2021 | 74/2 | 177–187 | 2 Figs. | 3 Tabs. | www.geologia-croatica.hr Journal of the Croatian Geological Survey and the Croatian Geological Society 1. INTRODUCTION The advancement of geochemical techniques, in particular the development of advanced mass spectrometry techniques, has allowed us to more accurately determine a number of stable iso- tope ratios and broaden the set of proxies for many geological and environmental processes. In addition to traditional stable iso- topes, such as C, O, and H, the discussion on isotopic signatures of different geological settings has thus extended to new ones, called non-traditional isotopes (e.g., Li, Mg, Ca, Cr, Mn, Fe, Cu, Zn, Se, Cl, etc.), which, unlike the traditional ones, cannot be analyzed by gas-source mass spectrometry (TENG et al., 2017). Although the idea of using an isotope signature in geological studies is not new (NÄGLER et al., 2000; BÖHM et al., 2006; CHAN et al., 2006; HIPPLER et al., 2006; FANTLE & TIP- PER, 2014; ETC.), as our knowledge of isotope fractionation in different environmental compartments grows, our ability to use them significantly increases. In this regard, sediments, as archives of environmental changes and conditions, play a crucial role. As major reservoirs of trace metals in the aquatic environment, their total content and isotopic fractionation have the potential to distinguish between the contributions of different phases to the sediment content, e.g. biogenic, authigenic, and lithogenic, as well as to provide infor- mation on redox conditions and diagenetic and weathering pro- cesses (CHAN et al., 2006; LITTLE et al., 2014; BLÄTTLER et al., 2015; LITTLE et al., 2017; AHM et al., 2018; CISCATO et al., 2018; BRUGGMANN et al., 2019; HU et el., 2019). This review summarizes recent findings on the application of non-traditional stable isotopes of: the alkali metal (Li), alka- line earth metals (Ca, Mg, and Sr) and the transition metals (Ni, Cr, and Cu) group in the interpretation of post-depositional pro- cesses and environmental conditions (i.e. redox conditions and Non-traditional stable isotope signatures in geological matrices as a tool for interpreting environmental changes – a review Željka Fiket*, Martina Furdek Turk, Maja Ivanić and Goran Kniewald Ruđer Bošković Institute, Division for Marine and Environmental Research, Bijenička cesta 54, 10000 Zagreb, Croatia; (*correspondence: zeljka.fiket@irb.hr) doi: 10.4154/gc.2021.12 Abstract The development of new analytical techniques enabled the precise determination of the expand- ed set of stable isotopes and provided new insight into existing geological issues. This review outlines recent studies of non-traditional isotope signatures in geological matrices, