of investigation. note the difference in vertical scales. light red arrows added to highlight trends by area. 8 j.a. bernau, b.b. bowen, e.l. kipnis, and j.c. lerback observations of decadal-scale brine chemistry change at the bonneville salt flats tion. stable water isotopes of hydrogen (δ2h) and oxygen (δ18o) were measured at the spatial lab at the university of utah with a picarro l2130-i cavity ring-down spectrometer. water isotope concentrations were calibrated to laboratory standards, and values were reported as per mil (‰) relative to the vsmow scale. between 2016 and 2021, ground and surface water samples were collected in acid-washed and deionized water-rinsed bottles and returned to the university of utah for major ion and stable isotope analyses. a subset of samples was analyzed for major ion concentrations by external laboratories using inductively coupled plasma mass spectrometry (icp-ms/aes) and ion chromatography as reported in kipnis and others (2020). samples collected in 2017 and 2020 were analyzed for major ion concentrations using a portable x-ray fluorescence spectrometer (pxrf) and calibrated to the methods described in kipnis and others (2020). calcium concentrations and total dissolved solids were excluded from pxrf results due to data calibration challenges. additional samples were collected in 2017 and 2018 for chemical analyses including radiocarbon (c14) and tritium (3h/3he) dating as reported in lerback and others (2019). samples collected in 2021 were measured for major ions using icp-ms and ion chromatography at the university of utah earth core facility. samples collected in 2022 were analyzed for major ion concentrations by external laboratories (chemtech-ford, utah public health laboratory). brine density values were measured in field and lab settings with a mettler-toledo densito 30px. laboratory measurements of density at 20°c were made following the procedures of bernau and others (2023a) at 20 °c. before 2022, when analyzing hydrogen (δ2h) and oxygen (δ18o) stable isotopes from brine samples, solutes were removed through cryogenic vacuum extraction before stable isotope measurement. starting in 2022, vacuum extraction was not used for brine samples analyzed for stable isotopes. data quality control due to the diversity in data vintages, care was taken in reviewing data quality. because of this, changes in density were primarily focused on as density measurements are relatively robust over time and are less susceptible to methodological changes (bernau and others, 2023a). salinity is directly proportional to brine density and the two terms are used interchangeably here. in addition to data quality, seasonal changes in brine salinity were considered. seasonal variations in brine density, as observed from monitoring well brines and brines extracted for potash production, indicate that samples collected during cooler and wetter winter to spring months are more likely to have depressed density measurements (figure 6). measurements with higher density from warmer, drier months (july to september, preferably august), are preferred for long-term evaluation. while the impact of temperature on density measurements was taken into account and corrected for whenever feasible (with a change of approximately 0.01 g/cm3 observed between temperatures of 10 and 30 °c), it should be noted that warmer brines can dissolve more halite, resulting in higher densities (bernau and others, 2023a). to mitigate the influence of dilution caused by flooding, it is advisable to utilize the highest recorded density at a site during a study period for long-term comparisons. brine chemistry reporting varied across studies. some studies only reported field density, others reported major ion chemistry and periodically tds, whereas others reported laboratory density measurements in addition to the measurements above. field densities with reported temperature (if available) were corrected to the density at 20°c using equation 11 in bernau and others (2023a). following the methods in bernau and others (2023a), major ion data were used to model density using the spece8 module of geochemists’ workbench® with the phrqpitz thermodynamic dataset (pitzer, 1973; harvie and others, 1980; plummer and others, 1988; bethke, 2013). finally, available measurements were utilized to establish the correlation between tds and density for bsf brines. using the measured tds data, a salinity for these brines was estimated. chemical model-based estimates of density at bsf tend to underestimate density, indicating that major ion concentrations are typically underreported (bernau and others, 2023a). when all of these measurement types and estimates of salinity were available, they were contrasted to delineate measurement quality. for long-term comparative analyses, laboratory density measurements were prioritized, then field density measurements, followed by chemically modeled density, and finally density estimated from tds. an additional step in data quality control was made using site-based knowledge to assess data quality and identify erroneous data to remove. for example, anomalously high density values (>1.22 g/cm3) are not possible at bsf given its brine composition and suggest measurement errors, such as suspended sediment increasing field density measurements. an additional consideration was unusually low density measurements. for example, some samples from the years 1991 to 1993 have unusually low reported densities. samples that contain higher levels of sodium m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 9 figure 6. seasonality in density measurements across different sample areas. data from 2000 to 2020 for extracted brines and 2015 to 2022 for well samples. well locations are shown on figure 1b. in general, brine salinity is highest and most consistent in august and september. extracted brine density increases in the summer with peak evaporation and decreases in the winter with decreased evaporation and increased precipitation. halite nucleus salinity reflects the impact of temperature on the solubility of halite. as the crust warms, it’s brine can dissolve more halite, and increase in density. the halite nucleus shallow aquifer salinity is highest at the end of summer and lowest in the spring. lower salinity in early spring may reflect the upward movement of deeper, less saline brine as the surface warms (bernau and bowen, 2021). brine densities in the transitional zone east shallow aquifer generally peak in september, potentially reflecting evapoconcentration. 10 j.a. bernau, b.b. bowen, e.l. kipnis, and j.c. lerback observations of decadal-scale brine chemistry change at the bonneville salt flats and chloride ions exhibit greater densities. when these samples reach halite saturation, as they are when they are in contact with halite for a sustained period, any changes in density are limited by the process of halite dissolution or crystallization. in the case of brines in contact with halite, there have been instances where the reported densities were below the density of halite saturation. this was exceptionally clear in samples from the crust-hosted aquifer with reported densities below 1.18 g/cm3, which is below halite saturation (1.195 g/cm3 minimum at halite saturation at bsf) (figure 5). these anomalously high and low measurements were omitted from analyses. the final step in data quality control used known spatial distributions of salinity to identify unusually high or low density measurements, which were then assessed and removed or kept if they were consistent with additional site measurements during that period. individual well-based analysis given the spatial heterogeneity of the brine chemistry, it is important to evaluate this system at specific sites over time in addition to characterizing the overall system. well sites were selected for individual well-based analysis if they had multiple measurements across several study periods (a site could include past borings and wells of similar depths at the same location). the kendall test and linear regression (kendall package and linear model function in the r© coding language) (wilkinson and rogers, 1973a; hipel and mcleod, 1994; r core team, 2021; mcleod, 2022) were applied to site data to identify locations with statistically significant (p-value <0.05) long-term trends in density change. the kendall test assumes that long-term trends are consistent and have not changed. because measurements were not uniformly distributed over time they could not be used for mann-kendall or break-point analyses. areas of investigation to better describe changes and aggregate data, studied areas were divided spatially and by aquifer depth (figure 1b and d). lateral divisions areas of investigation were divided based on surface mineralogy, potential groundwater flows, fabricated structures (the interstate highway and drainage ditches), and lateral salinity (density) gradients (figures 1b and 4a). similar terminology based on hydrological fluxes and sedimentology has been used to establish lateral divisions in other saline pans (munk and others, 2021). from west to east, these areas are the transition zone west (tzw), halite nucleus, transition zone east (tze), east of ditches, and i80 south (figure 1). the tzw area includes the region to the west of the persistent halite crust up to the silver island mountain front. this area consists of a mudflat (inner tzw) which transitions into a higher-elevation mudflat with dunes and intermittent vegetation (outer tzw). the location of the inner and outer tzw is reflected in usda soil maps, where the inner tzw area corresponds with a playa unit and the outer tzw area corresponds with a playa-saltair complex with 0 to 1% slope (soil survey staff, accessed march 2023). tzw waters flowed toward the saline pan in the past. before 1946, at least two now-dormant springs near the mountain front flowed between 0.02 and 3.4 m3/minute (utah division of water rights database). lines (1979) reported a decline in hydraulic gradient between the alluvial fan and the saline pan. mason and kipp (1998) also reported outer tzw desiccation fractures (some >1 m wide) and no hydraulic gradient between the saline pan and tzw. kipnis and bowen (2018) also noted a decline in alluvial-fan aquifer groundwater levels beyond historical norms after 1998. the halite nucleus consists of an area with a persistent halite crust (up to 1.5 m thick). the minimum halite extent, as mapped across several decades of aerial imagery, was interpreted as the halite nucleus’ boundary. the salt crust aquifer only occurs beneath the halite nucleus. nolan (1927) noted the halite nucleus extended south to the area of the current potash mine in 1925. the tze is between the halite nucleus and the eastern brine extraction ditches. this area is covered by ephemeral halite crust (precipitated from standing water) and efflorescent salts (primarily halite) overlying authigenic gypsum sand (bernau and bowen, 2021). the area to the east of the brine collection ditches is hydraulically connected to the greater gsld. the east of ditches area is defined as the region that is closest to the drainage ditches and is likely to be impacted by brine extraction. the final area of investigation is the zone immediately to the south of i-80 that is impacted by brine extraction ditches and is isolated from the saline pan by i-80, which prevents overland flow and limits subsurface brine movement (mason and kipp, 1998). before manmade structures were built at bsf, brines could flow from the southern and eastern parts of the gsld to bsf. evidence of this can be seen today in aerial and satellite imagery of seasonal ponds that develop to the 11 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 southeast of the potash mine (radwin and bowen, 2021). aquifer divisions previous studies characterized basin-fill, alluvialfan, and shallow brine aquifers by water chemistry and recharge rates (turk, 1973; lines, 1979; mason and kipp, 1998). aquifers and samples are delineated by depth (figure 1d); these include surficial samples, such as brine collected from extraction ditches and the laydown, and subsurface samples, which are the focus of this study. depth intervals are described from shallowest to deepest. the crust-hosted wells occur in the halite nucleus, where they are screened within 1 m of the surface. brine samples from this aquifer should be at halite saturation because the aquifer is hosted in halite. the shallow aquifer occurs directly under the crust-hosted aquifer in the halite nucleus and is in contact with the surface elsewhere. the shallow, moderate, and deep aquifers occur within lacustrine to saline sediments which consist of carbonaterich mud and gypsum (with gypsum only occurring in deep basinal muds, >50 m depth) (shuey, 1971; stephens, 1974; oviatt and others, 2020; utah division of water rights database). the shallow basinal mud aquifer (reported as the shallow brine aquifer in other publications) occurs from ≥0 to <10 m depth. this aquifer ranges in salinity from 1.04 to 1.21 g/cm3 (figure 4a), is fractured, and contains brine-shrimp fecal pellet intervals, contributing to a higher hydraulic conductivity than anticipated from its mean grain size of silty clay (turk and others, 1973; lines, 1979). the aquifer’s fractures occur in hexagonal patterns. turk and others (1973) proposed that the fractures formed through osmotic desiccation or synaeresis. where multiple wells exist in the same aquifer depth range at the same site, the shallower well or the well with a longer reporting span was used for multi-decadal single-well analyses. turk (1973), turk and others (1973), and lines (1979) estimated the total thickness of this aquifer to be between 4.5 and 8 m. the largest source of recharge to the shallow brine aquifer is meteoric water infiltration through the surface (mason and kipp, 1998). major aquifer discharge sources are the pumping of ditches along the eastern margin of the saline pan and subsurface flow south underneath i-80 (mason and kipp, 1998). wells screened in the moderate depth aquifer occur within ~10 to 30 m depth and only occur within the halite nucleus and tze. permeability in the moderate depth aquifer is far lower than the shallow aquifer (mason and kipp, 1998), possibly due to the absence of fractures and limited connection with overlying higher-permeability aquifers. alluvial-fan brackick water (bw) aquifer wells (>10 to 150 m depth; measured wells occur between 22 and 111 m depth) are screened in muds (which occur from surface to 8 to ~70 m depth across bw wells) to alluvialfan gravels (stephens, 1974; bernau and others, 2023b). the observation well (ow) alluvial-fan aquifer wells (figure 1c) do not have any reported logs with them, but the ow wells closer to the silver island mountain front reach depths known to intersect gravel lenses. two wells at bsf’s center occur within deep (~30 to ~250 m depth) basinal muds and possibly bedded gypsum; this lithology interpretation is based on deep brine well logs from the potash mine to the south (utah division of water rights database; bernau and others, 2023b). these deep basinal wells have 3-mlong screens at ~70 and 150 m depth. underlying deep basinal muds are basinal gravels, which occur at depths of >250 m. the potash mine uses wells in basinal gravels as a source of potassium-rich brine. the basinal gravel aquifer consists of gravels, conglomerates, and tertiary volcanic rocks (stephens, 1974). water rights reports and data reported in the potash mine reclamation plan (shaw environmental, 2020) of basinal gravel wells show the deep brine aquifer’s potentiometric surface declined ~20 to 30 m between the 1950s and 2010s. results here, chemical results are analyzed in a spatial and temporal context, progressing from west to east, covering the period from 1925 to 2023. datasets with insufficient information to differentiate trends are not discussed. for example, most trace elements had insufficient data to identify spatial or temporal changes. transitional zone west compositionally, many tzw samples differ markedly from other bsf samples (figure 7). they have higher relative proportions of sulfate (so4 2-), alkalinity (as hco3 -), calcium (ca2+), and magnesium (mg2+) than other areas because they have lower concentrations of sodium (na+) and chloride (cl-) (especially in the alluvial-fan aquifer wells). shallow aquifer tzw brines show a clear decrease in magnesium between the 1964–1972 and 1999–2006 periods, with increasing magnesium after the 1999–2006 period (figure s1). additionally, lithium (li+) concentrations are much lower in the tzw than in other areas (figure s2). analysis of individual wells in the tze shallow 12 j.a. bernau, b.b. bowen, e.l. kipnis, and j.c. lerback observations of decadal-scale brine chemistry change at the bonneville salt flats aquifer identified two areas of increasing salinity (figures 8a and b). the first segment occurs near the ow transect (figure 1c). the second segment occurs mid-way along the saline pan, where the inner tzw is wider because of a low-lying silver island mountains pass connecting this area to pilot valley. aggregate analysis of the tzw shallow aquifer (figure 9a) indicates that the outer tzw area has experienced an increase in salinity over time. in contrast, the inner tzw area does not exhibit consistent changes in salinity. the inner tzw aggregate results contrast with those of the individual well analyses, possibly from differences in sampling location across studies. over time, the outer tzw area has had marked declines in groundwater levels; several shallow aquifer wells were persistently dry during the 2013–2023 study period. stable water isotope measurements from the tzw shallow aquifer indicate that it is isotopically lighter (meaning it originates from less evaporated waters or precipitation from cooler periods) relative to other shallow aquifer areas (figures 10 and 11). similar to the shallow aquifer in other areas, a negative shift in tzw shallow aquifer deuterium excess values (dansgaard, 1964) suggests a change to more evaporated waters over time (figure 11). the alluvial-fan aquifer has had notable declines in groundwater levels since the early 1990s (figure 3e) (kipnis and bowen, 2018). in addition to hydraulic head changes, between 1993 and 2022 there have been marked changes in brine density and δ2h and δ18o values. furthermore, the spatial distribution of these values over time (figures 12–15) reflects changing groundwater sourcing from mountain front sourced waters to more evaporated waters from the saline pan area. the bw production well transect shows increasing density and a shift to heavier δ2h and δ18o isotopes over time (figure 12). the largest density increase is concentrated at the center of the active production field (where produced waters now exceed a density of 1.05 g/cm3), with smaller density increases occurring on the edge of the active field (wells at 1.8 and 4.2 km). there are some exceptions to the correlation between increased density and generally heavier water isotopes. waters from a well at 4.2 km had relatively low densities, but heavier isotopic values. this indicates that some waters originate from precipitation under warmer conditions or from evaporated waters (as suggested by water isotopes) that do not have an elevated salinity. observations from the ow well transect, which spans the inner to outer parts of the alluvial-fan aquifer (figure 1c), inform the interpretation of observed changes in bw chemistry (figures 13 to 15). figure 13a to c, a cross section of ow measurements over time shows an increased density gradient towards the basinward direction (to the east) with a transition in salinity between 1 and 2 km. similarly, figures 15a and d show a transitional zone between the alluvialfan and basinal δ2h and δ18o values in the year 1993 occured between 1 and 2 km along the transect. the δ2h, δ18o, and deuterium excess values observed in mountain front-adjacent alluvial-fan samples suggest these waters are sourced from winter precipitation that has undergone minimal evaporation and fractionation (figure 11). the deuterium excess values of all figure 7. piper diagram of geochemistry measurements across bsf at a) normal scale, b) magnified scale. a) the high salinity system dominated by na and cl makes differentiating between sites based on relative ionic content challenging. b) differentiation between sites when examined at the 90/10/10 percentile values. 13 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 a b figure 8. a). select well trends in density through time. samples from east of ditches, tze, and i-80 south areas are all from wells screened in the shallow aquifer. see figure 1b well locations. b) sites from fig. 8a with a statistically significant change in salinity. color is used to denote change in density. size denotes the relative fit of change with time (higher r2 shows a stronger correlation of change with time). all wells in b) are in the shallow aquifer except for the alluvial-fan aquifer well in the southwest corner of the map.  14 j.a. bernau, b.b. bowen, e.l. kipnis, and j.c. lerback observations of decadal-scale brine chemistry change at the bonneville salt flats figure 9. box and whisker plots of changes in (a) shallow aquifer brine density across different areas over time and in (b) brine extraction rates over time. individual well data has been averaged for each sampling period to minimize the effect of sampling bias on results. note that sampling has not been consistent at the same wells over time, making trends identified here different from those identified in figure 8a. extraction rates before 1991 are primarily estimated and have a high range of uncertainty. 15 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 other areas suggest their composition was influenced by evaporation. nevertheless, the absence of a proportional increase in deuterium excess values with increasing brine density implies that the salinity of the brine arises not only from evapoconcentration but also from the dissolution of salt (figure 11d). these data strongly support the theory that high salinity in the halite nucleus is maintained by salt crust dissolution. between 1993 and 2022, the hydraulic gradient shifted to mountain front-directed flow with a 10 m change in the hydraulic head at 0 km in the ow transect (figure 14a to f). hydraulic head was corrected for density to assess the effect of density on groundwater flow by using equation 1 (figure 14d and e) (post and others, 2007). (equation 1) figure 10. spatial changes in water δ2h and δ18o. a) all isotope data. b) average isotope values by area with standard deviation in error bars. tze alluvial outer wells indicate ow wells located >1 km east of transect start. lmwl is the local meteoric water line. lel is the local evaporation line. 16 j.a. bernau, b.b. bowen, e.l. kipnis, and j.c. lerback observations of decadal-scale brine chemistry change at the bonneville salt flats where ρ1 is the reference density to adjust the sample to (freshwater density, 1.0 g/cm3); ρ2 is the density of the well-water (calculated from the average density of water’s produced from a well in a period); h2 is the height of the water level above a datum (mean sea level); z is the elevation (above the sea-level datum) of the mid-point of the well’s screened interval; and h1 is the equivalent head relative to the datum. between the years 1993 and 2022, there was an approximately 12-meter change in hydraulic head at the center of the bw production field. interestingly, the basal head level observed in the ow wells (figure 14b) for 2022 closely matches the head observed at the center of the production field in figure 3e. this similarity in head change, despite a lateral offset of over 4 kilometers, indicates a high level of hydraulic connectivity across the bw alluvial-fan aquifer wells. correcting for the effect of density on hydraulic head has a significant impact on inferred water flow, as shown in figure 14. in 1993, it becomes evident that groundwater primarily flows towards the west when the density correction is applied. without this correction, groundwater flow would have been interpreted as moving towards an elevation of approxifigure 11. comparison of stable isotopes by area and aquifer over time. a) δ2h, b) δ18o, c) deuterium excess, and d) deuterium excess relative to density. 17 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 mately 1240 meters above sea level at 2 kilometers along the transect. in 2022, the density correction has a relatively smaller impact, but it still highlights that groundwater primarily flows in a westward direction, and downward vertical flow is less significant than what the uncorrected head measurement would suggest. changes in ow-brine density indicate that brine transports salt mass from the halite nucleus crust and underlying aquifers and towards the alluvial fan. changes are concentrated on the basinward side of the ow cross section; where salinity has increased by 0.04 g/cm3 at 2.6 km at ~1255 meters above sea level and by ~0.01 g/cm3 at ~1215 meters above sea level. the cross-sectional view of salinity change (figure 13c) shows a saline “nose” with brine migrating down and then west and upwards towards the alluvial fan (possibly along gravel lenses). ow-transect wells have the largest stable-isotope changes of any area in the dataset. the largest change occurs at ~0.8 km at ~1245 meters above sea level (figure 15c and f). this change reflects the movement of basinal water in the direction of the mountain front and is also shown in the 2022 cross-sectional view of isotope data (figure 15b and e), which shows a blurring of the delineation between alluvial and basinal-sourced water that was evident in the past (figure 15a and d). these ow changes suggest that the bw well at 4.2 km has tapped basinal waters but has not yet sourced waters from areas with elevated salinity. halite nucleus figure 5 illustrates declining halite nucleus brine density with depth. crust aquifer brine is halite saturated (>1.195 g/cm3); the shallow aquifer has high salinity (1.175 to >1.195 g/cm3), and the moderate and deep aquifers have lower salinities (1.09 to 1.175 g/ cm3 and 1.08 to 1.11 g/cm3, respectively). mason and others (1995) showed this decline in salinity with depth by measuring pore water chemistry at multiple depths. individual well plots show no change in density over time in the halite nucleus crust samples (figure 5). past research suggested that the crust aquifer potassium concentrations decreased between the 1960s and the 1970s (lines, 1979); subsequent analyses, however, show no long-term change in potassium from the 1960s baseline (mason and kipp, 1998; white, 2002) (figure s3). calcium concentrations decreased between the 1960s (from a high of ~1700 mg/ l) to an observed low in the 1991–1997 period (~1100 mg/l) and then increased (figure s4). there is a notable isotopic lightening in the halite nucleus crust aquifer for δ2h values between the periods of 1991–1997 and 2013–2023. shallow aquifer samples show a similar change in δ2h. this trend is not seen in the δ18o values, which slightly increased in both aquifers. two wells showed decreasing halite nucleus shallow aquifer salinity over time (figure 8). aggregate data show a long-term density decrease with brines becoming halite undersaturated between 1964 and 1997 (figure 9a). afterward, density remained stable and then increased during the 2013–2023 period. spatial differences in sampling location over time may influence this trend. in contrast to density measurements, the halite nucleus shallow aquifer sodium and chloride concentrations appear to consistently decrease over time (from 105 to 90 g/l and from 180 to 150 mg/l, respectively) (figures s5 and s6), highlighting the problematic nature of accurately measuring high-salinity brines (bernau and others, 2023a). calcium concentrations are positively correlated to changes in density over time in the halite nucleus shallow aquifer while sulfate concentrations are negatively correlated to density changes (figures 9, s4, and s7). moderate and deep halite nucleus aquifer wells figure 12. bw well cross section with salinity and water stable isotopes over time. location of a to a’ is shown on figure 1c. 18 j.a. bernau, b.b. bowen, e.l. kipnis, and j.c. lerback observations of decadal-scale brine chemistry change at the bonneville salt flats figure 13. ow cross section density. density in (a) 1993, (b) 2022, and (c) change in density between 1993 and 2022. each blue dot represents the midpoint of a well’s screened interval. the location of b to b’ is shown in figure 1c. (~40x vertical exaggeration). elevation is meters above sea level. local surface elevation (not shown) is 1284.9 to 1293.0 m. note that a & b use the same color scales with only the applicable range for each segment being shown. 19 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 have a limited sampling history over a small spatial extent, and no unequivocal changes in their composition were observed. the moderate depth aquifer had the highest observed sulfate concentrations of any area and aquifer (figure s7). transitional zone east two wells showed clear decreases in tze shallow aquifer salinity over time (figure 8). aggregate tze shallow aquifer measurements show salinity decreased through the 1991–1997 period and increased afterward. in contrast to density (and similar to the halite nucleus shallow aquifer), reported sodium and chloride concentrations have decreased over time. magnesium concentrations at tze reflect observed density changes. similarly, tze shallow aquifer potassium concentrations decreased up to the 1991– figure 14. ow cross section of hydraulic head over time. (a to c) measured hydraulic head measurements and (d to f) hydraulic head measurements corrected for density. note the impact of density correction on apparent flow direction (red arrows). note that a & b, and d & e use the same color scales, respectively, with only the applicable range for each segment being shown. each blue dot represents the midpoint of a well’s screened interval. the location of b to b’ is shown in figure 1c. (~40x vertical exaggeration). local surface elevation (not shown) is 1284.9 to 1293.0 m. elevation is meters above sea level. 20 j.a. bernau, b.b. bowen, e.l. kipnis, and j.c. lerback observations of decadal-scale brine chemistry change at the bonneville salt flats 1997 period; there were no changes in the following period. measured δ2h became lighter between the 1991–1997 and 2013–2023 periods (figure 11a). extracted brines here, changes in brine extracted from the eastern collection ditches are described. nacl removal from subsurface flow to the south (mason and kipp, 1998) is not considered in these values. sustained brine production at bsf began in 1939 (bingham, 1980); since then, the location and volume of brine extraction for potash production at bsf have varied over time (figure 9b). before 1966, the salduro loop ditch at the center of bsf (figure 1b) was used to harvest brine; this ditch was likely constructed after sustained production began in 1939, and it is present in 1946 figure 15. ow cross section of water stable isotopes (δ2h: a to c, and δ18o: d to f) over time. a & b and d & e use the same color scales, respectively, with only the applicable range for each segment being shown. each blue dot represents the midpoint of a well’s screened interval. the location of b to b’ is shown in figure 1c. note the vertical scale differs from figures 13 and 14. (~40x vertical exaggeration). local surface elevation (not shown) is 1284.9 to 1293.0 m. elevation is meters above sea level. 21 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 aerial imagery. hadzeriga (1964) reports an initial 1% kcl and 21% nacl weight content of brines used for potash production. the mass of potassium produced annually from the shallow aquifer since 1968 is reported in the potash mine’s estimated resources and reserves report (agapito, 2022). production from bsf area and the area south of i-80 was lumped together, as such these values can only be used to provide an estimate of brine extraction from bsf over time. the eastern ditches at bsf have been pumped intermittently since 1963 (lines, 1979). an estimated 0.26 mt nacl/year was extracted from the eastern ditches between 1966 and 1972 (stephens, 1974) (figure 9b). lines (1979) reports extraction of >0.25 mt nacl/ year in 1976. mason and kipp (1998) estimated 0.47 mt nacl/year extraction. white (2004) reports that brine extraction between 1995 and 1998 was between 0.45 and 0.94 mt nacl/year. brine extraction rates between 2001 and 2021 were highly variable (todd marks, bureau of land management, written communication, 2023), ranging from almost none to nearly 0.8 mt of nacl extracted/year (figure 3a). there was over twice as much annual average extracted mass in the 1996–2006 period (0.40 mt/year) than in the 2013 –2023 period (0.16 mt/year). during at least the last part of the 1996–2023 period, the northern part of the extraction ditches was inactive (figure 1b; potash mine personnel, verbal communication, 2022). examination of the ratio of total kcl produced from the shallow aquifer (bsf and area south of i-80) and nacl extracted from bsf over time shows that production became less reliant on bsf after the year 2000; with an average of 10 tonnes of nacl extracted from bsf for every ton of kcl produced by the mine before 2000, and 5.5 tonnes of nacl extracted from bsf for every ton of kcl produced afterward. while seasonal precipitation and evaporation changes can impact the salinity of extracted brine (figure 6), month-to-month analysis of density and extracted brine volumes show a clear decrease in extracted brine salinity with increased extraction rates (figure 3d). the decrease in salinity with increased extraction suggests that deeper, less-saline waters rise upward (with reduced hydrostatic pressure) and contribute to the shallow aquifer. from 2000 to 2004, the density of the extracted brine increased while the volume of brine extracted remained constant. from 2004 to 2015 the extracted brine density decreased; from 2004 to 2010 there was no net addition of nacl to the salt flats. following 2010, brine extraction greatly decreased, while laydown increased, leading to increased net volumes of brine contributed to bsf. apparently, a 5-year lag occurred between decreased extraction with increased net brine contributions and the onset of salinity increase within the system. this is a much longer period than the salinity recovery period in the early 2000s, which was associated with a much larger pulse of input solutes over a shorter period. following increasing density up to 2015, density remained high during the 2016–2020 period. east of ditches and i-80 south both the east of ditches and i-80 south areas have reduced hydraulic connection with the halite nucleus and are isolated from laydown brines. furthermore, brine extraction ditches impact both of these areas. accordingly, these areas provide an example of how the brine system may respond to extraction on decadal timescales without external solute sources. in contrast to the halite nucleus and tze shallow aquifers, brine samples from the east of ditches and i80 south shallow aquifers show decreases in density across individual wells and in aggregate (figures 8a and 9). similar to the halite nucleus crust aquifer and tze shallow aquifer, calcium concentrations in the east of ditches shallow aquifer decreased until the 1991–1997 period (from ~1500 to 1100 mg/l) and increased afterward (to ~1400 mg/l) (figure s4). discussion brine chemistry changes over time bsf brine chemistry changes lie within three groups: 1) no change, 2) long-term decrease or increase, and 3) change in long-term trend following the 1991–1997 period. no long-term changes in the deep and moderate-depth aquifers underlying the halite nucleus and tze were observed. similarly, the inner tzw area, when taken in aggregate, did not show any change in density over time. however, analyses of several wells from the tzw area indicate parts of this area increased in density over time. the shallow aquifer in the east of ditches and i-80 south areas show clear decreases in density over time, in contrast, the outer tzw area is the only shallow brine area to show long-term increases in density across several studies. the bw production well area in the tzw also shows consistent long-term increases in alluvial-fan aquifer density, but the east of ditches and i-80 south areas show long-term decreases in brine density, with the i-80 south area showing the largest salinity decrease. the changes in brine salinity in the i-80 south area also suggest that there is limited transport of solutes under i-80 from bsf to the i-80 south area. high connectivity between these areas would likely limit salinity decreases in the i-80 south 22 j.a. bernau, b.b. bowen, e.l. kipnis, and j.c. lerback observations of decadal-scale brine chemistry change at the bonneville salt flats area as transported salt would replenish removed solutes. the last group of chemical change, with changes in long-term trends after the 1991–1997 period, indicates a change in saline pan conditions. long-term density decreases ceased and salinity increased after the 1991–1997 period in the halite nucleus and tze shallow aquifer (figure 9a). on smaller timescales, brine extracted for potash production shows an increase in salinity following the onset of the laydown in 1998. later, an increase in salinity occurred when brine extraction greatly decreased as the laydown continued (figure 3a to c). the onset of the laydown coincides with a marked decrease in alluvial-aquifer groundwater levels. declining levels, in turn, reversed hydraulic gradients, enabling basinal brine movement away from the saline pan (as seen in density and isotopic changes) (figure 16). careful consideration of the underlying forces driving changes in brine chemistry, primarily density and water stable isotopes (e.g., laydown, decrease in brine extraction, or long-term groundwater consumption trends), is needed to identify the core controls on change that will influence management decisions. potential controls on brine chemistry change laydown a cessation in long-term density decreases and an increase in density in the halite nucleus and tze shallow aquifers that is concurrent with the laydown was observed. similarly, extracted brines for potash production showed increased density in the 3to 5-year period after the laydown began. these observations support white’s (2004) hypothesis that the laydown failed to lead to a 4–5 cm increase in halite thickness because it buffered salinity decreases in the shallow brine aquifer. although increases in aquifer density are concurrent with the onset of the laydown, changing potash brine extraction rates may have had a larger impact on observed changes in groundwater density than the laydown. an δ2h isotopic lightening in both halite nucleus crust brines and shallow aquifer waters over time was observed; in the absence of other information, this lightening could be attributed to isotopically lighter laydown waters. however, isotopic measurements from the east of ditches area (which is isolated from figure 16. conceptual model for pre-alteration and modern groundwater (and brine) movement at bsf shown in cross section (left) and as a simplified conceptual model (right). 23 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 laydown waters) show the same lightening over time. therefore, a different mechanism, possibly increased infiltration of meteoric precipitation, may explain observed changes. changing brine extraction rates bsf brine extraction volumes before 1995 are poorly reported, leading to high uncertainty in estimated extraction volumes (figure 9b). the average extraction volumes during the 1964–1972 and 1975– 1981 periods may have been similar to or higher than those reported between 1995 and 1998. measurements from 1999 to 2023 show stabilizing to increasing density in the halite nucleus and tze shallow aquifers. there is a corresponding decrease in brine extraction during this period, with the largest increase in the halite nucleus shallow aquifer salinity corresponding to the greatest decrease in extraction rates. this relationship suggests that reduced extraction rates may contribute to some (and possibly most) of the observed density increases in these areas. examination of increasing sulfate concentrations over time relative to decreasing density in the halite nucleus and i-80 south areas shallow aquifers indicates these two values are related (figures 9 and s7). this relationship suggests a mechanism for brine replacement after extraction. the moderate depth aquifer has elevated sulfate concentrations. when groundwater is extracted for mining, it lowers the constraining hydrostatic pressure, enabling deeper, less saline groundwater with higher sulfate concentrations to rise and replace extracted waters. there is a positive correlation between extracted brine density and potash production brine extraction volumes (figure 3d). using that relationship, anticipated changes in density based on extracted brine volumes over time were modeled. while not reflecting observed density measurements (modeled density of 1.10 to 1.15 g/cm3 between 2002 and 2010, when observed density was ~1.17 to 1.21 g/cm3), modeled density did replicate trends in brine density change for some periods. the model replicates observed trends between 2002 and 2006 where it shows increasing and then decreasing density; it also shows increasing and then stabilized density between 2015 and 2021. modeled and observed trends strongly differ between 2006 and 2015, where the model shows generally increasing density, while observations show generally decreasing density. these differences between observations and modeled changes suggest 1) extracted brine density may respond non-linearly to extraction rates (and accordingly, there is a lower limit for extracted brine salinity, potentially buffered from the dissolution of the halite crust); and possibly 2) under normal operating conditions (with some extraction and laydown volume of ~0.5 mt nacl/year) extracted brine salinity (and aquifer salinity) will decrease; additionally, 3) the large initial laydown pulse between 1998 and 2000 increased the salinity of extracted brines beyond their anticipated baseline. brine extraction can also depress local groundwater levels (turk, 1973). these declines may enable infiltration of surface waters into the subsurface before they can evaporate. isotopic lightening of the shallow aquifer (figure 11a) may be attributed to the incorporation of isotopically lighter winter precipitation into the aquifer before significant evaporation occurred. given these observations, especially that changes in salinity can be attributed to extraction rates, the relative role of the laydown in increasing tze and the halite nucleus shallow aquifer salinity remains unclear. it may be that a solute source (such as a halite crust or laydown brines) is necessary for density values to recover. recent work on the sedimentology of bsf salt crusts documented extensive evidence of halite dissolution, suggesting that the diminishing crust is a likely source of these solutes (bernau and bowen, 2021). declining tzw groundwater levels two areas of increasing salinity were identified in the inner tzw shallow aquifer, and a long-term aggregate trend in increasing salinity in the outer tzw shallow aquifer was also found. kipnis (2021) and lines (1979) noted decreasing groundwater levels in these areas. furthermore, several dry outer tzw shallow wells were noted during this investigation. two mechanisms for increasing tzw shallow aquifer salinity are proposed: 1) the tzw inner area receives brine from the halite nucleus aquifers and surface precipitation, and 2) salinity that was once concentrated at the surface by efflorescence and then recycled basinward now accumulates in groundwater as the capillary fringe falls below the ground surface. flow of halite nucleus brine to the inner tzw area may be enhanced by declining tzw hydraulic head. furthermore, there is the potential for laydown brines (which accumulate on the western halite nucleus edge) to enter the tzw area. several observations, such as declining western halite nucleus shallow aquifer salinity and areas of inner tzw with increasing salinity, support the interpretation that westward movement of shallow aquifer brine contributes to recent declines in bsf crust thickness. between the 2003 and 2016 salt crust thickness studies, the area on the southwestern side of bsf had some of the largest observed decreases in salt crust thickness. this area is also the closest part 24 j.a. bernau, b.b. bowen, e.l. kipnis, and j.c. lerback observations of decadal-scale brine chemistry change at the bonneville salt flats of bsf to the alluvial-fan production wells. during the same period, the northernmost parts of the bsf crust reduced in volume; those declines may be attributable to their distance from the thickest part of the halite crust. the alluvial-fan aquifer shows several significant changes over time. the decline in groundwater levels has reversed the hydrological gradient such that basinal waters now flow toward the mountain front. the movement of basinal waters toward the mountain front is demarcated by changes in brine density and δ2h and δ18o isotopes. formerly fresh production wells now produce waters that exceed the salinity of the ocean (~35 ppt or 1.03 g/cm3). these changes indicate that basinal brine (possibly the halite nucleus shallow and crust aquifer brine) is being removed from the saline pan area (figure 16). isotopic and density data should be considered to determine the relative sourcing of waters extracted by production wells to estimate how much brine will be removed from the halite nucleus area under different alluvialfan extraction scenarios. climate does not explain observed alluvialfan groundwater level declines between 1993 and 2010, there were frequent drought periods, with the majority of years experiencing precipitation levels below the 25% quartile for precipitation based on data from 1910 to 2020 (bernau, 2022). before the 1990s, alluvial-fan aquifer levels varied but regularly returned to the land surface. in the period since then, they have shown a clear long-term decline that strongly differs from observations of similar alluvial-fan aquifer wells in the gsld, which have remained relatively stable to slightly increasing over this period (nwis, https:// maps.waterdata.usgs.gov/mapper/index.html, sites 404757112582701 and 394905113354101). these data indicate that recent declines in the alluvial-fan aquifer are occurring because extraction rates have exceeded recharge rates for over two decades. impact of laydown on alluvial-fan aquifer drawdown before 1997, the alluvial-fan groundwater level remained within 0 to 7 m of the surface (kipnis and bowen, 2018; mason and kipp, 1998). following the laydown, groundwater levels never rose above 10 m below ground level and have continued to decline (figure 3e). this suggests that groundwater levels have not yet reached a new equilibrium where inflow is equal to pumping rates and a larger area will be drained over time. the relative proportion of basinal water in produced alluvial aquifer waters will increase as this area expands. the laydown has led to as much as a doubling in alluvial groundwater extraction, and as such, plays a major role in decreasing alluvial-fan groundwater levels. conclusions new chemical and groundwater level measurements and past research were compiled to examine multi-decadal changes at the bonneville salt flats. brine chemistry, most notably density and δ2h and δ18o water isotopes, has changed in response to anthropogenic activities (figure 16). shallow aquifer brine under and to the east of the crust declined in salinity between 1964 and 1997 and stabilized and increased in salinity afterward. increased salinity may be due to decreased extraction rates in the past two decades, especially as the largest increase in salinity, during the 2013–2023 period, is concurrent with the largest decrease in extraction. however, this period is also concurrent with the experimental salt restoration laydown program. the relative role of the laydown in increasing aquifer salinity remains unclear. alluvial aquifer groundwater levels have declined over time. this decline is linked to industrial water production, including the laydown. as a result, the hydraulic gradient has reversed, causing brine to flow away from the saline pan and towards the alluvial aquifer. this flow increases alluvial fan aquifer salinity and changes its isotopic composition. if alluvial-fan extraction rates remain the same, or if they rise with increases to the laydown, more salt will be removed from the bonneville salt flats halite nucleus, potentially at volumes exceeding the laydown. these multi-decadal chemical changes inform the understanding of groundwater movement and halite crust changes in this system, which informs management for the sustained use of this landscape. supplemental data supplemental data and figures are available at https://doi.org/10.5281/zenodo.8152647. acknowledgments the authors acknowledge that this study was conducted on traditionally newe/western shoshone and goshute lands. the authors would like to acknowledge craig peterson and russ draper with intrepid potash; todd marks, roxanne tea, steve allen and other past and current u.s. bureau of land 25 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 management west desert field office staff; hannah stinson, lily wetterlin, olivia watkins, candace penrod, gabë regenhardt, and dr. anna cassell from the university of utah; and paul inkenbrandt, hugh hurlow, and greg gavin of the utah geological survey for assistance with data collection and sample analysis. this manuscript was improved through reviews and discussion with greg carling, stephanie carney, scott hynek, michael hylland, and elliot jagniecki. torrie duncan and john good contributed to refining some of the figures in this report. cryogenic vacuum extraction of brine samples was made possible by suvankar chakraborty, dr. jim ehleringer, and dr. gabe bowen at the university of utah sirfer lab. bill white provided unpublished field notes contributing to this work. funding for this work was provided by an nsf coupled natural human systems award #1617473, the united states geological survey, the utah geological survey, the utah state legislature, the bureau of land management, and the university of utah global change and sustainability center graduate research grants. references agapito associates, inc., 2022, technical report summary of 2021 estimated resources and reserves at intrepid potash-wendover, 262 p. bernau, j.a., 2022, spatial and temporal scales of water and salt movement at the bonneville salt flats: salt lake city, university of utah, ph.d. dissertation, 195 p. bernau, j.a., and bowen, b.b., 2021, depositional and 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description of factorial models for analysis of variance: journal of the royal statistical society, series c, v. 22, no. 3, p. 392–399. downardgslwetlandvegn.pub 1 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2023 utah geological association publication 51 abstract great salt lake (gsl) wetlands support more than 300 species of migratory birds and provide many ecosystem functions, including flood and drought attenuation, dust mitigation, and water quality improvement. wetland vegetation is a key factor in providing those functions and can also tell us about how healthy a wetland is. from 2013 to 2022, 135  gsl wetlands were surveyed to develop a multi-metric index of gsl wetland condition. that wetland condition data, along with environmental variables like soil and water chemistry and physical disturbance, are summarized here as 1) an ecological characterization of the three main types of gsl wetlands, 2) a description of how the plant community differs across environmental and anthropogenic disturbance gradients, and 3) assessment of the major risks to gsl wetland health. gsl wetland plant species are generally resistant to environmental disturbance because of the anatomical and physical adaptations that allow them to survive in dynamic wetland environments. however, land use conversion and the rapid expansion of invasive species, the major threats to gsl wetland health, have seriously degraded wetland condition around gsl. in addition to being useful in wetland monitoring and assessment, the results presented here can also identify wetlands in need of enhanced protection or those with restoration potential as well as setting realistic wetland restoration goals for the region. great salt lake wetland vegetation and what it tells us about environmental gradients and disturbance becka downard utah geological survey, salt lake city, utah, beckad@utah.gov introduction: the great salt lake ecosystem great salt lake (gsl) and its surrounding wetlands are often described in superlative terms: great, immense, critical, and essential. gsl is the largest saline lake in north america and eighth-largest in the world. more than 1,500 square kilometers of wetlands thrive on the margins of gsl where freshwater flows toward the lake (figure 1). millions of birds representing 338 species rely on gsl wetlands to power their migrations across the western hemisphere (sorenson and others, 2020). studying the plant community that thrives in gsl wetlands highlights significant features of natural history, the impacts of wetland management and human disturbances the gsl ecosystem experiences, and how to best protect and restore the wetlands in the future. gsl natural history gsl is all that remains of historical lake bonneville which occupied much of northern and western utah 15,000 years ago but shrank as the regional climate became much drier (inkenbrandt, 2021). gsl is a terminal lake with no surface water outlets, water only leaves through evaporation. the solutes rivers bring to gsl have concentrated over time and currently the lake is more than three times saltier than ocean water, ranging in salinity between 125 and 185 g/l (u.s. geological survey, 2023). the bear, weber, and jordan rivers provide approximately 90% of the water to gsl. the gsl watershed occupies a total of 91,908 square kilometers, an immense area within which changes in climate, water availability, and water quality can impact the gsl ecosystem (zedler and kercher, 2004; ramsey and others, 2009). the rivers supplying gsl terminate in massive deltas composed of diverse wetland types, from sparsely vegetated saline playas to freshwater marshes and ponds. wetlands are defined by three characteristics: the presence of water for part of the year, soils with low oxygen (hydric soils), and plants adapted to flooding and low oxygen (mitsch and gosselink, 2015). within that definition, a variety of environmental conditions create diverse wetland types with their own suite of ecosystem functions, from water quality improvement to hydrologic and climate regulation (wetzel, 2006). the diversity of wetland types in gsl river deltas as well as their expansive size allows the ecosystem to support many species of birds, from tiny snowy plovers to massive american white pelicans (aldrich and paul, 2002). though gsl wetlands are a reliable place for migratory birds to feed and nest, they are hardly static. wetlands are dynamic habitats, shifting between 10.31711/ugap.v51i.140 2 b. downard great salt lake wetland vegetation and what it tells us about environmental gradients, and disturbance figure 1. great salt lake ecosystem includes the lake and 1,500 sq. km of wetlands. map by grant mauk. wetlands layer (u.s. fish and wildlife service, 2018), rivers and lakes layer (u.s. geological survey, 2020), digital elevation model (quantum spatial, inc, 2017 ) 3 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 flooded and dry over the growing season and bridging the transition between aquatic and terrestrial environments; gsl wetlands are especially dynamic. terminal lakes fluctuate in area much more than other lakes and this has big implications for gsl wetlands (u.s. geological survey, 2023). in high water conditions, the hypersaline waters of gsl can rise to inundate the wetlands, but as the lake retreats during drought, wetlands occupy the lakebed. these changes in area are significant; by one estimate, 180 square kilometers of lakebed are exposed for every foot in elevation that gsl falls (aldrich and paul, 2002). within the wetland complexes, changes in water availability shift the boundaries between terrestrial, wetland, and fully aquatic environments. gsl human history peoples of the ute, paiute, goshute, and shoshone nations utilized gsl wetlands for centuries, but european settlers have left the most distinct marks on the system (madsen, 2015). when john c. fremont saw the bear river delta in 1843, he described the sound of birds taking off as having “wings of thunder” because the birds were so numerous. european settlers arrived in the salt lake valley in 1847 and immediately began diverting tributaries of the jordan river to support agriculture. the transcontinental railroad was completed at the northern end of gsl in 1869 bringing industry and transportation of agricultural goods (baxter and butler, 2020). by the 1920s the bear river delta had been dewatered so severely that avian botulism was leading to massive bird dieoffs in the few locations migratory birds found habitat (wilson and carson, 1950). local communities pressed congress for the establishment of a federal wildlife refuge in the bear river delta and the first act of refuge building was the construction of a series of dikes to hold water in the river delta when it was available in the spring and manage drawdown more slowly during the irrigation season (downard and endter-wada, 2013). this intense impounded wetland water management practice was successful in preserving migratory bird habitat and has been adopted by state waterfowl management areas, private hunting clubs, and conservation areas (figure 1) (downard and others, 2014). according to both researchers and stakeholders, upstream consumptive water use and subsequent drought downstream is the primary threat to gsl wetlands and the lake itself (wurtsbaugh  and others, 2017; utah division of water quality, 2019). in the last century, the elevation of gsl has fallen approximately 11 feet due to diversion of surface water for human needs (wurtsbaugh and others, 2016). in october 2022, gsl fell to a record low elevation of 4,188.7 feet which exposed thousands of square kilometers of lakebed (u.s. geological survey, 2023). water quality threats, most notably legacy phosphorus bound to soils, also impact the gsl ecosystem and become more problematic as water availability decreases (utah division of water quality, 2014). invasive species, especially phragmites australis, complicate the water situation even further by changing how water flows across the very flat landscape and altering nutrient cycles in wetlands (kettenring and others, 2020). gsl wetlands ecology and ecosystem services the path surface water follows through gsl wetlands from river to lake is a complex mix of deliberate management actions and unintended consequences of upstream water diversions and nearby water discharges. gsl wetlands are divided into three classes—impounded, fringe, and playa wetlands— that shift in area according to where water is available and how long and deep flooding is. impounded wetlands are the most deeply flooded wetland class and are flooded for the longest part of the year. fringe wetlands may be flooded nearly as deeply as impounded wetlands, but water depth often fluctuates between flooded and dry stages over the growing season. playa wetlands are often not flooded, but saturated. the relatively permanent flow of water into impounded and fringe wetlands keeps them fresh to brackish, especially compared to the saline waters of playa wetlands and gsl. wetland vegetation is both a defining feature of wetlands and an indicator of the ecosystem functions wetlands perform and integrates the environmental stresses and anthropogenic disturbances a wetland faces over time (moor and others, 2017). differences in the growth form, life cycle, wetland indicator status, and habitat specificity of plant species present in wetlands vary over gradients of water regime, management history, and disturbance (lytle and poff, 2004). wetlands present a suite of challenges to plant life and wetland species have a number of common adaptations that allow them to grow and reproduce. a wetland plant in this region must deal with unpredictable water regimes, soil anoxia when water is present and drought stress when water is absent, a range of salinities, and periodic catastrophic flooding. wetland environmental gradients, especially water depth and salinity, act like a sieve, filtering the species that can occupy that space (van der valk, 1981). water regime—the pattern of flooding and drying in a wetland— is largely considered the most im4 b. downard great salt lake wetland vegetation and what it tells us about environmental gradients, and disturbance portant factor in determining the wetland plant community (mitsch and gosselink, 2015). in wetlands with relatively permanent flooding, perennial species with specialized adaptations to flooding like aerenchyma and floating seeds are dominant (cronk and fennessy, 2001). wetlands that fluctuate between flooded and dry states more frequently (i.e., those with more seasonal hydroperiods) have a unique suite of species as well, often rapidly growing species with dense networks of rhizomes that allow clonal species to share gases when wetlands are flooded and water when wetlands are dry (cronk and fennessy, 2001). temporarily or ephemerally flooded wetlands in turn tend to have communities dominated by annual species, those that can complete their lifecycle in a single growing season if conditions are right (keddy, 2010). in addition to the broad life history traits outlined above, botanists can also characterize how specific a species’ ecological requirements are (i.e., how conservative their habitat is) and the complementary measure of how tolerant it is to ecological or anthropogenic disturbance. highly conservative species with a coefficient of conservatism (cc score) of 10 are only found in a specific type of habitat and are sensitive to disturbance (lopez and fennessy, 2002). species that occupy a wider range of habitat types and tolerate more disturbance have lower cc scores. the most successful and widespread invasive species tend to be disturbance specialists—species that can exploit a disturbance that leaves exposed soils and elevated water nutrients—and have a default cc score of zero (hazelton and others, 2014). wetland condition is analogous to ecosystem health or biological integrity and is most often measured by the plant community because the species occupying a wetland integrate multiple impacts over time. ecologically, wetland condition is the ability of a plant community to maintain its structure and function, compared to wetlands in undisturbed locations. a wetland in good condition looks and functions similarly to pristine wetlands, whereas wetlands in poor condition have experienced enough disturbance that they no longer support the same plant community or ecosystem functions (davies and jackson, 2006). unlike birds or macroinvertebrates, plants cannot migrate when conditions get tough. some plant species can abide in places with high levels of disturbance where other species will be eliminated, and a multimetric index (mmi) captures the ways disturbance tolerators or more sensitive species shape the wetland community (magee and others, 2019). an mmi is a combination of multiple variables describing some aspect of the plant community that changes with increasing anthropogenic disturbance (i.e., it measures the overall health of a wetland). a discussion of stress and disturbance terminology is merited before jumping into the methods and results. stress, natural disturbance, and anthropogenic disturbance have similar effects on the wetland plant community but differ in origin and the time scale they operate at. in this paper, stress is a factor that limits a plant’s ability to grow and reproduce, like living in an environment with limited oxygen, extreme temperatures, or low nutrient availability (grime, 1989). stress is a relatively constant feature of the environment, while disturbance is more episodic (borics and others, 2013). flooding and drought, fire, herbivore grazing, and plant species invasions are common natural disturbances in wetlands that can alter the plant community (cronk and fennessy, 2001). anthropogenic disturbances include converting land uses from natural types to developed sites, diverting water from streams or adding points of discharge with water quality contaminants (miller and wardrop, 2006). though it is possible to define those three terms separately on paper, it is difficult to distinguish between the three in the wetlands because anthropogenic disturbances like water diversion and climate change can increase the frequency of natural disturbances and lead to long-term stress. further, plant communities respond similarly to stress and both types of disturbance, often becoming less diverse and dominated by fast-growing species (cronk and fennessy, 2001). this paper focuses on plant community adaptations to environmental stresses of the dominant species in each wetland class as well as the overall wetland response to anthropogenic disturbances. impounded wetlands form the heart of managed wetland complexes where dikes impound the terminus of a river or stream. impounded wetlands are the only gsl wetland class that has firm boundaries because they are defined by the presence of dikes or berms that are designed to increase the depth and length of time this wetland class is flooded. water depth is managed throughout the year with headgates. the primary goal of impounded wetland management is to grow submerged aquatic vegetation (sav) that supports migrating waterfowl (ducks, geese, and swans) (figure 2a), though emergent vegetation is also a component of the impounded wetland community (aldrich and paul, 2002). fringe wetlands are defined by emergent vegetation that forms deltas where water sources like streams, springs, and impounded wetland water control structures discharge onto the bed of gsl (figure 2b) (utah division of water quality, 2016). the mix of short and tall emergent species provides critical nesting habitat for waterfowl and ample food for waterbirds like white-faced ibis and egrets. the extent of fringe wetlands changes based on freshwater availa5 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 2. characteristic examples of a) impounded, b) fringe, and c) playa wetlands near gsl. a b c 6 b. downard great salt lake wetland vegetation and what it tells us about environmental gradients, and disturbance bility, expanding where water is perennial and contracting when water is diverted elsewhere. fringe wetlands are located outside the boundaries of impoundments and are commonly referred to as marshes and meadows. playa wetlands are ephemerally flooded or saturated, sparsely vegetated, often saline wetlands that support astounding populations of shorebirds, who probe the soils for macroinvertebrates (figure 2c). this class contains two types of features, playas and mudflats. playas are a geological feature that form in depressions often supported by shallow groundwater or precipitation (oviatt, 2014). mudflats are the exposed surfaces of drying lakes and wetlands. though the processes that form playas and mudflats are different, they support the same vegetation communities and will be considered together here. as gsl has retreated over the last decade, playa wetlands have expanded to occupy the exposed lakebed. depending on gsl elevation, playa wetlands account for as little as 40% or as much as 85% of the wetland acreage around gsl (u.s. fish and wildlife service, 2018). gsl wetland survey methods and analysis the wetlands around gsl are a critical resource for utahans and of great interest to many stakeholders, including the state agencies that pursued the projects described below. the data presented here are the result of more than ten years of vegetation monitoring in gsl wetlands with the overall objective of developing an mmi to measure wetland condition specific to this region. altogether we have detailed vegetation data from five separate surveys that sampled 135 wetlands from all three gsl wetland classes. a summary of the site selection, field methods, and data analysis are presented below with citations to the supporting field protocols and detailed analysis documentation. site selection survey sites were primarily selected via generalized random tessellation stratified (grts) samples. grts sample design creates spatially balanced samples that can be stratified by factors of interest and include factors that create unequal probabilities (stevens and olsen, 2004; kincaid and others, 2019). site selection for surveys conducted in 2019–2022 built on prior work with one key update: wetlands remained in the sample regardless of whether they had surface water during the time of sampling, in contrast to earlier surveys that required the presence of surface water to sample. forty impounded wetlands were surveyed in 2019, adapting protocols established in 2012. the grts samples were stratified so that an equal proportion of sites were drawn from the major watersheds of gsl (bear, weber, and jordan) and an equal proportion of each size class (small, medium, and large) was represented (utah division of water quality, 2020) (figure 3). the first fringe wetland surveys were conducted in 2013 and 2015 and gathered vegetation data from a targeted selection of sites. rather than a random sample, project leaders selected sites they believed would represent the best and worst condition wetlands to capture the full range of condition possible (utah division of water quality, 2016). in 2020, 15 sites from a grts sample with no stratification were assessed to bring the collective number of fringe sites surveyed to 50. finally, a 50-site grts sample of playa wetlands was surveyed in 2022. the playa sample was stratified by wetland system (palustrine or lacustrine) and an unequal probability factor was added to select more sites from huc12 watersheds with higher percentage of riverine wetlands (utah division of water quality, 2022). field methods for all surveys, data were collected from 100 meter transects, though the placement and segmentation of those transects was adapted for each wetland class to capture the most representative vegetation (utah division of water quality 2020, 2022). vegetation data was central to the analysis of each project, so each site visit was conducted during the index period that began on july 1 and ended on september 30, which captures the most representative and reliably identifiable vegetation. impounded wetlands were visited twice during the survey, once during the early summer and later in the season. for all surveys, the identity and absolute cover of each species present along the 100-m transect was recorded as well as cover of bare ground, open water, and filamentous algae. water (surface water or pore water) and composite soil chemistry as well as on-site disturbance data were gathered in addition to vegetation data. observations of physical disturbance within a 100-meter buffer surrounding the center point of each transect were recorded as well. further details of laboratory analyses, data quality control, and individual project objectives are elaborated on in each survey’s sampling and analysis plan (utah division of water quality 2020, 2022). landscape disturbance data were gathered after field work from statewide geospatial layers. small (100-meter) and large (1-kilometer) buffers were added to the center point of each wetland sampled and the prevalence of the following features were calcu7 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 3. location of wetlands surveyed for this ecological characterization. map by grant mauk. wetlands layer (u.s. fish and wildlife service, 2018), rivers and lakes layer (u.s. geological survey, 2020), digital elevation model (quantum spatial, inc, 2017). 8 b. downard great salt lake wetland vegetation and what it tells us about environmental gradients, and disturbance lated within those buffers: 1) agricultural and developed land uses; 2) impervious surface; 3) length of roadways; 4) water right points of diversion; 5) permitted point source and stormwater discharges; and 6) mineral mines and oil and gas wells (u.s. geological survey, 2019a and 2019b; utah geospatial resource center, 2020a, 2020b, 2020c, and 2020d). analysis data analysis occurred in three stages. first, a cumulative disturbance score was calculated for all gsl wetlands, which was in turn used to define reference condition for each wetland class. second, a large group of vegetation metrics were calculated and then screened for their utility in measuring condition and built into mmi’s. the third stage used the disturbance and condition indices to estimate the influence of individual anthropogenic disturbances on wetland condition. see downard (2021) for further details of the analyses. the disturbance score, modeled on the anthropogenic stress indices developed for the national wetland condition assessment (nwca) is a cumulative measure of disturbance a wetland experiences based on nine measures (lomnicky and others, 2019). the first four measures quantify land use impacts within the large one-kilometer site buffers: agricultural and developed land uses, extractive industry claims, and hydrologic modifications (impervious surface, roadways, diversions, and discharges). the fifth human disturbance metric is a standardized summary of the four large buffer metrics. two disturbance measures are captured within the small 100meter buffer: hydrologic modifications and vegetation removal by cattle grazing and herbicide use. the final two disturbance metrics that form the overall disturbance score were recorded from site visits—the number of heavy metals in soils that exceeded background concentrations and the relative cover of introduced species. soil metal background concentrations were established specifically for gsl wetlands. metals and metalloids selected for inclusion followed the recommendations of nahlik and others (2019) and used a regression approach developed by alfaro and others (2015). defining reference condition, the baseline against which wetland condition is compared to, is a critical step in any condition assessment. the simplest definition of reference condition is pristine, the state of a wetland that is not impacted by human activities (stoddard and others, 2006). wetland condition then measures how different a wetland is from reference (davies and jackson, 2006). however, un-impacted wetlands are nearly impossible to find, given the widespread nature of anthropogenic disturbance. instead gsl wetland reference condition was defined as least disturbed condition (ldc): the best available condition of wetlands is assumed to be those wetlands with the least amount of disturbance, accepting that human disturbance has impacted all wetlands to some degree. defining ldc for each wetland class was an iterative process of determining the threshold of each type of disturbance included in the disturbance score that separated ldc from the more disturbed wetlands, following the lead of herlihy and others (2019a). choosing reference condition based on distributional approaches, as done here, is common and controversial. assumptions about the impacts of disturbance, outliers and skewed data, and lack of minimally disturbed conditions can distort the results, thus the discussion of condition and risk should be interpreted with that knowledge in mind (reynoldson and others, 1997). to build an mmi of condition we calculated 211 potential vegetation metrics that captured some aspect of the plant community which were in turn sieved through a series of screens to test for applicability as a measure of wetland condition. each vegetation metric fell into one of six categories: taxa composition, life history traits, hydrophytic status, sensitivity or tolerance to disturbance, vegetation structure, and floristic quality (table 1). the plants database lists the status of all plant species as native or introduced, life history and growth form traits, and their wetland indicator status (u.s. department of agriculture natural resource conservation service, 2020). sensitivity and floristic quality measures of each species were retrieved from the nwca database (u.s. environmental protection agency, 2016). differences in metrics between wetland classes were assessed using two univariate statistical methods. first, an analysis of variance (anova) was conducted to determine if a metric varied by wetland class. if the anova was significant (p ≤0.05) then we conducted a pairwise t-test between combinations of wetland classes to determine which had significant differences. magee and others’ (2019) nwca data analysis provided guidance on sifting through potential mmi metrics by identifying those that span an appropriate range, are repeatable and responsive to disturbance. skewed metrics or those observed over a very narrow range were removed as well as metrics that varied significantly over a single growing season (repeatability screen) or failed to distinguish between high and low disturbance sites (responsiveness screen). the 35 metrics that passed all three screens were equally scaled and standardized then assembled into unique mmi’s of three, four, and five metrics. these candidate mmi’s were screened through tests 9 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 of redundancy, sensitivity, and repeatability. the mmi described below has component metrics that are not highly correlated with one another (redundancy screen), distinguish between high and low condition wetlands (sensitivity screen), and remained consistent over the index period (repeatability) (magee and others, 2019). the final great salt lake vegetation-based multimetric index (gsl-vmmi) is a combination of three metrics: dicot species richness, cover of highly tolerant species, and cover of facultative wetland species. each of those metrics increases with disturbance, thus wetlands in good condition have more monocot species than dicot species and higher cover of species that are less than highly tolerant of disturbance and obligate wetland species. thresholds for good, fair, and poor condition were established individually for each wetland class based on the condition scores for sites that were in least disturbed reference condition (magee and others, 2019). setting condition thresholds based on a distribution is suboptimal because it creates a moving target with each new survey. however, it is the most realistic option for this dataset. the final part of the analysis was to conduct a risk assessment calculating the influence of individual anthropogenic disturbances on wetland condition, measured as relative and attributable risk (herlihy and others, 2019b). relative risk is a ratio that expresses the likelihood that a wetland will be in poor condition when a particular disturbance is high. attributable risk represents the proportion of wetlands in poor condition that could improve if a particular disturbance is removed. thresholds for distinguishing between high and moderate levels of a particular disturbance were set by analyzing the distribution of a particular disturbance and setting “high” at a point that marked the 33rd percentile for disturbances with normal data distributions or the inflection point for disturbances with skewed distributions. risk estimates are calculated based on contingency tables that tabulate the number of wetlands in two condition categories— not poor condition and poor condition— and two disturbance categories—high disturbance and not high disturbance (kincaid and others, 2019; r core team, 2020). the risk analysis assessed both the metrics that were part of the overall disturbance scores and individual parts of composite metrics (e.g., diversions were assessed separately from discharges) as well as potential sources of disturbance that are of particular interest to gsl stakeholders, like soil phosphorus and individual soil metals. three significant assumptions go into the risk analysis: 1) there is causality between a disturbance and condition; 2) a disturbance is reversible; and 3) disturbances are independent (herlihy and others, 2019b). both risk calculations are bounded by 95% confidence intervals and require large datasets to detect statistically significant risks. even with this relatively large dataset, the error bars on the risk estimates are quite large. further, if any cell in the contingency table is empty (e.g., there are no sites in poor condition with high disturbance from mines) the estimate for both risk factors will be zero. the risk results should be taken with these grains of salt— big assumptions, big error bars, and missing estimates— in mind. category  metrics  taxa composi on  species richness, na ve species a, introduced species a, simpson’s diversity b, shannon-wiener diversity b, species evenness  life history b  annual species, perennial species, forb species, graminoid species, monocot species, dicot species  hydrophy c status b  obligate species, obligate + faculta ve wetland species, faculta ve wetland species, faculta ve species, faculta ve upland + upland species  sensi vity/tolerance to disturbance b  sensi ve species, intermediate + insensi ve species, tolerant species,  highly tolerant species  vegeta on structure b  emergent species, submerged species, floa ng species, algae, bare  ground  floris c quality c  mean coefficient of conserva sm (cc), total cc, cover-weighted mean  cc, floris c quality index, cover-weighted floris c quality index  a – metrics include total richness, relative richness, total cover, relative cover, mean cover, frequency, and importance b – metrics calculated for all species present, native species only, and introduced species only c – calculated for all species and native species only table 1. plant community attributes calculated based on wetland survey data. bold attributes are those selected in the final mmi. 10 b. downard great salt lake wetland vegetation and what it tells us about environmental gradients, and disturbance results: gsl wetland ecological characterization we used the data gathered in gsl wetlands to answer three questions. first, what plants characterize gsl wetlands and how are they similar or different between classes? second, within a given wetland class, what factors drive variation in the plant community? finally, over all gsl wetland classes, what disturbances represent the most significant risk to wetland condition? what plants characterize gsl wetlands? over the nine years of plant surveys, we found 123 unique species across three gsl wetland classes. average species richness in gsl wetlands is five species per site, so even though we have a large species list, only 13 species were common, defined by being found in at least 10% of all gsl wetlands surveyed here, and most species were rare (downard and others, 2018). the most common species varied according to the wetland class being surveyed (impounded, fringe, or playa wetlands), though species are not exclusive to wetland class and can be found in multiple wetland classes. the first step in characterizing the community is to calculate and plot an ordination of the data, which summarizes complex patterns by visually highlighting species and sites that group together (mccune and grace, 2002). figure 4 is a non-metric multidimensional scaling (nmds) output of gsl wetland plant communities calculated based on the relative cover of the most common gsl wetland species. each color-coded point represents a wetland we sampled and the location along the vertical and horizontal axes show how similar or different the sampled plant communities are: points closer to each other have more similar communities and points farther from each other are more different. the text and grey points represent the center of a plant species’ area and indicate the most important species in that part of the ordination. along the horizontal axis (nmds 1), sites are figure 4. non-metric multidimensional scaling of gsl wetland plant community data. the ellipses represent a multivariate 95% confidence interval around the centroid of each wetland class. 11 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 generally grouped according to wetland class. impounded wetlands occupy the negative side of nmds 1, fringe wetlands occupy the center, and playa wetlands are on the right side. this pattern also matches the differences in hydroperiod and salinity, with deepest flooding and freshest water on the left/impounded side and saturation with saline water on the right/ playa side. figure 5a shows the distribution of observed water depth measurements in each wetland class and figure 5b shows the conductivity of surface water (impounded, fringe, and a minority of playa wetlands) or pore water (playa wetlands) recorded during field work. impounded wetlands were flooded most deeply of the three wetland classes while playa wetlands rarely had recordable surface water. salinity was similar between impounded and fringe wetlands, but significantly higher in playa wetlands. impounded wetlands are dominated by submerged aquatic vegetation (sav) species (table 2). for sav to grow, these wetlands must be flooded for most or all of the growing season, which creates highly anoxic soil conditions that limits nutrient availability and drives the buildup of reduced forms of elements like selenium and mercury which can potentially be toxic (cronk and fennessy, 2001). deep flooding also reduces light availability and gas exchange, which makes photosynthesis difficult (mitsch and gosselink, 2015). adaptations to this challenging environment include being rootless (ceratophyllum demersem, chara spp), utilizing bicarbonate in photosynthesis cycles (stuckenia pectinata, c. demersem), and having long, thin leaves that maximize surface area for light and gas exchange (all species in table 2). dense sav growth drives many ecosystem functions; it provides structure for aquatic macroinvertebrates, sequesters metals and nutrients from soils temporarily, and oxygenates water through respiration (cronk and fennessy, 2001). all the plant and macroinvertebrate growth in impounded wetlands create critical feeding habitat for migratory birds, especially larger birds like waterfowl. emergent species of cattails (typha spp), bulrushes (bolboschoenus and schoenoplectus spp), and grasses dominate in fringe wetlands (table 3). some emergent species can grow in water up to one meter deep (typha latifolia), but really thrive in water that fluctuates between flooded and saturated or dry conditions that submerged species cannot tolerate (larson, 1993). the species listed in table 3 have life history strategies adapted to a variable water regime. the seeds of all four common fringe species require bare ground to germinate, though these species readily expand via clonal growth under flooded conditions. clonal growth via adventitious rhizomes in combination with aerenchyma in their tissues allow patches of emergent species to share resources like oxygen and water across large distances, which supports the expanding margin of fringe wetlands (cronk and fennessy, 2001). emergent marshes are some of the most productive habitats on earth, enabling them to sequester soil metals and nutrients (reddy and delaune, 2008). dense vegetation also provides critical nesting habitat for migratory birds while vegetation that produces large seeds (e.g., bulrushes) also provides nutrient dense food (sweetman and others, 2013, marty and kettenring, 2017). playa wetlands are largely defined by being mostly expanses of bare ground, but a couple species of halophytes—species that grow specifically in salty and alkaline locations—also thrive (table 4). most plant species cannot grow in saline environments because high salt concentration makes it difficult for plants to obtain water and acquire beneficial elements (cronk and fennessy, 2001). distichlis spicata survives in saline wetlands through the ability to exude salt from specialized pores while salicornia rubra has adopted succulence and the ability to concentrate salts in specialized cells (welsh and others, 2004; hauser, 2006). s. rubra is the only common annual species in gsl wetlands and reproduces strictly by seeds, allowing vegetation to appear seasonally based on water availability. d. spicata, on the other hand, most commonly reproduces through rhizomes, allowing it to share resources amongst clonal stems. while the plant species of playas do provide some food for migratory birds, the macroinvertebrates in the soils are the most crucial resource for shorebirds that can probe the soils (sorensen and others, 2020). the isolated nature of playas also makes them critical nesting habitat for shorebirds because they are farther from infrastructure and predators than fringe or impounded wetlands. how do gsl wetland plant communities differ? the simplest measure of a plant community is species richness, which is a count of how many species are present. overall, species richness tends to be low in gsl wetlands but there are significant differences in richness between classes (figure 6). impounded wetlands have the lowest mean species richness (2.32), playa wetlands have intermediate richness (4.32), and fringe wetlands have the highest richness (7.92 species). both high environmental stress and high disturbance environments tend to have low species richness (cornk and fennessy, 2001) and later analyses will try to parse the impacts of disturbance versus stress. whether plants present are native to the region or introduced from elsewhere is a clearer indicator of 12 b. downard great salt lake wetland vegetation and what it tells us about environmental gradients, and disturbance figure 5. a) median (solid line) and mean (dashed line) maximum water depth; and b) median (solid line) and mean (dashed line) water conductivity in three classes of gsl wetlands. a b 13 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 how disturbed an environment is. introduced species that can establish and expand in new wetland environment often have adaptations that take advantage of gaps in vegetation because they have wide ecological tolerances, grow rapidly, and reproduce prolifically (zedler and kercher, 2004). this is especially true for phragmites australis (hereafter, phragmites), which occupies tens of thousands of acres of gsl wetlands (kettenring and others, 2020). introduced species relative cover (the proportion of all plant cover that is from introduced species) differs significantly in gsl wetland classes, matching patterns in species richness—highest in fringe wetlands and lowest in impounded wetlands (figure 7). as we explore the sources and consequences of anthropogenic disturbance in wetland plant communities, fringe wetlands and introduced species will come up again. growth form of the dominant plant in a type of wetland (forb, grass, or shrub) is how wetlands are mapped in the national wetland inventory, a comprehensive dataset of nationwide wetland extent, and those nationwide patterns also distinguish between gsl wetland classes. impounded wetlands tend to be aquatic bed features, fringe wetlands are predominantly emergent, and playa wetlands are those with less than 30% vegetation cover (u.s. fish and wildlife service, 2019). gsl wetlands are almost entirely herbaceous which means that woody species are uncommon and a small part of the overall cover when present. herbaceous plants can be further divided into graminoids—grasses, sedges, and other plants with grass-like growth patterns—and forbs—all the other species that tend to have broader leaves. wetland plants can also be grouped based on the length of their life cycle. annual species only live for one year whereas perennial species persist over multiple years, species  taxonomy  growth form  na ve  cc score / tolerance  chara – s nkweed  algae – characeae  annual or perennial,  macro-algae  na ve, obligate  wetland  undetermined  ceratophyllum demersum – coontail  dicot – ceratophyllaceae  perennial, submerged  aqua c forb  na ve, obligate  wetland  3 - tolerant  stuckenia pec nata –   sago pondweed  monocot – potamogetonaceae  perennial, submerged aqua c forb  na ve, obligate  wetland  3 – tolerant  ruppia cirrhosa –   widgeongrass  monocot – ruppiaceae  perennial, submerged  aqua c forb  na ve, obligate  wetland  6 – intermediate  table 2. dominant plant species in impounded gsl wetlands. species  taxonomy  growth form  na ve  cc score / tolerance  bolboschoenus mari mus –  alkali bulrush  monocot – cyperaceae  perennial, emergent graminoid  na ve, obligate  wetland  5 – intermediate  schoenoplectus americanus –  threesquare bulrush  monocot – cyperaceae  perennial, emergent graminoid  na ve, obligate  wetland  5 – intermediate  phragmites australis – phragmites  monocot – poaceae  perennial, emergent graminoid  introduced, faculta ve wetland  0 – highly tolerant  typha la folia – broadleaf  ca ail  monocot – typhaceae  perennial, emergent forb  na ve, obligate  wetland  2 – highly tolerant  table 3. dominant plant species in fringe gsl wetlands. species  taxonomy  growth form  na ve  cc score / tolerance  salicornia rubra – pickleweed  dicot – chenopodiaceae  annual forb  na ve, obligate  wetland  4 – tolerant  dis chlis spicata – saltgrass  monocot – poaceae  perennial graminoid  na ve, faculta ve  4 – tolerant  table 4. dominant plant species in playa gsl wetlands 14 b. downard great salt lake wetland vegetation and what it tells us about environmental gradients, and disturbance figure 6. median (solid line) and mean (dashed line) species richness in three gsl wetland classes. unique letters above boxplots indicate statistically different measures according to pairwise t-tests (α = 0.05). figure 7. relative cover of native and introduced species in three wetland classes. asterisks in legend indicate statistically different measures according to anova (α = 0.05). 15 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 growing back in subsequent seasons from perennating structures like rhizomes and tubers. the sav that characterizes impounded wetlands are primarily perennial forbs, the emergent species that dominate fringe wetlands are perennial graminoids, and the most common halophytes in playa wetlands are annual forbs (figure 8). as discussed earlier, a limited group of species is adapted to life in wetlands. however, even with their adaptations, wetland species are not uniform in their ability to tolerate natural or anthropogenic disturbance. sensitive plant species (as determined by u.s. environmental protection agency, 2016) are a small component of cover across all gsl wetlands. disturbance tolerant species cover the most area in gsl wetlands (figure 9). matching patterns reflected in introduced species cover by wetland class, the relative cover of highly tolerant species in fringe wetlands is significantly higher than in other wetland classes. coefficient of conservatism (cc) scores, the continuous metric that compliments the categorical sensitivity/tolerance variable, can be built into simple or complex measures of the floristic quality of the community (colorado natural heritage program, 2022). mean cc, the simplest of such measures, is nearly identical in impounded and playa wetlands, but significantly lower in fringe wetlands (figure 10). the floristic quality index (fqi) multiplies mean cc by a coefficient of species richness, and in gsl wetlands that flips the floristic quality results: fringe wetlands figure 8. relative cover of annual and perennial forb and graminoid species in three gsl wetland classes. asterisks in legend indicate statistically different measures according to anova (α = 0.05). other growth forms include shrubs, trees, and macroalgae. figure 9. relative cover of sensitive, intermediate, tolerant, and highly tolerant species in all gsl wetlands and within three wetland classes. asterisks in legend indicate statistically different measures according to anova (α = 0.05). 16 b. downard great salt lake wetland vegetation and what it tells us about environmental gradients, and disturbance have significantly higher fqi than the other wetland classes. the mechanisms for this switch in quality scores is clear, as fringe wetlands have higher species richness, but the implications are murky. what factors are associated with differences in the wetland plant community? the differences in the plant community between wetland classes described above are the result of a complex mix of environmental gradients, management actions, and anthropogenic disturbance. these gradients may also drive variation with each wetland class. nmds ordinations were generated using the most common species in each wetland class (those found in at least 10% of sites sampled for each class) and then overlaid with gradients of soil chemistry, water depth, and physical disturbances (table 5) to visually assess important gradients to each community (okansen and others, 2007). only those factors with relatively high r2 coefficients and p-values less than 0.05 were plotted because there was higher likelihood that those gradients are truly aligned with the plant community. however, measures of significance with ordinations do not hold the same rigor as in univariate data analysis and should be interpreted with that in mind (mccune and grace, 2002). impounded wetland sites clustered in two distinct communities of submerged species along the horizontal nmds1 axis and the vectors reflect common impounded wetland management practices (figure 11a). stuckenia pectinata, a highly valued habitat species for waterfowl, grows in deeper water than other sav species (see water depth vector) which is often at the farthest downstream point of impoundments (see impervious surface vector). ruppia cirrhosa favors more saline waters than other sav species and the conductivity vector increases along the positive side of nmds1. lemna minor is an indicator of nutrient enrichment (reddy and delaune, 2008) and the soil phosphorus and water quality discharge vectors both increase toward the upper left quadrant of the ordination that l. minor occupies. the divergent soil metal vectors are intriguing. copper, zinc, and lead vectors increase on the negative range of nmds 1 while selenium and barium follow the positive range of nmds 1. copper and zinc are both common in stormwater runoff from roads and it is possible l. minor and c. demersum could be indicators of contamination from roads (ladislas and others, 2012). the ordination of common species in fringe wetfigure 10. median (solid line) and mean (dashed line) mean cc and floristic quality index scores in three gsl wetland classes. 17 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 lands does not have the clear clusters of sites that impounded wetlands displayed, but the centroids of species indicate associations of species (figure 11b). three factors associated with water management (diversions, water depth, and roads that are built on dikes) all increase toward the lower quadrant of the ordination occupied by two species of interest to wetland managers: s. pectinata and bolboschoneus maritimus. it is possible the horizontal nmds1 axis reflects the influence of or similar conditions to adjacent wetland classes—impounded wetlands on the right and playa wetlands represented by salicornia rubra on the left. two common species in fringe wetlands that are classified as highly tolerant, phragmites and typha latifolia, occupy different sides of the vertical axis (nmds 2) which suggests that multiple gradients are driving different types of highly tolerant communities, one dominated by typha spp. and another by phragmites. based on the results of the nmds, physical distance from infrastructure may isolate playa wetlands from anthropogenic disturbance, which is reflected in the fact that no physical disturbance factors were meaningfully aligned with the playa plant community (figure 11c). although playa wetland sites did not cluster in a clear pattern, the species centroids did show that the right side of the plot is dominated by the salt-loving species s. rubra, puccinellia nuttalliana, and d. spicata along with a vector indicating higher soil salinity. the lack of clear vegetation patterns within playa wetlands may be due to the sparse vegetation present in this class of wetlands or the ephemeral nature of a community dominated by annual species. what condition are gsl wetlands in and why? while the previous sections detail the ways gsl wetlands are different between and within wetland classes, this final section will look at gsl wetlands collectively through the lens of wetland condition. recall that gsl wetland condition is measured through the gsl-vmmi, a composite of three metrics: cover of highly tolerant species and facultative wetland species and dicot species richness. gsl wetlands that experience little anthropogenic disturbance tend to have more monocot species than dicot species and more cover of wetland obligate and less tolerant species. as condition decreases dicot species become more numerous and facultative wetland species and highly tolerant species occupy more wetland area. through the process of selecting a vmmi explained in the analysis section, we know that condition is correlated with a cumulative measure of anthropogenic disturbance, but understanding the specific drivers of wetland condition requires a more robust analysis. risk analysis links the discrete measures of anthropogenic disturbance to poor wetland condition. relative risk analysis identifies the individual factors that contribute to poor condition by estimating the likelihood of a wetland being in poor condition if it also experiences high levels of a particular disturbance. ecological relative risk is analogous to heart disease risk: a human with high blood pressure (i.e., high stress or disturbance) is more likely to also have gradient  high disturbance threshold  environmental factors  water depth  -  conduc vity – water  -  conduc vity – soil  -  soil organic ma er  -  soil phosphorus ≥ 39.8 mg/kg  aluminum – soil  -  arsenic – soil ≥ 11.22 mg/kg  barium – soil  -  copper – soil ≥ 83.92 mg/kg  lead – soil  -  manganese – soil  -  nickel – soil  -  selenium – soil ≥ 0.17 mg/kg  zinc – soil  -  soil metal (exceedances of  background for as, ba, cu,  pb, mn, ni, se, and zn)  ≥ 5 exceedances  physical disturbances  water conduc vity  -  grazing severity  severe  herbicide severity  severe  impervious surface within  100m  >1 roads within 100m  -  discharges within 100m  diversions within 100m  impervious surface (%)  within 1km  ≥ 25%  diversions within 1 km  ≥ 3  discharges within 1 km  ≥ 1  developed and agricultural  land within 1 km  ≥ 6%  mines within 1 km ≥ 1  introduced species cover  ≥ 15% rela ve cover  table 5. environmental and anthropogenic gradients considered in nmds and risk analysis and cutoffs that distinguish high from low stress for risk categorization. 18 b. downard great salt lake wetland vegetation and what it tells us about environmental gradients, and disturbance figure 11. non-metric multidimensional scaling with significantly aligned environmental vectors in a) impounded, b) fringe, and c) playa wetlands. a b c 19 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 heart disease (i.e., poor condition) (herlihy and others, 2019b). attributable risk identifies the disturbances that, if removed, will result in improved condition. the estimate represents the proportion of poor condition sites that are likely to improve if a disturbance is removed. in the analogy of heart health, attributable risk is the improvement in heart health driven by decreasing blood pressure. risk estimates are interpreted with 95% confidence intervals; relative risk factors are considered significant if the lower confidence interval is greater than one and significant attributable risk factors have a lower confidence interval greater than zero (van sickle and paulsen, 2008). table 5 lists the disturbance factors included in the risk analysis and the threshold that separates high levels of disturbance from moderate to low disturbance. when all wetland classes are considered together, introduced species and changes in land use near a wetland are both significant relative risks. wetlands with more than 6% developed or agricultural land within one kilometer of the sample location are 2.6 times more likely to be in poor condition (figure 12). when wetland class is considered, however, land use change is only a significant risk for fringe wetlands. high cover of introduced species (>15% relative cover) is a significant risk for all classes of wetlands but has especially high relative risk estimates in impounded and playa classes, 38.62 and 5.46 respectively (table 6). the high relative risk of introduced species cover is likely driven by phragmites, which is widespread around gsl and has been a concern of wetland managers due to its propensity to crowd out native species and inability to support migratory bird use (cranney, 2016; long and others, 2017). phragmites is a facultative wetland species and highly tolerant to disturbance, properties that correspond to two metrics in the gsl-vmmi, thus there is some circularity in the risk and condition estimates. higher soil arsenic and selenium concentrations are also a significant relative risk to all gsl wetlands (1.15), which is an interesting complement to existing concerns about selenium in the gsl open water ecosystem (brix and others, 2004). selenium bioaccumulates in the open water food web, from algae to brine shrimp to aquatic birds. the gsl-specific research into selenium did not look at soils or wetland macroinvertebrates but research elsewhere has found a figure 12. relative risk estimates for environmental and anthropogenic stressors in all gsl wetlands and in three wetland classes. bold red boxes and asterisks indicate significant relative risk factors (estimate ± 95% confidence interval > 1). 20 b. downard great salt lake wetland vegetation and what it tells us about environmental gradients, and disturbance high potential for selenium accumulation in soils that are regularly flooded (jones and others, 2017). the most significant attributable risk factor for gsl wetlands is introduced species cover. when considered altogether, seven percent of poor condition wetlands would improve if the introduced species risk were removed (figure 13). the attributable risk estimate is largest for impounded and playa wetlands (63% and 39% respectively). it is encouraging that introduced species removal may improve wetland condition because years of research and adaptive management directed at phragmites removal has made significant progress in alleviating pressure from that species (rohal and others, 2017; rohal, 2018). though the circularity between condition metrics that reflect the presence of phragmites and risk estimates as well as the assumption of reversibility that is built into this analysis need to be remembered. the two other significant attributable risk factors, land use changes and soil metals, are unlikely to be reversible, regardless of the impact of their removal. land use change, a significant attributable risk for all gsl wetlands together and fringe wetlands in particular, are almost certainly permanent landscape features. soil selenium and arsenic are also difficult to remediate, not only because soil remediation is challenging, but also because wetlands act as landscape sinks for both arsenic and selenium, continually capturing metals from across the watershed (adams and others, 2015). however, decreasing soil metal concentrations would result in fewer poor condition fringe wetlands and have potential impacts for migratory bird populations, which can bioaccumulate both metals. conclusions characterizing the dominant plant communities and exploring the various environmental and anthropogenic gradients relevant to each class of gsl wetlands show the unique suite of factors that have filtered the plant community down to the species best adapted to each class. in impounded wetlands, deep freshwater flooding made possible by water management infrastructure supports predominantly native submerged aquatic plant species. the dynamic water regimes in fringe wetlands create an ideal environment for perennial emergent species. playa wetlands   popula on  disturbance  risk es mate  lower ci  upper ci  rela ve risk     all gsl  land use change  2.70  2.07  3.51     all gsl  introduced species cover  1.45  1.38  1.53     all gsl  soil arsenic  1.15  1.02  1.29     all gsl  soil selenium  1.15  1.03  1.28     impounded  introduced species cover  38.62  12.81  116.42     impounded  soil zinc  4.25  1.00  18.03     fringe  land use change  1.69  1.27  2.25     fringe  introduced species cover  1.30  1.24  1.36     fringe  soil arsenic  1.28  1.17  1.39     fringe  soil selenium  1.40  1.31  1.50     playa  mines  3.71  2.41  5.69     playa  introduced species cover  5.46  2.90  10.27     playa  soil metals  2.91  1.35  6.26     playa  soil zinc  2.97  1.39  6.32  a ributable risk     all gsl  land use change  0.59  0.48  0.09     all gsl  introduced species  0.07  0.06  0.09     all gsl  soil arsenic  0.02  <0.01  0.04     all gsl  soil selenium  0.04  0.01  0.06     impounded  introduced species  0.63  0.13  0.84     fringe  land use change  0.39  0.20  0.54     fringe  introduced species  0.05  0.04  0.06     fringe  soil arsenic  0.03  0.02  0.05     fringe  soil selenium  0.08  0.06  0.10     playa  introduced species  0.37  0.09  0.57  table 6. significant relative and attributable risk estimates for disturbances in gsl wetlands. 21 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 are dominated by species adapted to extremes in salinity. water depth is associated with differences in community within each class according to ordination results. finally, risk analysis identified land use change and introduced species as the greatest risks to condition and the greatest opportunity for restoration. restoration implications the ecological characterization presented here, the multi-metric index of wetland condition, and the risk analysis all have implications for restoration practices around gsl. identifying the correct potential plant communities, which are specific to wetland class, is critical to any restoration project and should be carefully considered when selecting species to plant, water regimes that are possible, and ultimate restoration targets (tarsa and others, 2022). the gsl -vmmi has a role both in identifying wetlands in need of restoration (those in poor condition) and in monitoring if a restored wetland is on a trajectory for better health over time. finally, the results of the risk analysis should be considered when identifying appropriate sites for restoration efforts. most especially, significant relative and attributable risk factors like introduced species should be minimized or eliminated prior to initiating restoration efforts. future research needs gsl wetlands form vast complexes of intermingling classes, which is what drives much of the bird diversity the ecosystem supports. the entire intermountain west region has experienced two decades of drought that pushed gsl to its lowest elevation and saltiest state. even with the impact of climate change on precipitation patterns, humans diverting and using water to grow food and lawns has exacerbated the impacts of drought (wurtsbaugh and others, 2016). as mentioned in the introduction, distinguishing between natural and anthropogenic disturbances is difficult and this is especially true for wetland water availability. the experience of the bear river delta in the early 20th century provides a stark example of the impact that years of drought can have on the ability of wetland complexes to provide their ecosystems functions. however, we also know that many wetland species are adapted to periodic drying events. future research into the natural range of hydrologic variability that 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hurlow, k. gunderson, g. atwood, editors 2023 utah geological association publication 51 wetzel, r.g., 2006, wetland ecosystem processes, in batzer, d.p, and sharitz, r.r., editors, ecology of freshwater and estuarine wetlands: los angeles, california, university of california press, p. 285– 312. wilson, v.t., and carson, r., 1950, bear river—a national wildlife refuge: washington d.c., u.s. fish and wildlife service. wurtsbaugh, w.a., miller, c., null, s.e., wilcock, p., hahnenberger, m., and howe, f., 2016, impacts of water development on great salt lake and the wasatch front: logan, utah, watershed sciences faculty publications, paper 875. wurtsbaugh , w.a., miller, c., null, s.e., derose, r.j., wilcock, p., hahnenberger, m., howe, f., and moore, j., 2017, decline of the world's saline lakes: nature geoscience, v. 10, no. 11, p. 816– 821. zedler, j.b., and kercher, s., 2004, causes and consequences of invasive plants in wetlands— opportunities, opportunists, and outcomes: critical reviews on plant sciences, v. 23, no. 5, p. 431 –452. smithwavedynamics.pub 1 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2023 utah geological association publication 51 abstract great salt lake is a natural laboratory to test and refine ideas about the relationship between sediment transport by waves and the characteristics of shoreline carbonate sediments, in particular ooid sands and microbialite mounds. in this chapter, we present a year-long series of wave data collected from july 2021 through june 2022 and use these wave data to assess the performance of a us army corps of engineers wave model previously used to estimate bed shear velocity and intermittency of sediment transport in great salt lake (smith and others, 2020). we use this model-data comparison to identify the strengths and weaknesses of the existing model for both geological and ecological applications, and areas of improvement for future model development. we also use shallow sediment cores and unmanned aerial vehicle (uav)-based orthomosaics collected from shorelines near each buoy to assess how the wave climate along two parts of the lake shore influences the stratigraphic record and the surface morphology of the lakebed. wave dynamics and sediment transport in great salt lake: a model-data comparison benjamin smith1, robert mahon2, tyler lincoln3, cedric j. hagen4, juliana olsen-valdez3, john magyar1, and elizabeth trower3 1division of geological and planetary sciences, california institute of technology, pasadena, california, bpsmith@caltech.edu 2department of earth and environmental sciences, university of new orleans, new orleans, louisiana 3department of geological sciences, university of colorado boulder, boulder, colorado, lizzy.trower@colorado.edu 4department of geosciences, princeton university, princeton, new jersey introduction great salt lake (gsl), ut is a critical ecological and economic resource—a key waypoint in the pacific flyway (paul & manning, 2002) and a primary source of magnesium metal in north america (tripp, 2009). gsl is also home to an exceptional modern geobiological archive of at least1000 km2 of meterscale microbialite mounds (vanden berg, 2019; baskin and others, 2021) that play a key role in the gsl ecosystem (wurtsbaugh and others, 2011). these mounds act as a food source and substrate critical for reproduction cycles for the brine fly, ephydra gracilis, and the brine shrimp, artemia franciscana, which in turn are key food sources for the millions of shorebirds and waterbirds that visit gsl each year (collins, 1980; wurtsbaugh, 2009; belovsky and others, 2011). radiocarbon dating, though complex, suggests that gsl microbialites have been accumulating for >10,000 years (bouton and others, 2016a; newell and others, 2017, 2020; homewood and others, 2022). previous authors have described gradients in microbialite morphology with distance from the shoreline and/or water depth (eardley, 1938; carozzi, 1962; bouton and others, 2016a, 2016b; vanden berg, 2019), suggesting that hydrodynamics and sediment transport, in addition to geochemistry and microbial metabolic activity, play a role in microbialite construction. previous workers have also observed that elongated microbialite mounds or domes tend to have preferred orientations relative to a shoreline and/or wave crests (bouton and others, 2016a; chidsey and others, 2015; vanden berg, 2019). these observations hint at a potential link between sediment transport and microbialite morphology and orientation. in theory, gsl microbialites are an ideal system in which to test ideas about the role of hydrodynamics and sediment transport on microbialite morphology because microbialites occur along shorelines with different orientations and therefore experience different wave conditions. however, this work first requires a more robust understanding of wave dynamics in the lake and an ability to accurately model past wave conditions (e.g., prior to causeway construction and at higher lake levels) along different shorelines. beyond microbialites, wave dynamics in the lake also affect other sedimentological and ecological characteristics of the lake. for example, the formation (including grain size and shape) of ooids is influenced by the frequency and energy of sediment transport (trower and others, 2020). wave dynamics can also influence mixing of the typically-stratified south arm lake water, which in turn affects the ecosystem by varying the availability and mobility of nutrients (belovsky and others, 2011) and delivery of toxins like mercury, selenium, and arsenic from the deep brine layer to the upper water column (beisner and others, 2009; jones and wurtsbaugh, 2014). furthermore, 10.31711/ugap.v51i.139 2 b. smith, r. mahone, t. lincoln, c.j. hagen, j. olsen-valdez, j. magyar, and e. trower wave dynamics and sediment transport given the historically low lake levels witnessed in 2021 and 2022 (abbott and others, 2023), accurate models of wave climate are needed to better understand how different future lake levels (higher or lower) will influence the hydrodynamics of the ecosystem. previously, trower and others (2020) and smith and others (2020) applied a linear wave model (rohweder and others, 2008) to calculate wave characteristics and bed shear stress using gsl bathymetry (baskin and allen, 2005; baskin and turner, 2006; tarboton, 2017) and wind data from the university of utah mesowest database (horel and others, 2002). however, this model was not necessarily designed to perform optimally for an environment like gsl, which includes very shallow and low sloping shorelines and the sharp changes in shoreline slope associated with the east lake fault scarp off the western shore of antelope island (colman and others, 2002). the purpose of this paper is to assess the performance of a linear wave model in gsl using data from two wave buoys in parts of the lake with contrasting shoreline orientations. -we evaluated the performance of this linear wave model using data from these wave buoys. we also present observations from sediment cores collected near each wave buoy and measurements of microbialite orientations adjacent to one wave buoy to assess how wave climate affects the composition and morphology of carbonate sediments along differently-oriented shorelines. our ultimate goal is to assess whether the model performs sufficiently well to be more widely applied to predict future and/or reconstruct ancient wave hydrodynamics in gsl. methods wind and wave measurements sofarocean spotter wave buoys were deployed at two locations along the south arm of the lake from the period of july 13, 2021, through june 28, 2022 (figure 1). the two sites were selected based on public interest and scientific importance. buoy #1356 (black rock) was deployed near gsl state park and buoy #1328 (miera spit) was deployed near the southern end of antelope island state park. previous work documented relationships between sedimentary facies and the physical environment near these locations, including trower and others (2020) with ooids near black rock, and smith and others (2020) with rip-up clasts and other storm features near miera spit. both buoys were deployed in relatively shallow water, approximately 2.5 m for #1356 and 1.6 m for #1328. for each wave buoy, water depth was measured using hobo u20l-04 water level loggers attached to anchors. each logger recorded pressure and temperature every 30 minutes; pressure measurements from each logger were converted to water depth using a water density of 1100 kg/m3 and corrected for atmospheric pressure using data from the kcc01 mesowest weather station. sofarocean buoys were chosen for deployment because they offer a lower-cost alternative to other wave monitoring techniques, and they transmit in real time through the use of onboard accelerometers and iridium satellite communication capabilities. buoys operated in six-hour cycles, alternating 5 hours of rest with one hour of active data collection. at the end of each cycle, buoys performed onboard processing of accelerometer data, converting it to spectral wave parameters—e.g., significant wave height, peak and mean direction, and spectral moments using opensource algorithms made available by sofarocean. at the end of the study, higher-frequency data were retrieved from the buoy’s onboard memory and processed using sofarocean parsing and analysis scripts (parser_v1.12.0). the full dataset for both buoys is available in an online repository (mahon and others, 2023), including additional wave data (e.g., directional spread, etc.) and metadata (e.g., temperature) not directly described in this chapter. for each buoy, hobo water depth time series data were examined to determine whether the anchors moved during the study period, as evidenced by substantial step changes in water depth. water depths for #1356 (black rock) varied smoothly between 1.732.95 m, reaching a low point of 1.73 m in october 2021, corresponding to the new historical low of 4190.1 ft (figure 2). buoy #1356 (black rock) was retrieved in good working order with no evidence of anchor movement or onboard electronics failures, indicating that the buoy made reliable wave measurements over the full study period. in contrast, water depths for #1328 (miera spit) started at 1.78 m but dropped rapidly to 0.85m on july 15, 2021, and dropped further to 0.43m on august 17, 2021 (figure 2). the timing of these rapid water depth changes matches the timing of buoy location changes when the anchor was dragged inshore by waves. when buoy #1328 (miera spit) was retrieved at the end of the study period, it was partially beached with its ballast chain touching the bed. we surmised that data quality was suspect after the second abrupt change in water depth on august 17, 2021, when the anchor was moved during a storm. a second consideration to buoy data quality was the detection limit of very lowamplitude waves. under calm conditions, buoy sensors experience an internal electronic “ringing” which produces spurious derived wave data with unrelated 3 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 1. the gsl-bb system. a. location of gsl and lake bonneville in western utah. b. overview map of gsl showing the historic average elevation, and the new 2022 historic low (figure from clark and baxter, 2023.) c. corresponding landsat satellite imagery of gsl elevations showing the record high of gsl in 1986 at left vs. historic low in 2022. ai = antelope island. images (images are public domain.) d. known bonneville basin lake cycles. the blue line labeled b in the main graph marks the bonneville deep-lake cycle. vertical black bars represent older deep-lake cycles. the base of the main graph is the elevation of modern gsl. inset shows the shoreline history of lake bonneville (blue) and gsl (red) with named shorelines (also see figure 3). (inset figure from oviatt and shroder, 2016a). figure 2. logs of water depth at each buoy location over the year-long deployment. note abrupt drops in water depth at miera spit buoy on july 15, 2021 and august 17, 2021. these correspond with storm events and affect data quality after july 15th. 4 b. smith, r. mahone, t. lincoln, c.j. hagen, j. olsen-valdez, j. magyar, and e. trower wave dynamics and sediment transport directions and magnitudes. these data were distinguished by their long periods (up to 20s) which were unreasonable for waves in the lake. from the 1400 wave spectral records at buoy #1356 (black rock), 793 were deemed to accurately reflect present conditions based on their wave periods (<10s). for buoy #1328 (miera spit), 41 records were deemed suitable based on the abbreviated operation period and noise screening. wind speed and direction were obtained from several wind stations in the mesowest database over a period coinciding with buoy deployment (7-13-2021 to 6-28-2022, figure 3). for both wind/wave comparisons and fetch-limited calculations, stations were selected based on proximity to each buoy and completeness of wind data over the study period. wind conditions near buoy #1356 (black rock) were taken from station kcc02, located in the marina of gsl state park (figure 3). wind conditions near buoy #1328 (miera spit) were taken from the station at hat island (hatut) because other, more proximal stations either had incomplete records or were not operational over the study period. comparing estimated and observed wave parameters wave parameters can be estimated using a combination of linear wave theory, empirically derived equations, and measurements for wind speed and fetch. many studies use methods developed by the us army corps of engineers as outlined in the coastal engineering manual (u.s. army corps of engineers, 2002) and the shore protection manual (coastal engineering research center, 1984). in particular, this approach underlies recent work by trower and others (2020) and smith and others (2020) which used an arcgis plugin (rohweder and others, 2008) to estimate wave parameters based on mesowest wind data (horel and others, 2002) and a digital elevation model (tarboton, 2017). a comparison between observed and calculated wave parameters provides direct feedback on the appropriateness of commonly used approaches for gsl, as well as potential complications due to wave refractions/diffraction, interactions with lake bathymetry, and inaccuracy of the bathymetry model. the model-data comparisons focused on significant wave heights, peak wave heights, and estimated shear velocities at the bed. significant wave heights, defined as the average height of the upper one-third of wave crests, were calculated using a procedure used in the coastal engineering manual for fetchlimited conditions (u.s. army corps of engineers, 2002): (1) (2) (3) (4) (5) where is the drag coefficient, is the wind speed (m/s) adjusted for height and whether the observations were collected over land or water, is the friction velocity (m/s), is the wind fetch, is the non-dimensional wind fetch, is the nondimensional significant wave height, is the significant wave height (m), is a constant with a value of 0.0413, is gravitational acceleration (9.81 m/s2), and is a constant with a value of ½. peak wave periods under fetch-limited conditions were calculated using: (6) (7) where is the non-dimensional peak wave period, is a constant with a value of 0.751, is a constant with a value of ⅓, and is the peak wave period. shear velocities were calculated using estimates of maximum orbital velocity and wavelength as intermediate steps. wavelengths (l) were calculated as: (8) and maximum orbital velocities, , were calculated as: (9) shear velocities, , were calculated as: (10) where f is the friction factor set to 0.032. note that while significant wave height and wave period provide direct comparisons between model estimates and 5 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 wave data, shear velocity is the most important variable for understanding sedimentary processes. unmanned aerial vehicle photography ever since eardley (1938) produced the first map of sediments in gsl, sedimentary studies in gsl have used multi-scale mapping to characterize shallow-water features such as bedforms and microbialites (bouton and others, 2016a, 2016b; vanden berg, 2019; smith and others, 2020; baskin and others, 2021). orthomosaic photos from unmanned aerial vehicles (uavs) provide intermediate-scale maps that bridge field observations and satellite imagery. to facilitate comparisons with directional wave data, an orthomosaic was collected on june 29, 2022, from the black rock area using a dji mavic air 2 at 200 ft standoff height via dronelink mission planning software (figure 1). a total of 877 orthophotos were collected and stitched together using pix4dcloud, covering an area of 0.278 km2. linear features within individual microbialites were measured using jmicrovision (roduit, 2019) by tracing the long axes of 100 microbialite ridges from the mapped area as well as five lineations that crosscut the primary microbialite ridge orientation in the northwest corner of the mapped area. figure 3. wind data for the study period from four stations in the mesow est database. ktvy = tooele valley airport, kslc = salt lake city airport, kcc02 = gsl state park marina, hatut = hat island. 6 b. smith, r. mahone, t. lincoln, c.j. hagen, j. olsen-valdez, j. magyar, and e. trower wave dynamics and sediment transport core recovery and grain size analysis we collected three sediment cores using an sdi vibecore mini electric vibracore: one core from near miera spit (gsl22-ms) and two cores from near black rock (gsl22-br-w and gsl22-br-e) (figure 1). we also collected an additional push core from a second location near miera spit (gsl22-sai). we split each core using electric shears and collected ~15 ml sediment samples every 4 cm from one half of each core; the second half of each core was described and archived. sediment samples were briefly rinsed with tap water to prevent grains from sticking together (due to salt precipitation from evaporating pore fluids) without dissolving minerals, then air dried. grain size and shape of each sediment sample was analyzed using a retsch camsizer p4. cores and subsamples of cores for analysis were registered with igsns (international geo sample numbers) in the sesar (system for earth sample registration) database; parent core igsns are listed in the results section and subsamples from each core have unique igsns associated with their respective parent igsn. results observed wind and wave conditions near black rock and miera spit wind data over the study period were variable across mesowest stations near black rock. south of black rock, wind stations have a predominantly n/s orientation. data from salt lake city airport (station kslc) are predominantly nw/se while data from bolinder-tooele valley airport (station ktvy) are n/s. wind orientations at both stations are consistent with previously observed lake breezes due to diurnal heating/cooling of the lake and land, as well as funneling of winds through the salt lake and tooele valleys, respectively (ludwig and others, 2004). the station closest to black rock, kcc02, has strong n, e, and sw components. the near absence of strong and/or frequent winds from the south likely reflects obstruction from the nearby oquirrh mountains. overlapping time series of both wind and wave data suggested that wave orientations generally aligned with local wind orientations. near black rock, wave observations, especially those with heights >15 cm, were oriented n/nnw (figure 4). less frequently, strong winds from the wsw produced waves arriving from this direction, most notably during late december of 2021 through early january 2022. although wind data had multi-modal orientation, the rose diagram of wave directions had a strong modal peak oriented at 350°. in turn, the modal peak aligned with the long direction of the lake relative to the local shoreline. the largest wave heights were also observed in this direction, which was consistent with previous assumptions about fetch-limited wave conditions. in contrast, wave data from miera spit differed from those at black rock, even though they were collected over a relatively short interval (figure 5). significant wave events at miera spit were aligned with winds from the sw rather than the n/nnw. model-data comparison for black rock for model-data comparisons at black rock, we restricted our analysis to observations for which both wind and wave directions had an orientation of 350° +/10° based on the mode in wave directions (figure 4). the estimated significant wave heights, peak wave periods, and shear velocities (red curves in figure 6) were calculated as functions of wind speed along a fetch of 54 kilometers using eqns. 1-7. results show that the model slightly overestimates significant wave heights, and the effect is most pronounced at high wind speeds (figure 6a). in contrast, model predictions for peak wave periods fall within the data (figure 6b), although a normal q-q plot (figure 6d) shows that the residuals are not normally distributed about the fit. shear velocities predicted by model results also agree with those calculated with observed wave parameters (figure 6c), although another q-q plot also shows some structure in the residuals (figure 6f). no model-data comparison was performed for miera spit due to the shorter time interval and fewer wave measurements. core sedimentology maximum depths of each core were 24 cm (gsl22-sai; igsn: 10.58052/ieejt008b), 60 cm (gsl22-ms; igsn: 10.58052/ieejt008a), 77 cm (gsl22-br-w; igsn: 10.58052/ieejt008c), and 74 cm (gsl22-br-e; igsn: 10.58052/ieejt008d). for all cores, the maximum depth of coring represents the depth of a resistant hardground that we could not penetrate with our equipment. sediments in cores from near miera spit (gsl22-ms and gsl22-sai) were mainly composed of ooids with minor peloids (primarily artemia fecal pellets), grapestones, and mica flakes; both cores lacked carbonate mud (figure 7). below 23 cm depth, ooid sands in the gsl22-ms core were roughly bimodal mixtures of fine and coarse ooids. sediments in cores from near black 7 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 4. wind and wave observations near black rock collected from july 2021 to june 2022. a) time series of wind speeds from a nearby wind station (kcco2) plotted with significant wave heights from buoy #1356. the heatmap, which depicts wind speed and azimuth, shows that large wind events (light colors) are predominantly oriented n/ne with a secondary sw orientation. significant wave heights >15 cm (yellow triangles) coincide with the timing and orientation of strong winds while wave heights <15 cm (grey crosses) have more variable orientations b) rose diagram of wind measurements binned by 20° increments. measurements are multimodal with w, s, and e peaks. c) wave azimuths for significant wave heights. wave directions were largely unimodal with an azimuth of 350° +/10°. note that the dominant wave direction is both a frequent wind direction and a long fetch relative to the shoreline at black rock. 8 b. smith, r. mahone, t. lincoln, c.j. hagen, j. olsen-valdez, j. magyar, and e. trower wave dynamics and sediment transport figure 5. wind and wave observations near miera spit collected from july to august 2021. a) time series of wind speeds from a nearby wind station (hatut) plotted with significant wave heights from buoy #1328. the heatmap, which depicts wind speed and azimuth, shows that large wind events (light colors) are predominantly oriented e or sw. significant wave heights >15 cm (yellow triangles) somewhat coincide with the timing and orientation of strong winds, although the match is weaker than at black rock. wave heights <15 cm (grey crosses) have variable orientations. b) rose diagram of wind measurements binned by 20° increments. measurements are mostly bimodal with e and sw peaks. c) wave azimuths for significant wave heights. wave directions were largely unimodal with an azimuth of 210° +/10°. note that the dominant wave direction differs from black rock (figure 4), reflecting differences between wind and shoreline orientation between the two sites. 9 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 rock (gsl22-br-w and gsl22-br-e) were composed of dark greento dark orange-brownpigmented, gravel-sized microbial mat or partially mineralized microbialite fragments (referred to as “pustular grains” and “microbial popcorn” by chidsey and others (2015), grapestones, angular carbonate sand grains (not ooids), and peloids (also primarily artemia fecal pellets) (figure 7); with the exception of a few horizons, both cores lacked muddy matrix. both black rock cores shared a similar sequence of ~20 cm of gravel-sized microbial mat and microbialite fragments overlying 40-50 cm of grapestone-dominated sediment. core gsl22-br-e had an additional 10.5-cm-thick layer of fine ooid sand overlying the microbialite fragment layer. grapestone compositions included aggregates of ooids, peloids, and microbialite fragments. the grain size and shape data are distinctly different between the miera spit and black rock sediment cores (figure 8). median grain diameters (d50) in the miera spit cores range from 369-496 μm (gsl22ms) and 332-434 μm (gsl22-sai), with mean roundness in both cores ranging from 0.710.77. these values are similar to previously characterized gsl ooids (trower and others, 2020), although the gsl22-ms core includes d50 values that are greater than reported in other areas. the grain size and roundness trends with depth are very similar between the two miera spit cores, showing little variability, although ooids in the gsl22-ms core are consistently larger than those in the gsl22-sai core. in comparison, the black rock cores depict more variability, where the two black rock cores differ most in the upper 20 cm. these trends match the lithologic variability observed in the cores (figure 8): samples in the upper 10.5 cm of gsl22-br-e have median grain diameters (d50 = 281-322 μm) and mean roundness (0.72-0.78) characteristic of ooids, while the microbial mat and microbialite fragments were very coarse sand to very fine gravel sized (d50 = 1448-3528 μm) and angular (mean roundness = 0.32-0.37) and the grapestones were coarse to very coarse sand sized (d50 = 911-1839 μm) and angular (mean roundness = 0.36-0.43). grapestones in the gsl22-br-w core were consistently coarser than the gsl22-br-e core. figure 6. comparison of wind speed with significant wave heights (a), peak periods (b), and calculated shear velocities (c). black circles represent simultaneous measurements of wind speed from station kcc02 matched with wave measurements at buoy #1356. points represent a subset of wind and wave measurements with an azimuth of 350° +/10°. for a and b, red curves show wave parameters (vertical axes) as a function of wind speed as calculated by eqns. 1-7 with a constant fetch of 54 km. for c, shear velocities were calculated using eqns. 8-10. black circles show shear velocities calculated using buoy observations of wave height and peak periods, while the red curve uses wave heights and peak wave periods calculated from eqns. 1-7. the shaded gray bar shows the range of shear velocities most relevant to sediment transport near black rock: the lower and upper bounds represent thresholds for motion and suspension, respectively, for 370 μm sand. 10 b. smith, r. mahone, t. lincoln, c.j. hagen, j. olsen-valdez, j. magyar, and e. trower wave dynamics and sediment transport together, the core sedimentology data indicate that the miera spit area has historically been characterized almost exclusively by the production and deposition of ooid sand. in contrast, the black rock area was instead a grapestone factory prior to the more recent development of a continuous blanket of microbialites, overlain by a mobile and transient layer of ooid sand. although the sets of miera spit and black rock cores were both significantly more similar amongst each set than between sets, both sets of cores displayed more subtle but systematic differences in grain size associated with their different locations along each shoreline. uav imagery of microbialite forms uav orthomosaic imagery at black rock reveals the orientations of exposed microbialite ridges (figure 9). individual microbialite ridges trend nnw/ sse along long axes. multiple microbialites from the northwestern corner of the orthomosaic form an additional array of lineations roughly 20-25 m in length, trending ne/sw. discussion wave orientations differ between sites the differing wave orientations at the two sites reflect differences in the fetch between the two shorelines. at black rock, the predominant n/nne wave orientation aligns with the long direction of the lake, and thus the longest available fetches. at a broader level, the regional geology provides an underlying factor linking basin orientation and diurnal wind patterns. both the shape of the lake and the nne/sse wind directions follow the strong topography of nearby mountain ranges such as the wasatch in the east figure 7. annotated images of sediment cores. 11 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 8. grain diameter and roundness from the four collected cores. diameter and roundness metrics derived from camsizer output data, where the red line indicates the 10th percentile (triangles), black indicates the 50th percentile (circles), and blue indicates the 90th percentile (squares). 12 b. smith, r. mahone, t. lincoln, c.j. hagen, j. olsen-valdez, j. magyar, and e. trower wave dynamics and sediment transport (figure 3). these effects are especially strong towards the south end of the lake where onshore/offshore lake breezes are funneled through the tooele and salt lake valleys (figure 3). although the windrose for the marina (kcc02) is more complex than those in the valleys, the nearly unimodal wave directions from the north can be explained by short fetches for winds not aligned to the north. in contrast, wave orientations at miera spit have a different alignment because the shoreline is nearly perpendicular to its counterpart near black rock. winds from the n/nnw have short fetches obstructed by antelope island, gunnison island, and promontory point. instead, most waves had a sw orientation, which is consistent with the analysis of nearby wave ripples and bar forms by smith and others (2020). the regional wind patterns that generate these waves differ from the predominant nnw/sse winds along the basin axis, but are aligned with the predominant southwesterly to southeasterly orientation of the strongest winds in the eastern and southern bonneville basin from 1946–1993 compiled by jewell (2007). waves, microbialite morphologies, and paleoflow indicators at first glance, the near-unimodal orientation of waves at black rock provides a compelling test of microbialite ridge orientations as paleoflow indicators. however, even a first-order analysis precludes a 1:1 mapping of microbialite ridge orientation onto wave directions. first, there are at least two sets of sufigure 9. uav image of exposed microbialite ridges west of black rock, near buoy #1356. individual microbialite ridges trend nnw/sse along long axes. multiple individual microbialite ridges group together to form arrays of lineations that trend ne/sw, highlighted in teal in image inset. 13 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 perimposed linear microbialite orientations in the area we analyzed near the black rock buoy, and the expected orientation of either set of linear features relative to the dominant wave direction likely depends on their origin. for example, vanden berg (2019) documented incipient microbialites forming on the crests of wave ripples. if some lineations in mature microbialites reflect underlying nucleation on bedforms, then the lineations should be perpendicular to the direction of the waves. in contrast, the long axes of microbialite ridges near the black rock buoy have a strong onshore/offshore orientation that is nearly parallel to measured wave directions. thus, even when wave orientation reasonably influences microbialites, the orientation of lineations relative to wave features– and thus, their use as a paleoflow proxy–is complex. additionally, other origins for strong lineations (e.g., underlying faults and fractures) must be accounted for. a more complete understanding of potential paleoflow indicators has several spatial and temporal correlations. for example, why are the onshoreoffshore lineations at black rock rotated with respect to measured wave orientations? one possibility is that waves observed at the buoy refract as they interact with steep and irregular bathymetry near the shoreline. another possibility is that microbialites reflect a time-integrated signal of wave conditions, and that direct comparison to modern waves represents a recency bias. addressing this issue requires more data on the absolute ages of microbialites and their underlying sediments from the vibracore recoveries (figure 7). sediments from miera spit and the black rock shoreline have not been previously dated with radiocarbon. however, radiocarbon dating of microbialites from the northern shores of antelope island (i.e., bridger bay, buffalo point, and white rock bay) and from the north arm have suggested at least two pulses of microbialite formation from ~11.4 and 8 ka, and 3.8 and 1.7 ka (bouton and others, 2016b; newell and others, 2017, 2020). furthermore, radiocarbon dating of gsl ooids from northern antelope island (bridger bay) and the north arm (spiral jetty) suggest that ooids at the modern sediment surface have been slowly accumulating over the past ~6 ka (paradis, 2019). together, these data suggest that the microbialites on the black rock shoreline are likely at least 1.7 ka in age, and potentially thousands of years older. a dominant nnw orientation may therefore reflect basin orientation and tectonic effects on topography, which are stable over these time periods. the deviation from modern waves could be also explained by differences in wind patterns due to variations in holocene climate, as well as differences in lake level and surface area. nevertheless, because the rohweder and others (2008) wave model performs relatively well at matching the modern wave data, we suggest that this model could be a useful tool to evaluate how different wind patterns and/or lake level in the past might better explain the microbialite ridge orientations. influences of wave-driven sediment transport on sedimentary facies two notable sedimentological differences between the miera spit and black rock sites could be related to differences in transport mode and frequency: (1) the relatively large ooid diameters at miera spit (primarily upper medium sand sized, whereas most ooids elsewhere in the lake including at black rock are lower medium sand sized); and (2) the contrast between ooid-dominated sediments at miera spit and grapestone-dominated sediments (at depth in cores) at black rock. here, we suggest hypotheses about how transport mode and frequency may influence the distinct sedimentology of these sites and use our data to provide an initial evaluation of these hypotheses. trower and others (2020) noted that the aragonite saturation state (ωar) of gsl water is lower than that characteristic of seawater in modern marine ooidforming environments. this relatively low ωar value explains the relatively small sizes of gsl ooids because it results in relatively slow precipitation rates and therefore smaller equilibrium ooid sizes (trower and others, 2017). furthermore, gsl ooids are so small due to low lake water ωar that many of their sizes are close to the threshold below which impacts are completely viscously damped, resulting in no abrasion. due to this effect, increasingly frequent transport events (i.e., increasing intermittency, fint) cannot reduce ooid size beyond ~200 μm. many gsl ooid sizes are close to this threshold (figure 10). however, miera spit is unique in that ooids there have grown to larger sizes than observed in other locations. within the equilibrium ooid size framework, we would therefore predict that the larger ooid sizes at miera spit must be associated with either less frequent transport (lower fint), or more energetic transport (higher u*) (figure 10). model-based estimates of intermittency of movement suggest similar values in the range of 1-2% (fint = 0.01 0.02) at these two sites (smith and others, 2020; trower and others, 2020). due to the limited size of the miera spit dataset, we are unable to make a robust comparison of shear velocities at the two sites to assess whether differences in u* might be driving the larger ooid sizes at miera spit. an analysis of wind patterns along the eastern and southern margins of the bonneville basin over a 14 b. smith, r. mahone, t. lincoln, c.j. hagen, j. olsen-valdez, j. magyar, and e. trower wave dynamics and sediment transport longer observation duration than our study indicated that the strongest modern winds were from the swse over the period between 1946–1993 (jewell, 2007). these southwesterly to southeasterly winds would have resulted in higher shear velocities at miera spit than at black rock due to the differences in fetch. the dynamics of grapestone formation have been less thoroughly examined than those of ooids. some workers have suggested that microbially-mediated carbonate mineral precipitation plays a key role in grapestone formation (purdy, 1963; winland and matthews, 1974; fabricius, 1977; diaz and others, 2022). if this process is the key factor driving the formation of grapestones in gsl, we might expect to see microbial community differences between the miera spit and black rock areas. although we did not collect microbial diversity data as part of this study, ingalls and others (2020) did report some notable differences between ooid-dominated sediments from bridger bay (on the northern part of antelope island) and ooid-dominated sediments between black rock and gsl state park. in particular, relative to samples from the gsl state park/black rock site, samples from bridger bay lacked cyanobacteria (which are commonly implicated in driving carbonate precipitation) and had more abundant chloroflexi and deinococcus-thermus sequences (ingalls and others, 2020). however, there is no evidence directly linking this specific microbial community difference to sedimentary facies differences between those two sites. further microbial community analyses of the miera spit area might help to better evaluate this hypothesis. however, it is not clear that the modern surface microbial community at each site would be representative of the community that was present when sediments at the base of each core were forming, particularly given that the modern lake microbial community figure 10. plot of predicted equilibrium ooid sizes (deq) as a function of bed shear velocity (u*) following trower and others (2017) for a range of intermittencies (fint), compared with violin plots of shear velocities from the black rock buoy (horizontal violin) and pooled ooid size data from the three cores that contained ooid-dominated layers (vertical violins). solid black line shows threshold of motion, dashed black line shows threshold of suspension, and the dash-dot black line shows the viscous damping threshold (stokes number, st = 1) below which grains cannot abrade. the larger ooids at miera spit could be explained by lower intermittency (less frequent transport) and/or higher shear velocity. 15 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 has already been influenced by the recent historically low lake levels (frantz and others, 2022). alternatively, one could also speculate that physical, rather than biological, processes are responsible for the black rock grapestone factory. grapestone formation is commonly thought to reflect very infrequent but very energetic transport, providing long rest periods for grains to be cemented together with transport events that can still entrain these relatively coarse compound grains. this explanation appears to conflict with the relatively similar estimates of intermittency between the two sites and our interpretation of higher shear velocities at miera spit based on differences in ooid size. the grapestones in the black rock cores could therefore reflect an older and deeper lake stage than that represented by ooids in either location. this idea is supported by our observation of grapestones in surface sediments (collected via dredge), which we only found in deeper waters (2.5-3 m water depth) off the southern antelope island shoreline (figure 11). there, grapestones occurred in the troughs between microbialite mounds and ooids were rare. in contrast, we did not find any grapestones in microbialite troughs near black rock, suggesting that, currently, transport conditions are not as conducive to grapestone formation even in deeper figure 11. images of grapestones collected by dredge at the modern sediment surface in deeper water near the southern tip of antelope island. a) grapestones occurred in patches of mobile sediment in troughs between microbialites, as illustrated in this image from a submersible remotely-operated vehicle (rov). b-c) grapestone-rich sediment collected by dredge (b) and zoomed-in field image of grapestone-rich sediment (c). d) stereoscope image of grapestones from this location. 16 b. smith, r. mahone, t. lincoln, c.j. hagen, j. olsen-valdez, j. magyar, and e. trower wave dynamics and sediment transport water along that shoreline. morphological analyses of spits associated with older bonneville shorelines suggest that in the late pleistocene (i.e., prior to the gilbert episode (oviatt, 2014)), wave transport in the lake was dominated by strong northerly to northwesterly storms (schofield and others, 2004; jewell, 2007). infrequent but strong wave currents from these types of storms could be consistent with the optimal shoreline orientation and location for grapestone development in the past differing from that in the modern lake. geochronological constraints and petrographic analysis of buried grapestone sediments at black rock are needed to further test this hypothesis and evaluate the roles of microbial community versus hydrodynamics on grapestone formation in gsl. again, given the relatively good fit between the rohweder and others (2008) model and the wave buoy data for the black rock area, we suggest that the model would be a useful tool to reconstruct historical wave conditions in this area. strengths, weaknesses, and potential applications of model cross-validation of wave models with buoy data provides several key takeaways for future studies of gsl across past, present, and future. in optimal cases (i.e., when high-quality, continuous wind data are available near the shoreline of interest), fetch-limited wave models yield reasonable results for key variables such as shear velocities. however, the plots of the residuals in figures 6d-f suggest that there is unexplained structure in wave observations that are not captured by the model. a likely source of discrepancies is that some of the empirical constants in eqns. 1 -7 were calibrated for seawater, which is less dense than gsl lake water. in particular, the drag coefficient in eqn. 1 is sensitive to temperature and density variations of both air and water (le roux, 2009). a density effect could reasonably affect all three parameters in figures 6a-c since they all involve the drag coefficient. le roux (2009) also notes that fluids denser than seawater—for example, those with high suspended sediment loads—produce waves that are smaller than those predicted by eqns. 1-7. the overprediction of wave heights observed in fig. 6a is thus consistent with a density effect. the dependence on nearby wind stations is both a challenge and an opportunity. even with limited results from miera spit and black rock, fetch-based wave models reasonably predict spatial differences in shear velocities and frequencies of motion that are crucial for further studies of how sedimentary facies are distributed across gsl. however, the variety of wind conditions observed at different mesowest stations (figure 3) strongly suggests that the quality of modeled wave parameters strongly depends on the proximity of relevant wind data. for example, onshore data from salt lake city (kslc) and tooelebolinder airport differ considerably from stations on the shores of the lake itself, such as kcc02 and hatut. at a more granular level, predicting conditions along specific shorelines requires local wind data. at present, the marina has two wind stations relevant to gsl state park (as768 and kcc02), but relevant wind data for antelope island state park have been challenging to obtain since the loss of the bridger bay station in 2018. while data from other stations, such as hat island (hatut), may be appropriate for sedimentary research, more proximal data is needed should these wave models become important for gsl conservation efforts and policymaking (rohweder and others, 2008). even a rudimentary agreement between wave models and empirical data opens the door to using these models to study how past anthropogenic and climatic changes may have modulated the sedimentary facies we observe today. for example, construction of the causeway divided the lake into chemically distinct north and south arms; did this causeway also change effective fetches, especially for shorelines near black rock? since the causeway is recent within the context of the lake's holocene inception, modeling precauseway conditions might prove instructive for interpreting both surficial and cored sediment data collected from near gsl state park (figure 7). for natural climatic variations, previous research has suggested that lake bonneville and other paleolake shorelines were associated not only with higher lake levels, but different prevailing winds ( schofield and others, 2004; jewell, 2007). while linkages between lake level, paleoclimate, and lake chemistry strongly affect carbonate facies, fetch-limited wave models may provide a more holistic view of how paleolake levels related to sediment transport conditions beyond simple changes in water depth. finally, it is worth noting that the fetch-limited wave models used here and in previous work (smith and others, 2020; trower and others, 2020) have applications outside of sedimentary geology, such as for environmental forecasting and conservation. in fact, the arc plugin used in these studies (rohweder and others, 2008) was originally developed by the usgs for environmental conservation and management. while environmental forecasts and recommendations are beyond the scope of this work, we do point out that basic model-data comparisons–especially with respect to shear velocity and sediment mobility–are fundamental to future applications of lake modeling 17 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 with regards to gsl environmental conservation and policy making. acknowledgements permission to temporarily anchor buoys to the lakebed was granted by the state of utah department of natural resources under right of entry permit no. 410-00698. additional permission to access the lake shore at miera spit by land via antelope island state park was provided by utah division of parks and recreation special use permit p72-21. references abbott, b.w., baxter, b.k., busche, k., de freitas, l., frei, r., gomez, t., et al., 2023, emergency measures needed to rescue great salt lake from ongoing collapse. baskin, r.l., and allen, d.v., 2005, bathymetric map of the south part of great salt lake, utah, 2005: usgs. baskin, r.l., and turner, j., 2006, bathymetric map of the north part of great salt lake, utah, 2006: usgs. baskin, r.l., della porta, g., and wright, v.p., 2021, characteristics and controls on the distribution of sublittoral microbial bioherms in great salt lake, utah: implications for understanding 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biostromes, brine flies, birds and the bioaccumulation of selenium in great salt lake, utah: natural resources and environmental issues, 15(2), 1–13. wurtsbaugh, w.a., gardberg, j., and izdepski, c., 2011, biostrome communities and mercury and selenium bioaccumulation in the great salt lake (utah, usa): the science of the total environment, 409(20), 4425–4434. http://paperpile.com/b/briy5y/6k17 http://paperpile.com/b/briy5y/kmo6 http://paperpile.com/b/briy5y/kmo6 http://paperpile.com/b/briy5y/kmo6 http://paperpile.com/b/briy5y/2h1h http://paperpile.com/b/briy5y/2h1h http://paperpile.com/b/briy5y/fwrk http://paperpile.com/b/briy5y/fwrk http://paperpile.com/b/briy5y/fwrk http://paperpile.com/b/briy5y/fwrk https://doi.org/%2010.46428/sld https://doi.org/%2010.46428/sld http://paperpile.com/b/briy5y/8row http://paperpile.com/b/briy5y/8row http://paperpile.com/b/briy5y/8row http://paperpile.com/b/briy5y/8row http://paperpile.com/b/briy5y/8row http://paperpile.com/b/briy5y/8row https://doi.org/10.1130/abs/2020rm-346645 http://paperpile.com/b/briy5y/m9me http://paperpile.com/b/briy5y/c40v http://paperpile.com/b/briy5y/c40v http://paperpile.com/b/briy5y/c40v http://paperpile.com/b/briy5y/c40v http://paperpile.com/b/briy5y/idr0 http://paperpile.com/b/briy5y/2rwj http://paperpile.com/b/briy5y/2rwj http://paperpile.com/b/briy5y/2rwj https://jmicrovision.github.io http://paperpile.com/b/briy5y/4sjn http://paperpile.com/b/briy5y/4sjn http://paperpile.com/b/briy5y/4sjn http://paperpile.com/b/briy5y/nzan http://paperpile.com/b/briy5y/nzan http://paperpile.com/b/briy5y/nzan http://paperpile.com/b/briy5y/nzan http://paperpile.com/b/briy5y/nzan https://doi.org/10.1029/2020jf005733 http://paperpile.com/b/briy5y/r03h http://paperpile.com/b/briy5y/r03h http://paperpile.com/b/briy5y/sknr http://paperpile.com/b/briy5y/sknr http://paperpile.com/b/briy5y/sknr http://paperpile.com/b/briy5y/m6gr http://paperpile.com/b/briy5y/m6gr http://paperpile.com/b/briy5y/m6gr http://paperpile.com/b/briy5y/m6gr http://paperpile.com/b/briy5y/m6gr https://doi.org/10.%201029/2019jf005452 https://doi.org/10.%201029/2019jf005452 http://paperpile.com/b/briy5y/a4ao http://paperpile.com/b/briy5y/a4ao http://paperpile.com/b/briy5y/cj7b http://paperpile.com/b/briy5y/cj7b http://paperpile.com/b/briy5y/8oh6 http://paperpile.com/b/briy5y/8oh6 http://paperpile.com/b/briy5y/8oh6 http://paperpile.com/b/briy5y/aifm http://paperpile.com/b/briy5y/aifm http://paperpile.com/b/briy5y/aifm http://paperpile.com/b/briy5y/aifm http://paperpile.com/b/briy5y/tfoi http://paperpile.com/b/briy5y/tfoi http://paperpile.com/b/briy5y/tfoi http://paperpile.com/b/briy5y/tfoi http://paperpile.com/b/briy5y/tfoi c:\users\gnuwi\appdata\local\temp\msobbb1.tmp 1 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2023 utah geological association publication 51 abstract the elevation of great salt lake has fallen to historic lows in recent years, exposing once submerged microbialites along the lake’s shores. although prior studies have attempted to map microbialite locations, this has proved challenging, with mapped microbialite areas limited to accessible shoreline locations or via indirect sonographic evidence. meanwhile, the importance of great salt lake’s microbialites to the lake’s food chain has made quantifying the extent of microbialites exposed versus submerged at different lake elevations critical to lake management decisions. low lake levels combined with seasonal high-water clarity have enabled microbialite reefs to be spotted in aerial and satellite imagery, even in deeper areas of the lake. in this study, satellite images were used to identify and map microbialite reef areas in great salt lake and along its dry shores. in the south arm, submerged microbialites were easily recognized as dark green reefs against a light-colored benthic background (primarily ooid sand). stationary microbialite mounds were distinguished from rip-up clasts or other dark-colored mobile material by comparing potential microbialite regions across several high-visibility timepoints. in this way, we identified 649 km2 (251 mi2) of putative microbialite reef area: 288 km2 (111 mi2) in the north arm, 360 km2 (139 mi2) in the south arm, of which 375 km2 (145 mi2) was mapped at a high degree of confidence. we also produced geospatial shapefiles of these areas. this map, combined with currently available lake bathymetric data, permits the estimation of the extent of microbialite reef exposed vs. submerged in various parts of the lake at different lake elevations. at the end of fall 2022, when lake level dipped to 1276.7 masl (4188.5 ft-asl) in elevation, we estimate that ~40% of the south arm microbialite reef area was exposed. use of remote imagery to map microbialite distribution at great salt lake, utah: implications for microbialite exposure laura wilcock1,2, carie m. frantz1, and michael d. vanden berg3 1department of earth & environmental sciences, weber state university, ogden, utah; laura.wilcock@utah.edu 2department of geology and geophysics, university of utah, salt lake city, utah 3energy & minerals program, utah geological survey, salt lake city, utah introduction microbialites cover substantial portions of the great salt lake benthos, and host microbial communities are believed to be important to the great salt lake ecosystem. models of the lake’s ecosystem, therefore, must necessarily incorporate estimates of microbialite extent (belovsky and others, 2011; barrett, 2020), which need refining, particularly in the face of recent lake level decline and microbialite exposure. microbialites in modern great salt lake great salt lake is the largest saline lake in the western hemisphere. unlike other terminal lakes in the basin and range of the western united states, which tend to be alkaline, great salt lake is a na-mg -cl-so4-dominated system with relatively low levels of alkalinity (domagalski and others, 1989; jones and others, 2009). high rates of ca2+ and hco3 delivery, slightly alkaline surface waters, the lake’s hypersalinity (which promotes co2 degassing), and high levels of microbial activity produce conditions that approach or exceed aragonite saturation in much of the lake, despite relatively low lake water concentrations of ca2+ and co3 (pace and others, 2016; ingalls and others, 2020; bouton and others, 2020). these factors have made great salt lake (as well as its predecessors) a “carbonate factory,” with carbonates making up a major portion of lake sediments, especially since the draining of pleistocene lake bonneville (jones and others, 2009; vennin and others, 2019). carbonate deposits blanket the modern bed of the lake, and include organic-rich carbonate mud, oolitic sand, and microbialite reefs (eardley, 1938; chidsey and others, 2015; vanden berg, 2019; ingalls and others, 2020; bouton and others, 2020; baskin and others, 2022). microbialites are “organosedimentary deposits formed from interaction between benthic microbial communities…and detrital or chemical sediment” (burne and moore, 1987). they are typically formed by processes of trapping and binding by microbial mats (for example, frantz and others, 2015), 10.31711/ugap.v51i.136 2 l. wilcock, c.m. frantz, and m.d. vanden berg use of remote imagery to map microbialite distribution at great salt lake induction of mineral precipitation via metabolic activities of microbial communities (for example, dupraz and others, 2009), and/or inorganic calcification (for example, shen and others, 2022). an aside on terminology: the term “bioherm,” ostensibly coined by cumings and shrock (1928), broadly refers to any reeflike mound built by living organisms. “microbialite,” meanwhile, refers to a sedimentary rock built at least in part by the activities of microorganisms (burne and moore, 1987). thus, “microbialite reef” is subtly different from “bioherm,” indicating that microorganisms are involved in the construction of the reefs, but also acknowledging potential abiogenic contributions. great salt lake’s microbialites were first documented by eardley (1938) in his seminal tome describing the lake’s chemistry and sediments, describing in detail the ''extensive calcareous bioherms'' that were visible during a period of relatively low lake elevation in the mid-1930s (figure 1). he noted their dense mats (periphyton), dominated by the cyanobacterium aphanothece packardii (now identified as eufigure 1. a) modern great salt lake south arm surface elevations as measured at usgs water monitoring locations 1001000 and 10010024. green shaded areas indicate 1 ft elevation bands below 4200 ft-asl where microbialites were mapped (this study), with shade indicating the total percentage of microbialites that would be submerged at that lake elevation. b) lake bonneville-great salt lake hydrograph (black line) showing ages and elevations of dated microbialite materials from bouton and others, 2016a (light green circles) and newell and others, 2017 (dark green circles). hydrograph prior to 13 ka modeled after oviatt, 2015. hydrograph after 13 ka modeled after oviatt and others, 2021, with the dark gray horizontal bar indicating the uncertainty in lake elevation during the great salt lake phase. 3 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 halothece spp.; lindsay and others, 2019; frantz and others, 2023), and attributed their formation to microbially-mediated carbonate precipitation. even lower lake elevations in the early 1960s afforded a second look at the lake’s microbialites. carozzi (1962) examined their morphological variability and spatial distribution and linked their occurrence to underlying topographic highs. halley (1976) described in detail the high variability in their internal structure, with laminated and unlaminated microfabrics existing within different portions of single microbialites (hence, “microbialite” vs. a more descriptive term such as thrombolite or stromatolites). he also noted a general lack of relationship between the living periphyton and observed calcified microstructure and microfossils, notably remarking that, “the organisms on the surface of the great salt lake algal mounds are probably not those which are responsible for the internal structure.” by the late 1960s, the lake’s microbialites were once again submerged by a rise in lake level and all but forgotten until they reappeared in the early 2010s during the period of prolonged lake level fall after the 1986–1987 lake highstand. this ushered in a new era of great salt lake microbialite research in which the microbialites were investigated as contributors to the lake ecosystem (wurtsbaugh, 2009; belovsky and others, 2011; wurtsbaugh and others, 2011) and as geobiologic curiosities (pedone and folk, 1996; baskin, 2014; pace and others, 2016; lindsay and others, 2017). interest in the structures was further enhanced by the discovery of the microbialiteassociated pre-salt petroleum deposits of offshore brazil in the mid-2000s, with interest in great salt lake as a potential modern analog environment (chidsey and others, 2015; vanden berg, 2019). recent studies utilized new techniques and technology, including advanced microscopy (pace and others, 2016), molecular biology (lindsay and others, 2017), geospatial and marine acoustic technology (baskin, 2014; baskin and others, 2022), and drone imagery (vanden berg, 2019). while the bulk of academic focus on the lake’s microbialites (including that of this paper) has been on the extensive reefs that are submerged during “normal” levels of the modern lake, i.e., those below about 1280 meters above sea level (masl; 4200 feet above sea level, or ft-asl), microbialites and other putative microbial carbonates are also found in discrete locations at higher elevations, associated with earlier phases of the lake system (chidsey and others, 2015; vennin and others, 2019; homewood and others, 2022). however, in the remainder of this paper, we use “microbialites” to refer only to the reef-forming deposits below 1280 masl (4200 ft-asl) in great salt lake and its recently exposed shores. the megaand macrostructure (shapiro, 2000) of great salt lake’s microbialites includes roughly circular domes ranging in size from ~15–300 cm in diameter, rings of the same scale with collapsed interiors, linear ridges up to several meters long, and mounds that outline the cracks of 30–75 m desiccation polygons at the lake margin (vanden berg, 2019) (figure 2). the morphological diversity of the microbialites is presumably influenced by physical factors including substrate, bathymetry, tectonics, and hydrodynamics. correlations between these physical factors and microbialite growth suggest that microbialites tend to grow on underlying raised substrate (eardley, 1938; chidsey and others, 2015; bouton and others, 2016b; bouton and others, 2016a; vennin and others, 2019; vanden berg, 2019; kanik and others, 2020; baskin and others, 2022). at the mesoscale, the interior composition of the microbialites includes primarily clotted aragonite (posited to be of direct microbial origin (pace et al, 2016; vanden berg, 2019), as well as trapped and cemented ooids, artemia (brine shrimp) pellets, and some allochthonous grains (chidsey and others, 2015). many microbialites also include poorly-defined, laminated stromatolitic fabrics as a minor interior component. thus, the term microbialite since the structures comprise a mix of fabric types, instead of using more specific terms such as stromatolite, thrombolite, or leolite. radiocarbon (14c) dating of both solid carbonate and trapped organic material has yielded ages for microbialite material of 12.7–2.7 ka (figure 1a) (bouton and others, 2016b; bouton and others, 2016a; newell and others, 2017). the reservoir effect in the modern lake appears to be on the order of several hundred years (bowen and others, 2019; paradis and others, 2023), however, it may have been greater in the past (bowen and others, 2019), and carbonate formation in close association with groundwater may incorporate a reservoir effect of over 5000 years (homewood and others, 2022). thus, there is a rather high degree of uncertainty in microbialite radiocarbon ages. notwithstanding, to date, no modern ages have been measured from microbialite material, although dating is limited to only six microbialites from two locations at the northwest shore of antelope island, and none targeted periphyton-rich outer zones where modern carbonate precipitation appears to be happening (for example, pace and others, 2016). it also appears that microbialites form over thousands of years, with a range from 7.6–12.7 cal ka measured from organic material extracted from four zones within a single microbialite (newell and others, 2017). this covers a period when the surface elevation of great salt lake is poorly constrained within a rough range of l. wilcock, c.m. frantz, and m.d. vanden berg use of remote imagery to map microbialite distribution at great salt lake 4 figure 2. photographs of microbialites in and around great salt lake. (a–b) microbialites that grew at the boundaries of desiccation polygons at promontory point, north arm. note the bright/light surface color (photosynthetic microbial mats are absent) of partially submerged microbialites in halite-saturated north arm water. (c–d) microbialite reef at ladyfinger point on antelope island, showing transition from living periphyton to desiccated bright forms, (d) healthy mat and brine fly pupae visible on the surface of a collected microbialite sample; sample is roughly 14 cm across. (e–f) microbialites at bridger bay off antelope island, showing (f) collapsed centers; area shown is roughly 1 m across. (g–h) microbialite reef at buffalo point on antelope island, showing both exposed and partially-eroded structures, as well as (h) submerged structures with a dark, photosynthetic periphyton; area shown is roughly 0.8 m across. (i–j) large and elongate microbialites off of stansbury island, with thrombolitic crust. partially eroded crust visible in (j); area shown is roughly 1 m across. locations where each set of photographs were taken are shown as markers on the map in fig. 8. 5 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 1271–1285 masl (4170–4216 ft-asl) (oviatt and others, 2021) (figure 1a). regardless of their age and origin, microbialites play an important role in the modern great salt lake. the exposure of vast expanses of microbialites with historically low lake elevation levels is threatening their preservation and keystone function in the great salt lake ecosystem. great salt lake comprises distinct habitat types ranging from fresh– to brackishwater estuaries and wetlands where rivers enter the lake, to expansive mudflats and playas, to the hypersaline open water of gunnison bay (the north arm) and the south arm of great salt lake. great salt lake has historically supported a simple but hemispherically important ecosystem (figure 3). ten million birds rely on the lake, including 90% of the world’s eared grebes (podiceps nigricollis), two species of phalaropes (phalaropus lobatus and phalaropus tricolor), and large nesting colonies of american white pelicans (pelecanus erythrorhynchos) and california gulls (larus californicus) (conover and bell, 2020). the lake also supports an economically important brine shrimp cyst-harvesting industry, which supports global aquaculture (marden and others, 2020). great salt lake’s microbialites are a critical feature that supports this extreme ecosystem. microbialites, the lithified structures, are distinct from microbialite periphyton communities, which, in great salt lake, are robust, productive, and diverse microbial communities that blanket microbialite surfaces (pace and others, 2016; lindsay and others, 2017; kanik and others, 2020; ingalls and others, 2020). microbialite periphyton communities are conservatively estimated to be responsible for 30% of the lake’s primary productivity (wurtsbaugh and others, 2011; anderson and others, 2020; unpublished data by b. baxter and others, 2023), the remainder is attributed to planktonic algae. the significance of microbialites is as anchored, solid substrates with substantial relief above the surrounding sediment in the great salt lake benthos, providing islands of stability in otherwise mobile sediment where robust mats of photosynthetic microbes can develop. microbialites can contribute biomass to pelagic zones via sloughing, wave action, figure 3. a simplified great salt lake food web illustrating the importance of the lake’s microbialites and associated periphyton, which feed brine fly larvae and the occasional brine shrimp, which in turn feed higher trophic levels of the ecosystem. modified from frantz and others (2023) (licensed under cc4.0 and used with permission). 6 l. wilcock, c.m. frantz, and m.d. vanden berg use of remote imagery to map microbialite distribution at great salt lake and/or bioturbation (macintyre and melack, 1995; barrett, 2020; marden and others, 2020). brine shrimp (artemia franciscana) are filter feeders that prefer pelagic microalgae for nutrition, however, they will also graze on microbialite periphyton in shallow waters (caudell and conover, 2006; lindsay and others, 2019; brown and others, 2022). indeed, stable isotope and gut content dna evidence suggests that brine shrimp feed on microbialite surface communities during summer months (barrett, 2020; marden and others, 2020), presumably because the shrimp reduce the planktonic phytoplankton concentrations below the level at which they can efficiently feed (belovsky and others, 2011), necessitating a supplementary food source. microbialites are also a critical part of the brine fly (ephydra spp.) lifecycle, which depend on microbialites for habitat and food (collins, 1980; caudell and conover, 2006; belovsky and others, 2011; wurtsbaugh and others, 2011; conover and bell, 2020; brown and others, 2022), and are a critical nutritional source for both shorebirds and pelagic birds at great salt lake (conover and bell, 2020; sorensen and others, 2020). the overwhelming majority of brine flies appear to pupate on submerged microbialites (collins, 1980; wurtsbaugh, 2009), again, because they offer a stable benthic substrate. hatched brine fly larvae then feed primarily on microbialite periphyton communities (collins, 1980; barrett, 2020). in shore areas where submerged microbialites are nearby and salinity levels do not exceed 20% (which may be an upper survival limit for microbialite primary producers; lindsay and others, 2019), the dense clouds of hatched brine flies in late summer are remarkable; walking through a microbialite reef disturbs innumerable thousands of flies that rise from the surface of microbialites and ponded water in swarms. lake ecosystem models (for example, those described by belovsky and others, 2011; barrett, 2020) require accurate estimates of microbialite extent and relationships between lake elevation and the proportion of submerged vs. exposed microbialites. lake level fall and exposure of the lake’s microbialites great salt lake elevation levels have dropped to historic lows in recent years, the result of megadrought and overuse of water in the upstream watershed (null and wurtsbaugh, 2020), with profound consequences to the lake ecosystem. avian nesting grounds that were previously protected from predation as islands have become connected to outer lake shores, disrupting bird populations (kijowski and others, 2020; sorensen and others, 2020). increases in lake salinity have produced conditions that exceed levels at which keystone members of the ecosystem optimally survive and reproduce (baxter and butler, 2020; great salt lake salinity advisory committee, 2021). in addition, low elevation and consequent shoreline shift has exposed hundreds of kilometers of microbialite reefs, subjecting them—and their ecologically-important periphyton communities—to desiccation, negating their ecosystem function. recent work by frantz and others (2023) provided some hope in the face of current mass microbialite exposure, showing that exposed and desiccated microbialites can regain some of their periphyton community in relatively short order once re-submerged in healthy lake water. however, their study was limited to a brief period of recovery, well before thick, carbonate-rich mats began to reappear (which could take years to decades). their results also indicated that recovery is limited as lake level continues to fall and salinity continues to rise. in addition, they noted results that hint that individual microbialite areas harbor distinct strains of euhalothece, the primary microbialite phototroph; losing areas of reef may therefore disrupt natural microbial diversity and could make the lake’s microbialite-supported ecosystem less resilient to future change. furthermore, they showed that subaerially exposed microbialites are rapidly weathered. extended periods of exposure could reduce the height of microbialite reefs (and raise the surrounding sediment), diminishing their value as habitat for periphyton and brine fly larvae, even if lake levels rebound. mapping great salt lake’s microbialites the current threat to the lake’s microbialites with lake level fall, and consequent long-term impacts on the lake ecosystem, mean that management of great salt lake and its watershed requires a quantitative understanding of how different lake elevations affect microbialite exposure. this in turn depends on accurate maps of microbialite reef extent in great salt lake, as well as refined relationships between lake bathymetry and microbialite exposure. additionally, low lake levels and the exposure of the lake’s microbialites has presented new hazards for navigation of watercraft on the lake. accurate mapping of microbialite extent also has scientific value, as illustrated by several recent publications that have linked microbialite locations and extent to topographic features, faults, tectonics, wave energy, depth bands, and groundwater availability (bouton and others, 2016b; bouton and others, 2016a; vanden berg, 2019; baskin and others, 2022). 7 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 the first map of microbialite extent was from eardley (1938), who took advantage of a period of relatively low lake level in the mid-1930s to map them roughly from shore, as well as observing them at depths up to 1 m during “considerable travel” via a boat, the appropriately named hydrographer, near the shores of the lake and in transects between the lake’s islands. his paper includes both site and aerial photographs at various locations around the lake shore. importantly, he also noted that cores from previously conducted engineering studies indicated prior periods of microbialite formation in the lake in areas different from where he had observed them. he used a planimeter and his map to determine a rough microbialite reef area of 398 km2 (154 mi2) within the lake (figure 4). due to limited mapping technology and limited field observations, eardley’s map largely missed microbialite reefs on the western side of the lake, as well as deeper-water areas, whereas extents on the east side of the north arm are overestimated. overall, eardley underestimated the extent of great salt lake microbialites. for his 2014 ph.d. dissertation, baskin (baskin, 2014) produced the first major update to eardley’s map, utilizing single-beam sound-velocity soundings obtained during his work producing digital bathymetric surveys of the lake with the united states geological survey (usgs) (baskin and allen, 2005; baskin and turner, 2006). his method for identifying microbialites involved a calculation of rugosity from the sounding data that was truthed in select high-rugosity areas using dual-frequency 2d side scan sonar, sweptfrequency chirp sub bottom profiles, and videography (when lake visibility permitted), as well as in situ sampling in known microbialite locations. the extents identified in his dissertation were then updated and refined with the publication of baskin and others (2022). this newer publication identified an area of ~1000 km2 (~390 mi2) of putative microbialite reef, with >700 km2 (270 mi2) in the south arm and >300 km2 (~120 mi2) in the north arm (figure 4), nearly tripling the extent mapped by eardley (1938). in his thesis, baskin also noted the effect of the railroad causeway, completed in 1959, that bisected the lake and cut off the north arm from most of the lake’s freshwater input, causing it to become rapidly salt-saturated and killing off the euhalothece-based periphyton on north arm microbialites (this was also noted by post, 1977, and verified with dna sequencing by lindsay and others, 2017). although extensive, baskin’s map was largely based on indirect data; due to time and resource constraints he was only able to verify the presence of benthic microbialites in limited areas of his reported mapped extent. vanden berg (2019) produced an alternative map of microbialite extent using google earth imagery and limited field mapping, yielding a microbialite reef aerial extent of 680 km2. however, the map and extent estimates were limited by the availability of clear -water imagery and stated the need for further field verification. bouton and others (2020) further amended microbialite extent estimates by merging the eardley (1938) and baskin (2014) maps and adding additional refinement based on limited remote imagery of western antelope island from bouton and others (2016a), yielding an expanded (and overestimated) microbialite reef aerial extent of 1261 km2 (487 mi2). in sum, maps of microbialite reef extent in the literature to date have given conflicting and highly variable results (figure 4). recent low lake elevations and increasing resolution of satellite and aerial imagery have made microbialite mapping via remote imaging more powerful and accurate than ever before. water column visibility in the lake varies greatly with season, biological activity, and weather, however, during clear-water periods the secchi disk depth typically exceeds 3 m (10 ft), making the lake bottom visible from aerial view in all but the deepest portions of the lake (belovsky and others, 2011). microbialites are visible to depths in excess of 4 m (13 ft) in some high-visibility images, a fact that several studies have utilized to identify extents of microbialites against the lake bed (bouton and others, 2016a; vanden berg, 2019). advantages of using remote imagery over field-based mapping include the ability to quickly map large regions across the full extent of the lake (vs. transects or areas only accessible from shore), and that dry, shallow-water, and deep microbialites can all be mapped using the same method. the varied estimates of microbialite extent from prior literature (table 1) adds a large element of uncertainty to estimates of overall microbialite productivity, microbialite exposure, and other factors influencing the management of great salt lake. thus, our study attempted to improve on previous estimates by (1) mapping microbialites using satellite imagery, taking advantage of historic low lake level and improved spatial and temporal resolution of available images, (2) confirming (or refuting) the presence of suspected microbialite areas from prior mapping efforts via aerial imagery and field checks, and (3) generating shapefiles of microbialite reef extent that can be used in quantitative estimates of microbialite extent and exposure. here, we present our results, which include the most detailed map of great salt lake microbialite extent to date and a model of microbialite exposure at different lake elevations. l. wilcock, c.m. frantz, and m.d. vanden berg use of remote imagery to map microbialite distribution at great salt lake 4 figure 4. comparison of previously published microbialite reef extent maps for great salt lake with our mapped reef extent. (a) microbialite reef areas mapped by eardley (1938; in purple), baskin and others (2022; in blue), and this study (yellow), highlighting areas of overlap and major differences. (b) quantified comparison of mapped reef areas in the three studies. darker vs. lighter colors in the plot for this study indicate regions of high vs. low confidence. 9 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 methods mapping microbialites using satellite imagery data acquisition positive identification of microbialites through the application of remote sensing required highresolution imagery with sufficient temporal resolution to permit analysis of areas of interest during favorable periods (i.e., periods without obscuring cloud cover, with low lake elevations, and with good water clarity). imagery was collected through esri’s world imagery wayback (ewiw) archive. ewiw is a digital archive of published world imagery since 2014 that is stored as layer files that can be downloaded or viewed online through arcgis’s living atlas. the current extent of the great salt lake covers over 4000 km2 within the great salt lake basin (within the quad 40.6 –41.8°n, 111.8–113.2°w). imagery for the region is collected via multiple satellite constellations at different temporal sequences that are location-dependent. north and south arms of great salt lake required sets of time series imagery that often come from different capture dates (table 2). ewiw acquires imagery via landsat, usda naip, terracolor, digital globe, geoeye ikonos and aerogrid at 0.6–15 m spatial resolution depending on location and provider. dates were selected to provide optimal belowwater visibility, with favorable atmospheric conditions (especially low cloud cover), clear water periods (during the absence of water turbulence or algal blooms), and relatively low lake levels (permitting visibility in deeper areas of the lake), allowing good visual records of changing microbialite reef exposure (figure 5). the analysis over multiple time points was vital for distinguishing loose debris from true reef, as illustrated in figure 6. google earth pro (gep) was also utilized to compare and contrast visible reef zones with ewiw imagery. gep utilizes landsat and copernicus satellite constellations for imagery collection. dates of available archival gep imagery vary; imagery from 2016–2022 provided the best clarity for positive or negative identification of microbialites. imagery in gep varies based on location and scale, with each view of lake locations utilizing several remote sensing sources and acquisition dates. high-resolution historical imagery was collected from ewiw and downloaded as layer files. once imported into arcgis pro, each layer file was used for side-by-side comparison of microbialite structures. this side-by-side analysis of archived ewiw and gep imagery was used to digitize areas that could be positively identified as reef zones via remote sensing. identification and mapping to develop criteria for microbialite reef identification, we first compared characteristics of known reef zones (from field studies by the authors) to our remote sensing imagery (figure 7). we identified three reliable patterns for identifying microbialites in remote imagery. reference method shortcomings and uncertain es mapped microbialite extent (km2) south arm north arm total eardley, 1938 field verifica on from shore and by boat limited to primarily nearshore areas confirmed in the field, missed areas of deeper microbialite reef and areas in the western por ons of the lake 117 160 277 260 baskin and others, 2022 rugosity from acous c soundings during bathymetric surveys, par ally confirmed in the field indirect measure with limited field confirma on 700 654 300 446 1000 1099 vanden berg, 2019 remote imagery limited image availability, limited field verifica on 56 92 147 bouton and others, 2020 merged prior maps with addional areas from remote imagery reported in bouton and others, 2016a inherited uncertain es from prior work, assumed variable regions were due to burial vs. rip-up clasts 1261 this study (high confidence) remote imagery limited field verifica on, some deepwater areas could not be mapped 288 360 648 table 1. summary of prior attempts to map lakebed microbialites in great salt lake. where given, reported values are non-italicized while values inferred from traced shapefiles are italicized. 10 l. wilcock, c.m. frantz, and m.d. vanden berg use of remote imagery to map microbialite distribution at great salt lake first, “healthy,” submerged microbialites appear dark green in remote imagery and stand out against the brighter carbonate sediment background (figs. 7a & 7d). in some instances of dark green submerged substrate, microbialite reefs were indistinguishable from loose microbialite debris (figure 6) in single images; for such regions, we compared images from at least three different dates to look for evidence of mobility, with only stationary features mapped as reef. second, as lake levels fall, microbialites become exposed and “bleach” (frantz and others, 2023), causing exposed reef areas to appear bright in partiallyexposed reef areas. our second identified pattern was that of white reef areas (bleached microbialites) with patterned high-relief mounds (for example, figs. 7c & d). image capture date world imagery date loca on provider resolu on (m) accuracy (m) 2014-06-29 2015-07-08 south arm naip 1 6 2014-08-31 2015-07-08 north arm naip 1 6 2016-06-26 2017-05-03 south arm naip 1 6 2016-07-15 2017-05-03 north arm naip 1 6 2016-05-07 2018-01-08 north arm digital globe 0.5 10.2 2013-08-29 2018-01-08 south arm digital globe 0.5 10.2 2022-05-07 2022-11-02 south arm maxar (geo1) 0.46 5 2021-10-15 2022-11-02 north arm maxar (wv02) 0.5 5 2021-04-08 2022-12-14 south arm maxar (geo1) 0.46 5 2021-10-15 2022-12-14 north arm maxar (wv02) 0.5 5 table 2. summary of remote imagery utilized for this study. image capture date is the date satellite images were captured, while world imagery date is a date of availability in arcgis for the set of images. figure 5. comparison of satellite images of a specific location at northern antelope island (41.06°, -112.26°) using different image dates. in all images, the thin, white dashed line shows the area outlined as microbialite reef in this study. (a) microbialite reef can be seen as a dark green submerged region in june 2014 (esri world imagery wayback). (b) in may 2016, visibility of the reef was limited due to poor water clarity and higher lake elevation (esri world imagery wayback). (c) in september 2018, part of the visible reef was obscured due to image distortion and resolution issues (google earth pro). (d) waves on the lake in may 2020 obscure the reef (google earth pro). (e) waves and light reflection again obscure parts of the reef, with image stitching artifacts obscuring other portions (esri world imagery wayback). (f) exposed microbialite reef appear as bright/light regions during low lake level in may 2022 (esri world imagery wayback). 11 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 finally, microbialites tend to form on the perimeters of “megapolygons”—polygonal structures roughly 30–75 meters in diameter (vanden berg, 2019) (figs. 7a & 7b). thus, megapolygons are our third identified pattern. in contrast, zones of smaller desiccation–related polygons, averaging only 4–9 meters, are present along shoreline areas at higher elevation and are not associated with microbialites (vanden berg, 2019). these smaller polygons can be ephemeral, appearing and disappearing with changes in lake level. in some areas, particularly north and northwest of hat island (112.586°w 41.071°n), we identified broad regions of megapolygons (some quite faint) at elevations above 4195 ft-asl, however, we excluded these from our map due to lack of field verification and their anomalously high elevations; if microbialites are found associated with these megapolygons, they might belong to an older generation. regions positively identified as containing microfigure 6. examples of mobile debris. (a) field photograph of loose carbonaceous microbial mat debris between actual microbialite mounds at buffalo point in august 2021. (b) google earth pro remote image showing a dark green region of potential microbialite reef in the southwest arm of the lake (40.983°, -112.709°) on 2019-08-17, and (c) google earth pro remote image of the same location on 2015-06-27 showing shifted mobile debris. scale bars in (b) and (c) are both 1 km. figure 7. examples of field-verified microbialite reef areas identified from remote imagery in the south arm of great salt lake. (a) dark green submerged microbialites and bright bleached megapolygons indicate the presence of microbialites in a nearshore area in the south arm of the lake (41.073°, -112.573°). (b) submerged desiccated microbialiteedged polygons in the north arm of the lake (41.249°, -112.533°). (c) bright exposed and desiccated microbialites stand out against green lake water at a site near the antelope island marina (41.064°, -112.237°). (d) partially submerged microbialites between buffalo point and white rock bay (41.033°, -112.275°). scale bar in all images is 100 m. image locations are shown as markers on the map in fig. 8. images from google earth pro. 12 l. wilcock, c.m. frantz, and m.d. vanden berg use of remote imagery to map microbialite distribution at great salt lake bialite reefs were mapped in esri’s arcgis pro. separate feature classes were created for the north and south arms and were digitized using arcgis pro by tracing shapes over downloaded imagery. feature classes as well as bathymetric layers were imported and projected as utm nad83 zone 12n to minimize distortion and maximize location accuracy. areas were initially digitized in large zones before being refined to greater resolution in a second stage of processing. comparison to prior work areas mapped by prior studies were given extra attention in our analysis, with maps by eardley (1938), vanden berg (2019), bouton and others (2020), and baskin and others (2022) providing a framework for the mapping efforts described in this study (figure 4). some regions identified as reef zones by baskin and others (2022) were not able to be conclusively analyzed using remote imagery due to their occurrence in deeper areas of the lake. we included some of these regions from baskin in our map as low-confidence regions. field verification many identified reef sites were confirmed with field verification, particularly in accessible shoreline areas (figure 8); these regions are denoted as highconfidence regions in our map. the western shores of the lake are difficult to access due in part to military restrictions and private land ownership, thus most sites on the west side of the lake have not been field verified. identified reef sites not yet confirmed with field verification are denoted as low-confidence regions except for those associated with megapolygons, which were classified as high-confidence even in the absence of field verification. lake elevation-exposure model in order to develop a model of microbialite exposure at different lake elevations, we used shapefiles for the mapped microbialites and determined overlap with lake bathymetry shapefiles (1 ft intervals) imported from baskin and allen (2005) and baskin and turner (2006). however, caution should be exercised when using the historical bathymetry data, especially in the nearshore environment: modern observations during extreme low lake level indicate that these contours are significantly incorrect in several nearshore environments around the lake. inaccuracies in the bathymetric data will create inaccuracies in the exposure models presented in this study, but currently this is the only published bathymetric data available. microbialite reef area shapes were combined in distinct layers for the north vs. south arm of the lake, since the two arms can have independent water surface elevation levels and can be managed separately for ecosystem function. digitized microbialite reef zones were split based on bathymetric data. these clipped zones were used to identify areas of exposure as lake levels decline. areas of mapped microbialite reef at elevations above bathymetric lines were considered exposed at that lake elevation, whereas areas of microbialite reef at or below bathymetric lines were considered submerged. the curve fit least-squares function in the scipy.optimize python package (virtanen and others, 2020) was used to generate logistic regression models parameterized to fit the area vs. bathymetry elevation values for each arm of the lake using the least squares method. results microbialite reef extent our remote imagery-based mapping of microbialite extent indicates 360 km2 (139 mi2) of microbialite reef between 1271.6 and 1280.5 masl (4172– 4201 ft-asl) in the south arm of great salt lake, of which 45% are high-confidence regions. in the north arm of the lake, we mapped 288 km2 (111 mi2) of microbialite reef in the same elevation band, of which 74% are high-confidence regions confirmed with field observation (figure 8). the distributions of mapped microbialites by elevation were similar in the north and south arms (figure s1), although our mapped region in the north arm was limited by limited field verification, poor water visibility, and image resolution. our mapped extent was somewhat similar with the baskin and others (2022) map, with several important differences. first, we were able to map microbialites in exposed shore environments that were inaccessible by boat and therefore unable to be mapped sonographically by baskin, thus, our map extends to higher elevations than the baskin and others (2022) map (for example, bottom left of figure 9c). second, in some regions, areas mapped by baskin extended deeper into the lake than what we found, for example, on the western shore of the lake (figure 9b–c). third, our map is more spatially refined (figure 9e). also, some regions mapped by baskin were exposed as dry shoreline in recent years, with no apparent microbialites present (for example, figure 9d). most (95%) of the microbialites that we mapped lie in an elevation band between 1274.0 and 1278.6 13 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 8. mapped extent of microbialites in great salt lake (this study) showing regions of high confidence of microbialite occurrence (areas confirmed with field verification or presence of megapolygons) and regions mapped at low confidence of microbialite occurrence (areas of apparent microbialite reef in remote imagery). stars indicate areas where field verification of microbialite reef existence (or non-existence) was verified. triangles mark the approximate locations of photograph sets shown in fig. 2. circles mark the locations of remote imagery shown in fig. 7. basemap imagery provided by earthstar geographics. 14 l. wilcock, c.m. frantz, and m.d. vanden berg use of remote imagery to map microbialite distribution at great salt lake masl (4180–4195 ft-asl) (figure 10). several notable deeper-water outlier areas are bounded by active quaternary fault zones (figure s2). elevation-exposure model our findings for microbialite exposure at different lake elevations are summarized in table 3 and figs. 11–12. fitting a logistic regression line (equation 1) using the least-squares method to the lake elevation (elev, in masl or ft-asl) vs. microbialite exposure data (in km2 or mi2) gave r2 values ≥ 0.995 for all models (figure 12). equation 1: in equation 1, is the area (in km2 or mi2) of microbialites exposed at a given lake elevation (elev, in masl), where l, k, x0, and b are model parameters defined in table 4. discussion refined map of microbialite reef extent for great salt lake our remote imagery-based map of microbialite extent yielded an extent of microbialites between the lower and upper bounds of prior work (figure 4): at both low and high confidence levels, we mapped significantly more microbialite area than eardley (1938), but substantially less than what was mapped by baskin and others (2022). because it relied on limited field observation and rough mapping tools available at the time, the eardley (1938) map represents an understandable underestimate of microbialite extent. meanwhile, the baskin and others (2022) map covered the entire lakebed in relatively high resolution, however, by relying on indirect measurements of lake-bottom rugosity, it could have overestimated true microbialite extent. in general, our map refines the spatial extent of reefs identified by baskin: 86% of our mapped regions were also mapped by baskin, for both our high and low confifigure 9. example detail areas where mapped microbialite extents in this study differed significantly from baskin and others (2022). (a) mapped microbialite extents in baskin (blue) vs. this study (yellow) showing areas of detail (b–e). (b) region along the northeastern lakeshore mapped as having microbialites by baskin where we were unable to find evidence of microbialites in remote imagery or via field checks. base image from maxar 2015-07-08. (c) area along the western shore of the lake where we identified a region of higher elevation microbialites visible in remote imagery but unmapped by baskin. the baskin map also extends into deeper water than we were able to confirm. base image from maxar 2015-04-27; mid-image color changes is an imagery artifact. (d) area at the southwestern shore of the lake where the baskin map includes microbialites where we only observed regions of mobile clasts. base image from maxar 2021-10-16. (e) region off the northwest shore of antelope island where high-resolution imagery from esri world imagery wayback and google earth pro allowed for more precise mapping of microbialite reef zones in our study relative to the baskin map. base image from maxar/earthstar geographics 2022-05-08. white scale bars in areas of detail (b–e) are all 1 km. 15 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 10. histograms of microbialite reef area identified at high and low confidence in different 1 ft elevation bands (labels show the lower bound of the band). (a) north arm (na). (b) south arm (sa). (c) both arms. l. wilcock, c.m. frantz, and m.d. vanden berg use of remote imagery to map microbialite distribu on at great salt lake 16 eleva on band ( ‐asl) area of mapped microbialite reef (km2) total area exposed at lower eleva on bound (km2) high confidence low confidence high confidence all mapped (high + low conf.) north arm south arm north arm south arm    north arm south arm whole lake north arm south arm whole lake 4172  –  4173  0.00  0.01  0.00  0.79  212.0  162.6  374.6  288.45  360.4  648.8  4173  –  4174  0.00  0.00  0.00  0.03  212.0  162.6  374.6  288.45  359.6  648.0  4174  –  4175  0.00  0.02  0.00  1.06  212.0  162.6  374.6  288.45  359.5  648.0  4175  –  4176  0.01  0.01  0.00  2.90  212.0  162.5  374.6  288.45  358.4  646.9  4176  –  4177  0.04  0.06  0.00  1.80  212.0  162.5  374.5  288.43  355.5  644.0  4177  –  4178  0.09  0.08  0.00  2.04  212.0  162.5  374.4  288.39  353.7  642.1  4178  –  4179  0.09  0.15  0.00  5.55  211.9  162.4  374.3  288.30  351.6  639.9  4179  –  4180  0.15  1.17  0.00  9.65  211.8  162.2  374.0  288.21  345.9  634.1  4180  –  4181  0.59  1.81  0.01  10.07  211.6  161.1  372.7  288.07  335.0  623.1  4181  –  4182  6.33  4.35  0.01  12.28  211.1  159.3  370.3  287.47  323.2  610.6  4182  –  4183  7.30  1.45  0.82  14.94  204.7  154.9  359.6  281.12  306.5  587.6  4183  –  4184  9.36  7.41  1.16  15.21  197.4  153.5  350.9  273.00  290.1  563.1  4184  –  4185  16.90  3.66  3.78  14.49  188.1  146.0  334.1  262.48  267.5  530.0  4185  –  4186  14.32  14.12  5.30  12.67  171.2  142.4  313.5  241.80  249.4  491.2  4186  –  4187  15.62  11.27  6.19  9.90  156.8  128.3  285.1  222.18  222.6  444.8  4187  –  4188  13.67  29.04  6.05  9.06  141.2  117.0  258.2  200.37  201.4  401.8  4188  –  4189  12.73  20.46  6.09  10.36  127.5  88.0  215.5  180.64  163.3  344.0  4189  –  4190  13.99  20.04  8.26  12.43  114.8  67.5  182.3  161.82  132.5  294.3  4190  –  4191  14.65  11.45  8.91  12.21  100.8  47.5  148.3  139.58  100.0  239.6  4191  –  4192  23.07  10.32  9.15  12.43  86.2  36.0  122.2  116.02  76.4  192.4  4192  –  4193  23.11  10.01  8.31  16.06  63.1  25.7  88.8  83.80  53.6  137.4  4193  –  4194  21.04  6.50  5.52  7.30  40.0  15.7  55.7  52.38  27.5  79.9  4194  –  4195  14.96  5.71  4.63  4.28  19.0  9.2  28.1  25.82  13.7  39.5  4195  –  4196  2.91  0.71  2.15  0.24  4.0  3.5  7.5  6.22  3.7  10.0  4196  –  4197  0.77  2.19  0.07  0.03  1.1  2.8  3.8  1.16  2.8  4.0  4197  –  4198  0.25  0.54  0.00  0.00  0.3  0.6  0.9  0.32  0.6  0.9  4198  –  4199  0.07  0.00  0.00  0.00  0.1  0.0  0.1  0.07  0.0  0.1  4199  –  4200  0.00  0.01  0.00  0.00  0.0  0.0  0.0  0.00  0.0  0.0  4200  –  4201  0.00  0.01  0.00  0.00  0.0  0.0  0.0  0.00  0.0  0.0  table 3. mapped microbialite reef area in different elevation bands, and area of microbialite exposure when lake level reaches the lower elevation bound. 17 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 11. map of microbialite reef areas (this study) correlated with lake bathymetry, highlighting the areas of microbialite reef exposed at different lake surface elevations (in ft-asl). basemap imagery provided by earthstar geographics. 18 l. wilcock, c.m. frantz, and m.d. vanden berg use of remote imagery to map microbialite distribution at great salt lake figure 12. relationship between lake elevation and total cumulative microbialite exposure in great salt lake. data points for each elevation band that we mapped are shown as points along with corresponding logistic regression best-fit lines. shaded areas represent the range of standard error for the regression models. the dashed vertical line marks the lake elevation at the autumn 2022 minimum (4188.5 ft-asl). (a) microbialites mapped in the north arm of great salt lake at high (light) and high+low (dark) confidence. (b) microbialites mapped in the south arm of great salt lake at high (light) and high+low (dark) confidence. (c) values for the whole lake, with mapped microbialites at high (light) and high+low (dark) confidence. 19 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 dence maps. however, baskin mapped ~350 km2 (135 mi2) more microbialite areas than we could confirm, largely in deep-water areas of the lake. there are several key differences between our map and the baskin map that warrant future field verification. first, our technique allowed for mapping of microbialites in shore environments that were not navigable and therefore unmapped by baskin, for example, in an area north of lakeside where we identified desiccation megapolygons (figure 9b). second, areas mapped by baskin frequently extended deeper into the lake than our remote imagery-based approach permitted, for example, on the western shore of the lake, and in the area between antelope island and fremont island (figs. 4, 9b–c). we did not include these deeper-area regions of putative reef mapped by baskin in our map or elevation-exposure model, however, we cannot rule out that they exist. also, our map only accounts for consistently unburied microbialites, which are more likely to contribute to lake productivity than intermittently buried microbialites, which could have been included in the baskin and others (2022) map. heavily eroded microbialites may also have been missed by our map. lake elevation and microbialite exposure during the autumn 2022 historic lake lowstand of 1276.7 masl (4188.5 ft-asl), we estimate (from microbialites mapped at both high and low confidence in this study) that >294 km2 (114 mi2, or >45%) of the lake’s microbialites were exposed, >133 km2 (51 mi2) in the south arm (>37% exposure), and 162 km2 (63 mi2) in the north arm (>56% exposure). microbialites in the lake’s north arm no longer support a robust microbialite surface community because of the arm’s high salinity levels (lindsay and others, 2019), thus, their exposure or submergence likely does not have much influence on the support of higher tropic levels in the great salt lake food web. in the south arm, recent evidence suggests that microbialite photosynthetic (periphyton) communities can survive months of subaerial exposure, and that re-submerged microbialites appear to be rapidly recolonized by lake water microorganisms (frantz and others, 2023). however, subaerially exposed microbialites cannot contribute to the benthic or planktonic food chains in the lake. additionally, areas of microbialites that experienced frequent exposure in the past half century never fully redeveloped a healthy periphyton (marked by thick gelatinous mats) even when re-submerged for periods of several seasons to years, indicating that the damage caused by prolonged exposure is long-lasting. it is also important to note that microbialites in the hypersaline north arm of the lake also lack the robust mats of primary producers that are present in “healthy” microbialites (lindsay and others, 2017); this is one of the reasons we clearly separate our maps of north vs. south arm microbialites. finally, exposed microbialites are subjected to rapid weathering, and it could take decades or even centuries for the raised mounds that represent stable oases in an otherwise shifting lake benthos to re-form. thus, the consequences of long-term subaerial exposure of the lake’s microbialites are profoundly concerning for the lake ecosystem. even in the short term, there are ecosystem consequences of microbialite exposure. if microbialite periphyton communities conservatively represent 30% of primary production in great salt lake, the expologis c regression model parameters metric units (masl, km²) logis c regression model parameters imperial units ( -asl, mi²) arm confidence r² l k x₀ b l k x₀ b north arm high 0.9954 229 ± 6 1.14 ± 0.07 1277.0 ± 0.1 -14 ± 5 88 ± 2 0.35 ± 0.02 4189.6 ± 0.2 -5.4 ± 1.8 north arm high+low 0.9967 308 ± 6 1.23 ± 0.06 1277.0 ± 0.1 -16 ± 5 119 ± 2 0.38 ± 0.02 4189.8 ± 0.2 -6.3 ± 1.9 south arm high 0.9992 163 ± 1 1.71 ± 0.04 1276.7 ± 0.0 0 ± 1 63 ± 0 0.52 ± 0.01 4188.5 ± 0.1 -0.1 ± 0.4 south arm high+low 0.9988 376 ± 5 1.12 ± 0.04 1276.4 ± 0.0 -14 ± 3 145 ± 2 0.34 ± 0.01 4187.5 ± 0.1 -5.5 ± 1.2 whole lake 0.9988 390 ± 4 1.32 ± 0.04 1276.8 ± 0.0 -12 ± 3 151 ± 2 0.40 ± 0.01 4189.1 ± 0.1 -4.7 ± 1.3 whole lake high+low 0.9986 684 ± 9 1.14 ± 0.04 1276.7 ± 0.0 -31 ± 7 264 ± 4 0.35 ± 0.01 4188.6 ± 0.1 12.0 ± 2.6 high table 4. logistic regression model results for microbialite exposure area at different lake elevations. to aid in the use of models for management, values are presented for use of both metric units (masl for lake elevation, km2 for area of exposed microbialites) and imperial units (ft-asl for lake elevation, mi2 for area of exposed microbialites). 20 l. wilcock, c.m. frantz, and m.d. vanden berg use of remote imagery to map microbialite distribution at great salt lake sure of ~ 40% of them in the lake’s south arm may have equated to a > 10% reduction in overall lake primary production in summer 2022 compared to “healthy” lake elevations (when microbialites are fully submerged). if one assumes that the bulk of microbialite-supported primary productivity occurs in relatively shallow water (i.e., the year-round photic zone), it is possible that the relative aerial extent of microbialites that occupy this zone has been relatively stable over the past several years of lake level fall, however, further lake level decline would substantially decrease the area of productive microbialites. also significant to the ecosystem is the substantial decrease in ephydra pupa anchor sites that occurs when microbialites become subaerially exposed. the greatest change in submerged microbialites occurs between 1275.6 and 1278.0 masl (4185–4193 ft-asl; figure 12) because of the large expanses and high density of microbialites in this zone (figure 10). the lower bound for the lake elevation target range for management of 1279.5 masl (4198 ft-asl) (utah dnr forestry, 2013) ensures that nearly all of the lake’s microbialites are submerged. at 1278 masl (4193 ft-asl), 88% are submerged, while at 1275.6 masl (4185 ft-asl), only 24% remain submerged. additionally, at lake elevation levels below ~1277 masl (4190 ft-asl), microbialite community health becomes threatened not only by exposure, but by salinity. at salinity levels above 15%, the primary productivity of euhalothece—and, thus, microbialite-associated productivity—declines (lindsay and others, 2019); this corresponds to a lake elevation of roughly 1277 masl (4191 ft-asl). thus, due both to microbialite exposure and high salinity levels, elevations above 1277 masl (4191 ft-asl) should be a minimum for lake management with respect to microbialite-supported ecosystem survival, whereas elevations above 1278.6 masl (4195 ft-asl) keep nearly all of the lake’s microbialites submerged. limitations of this study although we believe our map is a significant improvement over previously published maps of microbialite extent, it has several limitations and caveats. first, our map is limited to visible reef areas. in regions where remote imagery is low resolution, we were unable to confidently map microbialites. we were also unable to conclusively confirm or refute microbialite reef areas in deep-water portions of the lake (generally, below 1275 masl, or 4183 ft-asl, although this varied somewhat by remote imagery available), where water obscures reflected light. these deep-water portions of the lake represent an area of 1800 km2 (~700 mi2) and include 232 km2 (90 mi2) of microbialite reef mapped by baskin and others (2022); we cannot rule out the existence of microbialites above surrounding lake sediment at depths below 1275 masl (4183 ft-asl), but we were only able to confirm the probable existence of microbialites in 53 km2 (20 mi2) of that area based on remote imagery and the methods of our study. this could account for some, but not all discrepancies between the baskin map and ours. this caveat to our study could be remedied with a comprehensive field verification campaign. deep-water areas may need to be verified by divers. our study could also be used to help refine baskin’s benthic rugosity-based mapping algorithm (baskin, 2005). second, we excluded regions of reef that were not consistently visible in remote imagery. we did this to exclude areas of shifting microbialite debris/rip-up clasts. however, the change in visibility could also be due to shifting ooid sands covering up and then reexposing areas of active reef (as noted by bouton and others, 2016). these regions of reef could still, when exposed, contribute to primary production in the lake. roughly 59 km2 (23 mi2) of the lakebed we analyzed in this study comprised regions of variable brightness, i.e., either mobile clasts or varied exposure/covering by surrounding sediment, and it was not possible to distinguish mobile clasts from shifting sediment obscuring true reef areas. third, our model of microbialite exposure vs. lake elevation is based on the bathymetry of baskin and allen (2005) and baskin and turner (2006), which was limited spatially to 1-km transects in the navigable portions of the lake (baskin, 2005; baskin, 2006). thus, the bathymetry, especially in the elevation band of 1276.5–1278.6 masl (4188–4195 ft-asl), which corresponds to one of the greatest expanses of microbialite reef (figure 10), is poorly constrained, limiting the accuracy of our model. bathymetry in this band can be improved with detailed lidar mapping, work that is currently being explored and, we hope, done more extensively in the near future. finally, prolonged subaerial exposure of the lake’s microbialites results in their rapid weathering (frantz and others, 2023), thus, microbialite extents at higher elevation bands are subject to change (decrease) during periods of low lake elevation. additional research is required to quantify and model rates of microbialite weathering. summary we mapped 649 km2 (251 mi2) of microbialite reef in great salt lake by leveraging low lake levels and recent availability of high-resolution remote imagery. of that, 375 km2 (145 mi2) were either field21 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 verified or were identified as megapolygons, which are linked to microbialites in great salt lake (vanden berg, 2019). we believe that our map of microbialite extents refines previously published maps. we have also produced shapefiles of microbialite extent at different lake elevations (supplemental materials). our model of microbialite exposure vs. lake elevation can be used to inform great salt lake management: 1278.6 masl (4195 ft-asl) should be considered as a critical minimum lake elevation (with the understanding that higher lake levels provide greater protection) with respect to microbialites; at this depth, 98% of the lake’s microbialites are submerged. during the historic lowstand in autumn 2022 of 1276.7 masl (4188.5 ft -asl), we estimate that >37% of the microbialites in the south arm of the lake were subaerially exposed, representing substantial damage to benthic primary productivity (which was likely already threatened by high salinity levels) and ephydra larva habitat. acknowledgements lw was supported by nsf rise #1801760 to elizabeth balgord. cf was supported by nsf ear #1826869. mvb was supported by the utah geological survey. we thank ryan frazier and michael hernandez for their extensive help with gis aspects of this project, and several reviewers for detailed and helpful comments on a prior version of this manuscript. references anderson, n.l., barrett, k.l., jones, s.e., and belovsky, g.e., 2020, impact of abiotic factors on microbialite growth (great salt lake, utah, usa): a tank experiment: hydrobiologia, v. 847, no. 9, p. 2113–2122, doi: 10.1007/s10750-020-04235-9. barrett, k.l., 2020, microbialite communities and food web linkages in great salt lake: notre dame, university of notre dame, ph.d. dissertation, 195 p. baskin, r., 2005, calculation of area and volume for the south part of great salt lake, utah: usgs open-file report 2005–1327, 7 p. baskin, r.l., 2006, calculation of area and volume for the north part of great salt lake, utah: usgs open-file report 2006–1359, 6 p. baskin, r.l., 2014, occurrence and spatial distribution of microbial bioherms in great salt lake, utah: salt lake city, university of utah, ph.d. dissertation, 203 p. baskin, r.l., and allen, d.v., 2005, bathymetric map of the south part of great salt lake, utah, 2005: u.s. geological survey scientific investigations map 2894, scale 1:24,000. baskin, r.l., della porta, g., and wright, v.p., 2022, characteristics and controls on the distribution of sublittoral microbial bioherms in great salt lake, utah: implications for understanding microbialite development: the depositional record, v. 8, no. 1, p. 39–66, doi: 10.1002/dep2.159. baskin, r.l., and turner, j., 2006, bathymetric map of the north part of great salt lake, utah, 2006: u.s. geological survey scientific investigations map 2954, scale 1:24,000. baxter, b.k., and butler, j.k. 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publishing, p. 263–309, doi: 10.1007/978-3-03040352-2_9. utah dnr forestry, 2013, great salt lake comprehensive management plan: utah department of natural resources division of forestry, fire & state lands record of decision 13-0315–1, 391 p. vanden berg, m.d., 2019, domes, rings, ridges, https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx 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http://linkinghub.elsevier.com/retrieve/pii/0277379114005071 http://linkinghub.elsevier.com/retrieve/pii/0277379114005071 https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx 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https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx https://www.zotero.org/google-docs/?w9ogrx 24 l. wilcock, c.m. frantz, and m.d. vanden berg use of remote imagery to map microbialite distribution at great salt lake and polygons: characteristics of microbialites from utah’s great salt lake (l. birgenheier & h. harper, eds.): the sedimentary record, v. 17, no. 1, p. 4–10, doi: 10.2110/sedred.2019.1.4. vennin, e., bouton, a., bourillot, r., pace, a., roche, a., brayard, a., thomazo, c., virgone, a., gaucher, e.c., desaubliaux, g., and visscher, p.t., 2019, the lacustrine microbial carbonate factory of the successive lake bonneville and great salt lake, utah, usa (a. brasier, ed.): sedimentology, v. 66, no. 1, p. 165–204, doi: 10.1111/sed.12499. virtanen, p., gommers, r., oliphant, t.e., haberland, m., reddy, t., cournapeau, d., burovski, e., peterson, p., weckesser, w., bright, j., van der walt, s. j., brett, m., wilson, j., millman, k. j., mayorov, n., nelson, a. r. j., jones, e., kern, r., larson, e., carey, c. j., polat, i., feng, y., moore, e. w., vanderplas, j., laxalde, d., perktold, j., cimrman, r., henriksen, i., quintero, e. a., harris, c. r., archibald, a. m., ribeiro, a. h., pedregosa, f., van mulbregt, p., and scipy 1.0 contributors, 2020, scipy 1.0: fundamental algorithms for scientific computing in python: nature methods, v. 17, no. 3, p. 261-272. wurtsbaugh, w.a., 2009, biostromes, brine flies, birds, and the bioaccumulation of selenium in great salt lake, utah: saline lakes around the world: unique systems with unique values. natural resources and environmental issues, vol xv, v. 15, p. 1–15. wurtsbaugh, w.a., gardberg, j., and izdepski, c., 2011, biostrome communities and mercury and selenium bioaccumulation in the great salt lake (utah, usa): science of the total environment, v. 409, no. 20, p. 4425–4434, doi: 10.1016/ j.scitotenv.2011.07.027. 25 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 supplementary information . mapped microbialite area shapefiles, data tables, python code used for analysis, and supplemental figures are available at open science framework: https://osf.io/uf9yg/. 26 l. wilcock, c.m. frantz, and m.d. vanden berg use of remote imagery to map microbialite distribution at great salt lake zamoragroundwater flow and salinity contribution.pub 1 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2023 utah geological association publication 51 abstract groundwater discharge to great salt lake (gsl) is difficult to quantify but represents a potentially significant source of water and salinity to the lake’s overall water budget and chemistry, respectively. understanding groundwater and its role in the overall health of gsl is critical due to the current historically low lake levels. we compiled existing groundwater level data in wells in the basin-fill aquifer around gsl and used spatial analysis methods to 1) create potentiometric-surface maps in the areas adjoining gsl, 2) calculate groundwater contributions to gsl, and 3) estimate salinity inputs from groundwater to gsl. we observed groundwater-level declines in most of the basin-fill wells from the 1980s to 2010s. these declines are consistent with historical groundwater-level trends in the salt lake, tooele, curlew, and weber valleys and are a consequence of aquifer overdraft associated with less than average precipitation in the basin and increased groundwater withdrawals in the gsl watershed. using the darcy flux equation, we calculated a groundwater flux to gsl of 313,500 acre-feet per year, substantially greater than previous estimates derived from water balance studies but consistent with estimates derived from geochemical modeling of gsl water chemistry. we calculated a salt contribution from groundwater to gsl of 1.18 million metric tons per year, which represents about 10% of the solutes derived from surface flows to gsl in 2013. estimate of groundwater flow and salinity contribution to great salt lake using groundwater levels and spatial analysis hector a. zamora and paul inkenbrandt utah geological survey, salt lake city, utah, hector.zamora.hg@gmail.com, paulinkenbrandt@utah.gov introduction background and objectives groundwater discharge is an essential component in limnological systems’ hydrologic and chemical balances (healy and others, 2007; rosenberry and winter, 2009). despite its importance in developing accurate hydrochemical balances in lakes, groundwater contribution is often neglected or underestimated because it is difficult to quantify (rosenberry and others, 2015). understanding the groundwater component and its associated solute input is imperative for managing the environmental and economic resources of lakes affected by extensive anthropogenic water use and drought, such as great salt lake (gsl) in northern utah. potentiometric gradient and aquifer hydraulic conductivity are key inputs to calculate groundwater flow. this report focuses on the potentiometric data input for estimating groundwater flow. future work will better constrain hydraulic conductivity and geochemistry at the lake interface to improve groundwater flow estimates. gsl is a hypersaline terminal lake and a sink for surface and groundwater across a large part of the eastern great basin (spencer and others, 1985; duffy and al-hassan, 1988; arnow and stephens, 1990). salinity inputs and evaporation impact gsl’s ecosystems and mineral resources (carling and others, 2013; jagniecki and others, 2021). surface-water flows to gsl and associated salt loading are well constrained (shope and angeroth, 2015). however, the quantity of groundwater discharge remains relatively unknown, and groundwater is a potentially significant source to gsl’s overall salt load (kirby and others, 2019; bunce, 2022). the importance of understanding groundwater dynamics and its role in the overall health of gsl becomes prominant by the current, historically low lake levels. groundwater inflows to gsl will become critically important as surfacewater discharges decrease due to increasing water demands (null and wurtsbaugh, 2020), rising air temperatures, and changing snow cover conditions in the basin (hall and others, 2021). given the significance and uncertainty of groundwater to gsl’s system, the objectives of the present study are to 1) compile historical groundwater levels and use them to create generalized potentiometricsurface maps, 2) roughly estimate groundwater flow to gsl using a combination of spatial analysis techniques and darcy’s law, and 3) combine the results from the previous objective with existing groundwater chemistry data to estimate salinity inputs to gsl derived from groundwater. water and salt dynamics play a fundamental role in shaping not only gsl’s 10.31711/ugap.v51i.141 2 h.a. zamora and p. inkenbrandt estimate of groundwater flow and salinity contribution to great salt lake unique ecological, recreational, and mineral resources, but also the future development in the sprawling urban centers on the east shore of the lake. we provide the first systematic basin-wide assessment of groundwater levels in areas adjoining gsl to quantify groundwater contributions and their salt loading to the lake system. this information is required to constrain water and salt budgets needed by managing agencies to make informed decisions regarding the future health and productivity of ecosystems and industries in the lake. our findings serve as a basis for future work to better define the role of groundwater in gsl’s overall water volume and solutes budgets. study area the study area covers about 11,000 square miles in northern utah and southern idaho (figure 1). it is bounded on the east by the north-south-trending wasatch range and on the west by the great salt lake desert. several mountain ranges (hansel, promontory, oquirrh, stansbury, hogup, etc.) and associated valleys (curlew, hansel, malald–lower bear river, weber, salt lake, tooele, skull, etc.) bound the north and south flanks of the study area and drain into gsl (figure 1). gsl is the largest salt-water lake by area in the western hemisphere and the eighth largest in the world (hammer, 1986). it is 75 miles long by 28 miles wide and covers approximately 1700 square miles with a maximum depth of 33 feet at the average water-surface elevation of 4200 feet above sea level (fasl). the jordan, bear, and weber rivers deliver on average 2.9 million acre-feet of water to gsl, approximately 95% of the total stream inflow (stephens and gardner, 2007; mohammed and tarboton, 2012). previous studies (arnow and stephens, 1975; waddell and fields, 1976; loving and others, 2000; bunce, 2022) have estimated groundwater discharge to gsl ranges between 3% and 10% of the total inflow. the jordan, bear, and weber rivers also delivered an estimated 14.3 million metric tons of total dissolved solids (tds) in 2013 (shope and angeroth, 2015). groundwater potentially contributes a significant input to gsl’s overall salt load (hahl, 1968; spencer and others, 1985; loving and others, 2000), but the relationship between groundwater and gsl’s salinity has not been well defined. salinity of the lake ranges from 5% to 29% and creates diverse opportunities for ecological, recreational, and mineral uses. gsl is part of the pacific flyway and provides important nesting and foraging habitat for over 250 species of birds as they travel between north and south america (the nature conservancy, 2022). between 1.6 and 2.5 million metric tons of salt are commercially removed from the lake every year (stephens and gardner, 2007). mineral extraction, brine shrimp cyst production, and recreation in gsl can generate an estimated economic value of $1.32 billion per year (bioeconomics, 2012). as a terminal lake, gsl loses water primarily through evaporation. therefore, changes in streamflow conditions severely impact lake levels and salinity (mohammed and tarboton, 2012). stream diversions for agricultural and municipal uses reduce the amount of water flowing into gsl by 39% (null and wurtsbaugh, 2020). these diversions and a warming climate led to the present-day lake-level decline (wurtsbaugh and others, 2017; wang and others, 2018). as of november 2022, the area covered by gsl was reduced to about 900 square miles at a historical low water-surface elevation of 4188.5 fasl. consequently, salinity and the surface area covered by dry lakebed increased. increased salinity levels stress microbialite, brine fly, and brine shrimp populations, jeopardizing the entire ecological community that depends on them. dry lake beds are a major source of dust pollution and have the potential to accelerate snowmelt when dust is blown onto the snow (reynolds and others, 2014; skiles and others, 2018). methods data compilation we compiled historical groundwater level data from various datasets including the u.s. geological survey (usgs) national water information system (nwis, https://waterdata.usgs.gov/, u.s. geological survey, 2022a), the utah division of water rights (dwri) well drilling records, the utah geological survey (ugs) geologic hazards program subsurface geotechnical database, and the ugs wetlands section field data. data from the dwri was derived from a combination of well information tables and the water rights points of diversion (wrpod) feature class provided on the dwri website (https:// waterrights.utah.gov/). the wrpod feature class only includes wells with a well identification number (win) and excludes wells without an assigned win. the ugs geologic hazards program subsurface geotechnical database consists of 5141 boreholes in watersheds that contribute to the lake. the ugs wetlands section field data consists of 362 shallow boreholes in the wetlands proximal to gsl. in addition to groundwater level data, well properties (latitude, longitude, surface elevation, screen depth, and well depth) were also gathered where available. data compilation is limited to wells in the basin-fill aquifer and boreholes within the study area (figure 2). https://waterdata.usgs.gov/ https://waterrights.utah.gov/ https://waterrights.utah.gov/ 3 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 1. physiographic overview of gsl study area. 4 h.a. zamora and p. inkenbrandt estimate of groundwater flow and salinity contribution to great salt lake figure 2. geographical location of historical groundwater-level sites. 5 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 the compiled groundwater level data show spatial and temporal complexities within the study area. groundwater level measurements began in the 1900s but were limited to only a few sites (33). the number of data collection sites increased over time to reach a maximum of 3136 during the 1990s but remained variable (figure 3). furthermore, while many sites were visited regularly, others were visited less often, or only once in many cases. we reprojected, combined, and aggregated the data using python 3 and created two different datasets to deal with these spatial and temporal intricacies. the two datasets (tables s1 and s2) are included as supplementary information in geodatabase format (contact authors for database). the first dataset (table s1) contains data available from all sources (nwis, dwri, and ugs). groundwater level data for each site were grouped by decade and a mean water level elevation value was calculated if more than a single measurement was recorded during the time frame. the second dataset (table s2) contains only usgs nwis data for all wells available within the study area. here, the groundwater elevation value is the mean of all available data for each site. this dataset also contains estimated aquifer properties including saturated thickness, hydraulic conductivity, and transmissivity for individual sites. saturated thickness values were estimated by subtracting mean depth to water values from total well depth values. hydraulic conductivity values were extracted from layer 2 of the usgs’s groundwater model of the great basin carbonate and alluvial aquifer (brooks, 2017). finally, transmissivity values were calculated by multiplying saturated thickness times hydraulic conductivity. we downloaded three high-resolution aerial photographs from the european space agency’s sentinel -2 satellites for august 25th, august 28th, and september 9th, 2022, using the usgs global visualization viewer (glovis, https://glovis.usgs.gov/). sentinel-2 satellites carry an optical payload with visible, near infrared and shortwave infrared sensors encompassing 13 spectral bands: 4 bands at 10-meter, 6 bands at 20-meter, and 3 bands at 60-meter spatial resolution, with a swath width of 290 kilometers (european space agency, 2022). we created a single mosaicked image in arcgis pro® and used it to trace the extent of the lake. we assigned a surface elevation of 4190 feet to this extent based on usgs lake stage observations at the saltair boat harbor (usgs site id 10010000, https://webapps.usgs.gov/gsl/) on august 1st, 2022 (refer to figure 1 for location). lidar datasets are available for northern utah but do not cover the entire study area. for consistency, we downloaded six digital elevation model (dem) tiles from the terra advanced spaceborne thermal emission and reflection radiometer (aster) version 3 using the usgs earth explorer (https:// earthexplorer.usgs.gov/). these dem tiles have a one -third arc-second resolution (~10 meters) and are referenced to the north american vertical datum of 1988 (navd 88). we created a single mosaicked image in arcgis pro® and used it to extract surface elevations in well sites where this information was missing. we also used the aster dem as an explanatory variable input in the data interpolation process (discussed below). potentiometric surface interpolation we used the empirical bayesian kriging regression prediction (ebk-rp) tool in arcgis pro® to interpolate the water level elevation values in table s1. ebk-rp is a geostatistical interpolation method that uses ebk with explanatory variable rasters known to affect the value of the data being interpolated. the tool combines kriging with regression analysis to make predictions that are more accurate than either kriging or regression can achieve on their own (esri, 2022a). as the potentiometric surface generally follows topography, and because it is a component in the calculation of groundwater level elevation, we included the aster dem as an explanatory variable. we created two generalized potentiometric-surface maps: one for the 1980s (figure 4) and one for the 2010s (figure 5). these two decades were chosen because 1) not enough information is available to create a potentiometric surface for the 2020s, thus the 2010s data represent the most recent groundwater conditions for the study area, 2) average decadal gsl surface levels were close to the historical average (4200 fasl) in the 1980s, and 3) both decades contain about the same number of sites available for interpolation (figure 3). we also used table s1 to create a water figure 3. number of groundwater-level sites by decade. https://glovis.usgs.gov/ https://webapps.usgs.gov/gsl/ https://earthexplorer.usgs.gov/ https://earthexplorer.usgs.gov/ 6 h.a. zamora and p. inkenbrandt estimate of groundwater flow and salinity contribution to great salt lake figure 4. potentiometric surface elevation in study area for the 1980s. 7 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 5. potentiometric surface elevation in study area for the 2010s. 8 h.a. zamora and p. inkenbrandt estimate of groundwater flow and salinity contribution to great salt lake level difference map by subtracting water level elevation values (figure 6). this calculation was possible only for sites where data was available for both decades. we used the kriging tool in arcgis pro® to interpolate the water elevation, saturated thickness, and transmissivity values in table s2 (figure 7). the three output rasters are floating point, have the same areal extent, and have a cell size of 100. units are consistent for time (day) and length (feet) for all data. we used these rasters as input for the darcy flow tool in arcgis pro® to estimate groundwater seepage velocity values around gsl (discussed below). groundwater seepage velocity and darcy flux we applied two different approaches to estimate groundwater flow to great salt lake. both approaches are basic darcy flow estimates and should be considered rough approximations. the first approach used tools available in gis software and the second approach used basic linear discretization of the valleys around the lake. gis-based calculations of darcy flow for the first approach, we used the darcy flow tool in arcgis pro® to estimate groundwater seepage velocity around the lake. this method uses darcy’s law to model two-dimensional, vertically mixed, horizontal, and steady state flow, where groundwater head is independent of depth. darcy’s law states that darcy velocity in porous material is calculated from the hydraulic conductivity and the hydraulic gradient as: (1) where: q = darcy velocity, darcy flux, or specific discharge (v/t/a or l/t) k = hydraulic conductivity (l/t) ∇i = hydraulic gradient (dimensionless) ∇h = change in hydraulic head over length (l) ∇l = change in length (l) hydraulic conductivity (k) may be calculated from the transmissivity and thickness as: (2) where: t = transmissivity (l2/t) b = aquifer thickness (l) the specific discharge (q) is defined as the volume of water flow per unit time through a cross-sectional area normal to the direction of flow (bear, 1979). specific discharge is directly proportional to hydraulic conductivity. the aquifer flux is defined as: (3) where: u = aquifer flux (v/t/l) t = transmissivity (l2/t) ∇i = hydraulic gradient (dimensionless) the aquifer flux (u) represents the discharge per unit width of the aquifer. the average fluid velocity within the pores, or seepage velocity, is the darcy velocity (q) divided by the effective porosity of the medium: (4) where: v = groundwater seepage velocity (l/t) q = darcy velocity, darcy flux, or specific discharge (v/t/a or l/t) n = effective porosity (%) the groundwater seepage velocity (v) is calculated on a cell-by-cell basis in the darcy flow procedure (esri, 2022b). for cell {i, j}, the aquifer flux (u) is calculated through each of the four cell walls, using the difference in heads between the two adjacent cells and the harmonic average of the transmissivities (konikow and bredehoeft, 1978), which are assumed to be isotropic (esri, 2022b). the darcy flow tool requires four raster datasets as input: groundwater head elevation (fasl), saturated thickness (feet), formation transmissivity (square feet per day), and effective formation porosity. we created the first three datasets by interpolating the aquifer property data in table s2 (figure 7), and we assigned the effective formation porosity a value of 0.35 for the basin-fill aquifer. two raster datasets result from this calculation: an output magnitude raster and an output direction raster. in the output magnitude raster, each cell value represents the magnitude of the seepage velocity vector (average linear velocity) at the center of the cell and is calculated as the average value of the seepage velocity through the four faces of the cell (esri, 2022b). in the output flow direction raster, each cell value represents the direction of the seepage velocity vector (esri, 2022b). we extracted the mean groundwater seepage velocity around gsl from the output magnitude raster using the zonal statistics as table tool in arcgis pro® with the gsl perimeter shapefile (both 4190 and 4200 fasl, table 1) as the feature zone data (dataset that defines zone of interest). we conducted zonal statistics using both polylines and polygons of gsl at both elevations to compare how the cells’ statistics were aggregated. we also divided the 4190 fasl gsl perimeter shapefile into three different 9 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 6. groundwater level difference between the 1980s and 2010s. 10 h.a. zamora and p. inkenbrandt estimate of groundwater flow and salinity contribution to great salt lake figure 7. raster datasets used as input to calculate groundwater seepage velocity around gsl in arcgis pro®: a) usgs water level sites, b) potentiometric surface, c) saturated thickness, and d) transmissivity. all data needed to create these raster datasets are available in table s2. 11 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 sections: west/south, north, and east (figure 8). we extracted the groundwater seepage velocities from these different sections (using the zonal statistics tool as above) to estimate the potential groundwater contribution by area (table 2). finally, we multiplied the groundwater seepage velocity by the mean aquifer thickness (232 feet, calculated from table s2) and the perimeter length of gsl (obtained from the gsl perimeter shapefiles) to obtain an estimated groundwater contribution (q) in acre-feet per year (tables 1 and 2). straight line calculations of darcy flow for the second approach, we estimated the groundwater discharge from each adjoining basin-fill valley using darcy’s law: (5) for this approximation, we used a mean k value of 12.43 feet per day (from table s2). we created several polygons to roughly constrain the areal extent of the basin fill in the valleys (figure 8). we calculated the cross-section area (a) for the flux calculations by multiplying the width of the polygon (perpendicular to the flow direction) by the mean aquifer thickness of 232 feet (from table s2). we calculated the hydraulic gradient (∇i) by extracting h1 and h2 values from the groundwater elevation raster, near the upper and lower ends of the polygons (figure 8) and dividing the difference between the two piezometric heads by the length of the polygon (parallel to groundwater flow). the geometrical properties and estimated groundwater discharge for each basin-fill area, (q) in acre-feet per year, are presented in table 3. quantifying error to better understand the amount of variability that can be introduced by hydraulic conductivity, we iteratively calculated the groundwater flow for each area using the straight-line approach. we created a lognormal distribution of hydraulic conductivity, using 1.094 (log 10 of 12.43 ft/day) as the mean and 1.3 as the standard deviation, which is the standard deviation of the hydraulic conductivity of the upper basinfill aquifer from the usgs groundwater model. we randomly sampled hydraulic conductivity 100,000 times from this distribution and used the resulting summary statistics to constrain variability of groundwater flow estimates from the calculations. salt loading we calculated mean tds values around gsl using data available from previous studies in the area (kirby and others, 2019). from kirby and others’ (2019) data, we obtained three different tds values using 1) their calculated tds values to compute a mean tds for the whole study area, 2) their tds raster and the zonal statistics as table tool in arcgis pro® to estimate the mean tds value around gsl at 4190 fasl (polyline), and 3) their tds raster and the zonal statistics as table tool in arcgis pro® to estimate the mean tds value around gsl at 4200 fasl (polyline, table 4). we used the sectioned 4190 fasl gsl shapefile (west/south, north, and east) in figure 8 to extract mean tds values by segment from the tds raster using the zonal statistics as table tool in arcgis pro® (table 5). this approach helped to estimate the salt loading contribution by section. we combined the mean tds values, in milligrams per liter (mg/l), with the groundwater discharge (q, acrefeet) values in tables 1 and 2 to estimate a salt loading to the lake in metric tons per year (tables 4 and 5). results generalized potentiometric surface the generalized potentiometric surface maps for the 1980s and 2010s show that, at the scale of the study area, groundwater flow patterns are relatively constant over time (figures 4 and 5). groundwater flows from the high-elevation mountains surrounding the study area towards the adjacent valleys and into gsl level ( asl) length ( ) aquifer thickness ( ) area ( 2) mean seepage velocity ( /day) q (af/year) 4190a  1,388,561  232  322,666,029  0.15  411,000 4190b  1,388,561  232  322,666,029  0.12  324,000 4200a  1,768,800  232  411,023,839  0.09  310,000 4200b  1,768,800  232  411,023,839  0.11  379,000 a es mate made using gsl polyline  b es mate made using gsl polygon  table 1. groundwater flux estimates for gsl in acre-feet per year. 12 h.a. zamora and p. inkenbrandt estimate of groundwater flow and salinity contribution to great salt lake figure 8. polygons used in darcy’s law equation to estimate groundwater discharge from each basin-fill valley adjoining gsl. colors around gsl show divisions used to estimate groundwater and salt contributions by section. 13 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 gsl. for both decades, the steepest groundwater gradients occur in the vicinity of the raft river mountains, across the west flank of the oquirrh mountains (southeastern tooele valley), across the east flank of the oquirrh mountains (western salt lake valley), and along the wasatch range (figures 4 and 5). there are noticeable changes in the potentiometric surface between the two decades. for example, in salt lake valley, the 4300-foot contour has moved farther south (upstream), and we observed cones of depression north of the oquirrh mountains and in the taylorsville and bountiful areas (figures 9 and 10). we also observed some recovery in the weber river delta area where groundwater recharge projects have been operating since the 2000s (hurlow and others, 2011; figures 11 and 12). figure 6 also shows groundwaterlevel recovery in some wells. however, groundwaterlevel recovery is an exception rather than the rule because 129 out of 147 wells having data available in both decades show a groundwater-level decline (figure 6). groundwater-level declines are particularly high in localized areas of the north malad river valley in idaho (-86 feet), curlew valley (-35 to -49 feet), near the bear river (-51 feet), and salt lake valley (-44 feet). the data in figure 6 show there has been a mean decrease in water levels of 10.14 feet from the 1980s to 2010s. seepage velocity and darcy flux mean groundwater seepage velocities into gsl are 0.15 (polyline) and 0.12 (polygon) feet per day using the 4190 fasl shapefiles (table 1). mean groundwater seepage velocities along the gsl perimeter are 0.09 (polyline) and 0.11 (polygon) feet per day using the 4200 fasl shapefiles (table 1). these mean seepage velocities result in groundwater fluxes ranging from 310,000 to 411,000 acre-feet per year (table 1). the average of these estimates is 356,000 acre-feet per year. the mean groundwater seepage velocities by section (4190 fasl) are 0.05 (west/ south), 0.30 (north), and 0.10 (east) feet per day (table 2). these mean seepage velocities result in groundwater fluxes of 50,291 (west/south), 274,660 (north), and 85,829 (east) acre-feet per year for a total of 411,000 acre-feet per year (table 2). using the darcy flux equation on the linear traverses in adjoining valleys (figure 8), groundwater contributions to gsl range from 3500 (skull valley) to 63,900 (tooele valley) acre-feet per year for a total of 313,500 acre-feet per year (table 3). the groundwater discharge from the wasatch range (sum of salt lake, weber, and malad valley and brigham city) is 146,400 acre-feet per year, 47% of the total groundwater contribution to the lake. salt loading the mean tds values along gsl are 2116, 2538, and 3594 mg/l with an average value of 2749 mg/l using three different methods as explained in the “methods” section (table 4). these tds values and the calculated groundwater contributions in tables 1 and 3 result in yearly dissolved solid (salt) fluxes ranging from 810,000 to 1,820,000 metric tons, and an average of 1,060,000 metric tons, into gsl (table 4). table 5 reports the mean tds values by perimeter section as 4627 (west/south), 3055 (north), and 3023 (east) mg/l. these tds values and the calculated groundwater contribution by section in table 3 result in yearly salt fluxes of 290,000 (west/south), 1,040,000 (north), and 320,000 (east) metric tons (table 5), resulting in 1,650,000 metric tons per year. error the uncertainty of the darcy flux calculations on the linear traverses through adjoining valleys is high. the 5th percentile for flow values is 2350 acre-feet/yr and the 95th percentile is 44.5 million ac-ft/yr. see table 3 for a complete list of variations associated with potential variability in hydraulic conductivity. discussion generalized potentiometric surface groundwater levels declined in most of the basinfill aquifer from the 1980s to the 2010s (figure 6). these declines are consistent with observed historical sec on length ( ) thickness ( ) min seepage velocity ( /day) max seepage velocity ( /day) median seepage velocity ( /day) mean seepage velocity ( /day) min q (af/year) median q (af/year) mean q (af/year) west/south  485,343  232  0.0003  3.3  0.005  0.05  292  4,900 50,000 north  469,067  232  0.0004  16.8  0.016  0.30  332  14,600 275,000 east  434,151  232  0.0013  14.4  0.027  0.10  1,135  23,000 86,000 1,759  42,600 411,000 table 2. groundwater flux estimates for gsl by section at 4190 fasl in acre-feet per year. table 3. groundwater contribution from adjoining areas in acre-foot per year h.a. zamora and p. inkenbrandt estimate of groundwater flow and salinity contribution to great salt lake 14 polygon  area  avg k   ( /day)  width   ( )  thickness   ( )  h1   ( )  h2   ( )  length  ( )  gradient  area  ( 2)  q  (af/year)  5th %‐ le  (ac‐ /yr)  median q   (ac‐ /yr)  95th %‐ le  (ac‐ /yr)  1 skull valley  12.43  54,454  232  4823  4289  203,975  0.0026  12,653,714  3,500  30  3,400  486,500  2 tooele valley  12.43  40,416  232  5179  4292  32,359  0.0274  9,391,642  26,800  200  26,600  3,780,500  3 tooele valley  12.43  71,767  232  5234  4282  44,598  0.0213  16,676,811  37,100  280  36,900  5,227,800  4 salt lake valley  12.43  45,308  232  5255  4254  71,213  0.0141  10,528,418  15,400  110  15,300  2,173,300  5 salt lake valley  12.43  92,678  232  4700  4277  24,740  0.0171  21,535,992  38,400  280  38,100  5,407,400  6 weber valley  12.43  191,539  232  4542  4267  21,629  0.0127  44,508,754  59,000  440  58,600  8,310,400  7 brigham city  12.43  20,500  232  4686  4298  11,987  0.0324  4,763,675  16,100  120  16,000  2,264,400  8 malad valley  12.43  42,798  232  4600  4377  14,285  0.0156  9,945,158  16,200  120  16,100  2,279,900  9 malad valley  12.43  22,591  232  5224  4430  190,199  0.0042  5,249,569  2,300  20  2,300  321,800  10 blue creek  12.43  31,015  232  4950  4476  129,879  0.0036  7,207,091  2,700  20  2,700  386,300  11 n. promontory  12.43  29,403  232  5069  4554  18,322  0.0281  6,832,504  20,000  150  19,900  2,820,300  12 n. promontory  12.43  10,239  232  5090  4247  42,539  0.0198  2,379,281  4,900  40  4,900  692,400  13 hansel valley  12.43  27,543  232  4934  4495  53,115  0.0083  6,400,287  5,500  40  5,500  776,800  14 curlew valley  12.43  29,051  232  5107  4321  149,383  0.0053  6,750,708  3,700  30  3,700  521,600  15 park valley  12.43  34,923  232  4710  4266  64,256  0.0069  8,115,210  5,800  40  5,800  823,500  16 park valley  12.43  69,117  232  4658  4276  35,522  0.0108  16,061,019  18,000  130  17,900  2,536,400  17 park valley  12.43  32,002  232  5237  4212  32,206  0.0318  7,436,445  24,700  180  24,500  3,475,600  18 gsl desert  12.43  58,076  232  4931  4292  138,707  0.0046  13,495,374  6,500  50  6,400  913,000  19 gsl desert  12.43  58,197  232  4854  4223  99,404  0.0063  13,523,491  8,900  70  8,900  1,260,700                                315,500  2,350  313,500  44,458,600  15 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 groundwater level trends in the salt lake, tooele, and curlew valleys where levels dropped between 15 and 40 feet from 1975 to 2005 (burden and others, 2005). groundwater levels declined by an average of 27 feet from 1953 to 1985 in the ogden area (hurlow and others, 2011). our data shows that, in general terms, this trend continued into the 2010s. these declines are a consequence of aquifer overdraft associated with less than average precipitation in the basin and increased withdrawals for municipal, industrial, and agricultural use (burden and others, 2005). young and others (2021) estimated that the gsl basin lost 8.8 ± 2.3 million acre-feet of groundwater storage whereas gsl lost 4.5 ± 0.8 million acre-feet of surface water during the 2012–2016 drought. overdraft conditions in basin-fill aquifers can cause several problems. groundwater removal through pumping for anthropogenic use (and associated groundwater level declines) may lead to reduction of water in streams and lakes, land subsidence or ground failure due to soil compaction, increased costs for users due to higher pumping lifts, and deterioration of water quality from saltwater intrusion (u.s. geological survey, 2022b). consumptive water uses in the gsl watershed have already depleted surface inflows to the lake by ~39% (null and wurtsbaugh, 2020). these inflow reductions are in part responsible for the recent ~10-foot drop in gsl’s surface water level (figure 1; null and wurtsbaugh, 2020). the downward trend in water-surface elevation is expected to continue as human population and water consumption increase under changing climatic conditions in the state. land subsidence and earth fissures due to long-term groundwater pumping in excess of recharge have been reported in cedar valley in southern utah, where average basin-wide subsidence is estimated to continue at a rate of 0.04 to 2.4 inches per year under the current rates of groundwater decline (3 feet per year; lund and others, 2011). several instances of land subsidence have been reported in woods cross city (see figure 10 for location), but current subsidence rates are unknown. several longterm monitoring sites on gsl’s east shore show a significant increase in tds over time (clark and others, 1990), but saltwater intrusion into the freshwater aquifer has not been explicitly documented. however, groundwater-level declines on the east shore may create the conditions required to induce saltwater intrusion. such phenomena have been observed in lake urmia, a terminal saline lake in iran, where groundwater-level declines of 13 feet in the freshwater aquifer induced saltwater intrusion (ahmadi and others, 2022). the general trends in groundwater levels shown in figure 6 and the 100-foot contour intervals shown in figures 4 and 5 allow visualization of groundwater conditions in the study area. unfortunately, 100-foot contours do not provide good spatial resolution on fine-scale responses to groundwater withdrawals in the basin-fill aquifers. similarly, mean groundwater levels on a decadal timescale provide a glimpse of the hydrological conditions at specific sites, but do not offer a detailed temporal resolution of groundwaterlevel responses to withdrawal or recharge. groundwater-level maps at finer spatial and temporal scales than presented here are needed for each individual valley adjoining gsl to track withdrawal responses on a yearly (or seasonal) basis. thorough and more comprehensive groundwater-level maps are particularly needed along the east shore of gsl to monitor potential saltwater intrusion to the freshwater aquifer. q (af/year) tds (mg/l) dissolved load flux (metric ton/year) 313,500  2,116  a  820,000  313,500  2,538  b  980,000  313,500  3,594  c  1,390,000  411,000  2,116  a  1,070,000  411,000  2,538  b  1,290,000  411,000  3,594  c  1,820,000  324,000  2,116 a  850,000  324,000  2,538 b  1,010,000  324,000  3,594  c  1,440,000  309,966  2,116  a  810,000  309,966  2,538  b  970,000  309,966  3,594  c  1,370,000  379,000  2,116  a  990,000  379,000  2,538  b  1,190,000  379,000  3,594  c  1,680,000  347,493 2,749 1,060,000 table 4. salt flux estimates for gsl in metric tons per year. average values are in boldface. aaverage of calculated tds values from available well data in the study  area (kirby and others, 2019)  bes mate obtained using tds raster from kirby and others (2019) and  gsl shapefile (polyline) at 4190 fasl   ces mate obtained using tds raster from kirby and others (2019) and  gsl shapefile (polyline) at 4200 fasl   sec on q (af/year) tds (mg/l) tds (metric ton/year) west/south  50,000  4627  290,000  north  275,000  3055  1,040,000  east  86,000  3023  320,000  1,650,000 table 5. salt flux estimates for gsl by section at 4190 fasl in acre-feet per year 16 h.a. zamora and p. inkenbrandt estimate of groundwater flow and salinity contribution to great salt lake figure 9. potentiometric surface elevation in the salt lake valley area for the 1980s. 17 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 10. potentiometric surface elevation in the salt lake valley area for the 2010s. 18 h.a. zamora and p. inkenbrandt estimate of groundwater flow and salinity contribution to great salt lake figure 11. potentiometric surface elevation in the weber delta area for the 1980s. 19 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 12. potentiometric surface elevation in the weber valley area for the 2010s. 20 h.a. zamora and p. inkenbrandt estimate of groundwater flow and salinity contribution to great salt lake groundwater flow and salt loading the total groundwater flux to gsl calculated in this study (an average of 356,000 acre-feet per year from the seepage calculations, 313,500 acre-feet per year from the darcy flux calculations) is substantially higher than previous estimates derived from water balance studies (75,000 acre-feet per year; waddell and fields, 1976; loving and others, 2000). arnow and stephens (1990) estimated that between 6250 and 100,000 acre-feet of groundwater enter the lake per year. however, our results are consistent with recent estimates derived from geochemical modeling of gsl water chemistry (10% of the total inflow or ~300,000 acre-feet per year), assuming surface water contributions of 2.9 million acre-feet per year to the system (stephens and gardner, 2007; bunce, 2022). estimates of groundwater seepage using seepage meters averaged 0.77 cm/day from july 8–15, 2010, (anderson, 2012) at locations of suspected groundwater seepage. extrapolating that rate to the entire lake at its current coverage area would suggest that about 560,000 acre-feet of groundwater seeps into the lake per year. however, the seepage estimates could vary spatially, seasonally, and temporally, and were likely biased towards higher seepage rates due to how the measurement sites were selected. based on our preliminary estimates, the largest groundwater contribution originates in the north and east sections of gsl (table 5) where steep hydraulic gradients occur (figures 4 and 5). the groundwater flow derived in the north section of gsl was unexpectedly high (table 2). this high number is likely explained by the high transmissivity values calculated for the areas around the park and curlew valleys (figure 7d). groundwater flow conditions in park valley are poorly known. extensive groundwater pumping in curlew valley has substantially reduced groundwater levels and discharge from the locomotive springs complex during the past 40 years (hurlow and burk, 2008), therefore the groundwater flux from this valley to gsl has likely declined significantly. the average salt contribution from groundwater to gsl calculated in this study (1.18 million metric tons per year) represents about 10% of the solutes delivered by the jordan, bear, and weber rivers to gsl in 2013 (14.3 million metric tons; shope and angeroth, 2015). the highest tds concentrations are found in the west/south sections of gsl (table 5) where hydraulic gradients are shallow, evaporation rates are high, and recharge likely occurs at a slow rate. our new estimates of groundwater discharge and its salinity contribution will likely require revision of gsl’s water and salt budgets. however, two important considerations limit how these should be evaluated and used. 1. error on the estimates of groundwater flux is large. based on sensitivity analyses from iterative calculations, the estimates are most sensitive to values of hydraulic conductivity. hydraulic conductivity is lognormally distributed and can range by orders of magnitude over a study area as large as ours, resulting in estimates of flux that range over orders of magnitude. of the iterative calculations conducted, 90% of resultant estimated flux fell between 2350 and 44,000,000 acre-ft/yr, indicating a need to better constrain aquifer properties. 2. because we only considered wells completed in the basin-fill aquifer, our calculations do not include flow paths that are entirely within bedrock (but do include groundwater that discharges from bedrock to basin fill in the subsurface). significant discharge from bedrock springs occurs in the southeastern part of the malad–lower bear river valley, along the margins of the promontory mountains, and in the northwest part of tooele valley. these springs contribute groundwater flow to the gsl playa and, perhaps, different salt loading having different compositions and concentrations than groundwater in the basin-fill aquifer. other aquifer properties, including porosity, saturated aquifer thickness, and cross-sectional area, also influence flux estimates. further information is needed to constrain the differences in porosity values around gsl’s shorelands and in the basin-fill aquifer. these data can potentially improve the seepage velocity estimates around gsl. additionally, the aquifer thickness values we used (difference between depth to water and total well depth) have two sources of uncertainty. first, most of the wells in the basin fill only partially penetrate the saturated thickness of the aquifer. thus, aquifer thickness values around gsl are likely larger than we estimated and could result in larger groundwater fluxes than presented here. second, most of the wells used in this study were drilled to target the most productive aquifer depths. for some areas, this would result in larger groundwater and salinity contributions than expected due to bias toward higher aquifer-property values. groundwater levels change over time due to natural and anthropogenic influences, resulting in variable hydraulic gradients and saturated thicknesses. using previous work and ongoing groundwater-level observations, we can constrain these estimates fairly well. cross-sectional area of groundwater flow paths is more complicated, especially if one assumes that the area matches that of the lake margin. the lake perimeter varies dramatically 21 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 depending on lake level. increases in lake level will increase the lake perimeter length and the resultant cross-sectional area of estimation. hypothetically, an increase in lake level would result in a decreased groundwater gradient, but due to lack of lake-margin groundwater data, we are unsure of this relationship. regarding salinity inputs from groundwater, the spring systems around gsl need further consideration. springs in the area have been measured to reach tds concentrations of up to ~76,000 mg/l (bunce, 2022) and are point-sources of solutes to gsl. it is also possible that there are density differences in the groundwater system around and below gsl. these density variations have the potential to create flow boundaries that we did not account for (rosen, 1994; sheibani and others, 2020). further studies are needed to understand spring dynamics and density-driven flows in order to provide further insight on their overall role in the water and salinity budgets in gsl. conclusions and recommendations we provide the first systematic, basin-wide assessment of groundwater levels in areas adjoining gsl to quantify groundwater contributions and their salt loading to the lake system. we observed groundwater-level declines in most of the basin-fill aquifer from the 1980s to the 2010s (figure 6). these declines are consistent with historical groundwater level trends in the salt lake, tooele, curlew and weber valleys and result from aquifer overdraft associated with less than average precipitation in the basin and increased withdrawal for human consumption. we calculated a mean groundwater flux to gsl of 356,000 acre-feet per year using a seepage velocity method in arcgis and 313,500 acre-feet per year using the darcy flux equation for linear traverses through adjoining valleys. both estimates are substantially greater than previous estimates derived from water balance studies, but are consistent with estimates derived from previous in situ seepage measurements and geochemical modeling of gsl water chemistry. we calculated a salt contribution from groundwater to gsl of 1.06 million metric tons per year which represents about 10% of the solutes derived from surface flows to gsl in 2013. these estimates have very large uncertainty, and the input parameters need to be better understood and constrained. groundwater monitoring wells and a formal groundwater model are recommended to constrain groundwater parameters. the data presented here have the potential to improve current water and salt budgets for gsl’s system. however, further work is needed to improve these estimates and better delineate surface/ groundwater dynamics in the area. in order to do so, we recommend the following: 1. estimates of the hydraulic properties of the basin-fill aquifer should be refined by compiling results from high-quality well tests and aquifer tests and, perhaps, conducting new aquifer tests. 2. a monitoring-well network should be established with local/state/federal participation within each valley adjoining gsl. this system of wells should be thorough and accessible to visit and measure water level fluctuations on a seasonal or, at least, yearly basis. this information could be used to create detailed, yearly potentiometric surface maps and track/compare changes in groundwater levels over the years. 3. nested piezometers and/or monitoring wells should be installed along different sections of gsl. these piezometers at different depths could be used to calculate hydraulic gradients and to monitor water level/salinity trends in areas of the aquifer susceptible to brine intrusion. core or cuttings recovered during the installation of these piezometers/monitoring wells could be used to estimate porosity values in the subsurface near gsl. 4. sample springs around gsl and measure their flow. geochemical and isotopic data on springs can provide information regarding sources, flow paths, residence time of groundwater and help to better understand their role in the water/salt budget of gsl. acknowledgments we would like to thank our sister-agency the utah division of forestry, fire and state lands for funding this study. we also thank reviewers marek matyjasik and craig miller, and editor carie frantz for their contributions to this manuscript. references ahmadi, h., hemmati, m., and motallebian, m., 2022, numerical modeling of saltwater wedge under intruding and receding conditions (case study: kahriz aquifer, lake urmia): water resources, v. 49, no. 2022, p. 249–258, doi: https:// doi.org/10.1134/s0097807822020099. anderson, r.b., 2012, quantity and quality of groundwater 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/hrv (za stvaranje adobe pdf dokumenata najpogodnijih za visokokvalitetni ispis prije tiskanja koristite ove postavke. stvoreni pdf dokumenti mogu se otvoriti acrobat i adobe reader 5.0 i kasnijim verzijama.) /hun 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can be opened with acrobat and adobe reader 5.0 and later.) >> /namespace [ (adobe) (common) (1.0) ] /othernamespaces [ << /asreaderspreads false /cropimagestoframes true /errorcontrol /warnandcontinue /flattenerignorespreadoverrides false /includeguidesgrids false /includenonprinting false /includeslug false /namespace [ (adobe) (indesign) (4.0) ] /omitplacedbitmaps false /omitplacedeps false /omitplacedpdf false /simulateoverprint /legacy >> << /addbleedmarks false /addcolorbars false /addcropmarks false /addpageinfo false /addregmarks false /convertcolors /converttocmyk /destinationprofilename () /destinationprofileselector /documentcmyk /downsample16bitimages true /flattenerpreset << /presetselector /mediumresolution >> /formelements false /generatestructure false /includebookmarks false /includehyperlinks false /includeinteractive false /includelayers false /includeprofiles false /multimediahandling /useobjectsettings /namespace [ (adobe) (creativesuite) (2.0) ] /pdfxoutputintentprofileselector /documentcmyk /preserveediting true /untaggedcmykhandling /leaveuntagged /untaggedrgbhandling /usedocumentprofile /usedocumentbleed false >> ] >> setdistillerparams << /hwresolution [2400 2400] /pagesize [612.000 792.000] >> setpagedevice radwinlakebedevolution.pub 1 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2023 utah geological association publication 51 abstract the great salt lake has been rapidly shrinking since the highstand of the mid-1980s, creating cause for concern in recent decades as the lake has reached historic lows. many investigators have assessed the evolution of lake elevation, geochemistry, anthropogenic impacts, and links to climate and atmospheric processes; however, the use of remote sensing to study the evolution of the lake has been significantly limited. harnessing recent advancements in cloud-processing, specifically google earth engine cloud computing, this study utilizes over 600 landsat tm/oli and sentinel msi satellite images from 1984-2023 to present time-series analyses of remotely sensed great salt lake water area, exposed lakebed area, surface cover types, and chlorophyll-a analyses paired with modelled estimates for water and exposed lakebed area. results show that since the highstand of 1986-1987, the water area has declined by 45% (~3,000 km2) and the exposed lakebed area has increased to ~3,500 km2 from ~500 km2. the area of unconsolidated sediments not protected by vegetation or halite crusts has risen to ~2,400 km2. significant halite crusts are observed in the north arm, having a max extent of ~150 km2 between 2002 and 2003, while only small extents of halite crusts are observed for the south arm. vegetation is more prevalent in the bear river bay and south arm, with surface area increases over 400% since 1990. gypsum is widely observed independent of halite crusts. the results highlight multiple instances of land-use/water-management that led to observable changes in water/exposed lakebed area and halite crust extent. this study demonstrates the important benefits of maintaining a lake elevation above ~4,194 ft to maximize lake and halite crust area, which would help mitigate possible dust events and maintain a broad lake extent. evolution of great salt lake’s exposed lakebed (1984-2023): variations in sediment composition, water, and vegetation from landsat oli and sentinel msi satellite reflectance data mark h. radwin and brenda b. bowen university of utah, salt lake city, utah, markradwin@gmail.com introduction in recent years the great salt lake in northern utah has attracted the attention of local legislators and a global audience as the lake reached historic lows and caused concerns for public health and the health of the overall great salt lake ecosystem. once part of the vast pleistocene lake bonneville, the great salt lake has shrunk to the modern state from an evaporative evolution in a closed basin with natural inputs from three major rivers (bear, jordan, and weber rivers). however, in recent centuries, anthropogenic activity has considerably affected the great salt lake. this influence extends to direct physical alterations of the lake's landscape, modifications to its hydrology that alter water flow and distribution, the introduction of invasive plant species, and extensive resource extraction. in 1959 a railroad causeway was completed, separating the lake into a north and south arm, which has been modified over the years with various breaches, culverts, and berms to control flow between the flow between the two arms (figure 1). additionally, the railroad causeway, mineral operator evaporation ponds, and other various impoundments have significantly separated and controlled the flow from bear river bay to the south arm. with the north arm largely cut-off from major river inputs, it has evolved to be much more saline and commonly surpasses halite saturation, leading to precipitation of lake-bottom and shoreline halite crusts as well as a different color of water due to halophilic microorganisms. the lake has been used by wildlife as a crucial bird migratory location and anthropogenically for resource extraction. in the 1980s the lake rose nearly 8 ft due to an unusually heavy period of precipitation between 1982 and 1987, but has been steadily shrinking since, reaching a historic low in 2022. with the ongoing reduction in the lake's size, there is an escalating risk of moderate-to-severe dust storms associated with lakebed exposure and substantial changes in the ecosystem, which could adversely impact bird migrations. additionally, the overall stability of the regional ecosystem is becoming increasingly compromised. many aspects of the great salt lake’s evolution are well documented. since the mid-1800s the us geological survey has been recording lake elevations and water quality metrics, and since the mid-1900s 10.31711/ugap.v51i.134 2 m.h. radwin and b.b. bowen evolution of great salt lake’s exposed lakebed (1984-2023) the utah geological survey has been recording geochemical measurements (arnow, 1984; gwynn, 2007; rupke and mcdonald, 2012; naftz and others, 2013) . in recent decades investigators have started assessing the contributing factors to the decline of the great salt lake through water balance models, finding anthropogenic reduction of inflow and drought conditions (precipitation/inflow) to be the leading drivers of lake decline, with climate (evaporation) being a secondary factor (mohammed and tarboton, 2012; wurtsbaugh and others, 2016; wurtsbaugh and others, 2017; wine and others, 2019; null and wurtsbaugh, 2020; wurtsbaugh and sima, 2022). more recent reports have constrained the impact of natural and human consumptive use to be responsible for 67-73% of the great salt lake water loss figure 1. map of the great salt lake system and surrounding localities, including boundaries for the north arm, south arm, and bear river bay. also defined are the boundaries of the north arm mineral operator evaporation pools included in analyses between 1984-1994 and the evaporation pool masked for halite analyses in the bear river bay. the dashed rectangular line indicates the area captured by the sentinel-2 msi satellite and the basemap is landsat 8 oli imagery from june 1st (south image) and 2nd (north image). 3 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 (ahmadi and others, 2023). other studies have observed relationships between atmospheric oscillations in the pacific and multidecadal drought conditions which directly affect great salt lake levels, and determined that although climate change will lower lake levels through higher temperatures, evaporation, and changes in the snowmelt cycle, those impacts will be overshadowed by anthropogenic water withdrawal and drought conditions (wang and others, 2012; mohammed and tarboton, 2012; wine and others, 2019; hall and others, 2021; ahmadi and others, 2023). further, climate models suggest there will be an increase in precipitation with a warmer climate, but increases in precipitation will be negated by a greater increase in evaporation (ahmadi and others, 2023). related to the impacts of a shrinking great salt lake, others have investigated dust sources around the shoreline, impacts from dust events and dust-onsnow, pollutant contamination of dust-derivedsediments, regional land cover changes, and the atmospheric characteristics of dust events, all finding great salt lake sediments to be a significant dust source in northern utah (hahnenberger and nicoll, 2012; hahnenberger and nicoll, 2014; skiles and others, 2018; perry and others, 2019; nicoll and others, 2020; carling and others, 2020). although these aspects of the lake are well documented, the use of remote sensing to document the changing great salt lake system is only limited to water-surfacetemperature, algal blooms, outdated classification maps, and the common use of side-by-side true-color satellite image comparisons (hung and wu, 2005; bradt and others, 2006; crosman and horel, 2009; hansen and others, 2016). here, multispectral remote sensing data of the great salt lake from 1984 to 2023 are used to assess the evolution of sediment types and sediment area, vegetation area, water area, and relative chlorophyll-a concentrations between the north arm, south arm (including the farmington bay), and bear river bay. the nasa/usgs landsat 5 thematic mapper (tm), landsat 8 operational land imager (oli), and landsat 9 oli, as well as the esa sentinel 2 a&b multispectral instrument (msi) satellite platforms are chosen for this study, where the landsat imagery extends back to the 1980’s while the sentinel imagery extends back to 2019 for this region. combining these datasets results in over 600 near-cloud-free satellite scenes of the region from 1984 to 2023. historically, this volume of data prevented analyses due to the sheer amount of work and processing power involved, but has recently become feasible through automation and cloud-processing platforms. the results will help to understand the evolution of exposed sediments, halite crust formation, changes in vegetation, and the relationships between land-use, climate, and increasing sediment area. this work builds off of recent remote sensing studies in the bonneville basin which utilized landsat 5 tm and 8 oli multispectral imagery to map halite, gypsum, and carbonate-muds (lacustrine detritus; bowen et al., 2017; radwin & bowen, 2021). methods data sources and cloud processing the landsat tm/oli and sentinel msi platforms were used for this analysis as these sensors can capture the entire extent of either arm of the great salt lake and bear river bay during a single swath path, have suitable spatial and spectral band wavelengths for investigating surface features and types, theoretically allow for at least one image acquisition per month, and have longevity with multispectral data extending back to the 1980s. the landsat 5 tm, 8 oli, and 9 oli platforms have a spatial resolution of 30 m/pixel and seven bands (six for tm) ranging the vswir spectrum (~350-2500 nm), with a revisit time of 16 days (table 1). the landsat 5 platform was operational from 1984 to 2012, and the landsat 8 and 9 platforms have been operational since 2013 and 2021, respectively. the sentinel-2 msi platform, operational since 2015, has a spatial resolution that ranges from 10-60 m/pixel (max 20 m/pixel used in this study) and 12 bands ranging the vswir spectrum (table 1), with a revisit time of 10 days (5-days including both a&b satellites). although the sentinel-2 platform has been active since 2015, images for utah were not acquired until very late 2018. additionally, the extent of the sentinel swath fails to image the entirety of the farmington bay region (figure 1 dashed white line), but this is accounted for when comparing to landsat observations by cropping the landsat observations for vegetation to the extent of sentinel 2 tiles. image acquisition and processing is done in the cloud with google earth engine (gee), implemented via the gee python 3 api in conjunction with the geemap python package for interactive mapping and data export (amani and others, 2020; tamiminia and others, 2020; wu, 2020). pre-processed, atmospherically corrected landsat level 2 (tier 1, collection 2) and sentinel-2 level-2a (harmonized) reflectance image collections are defined from the base gee collections, which are then filtered to near-cloud-free images covering the great salt lake region. landsat 5 tm bands are renamed to match landsat 8 & 9 oli specifications, and all the landsat images are merged into the 4 m.h. radwin and b.b. bowen evolution of great salt lake’s exposed lakebed (1984-2023) same collection. cloudy image filtering is accomplished using image cloud percentage metadata provided by the usgs and esa (drusch and others, 2012; foga and others, 2017; tiede and others, 2021), where images with less than 10% of the scene covered by clouds are chosen to process for both landsat and sentinel imagery. for sentinel-2, many images were found to have a significant percentage of badpixels (no data), thus a bad-pixel filter was applied to remove those images. the size and swath path of landsat imagery results in only one complete arm of the lake being imaged for each swath, meaning each arm of the lake is observed on different dates. however, sentinel-2 can image both arms on the same date. for this reason, all imagery results are split between north and south arm. the landsat tile specifications are rows 31 and 32, and paths 38 and 39, while the sentinel tile specifications are 12tum and 12tul. images with the same date are combined to a single image, but images without a paired same-date southern or northern swath image are discarded as that indicates the other scene isn’t suitable and the entire area couldn’t be observed. landsat 1 true-color images from 1972, 1974, and 1979 are used for manual delineation of lake extent to provide a reference prior to the wet 1980’s. landsat 5 images from 1984 are used for manually delineating the extent of the entire great salt lake system, also referencing recent imagery, to be used for masking the data to a boundary and for exposed lakebed area calculations (figure 1). the exposed lakebed is here defined as the area extending from the shoreline to the imposed great salt lake system boundary (figure 1) that encompasses lacustrine derived sediments, evaporites, and vegetation. select mineral operator evaporation ponds within the project-defined boundary of the great salt lake system are not included in the analyses. these areas include the evaporation ponds to the southwest, west and south of stansbury island, and to northeast in the bear river bay, which were established prior to 1984, in addition to evaporation ponds to the northlandsat 5 tm landsat 8 & 9 oli sentinel 2 msi band number band name spectral range (nm) resolution (m) band number band name spectral range (nm) resolution (m) band number band name spectral range (nm) resolution (m) 1 blue 450-520 30 1 coastal aerosol 433–453 30 1 coastal aerosol 433-453 60 2 green 520-600 30 2 blue 450–515 30 2 blue 458-523 10 3 red 630-690 30 3 green 525–600 30 3 green 543-578 10 4 nir 760-900 30 4 red 630–680 30 4 red 650-680 10 5 swir 1 15501750 30 5 nir 845–885 30 5 vegetation red edge 698-713 20 6 tirs 1040012500 120 (30) 6 swir 1 1560– 1660 30 6 vegetation red edge 733-748 20 7 swir 2 20802350 30 7 swir 2 2100– 2300 30 7 vegetation red edge 773-793 20 8 panchromatic 500–680 15 8 nir 785-900 10 9 cirrus 1360– 1390 30 8a narrow nir 855-875 20 10 tirs 1 1060011200 100 9 water vapor 935-955 60 11 tirs 2 1150012500 100 10 swir – cirrus 1360-1390 60 11 swir 1565-1655 20 12 swir 2100-2280 20 table 1. spectral band specifications for landsat tm, oli, and sentinel msi multispectral sensors. 5 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 west. however, the area of the evaporation pond to the northwest, in the north arm, is included in analyses up until the evaporation pond was constructed in 1994. similarly, the evaporation ponds situated in the southern region of bear river bay are employed for satellite monitoring, with the exception of halite detections. this is to circumvent any false readings caused by halite linked with mining activities. other evaporation ponds that exist within the study area are not masked-out and are included in analyses, albeit the remaining ponds are small in comparison to the evaporation ponds removed from analyses. the size of the north arm evaporation pond accounts for ~7.6% of the area within the north arm boundary (figure 1). spectral indices the general mineralogy of the great salt lake exposed lakebed sediments are similar to the sediments in the proximal bonneville basin (a.k.a., great salt lake desert) as both landscapes share a provenance (lake bonneville) and are connected by a spillway. the general mineralogical suite can be simplified to carbonate-rich lacustrine sediments, that comprise the majority of the sediments, which are overlain or interfingered with gypsum (caso4·2h2o) and halite (nacl) evaporite deposits that vary spatiotemporally. the carbonate-rich lacustrine sediments are subsequently referred to as carbonate-muds, as they are typically an intimate mixture of carbonates (including authigenic coatings/cements/nodules, oolitic sands, skeletal fragments, and intraclasts of calcite or aragonite; caco3), quartz grains (sio2), and phyllosilicates (clays/muds), but also may contain magnesite (mgco3), mirabilite (na2so4·10h2o), and other lesser-occurring but still prevalent minerals (lines, 1979; pace and others, 2016; newell and others, 2017; perry and others, 2019; dunham and others, 2020; ingalls and others, 2020; smith and others, 2020; jagniecki and others, 2021; homewood and others, 2022). the grain size distribution as well as proportion of mineralogical components varies spatially for exposed carbonate-muds, but only the surface mineralogy type is considered here (perry and others, 2019). gypsum deposits are found precipitating from springs found within the great salt lake system, but much of the gypsum within the system is likely redistributed rather than actively precipitating, as the great salt lake chemistry is calcium limited and now an mgso4 system (hardie and eugster, 1970; jagniecki and others, 2021). to map these three sediment type classes, each satellite image is processed to mask out other landcover, leaving only surficial sediments, and then each sediment type is differentiated using multispectral indices adapted from work in the bonneville basin mapping similar surface types (radwin and bowen, 2021). to map the extent of water and vegetation, which is used to isolate surficial sediments, the normalized difference water index (ndwi) and normalized difference vegetation index (ndvi) are utilized (mcfeeters, 1996; gandhi and others, 2015; huang and others, 2021). for this study, the halite index takes the form of red – swir1 / red + swir1 and the index for gypsum and carbonate-muds takes the form of swir1 – swir2 / swir1 + swir2. the halite index exploits a significant drop in reflectance from the red (~650 nm) to the swir1 (~1600 nm) bands observed in local halite spectra, which is not observed for the other sediment types (radwin and bowen, 2021). likewise, the gypsum index exploits a slight decrease in reflectance between the swir1 (~1600 nm) and swir2 (~2200 nm) bands observed for local gypsum spectra, which is not typically observed for the local intimate-mixture of carbonates, quartz, or phyllosilicates (carbonate-muds). all resulting images from surface type indices are masked to the surface type of interest using image histogram thresholds. for landsat ndwi results, the threshold is sensitive to sensor-type as well as radiometric differences between scenes, and is determined for each image using an adapted otsu image segmentation technique, which is then offset by +0.15, +0.175, and +0.175 for the north arm, south arm, and bear river bay, respectively (otsu, 1979; ji and others, 2009). the dynamic thresholding is noted to drastically help the accuracy of water detection for landsat imagery, particularly at the water-shore interface. other index results use a static threshold for all images, with differing values for landsat and sentinel to account for differences between sensors. all static thresholds are determined through incrementally assessing how thresholds perform delineation of surface type boundaries, with the goal of having the threshold provide the greatest separation from background values without including background values in the results. for landsat indices, the thresholds chosen are: ≥ 0.345 for halite, ≥ 0.153 for gypsum, < 0.153 for carbonate-muds, and ≥ 0.105 for ndvi. for sentinel, the thresholds chosen are: ≥ 0.58 for halite, ≥ 0.3 for gypsum, < 0.3 for carbonate-muds, ≥ 0.185 for ndvi, and ≥ 0.06 for ndwi. rather than employ a separate index to map carbonate muds, the gypsum index is also used where all unmasked sediments below the threshold used for gypsum are classified as carbonate-muds or other by process-of-elimination. dynamic thresholding for sentinel ndwi images is not applied as there are data-issues associated with bad/no-data pixels that hinder the dynamic threshold 6 m.h. radwin and b.b. bowen evolution of great salt lake’s exposed lakebed (1984-2023) processing for dozens of images with no apparent fix. however, the ndwi threshold for sentinel appears to be less sensitive compared to landsat results. for sentinel-2 msi gypsum, carbonates, and chlorophylla indices, the 10 m/pixel input bands are resampled to 20 m/pixel to match the resolution of the swir bands. to assess relative chlorophyll-a concentrations, the kivu and 2bda indices are used for landsat and sentinel imagery, respectively (gitelson and others, 2003; buma and lee, 2020). different indices are chosen as the sentinel msi sensor is better suited for chlorophyll detection having red-edge bands. the kivu index takes the form of blue – red / green while 2bda takes the form of red-edge-1 / red. processing workflow all images are masked to the correct arm of the lake system prior to processing spectral indices. a systematic workflow is implemented to process each surface type index, where the order of processing follows: 1) water (ndwi), 2) vegetation (ndvi), 3) halite, 4) gypsum, and 5) carbonate-muds/other (figure 2). it is important to note that the results of each index are used to mask the image of the following index, to ensure no pixels are classified twice. for example, the input image for the halite index is masked to be absent of water (ndwi) or vegetation (ndvi) pixels and is theoretically just surficial sediments. the order of processing is chosen as ndwi and ndvi are more standard and broadly applicable spectral techniques that can be used regardless of the surrounding geologfigure 2. workflow chart of methods used to define and process satellite imagery using google earth engine python api. 7 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 ical/mineralogical context, while the sediment indices rely on the isolation of surficial sediments with the mineralogical framework found in the great salt lake and lake bonneville system for the intended performance. the halite index is processed as the first sediment index as it exploits a significant spectral characteristic not found in the other sediments and is believed to be the most sensitive of the mineralogical indices used here. thus, the order of index calculation and image masking follows the most broad-to-limited applicability for the chosen spectral indices and mitigates water or vegetation false-positives for mineralogical differentiations. changing the order of ndwi and ndvi should not have much impact, but the order of the sediment indices matters as the gypsum index can wrongly detect halite pixels as gypsum. for the chlorophyll-a sensitive indices, the ndwi results are used to isolate the data to water pixels prior to processing. all final index results are exported as single-band images and all of the unmasked/output pixels are used to determine the surface area extent of each class. area calculation of each class result requires the use of gee specific area functions to account for the projection of each pixel and calculate the geodesic area of each unmasked pixel. area calculations without accounting for projection are greatly overestimated. all pixel-areas of unmasked pixels for each index result are summed together to estimate the total area of the class. these results are stored and exported as tables for analysis. in contrast, for the kivu and 2bda chlorophyll-a indices, the mean value of all unmasked pixels is calculated for each arm to represent the relative chlorophyll-a concentrations. other data, issues, and error from the index results, the exposed lakebed area is estimated by summing the area of vegetation, halite, gypsum, and carbonate-muds/other, while the calculated exposed lakebed area is estimated by subtracting the water area from the total area of the respective arm of the lake. a calculated exposed lakebed area is also presented for true-color images before 1984, where the water area is manually delineated and subsequently subtracted from the total area. the modelled exposed lakebed area is calculated by subtracting the modelled water area from the total area for each region. erodable exposed lakebed area is calculated by summing the area of the gypsum and carbonate-muds/other classes, as these sediment types are unconsolidated and potentially susceptible to eolian transport. it is assumed that halite and vegetation around the rest of the exposed lakebed aids in retraining sediments from eolian transport by adding a protective surface (reynolds and others, 2007). however, it has been observed that salt crusts may also contribute to dust events if enough desiccation and/or wind occurs (bucher and stein, 2016). additional products presented derived from spectral results are lake area extent boundaries (shapefiles) from select images, as well as a historical halite classification map derived from summing all halite index results for the north arm. lake area extent shapefiles are produced using the output ndwi rasters in arcgis pro, where the rasters are converted to shapefiles, boundaries are dissolved, and all features except the main water body are removed. the historical halite classification map is also produced in arcgis pro by summing all pixel-cell values for all north arm halite images, which effectively produces a historical occurrence map of halite crusts across the lakebed since 1984. the halite values were then classified by value to ten quantiles to form a decile classification map to better assess distribution patterns. daily precipitation data are acquired from noaa station usw00024127 at the salt lake city international airport, which is situated proximal to the southern end of the lake. river discharge data for the bear, jordan – west, jordan – east, and weber rivers are taken from usgs stations 10126000, 10171000, 10170500, and 10141000, respectively. each station is proximal to the lake and roughly represents the river-water influx into the lake system. yearly-running-averages of palmer drought severity index (pdsi) data, a relative dryness/drought indicator using temperature and precipitation data, is acquired for the great salt lake region from 1982-2020 from climate engine using the gridmet drought (4km resolution) dataset. a polygon is used to define the general area of the great salt lake system in climate engine and the mean pdsi value of all gridded pixels within the polygon is calculated then exported. although official cloud percentage metadata are used to filter out cloudy scenes, it is noted that over 30 scenes show excessive amounts of clouds and are removed from analyses. this poor performance of the cloud detection algorithm is shared between both landsat and sentinel products but is infrequent as it occurs in only about five percent of the total amount of images. other issues such as snow, smoke, and surficial-cyanobacteria-growth are observed for a handful of images and those are also excluded from analyses. however, over 15 other landsat images were excluded from analyses due to strange image artifacts, encompassing much of the water body, resulting in a plethora of missing pixels for some or all of the spectral bands. in total, the observations from 80 images are excluded from analyses. given the constraints and limits of manually being able to differentiate surface types from multispec8 m.h. radwin and b.b. bowen evolution of great salt lake’s exposed lakebed (1984-2023) tral satellite imagery, as well as the vast spatial and temporal scope of the study area, one of the only error assessments available is to assess the performance of water-body detection with manually derived comparisons. three locations around the lake are chosen for two separate landsat scenes, and for each region the waterbody is manually delineated and the area is calculated and compared to the area reported by ndwi for the same locations. the magnitude of difference between the results is used as a rough error metric, indicating a difference of <1% for deep waters and difference of ~4% for shallow waters such as the farmington bay. it is observed in many resulting images that when the water in farmington bay is shallow, ndwi has difficulty and typically underestimates the water area. the performance of the vegetation and halite results appear to be very robust in that there is clear separation from background values when assessing the resulting images. additionally, it is worth noting that both indices use a conservative threshold and thus may slightly underestimate the total area of these classes, as it is observed for many images that there is a slight halo around regions of classified pixels with values that could be also included in the class -of-interest as they are well-separated from background values. a temporal model of great salt lake water area is also included as a comparator for water-body detection performance and as an additional source of data. the model is based on a univariate spline interpolation of published values for area vs elevation of the lake, and estimates area via lake elevation data from usgs water-station sites 10010000 and 10010100 (robert, 2005; robert, 2006). the modelled values are a rough estimate as the initial resolution of the lake area data is for every 0.5 ft of lake elevation. however, the interpolation strongly matches the usgs curve as the interpolation utilizes 15 breakpoints (4169, 4171, 4173, 4178, 4183, 4188, 4194, 4200, 4201, 4203, 4205, 4207, 4209, 4211, and 4214 ft; figure s1). sources of differences between the modelled and observed area values primarily stem from differences in the boundaries utilized. the usgs north arm area data does not include the evaporation pond to the west, which is included in this study in analyses until 1994, and the usgs south arm area data includes the large evaporation pond west of stansbury island, which is not included in analyses from this study and accounts for significant differences between the model and ndwi up until the year 2000 (when the water elevation dropped below the level which would naturally inundate the evaporation ponds). nonetheless, the model provides a useful comparison and shows robust agreement with the ndwi results. results & discussion water and exposed lakebed area evolution the resulting time series data show a stark evolution in the surface area of the lake that closely follows the trends from lake elevation data as well as the modelled surface area (figure 3). annual oscillations in lake level are observable for years with more than about three images, as confirmed by the lake elevation and modelled data (figure 3a). sentinel imagery have a much higher temporal resolution and capture annual oscillations in greater detail. after the year 2000, the image-derived and modelled water areas have strong agreement, where the weaker agreement is due to the modelled area including a portion of the south arm salt pool areas for years prior to 2000. the observations between the landsat and sentinel platforms appear to agree well and show relatively little difference. the water surface area for both arms of the lake increase drastically from the 1970’s into the mid 1980’s where the lake filled due to significant precipitation, then slowly decreases in time with only a handful of wet years to follow. for the bear river bay, the water surface area decreased alongside falling lake levels until the early 2000’s when the modelled and observed water area began to diverge. the modelled water area suggests the bear river bay should have been absent of standing water around 2005, but the observed area indicates an anthropogenically maintained water surface area between 200 and 500 km2, with an average of 300 km2. after 2005 the bear river bay water surface area no longer followed trends in lake elevation change. since the maximum extent of 1986-1987, which closely resembles the estimated mean lake area in the absence of anthropogenic consumption (wurtsbaugh and others, 2017), the total observed lake area has decreased from ~5,700 km2 to ~2,590 km2 during the summer months, a loss of ~45%. when considering just the north and south arms, the observed lake area has decreased ~30% from ~3,400 km2 to ~2,380 km2 since 1979. the south arm water area has responded greater to lake elevation change, losing >250 km2 more than the north arm since 1986, with the drying up of the shallow farmington bay being partly responsible. during lowstands the south arm water area is seen to oscillate in greater amplitude compared to the north arm water area, which is coincident with the south arm being directly influenced by seasonal fluxes in river input and association with more shallow lakebed. however, during highstands, when the lake arms are closer to equilibrium in elevation, the water areas fluctuate similarly. in contrast, the water area 9 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 3. evolution of lake surface area (a), lake elevation (b), exposed lakebed surface area and daily precipitation (c), monthly river discharge rate and palmer drought severity index (pdsi) (d), and lake-input depletion data (e) (from wurtsbaugh and others, 2017). the dashed horizontal lines on panel a) indicate the estimated natural mean area of each arm of the lake (corresponding to ~4,207 ft lake elevation) in the absence of anthropogenic consumption (wurtsbaugh and others, 2017). the dashed horizontal line on panel b) indicates the 4,194 ft topographic threshold. also included are color bars indicating times of the anomalous wet period (light blue), west desert pumping project (light purple), and maximum halite crust extents (light green). error bars of -2.5% and +5% are used for lake surface area measurements, as it is more likely to underestimate the observation than overestimate. for the south arm, between 1995 and 2015, the error bars show -2.5% and +10% due to the shallow farmington bay waters. the analyses do not include the north arm salt pool after 1994. arrows indicate specific events in time. the lake and exposed lakebed surface area panels (a, c) include remotely sensed area estimates and the modelled area derived from published surface area vs elevation calculations (robert, 2005; robert, 2006). lake elevation data are from usgs waterstations 10010000 and 10010100. daily precipitation data are from noaa station usw00024127 at the salt lake city international airport. monthly river discharge rate data for the bear, jordan – west, jordan – east, and weber rivers are from usgs water-stations 10126000, 10171000, 10170500, and 10141000, respectively. pdsi data acquired from climate engine for the great salt lake region. 10 m.h. radwin and b.b. bowen evolution of great salt lake’s exposed lakebed (1984-2023) fluctuations within the bear river bay appear to follow greater seasonal and inter-seasonal variations, associated with seasonal flow variations for the bear river and water-management actions. for years with significant rains (figure 3c), where the water elevation has been able to rebound multiple feet, the water area can be seen to dramatically increase, typically by 500-750 km2, between the north and south arms. for example, the wet year of 2011 increased the lake elevation by ~4 ft and north + south arm area by ~670 km2 (figure 4). as the topography of the lake-bottom becomes significantly steeper below ~4,194 ft, water elevation changes below this elevation have significantly less impact to water area (figure s1; robert, 2005; robert, 2006). starting in 2003 the mean lake elevation began to fluctuate near 4,194 ft, which lasted until about 2020, and whenever lake elevation is seen to drop below ~4,194 ft there are noticeably less significant changes in water surface area. knowing that the lake area is figure 4. water boundaries of both the north and south arms for all the major lowstands since 1986 (1995, 2005, 2011, 2016, and 2022) compared to the highstand boundary of 1986. the boundaries show the outermost boundary and do not include interior boundaries such as the boundaries along island perimeters. the southwestern north arm evaporation pool is only included for the 1986 boundary and the bear river bay is not included. a lake elevation plot is inset in the upper right as reference, with the data being from usgs water-stations 10010000 and 10010100. the basemap is the esri world hillshade map with an esri highway layer. the 2005 and 2011 boundaries are close to the ~4,194 ft threshold. 11 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 more sensitive to lake elevation above ~4,194 ft indicates that lake managing efforts should aim to keep the lake at least above ~4,194 ft to maximize the area of the lake and sediment coverage. ideally, when considering maximizing water area (sediment coverage), the water elevation should be kept above ~4,200 ft so fluctuations don’t drop near the ~4,194 ft threshold. maximizing sediment coverage will be increasingly important in the future to mitigate more-significant dust events. a recent report for policymakers determined the optimal range of lake elevation is between 4198 and 4205 ft, with a transitionary zone between 4195 and 4198 ft, based on impacts to air quality, ecosystem, mineral production, recreation, and brine shrimp viability (ahmadi and others, 2023). these proposed elevations align with the presented minimum threshold of ~4194 ft, and if implemented would result in a north + south arm lake area of ~3,100 to ~4,700 km2, roughly 700 to 2,300 km2 greater than the lake area in 2022. associated with the lake surface area change, the total observed exposed lakebed area has increased ~2,985 km2, from ~504 km2 to ~3,489 km2 over 36 years (figure 3c). assessed as a simple trend, this suggests the rate of exposed lakebed area growth has been roughly 80 km2 per year. compared to 1979, before the significantly wet period, the exposed lakebed area for the north and south arms has increased ~1,000 km2, from ~1,600 km2 to ~2,600 km2. since 1986-1987 the south arm has exposed nearly 50% more exposed lakebed compared to the north arm, as the south arm has had a stronger response to water level dropping. however, much of this additional exposed lakebed, particularly in farmington bay, has been altered from a saline mudflat to a vegetated wetlands ecosystem with the rapid encroachment of phragmites. erodable exposed lakebed, exposed lakebed without vegetation or halite crusts to entrain the sediments, has increased from ~330 km2 to ~2,750 km2 since 1986-1987 for the total lake system. erodable exposed lakebed increased by ~900, ~1,110, and ~390 km2 for the north arm, south arm, and bear river bay, respectively, since 1986-1987. the bear river bay has had much less of an increase in erodable exposed lakebed due to anthropogenic maintenance of surface waters and the smaller size of the subsystem area. although vegetation and halite help to protect a sizable portion of the exposed lakebed surface, erodable exposed lakebed has consistently dominated more than 80% of the exposed lakebed surface, except for bear river bay where the average proportion of erodable lakebed surface has been roughly 60%. a caveat associated with vegetation growth protecting the surface is that much of the vegetation in the south arm is due to invasive phragmites, which consume significant amounts of water compared to native vegetation (kulmatiski and others, 2011). precipitation and river discharge data (figure 3cd) help explain major changes to water and exposed lakebed area, where years with significant rains typically result in a much greater river discharge which significantly increase water area and decrease exposed lakebed area. however, years with higher amounts of precipitation but no increase in river discharge (i.e., 2002, 2003, and 2015), associated with river diversion/extraction for agricultural and other uses (figure 3e), are seen to have little effect on the lake/exposed lakebed area (wurtsbaugh and others, 2017; ahmadi and others, 2023). thus, although precipitation directly impacts river discharge, if consumption of the river waters is too great there may be no increase in water/exposed lakebed area and perhaps a decrease. utilizing a yearly-running-mean of the palmer drought severity index (pdsi) emphasizes wetter and drier periods, effectively separating periods with low and high river discharge connected to climatic cycles (figure 3d). the pdsi values of the mid-1980’s and late 1990’s are indicative of wetter periods (>1), which is clear from precipitation and river discharge data, but the mid-2010’s are indicated to be transitional (~0) although discharge into the lake was relatively low. in general, trends from pdsi follow trends from lake elevation and area well up until ~2013, where infrequent but significant precipitation caused the pdsi to slightly rise but the lake elevation and area continued to decline. exposed lakebed evolution results from the spectral indices for vegetation, halite, gypsum, and carbonate-muds highlight key similarities and differences between the north arm, south arm, and bear river bay (figure 5). sentinel and landsat surface classifications agree well, although there are noticeable differences during 2022 where sentinel appears to underestimate the vegetation and evaporite extent. the most significant difference between the sediments shared between the lake regions is that the extent of evaporite formation is magnitudes greater in the north arm (figure 5a-e). carbonate-muds comprise the majority (>75%) of the exposed lakebed for all lake regions and vegetation is typically the second most prevalent land cover type. through the temporal evolution of exposed lakebed area, the percentage of each surface type appears to stay relatively consistent through time, in that there haven’t been any significant changes to the proportion of sediment types as the lake has rapidly 12 m.h. radwin and b.b. bowen evolution of great salt lake’s exposed lakebed (1984-2023) dropped. this is also observed through temporally assessing the percentage of erodable exposed lakebed (figure 5f – dashed line) which consistently oscillates between ~75-95% of the exposed lakebed area for the north and south arms. seasonal oscillations in extent for vegetation and evaporites coincide with wet/cold and dry/warm seasons, as seen by the annual fluctuations of exposed lakebed land cover proportions by ~5-20%. seasonal variation in halite extent appears to be greatest for the bear river bay, as there are spikes of halite detection during the winter months when the surface waters are at a minimum extent (figure 5e). however, the halite variations in the bear river bay are likely overestimated by the sensor as the values appear unreasonably high. overall, evaporites appear to be lesser occurring in bear river bay as compared to the north and south arms, and minimally contributes to the bear river bay lakebed outside of winter months. halite crust formation has been a significant part of the evolution of the north arm exposed lakebed area (figure 5a,f). halite crust in the north arm is formed from either evapoconcentrating pore-waters of surficially saturated sediments or by precipitation of halite in the supersaturated lake waters and accumulation on the lake-bottom. spanning much of the north arm lake-bottom is a robust and thick (>1 ft) halite crust, which becomes partially exposed around the perimeter of the water when the lake recedes (rupke and others, 2016; rupke and boden, 2020). additionally, during the warmer months the waters and saturated sediments on and/or near the fringe of figure 5. percentage of each surface type for the north arm (a-b), south arm (c-d), and bear river bay (e) exposed lakebed areas, split between landsat (a, c, e) and sentinel (b, d) observations, as well as the percentage of erodable exposed lakebed area and detected halite area for the north and south arms (f). the dashed line on f) indicates the percentage of erodable exposed lakebed. 13 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 the water-sediment-interface commonly reach halite supersaturation through evapoconcentration, or are already supersaturated, and form halite crusts that vary in extent depending on a variety of factors (jagniecki and others, 2021). from the satellite observations, the greatest observed extent of halite in the north arm is roughly 150 km2, in contrast to the greatest observed extent of halite in south arm of roughly 30 km2. however, the temporal evolution of halite crust extent in the north arm is complex and the average extent of halite since 1990 is ~78 km2. seasonal fluctuations in halite crust area can vary in magnitude but it is common to see changes greater than 50 km2 during the wet and cold months when halite crusts dissolve and/or when sediments wash in and mask the crust surface. gypsum extent appears to be independent of halite formation, as gypsum extent is observed to vary regardless of halite. however, gypsum extent follows seasonal variations where the greatest extent is during the colder and wetter months, and is most prevalent in the north arm, despite that active gypsum precipitation of significant amounts is unlikely to occur from lake waters. these observations may be attributed to seasonal coverage/reworking by loose sediments or halite crusts, detecting other hydrated sulfate rich minerals (such as mirabilite), or annual cycles of gypsum precipitation from springs or interstitial brines (jagniecki and others, 2021). significant aggregates of mirabilite, if present, are likely classified as gypsum, as their mineralogy and spectral characteristics are similar (kokaly and others, 2017). as gypsum observations are greatest in winter when mirabilite is known to form in the great salt lake system, it is reasonable to interpret that the observations are indeed incorporating detections of mirabilite, which suggests the variations are less in part due to variations in gypsum distribution but rather variations in the combined distributions of gypsum and mirabilite. it is likely a significant portion of the surficial gypsum at a given time is retained from previous years due to redistribution to drier, more protected zones. redistributed gypsum may also be a by-product of evaporative mining in the system. although the south arm forms few halite crusts, gypsum spatiotemporally accounts for an appreciable portion of the exposed lakebed surface, which may be a valid observation or indicate the gypsum threshold is too low as the reported amounts of gypsum are unexpectedly significant. observations from bear river bay indicate a minimal presence of gypsum, a finding that is consistent with the bay's fresher water qualities but may also be associated with local geology, biological mediation, and/or hydrologic processes. vegetation in the great salt lake system spread dramatically starting in the early 1990s, where vegetation in the south arm and bear river bay started growing with rates of ~9 and ~11 km2 per year, respectively (figure 6a). the areas of greatest vegetation growth are associated with the farmington and bear river bays, with the bear river bay hosting the most vegetation. the bear river bay hosts a variety of agricultural, wetland, and floodplain vegetation types while the farmington bay mainly hosts wetland vegetation types. the greatest seasonal variations in vegetation area are attributed to bear river bay, which can vary over 300 km2 (up to >90%) from summer to winter, with the south arm also showing significant seasonal variations. in 2020 the area of vegetation in the bear river bay spiked over 500 km2, 340+ km2 (>300%) greater than pre-1995 observations of vegetation area. the area of vegetation in 2022 is ~400+% greater than the area of vegetation between 1984-1994. vegetation in the north arm shows no significant growth up until around 2010, when vegetation started growing rapidly and quadrupled in area in about 6 years. however, since 2019 the extent of vegetation in the north arm has dropped dramatically. ndvi comparisons between landsat and sentinel agree extremely well, possibly better than any of the other indices used in this study. overall, satellite observations suggest vegetation is rapidly encroaching on the exposed lakebed of the bear river bay and farmington bay. chlorophyll-a analyses represent the mean relative chlorophyll-a concentration for each arm of the lake and shows much different temporal results for both arms of the lake (figure 6b-c). although the microbiology of both arms is greatly different and that many of the organisms don’t produce chlorophyll-a but produce carotenoids (a different biotic pigment), it is expected the chlorophyll-a indices should still capture changes in pigment (weimer and others, 2009; roney and others, 2009; baxter, 2018). the north arm shows a continual decrease in relative chlorophyll-a concentrations through time, having the greatest decreases between ~1992-1995 and ~20122013 (figure 6b). in contrast, the south arm shows a relatively consistent average chlorophyll-a concentration that fluctuates seasonally with variations in temperature, nutrient flux, and turbidity (figure 6c). sentinel 2bda results, which are likely more sensitive to true chlorophyll-a changes due to the inclusion of a red-edge band, capture large seasonal chlorophyll-a fluctuations in the south arm that are much greater in amplitude than changes in the north arm. given that the salinity of the north arm is much greater than the south arm due to a lack of inputs, and that turbidity is much lower in the south arm, the biotic regime is known to be much different and ex14 m.h. radwin and b.b. bowen evolution of great salt lake’s exposed lakebed (1984-2023) plains the differences between the lake arms. reasoning to explain the continual decline of chlorophyll-a in the north arm is that in the 1980s when the lake filled the salinity dropped drastically, nutrient flux increased, and turbidity increased all leading to conditions favorable for microorganism growth. as the north arm has evolved to be more saline, the microorganism community transitioned to saline-favorable organisms and subsequently the less halotolerant microorganisms died (almeida-dalmet and others, 2015; baxter, 2018). additionally, it has been observed that the modern community of microorganisms in the north arm is more resistant to changes in salinity and temperature than in the south arm (almeidadalmet and others, 2015), which may explain the slower rate of observed changes between 1995 and 2013 as well as the smaller magnitude of seasonal changes in the north arm. the landsat and sentinel results agree well for changes in chlorophyll-a concentration in the north arm but appear inversed for the south arm, which may be due to environmental noise or the limitations of the landsat tm and oli sensors to observe changes in chlorophyll-a response of the microorganisms present in the south arm. evolution of halite crusts exposed halite crusts in the north arm were nonexistent during the highstand following the mid-1980s but started forming or becoming exposed in the early 1990s (figure 5f). overall, it appears halite crusts grow in extent as lake levels recede to lowstands (1995, 2004, 2010, 2015, and 2022) and when there is moderate-to-significant annual variations in water surface area (annual redistribution of saline waters to figure 6. evolution of vegetation area (a) for each region of the great salt lake and mean relative chlorophyll-a concentrations for the north (b) and south (c) arms from both landsat and sentinel data. the dark green line of panel a), labeled “south arm – landsat (sentinel bounds)” shows the area of vegetation in the south arm for landsat data that are clipped to the extent/boundary of the sentinel-2 imagery for direct comparison. 15 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 sediment-pore-spaces). in contrast, the halite crusts appear to shrink during periods of wet seasons or extended exposure. following a wet season that dissolved most of the halite crusts in 1998, the lake reached a highstand in 2000 then slowly receded where halite crusts subsequently reached a maximum extent of ~163 km2 in october 2002 and 2003. following 2003 the extent of halite crusts slowly dropped until another wet year during 2011, which quickly diminished halite extent and was followed by receding lake levels through 2016. halite crusts grew again in 2013 to extents similar between 2005-2010 but then started shrinking to the lowest extent in roughly 20 years in 2017. since 2017 halite crusts have been slowly growing again, increasing in size leading up to the lowstand of 2022, but are roughly half the size of crusts observed between 2005-2010 and a quarter of the maximum extent. changes in halite crust extent are observed to partially correspond to significant water management changes. in 2012 the western culvert allowing for flow between the south and north arm was closed and similarly in 2013 the eastern culvert was closed (figure 3b and 5f). the closure of both culverts led to a drop in lake elevation for the north arm of greater than 5 ft as the north arm no longer had any major water input. the rapid drop in lake elevation would have led to exposure of nearshore salt crusts that were previously under water, which is likely responsible for the increase in halite crust area in 2013. subsequent rain and sheetflow events would have progressively dissolved the exposed lake-bottom halite crust, as seen from 2013 to 2016. in late 2016 a causeway bridge was opened to resume flow into the north arm, which resulted in a rapid increase in water elevation and dilution of the north arm water salinity (jagniecki and others, 2021). the significant decrease in halite extent during the early summer of 2017 is likely due to the mixed contribution of rapid water level increase and the influx of fresher waters. rapid water level rise, where the lake rose several feet over the course of a few months, would have inundated and/or dissolved nearshore halite crusts, and fresher water influx undersaturated the water with respect to halite leading to halite dissolution. waters appear to have reached halite saturation by late 2017 into early 2018 as halite crusts reappear (figure 5e). these observations indicate that water management, specifically managing the flow from the south arm to the north arm, has a large impact on halite crust formation processes. aside from direct precipitation (meteoric rain and snow), inundation, and water management, mirabilite formation driven by cold temperatures may be partly responsible for the decreases in halite crust extent, specifically for years where the winter months provided little precipitation but the halite extent dropped significantly. reports have identified that during the colder months mirabilite precipitates from the north arm water column and effectively lowers the salinity of the water to the point where the water becomes undersaturated with respect to halite (jagniecki and others, 2021). this process may cause the lake water and sediment-pore-water to dissolve halite crusts along the shoreline during the winter months, even in the absence of precipitation. spatial distribution of surface types although the time series results provide valuable information regarding the overall evolution of the lake system, the classification map results help understand the distribution of the surface types, which is useful for interpreting the processes responsible for shifts in exposed lakebed composition and cover. the classification results for the north arm show that during the highstand of the 1980s when lake levels were very high there is little exposed lakebed exposed, but what lakebed is exposed is associated with a significant amount of vegetation (figure 7a). following this time, the lake declined rapidly into the 1990s where significant exposed lakebed area appears with sizable halite crusts focused on the northwest sector of the exposed lakebed and much less vegetation (figure 7b). the halite crusts during this period extend roughly 1-4 km from the shoreline and show a close association to proximal gypsum deposits that are likely underlying much of the halite. gypsum appears most prevalent in the north arm during the 1990s but also reappears in similar extent in later years (figure 7b,c,h). the classification maps from 2002 and 2006 show some of the greatest extents of halite, where the map from 2002 shows halite at its near-maximum extent with crusts on average extending 5 km from the shoreline on the western side (figure 7d). additionally, during this period sizable crusts are observed on the eastern side near the location of the spiral jetty. although the lake area in 2011 rebounded to near the 2002 extent, the distribution of halite crusts during and after 2011 is dramatically less and is limited to about 1-2 km from the shoreline (figure 7f). this suggests that the majority of exposed halite crusts in the north arm are formed as part of the lake-bottom crust rather than evapoconcentration of saturated sediment-pore-water, and that the lake-bottom crust didn’t have suitable time or conditions to grow near the 2002 extent during the highstand of 2011. in 2017 the vegetation in the north arm is seen to grow dramatically and the water area decreased, along with a thin 16 m.h. radwin and b.b. bowen evolution of great salt lake’s exposed lakebed (1984-2023) figure 7. north arm classification maps illustrating surface type distributions during 1986 (a), 1992 (b), 1996 (c), 2002 (d), 2006 (e), 2011 (f), 2017 (g), and 2022 (h). 17 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 extent of halite along the western perimeter of the shoreline (figure 7g). the 2022 classification is similar, with a further decline in water extent but perhaps somewhat more halite distributed about the shoreline as well as on the outside perimeter of the salt ponds to the southwest (figure 7h). however, the vegetation growth observed in 2017 is absent in 2022 and there is significantly more distributed gypsum to the northwest. in general, as the water elevation and water area decreased, the halite has been focused around the western side of the north arm shoreline and crusts have slowly fallen in elevation and extent alongside the lake. it is likely that the shallow slope of the lakebed on the western side of the north arm has contributed to the greater observed extents of halite crust. a historical halite distribution map, produced by summing all halite images in the north arm and classifying the image using deciles (ten quantiles), emphasizes the lateral migration of halite crusts through time as the areas where halite crusts repeatedly formed on the western side have much greater values (recurrence of detections) than the surrounding landscape but extend nearly 10-15 km from the modern shoreline (figure 8). recent crusts, which rim the water boundary, show up within the lower decile classes, reflecting less recurrent observations of halite in those areas since 1984. spherical-to-ellipse shaped zones with high pixel values (≥80th decile) on the western side may be local lows that promoted halite crust formation through ponding. although the western side has been the predominant location for halite crust formation, the map shows that halite crusts have formed along the entire shoreline since initial exposure in 1990. years with the most halite appear to correspond to years where there has been a sustained drop from higher-to-lower water elevations exposing the robust lake-bottom crust and/or where the water elevation is above ~4,194 ft such that the exposed lakebed slope is shallower. when the water levels seasonally fluctuate above ~4,194 ft a broader area of sediments can become saturated with saline waters, which should result in more expansive halite crust formation during the summer months when evapoconcentration of sediment-pore-water can form a thin halite crust on the surface. this effect may explain why recent halite crusts have been much smaller than the crusts observed between ~1995-2013, as the water elevations have been on average below ~4,194 ft and the seasonal water area fluctuations are much less. alternatively, the opening of the new causeway breach in 2016, which allows for much greater southto-north flow, could be responsible for the smaller recent crusts, as the waters significantly dropped in salinity and have been slowly recovering. both seem reasonable explanations that can occur in conjunction, however, it appears lake-bottom crust temporally composes the majority of exposed north arm halite crust. thus, the new causeway has likely had a greater impact on recent halite crust formation/exposure than changes to seasonal redistribution of saline waters to sediment-pore-spaces. in contrast to the north arm, classification results for the south arm show a much different distribution of sediment types and vegetation. in 1986 the south arm was very full (figure 9a) but decreased significantly into the 1990s, leading to lakebed exposure and the start of vegetation growth in the farmington bay region (figure 9b-c). relatively little halite is observed in the south arm during the 1990s except for a small crust and associated gypsum to the south. the 2002 and 2006 classification maps (figure 9d-e) show the initial decline of farmington bay waters and indicate some small halite crusts to the south. in 2006 there is a significant increase in gypsum extent that appears to be linked to the gypsum distributions through 2017 (figure 9e-g). the 2011 map shows a significant increase in water and vegetation area, but also highlights water detection issues in the shallow and turbid farmington bay as some of the water area is classified as carbonate-muds and vegetation (figure 9f). the maps from 2017 and 2022 (figure 9g-h) show drastic reductions in water area for the farmington bay alongside slight vegetative growth and a somewhat significant halite crust to the south that is roughly 5 km long and 2 km wide. the water elevation and area during 2022 was the lowest ever recorded. the 2022 classification map also shows errors for shallow and turbid water detection as the outer lateral sides of the water in the farmington bay (now in a channel) are detected as carbonate-muds (figure 9h). conclusions this study processed over 600 reflectance satellite images to better understand the evolution of the water, vegetation, halite, gypsum, and carbonate-mud land cover types in the great salt lake system from 1984 to 2023. the results highlight the magnitude and pace of changes in the system, showing that the exposed lakebed area and halite crust area has responded significantly to lake elevation changes through time. since 1986-1987 the total lake area has decreased by ~45%, from ~5,700 km2 to ~2,590 km2 during the summer months, where the south arm has decreased in greater extent than the north arm. likewise, the exposed lakebed area has increased by ~2,985 km2 over 36 years and reached an area of over ~3,489 km2 in 2022. the bear river bay followed a natural decline in water area up until ~2000, when the 18 m.h. radwin and b.b. bowen evolution of great salt lake’s exposed lakebed (1984-2023) water area diverged from the natural evolution to be anthropogenically maintained near an average surface area of ~300 km2. the critical elevation of ~4,194 ft, where there is a shift in the topographic slope of the lake-bottom, has a sizable impact on the magnitude of water/ exposed lakebed area changes and should be of importance to land-managers and law makers associated with the management of the great salt lake. above ~4,194 ft the lake responds much more significantly to changes in elevation, such that the water area increases significantly even for small changes in water elevation. this is not only important to maintain a healthy size of the lake but to promote evapoconentrative halite formation in saturated sediment-porewaters, as it appears halite crusts have formed in more distributed amounts when the lake fluctuates in elevation above ~4,194 ft and effectively saturates more sediments with a saline brine. furthermore, a greater extent of lake waters promotes expanded lake-bottom halite crust formation, which would be exposed when lake levels recede. years with significant river discharge into the south arm, which can sometimes be associated to years with lower river water consumption rather than higher amounts of precipitation, are observed to rapidly and significantly increase the water surface area, typically by 500-750 km2. this suggests that water conservation efforts, that would lead to a greater annual river discharge into the great salt lake, have the potential to significantly increase the surface area of the lake. halite crusts are predominantly observed in the north arm, where the extent of crusts has undergone a complex evolution since the 1980’s. the maximum extent of halite occurred between 2002 and 2003 in the north arm, with crusts extending over 150 km2. figure 8. decile classification raster produced from the summation of all north arm halite pixel cells between 1984 and 2023, showing the historical halite distribution and areas with most-or-least recurrent halite crusts. values are separated into ten quantiles (deciles), where the largest decile indicates the greatest summation of halite values and the most common historical sites of halite formation. modern halite crust locations, confined near the water boundary, have had significantly fewer recurring observations and are classified in lower deciles. the basemap is landsat 8 oli imagery from june 1st (south image) and 2nd (north image). 19 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 9. south arm classification maps illustrating surface type distributions during 1986 (a), 1992 (b), 1995 (c), 2002 (d), 2006 (e), 2011 (f), 2017 (g), and 2022 (h). 20 m.h. radwin and b.b. bowen evolution of great salt lake’s exposed lakebed (1984-2023) since the peak extent halite crusts have significantly shrunk, related to changes in land-use, lake elevation, and annual fluctuations. the most important control over halite extent appears to be associated with the lake-bottom crust formation/accumulation, topography, and magnitude of seasonal fluctuations. periods of elevated water levels facilitate the restoration and expansion of the lake-bottom halite crust. subsequent receding water levels then enable the exposure of these crusts. greater seasonal elevation fluctuations and shallower topography leads to broader sediment saturation and evapoconcentrative halite crust formation. other important controls that may have impacted the extent and distribution of halite crusts is management of the causeway, where management has affected the salinity and water levels of the north arm waters. additionally, results from this study are consistent with recent findings in the north arm of forced halite dissolution by mirabilite precipitation in cold temperatures, such that there are significant seasonal variations of halite extent even for winters with relatively little precipitation to dissolve the expansive halite crusts. overall, remote sensing techniques to monitor the great salt lake system have been established in this study and provide valuable observations that should be used in conjunction with other monitoring campaigns in the future. future studies should utilize ground truth missions using spectroradiometers and drone surveys to quantify errors using these spectral techniques as well as provide further information on the modern land cover types. similarly, the use of multispectral and active-radar satellites in future studies may help differentiate vegetation types in the great salt lake system. supplemental data the results of the analyses, including supplemental data such as a list of outlier images not used for analyses and the ndwi thresholds for each landsat image, as well as satellite imagery based videos animating the evolution of the lake, are stored on an online database: https://doi.org/10.5281/ zenodo.7996314 or https://zenodo.org/ record/7996314 code utilized in this study for data retrieval and modelling can be found on github: https://github.com/radwinskis/great-salt-lake2023-study-code figure s1. hypsometric data and curves for the north arm, south arm, and bear river bay as shown by published usgs data (dots) and interpolations of the usgs data (solid lines), illustrating the changes in surface area compared to changes in elevation, which is related to the topography of the lake-bottom. the ~4,194 ft threshold is easily seen where the slope of the lines change between 4190 and 4200 ft. from ~4,195 to ~4,201 ft the slope is much steeper, which indicates between these elevations the topography is much shallower. the interpolated lines are formed using 15 breakpoints shown as gray vertical lines. the data and interpolations show to fit very well, supporting the use of interpolation to model lake surface area. https://doi.org/10.5281/zenodo.7996314 https://doi.org/10.5281/zenodo.7996314 https://zenodo.org/record/7996314 https://zenodo.org/record/7996314 https://github.com/radwinskis/great-salt-lake-2023-study-code https://github.com/radwinskis/great-salt-lake-2023-study-code 21 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 references ahmadi, l., albers, e., bingham, b., brooks, p., endter-wada, j., hasenyager, c., lin, j., mcentire, a., 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can be opened with acrobat and adobe reader 5.0 and later.) >> /namespace [ (adobe) (common) (1.0) ] /othernamespaces [ << /asreaderspreads false /cropimagestoframes true /errorcontrol /warnandcontinue /flattenerignorespreadoverrides false /includeguidesgrids false /includenonprinting false /includeslug false /namespace [ (adobe) (indesign) (4.0) ] /omitplacedbitmaps false /omitplacedeps false /omitplacedpdf false /simulateoverprint /legacy >> << /addbleedmarks false /addcolorbars false /addcropmarks false /addpageinfo false /addregmarks false /convertcolors /converttocmyk /destinationprofilename () /destinationprofileselector /documentcmyk /downsample16bitimages true /flattenerpreset << /presetselector /mediumresolution >> /formelements false /generatestructure false /includebookmarks false /includehyperlinks false /includeinteractive false /includelayers false /includeprofiles false /multimediahandling /useobjectsettings /namespace [ (adobe) (creativesuite) (2.0) ] /pdfxoutputintentprofileselector /documentcmyk /preserveediting true /untaggedcmykhandling /leaveuntagged /untaggedrgbhandling /usedocumentprofile /usedocumentbleed false >> ] >> setdistillerparams << /hwresolution [2400 2400] /pagesize [612.000 792.000] >> setpagedevice lerback.pub 1 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2023 utah geological association publication 51 abstract playa margin wetlands in the bonneville basin are sustained by groundwater-fed brackish springs, which transport salts and other solutes into the playa basin. these wetlands are sensitive to changing water availability and quality, which are impacted by changing climate and land use, and whose sediments also provide important records of changing environmental conditions. gastropods building their shells in these springs provide important recorders of water chemistry and may reflect changing aqueous conditions. in this paper, we analyze spring water chemistry, gastropod ecology and gastropod shell chemistry of blue lake (bl) and horseshoe springs (hrs), two groundwater-fed wetlands in the great salt lake watershed. we report the physical parameters including ph, temperature, and specific conductivity across the spring pond at horseshoe springs. there was a slight but statistically significant variation in these physical characteristics between the deeper and shallower parts of the pool, providing evidence that there are different subsite microclimates, which may impact the populations and the isotopic composition of gastropod shells. we measured gastropod population diversity amongst nearly 12,000 shells sampled at horseshoe springs, finding low population diversity (shannon’s diversity index of 0.432), although the populations of shallow and deep snails are slightly different. the dominant snail at hrs is the pyrgulopsis which is imperiled, and we also note that we did not find living snails here. we evaluated the bulk shell variation of stable carbonate isotopes (δ13c, and δ18o) across sites and genera. we show that there were no significant subsite-level differences in gastropod δ13c compositions, suggesting that water depth and productivity were not impacting the isotopic signal. we found subsite and genera-specific differences in snail δ18o compositions, which we interpret to be more dependent on the geography and microclimate of where the snail lived rather than the genera’s physiology (pulmonate versus gil -breathing). we report concentrations of alkali metals (li, na, k, rb, cs), alkali earth metals (be, mg, ca, sr, ba), and metals and metalloids (al, sc, mn, fe, cu, ni, zn, as) at spring site waters and in bulk shells as potential baseline data for interpreting future or past environmental changes as recorded in shell material. we found trace element concentration and certain elemental ratio differences between genera at the same site (particularly of note were li, zn, mn and al) that will be important to constrain if these shells are to be applied as a paleoenvironmental proxy and are sometimes attributed to land use change. keywords: hydrology, isotopes, carbon, groundwater, gastropod, critical zone bonneville basin critical zones: spring chemistry and gastropod ecology in playa-margin wetlands jory c. lerback 1, brenda b. bowen2,4, sam bagge2, mikelia heberer2, ryan cocke2, and hayley l. bricker3 1nuclear and chemical sciences division, lawrence livermore national laboratory, livermore, california, lerback1@llnl.gov 2department of geology and geophysics, university of utah, salt lake city, utah 3department of earth, planetary and space sciences, university of california, los angeles, california 4global change and sustainability center, university of utah, salt lake city, utah introduction earth’s critical zone encompasses the interactions between the biosphere, atmosphere, hydrosphere, and lithosphere from the top of vegetation to the bedrock (u.s. national research council, 2001; anderson and others, 2007; white and others, 2015). the critical zone is linked to anthropogenic activity, from soil formation’s relationship with agricultural production to landscape modifications impacting the hydrologic cycle and water resources (brantley and others, 2007; fan and others, 2019; fovet and others, 2021; minor and others, 2020). groundwater-fed wetlands are, like many critical zone ecosystems, susceptible to changes in climate, water quality, air quality, and other effects of human impacts including recreation, agriculture and urbanization (miguez-macho & fan, 2012; singha & navarre‐sitchler, 2022; torgeson and others, 2022). groundwater-fed wetlands also provide sedimentary records of critical zone (particularly hydrological) processes over time, termed paleo-critical zones by ashley (2020), which help to calibrate and extend the temporal scales by which we understand the feedbacks between groundwater and climate. playa margin wetlands in the bonneville basin are sustained by brackish to saline springs, transporting salts into the playa basin (lerback and others, 2019). louderback and rhode (2009) estimate the discharge rate of 1.6 cubic meters per second or 5x1010 l/yr at 10.31711/ugap.v51i.144 2 j.c. lerback, b.b. bowen, s. bagge, m. heberer, r. cocke, h.l. bricker bonneville basin critcal zones blue lake (bl), one of the two springs in this study. lerback and others (2019) reported a slightly lower annual discharge rate of 1.3x1010 l/yr, and report measured na concentration ranging between 14001600 mg/l from 2016-2018. using these values, and assuming no recirculation of playa solutes in these springs, we estimate that bl annually brings from 2.0�1013 7.5�10 13 mg of sodium to the playa sediments annually (between 22,000-83,000 tons). considering these salt accumulation rates and long timescales since glacial lake bonneville exposed these spring sites, brackish playa margin springs like bl may be an important component of bonneville basin solute budgets. thus, understanding the chemical history and sustainability of these spring wetlands is useful in future work describing the dynamic solute and water budgets sustaining the ecosystems and industries of the bonneville basin. in this paper, we describe two spring-fed wetlands in the relatively under-studied western side of the great salt lake watershed and describe their spring water chemistry by measuring their physical parameters (total dissolved solids, ph, dissolved oxygen, and temperature) and chemical compositions (alkali metals, alkaline earth metals, select metals and metalloids) (figure 1). we use these parameters to establish baseline chemistry for future monitoring of spring ecosystem functioning. importantly, these wetlands foster gastropod (snail) populations, including some endemic genera. gastropod community diversity can serve as a bioindicator of environmental changes, where an environmental change could lead to inhospitable conditions for a relatively homogenous gastropod population (magurran, 1988; hershler and others, 2014). thus, we survey the gastropod communities in these two springs, and provide a baseline of population composition and diversity. gastropod shell chemistry has been shown to record groundwater chemistry and changing aqueous conditions in the present, setting the stage for evaluating near-future environmental changes, and in the past to contextualize modern environmental change (abell, 1985; abell and williams, 1989; rosenthal and katz, 1989; ayliffe and others, 1996). we provide some context for using gastropod shells as proxies for environmental change by investigating the variability of modern shell chemistry, using δ13c and δ18o and the trace elemental composition of shells (alkali metals, alkaline earth metals, select metals and metalloids) in comparison to water. δ13c has been used to reflect changes in the carbon cycle, such as changes in carbon inputs (land-plant versus aquatic humus), photosynthesis, dissolved oxygen content (keith and others, 1964; aravena and others, 1992; jin and others, 2021). δ18o is often used to interpret the water temperatures at the time of carbonate formation (anadon and others 2006; immenhauser and others, 2016). while previous work highlights the complexities of using freshwater gastropods as direct stable isotopic proxies (shanahan and others, 2005), we provide some additional context of differences by genera to understand differences in shell-building processes and potential disruptions to the isotopic utility as paleoenvironmental indicators. shell chemical compositions, particularly trace elements, also have potential conservation applications as the rapidly building shells incorporate trace elements being introduced to the environment. if new material (particularly if containing heavy metals) is introduced (deposited and bioavailable) to the springs due to land use change, urbanization, air quality, or industry, the shell chemistry and ecology may record these changes, serving as sentinels of environmental change (rainbow, 2007; baroudi and others, 2020). additionally, recent work highlights the potential for shells from gill-breathing gastropods preserved within spring sediments to record changes in groundwater chemistry through time using radiocarbon isotopes (lerback and others, 2023). materials and methods site description this study describes two perennial spring wetland sites in northwestern utah, on traditional and ancestral lands of the newe/western shoshone, goshute, and ute peoples. the springs in this study are blue lake (bl) springs (40.502, -114.033) and horseshoe springs (hrs) (40.614, -112.709) in toole county, utah. as reported by lerback and others (2023), bl and hrs spring systems are brackish (with specific conductance measurements above 7000 μs/cm) and mesothermal, with average temperatures between 20° c and 30°c depending on measurement location within the spring pools. these temperatures are higher than mean annual air temperatures of 12°c (lerback and others, 2023). gastropod physiology gastropod genera sampled in this study include melanoides, pyrgulopsis, physella, tryonia, planorbella, and succineidae (figure 2). melanoides melanoides shells found in this study are of the species tuberculate. this paper will refer to mela3 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 1. site description. a) location map of spring wetland sites. b) satellite image of blue lake (bl) wetland at the playa margin. c) satellite image of horseshoe springs (hrs) wetland at the playa margin with the southwest part of the great salt lake. d) schematic of blue lake (bl) from a topdown view with subsites marked. e) schematic of horseshoe springs (hrs) from a top-down view (left) and crosssectional schematic view (right, not to scale) with subsites marked. 4 j.c. lerback, b.b. bowen, s. bagge, m. heberer, r. cocke, h.l. bricker bonneville basin critcal zones noides tuberculate as the genus only for consistency with the other genera described here. melanoides is a prosobranch (gill-breathing), fully aquatic freshwater spring snail native to tropical africa and asia. melanoides was first introduced to north america via the aquatic trade during the 1930s and has continued to spread across the continent into warm regions such as the great basin (murray, 1971). melanoides is an invasive species (dudgeon, 1986; facon and others, 2003; raw and others, 2016). this species likes to burrow into the spring substrate during daylight hours and can therefore be difficult to detect in locations where it has recently been introduced (subda rao and mitra, 1982). they may vary in size from 20 to 40 mm with a lifespan of about 2 – 3.5 years (berry and kadri, 1974; dudgeon, 1982; livshits and fishelson, 1983; pointier, 1989). melanoides is less sensitive to salinity conditions than it is to temperature range, with an optimal growth range of 18 – 31oc (murray, 1971; russo, 1973; roessler and others, 1977; neck, 1985; bolaji and others, 2011). when optimal conditions are consistent and abundant, melanoides may reach population densities of up to 6452 m-2, as was found in a study conducted at fish springs national wildlife refuge by rader et al. 2003. this is attributed to the species reproducing asexually (parthenogenetic), reproducing more than once in its lifetime as well as early in the life cycle (iteroparous), and by developing offspring internally (viviparous). pyrgulopsis pyrgulopsis sp. are one of the largest genera in the family hydrobiidae, which are a family of prosobranch (gill-breathing) snails. they are the second most common hydrobiidae genera in north america, specifically in utah, nevada, and idaho, and are typically found in moist wetland areas such as the benthos of lakes and springs (hershler, 1994). measuring about 1 – 8 mm in shell length, individuals typically cluster with densities greater than 1000 m-2 (hershler, 1994). they may grow to a length of 2.5 mm (hershler and sada, 1987). the temperature range of living specimens falls between 22 – 35oc (hershler, 1994). individuals are typically found near spring groundwater discharge areas (hershler and others, 2014). pyrgulopsis sp. are very sensitive to climatic and environmental changes, which stem from members of this genus diversifying due to their regional separation and isolation; although individual species may live in a range of environments (e.g. temperatures, salinities, co2 concentrations), perturbations to these constant conditions can greatly disturb populations (pearson and others, 2014). they are considered imperiled (turgeon and others, 1998). figure 2. gastropod genera found in this study. a) melanoides, b) pyrgulopsis, c) physella, d) planorbella, e) succineidae, and f) tyronia. 5 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 tyronia tyronia sp. is another genera part of the hydrobiidae family and is restricted to north america. like its fellow hydrobiid genera, tyronia sp. are fully aquatic and typically prefer to inhabit thermal springs. their dispersion is slow and may be linked to drainage history, making them key biogeographical indicator genera (hershler and others, 1999). some species may be quite salinity tolerant (hershler and others, 1999). shells can range between 1.2 – 7 mm in length (hershler and sada, 1987). physella physella sp. are part of subfamily physinae, which are pulmonated (lung-breathing) freshwater spring snails. they are difficult to identify based on morphology alone (young et al., 2021). physella sp. are capable of self-fertilization (parthenogenesis), which may contribute to rapid evolution amplified by isolation or thermally different habitats (perrin, 1986). they typically reproduce annually (russellhunter, 1978). observationally, physella sp. have been known to inhabit waters with temperatures of 8 – 35oc. shell length can grow to 14 mm in very warm water temperatures, indicating that growth is temperature-dependent (mcmahon, 1975). planorbella planorbella sp. are part of family planorbidae and is a freshwater gastropod genus restricted to north america (baker, 1945). members of this genera are hermaphroditic and may rely on self-fertilization for reproduction (martin and others, 2020). they have been shown in studies to be optimally active between 26 28oc, with minimum and maximum optimal thresholds appearing to occur at 18oc and 33oc, respectively. (el-emam and madsen, 1982). succineidae succineidae are a family of minute taxa of pulmonated (lung-breathing) land snails that typically inhabit wetland areas worldwide (pilsbry, 1948; patterson, 1971). genera are typically found on vegetations near streams or marshes, or where dew might be present. there may be extreme differences in morphological features such as size and shell shape, between genera within succineidae. they are hermaphroditic and can reproduce through mutual fertilization or self -fertilization. experimental design we measured physical and chemical parameters using a multiparameter probe in the spring to understand circulation within the spring pond. we measured probe depth (m), total dissolved solids (tds in ppt), ph, dissolved oxygen (do in mg/l) and temperature (°c) using an aqua troll 600 multiparameter sonde. we recorded these basic physical and chemical parameters of the spring water along a nwse transect at hrs to understand the structure of flow and potential circulation within the northern spring pool and stream outlet (figure 3). the probe recorded the time, which we marked and calibrated to locations marked at 0, 150s, 200s, 250s, 300s, and 345s. time is used as a proxy for the distance along the transect, as it is not a linear transect. water samples were collected at the water surface for trace element analyses in high-density polyethylene (hdpe) bottles that were washed with 5% hcl and rinsed three times with deionized water. samples were filtered with a 0.45 μm polypropylene syringe filter and stored with minimal headspace. at hrs, we sampled bulk sediment at sites hrs1, hrs-2, hrs-3, and hrs-4 to measure the diversity and density of snail populations. at these four subsites in shallow and deeper waters, we filled one 0.25 l container with bulk sediment. samples were cleaned at the university of utah by sieving and rinsing bulk samples with deionized water and soaking the shells in 3% hydrogen peroxide for one hour which served to separate the sediment and organic material from the shells. snails in sediment samples at each site were identified to the genus level and counted to understand the diversity of taxa across subsites at hrs. gastropods were collected under a research agreement with the utah division of wildlife resources (4coll10642). we used shannon’s diversity index to test the diversity of gastropods at hrs, which is a common metric of ecological diversity that takes into consideration the richness and evenness of each of the genera or species collected (clarke and others, 2014). the equation from shannon (1948) is shells prepared for chemical analysis were cleaned with 3% hydrogen peroxide for one hour to remove organic material, rinsed with deionized water, and then sonicated to further remove organic matter and sediment. bulk shells were homogenized individually using a mortar and pestle. four shells were selected to be subsampled along transects from tip to (1) 6 j.c. lerback, b.b. bowen, s. bagge, m. heberer, r. cocke, h.l. bricker bonneville basin critcal zones aperture (“intrashell transects”) to assess variation in shell chemistry over the snail’s lifetime. intrashell transects were collected at four evenly spaced subsites along the long axis of the shell using a microdrill. melanoides shells from bl-spring were selected for the intrashell transects due to their relatively larger size. a total of sixty-four whole gastropod shells (37 from bl and from 27 hrs) and four sub-sampled shells from bl were analyzed at the sirfer laboratory at the university of utah. samples were reacted with orthophosphoric acid and analyzed as co2 after cryogenic purification. samples were analyzed on a finnigan mat 252 mass spectrometer. data are reported using delta notation relative to the vienna pee dee belemnite (vpdb) standard for carbonates and water δ13c and the vienna standard mean ocean water (vsmow) for water δ18o, where analytical precision for δ13c and δ18o was ~0.1‰. δ18o-vsmow values were converted to δ18o-vpdb to directly compare δ18o of water and shells. an additional 22 shell samples (17 from bl and from five hrs) were added to the dataset here from lerback and others (2023), where their data were made using the same methods. analyses including calculation of mean and standard deviations (sd), and statistical tests including analyfigure 3. physical parameters of water along a transect at hrs. a) schematic maps of hrs with probe transect locations and select times marked in the left panel, and the bathymetric distinction made between shallow and deep areas of the spring pool on the crosssectional view in the right panel. b) probe depth along transect. c) total dissolved solids (tds) along transect, d) ph along transect. e) dissolved oxygen (do) along transect. f) water temperature along transect. 7 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 sis of variance (anova) and tukey’s honestly significant difference (hsd) were performed in r v3.6 in rstudio. comparisons between water and shell isotope values evaluated the fractionation (ε), which is the difference between shell and water isotope values. eight shells (four succineidae from bl, three pyrgulopsis and one tryonia from hrs) and four water samples (two from bl and two from hrs) were analyzed for trace element concentrations at the university of utah's strontium isotopes laboratory using an agilent 7500ce quadrupole inductively coupled plasma mass spectrometer (icp-ms). results and discussion spring water we evaluated the multiparameter probe data from hrs to understand how the aquatic environment varies within the spring system. we divided the hrs multiparameter probe data from hrs into shallow and deep sections of the pond at time 150s (of a total of 345s recorded) due to the relatively steady depth of the probe measurements (0.5 m depth in the nw section, and 1-1.5 m depth in the se part of the transect where the shore edge became steep) and the variable depths of the spring pond (table 1). over the transect, tds ranged from 1.7 in the shallow pool to 5.0 ppt in the deeper pool. average tds values were 3.6 ppt (sd = 1.2) for the shallow pool and 4.8 ppt (sd = 0.1) for the deeper pool. this difference was significant where (t (151) = 12.6; p < 0.01). the decrease in tds values downstream (from shallower to deeper) within the spring may result from fresh water discharging to the shallower spring pool or evapoconcentration in the deep pool. the water ph increased through the transect, with a total mean of 7.8 (sd = 0.08), ranging from a low of 7.7 in the deep section to a high of 8.2 in the shallow areas. the average of the deep section was 7.8 (sd = 0.01), and the shallow section was 7.9 (sd = 0.07), with a significant difference (t(162) = -22.5; p < 0.01). do content gradually increased from 7.45 mg/l (sd = 0.22) to 8.1 (sd = 0.78) as the water flowed into the shallower region. this difference was significant where (t (170) = 15.5; p < 0.01). the temperature was relatively elevated through the transect around 19.5℃ (sd = 1.19). the water temperature in the shallow section of the spring was an average of 18.8℃ (sd= 1.59. the deep section of the pond water 20.0℃ (sd = 0.08), which is warmer than the shallower section (t(148) = 9.0; p < 0.001). data for these measurements are provided in appendix 1. we measured alkali metals (li, na, k, rb, and cs), alkaline earth metals (be, mg, ca, sr, and ba), and select metals and metalloids (al, mn, fe, zn) in spring waters (figure 4 and appendix 2). overall, alkali metal concentrations were more abundant at bl than at hrs and were highest at the bl-marsh site, which is likely due to evapoconcentration in the shallow standing water. al, mn, and fe concentrations in hrs and bl were the 0.04 mg/l detection limit. although this water is not designated for human consumption, it is worth contextualizing these values as below the national secondary drinking water standards of 0.05-0.2 mg/l, 0.05 and 0.3 mg/l, respectively (u.s. environmental protection agency, 2009). although these natural brackish springs are not used for drinking water, we note that these analyses did not have high enough resolution to detect whether the concentrations were below the national primary drinking water regulations maximum contaminant level for as of 0.01 mg/l (u.s. environmental protection agency, 2009). gastropods ecological diversity biodiversity is important for protecting the stability of the community which can aid in the overall recovery time from ecological harm that may threaten an ecosystem (e.g., natural disasters, famine, and diseases) (magurran, 1988). we counted nearly 12,000 gastropod shells across four subsites at hrs, and we note that we did not find any living specimens with organic tissues which needed to be cleaned. we believe that the sampled shells are relatively modern parameter  shallow  deep  total  probe depth (m) mean = 0.54 mean = 0.95 mean = 0.77 tds (ppt) mean = 3.6, sd = 1.21 mean = 4.8, sd = 0.14 mean = 4.3, sd = 1.01 ph mean = 7.9, sd = 0.07 mean = 7.8, sd = 0.02 mean = 7.8, sd = 0.08 do (mg/l) mean = 8.4, sd = 0.73 mean = 7.4, sd = 0.23 mean = 7.8, sd = 0.69 temperature (°c) mean = 18.8, sd = 1.59 mean = 20, sd = 0.08 mean = 19.5, sd = 1.2 table 1. water physical parameters: probe transect data summary. 8 j.c. lerback, b.b. bowen, s. bagge, m. heberer, r. cocke, h.l. bricker bonneville basin critcal zones (i.e., not representing shells last alive thousands or hundreds of years ago) based on their sampling location at the surface of the spring sediments but recognize some might represent older shells that could have been brought to the surface by sediment disturbances, e.g., fish burrows in the spring sediments. we counted the number of individuals in each genus to measure the diversity of genera in the ecosystem (table 2). the shannon’s diversity index at hrs (combining subsites) was 0.432. while the index theoretically ranges from zero to infinity, this value is low compared to other studies where shannon’s diversity index often ranges from 1.5-3.5 (magurran and mcgill, 2011; ifo and others, 2016). we measured gastropod population diversity differences between the subsites (figure 1e), which we further grouped into the shallow and deep sections following the distinctions shown in figure 3a. subsites hrs-2 and hrs-3 are considered shallow, and subsites hrs-0, hrs-1, and hrs-4 are considered deep (although hrs-0 was not sampled to characterize gastropod diversity). the shallow sediment samples yielded a higher density of shells, where 76% (n = 9087) of individual shells counted were from the shallow samples and both shallow and deep samples had the same volume of sediment collected. pyrgulopsis and tyronia were the most common genera found, with a few physella (n = 18) found in both shallow and deep subsites. while different genera, pyrgulopsis and tyronia are both members of the same gastropod family, and their joint presence may be due to shared preference for similar environmental conditions. a chi-squared (χ2) analysis of the ob figure 4. trace element concentrations of springwaters. marsh waters have relatively higher elemental concentrations due to evaporation and relatedly less water input. genera hrs‐1  (deep) hrs‐2  (shallow) hrs‐3  (shallow) hrs‐4  (deep) pyrgulopsis 2404 2149 5253 334 tyronia 112 72 1596 47 physella 1 0 17 0 table 2. gastropod population: count of individuals in sampled community. 9 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 served number of pyrgulopsis, tyronia, and physella genera counted from the shallow versus deep subsites resulted in a χ2 of 219 (p < 0.05), thus indicating a statistically significant difference (albeit small) between populations in the small and deep parts of the hrs pond. δ13c and δ18o shells were analyzed for δ13c and δ18o and compared to total dissolved inorganic carbon (tdic) δ13c and water δ18o to understand the potential effects of environmentally related and biomediated carbon isotope fractionation (appendix 3). δ13c and δ18o metrics can be representative of environmental conditions during the time of snail activity and can be used as an indicator of source carbon and source water, useful for evaluating environmental changes in the past (recorded in sedimentary records), and in future collections. shell samples were aggregated by subsite at bl and hrs (figure 5a, table 3). at bl, the subsite avfigure 5. stable isotope measurements of gastropod shells. a) δ13c and δ18o measurements of bulk, homogenized shells by site and subsite. b) εc and εo of bulk, homogenized shells by genera. c) δ13c and δ18o variation within intrashell transects on melanoides shells. d) modelled shell formation (ambient water) temperatures based on δ18o composition by site and genera. 10 j.c. lerback, b.b. bowen, s. bagge, m. heberer, r. cocke, h.l. bricker bonneville basin critcal zones erage δ13c values of shells ranged from -2.1 to -3.5‰ (a range of by 1.4‰), whereas the δ13ctdic values in water samples at bl averaged -4.2‰ (sd = 0.43, n=4) (lerback and others, 2023). an anova shows that there are no significant differences between subsite δ13cshell values (f = 0.765, p = 0.519). at bl, the subsite shell averages of δ18o values ranged from 11.9 to -17.3‰ (a range of by 5.4‰). subsites had statistically different values (anova f= 36.18, p < 0.01), where a tukey’s hsd test showed that pairwise, the bl-marsh was different than every other site at p < 0.01 with a difference of more than 4‰). this is likely because the marsh had shallow standing water which may have had significant evaporative effects. water samples (converted to vpbd from vsmow) of bl discharge averaged -45.5‰ (sd = 0.1, n=12). data for water stable isotopes at these sites is provided in appendix 4. at hrs, the subsite averages of δ13c ranged from -4.5 to -5.2‰ (a range of 0.8‰), whereas δ13ctdic values from water at hrs yielded a value of 7.2‰ (n = 1) (lerback and others, 2019). the δ13ctdic values were not statistically different across sites (anova f = 0.479, p = 0.75). at hrs, the shell subsite averages of δ18o (vpdb) values ranged from -14.7 to 15.9‰ (a range of 2.6‰). the δ18o of shells at hrs was statistically different between sites (anova f = 32.28, p < 0.01), where a tukey’s hsd test showed that pairwise, the hrs-0 was different than the other sites at p < 0.01, with a difference of 1‰).water at hrs had a measured value (converted to vpbd from vsmow) of -45.5‰ (sd=0.1, n=8). δ13c values of bulk sediment were -26.6 and 20.7‰ at bl-pond and hrs-0, respectively (lerback and others, 2023), which are within the range of values expected of plant material in the region (hart and others, 2010). shell δ13c values are more reflective of water δ13c (tdic) than sediment, similar to findings by fritz and poplowski (1974). stable isotope data of shells were also aggregated by genera (figure 5b, table 4) δ13c (vpdb) values in shells ranged from -1.7‰ of a tryonia at bl to 9.0‰ from a planorbella at hrs. significant differences existed in δ13c between genera (anova f = 7.172, p < 0.01), where the tukey hsd found a difference of greater than 1‰ between melanoides and all other genera (p < 0.1) but no significant differences between the four other genera (p > 0.1). δ18o (vpdb) values in shells range from -10.0‰ of a succineidae at bl-marsh at bl to -17.5‰ of a melanoides at bl-spring. we found differences between the genera δ18o (anova f = 45.41, p < 0.01), where the tukey hsd showed a pairwise difference between melanoides and tyronia (difference of 1.3‰, p < 0.01), and between melanoides and pyrgulopsis (difference of 1.9‰, p < 0.01). there was a difference of at least 4‰ between succineidae and all other genera including melanoides (p < 0.01). lastly, there was also a difference of 1.3‰ between planorbella and pyrgulopsis (p < 0.1). the fractionation between the measured shell and the spring waters is represented by epsilon (ε) for shell-tdic in carbon and for shell-h2o in oxygen stable isotope values (element denoted with a subscript). water δ18o for bl and hrs was reported relative to a vsmow standard, which we convert to the vpdb standard before calculating the fractionation values for oxygen (εo). planorbella and succineidae have εc near 0, while the other shells (melanoides, physella, and tyronia) show εc of greater than 1.5‰. all genera but succineidae, had an average εo of 29.6 (sd = 1.2, n = 57). the elevated εo of +34‰ (sd = 0.44, n = 6) found in succineidae is unsurprising because succineidae is a genus only found at blsite average of δ13c sd of δ13c count of δ13c average of δ18o sd of δ18o count of δ18o bl‐lake ‐3.01 0.41 8 ‐17.13 0.16 8 bl‐marsh ‐3.48 1.44 10 ‐11.91 1.85 10 bl‐pond ‐2.85 1.10 21 ‐16.41 1.12 21 bl‐spring ‐2.13 0.33 14 ‐17.31 0.46 14 hrs‐subsite‐00 ‐5.22 1.06 15 ‐15.89 0.20 15 hrs‐subsite‐01 ‐5.06 0.23 4 ‐14.82 0.41 4 hrs‐subsite‐02 ‐4.72 0.25 4 ‐14.77 0.14 4 hrs‐subsite‐03 ‐4.46 0.45 4 ‐14.68 0.35 4 hrs‐subsite‐04 ‐4.77 0.38 4 ‐14.79 0.17 4 table 3. gastropod shell δ13c and δ18o data by site. 11 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 marsh, where evaporation is occurring in shallow standing water (morgan, 1970). shanahan and others (2005) posit that succineidae and physella are pulmonates (lung-breathing), enabling them to live in these shallow areas where more evaporation is occurring and seasonally changing the water δ18o. following the methodologies provided by shanahan and others (2005), we also evaluated the variation within the whorls of a single shell. the withinshell (intrashell) isotopic variation here may be due to biomediated fractionation or due to seasonal variation in waters during stages of growth. we examined the isotopes in intrashell transects along the long axis of growth from shell aperture to tip (figure 5c, appendix 5). these variations are otherwise averaged by homogenizing bulk shells. variation of δ13c was less than 1‰ through the shell, where three out of four shells increased by almost 1‰ overall. for δ18o (vsmow) values, the maximum variation observed was a linear decrease in one shell from -16.5‰ to 17.5‰ along the transect, whereas the other three shells did not vary more than 0.3‰. the variation within shells indicates that intrashell variation is minimal compared to the variation between bulk shells. this likely reflects the stable conditions provided by the mesothermal, seasonally stable discharge at bl. we used the measured shell δ18o values and measured average water δ18o (which are assumed in paleoclimate studies) to calculate the expected shell formation temperatures (figure 5d, appendix 6). like shanahan and others (2005), we compare these estimated temperatures to measured temperatures of the springs reported by lerback and others (2023). the fractionation equation from kim and o'neil (1997) estimates temperatures for synthetic calcite, where kim and others (2007) estimate the formation temperature for synthetic aragonite, and finally, the equation by white and others (1999) estimates the formation temperature for aragonitic molluscs. each of these equations predicts a temperature lower than the measured discharge temperatures of 28℃ at bl and 21℃ at hrs, but predicted temperatures for the genera melanoides, physella, planorbella, and tyronia by white and others (1997) are within the range found at bl between the spring subsite discharge and the lake subsite, and this equation also is closer to the mesothermal temperatures at hrs as well. because physella, planorbella and succinidae are all pulmonates, we might expect these to have more similar predicted temperatures, and different as compared to the gil-breathing genera. however, physella and planorbella predicted temperatures are more closely aligned with the gil-breathing genera, so the δ18o may be more dependent on the microclimate associated with the location than the genera-specific vital effects that shanahan and others (2005) discussed. trace elements we measured alkali metals (li, na, k, rb, and cs), alkaline earth metals (be, mg, ca, sr, and ba), and select metals and metalloids (al, mn, fe, zn) in gastropod shells from hrs and bl to understand how shell chemistry may represent water or environmental chemistry (figure 6, appendix 7). on average, higher concentrations of li, k, rb, be, sr, mn, zn, and as were found in shells from bl than hrs. we provide these data to develop some baseline values as shells can be used as bioindicators of environmental change but note these associations need to be studied in more detail for genera-specific biases. notably, the single tyronia shell from hrs had concentrations of na, rb, cs, mg, ba, al, and mn distinctly higher than shells from the same location, although the data are too sparse to draw statistical significance. however, the difference in trace element concentration may indicate there may be phylaand genera-specific differences in how elements will bioaccumulate in body materials (including shells) (langston and others, 1998; rainbow, 2007). bolotov and others, (2015) showed that freshwater bivalve trace element concentrations are significantly impacted by biological shellbuilding processes and geography (water elemental concentrations as related to the proximity of chemical sources). land snail shell incorporation of environmental trace elements have also been discussed as bigenera average of δ13c sd of δ13c count of δ13c average of δ18o sd of δ18o count of δ18o melanoides 1.71 0.46 26 ‐0.95 1.19 15 physella ‐0.70 na 2 ‐0.64 na 2 pyrgulopsis 2.13 0.31 27 1.03 0.81 19 planorbella ‐0.10 1.10 6 ‐0.31 0.75 5 succineidae ‐0.06 1.33 6 5.41 0.44 6 tryonia 2.30 0.26 16 0.41 0.89 16 table 4. gastropod shell δ13c and δ18o data by genera. 12 j.c. lerback, b.b. bowen, s. bagge, m. heberer, r. cocke, h.l. bricker bonneville basin critcal zones omonitors for changing environmental conditions (de vaufleury and pihan, 1999; madejon and others, 2013; pauget and others, 2013), but need to be more carefully studied because of high soil variability and complexity of ecosystems. we show that spring snails throughout hrs ponds live in relatively chemically homogeneous environments, and thus we believe aquatic snail shells may be represent more consistent environmental proxies (within same-genera groups) than land snails. we calculated the elemental ratios (mol/mol) of shells and water that have been evaluated in gastropod shells for paleoclimate reconstructions (figure 7). these include mg/ca which have been used for temperature reconstruction in marine and lake foraminifera collections (nurenberg and others, 1996; lea and others, 1999; elderfield and ganssen, 2000; dekens and others, 2002; anand and others, 2003; tripati and others, 2003khider and others, 2015; gray and evans, 2019; saenger and evans, 2019). the study of mg/ca ratios in bivalves and gastropods have been limited and focus primarily on marine systems (wanamaker and others, 2008; garcía-escárzaga and others, 2015). ulrich and others (2021) found that there was strong association between biomineral elemental chemistry and shell-building genera relatedness and that amongst the marine gastropods that were studied, element incorporation patterns arose at the class level. our data test whether the mg/ca relationship works in the select freshwater gastropod taxa. at bl, waters had higher mg/ca ratios than were observed in succineidae shells (approximately 0.8 mmol/mol in water, and 0.0 mmol/mol in shells), whereas in hrs shell the mg/ca values were all very low, less than 2.0 mmol/mol. dellinger and others (2018) show data for marine mollusks that confirm lower mg/ca ratios (0.3-8 mmol/mol) than can be expected for more aragonitic materials. using the calibration equation from anand and others (2003) (which was derived for marine foraminifera which incorporates source water figure 6. trace element concentrations measured in gastropod shells. 13 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 mg/ca geochemistry), the expected mg/ca ratio for these shells is 3.011 and 2.515 mmol/mol, respectively (using the water temperatures of 23°c at bl and 21°c at hrs reported by lerback and others, 2023). given the measured ratios of 3.1 and 4.3 mmol/mol for bl and hrs, these spring water gastropod genera record very different mg/ca ratios that may reflect the gastropod shell-building processes partitioning of elemental species, especially in comparison to the marine foraminiferal calibrations that have been previously reported. therefore, freshwater gastropods are likely unsuitable for temperature reconstruction with the mg/ca paleothermometer without further study. gastropod elemental ratios of fe/mn and mn/sr have been used as a proxy for a variety of processes, such as changes in water chemistry including sediment input, redox conditions, and water balance fluctuations (rosenthal and katz, 1989; wanamaker jr and others, 2008; korponai and others, 2010). more specifically, fe/mn and mn/sr have been used in sedimentary records as indicators of redox conditions at the time of deposition and may also indicate potential alteration via diagenesis (templeton and others, 2000). examples of such factors affecting fe/mn and mn/sr ratios are changing mixing regimes, erosional input of sediment, aquatic productivity, soil leaching, and eutrophication. water fe/mn ratios were not reported because concentrations were below the instrument detection limit. the fe/mn values of succineidae shells found at bl range from 0.21 to 0.85 mol/ mol, whereas the fe/mn values of pyrgulopsis and single tryonia collected at hrs range from 2.08 to 31.56 mol/mol (figure 7, table 8). we observed within-genera variability even within the same site; one pyrgulopsis shell from hrs has fe/mn and mn/sr values that show an elevated signal compared to the two other pyrgulopsis samples, falling nearer that tryonia sample from the same site. we also observed some subsite variation; the succineidae samples from bl show a depleted fe/mn signal relative to sampled marsh surface water, and an elevated mn/sr signal when compared to both marsh surface and pond surface water. this elevated mn/sr may reflect preferential uptake or more bioavailability of the trace element mn as compared to the more abundant elements. figure 7. trace element ratios in waters and shells. 14 j.c. lerback, b.b. bowen, s. bagge, m. heberer, r. cocke, h.l. bricker bonneville basin critcal zones conclusions wetland critical zones are important sites to monitor for changes to water resources and biodiversity and are sites for sedimentological preservation of wetland critical zone processes. the water chemistry shapes wetland critical zones over long time periods, and the chemistry provides solutes that contribute to the saline ecosystems of the bonneville basin. aspects of the water chemistry are preserved in gastropod shells, which can (1) provide an ongoing observational metric as “sentinel” organisms, rapidly capturing chemical changes to the system, and (2) also which are preserved in sediments, providing historical records of the magnitude of change in spring chemistries which contextualize modern environmental changes. we provide chemical characteristics of modern water from two mesothermal, playa-margin springs in the western bonneville basin, and evaluate gastropod chemistry as providing potential records of spring chemistry changes through time. the shannon’s diversity index at four hrs subsites indicates that the ecological diversity is low, with one snail genera, pyrgulopsis (an imperiled genera) being the dominant shell found (magurran, 1988; magurran and mcgill, 2011). while more research (particularly into paleoenvironmental conditions and past population distributions) can clarify a baseline diversity in these springs, the low diversity index reported here may reflect that the ecosystem’s overall stability may be sensitive to environmental changes, including land use change that impacts the geochemical profile of the water. the modern chemistry of the shells can also be used as a baseline to compare with sedimentary shell records or future shell collections as bioindicators of environmental change. we present isotopic and trace element data among subsites and genera to constrain the variables relevant to scientists interested in using shells as proxies for spring water changes through time. we found genera and subsite differences in δ13c and δ18o variations and trace element chemistry of modern water and shells, which will need to be better constrained in order to effectively use these chemical relationships to interpret past environments. our data did not find evidence for significant physiological differences based on pulmonate versus gil-breathing genera in the stable isotope data. overall, this paper describes wetland critical zone chemistry and metrics of biodiversity in a system of ecological importance in the great salt lake watershed. as gastropods deposited in spring sediments can be used as recorders of environmental change, we evaluate the factors that may impact geochemical preservation and thus environmental reconstructions. acknowledgments thanks to kate holcomb, lisbeth louderback, don sada, and saxon sharp for assistance with gastropod identification. thank you to reviewers jay quade and jesse bateman for their helpful comments, which greatly improved this paper. this work is partially supported by the global changes and sustainability center at the university of utah, the nsf/gsa graduate student geoscience grant #12745-20, 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(eds.), great salt lake and the bonneville basin: geologic history and anthropocene issues: utah geological association publication 51. http://pubs.er.usgs.gov/publication/70162654 http://pubs.er.usgs.gov/publication/70162654 https://doi.org/10.1111/j.1365-3091.1975.tb00244.x https://doi.org/10.1111/j.1365-3091.1975.tb00244.x appendix 1. water physical parameter data from probe transect. time group depth (m) temperature (°c) total dissolved solids (ppt) ph rdo concentration (mg/l) specific conductivity (µs/cm) 1 shallow 0.373 11.27 2.5 8.1 11.68 3852.72 2 shallow 0.383 11.27 2.64 8.13 11.78 4067.68 3 shallow 0.393 11.28 2.78 8.15 11.88 4282.65 4 shallow 0.403 11.28 2.92 8.17 11.99 4497.62 5 shallow 0.362 15.12 2.52 8.15 10.2 3869.74 6 shallow 0.363 15.32 2.55 8.16 10.14 3916.6 7 shallow 0.365 15.53 2.58 8.17 10.08 3963.46 8 shallow 0.461 16.38 2.79 8.03 9.7 4286.13 9 shallow 0.465 16.5 2.78 8.02 9.64 4282.09 10 shallow 0.469 16.62 2.78 8.01 9.58 4278.05 11 shallow 0.5 17.22 2.73 7.98 9.21 4197.18 12 shallow 0.504 17.26 2.73 7.97 9.19 4199.55 13 shallow 0.508 17.3 2.73 7.97 9.17 4201.92 14 shallow 0.526 17.5 2.5 7.95 9.03 3849.93 15 shallow 0.527 17.52 2.49 7.95 9.01 3828.27 16 shallow 0.528 17.54 2.47 7.95 9 3806.61 17 shallow 0.53 17.55 2.46 7.95 8.99 3784.95 18 shallow 0.546 17.76 2.41 7.93 8.83 3700.48 19 shallow 0.547 17.77 2.4 7.93 8.82 3688.59 20 shallow 0.548 17.78 2.39 7.93 8.81 3676.71 21 shallow 0.53 17.95 2.39 7.93 8.69 3684.06 22 shallow 0.529 17.97 2.39 7.93 8.68 3683.68 23 shallow 0.528 17.98 2.39 7.93 8.67 3683.31 24 shallow 0.554 18.06 2.44 7.93 8.61 3746.17 25 shallow 0.555 18.07 2.44 7.93 8.6 3750.11 26 shallow 0.555 18.07 2.44 7.93 8.6 3754.06 27 shallow 0.497 18.2 2.37 7.93 8.51 3643.82 28 shallow 0.494 18.21 2.37 7.93 8.51 3638.84 29 shallow 0.491 18.22 2.36 7.93 8.5 3633.87 30 shallow 0.488 18.23 2.36 7.93 8.5 3628.89 31 shallow 0.532 18.3 2.44 7.93 8.48 3758.54 32 shallow 0.533 18.31 2.45 7.93 8.48 3763 33 shallow 0.534 18.32 2.45 7.93 8.47 3767.47 34 shallow 0.511 18.44 2.46 7.93 8.4 3780 35 shallow 0.511 18.45 2.46 7.93 8.39 3783.48 36 shallow 0.511 18.46 2.46 7.93 8.39 3786.96 37 shallow 0.524 18.53 2.37 7.93 8.37 3649.47 38 shallow 0.524 18.54 2.37 7.93 8.37 3641.83 39 shallow 0.524 18.54 2.36 7.93 8.37 3634.18 40 shallow 0.531 18.63 2.38 7.92 8.35 3661.48 41 shallow 0.531 18.63 2.38 7.92 8.34 3660.27 42 shallow 0.532 18.64 2.38 7.92 8.34 3659.06 43 shallow 0.533 18.65 2.38 7.92 8.34 3657.85 44 shallow 0.49 18.69 2.68 7.92 8.31 4119 45 shallow 0.488 18.7 2.69 7.92 8.31 4144.63 46 shallow 0.486 18.7 2.71 7.92 8.31 4170.27 47 shallow 0.516 18.74 2.43 7.91 8.27 3738.22 48 shallow 0.516 18.74 2.42 7.91 8.27 3723.76 49 shallow 0.517 18.75 2.41 7.91 8.27 3709.3 50 shallow 0.563 18.78 2.8 7.91 8.24 4312.73 51 shallow 0.566 18.79 2.82 7.91 8.24 4336.27 52 shallow 0.569 18.79 2.83 7.91 8.24 4359.8 53 shallow 0.522 18.86 2.86 7.93 8.21 4407.27 54 shallow 0.52 18.87 2.87 7.93 8.21 4422.75 55 shallow 0.518 18.87 2.88 7.93 8.21 4438.23 56 shallow 0.516 18.87 2.89 7.93 8.21 4453.71 57 shallow 0.535 18.82 2.96 7.92 8.24 4550.99 58 shallow 0.535 18.82 2.96 7.92 8.24 4555.91 59 shallow 0.534 18.82 2.96 7.92 8.24 4560.83 60 shallow 0.572 18.96 2.32 7.92 8.22 3570.09 61 shallow 0.574 18.97 2.29 7.92 8.22 3517.96 62 shallow 0.577 18.97 2.25 7.92 8.22 3465.82 63 shallow 0.554 19.02 2.28 7.92 8.26 3510.2 64 shallow 0.553 19.03 2.27 7.92 8.26 3493.2 65 shallow 0.553 19.03 2.26 7.92 8.26 3476.19 66 shallow 0.577 19.06 1.8 7.9 8.29 2773.32 67 shallow 0.578 19.06 1.78 7.9 8.3 2738.19 68 shallow 0.579 19.07 1.76 7.9 8.3 2703.05 69 shallow 0.579 19.07 1.73 7.9 8.3 2667.92 70 shallow 0.538 19.13 2.38 7.88 8.29 3666.19 71 shallow 0.536 19.13 2.41 7.88 8.29 3705.42 72 shallow 0.535 19.13 2.43 7.88 8.29 3744.64 73 shallow 0.586 19.19 2.63 7.88 8.25 4041.07 74 shallow 0.588 19.19 2.65 7.88 8.25 4078.32 75 shallow 0.59 19.19 2.68 7.88 8.25 4115.56 76 shallow 0.565 19.23 2.48 7.87 8.22 3819.99 77 shallow 0.565 19.23 2.47 7.87 8.21 3806.91 78 shallow 0.565 19.23 2.47 7.87 8.21 3793.82 79 shallow 0.577 19.25 2.74 7.88 8.2 4218.14 80 shallow 0.577 19.25 2.75 7.88 8.2 4234.77 81 shallow 0.577 19.25 2.76 7.88 8.2 4251.4 82 shallow 0.577 19.25 2.77 7.88 8.2 4268.03 83 shallow 0.58 19.27 3.44 7.88 8.19 5288.08 84 shallow 0.58 19.27 3.48 7.88 8.19 5352.31 85 shallow 0.581 19.27 3.52 7.88 8.19 5416.53 86 shallow 0.581 19.22 4.24 7.87 8.19 6517.24 87 shallow 0.581 19.22 4.29 7.87 8.19 6596.06 88 shallow 0.581 19.21 4.34 7.87 8.19 6674.87 89 shallow 0.54 19.26 4.82 7.88 8.17 7421.89 90 shallow 0.538 19.26 4.86 7.88 8.16 7481.36 91 shallow 0.535 19.26 4.9 7.88 8.16 7540.84 92 shallow 0.519 19.54 4.95 7.87 8.07 7613.5 93 shallow 0.517 19.56 4.96 7.87 8.07 7627.92 94 shallow 0.515 19.57 4.97 7.87 8.06 7642.34 95 shallow 0.514 19.59 4.98 7.87 8.06 7656.77 96 shallow 0.554 19.7 4.98 7.86 8.02 7661.11 97 shallow 0.556 19.71 4.98 7.86 8.01 7660.43 98 shallow 0.558 19.72 4.98 7.86 8.01 7659.74 99 shallow 0.571 19.77 4.99 7.88 7.97 7680.29 100 shallow 0.573 19.77 4.99 7.88 7.97 7681.31 101 shallow 0.574 19.78 4.99 7.88 7.96 7682.32 102 shallow 0.578 19.8 4.97 7.86 7.94 7650.34 103 shallow 0.579 19.81 4.97 7.86 7.94 7649.01 104 shallow 0.579 19.81 4.97 7.86 7.94 7647.67 105 shallow 0.585 19.84 4.98 7.86 7.86 7658.82 106 shallow 0.586 19.84 4.98 7.86 7.86 7658.77 107 shallow 0.586 19.84 4.98 7.86 7.85 7658.71 108 shallow 0.587 19.84 4.98 7.86 7.85 7658.65 109 shallow 0.571 19.83 4.98 7.88 7.87 7657.5 110 shallow 0.57 19.83 4.98 7.88 7.87 7657.75 111 shallow 0.57 19.83 4.98 7.88 7.87 7657.99 112 shallow 0.589 19.81 4.97 7.88 7.93 7648.08 113 shallow 0.59 19.81 4.97 7.88 7.93 7647.5 114 shallow 0.59 19.8 4.97 7.88 7.94 7646.92 115 shallow 0.55 19.84 4.97 7.86 7.94 7647.9 116 shallow 0.548 19.85 4.97 7.86 7.94 7647.75 117 shallow 0.547 19.85 4.97 7.86 7.95 7647.61 118 shallow 0.58 19.89 4.97 7.87 7.89 7643.28 119 shallow 0.581 19.89 4.97 7.87 7.89 7643.09 120 shallow 0.582 19.89 4.97 7.87 7.89 7642.9 121 shallow 0.583 19.9 4.97 7.87 7.89 7642.71 122 shallow 0.594 19.87 4.97 7.87 7.89 7642.89 123 shallow 0.595 19.87 4.97 7.87 7.89 7642.82 124 shallow 0.596 19.87 4.97 7.87 7.89 7642.74 125 shallow 0.517 19.92 4.97 7.87 7.95 7639.41 126 shallow 0.513 19.93 4.97 7.87 7.96 7639.25 127 shallow 0.509 19.93 4.97 7.87 7.96 7639.08 128 shallow 0.502 19.85 4.96 7.85 8.08 7632.75 129 shallow 0.5 19.85 4.96 7.85 8.09 7632.33 130 shallow 0.498 19.85 4.96 7.85 8.1 7631.92 131 shallow 0.502 19.94 4.96 7.85 8.09 7636.98 132 shallow 0.502 19.95 4.96 7.85 8.09 7637.16 133 shallow 0.503 19.95 4.96 7.85 8.09 7637.33 134 shallow 0.503 19.96 4.96 7.85 8.09 7637.5 135 shallow 0.559 19.94 4.95 7.83 8.1 7608.89 136 shallow 0.562 19.95 4.94 7.82 8.1 7607.48 137 shallow 0.565 19.95 4.94 7.82 8.1 7606.07 138 shallow 0.587 20.04 4.94 7.8 8 7601.13 139 shallow 0.59 20.04 4.94 7.8 8 7600.29 140 shallow 0.592 20.04 4.94 7.8 7.99 7599.45 141 shallow 0.576 20.03 4.94 7.79 7.93 7592.73 142 shallow 0.575 20.03 4.94 7.79 7.92 7592.35 143 shallow 0.574 20.03 4.93 7.79 7.91 7591.96 144 shallow 0.588 20.02 4.93 7.78 7.86 7586.62 145 shallow 0.588 20.02 4.93 7.78 7.86 7586.22 146 shallow 0.588 20.01 4.93 7.78 7.86 7585.82 147 shallow 0.589 20.01 4.93 7.78 7.85 7585.42 148 shallow 0.629 20.01 4.93 7.77 7.77 7578.26 149 shallow 0.631 20.01 4.93 7.77 7.77 7577.78 150 deep 0.634 20.01 4.93 7.77 7.76 7577.3 151 deep 0.704 20.02 4.93 7.77 7.69 7576.95 152 deep 0.708 20.02 4.92 7.77 7.68 7576.81 153 deep 0.713 20.02 4.92 7.77 7.67 7576.66 154 deep 0.774 20 4.92 7.77 7.62 7576.71 155 deep 0.779 19.99 4.92 7.77 7.62 7576.73 156 deep 0.784 19.99 4.92 7.77 7.62 7576.75 157 deep 0.856 19.99 4.92 7.77 7.57 7572.97 158 deep 0.861 19.99 4.92 7.77 7.57 7572.77 159 deep 0.866 19.99 4.92 7.77 7.56 7572.56 160 deep 0.871 19.98 4.92 7.77 7.56 7572.35 161 deep 0.977 19.98 4.92 7.77 7.52 7569.69 162 deep 0.984 19.98 4.92 7.77 7.52 7569.47 163 deep 0.991 19.98 4.92 7.77 7.52 7569.25 164 deep 1.028 19.99 4.92 7.77 7.5 7567.25 165 deep 1.032 19.99 4.92 7.77 7.5 7567.09 166 deep 1.035 19.99 4.92 7.77 7.5 7566.94 167 deep 1.154 19.98 4.92 7.78 7.49 7572.95 168 deep 1.161 19.98 4.92 7.78 7.49 7573.25 169 deep 1.167 19.98 4.92 7.78 7.49 7573.55 170 deep 1.123 19.98 4.92 7.77 7.49 7572.76 171 deep 1.123 19.98 4.92 7.77 7.49 7572.84 172 deep 1.123 19.98 4.92 7.77 7.48 7572.92 173 deep 1.122 19.98 4.92 7.77 7.48 7572.99 174 deep 1.268 19.98 4.93 7.77 7.46 7583.29 175 deep 1.275 19.98 4.93 7.77 7.46 7583.81 176 deep 1.281 19.98 4.93 7.77 7.45 7584.33 177 deep 1.225 19.97 4.93 7.77 7.46 7581.18 178 deep 0.745 20.23 4.91 7.72 7.04 7546.57 179 deep 0.773 20.24 4.9 7.72 7.03 7545.79 180 deep 0.802 20.24 4.9 7.72 7.02 7545.01 181 deep 0.784 20.14 4.91 7.76 7.03 7551.84 182 deep 0.793 20.13 4.91 7.76 7.03 7552.02 183 deep 0.802 20.12 4.91 7.76 7.03 7552.19 184 deep 0.959 20.08 4.92 7.75 7.06 7570.66 185 deep 0.966 20.08 4.92 7.75 7.06 7571.86 186 deep 0.973 20.07 4.92 7.75 7.06 7573.06 187 deep 0.979 20.07 4.92 7.75 7.07 7574.27 188 deep 0.908 20.07 4.93 7.75 7.07 7578.24 189 deep 0.907 20.07 4.93 7.75 7.07 7578.83 190 deep 0.906 20.07 4.93 7.75 7.07 7579.42 191 deep 0.954 20.05 4.92 7.74 7.09 7576.65 192 deep 0.955 20.05 4.92 7.74 7.09 7576.53 193 deep 0.956 20.05 4.92 7.74 7.09 7576.41 194 deep 1.06 20.08 4.94 7.75 7.1 7594.5 195 deep 1.067 20.08 4.94 7.75 7.1 7595.42 196 deep 1.074 20.08 4.94 7.75 7.1 7596.35 197 deep 1.112 20.1 4.93 7.75 7.1 7590.61 198 deep 1.116 20.1 4.93 7.75 7.1 7590.65 199 deep 1.12 20.1 4.93 7.74 7.1 7590.69 200 deep 1.124 20.1 4.93 7.74 7.1 7590.72 201 deep 1.135 20.11 4.92 7.75 7.11 7572.66 202 deep 1.136 20.11 4.92 7.76 7.11 7571.53 203 deep 1.138 20.11 4.92 7.76 7.11 7570.39 204 deep 1.121 20.08 4.93 7.77 7.17 7583.62 205 deep 1.121 20.08 4.93 7.77 7.17 7584 206 deep 1.12 20.08 4.93 7.77 7.17 7584.38 207 deep 1.143 20.07 4.93 7.78 7.25 7584.39 208 deep 1.144 20.06 4.93 7.78 7.26 7584.7 209 deep 1.145 20.06 4.93 7.78 7.26 7585.02 210 deep 1.146 20.06 4.93 7.78 7.27 7585.33 211 deep 1.088 20.04 4.93 7.79 7.32 7583.78 212 deep 1.086 20.04 4.93 7.8 7.32 7583.68 213 deep 1.083 20.04 4.93 7.8 7.33 7583.58 214 deep 1.108 20.03 4.93 7.79 7.38 7588.5 215 deep 1.108 20.03 4.93 7.79 7.38 7588.74 216 deep 1.109 20.03 4.93 7.79 7.38 7588.97 217 deep 1.075 20.04 4.93 7.77 7.37 7590.77 218 deep 1.074 20.04 4.93 7.77 7.38 7590.97 219 deep 1.073 20.04 4.93 7.77 7.38 7591.17 220 deep 1.095 20.06 4.93 7.77 7.36 7590.4 221 deep 1.095 20.06 4.93 7.77 7.36 7590.37 222 deep 1.096 20.06 4.93 7.76 7.36 7590.35 223 deep 1.096 20.07 4.93 7.76 7.36 7590.33 224 deep 1.11 20.05 4.94 7.76 7.34 7592.75 225 deep 1.111 20.05 4.94 7.76 7.34 7592.85 226 deep 1.113 20.05 4.94 7.76 7.34 7592.96 227 deep 1.008 20.05 4.94 7.76 7.31 7595.79 228 deep 1.002 20.05 4.94 7.76 7.3 7595.99 229 deep 0.997 20.05 4.94 7.76 7.3 7596.19 230 deep 0.919 20.05 4.68 7.76 7.27 7197.13 231 deep 0.912 20.05 4.66 7.76 7.27 7175.5 232 deep 0.906 20.05 4.65 7.76 7.27 7153.87 233 deep 0.928 20.05 4.62 7.76 7.27 7114.05 234 deep 0.928 20.05 4.62 7.76 7.27 7104.11 235 deep 0.928 20.05 4.61 7.76 7.27 7094.18 236 deep 0.928 20.05 4.6 7.76 7.27 7084.24 237 deep 0.99 20.05 4.61 7.83 7.47 7099.62 238 deep 0.994 20.05 4.62 7.84 7.48 7100.44 239 deep 0.998 20.05 4.62 7.84 7.49 7101.27 240 deep 0.833 20.04 4.63 7.84 7.64 7130.42 241 deep 0.825 20.04 4.64 7.84 7.65 7132.41 242 deep 0.817 20.04 4.64 7.84 7.66 7134.39 243 deep 0.909 20.03 4.62 7.78 7.76 7105.67 244 deep 0.911 20.03 4.62 7.77 7.77 7104.64 245 deep 0.912 20.03 4.62 7.77 7.78 7103.6 246 deep 0.987 20.03 4.63 7.79 7.77 7116.45 247 deep 0.993 20.03 4.63 7.79 7.77 7116.51 248 deep 0.999 20.03 4.63 7.79 7.77 7116.57 249 deep 1.006 20.03 4.63 7.79 7.77 7116.64 250 deep 1.04 20.03 4.64 7.78 7.82 7131.62 251 deep 1.043 20.02 4.64 7.78 7.82 7132.73 252 deep 1.046 20.02 4.64 7.78 7.82 7133.84 253 deep 1.044 20.02 4.64 7.78 7.82 7136.56 254 deep 1.044 20.01 4.64 7.78 7.82 7136.98 255 deep 1.044 20.01 4.64 7.78 7.82 7137.4 256 deep 0.87 20.01 4.63 7.77 7.84 7121.68 257 deep 0.86 20.01 4.63 7.77 7.84 7120.84 258 deep 0.85 20.01 4.63 7.77 7.84 7120 259 deep 0.792 20 4.61 7.76 7.82 7094.36 260 deep 0.786 20 4.61 7.76 7.82 7092.63 261 deep 0.779 20 4.61 7.76 7.82 7090.9 262 deep 0.773 20 4.61 7.76 7.82 7089.17 263 deep 0.85 20.01 4.6 7.76 7.77 7078.52 264 deep 0.853 20.01 4.6 7.76 7.77 7077.45 265 deep 0.856 20.01 4.6 7.76 7.77 7076.38 266 deep 0.979 20.01 4.61 7.76 7.76 7099.24 267 deep 0.988 20.01 4.62 7.76 7.76 7100.3 268 deep 0.996 20.01 4.62 7.76 7.75 7101.37 269 deep 0.981 20.01 4.6 7.76 7.73 7081.11 270 deep 0.982 20.01 4.6 7.76 7.73 7080.49 271 deep 0.984 20.01 4.6 7.76 7.73 7079.86 272 deep 0.963 20 4.6 7.76 7.69 7077.37 273 deep 0.961 20 4.6 7.76 7.69 7076.78 274 deep 0.959 20 4.6 7.76 7.69 7076.19 275 deep 0.957 20 4.6 7.76 7.68 7075.59 276 deep 1.056 19.98 4.62 7.75 7.67 7104.38 277 deep 1.061 19.98 4.62 7.75 7.66 7105.89 278 deep 1.066 19.97 4.62 7.75 7.66 7107.39 279 deep 1.029 19.96 4.65 7.76 7.66 7147.6 280 deep 1.029 19.96 4.65 7.76 7.66 7150.33 281 deep 1.029 19.96 4.65 7.76 7.65 7153.06 282 deep 1.057 19.95 4.61 7.76 7.65 7098.87 283 deep 1.057 19.95 4.61 7.76 7.65 7096.65 284 deep 1.058 19.95 4.61 7.76 7.65 7094.42 285 deep 0.932 19.95 4.62 7.75 7.63 7101.76 286 deep 0.926 19.95 4.62 7.75 7.62 7100.99 287 deep 0.92 19.95 4.62 7.75 7.62 7100.22 288 deep 0.913 19.95 4.61 7.75 7.62 7099.45 289 deep 0.83 19.95 4.83 7.75 7.59 7426.1 290 deep 0.823 19.95 4.84 7.75 7.59 7443.64 291 deep 0.816 19.95 4.85 7.75 7.58 7461.18 292 deep 0.431 19.95 4.85 7.76 7.54 7465.86 293 deep 0.409 19.95 4.86 7.76 7.53 7472.57 294 deep 0.387 19.95 4.86 7.76 7.53 7479.27 295 deep 0.519 19.95 4.88 7.75 7.46 7502.56 296 deep 0.519 19.95 4.88 7.75 7.45 7503.16 297 deep 0.871 19.95 4.89 7.76 7.29 7522.43 298 deep 0.879 19.95 4.89 7.76 7.29 7522.81 299 deep 0.879 19.94 4.89 7.76 7.28 7524.61 300 deep 0.879 19.94 4.89 7.76 7.28 7524.72 301 deep 0.879 19.94 4.89 7.76 7.28 7524.82 302 deep 0.776 19.94 4.89 7.77 7.29 7520.93 303 deep 0.768 19.94 4.89 7.77 7.29 7520.66 304 deep 0.761 19.94 4.89 7.77 7.29 7520.39 305 deep 0.754 19.94 4.89 7.77 7.29 7520.13 306 deep 0.907 19.91 4.89 7.77 7.3 7527.86 307 deep 0.913 19.91 4.89 7.77 7.3 7528.18 308 deep 0.919 19.91 4.89 7.76 7.3 7528.49 309 deep 0.791 19.91 4.89 7.77 7.32 7528.08 310 deep 0.787 19.9 4.89 7.77 7.32 7528.22 311 deep 0.783 19.9 4.89 7.77 7.32 7528.36 312 deep 0.805 19.9 4.89 7.76 7.32 7523.26 313 deep 0.803 19.9 4.89 7.76 7.32 7522.95 314 deep 0.802 19.9 4.89 7.76 7.32 7522.63 315 deep 0.829 19.89 4.92 7.76 7.31 7563.5 316 deep 0.831 19.89 4.92 7.76 7.3 7565.7 317 deep 0.834 19.88 4.92 7.76 7.3 7567.89 318 deep 0.836 19.88 4.92 7.76 7.3 7570.08 319 deep 0.85 19.86 4.92 7.76 7.29 7569.92 320 deep 0.851 19.86 4.92 7.76 7.29 7570.78 321 deep 0.853 19.86 4.92 7.76 7.29 7571.64 322 deep 0.859 19.84 4.92 7.77 7.29 7573.14 323 deep 0.86 19.84 4.92 7.77 7.29 7573.11 324 deep 0.861 19.84 4.92 7.77 7.29 7573.08 325 deep 0.726 19.82 4.91 7.77 7.29 7558.29 326 deep 0.718 19.82 4.91 7.77 7.29 7557.48 327 deep 0.711 19.82 4.91 7.77 7.29 7556.68 328 deep 0.858 19.84 4.91 7.76 7.28 7558.94 329 deep 0.863 19.84 4.91 7.76 7.28 7558.79 330 deep 0.869 19.85 4.91 7.76 7.28 7558.63 331 deep 0.874 19.85 4.91 7.76 7.28 7558.48 332 deep 1.114 19.86 4.91 7.75 7.26 7556.52 333 deep 1.13 19.86 4.91 7.75 7.26 7556.5 334 deep 1.146 19.86 4.91 7.75 7.26 7556.49 335 deep 1.182 19.87 4.92 7.76 7.25 7563.37 336 deep 1.188 19.88 4.92 7.76 7.25 7563.71 337 deep 1.194 19.88 4.92 7.76 7.25 7564.04 338 deep 1.184 19.87 4.91 7.76 7.27 7561.04 339 deep 1.183 19.87 4.91 7.76 7.27 7561.01 340 deep 1.182 19.87 4.91 7.76 7.27 7560.98 341 deep 0.987 19.87 4.91 7.76 7.26 7559.3 342 deep 0.976 19.87 4.91 7.76 7.26 7559.13 343 deep 0.964 19.87 4.91 7.76 7.26 7558.95 344 deep 0.953 19.87 4.91 7.76 7.26 7558.78 345 deep 0.988 19.86 4.91 7.76 7.27 7557.15 appendix 2. water trace element data (mg/l). site subsite li na k rb cs be mg ca sr ba al sc mn fe cu ni zn as bl marsh 3.81 3897 274 1.25 0.102 0.0002 122.1 278 5.5 0.11 0.02 <0.0002 0 0.042 <0.5 0 0.02 <0.04 bl pond 1.71 1501 114 0.56 0.046 <0.000006 52.1 109 2.5 0.06 0.02 0.0002 0 <0.008 <0.5 0 0.01 <0.04 hrs deep 0.59 1496 57 0.05 0.004 <0.00005 53.9 118 1.1 0.07 <0.005 <0.0004 <0.003 <0.04 <0.0007 <0.0002 0.01 <0.04 hrs surface 0.59 1498 58 0.05 0.004 <0.00005 53.8 117 1 0.07 0.01 <0.0004 <0.003 <0.04 0 <0.0002 0.02 <0.04 appendix 3. bulk shell stable isotope data. sample name site sample type genera sub-site δ 13 c (‰-vpdb) δ 18 o (‰-vsmow) data source lab bl-m-t6-1 bl shell succineidae bl-marsh -3.4 -10.017 this study sirfer bl-m-t6-2 bl shell succineidae bl-marsh -3.33 -10.017 this study sirfer bl-m-t6-6 bl shell succineidae bl-marsh -4.23 -10.546 this study sirfer bl-m-t6-3 bl shell succineidae bl-marsh -3.07 -10.703 this study sirfer bl-m-t6-4 bl shell succineidae bl-marsh -5.18 -10.812 this study sirfer bl-m-t6-5 bl shell succineidae bl-marsh -6.49 -11.096 this study sirfer bl-p-t5-8 bl shell melanoides bl-pond -1.84 -12.646 this study sirfer bl-m-t2-1 bl shell pyrgulopsis bl-marsh -2.34 -13.182 this study sirfer bl-m-t2-3 bl shell pyrgulopsis bl-marsh -2.52 -13.894 this study sirfer bl-m-t2-4 bl shell pyrgulopsis bl-marsh -2.14 -14.243 this study sirfer hrs-01-pp-o1 hrs shell pyrgulopsis hrs-subsite-01 -5.24 -14.343 this study sirfer hrs-03-pp-o1 hrs shell pyrgulopsis hrs-subsite-03 -4.85 -14.361 this study sirfer hrs-03-pp-o2 hrs shell pyrgulopsis hrs-subsite-03 -3.95 -14.406 this study sirfer hrs-02-tp-o2 hrs shell tryonia hrs-subsite-02 -4.39 -14.568 this study sirfer hrs-04-tp-o1 hrs shell tryonia hrs-subsite-04 -4.42 -14.578 this study sirfer bl-m-t2-2 bl shell pyrgulopsis bl-marsh -2.09 -14.634 this study sirfer hrs-04-tp-o2 hrs shell tryonia hrs-subsite-04 -4.48 -14.733 this study sirfer hrs-01-tp-o2 hrs shell tryonia hrs-subsite-01 -4.76 -14.776 this study sirfer hrs-02-pp-o1 hrs shell pyrgulopsis hrs-subsite-02 -4.67 -14.795 this study sirfer hrs-01-pp-o2 hrs shell pyrgulopsis hrs-subsite-01 -5.23 -14.83 this study sirfer hrs-02-pp-o2 hrs shell pyrgulopsis hrs-subsite-02 -4.86 -14.851 this study sirfer hrs-02-tp-o1 hrs shell tryonia hrs-subsite-02 -4.97 -14.858 this study sirfer hrs-04-pp-o1 hrs shell pyrgulopsis hrs-subsite-04 -5 -14.879 this study sirfer hrs-03-tp-o1 hrs shell tryonia hrs-subsite-03 -4.82 -14.88 this study sirfer hrs-04-pp-o2 hrs shell pyrgulopsis hrs-subsite-04 -5.18 -14.959 this study sirfer hrs-03-tp-o2 hrs shell tryonia hrs-subsite-03 -4.2 -15.059 this study sirfer bl-p-t4-2 bl shell planorbella bl-pond -4.82 -15.295 this study sirfer hrs-01-tp-o1 hrs shell tryonia hrs-subsite-01 -5.01 -15.349 this study sirfer hrs-t2-4 hrs shell pyrgulopsis hrs-subsite-00 -5.14 -15.565 this study sirfer hrs-t4-1 hrs shell planorbella hrs-subsite-00 -8.96 -15.67 this study sirfer hrs-t2-1 hrs shell pyrgulopsis hrs-subsite-00 -5.19 -15.718 this study sirfer hrs-t2-5 hrs shell pyrgulopsis hrs-subsite-00 -5.19 -15.778 this study sirfer hrs-t2-2 hrs shell pyrgulopsis hrs-subsite-00 -5.08 -15.788 this study sirfer hrs-t2-3 hrs shell pyrgulopsis hrs-subsite-00 -5.03 -15.883 this study sirfer hrs-t1-3 hrs shell tryonia hrs-subsite-00 -4.95 -15.98 this study sirfer bl-p-t2-2 bl shell pyrgulopsis bl-pond -1.92 -16.045 this study sirfer hrs-t1-4 hrs shell tryonia hrs-subsite-00 -4.98 -16.068 this study sirfer hrs-t1-1 hrs shell tryonia hrs-subsite-00 -4.57 -16.085 this study sirfer hrs-t1-5 hrs shell tryonia hrs-subsite-00 -5.02 -16.131 this study sirfer hrs-t1-2 hrs shell tryonia hrs-subsite-00 -4.96 -16.158 this study sirfer bl-p-t2-1 bl shell pyrgulopsis bl-pond -1.72 -16.282 this study sirfer bl-p-t3-2 bl shell physella bl-pond -5.02 -16.341 this study sirfer bl-s-t1-1 bl shell tryonia bl-spring -1.92 -16.374 this study sirfer bl-p-t1-1 bl shell tryonia bl-pond -1.71 -16.524 this study sirfer bl-p-t4-5 bl shell planorbella bl-pond -4.5 -16.641 this study sirfer bl-p-t3-1 bl shell physella bl-pond -4.82 -16.819 this study sirfer bl-p-t4-1 bl shell planorbella bl-pond -3.65 -16.823 this study sirfer bl-p-t5-6 bl shell melanoides bl-pond -2.24 -16.858 this study sirfer bl-l-t5-4 bl shell melanoides bl-lake -2.93 -16.893 this study sirfer bl-p-t4-4 bl shell planorbella bl-pond -3.71 -16.916 this study sirfer bl-p-t5-3 bl shell melanoides bl-pond -2.48 -17.015 this study sirfer bl-p-t5-4 bl shell melanoides bl-pond -2.15 -17.058 this study sirfer bl-p-t5-1 bl shell melanoides bl-pond -2.44 -17.084 this study sirfer bl-p-t5-7 bl shell melanoides bl-pond -2.49 -17.125 this study sirfer bl-p-t5-2 bl shell melanoides bl-pond -2.46 -17.135 this study sirfer bl-l-t5-3 bl shell melanoides bl-lake -2.99 -17.162 this study sirfer bl-l-t5-1 bl shell melanoides bl-lake -3.02 -17.238 this study sirfer bl-l-t5-2 bl shell melanoides bl-lake -2.38 -17.238 this study sirfer bl-s-t5-4 bl shell melanoides bl-spring -2.23 -17.343 this study sirfer bl-s-t5-1 bl shell melanoides bl-spring -2.1 -17.516 this study sirfer bl-s-t5-5 bl shell melanoides bl-spring -2.29 -17.519 this study sirfer bl-s-t1-2 bl shell tryonia bl-spring -1.91 -17.541 this study sirfer bl-s-t5-2 bl shell melanoides bl-spring -2.13 -17.549 this study sirfer bl-p-t4-3 bl shell planorbella bl-pond -3.08 this study sirfer bl-spring-shell-2019-1 bl shell melanoides bl-spring -2.21 lerback and others, 2023 nosams bl-spring-shell-2019-2 bl shell melanoides bl-spring -2.26 lerback and others, 2023 nosams bl-spring-shell-2019-3 bl shell melanoides bl-spring lerback and others, 2023 nosams bl-spring-shell-2019-4 bl shell melanoides bl-spring -3.08 lerback and others, 2023 nosams bl-pond-shell-2019-1 bl shell melanoides bl-pond -2.17 lerback and others, 2023 nosams bl-pond-shell-2019-2 bl shell melanoides bl-pond -2.09 lerback and others, 2023 nosams bl-pond-shell-2019-3 bl shell melanoides bl-pond -2.37 lerback and others, 2023 nosams bl-pond-shell-2019-4 bl shell melanoides bl-pond -2.14 lerback and others, 2023 nosams bl-lake-shell-2018-1 bl shell melanoides bl-lake -3.06 lerback and others, 2023 nosams bl-lake-shell-2018-2 bl shell melanoides bl-lake -2.85 lerback and others, 2023 nosams bl-lake-shell-2018-3 bl shell melanoides bl-lake -2.98 lerback and others, 2023 nosams bl-lake-shell-2018-4 bl shell melanoides bl-lake -3.89 lerback and others, 2023 nosams hrs-pond-shell-2020-1 hrs shell pyrgulopsis hrs-subsite-00 -5.18 lerback and others, 2023 nosams hrs-pond-shell-2020-2 hrs shell pyrgulopsis hrs-subsite-00 -4.43 lerback and others, 2023 nosams hrs-pond-shell-2020-3 hrs shell pyrgulopsis hrs-subsite-00 -4.77 lerback and others, 2023 nosams hrs-pond-shell-2020-4 hrs shell pyrgulopsis hrs-subsite-00 -4.79 lerback and others, 2023 nosams bl-spring-shell-2020-5 bl shell pyrgulopsis bl-spring -1.82 lerback and others, 2023 nosams bl-spring-shell-2020-6 bl shell pyrgulopsis bl-spring -1.97 lerback and others, 2023 nosams bl-spring-shell-2020-7 bl shell pyrgulopsis bl-spring -1.94 lerback and others, 2023 nosams bl-spring-shell-2020-8 bl shell pyrgulopsis bl-spring -1.83 lerback and others, 2023 nosams bl-pond-sediment-2020-1 bl sediment bl-pond -26.61 lerback and others, 2023 nosams hrs-pond-sediment-2020-1 hrs sediment hrs-subsite-00 -20.65 lerback and others, 2023 nosams appendix 4 . water stable isotope data. site type δ18o (‰-vsmow) δ18o (‰-vpdb converted) δ13c (‰-vpdb) δ13c source lab source subsite bl water -3.73 measured nosams lerback and others, 2023 bl-lake bl water -4 measured nosams lerback and others, 2023 bl-pond bl water -4.5 measured nosams lerback and others, 2019 supplementary data. "lookout point, 6/2/2017" bl-spring bl water -4.65 measured nosams lerback and others, 2023 bl-spring hrs water -7.02 measured nosams lerback and others, 2023 hrs-subsite-00 bl water -15.9 -45.42 sirfer lerback and others, 2019 bl-spring bl water -16.05 -45.56 sirfer lerback and others, 2019 bl-spring bl water -15.92 -45.43 sirfer lerback and others, 2019 bl-spring bl water -15.99 -45.5 sirfer lerback and others, 2019 bl-spring bl water -16 -45.51 sirfer lerback and others, 2019 bl-spring bl water -16.1 -45.61 sirfer lerback and others, 2019 bl-spring bl water -15.86 -45.37 sirfer lerback and others, 2019 bl-lake bl water -15.74 -45.26 sirfer lerback and others, 2019 bl-lake hrs water -16.23 -45.73 sirfer lerback and others, 2023 hrs-subsite-00 hrs water -16.06 -45.57 sirfer lerback and others, 2023 hrs-subsite-00 hrs water -15.98 -45.49 sirfer lerback and others, 2023 hrs-subsite-00 hrs water -16.04 -45.55 sirfer lerback and others, 2023 hrs-subsite-00 hrs water -16 -45.51 sirfer lerback and others, 2023 hrs-subsite-00 hrs water -15.89 -45.41 sirfer lerback and others, 2023 hrs-subsite-00 hrs water -16.01 -45.52 sirfer lerback and others, 2023 hrs-subsite-00 appendix 5. gastropod intrashell δ 13 c and δ 18 o transect data. sample name drill site number δ 13 c δ 18 o bl-s-t5-9-4 4 -2.91 -17.46 bl-s-t5-9-3 3 -2.41 -17.48 bl-s-t5-9-2 2 -2.21 -17.32 bl-s-t5-9-1 1 -2.24 -17.49 bl-s-t5-8-4 4 -2.64 -17.25 bl-s-t5-8-3 3 -2.54 -17.33 bl-s-t5-8-2 2 -2.13 -17.37 bl-s-t5-8-1 1 -1.91 -17.51 bl-s-t5-7-4 4 -2.54 -17.67 bl-s-t5-7-3 3 -2.08 -17.33 bl-s-t5-7-2 2 -2.24 -17.65 bl-s-t5-7-1 1 -2.25 -17.26 bl-s-t5-6-4 4 -2.97 -16.57 bl-s-t5-6-3 3 -2.84 -16.96 bl-s-t5-6-2 2 -2.41 -17.18 bl-s-t5-6-1 1 -2.2 -17.36 appendix 6. calculated formation temperatures of gastropod shells sample name genera sub-site t (℃; kim and o'neill, 1997) t (℃; kim and others, 2007) t (℃; white and others, 1999) bl-m-t6-1 succineidae bl-marsh -11.4 -8.6 -6.8 bl-m-t6-2 succineidae bl-marsh -11.4 -8.6 -6.8 bl-m-t6-6 succineidae bl-marsh -9.3 -6.5 -4.5 bl-m-t6-3 succineidae bl-marsh -8.7 -5.9 -3.8 bl-m-t6-4 succineidae bl-marsh -8.3 -5.4 -3.3 bl-m-t6-5 succineidae bl-marsh -7.2 -4.3 -2.1 bl-p-t5-8 melanoides bl-pond -0.8 2.2 5 bl-m-t2-1 pyrgulopsis bl-marsh 1.4 4.5 7.5 bl-m-t2-3 pyrgulopsis bl-marsh 4.5 7.7 11 bl-m-t2-4 pyrgulopsis bl-marsh 6 9.2 12.7 bl-m-t2-2 pyrgulopsis bl-marsh 7.7 11 14.6 bl-p-t4-2 planorbella bl-pond 10.7 14.1 18 bl-p-t2-2 pyrgulopsis bl-pond 14.1 17.7 21.9 bl-p-t2-1 pyrgulopsis bl-pond 15.2 18.8 23.2 bl-p-t3-2 physella bl-pond 15.5 19.1 23.5 bl-s-t1-1 tryonia bl-spring 15.7 19.2 23.7 bl-p-t1-1 tryonia bl-pond 16.4 20 24.5 bl-p-t4-5 planorbella bl-pond 16.9 20.5 25.1 bl-p-t3-1 physella bl-pond 17.8 21.4 26.1 bl-p-t4-1 planorbella bl-pond 17.8 21.4 26.1 bl-p-t5-6 melanoides bl-pond 17.9 21.6 26.3 bl-l-t5-4 melanoides bl-lake 18.1 21.8 26.5 bl-p-t4-4 planorbella bl-pond 18.2 21.9 26.6 bl-p-t5-3 melanoides bl-pond 18.7 22.4 27.1 bl-p-t5-4 melanoides bl-pond 18.9 22.6 27.4 bl-p-t5-1 melanoides bl-pond 19 22.7 27.5 bl-p-t5-7 melanoides bl-pond 19.2 22.9 27.7 bl-p-t5-2 melanoides bl-pond 19.3 23 27.8 bl-l-t5-3 melanoides bl-lake 19.4 23.1 27.9 bl-l-t5-1 melanoides bl-lake 19.8 23.5 28.4 bl-l-t5-2 melanoides bl-lake 19.8 23.5 28.4 bl-s-t5-4 melanoides bl-spring 20.3 24 28.9 bl-s-t5-1 melanoides bl-spring 21.1 24.9 29.9 bl-s-t5-5 melanoides bl-spring 21.1 24.9 29.9 bl-s-t1-2 tryonia bl-spring 21.2 25 30 bl-s-t5-2 melanoides bl-spring 21.3 25.1 30.1 hrs-01-pp-o1 pyrgulopsis hrs-subsite-01 6.3 9.5 13 hrs-03-pp-o1 pyrgulopsis hrs-subsite-03 6.3 9.6 13.1 hrs-03-pp-o2 pyrgulopsis hrs-subsite-03 6.5 9.8 13.3 hrs-02-tp-o2 tryonia hrs-subsite-02 7.3 10.6 14.1 hrs-04-tp-o1 tryonia hrs-subsite-04 7.3 10.6 14.2 hrs-04-tp-o2 tryonia hrs-subsite-04 8 11.3 15 hrs-01-tp-o2 tryonia hrs-subsite-01 8.2 11.5 15.2 hrs-02-pp-o1 pyrgulopsis hrs-subsite-02 8.3 11.6 15.3 hrs-01-pp-o2 pyrgulopsis hrs-subsite-01 8.4 11.8 15.4 hrs-02-pp-o2 pyrgulopsis hrs-subsite-02 8.5 11.9 15.6 hrs-02-tp-o1 tryonia hrs-subsite-02 8.5 11.9 15.6 hrs-04-pp-o1 pyrgulopsis hrs-subsite-04 8.6 12 15.7 hrs-03-tp-o1 tryonia hrs-subsite-03 8.6 12 15.7 hrs-04-pp-o2 pyrgulopsis hrs-subsite-04 9 12.4 16.1 hrs-03-tp-o2 tryonia hrs-subsite-03 9.4 12.8 16.6 hrs-01-tp-o1 tryonia hrs-subsite-01 10.8 14.2 18.1 hrs-t2-4 pyrgulopsis hrs-subsite-00 11.7 15.2 19.2 hrs-t4-1 planorbella hrs-subsite-00 12.2 15.7 19.8 hrs-t2-1 pyrgulopsis hrs-subsite-00 12.4 15.9 20 hrs-t2-5 pyrgulopsis hrs-subsite-00 12.7 16.2 20.3 hrs-t2-2 pyrgulopsis hrs-subsite-00 12.8 16.3 20.4 hrs-t2-3 pyrgulopsis hrs-subsite-00 13.2 16.7 20.9 hrs-t1-3 tryonia hrs-subsite-00 13.6 17.2 21.4 hrs-t1-4 tryonia hrs-subsite-00 14.1 17.6 21.8 hrs-t1-1 tryonia hrs-subsite-00 14.1 17.7 21.9 hrs-t1-5 tryonia hrs-subsite-00 14.4 17.9 22.2 hrs-t1-2 tryonia hrs-subsite-00 14.5 18 22.3 appendix 7. gastropod shell trace element data (mg/kg) site taxa li na k rb cs be mg ca sr ba al sc mn fe cu ni zn as bl succineidae 3.3 2573 370 0.12 0.008 0.0392 56.7 407754 2486.8 40.07 4.86 0.0268 12.19 2.611 5.4 0.21 2.44 0.11 bl succineidae 5.93 2268 456 0.19 0.037 0.195 130.5 425888 2666.1 55.38 13.16 0.032 9.67 8.333 3.7 0.22 3.91 0.36 bl succineidae 2.86 2464 561 0.2 0.021 0.0516 82.2 430344 2473.6 42.62 11.31 0.0398 12.02 10.424 3.4 0.15 2.11 0.42 bl succineidae 4.14 2340 427 0.13 0.007 0.064 40.6 401159 2463.8 36.12 3.7 0.0271 12.63 5.667 11.2 0.35 3.55 0.26 hrs pyrgulopsis 1.6 2170 21 0.03 0.017 0.007 80.6 398182 1263.9 99.39 5.95 0.09 3.12 100.026 <11 1.9 2.09 0.26 hrs pyrgulopsis 2 2667 22 <0.03 0.007 0.0105 51.1 415211 1200.8 106.43 7.12 0.0949 6.75 14.235 <11 2.2 1.05 0.11 hrs pyrgulopsis 1.47 1960 18 0.03 0.015 0.0077 63.9 347402 1131.3 83.02 12.29 0.0812 2.51 67.234 <9 1.67 0.74 0.26 hrs tryonia 1.44 2879 45 0.25 0.093 0.0134 230.9 425229 1317.5 129.36 81.88 0.1146 8.45 64.693 <20 2.25 2.11 0.14 1 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2023 utah geological association publication 51 abstract in the northeastern great basin, usa, thirteen new optically stimulated luminescence (osl) ages and one infrared stimulated luminescence (irsl) age show that two deep pluvial lakes preceded the bonneville lake cycle in cache valley during marine oxygen-isotope stages (mis) 6 (123-191 ka) and 4 (56-71 ka), respectively. our new data define quantitative hydrographs of the little valley and cutler dam lake cycles in both cache valley and the main bonneville basin. in western cache valley, excavation of a faulted, east-plunging spit has sequentially exposed these deposits and overlying mis 3 fielding humid-over-arid double geosols that end westward at a strand of the east-dipping dayton-oxford normal-fault zone. lithologically identical double paleosols in eastern cache valley overlie a variety of deposits, including dated little valley lake beds, and persist above the bonneville shoreline. six new ages show that the little valley lake cycle in cache valley began before 169 ka and ended after 143 ka, and its highest shoreline was above 1493 m. the >25 kyr duration of this pluvial lake cycle rivals the combined durations of the two subsequent lake cycles, during mis 4 and mis 2. the cutler dam lake rose at least to ~1450 m by ~67 ka in cache valley. in the type area in the main bonneville basin, west of cutler narrows, four averaged irsl dates from cutler dam lake beds show that the lake level there had dropped to ~1340 m by ~59 ka. the little valley lake rose at least 40 to 50 m above the local provo shoreline whereas the cutler dam lake missed reaching the provo shoreline by ~13 m. beneath central cache valley, southeast of the study area, there are two laterally extensive, confining layers of silty clay with an intervening sandy gravel layer, all overlying thick gravelly sediment. both confining layers enclose additional thin and discontinuous gravel layers with adjacent oxidized clays. these alternating coarse and fine sediments are probably correlative with the exposed mis 6 to mis 1 deposits and, possibly, older lake cycles. implications and hydrographs for two pre-bonneville pluvial lakes and double geosols from 14 osl-irsl ages in cache valley, ne bonneville basin robert q. oaks, jr., susanne u. jänecke, tammy m. rittenour, thad l. erickson, and michelle s. nelson geosciences department, utah state university, logan, utah, boboaks@comcast.net introduction cache valley is a narrow, elongate north-trending graben that straddles the utah-idaho border (williams, 1948, 1958, 1962; evans and oaks, 1996; janecke and evans, 1999; oaks, 2000; janecke and others, 2003; carney and janecke, 2005). it is separated from the main bonneville basin by a bedrockcored horst upthrown between the wasatch (west) and west cache (east) fault zones. cutler narrows connects the two basins (figure 1b). the bear river fully entered cache valley through oneida narrows (figure 1a) ~45 to 55 ka (pederson and others, 2016) due to diversion by volcanic eruptions in gem valley in se idaho (bright, 1963, 1967; link and others, 1999; janecke and oaks, 2014; utley, 2017). previous work pre-bonneville lakes in cache valley the last three lake cycles of the eastern great basin coincide with even-numbered marine-isotope stages (mis) (lisiecki and raymo, 2005). these are the little valley (~123 to 191 ka, mis 6), cutler dam (~ 56 to 71 ka, mis 4), and bonneville (~14 to 29 ka, mis 2) lake cycles (scott, 1988; scott and others, 1982, 1983; mccoy, 1981, 1987; oviatt and mccoy, 1988, 1992; oviatt and others, 1987, 1992; kaufman and others, 2001; hart and others, 2004). welldeveloped interglacial paleosols separate some but not all of the lake beds. a dated and formally defined paleosol is called a geosol. 10.31711/ugap.v51i.142 2 figure 1. a) major features of the greater cache valley region, n-central utah and se idaho. green box outlines area in figure 1 b. red line ne from college ward, south central cache valley, shows location of figure 6. jh = junction hills; cbd = cache butte divide. b) landscape of cache valley area showing sites of prebonneville deposits dated with aar, osl, and irsl. type area of cutler dam unit is along bear river, sw of cutler narrows. bonneville shoreline is lowest white; provo shoreline is between blue and green shading. white box outlines area in figure 2. r.q. oaks, jr., s.u. jänecke, t.m. rittenour, t.l. erickson, and m.s. nelson implications and hydrographs for 2 pre-bonneville pluvial lakes 3 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 bright (1963, 1967) and mccoy (1981) identified lacustrine gravels below a paleosol beneath gravels of lake bonneville at the ramsbottom gravel pit in idaho, ne cache valley (figure 1b). from that site and nearby smart mountain, idaho, scott and others (1982, 1983) derived amino-acid racemization (aar) data from snail shells in the older lake beds beneath the paleosol that were correlative with aar data from the little valley lake cycle in the main bonneville basin. highest altitudes of pre-bonneville lakes oviatt and others (1987), oviatt and mccoy (1988, 1992), and kaufman and others (2001) concluded that the cutler dam unit, in exposures up to 15 m thick sw of cutler narrows, was deposited in marshy to shallow lacustrine conditions with ostracods indicative of fluctuating brackish conditions. the highest outcrop is at ~1340 m. the highest probable little valley gravels in the main bonneville basin, which were not dated, are about half-way between the local bonneville and provo shorelines (scott and others, 1982, 1983). incision of cutler narrows the bear river flows sw through cutler narrows, the deep and narrow canyon of the bear river across the narrowest part of the cache butte divide (figure 1b). this canyon is cut into hard paleozoic bedrock, is up to 392 m deep, and coincides with the highest bedrock along the cache butte divide (maw, 1968). its bedrock channel is 1.8 km long. nearshore gravels of the cutler dam lake cycle in cache valley are ~110 m higher than somewhat younger marshy deposits in the main bonneville basin. from that, oaks and others (2019, 2020) concluded that most of the bedrock excavation of the lower part of cutler narrows, from an elevation between the highest levels attained by cutler dam and little valley pluvial lakes down to the present level near 1314 m, coincided with eastward flow during the bonneville flood, ~17.4 ka (marrero, 2009). from their analysis of digital-elevation models (dems), nelson (2012) and chen and maloof (2017) proposed that the stansbury oscillation (~26 to 24 ka in oviatt, 2015), may have reached into lower parts of cache valley through cutler narrows, across an area of ~300 km2. if so, cutler narrows was already deeply incised to below the stansbury level before lake bonneville existed, allowing lake bonneville to oscillate as a 5-10 m deep lake in lower cache valley. methods introduction our study emphasizes a staker-parson gravel pit that we call the newton hill pit, in west-central cache valley (figures 1, 2). our emphasis is primarily on pre-bonneville lakes, so the literature on lake bonneville is cited only where pertinent. all altitudes are above mean sea level. those within the newton hill pit are tied to an altitude at a nearby section corner and based on electronic distance meter (edm) and hand-level surveys. altitudes of the original surface there and altitudes elsewhere are based on u.s geological survey 7.5-minute topographic maps, gps readings, caltopo lidar, and google earth pro. we report present altitudes without correction for post-bonneville rebound or tectonics because bonneville rebound is <10-20 m in our study area in cache valley and rebound of pre-bonneville deposits cannot be computed without better pre-bonneville hydrographs. age control we obtained 12 osl (optically stimulated luminescence of quartz) and irsl (infrared stimulated luminescence of feldspar) ages from the newton hill pit, one from the se part of hyde park, utah, and one from muley hill in millville, utah. the latter two are in the east side of cache valley (figures 1a, 1b; table 1). a metal tube was pounded horizontally into the sediment except at muley hill, where matrix sand was collected from gravel beds using double black plastic bags under red light at night. surrounding sediment was obtained to establish both background data and moisture content for each sample. lab analyses at the utah state university osl lab by michelle nelson were done under the supervision of tammy rittenour, with standard procedures outlined in the notes of appendix 1. recalibration and new standards for osl dating changed the osl and irsl dates reported earlier by us (oaks and others, 2014, 2019, 2020). one previous pluvial lake bed dated at ~96 ka (n = 1; the newton hill beds), instead formed during the earlier little valley lake cycle (sample usu-1083; table 1; appendix 1). construction of map and geologic cross sections the evolving exposures of the pit walls were surveyed with a leica model tc600 laser total station in 4 r.q. oaks, jr., s.u. jänecke, t.m. rittenour, t.l. erickson, and m.s. nelson implications and hydrographs for 2 pre-bonneville pluvial lakes figure 2. digital-elevation model of lidar data of the newton hill area. a western strand of the daytonoxford fault zone intersects the pit (do). farther west, several newton fault scarps are left unlabeled to show their clear topographic expression. b = bonneville shoreline, p = provo shoreline. contour interval 20 m. blue is lower, brighter colors higher. m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 table 1. osl & irsl sample information and ages for staker-parson gravel pit (se flank of newton hill), se hyde park, and ne millville, cache county, utah. see appendix 1 for details for these samples. usu‐  sample  number  age in ka  and  method  hand‐level  from edm con‐ trol in feet  edm  al tude  in feet  sample  al tude  in meters  stra ‐  graphic  unit  loca on  1983 nad  comments; ~65 m w correc on  from 1983 gps data to 1927 north  american datum for usgs topo‐ graphic maps in 1960s  date and  collectors  859  15.42  + 1.39 osl n.d. map ~4790  4737  depth ~53  1444  late qlbp  ~ n 41° 52.614'  ~w 111° 57.426'  nw edge of pit; silt & sand beds  dip e; below ~4800' qlbp highest  shore  9‐15‐2010  tr & mn  3243  ne millville,  utah  20.98  + 3.04 osl gps ~5083  map ~5085  google ~5087  n.d. depth 3'  ~1550  early qlbb  n 41° 41.2048'  w 111° 48.2467'  3' below crest of muley hill, mill‐ ville, utah; dissected older delta  between provo and bonneville lake  stands  11‐19‐2019  ro & tc &  tr  1082  21.35  + 3.48 osl ~4665  map ~4775  <4672  depth ~115  ~1422  early qlbb  n 41° 52.5244'  w 111° 57.3198'  center of pit; laminated silty sand  over qlv gravel; 10' above usu‐ 1083  12‐2‐2011  ro & te  854  21.72  + 2.78 osl ~4748  map ~4785  n.d. depth ~37  ~1447  early qlbb  n 41° 52.4478'  w 111° 57.3978'  temporary road near s‐center  edge of pit; silty sand & clay above  qfg geosol, below qlbp gravel  9‐7‐2010  tr & ro  855  39.28  + 3.72 osl ~4739  map ~4810  n.d. depth ~71  ~1444  qfg  n 41° 52.478'  w 111° 57.393'  s‐center of pit; red colluvium:  sandy gravelly mud at top of loess  geosol  9‐7‐2010  tr & ro  1084  53.51  + 6.44 osl n.d. map ~4875  4865  depth ~10  1483  qfg? qcd?  n 41° 52.5045'  w 111° 57.5009'  high w pit margin; white reworked  ash and fine sand in nnw‐sse  channel, under e‐dipping gravel &  soil, over 4° w‐dipping qlv gravel  12‐5‐2011  ro  856  66.82  + 5.94 osl ~4729  map ~4810  n.d. depth ~81  ~1441  qcd  n 41° 52.479'  w 111° 57.388'  s‐center of pit; gravel below qfg  red paleosol base; 9.8' below usu‐ 855  9‐7‐2010  tr & ro  858  67.70  + 6.46 osl ~4709  map ~4790  n.d. depth ~81  ~1435  qcd  n 41° 52.473'  w 111° 57.382'  s‐center of pit; very fine to medi‐ um sand below gravel, ~25   be‐ low qfg geosol base  9‐15‐2010  tr & mn  2895  se hyde  park, utah  142.8  + 13.1 osl n.d. map ~4865  google ~4898  n.d. depth 9.25  ~1493  qlv  n 41° 47.8341'  w 111° 47.8214'  n‐s ver cal wall; fine to coarse  sand within pale green marl below  qfg white caliche geosol below  qlbb lag gravel under fine to very  fine sand with snails  7‐27‐2018  ro  1083  144.3  + 14.5 osl ~4655  map ~4780  <4673  depth ~125  ~1419  qlv  n 41° 52.5243'  w 111° 57.3310'  center of pit; gravel 8.4' below  base of qlbb sand of usu‐1082  12‐2‐2011  ro & te  3202  150.0  + 25.9 osl ~4690'  map ~4885  n.d. depth ~195 ~1430  qlv  n 41° 52.5570'  w 111° 57.4022'  w‐center of pit; pebbly sand 3.0’  below base of qfg red geosol, with  thin qcd between  10‐28‐2019  ro  2490  155.6  + 21.4 irsl ~4735  map ~4840  n.d. depth ~105  ~1443  qlv  n 41° 52.5203'  w 111° 57.4165'  w‐center of pit in wsw cut; sand  and gravel in cobble gravel, 22'  lower than base of overlying chan‐ nel to w  9‐26‐2016  ro & te  857  161.5  + 16.8 osl n.d. gps 4824 map ~4865  n.d. .  depth ~44  ~1470  qlv  n 41° 52.492'  w 111° 57.477'  sw pit in wsw cut; sand & pebble  groundmass in cobble gravel; edm  4821 later at graded site  9‐15‐2010  tr & mn  2491  169.4  + 28.6 osl ~4678   map ~4805  n.d. depth ~127  ~1426  qlv  n 41° 52.5548'  w 111° 57.3882'  nw pit near s end of headwall;  pebbly sand below qcd calcareous  sandy mud intertonguing upward  with sandy pebbly cobble gravel  clinoforms above  9‐26‐2016  ro & te  osl = optically stimulated luminescence on quartz sand; irsl = infrared stimulated luminescence on feldspathic sand; ka = thousands of years ago; google = google earth pro; edm = total station, electronic distance measurements with laser; gps = global-positioning-system measurement; hl = hand level used from edm base station; n.d. = no data; map: original surface altitudes are interpolated from 1964 u.s. geological survey 7.5' newton [c.i. = 5'] and trenton [c.i. = 20'] topographic quadrangles; qlbp = provo highstand lake stage; qlbb = bonneville highstand lake stage; qfg = fielding emergent interval with multistory humid over arid geosols, and perhaps higher n-s channel; qcd = cutler dam lake stage; qlv = little valley lake stage; mn = michelle s. nelson; ro = robert q. oaks, jr.; tc = tomas capaldi; te = thad l. erickson; tr = tammy m. rittenour. note: qcd and early qlbb lakes in cache valley may have been separated at cutler narrows from lower coeval lakes in the main bonneville basin.   5 6 r.q. oaks, jr., s.u. jänecke, t.m. rittenour, t.l. erickson, and m.s. nelson implications and hydrographs for 2 pre-bonneville pluvial lakes 2016. thereafter, new contacts were surveyed with an abney hand level from the edm base station. these data, combined with our 12 osl and irsl ages from the central and western parts of the pit, were used to construct a map and four composite stratigraphic sections across much of the newton hill pit (figures 3, 4). correlations are tied to: (1) continuous and isolated exposures of the fielding double geosols (oviatt and mccoy, 1988) at the top of dated cutler dam lake beds in the central part of the pit, and westward atop dated little valley beds; (2) a thick green marl low within bonneville deposits; and (3) an overlying pink marl. quantitative hydrographs our new data (table 1) and prior aar data (appendices 2, 3, 4) and thermoluminescence (tl) data, tied to altitudes (appendix 5), constrain our quantitative hydrographs (figure 5) of the cutler dam and little valley lake cycles in both cache valley and the main bonneville basin. these hydrographs update schematic plots of scott and others (1982, 1983), mccoy (1987), oviatt and others (1987), and hart and others (2004), for these two prebonneville lake cycles. our results align with far more detailed hydrographs of the bonneville lake cycle in the main bonneville basin of currey and oviatt (1985), oviatt and others (1992), nelson (2012), and oviatt (2015, 2020). our data also constrain the prebonneville, post-cutler dam age of newly identified red-over-white double fielding geosols in the newton hill pit and lithologically similar paleosols in eastern cache valley. results overview of newton hill gravel pit on the se flank of newton hill, central cache valley, utah (figure 3), our ongoing studies have delineated the internal architecture of an east-plunging, nose-shaped compound spit deposited atop an eastsloping, eroded face of little valley gravel during the cutler dam and bonneville lake cycles. the most continuous exposures lay between ~1408 m and ~1462 m, mostly below the prominent, higher provo shoreline (janecke and oaks, 2011a, 2011b) at ~1463 m at this locality. scattered exposures continued to ~1487 m. exposures in the south-central part of the pit in 2006 were so extensive that the key stratigraphic relations and the overall architecture of the deposits were unambiguous (figure 6). the spit’s original crest flattened uphill westward into a wave-cut and wave-built platform at the higher provo shoreline of lake bonneville (figure 2). the crest of the spit was parallel to and slightly north of the southern boundary of the gravel pit (figure 6d). pre-bonneville sediment is mostly exposed in the central and western half of the gravel pit. little valley lake beds stratigraphic relationships in the newton hill pit, little valley gravel is overlain by the upper red geosol at sample site usu2490 (figure 7a). at sample site usu-2491, there is no geosol between little valley pebbly sand and overlying cutler dam sandy mud (figure 7b). at usu-1083 (figures 4b, 4d) and at usu-857 (figure 4a), little valley gravel is overlain by bonneville deposits, with no geosol between. at usu-2895 little valley marl is overlain by a fielding-like caliche paleosol beneath offshore bonneville deposits. at usu-3202 little valley gravel is overlain by thin sediment of cutler dam lake cycle, then the upper fielding geosol, beneath laminated fine-grained bonneville deposits (figures 4a, 4d). although undated, at usu-1084 probable little valley gravel underlies a local channel with ashy sand under surficial gravels with modern soil. we did not find the base of the little valley deposits, nor identify pre-little valley units. downward excavation ceased in the central part of the newton hill pit because of a noncommercial green marl 4 to 6 m thick according to two pit operators. the little valley deposits are primarily pebble to cobble gravels and sandy gravels with low dips (figure 7a). discontinuous exposures west of the dayton-oxford fault strands reached at least 8 m thick. locally there are thin marls and sand beds. in hyde park, utah, in eastern cache valley (figure 1), at sample site usu-2895, a pale green little valley marl with a thin, calcareous, fineto coarse sand lens is overlain by a white bk paleosol 0.55 m thick, in turn overlain by a thin lag cobble gravel followed upward by 2.0 m of bonneville light brown, thinly laminated, silty very fine sand with snails (cf. nearby exposure at figure 8a). elsewhere in eastern cache valley, weakly laminated to structureless marls and minor fine sands dominate probable little valley deposits. these undated older lake beds underlie the double fielding geosols and bonneville deposits, and persist at least up to ~1530 m, which is about 40 to 45 m below the local bonneville shoreline (figure 8b). m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 7 figure 3. map of staker-parsons gravel pit se of newton hill shows locations of osl and isrl age dates; contours of the tops of the extensive red fielding geosol in the s and w, the pink/white/green shrinking marl in the n and se, the laminated green clay between them; locations of geologic cross sections a a' to d d' in figure 4, and locations of figures 6a, b, c and 7a, b. r.q. oaks, jr., s.u. jänecke, t.m. rittenour, t.l. erickson, and m.s. nelson implications and hydrographs for 2 pre-bonneville pluvial lakes 8 figure 4. geologic cross sections a a' to d -d’ show extents of identified geologic units, original surface, osl and irsl age dates, pre-bonneville lake deposits, and intersections with other geologic cross sections. see figure 3 for locations. 9 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors figure 5. hydrographs showing changes in shoreline levels in the main bonneville basin and cache valley since 200 ka compared with simultaneous climatic changes. dates with error bars, ages of ashes and chrons, and sources are from table 1 and appendices 2 and were revised from oaks and others (2019). 2024 utah geological association publication 51 10 r.q. oaks, jr., s.u. jänecke, t.m. rittenour, t.l. erickson, and m.s. nelson implications and hydrographs for 2 pre-bonneville pluvial lakes figure 6. a) original exposure of cutler dam (qcd) gravel overlain by the double fielding geosols (qfg), beneath deep-water bonneville and younger provo deposits (qlb). b) exposures w from the above site showed lateral continuity of this sequence in the hanging wall of the dayton-oxford fault. the fault dips toward viewer. figure c) details of qcd, qfg, and qlb at sample site usu-856. d) map showing camera positions of figures 6a, b, c. locations shown in figure 3. m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 11 figure 7. a) little valley (qlv) deposits. truncated channel in upper left has curved sand laminae dipping toward the deepest part. fielding red geosol here extended over qlv. staff = 1.50 m. sample usu-2490 is the same altitude as the highest exposures of cutler dam (qcd) deposits ~170 m se, but below the highest (projected) qcd ~1450 m ~120 m south b) northwest edge of newton hill pit shows erosional unconformity (yellow) between little valley lake beds (qlv) and overlying cutler dam beds (qcd). there is no paleosol along this contact. gravel and fines of the cutler dam lake cycle preserve bottomset, foreset, and topset beds (orange base) that formed in the east-plunging spit. deposits are cut by two subsequent faults or slumps (red). marker beds within the spit are color-coded. see figure 3 for locations. both photos 9-26-2016. r.q. oaks, jr., s.u. jänecke, t.m. rittenour, t.l. erickson, and m.s. nelson implications and hydrographs for 2 pre-bonneville pluvial lakes and double geosols 12 figure 8. a) double paleosols in eastern cache valley (figure 1b) that lithologically match our dated qfg in the newton hill pit. here they are overlain by bonneville lag gravel (qlbg) and sand (qlbs). underlying alluvial-fan deposits (qafo) were not dated, so subaerial exposure and soil-forming could have begun before mis 3. location is at yellow dot in 8b. b) lateral extent of exposures of double paleosols in east cache valley. latitude and longitude indicate the midpoint of this image (+). black dot marks site of ~143 ka osl age (usu-2895) sampled within fine-grained little valley lake beds (qlv) beneath a calcrete. qt fluvial terrace is offset 9 m across a strand of the east cache fault zone at the black arrow. 13 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 age control (n=6) five exposures of pre-bonneville lake gravels in the newton hill pit and one exposure in hyde park, utah returned osl and irsl ages coeval with the little valley lake cycle (mis 6). the oldest age of ~169 ka (usu-2491) is from the north-central part of the newton hill pit, whereas the youngest age of ~143 ka (usu-2895) is from hyde park at ~1493 m. the latter is also the age determination from the highest elevation. the little valley lake cycle flooded cache valley to elevations well above ~1493 m, possibly as high as ~1530 m, and attained altitudes many tens of meters higher than expected (cf. scott and others, 1983). the youngest beds dated in the newton hill pit (usu-1083; table 1) are essentially the same age as that from hyde park. cutler dam lake beds stratigraphic relationships in the south-central part of the newton hill pit, east-sloping foresets of sandy, well-rounded, pebble to cobble gravels underlie the fielding geosols. the foresets there were >6 m high and extended horizontally about 200 m (figure 6a). to the north, exposures of these spit gravels are about 6 to 10 m thick and flatten into finer bottomset beds (figure 7b). there are sharp erosional contacts locally within the foresets (figure 6a). the highest exposures reach ~1443 m, but early photos (figure 6b) and projection in figure 4d suggest that the highest lake beds may have reached ~1450 m (appendix 5). bedding in the s-central part of the pit and the shape of the overlying pink marl (figure 3) indicate that the spit probably was mainly east-plunging, yet part of this spit also extended northward (figures 4a, 4b, 4d, 6b). gravels to the north intertongue with underlying green, silty, fine-sandy laminated marl 2 m thick (figure 7b). most gravel lenses there thin downward and pinch out to the north between intercalated marl layers that thin upward and pinch out to the south. two fault or slump surfaces offset the gravels in the north. these offset the contact between topsets and foresets (figure 7b). age control (n=2) two samples from this deposit in the southcentral part of the newton hill pit yielded osl dates of ~67 ka (usu-856, -858). these are coeval with the cutler dam lake cycle and mis 4 (figures 3, 4, 5; table 1). ashy channel fill near the former west margin of the newton hill pit, a white reworked ashy fine sand filled a scour below thin surficial gravel and modern soil, ~1483 m. satellite imagery (8-11-2011) in google earth pro shows this narrow channel trended nnw-sse. probable little valley beds below this channel dip ~4o west and roll over eastward to dip gently east. the upper part of the probable little valley beds are truncated eastward at the pre-cutler dam erosional face (figure 7a). this subaerial channel fill yielded an osl age of ~54 ka (usu-1084), during mis 3 (table 1; figures 3, 4c). this is older than the upper fielding geosol but younger than the cutler dam gravels exposed lower in the newton hill pit and the shallowwater cutler dam muds in the type area southwest of cutler narrows (figure 5). double fielding geosols in newton hill pit stratigraphic relationships in the original south-central part of the pit, two successive geosols developed above and partly within the top of underlying gravel foresets of the cutler dam lake cycle (figure 6a). this unit consists of an upper, humid-climate, red-weathering, loessdominated interval and a lower, arid-climate, white caliche interval. the contact between the two geosols is primarily erosional, but locally gradational. in one place the upper geosol is separated from overlying deep-water bonneville deposits by a thin gravel wedge up to 1 m thick (figure 6c). the lower of the two geosols typically has only an eroded lower bk horizon, up to 1.5 m thick, above the cutler dam foreset gravels. this geosol pinches out east and west of the south-central part of the pit, and does not reach the east strand of the dayton-oxford fault westward in the pit (figures 6b, 6c). calcite in the lower geosol penetrated down into the cutler dam foreset gravels beneath (figure 6a). it has amalgamated subhorizontal stringers of carbonate and amorphous nodules. pieces of the eroded caliche are common in the lower part of the red geosol above (figure 8a). the eroded upper contact of the caliche has distinct channels up to 15 cm deep filled with, and overlain by, as much as 2.5 m of the red geosol. the upper geosol is mainly loess and slightly pebbly loess, although locally it contains abundant colluvium. it has considerable organic material, exhibits downward displacement of clay, has a distinctive reddish soil hue (10r5.5/4), displays little cementation, and has a few vertical calcite stringers, but lacks caliche nodules except those reworked into the base 14 r.q. oaks, jr., s.u. jänecke, t.m. rittenour, t.l. erickson, and m.s. nelson implications and hydrographs for 2 pre-bonneville pluvial lakes (figure 6c). its top has a less prominent erosion surface than its base. this upper geosol thickens to 5 m or more westward, near the hanging wall of the dayton-oxford fault (figures 6b, 7a), and locally on the north flank of the cutler dam spit, where the caliche geosol is absent (figures 4a, 4d). there its upper part is colluvial gravelly mud overlying 2 to 3 layers of gravelly loess with weak subsoils. locally in the north it pinches out eastward beneath a gray modern soil at the original surface of the pit. where absent in the east part of the pit, and locally in the north part of the pit, the upper contact of the double geosols is marked by a lag gravel or the green marl (figures 4a, 4d) at the base of the bonneville deposits above cutler dam foresets. surveyed contacts of the top of the reddish geosol suggest that it probably rose at least to ~1463 m in the west part of the pit (figures 4a, 4b, 4c). it descended to below ~1441 m in the se part of the pit, and to below ~1444 m locally northward (figures 3, 4d). erosion probably removed these geosols from the lower and higher parts of the present newton hill pit before bonneville deposits were laid down. the absence of the fielding geosols in the footwall of the dayton-oxford fault makes it challenging to estimate the throw across the fault, although it must be >2 m. age control (n = 1) in the s-central part of the pit, the middle part of the red geosol, ~1444 m, contains a lens of sandy sediment that yielded an osl age of ~39 ka (usu-855) (figures 3, 4b, 6c). this dates to the penultimate interglacial, mis 3c (lisiecki and raymo, 2005). double geosols in eastern cache valley in hyde park and north logan, utah, in eastern cache valley, we found numerous examples of prebonneville double paleosols in many trenches for utilities and in basement and landscape excavations (figure 8b). these paleosols are essentially identical to those in the newton hill pit, with a red clay-rich paleosol over an eroded white caliche paleosol. several of the lower exposures have only the eroded lower white bk paleosol, locally with a very thin, eroded, red paleosol above. detailed local mapping with an abney hand level near sample site usu-2895 demonstrated an undulose paleotopography beneath the paleosol with lateral changes in the underlying sediments uphill and laterally. all exposures lie above the highest cutler dam deposits in the newton hill pit. in eastern cache valley, either the double paleosol, loess deposits, or a gravel lag underlie the bonneville offshore sand with snails (west, lower) and bonneville gravel or post-bonneville colluvial gravel (east, higher), respectively (figure 8b). the white caliche paleosol overlies dated little valley marl (~143 ka; usu-2895) at ~1493 m in hyde park, and both paleosols overlie undated alluvial-fan debris flow deposits at ~1526 m in exposures farther east (figure 8a). exposures of these widespread double paleosols were recorded through a vertical range of at least 124 m and a horizontal separation of at least 2.7 km nnw -sse (figure 8b). the highest exposure, at ~1607 m, is above the bonneville shoreline (41.78501, 111.77766). our current concept of the spatial and stratigraphic relations of the lake cycles and intervening paleosols is shown in figure 9. bonneville lake beds stratigraphic relationships bonneville deposits originally blanketed the spit in the area of the newton hill pit (figure 2). in the southern exposures, topsets and foresets of sandy pebble to cobble gravels of the bonneville lake cycle (figure 6a) grade downward into finer bottomsets that overlie more than 3 m of transgressive deepwater marls and laminated silty sand (figure 6c). northward, where the pre-bonneville relief was lower, deposition included lower green marls and a single higher pink marl that form distinctive marker beds (figures 3, 4) between thicker bonneville gravels (figure 10). the pink marl is a calcareous, very fine sandy, clay-rich silt. it is plastic, weakly laminated, and thin (tens of cm thick). it is either pink throughout (oxidized reddish orange (5r7/2) or greenish-gray to whitish color at the base. it might be gilbert’s “white marl”, which dates from the highstand of lake bonneville. its red stain may be due to iron supplied by the proximal bear river. locally, a lower green bonneville marl directly overlies cutler dam deposits where the fielding geosols are absent (figures 4a, 4d), but there are other traceable pale greenish marls higher in the bonneville sequence. several marls produced low-angle slip surfaces that repeat layers within the bonneville deposits in small slumps and slides (figure 10). these might have been triggered by earthquakes, the bonneville flood, or both. age control (n=4) near the south-center margin of the pit, gently east-dipping, gray, laminated silty sand yielded an osl age of ~22 ka (usu-854) (figures 3, 4b, 4c). 15 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 figure 10. exposure of distinctive marls within bonneville (qlb) nearshore gravels in newton pit. the pink marl is widespread in the newton hill pit whereas the underlying laminated green marl is more restricted. thad erickson = 1.8 m. nw part of pit. the stratigraphic position and gravel of the slump above the pink marl suggests a possible trigger by the bonneville flood. the pink marl records the deepest water depths. figure 9. schematic cross section of the relative geometries of deposits of three pluvial lakes in cache valley, intervening double soils, the modern geosols, and the modern surface soil on the double paleosols above the bonneville shoreline in eastern cache valley and the newton hill pit. qlb = bonneville lake cycle; qlbb = bonneville shoreline; qlbp = provo shoreline; qfg = double fielding geosols; qcd = cutler dam lake cycle; qlv = little valley lake cycle; mis = marine oxygen-isotope stage. although we found no distinct mis 5 paleosol developed on qlv, it might be incorporated in the base of qfg above qcd deposits. above the bonneville shoreline, modern soil is developing on and augmenting exposed qfg. horizontal scale is tens of kilometers. concept from oviatt and others (1987). any paleosols within lake cycles are omitted. altitudes are not corrected for rebound. 16 r.q. oaks, jr., s.u. jänecke, t.m. rittenour, t.l. erickson, and m.s. nelson implications and hydrographs for 2 pre-bonneville pluvial lakes these deep-water deposits sharply overlie an unconformity and the upper fielding geosol (figure 6b). a laminated silty sand lens in cobble gravel from the lower part of bonneville deposits, in the central part of the pit, yielded a slightly younger osl age of ~21 ka (usu-1082). there the bonneville deposits directly overlie eroded little valley gravels (usu-1083) with no paleosol between (figures 4b, 4d). a sand bed intertongued with northeast-dipping gravel beds in the north-central part of the pit yielded a post-flood provo age of ~15 ka (usu-859) (figure 3). at muley hill in millville, utah, gravel atop an eroded delta, at ~1550 m, between the bonneville and provo levels, yielded an age of ~21 ka (table 1). subsurface evidence of pluvial lakes drillers’ logs from >1000 water wells across the center of cache valley southeast and east of newton hill document two gravelly layers and two clay-rich layers in the subsurface. an upper confining silty clay (marl?) unit ~18 m thick, an intervening, persistent gravel unit ~9 m thick, and a lower confining silty clay unit ~9 m thick, overlie thick underlying gravel and sand (williams, 1962; bjorklund and mcgreevy, 1971; clyde and others, 1984; kariya and others, 1994; robinson, 1999; thomas and others, 2011). figure 11 shows these relations along part of u.s. highways 89/91 (figure 1). within the upper confining layer there are typically two horizons of nonpersistent gravels associated laterally with oxidized brown silty clays. gray, blue, or black silty clays lie both above and below these gravel and oxidized intervals. the lower confining layer also encloses lenses of gravels and related oxidized horizons. these clays overlie sandy gravels of pleistocene age and underlying older gravels in the salt lake formation that cumulatively reach ~150 m to 300 m thick between smithfield, wellsville, and hyrum (robinson, 1999) (figure 1a). these coarse sediments are the principal aquifer in cache valley (figure 11). the unoxidized clays probably are deep-water lake deposits. they likely are coeval with the three lacustrine deposits in the newton hill pit, and perhaps earlier pluvial lakes in the main bonneville basin (williams, 1962), including older lake cycles identified in the saltair and burmester cores (eardley and gvosdetsky, 1960, eardley and others, 1973; williams, 1994; oviatt and others, 1999). the gravels and oxidized muds at distinct levels within the unoxidized muds either indicate interglacial epochs or major oscillations within long pluvials (williams, 1962; this study). in the southwest part of figure 11, a persistent gravel within the upper confining layer may be a chance intersection laterally along a former stream channel. figure 11. geologic cross section showing alternating pluvial fines (blue) and interglacial gravel and sand deposits (orange) beneath the low part of cache valley. this section is through college ward in central cache valley, utah, along u.s. highways 89/91. qlb = bonneville; qcd? = cutler dam; qlv? = little valley. see figure 1a for location. question marks indicate that correlations with other lake cycles are possible. williams (1962) first documented these repeating coarse and fine intervals of lacustrine and fluvial deposits in drill holes in five geologic cross sections across the utah part of cache valley. 17 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 elsewhere in cache valley, the underlying salt lake formation has many different lithologies. these include conglomerates, tuffaceous green (zeolitebearing) to dark and light gray shales, sandstones, siltstones, thick to thin, pale brown very fine crystalline (micritic) limestones, oolitic limestones, and diabase (adamson and others, 1955; goessel and others, 1999; oaks and others, 1999; janecke and evans, 1999; janecke and others, 2003). these distinctive lithologies are repeated by extensional folds and normal faults, so that the salt lake formation commonly exhibits tilts. dips as high as 78° distinguish the salt lake formation from the overlying quaternary deposits (oaks, 2000). discussion correlations related to the hydrographs overview of the hydrographs data for the hydrographs in figure 5 are in table 1 and appendices 1 to 5. the hydrographs for cache valley and the main bonneville basin show good correlation of lake highstands and lowstands in both basins and also with the o-isotope marine record of climatic fluctuations for mis 6 through mis 1. the little valley lake rose higher than the local provo shoreline, whereas the cutler dam lake cycle did not rise quite as high (figure 9). the mis 5 interglacial persisted for ~55 kyr, twice as long as the ~27 kyrlong mis3 interglacial. the isotopic data also suggest that mis 5 was warmer than mis 3 (figure 5). yet there are no distinct, widespread soils associated with mis 5 in cache valley, and the only possible exceptions elsewhere are the promontory/dimple dell geosols in the little valley pit and other parts of the main bonneville basin (scott and others, 1983). it is noteworthy that well-dated fielding humid soil and the underlying arid soil, both of which are widespread in cache valley, formed during the relatively short and mild interglacial mis 3 before the bonneville lake cycle, yet they are exceptionally thick and robust paleosols (figures 3, 5, 6, 8). the lake was at least 150 m deep during the 68 to 67 ka part of the cutler dam lake cycle, in the early part of mis 4, yet it coincided with a relatively minor oscillation in the climate record (lisiecki and raymo, 2005). this seems anomalous compared with the climatic and hydrologic conditions that favored deep lakes during mis 6 and mis2. the oscillations of benthic marine isotopes are only about 60% as intense during mis 4 as during mis 6 and mis2 (figure 5). perhaps deep lakes can form with less milankovitch forcing than glaciers. alternatively, added water may have begun to flow across a waterfall in oneida narrows into cache valley then, followed ~20 kyr later by the complete, final diversion of the bear river into the bonneville basin (pederson and others, 2016). a complex history of incision of oneida narrows is suggested by one or more widespread subsurface gravels below a mud layer under surficial gravel from oneida narrows through several kilometers downstream in drillers’ logs of water wells (oaks, 2010). little valley lake cycle the age and duration of the little valley lake cycle is constrained by our six new absolute ages, five published aar estimates, one published tl age, and one extrapolation from the estimated rate of formation of the overlying promontory paleosol (figure 5; table 1; appendices 2, 3, 4). combination of all the data for the main bonneville basin (blue dashes in figure 5) suggests that the little valley lake cycle might have persisted 20-30 ky into interglacial mis 5. however, an end closer to 123 ka, at the end of mis stage 6, is more likely based on the climate record and our new absolute ages (preferred model in figure 5). in the main bonneville basin, altitude control for the little valley lake cycle is limited, with some corrected for rebound, others not (mccoy, 1981, 1987; scott and others, 1982, 1983). the highest probable but undated little valley gravels in the main bonneville basin are at ~1512 m in the geneva quarry at point of the mountain, south of salt lake city (scott and others, 1983) and at ~1517 m in the little valley pit, where they were initially misidentified as “alpine” by morrison (1965, 1966) and reinterpreted by scott and others (1983). these older lake beds are about half way between the local bonneville and provo shorelines (scott and others, 1983). in cache valley the highest dated little valley deposits, at ~1493 m in hyde park are sandy, weakly laminated marl, and undated deposits traced uphill from dated beds in the upper newton hill pit, at ~1483 m. these also lie between the bonneville and provo shorelines. thus, the highest level attained by the little valley pluvial lake is not certain, but elevation ranges are high and roughly similar in both basins (appendix 5). active tectonics in both basins may have raised or lowered individual sites, which is especially critical for older lakes. further discovery of higher shoreline exposures and absolute ages are needed to determine if the actual highest water levels of the little valley lake cycle were the same or different across cutler narrows. 18 r.q. oaks, jr., s.u. jänecke, t.m. rittenour, t.l. erickson, and m.s. nelson implications and hydrographs for 2 pre-bonneville pluvial lakes cutler dam lake cycle our ages of cutler dam deposits in cache valley confirm that this pluvial lake rose at least 110 m above that of marshy sediments in the type area (oviatt and mccoy, 1988, 1992) in the main bonneville basin sw of cutler narrows (figures 1b, 6b; appendix 5). two irsl ages from cutler dam deposits and two from the base of the fielding geosol in the type area averaged ~59 ka (kaufman and others, 2001). this is younger than the average of ~67 ka for two osl ages near the higher level in cache valley. although the error limits of the ages from both sites overlap slightly (figure 5; table 1; appendix 2), the central ages differ. these data may indicate a drop to the lower level near the end of mis 4, consistent with the climatic data (figure 5). additional osl age control from distal cutler dam beds in the newton hill pit would further constrain the hydrograph in figure 5. re-dating lake beds between those of the little valley and bonneville lake cycles in hansel valley (robison and mccalpin, 1987) with osl might show that they are coeval with the cutler dam lake cycle, which seems likely. fielding double geosols our dated samples of the upper fielding geosol and the ashy sand channel fill indicate that subaerial deposition replaced the cutler dam lake after the end of mis 4, at ~56 ka. our double geosols are similar to the sequence described by kaufman and others (2001, p. 324) in the type area sw of cutler narrows. their figure 2 showed three successive geosols that comprise their fielding geosol, described in the figure as: “massive red-brown silt and clay; at least three petrocalcic horizons, each topped by a snail-rich horizon; oxidized rootlets on blocky weathered surfaces”. the similarity of our double geosols to descriptions of the promontory and dimple dell double paleosols in the little valley pit (morrison,1965) is also striking. there, a lower caliche geosol and an upper red (10yr) loess-derived geosol lie between little valley and bonneville deposits. all exposures there are above the highest known cutler dam lake beds in cache valley. despite the similar lithologic features and nearly identical stratigraphic relationships, the promontory and dimple dell geosols are interpreted to be much older, ~104 ka (scott and others, 1983; their table 5). if so, the promontory and dimple dell palesols are significantly older than the fielding geosols. absolute ages are needed to resolve this puzzle. an osl age is needed in hyde park within the double paleosols there, to determine if these paleosols are definitely coeval with, or differ in age from, the dated upper fielding geosol in the newton hill pit. bonneville lake cycle the final diversion of the bear river into cache valley ~ 45 to 55 ka (pederson and others, 2016) was too late to raise the cutler dam lake, and all earlier lakes, above a divide ~2 km north of red rock pass, at the north end of cache valley (gilbert, 1890) (figure 1a). its final addition raised lake bonneville higher than earlier lakes, to overtop that divide (bright, 1963; hochberg, 1996; bouchard and others, 1998; link and others, 1999; janecke and oaks, 2014; pederson and others, 2016; utley, 2017). an earlier overflow across oneida narrows (oaks, 2010) may have raised the cutler dam lake above that expected from the o-isotope data (figure 5). our two osl ages of ~21 ka in the newton hill pit, at ~1422 m and ~1447 m, lie within the wide envelope of 14c dates with confidence intervals for the rising limb of the bonneville transgression in the main bonneville basin (cf. oviatt, 2015, 2020). however, both are minimum depths for the lake level at those times. furthermore, well-rounded gravels at muley hill, with an age of ~21 ka, at ~1550 m elevation, is close to the local bonneville shoreline at ~1573 m (figure 9), and above the oviatt envelope of dates. janecke and others (2013) obtained a 14c age ~22 ka in nearshore sands at ~1500 m in a gravel pit at the mouth of green canyon in eastern cache valley, between logan river and city creek (figure 1a), somewhat above the oviatt envelope. thus, although the age-altitude data from cache valley plotted in figure 5 might suggest a slightly earlier rise of lake bonneville during its transgression, the data do not differ enough from those compiled in oviatt (2015, 2020) to be compelling. more precise and diverse age control is needed to improve the earlier curve for lake bonneville, which was compiled from 14c age determinations. we believe that a prolonged bonneville highstand during oscillatory (?) overflow to the north, is needed to explain high, steep, wave-cut bedrock cliffs at the bonneville shoreline throughout the bonneville basin (janecke and others, 2019). significant time is also required to backfill gem valley, oneida narrows, lower bear river-mink creek canyon, and finally deposit the large bonneville delta north of preston, idaho, with a surface area of >125 km2 in cache valley (figures 1a, 1b). the bonneville delta of the bear river back-filled a reach that was ~55 km long, between gem valley and northeast cache valley (janecke and oaks, 2011b). 19 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 implications for incision of cutler narrows it is unclear if cutler narrows was incised well below the ~1450 m cutler dam gravels of cache valley before the bonneville flood because the evidence is incomplete and inconclusive. sr isotopes indicate likely entry of water of the bear river west of cutler narrows during both the little valley and cutler dam lake cycles (hart and others, 2004). the flow could have been through a fully incised cutler narrows, with lakes at the same or similar levels on both sides, or as flow across a lip near or slightly below the ~1450 m cutler dam gravel of cache valley that separated lakes with different levels. although oviatt and mccoy (1988, 1992), oviatt and others (1987), and kaufman and others (2001) found no deep-water cutler dam deposits in ~15 m of shallow-water cutler dam deposits west of cutler narrows, such could be present in the subsurface there. several arguments suggest that deep incision almost to the modern level of the bear river is a reasonable interpretation of the existing data. these arguments include: (1) the >4 ma age of the east side of the horst block, so that considerable time was available to incise the canyon at cutler narrows; (2) the short and low canyon in cutler narrows, compared to dozens of deeper and longer canyons cut by streams with a fraction of the discharge nearby (e.g. logan canyon), which include some carved by now minor and intermittent streams (e.g. weston canyon); and (3) subsurface fluvial (?) sand and gravel deposits, hundreds of meters thick, that alternate with clay and silt that settled from lakes (williams, 1962). this facies pattern continues from the quaternary units down into the underlying pliocene salt lake formation (~12 to ~2[?] ma; goessel and others, 1999; oaks and others, 1999; janecke and others, 2003) (figure 11). the thick and laterally continuous fluvial (?) gravels beneath the center of cache valley suggest protracted external drainage because continuous playa and lake deposits would have formed if there had been a long-lived barrier in cutler narrows. williams (1962) also argued that external drainage during most of the pleistocene is required to produce the consistently thin quaternary deposits beneath cache valley. to determine if pre-little valley lakes extended through cutler narrows and how high they reached relative to those in the main bonneville basin, absolute ages are needed from more lake beds between the provo and bonneville shorelines in both basins. a continuous core where the quaternary deposits are thickest in cache valley, perhaps near the location of figure 11, could provide further age control. altogether, we conclude that the narrow, low horst between cache valley and the main bonneville basin was probably breached early because it is neither high enough nor wide enough to separate high pluvial lakes for an extended period of time (figure 1). much, possibly nearly all, of the excavation of cutler narrows in bedrock probably took place before the little valley lake cycle (oaks and others, 2014, 2019, 2020; cf. maw, 1968, hunt, 1982). complete resolution could come from finding: (1) ~59 ka cutler dam shallow-water lake beds in cache valley near the same elevation as the cutler dam beds in the type section; or (2) high-elevation cutler dam beds in the main bonneville basin that date from ~67 ka; or (3) that the dated cutler dam gravels between 1450 1410 m in cache valley are coeval with the lowelevation shallow-water deposits in the main bonneville basin. conclusions our 14 new osl and irsl ages establish the first evidence of cutler dam lake deposits and double fielding geosols, and provide the first absolute ages of little valley deposits in cache valley. our quantitative hydrographs show firm correlation of deposits in cache valley with the little valley (mis 6), cutler dam (mis 4), fielding (mis 3), and bonneville (mis 2) units in the main bonneville basin. none of our contacts between dated sediment of the little valley and cutler dam lake cycles preserve paleosols. in contrast, our double fielding geosols lie between well-dated cutler dam and bonneville deposits up to the highest near-shore gravel deposits of the cutler dam lake cycle in the newton hill pit (figures 6, 7a). higher in the newton hill pit and in hyde park (figure 8a) double paleosols lie between the little valley and bonneville deposits. above the bonneville shoreline in north logan (figures 8b, 9) they lie above pre-bonneville loess and alluvial-fan deposits. these paleosols consistently exhibit an eroded arid-climate white calcic bk horizon overlain by a loessic humid-climate red soil, and thus are provisionally correlated here with the dated fielding geosols in the newton hill pit despite the absence of additional geochronology. drillers’ logs of water wells identify two thick, confining clay-rich layers separated by a continuous gravel layer. these overlie thick gravels of the gravels of the principal aquifer of cache valley (figure 11). each confining clay sequence contains local gravels with adjacent oxidized clays that may indicate emergence due to oscillations within protracted lake cycles or interglacial episodes between pluvials. lake deposits older than little valley may be present here. the majority of incision of cutler narrows proba20 r.q. oaks, jr., s.u. jänecke, t.m. rittenour, t.l. erickson, and m.s. nelson implications and hydrographs for 2 pre-bonneville pluvial lakes bly predates the little valley lake cycle. although the evidence for when cutler narrows was cut below the ~1450 m cutler dam deposits in cache valley is incomplete, we believe that the evidence supports early incision to near its present depth. acknowledgements we thank reviewers jeremiah bernau, paul inkenbrandt, daren nelson, jack oviatt, an anonymous reviewer, and editor hugh hurlow for constructive suggestions for improvements to the text, figures, and tables, for pertinent references we had missed, and for valuable insights. james p. evans aided us in many ways in addition to careful reviews. we benefited greatly from staker-parsons’ permission to visit their newton hill pit repeatedly through 17 years, and thank sam waterman and steve tucker for their keen interest in our work, knowledge of layering in the pit before our first visit in 2006, and willingness on occasion to clean faces or excavate areas of interest between loading customers’ trucks. our colleague alexis ault discovered the double palesols in se hyde park, utah. tomas capaldi assisted in collecting sample usu-3243. hugh hurlow and mike vanden berg deftly and kindly shepherded this manuscript to completion. references adamson, r.d., hardy, c.t., and williams, j.s., 1955, tertiary rocks of cache valley, utah and idaho: utah geologic association guidebook 10, p. 1-22. auclair, m., lamothe, m., and huot, s., 2003. measurement of anomalous fading for feldspar irsl using sar: radiation measurements, v. 37, p. 487-492. bjorklund, l.j., and mcgreevy, l.j., 1971. groundwater resources of cache valley, utah and idaho: utah department of natural 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implications of their unexpected high altitudes for excavation of cutler narrows from a level above https://geodata.geology.utah.gov/pages/download.php?direct=1&noattach=true&ref=5845&ext=pdf&k= https://geodata.geology.utah.gov/pages/download.php?direct=1&noattach=true&ref=5845&ext=pdf&k= https://geodata.geology.utah.gov/pages/download.php?direct=1&noattach=true&ref=5845&ext=pdf&k= http://geology.isu.edu/snakeriver/proceedings.pdf http://geology.isu.edu/snakeriver/proceedings.pdf 23 m.d. vanden berg, r. ford, c. frantz, h. hurlow, k. gunderson, g. atwood, editors 2024 utah geological association publication 51 1494 m (4901'), down to the present 1314 m (4310') mainly during the bonneville lake cycle: utah geological survey symposium on lake bonneville, 3-4 october, 2018. oaks, r.q., jr., janecke, s.u., rittenour, t.m., erickson, t.l., and nelson, m.s., 2020, pre-bonneville lakes in cache valley and excavation of cutler narrows across the cache butte divide in the west, north-central 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in‐situ    h2o  (%)3  grain  size  (µm)  k (%)4  rb  (ppm)4  th  (ppm)4  u  (ppm)4  cosmic  (gy/ka)  osl/  irsl5  859    16.2    22 (57)  1.76  ± 0.07  27.11  ± 2.17  15.42  ± 1.39  5.9  (15%)  90‐150  1.14  ±0.03  52.0  ±2.1  6.2  ±0.6  1.7  ±0.1  0.05  ±0.01  osl    3243    1  22 (30)  0.80  ± 0.04 6  16.74  ± 1.92  20.98  ± 3.04  1.81  150‐  250  0.59 ±0.01  0.26 ±0.01  0.51  ±0.01  15.0 ±0.15  5.9 ±0.06  11.7  ±0.12  1.76 ±0.2  0.67 ±0.07  1.32  ±0.1  1.0 ±0.1  0.7 ±0.04  1.0  ±0.1  0.25  ±0.02  osl  1082  35.1  11 (42)  2.17  ± 0.09  46,31 ±  12.75 4  21.35  ± 3.48  7.4  150‐  250  1.48  ±0.04  66.5  ±2.7  8.8  ±0.8  1.9  ±0.1  0.02  ±0.00  osl  854  11.3  24 (37)  2.98  ± 0.12  64.72  ± 9.88 4  21.72  ± 2.78  14.4  90‐150  1.91  ±0.05  97.2  ±3.9  12.3  ±1.1  2.4  ±0.2  0.08  ±0.01  osl  855  21.6  24 (49)  3.90  ± 0.16  153.29 ±  15.01  39.28  ± 3.72  10.2  63‐150  2.41  ±0.06  119.5  ±4.8  14.6  ±1.3  3.4  ±0.2  0.04  ±0.00  osl  1084  3.0  13 (32)  2.74  ± 0.11  146.65 ±  19.32  53.51  ± 6.44  12.7  75‐150  1.72  ±0.04  74.3  ±3.0  10.5  ±1.0  1.8  ±0.1  0.19  ±0.02  osl  856  24.7  20 (42)  1.77  ± 0.07  118.71  ± 8.36  66.82  ± 5.94  1.9  125‐  250  1.03  ±0.03  40.9  ±1.6  6.8   ±0.6  1.9  ±0.1  0.03  ±0.00  osl  858  24.7  16 (57)  1.56  ± 0.06  92.17  ± 9.47  67.70  ± 6.46  3.2  150‐  250  1.27  ±0.03  34.7  ±1.4  4.4  ±0.4  1.1  ±0.1  0.03  ±0.00  osl  2895  2.8  16 (29)  1.22  ± 0.05  173.72  ± 14.76  142.8  ± 13.1  ‐  150‐  250  0.69  ±0.02  24.9  ±1.0  3.5  ±0.3  1.2  ±0.1  0.19  ±0.02  osl  1083  38.1  15 (34)  1.14  ± 0.05  164.12  ± 19.32  144.3  ± 14.5  3.3  150‐  250  0.85  ±0.02  29.7  ±1.2  3.9  ±0.4  0.9  ±0.1  0.02  ±0.00  osl  a-2 usu‐  sample  number    depth  (m)  number of  aliquots1  dose  rate  (gy/ka)  de 2 ±  2σ         (gy)  age3  ±  2σ (ka)  in‐situ    h2o  (%)3  grain  size  (µm)  k (%)4  rb  (ppm)4  th  (ppm)4  u  (ppm)4  cosmic  (gy/ka)  osl/  irsl5  3202  59.4  19 (34)  1.39  ± 0.06 7  208.47  ± 31.80  150.0  ± 25.9  7.6  63‐250  1.09  ±0.03  1.05  ±0.03  0.43  ±0.01  49.0  ±2.0  31.4  ±1.3  13.4  ±0.5  5.8  ±0.6  5.27  ±0.5  1.81  ±0.2  1.3  ±0.1  1.1  ±0.1  0.7  ±0.1  0.01  ±0.00  osl  2490  32.0  15 (17)  2.29  ± 0.10 8,9  234.56  ± 25.82  155.6  ± 21.4  3.8  125‐  250  0.73  ±0.02  1.06  ±0.03  23.6  ±0.9  25.2  ±1.0  4.0  ±0.4  3.6  ±0.3  1.0  ±0.1  1.0  ±0.1  0.02  ±0.00  irsl  857  13.4  20 (63)  0.94  ± 0.04  152.56  ± 19.30  161.5  ± 16.8  3.7  90‐250  0.66  ±0.02  22.0  ±0.9  2.6  ±0.2  0.6  ±0.1  0.07  ±0.01  osl  2491  38.7  23 (36)  1.13  ± 0.05  191.02  ± 28.12  169.4  ± 28.6  3.8  125‐  250  0.74  ±0.02  19.6  ±0.8  4.0  ±0.4  1.0  ±0.1  0.01  ±0.00  osl  1 number of aliquots used in age calculation and number of aliquots analyzed in parentheses.  2 equivalent dose (de) calculated using the central age model (cam) of galbraith and roberts (2012), unless otherwise noted.  3 assumed 10±3% for moisture content over burial history for in‐situ values <10%, excluding usu‐859.  4 radioelemental concentrations determined using icp‐ms and icp‐aes techniques; dose rate is derived from concentrations by conversion factors from guérin  et al. (2011).  5 osl age analysis using the single‐aliquot regenerative‐dose procedure of murray and wintle (2000) on 1‐2mm small‐aliquots of quartz sand. irsl age analysis  using the two‐temperature step (50°c, 225°c) pir irsl protocol of buylaert et al. (2009) on 1‐2 mm small‐aliquots of potassium‐rich feldspar. irsl age on  each aliquot corrected for fading following the method by auclair et al.  6 grain‐size based internal beta dose rate determined assuming 12.5% k and 400ppm rb using mejdahl (1979). alpha contribution to irsl dose rate determined  using an efficiency factor, or ‘a‐value’, of 0.09±0.01 after rees‐jones (1995).  7 dose rate includes weighted average of radioelemental chemistry based on sand fraction (top value, 35%) and gravel fraction (bottom value, 65%). a-3 appendix 2. data for hydrographs. see appendix 3 for aar correlations supporting ~417 ka age for qpp (oviatt and others, 1999). a-4 appendix 3. amino-acid-racemization data interpolated betwen known ages of ashes, chrons and 14c and tl data in the bonneville basin prior to our study. see appendix 4 for data. a-5 appendix 4. sources for original aar data used for appendices 2 and 3. kaufman and others (2001), west side canal; thatcher valley id (qmc); and little valley pit unit  sample  fossils  (#  of  samples)  racemization  (d/l) table 5  (aile/ile)  14c age ka  calibrated  table 2  tl age ka 8 hr uv  table 4  irsl age ka 3+ hr  sun table 4  qlb  k‐1 (#6)  c (8)  li (8)  0.202+0.009  0.154+0.014  14.5+0.4 a  12.2+1.3  12.0+1.2  qlb  k‐2 (#5)  h (8)  li (13)  0.184+0.011  0.221+0.016  ~24 b  19+2  23+2  qlb  (?)  (#4)        63.0+6.0  25.4+3.6  32.4+3.1  (16 hr sun)  qcd  (#3)      ~43 c      qcd  k‐3 (#2)  c (14)  li (9)  0.254+0.014  0.235+0.014    55.6+5.2  59.3+5.2  qcd  k‐4 (#1)          59.0+5.5  qmc  k‐5  c (12)  0.347+0.136  1.5 m above  st helens ash  ~110 ka  qlb  k‐6 (lv)  c (12)  0.237+0.015  20.2+0.3 d      qlv  k‐7 (lv)  c (23)  0.414+0.021        a fluminicola sp.  b limnocythere spp.   c heliosoma sp.  d arenicola sp.  e figure 5; interpolation based on bouchard and others (1998) in gem valley near thatcher id  oviatt and others (1999), burmester core   oviatt and others (1994b), leamington canyon unit  sample  fossils  (#  of  samples)  racemization  (d/l)  (alle/ile)  ages in ka and basis  qlb  o‐5  a (11)  0.16+0.03  14 ages between 14 & 21 (their table 1)  qlb  o‐6  l (7)  0.12+0.02    qlv  o‐7  a (4)  0.40+0.06  230th >90 ka & ~140 ka (their table 1)  unit  sample  fossils  (#  of  samples)  racemization  (d/l)  (alle/ile)  ages in ka and basis  qlb  o‐1  c, li (5)  0.25+0.01  ~20; numerous 14c ages  qlv  o‐2  c, li (52)  0.35+0.03  ~150+20; scott and others (1983); 230th & extrapolation from  ca accumulation rate  qpp  o‐3  c, li (15)  0.48+0.02  ~417+55; interpolation between qlv& qlcb  qlcb  o‐4  c, li (15)  0.55+0.02  ~620; below ~602 lava creek b ash, above ~760 brunhes chron  base  a-6 oviatt and others (1987), west side and hammond canals below cutler dam unit  sample  fossils racemization (# of samples) (d/l) (alle/ile)  qlb f  o‐8  l  0.11+0.03  s  0.12+0.01  a  0.15+0.03  qlb  o‐9 (#3, 5)  l (1)  0.06    s (1)  0.011  a (5)  0.10+0.005  qcd  o‐10 (#5, 6)  l (2)  0.12+0.01      s (3)  0.15+0.01      h (3)  0.11+0.01           v (3)      0.14+0.01  qlv f  o‐11  l  0.27+0.03    a  0.32+0.03  f average for bonneville basin from mccoy (1981) and this paper  scott and others (1983) cache valley (table 2) unit  sample  location  fossils racemization (# of samples) (d/l) (alle/ile)  qlb  s‐1 g  r  l (?)  0.08+0.01  qlb  s‐2  r  l (?)  0.14+0.00  qlb  s‐3  sm  l (1)  0.11  qlb  s‐4  sm  a (1)  0.14  qlv  s‐5 g  r  l (?)  0.24+0.01  qlv  s‐6 g  r  a (?)  0.42+0.06  qlv  s‐7  sm  a (?)  0.33+0.01  r = ramsbottom pit; sm = smart mountain  scott and others (1983) bonneville basin and cache valley, combined averages (table 1) locations: b, bc, g, jn, k, lc, lv, mc, mo, mu, or, p, r, sm, w (see appendix 3) unit  sample  fossils racemization (# of samples) (d/l) (alle/ile)  qlb  s‐8  l (50) 0.11+0.03  qlb  s‐9  a (35) 0.15+0.04  qlb  s‐10 gh  l (33) 0.15+0.04  qlb  s‐11 gh  a (28) 0.19+0.04  qlv  s‐12     l (2)  0.30+0.02  qlv  s‐13  a (13) 0.34+0.03  qlv  s‐14 gh  l (10) 0.33+0.08  qlv  s‐15 gh  a (28) 0.44+0.06  g 1980 preparation differed from the other samples and resulted in higher values  h  table 5 and page 280: assumed constant rate of addition of calcium to promontory paleosol based on rate in  post‐qlb soils = 70 ka to 120 ka plus 20 ka for burial by qlb = 90 ka to 140 ka for top of qlv.  230th qlv age:  > 105  ka from kaufman and broecker (1965, p. 4035). oviatt and others (1999) assumed ~150+20 ka for average age of  qlv (see above).  scott and others (1988) unit  sample  location  fossils racemization (# of samples) (d/l) (alle/ile)  qlv  s‐16  g  a (?)  0.47+0.02  a-7 mccoy (1981; 1987) unit  sample  location  # of  sample s  fossils racemization  (d/l) (alle/ile)  qlb  m‐1  jv, lc, lv, m, p, sm, tm, u  l (22)  0.11+0.01  qlb  m‐2  jv, lc, lv, m, sm  a (190  0.16+0.01  qlb  m‐3  lc  s (2)  0.14+0.02  qlb  m‐4 g  b, f, jn, l, o, p, pc, r, s, t  l (12)  0.15+0.02  qlb  m‐5 g  b, h, jn, k, lc, lv, o, p, s, t  a (12)  0.19+0.02  qlb  m‐6  s  v (3)  0.15+0.00  qlv  m‐7  b, lv, sm  a (13)  0.32+0.03  qlv  m‐7  sm  a (3)  0.33+0.01  qlv  m‐8 g  g, jn, k, lv, r  l (10)  0.36+0.04  qlv  m‐8 g  r  l (2)  0.25+0.01  qlv  m‐9 g  g, jn, k, lv, r, w  a (28)  0.43+0.02  qlv  m‐9 g  r  a (2)  0.42+0.02  qlv  m‐10  lv  l (2)  0.29+0.07  qpp  m‐11  lv  a (22)  0.42+0.06  qpp  m‐12 g  lv  a (12)  0.55+0.05  qpp  m‐13  lv  o (1)  0.58+0.05  qlcb  m‐14 g  lv  l? (2)  0.81+0.04  qlcb  m‐15 g  jn  p (5)  0.64+0.07  correlations of matched samples for same fossils and same author(s) in appendix 3. color  qlb  qcd  qmc  qlv  qpp  qlcb  unused  unmatched  green  k1li  k3li          k2li  k2h  green  k1c  k3c  k5c  k7c      k6c    yellow  o9l  o10 l  o11l        o8l    yellow  o5a  o7a              yellow  o8a    o11a        o9a    yellow  o9s  o10 s              yellow  o1c, li      o2c,li  o3c,li  o4c,li      red  s1(r)l      s5(r)l  s2(r) l      s3(sm)l  red  s4(sm)a      s7(sm)a          red  ssl      s14l          red  s9a      s13a      s11 a  s6(r)a  red  s10l      s12l          red  s11a      s15a      s16 a    blue  m1l      m19l        m13o m15p  blue  m2a      m7a  m11a      m34s m6v  blue  m4l      m8l    m14l?    m8(r)l  blue  m5a      m9a  m12a      m7(sm)a  a-8 appendix 5. shoreline altitudes of lake cycles in main bonneville basin compared to coeval shorelines in cache valley. altitudes of samples for these lakes are uncorrected for isostatic rebound. main  bonneville  basin  cache  valley  bay  lake  cycle  location;  source  age  in ka  shoreline  altitude  location; source  age in ka  shoreline altitude  altitude  difference in  cache  valley  little  valley  point of mountain  scott and others,  1988  ~124  ~4954'  ~1510 m  hyde park cut wall  ~142.8  >4889'  >1490 m  < ‐65'  < ‐20 m  little valley  big cottonwood  canyon, scott and  others, 1983  ~175  ~4960'  ~1512 m  newton hill pit  ~144.3  ~169.4  >4865'  >1483 m  < ‐95'  < ‐29 m  little valley?  alpine under  promontory  geosol in little  valley pit  morrison,  1965b, 1966  uncertain,  probably  little  valley  ~4986'  ~1519 m'  millville eroded  delta between  bonneville  highstand and  provo delta;  includes muley hill  unknown;  might be  little valley  ~4975'  ~1516 m  ~ ‐11'  ~ ‐3 m  hansel  valley  west gully;  robison &  mccalpin, 1987  ~82  ~76  ~4400'  ~1341 m  newton hill pit  none  none at pit level  none  cutler dam  westside canal;  kaufman and  others, 1971  ~59.4  ~4396'  ~1340 m  newton hill pit  ~66.82  ~67.70  ~4733' [4757']  ~1443 m [1450 m]  + 337' [361']  + 103 m [110 m]  early  bonneville  oviatt, 2015 data  oviatt, 2020  curve  ~21  ~4954'  ~1510 m  muley hill, millville  ~20.98  ~5085'  ~1549 m  ~+131'  ~+40 m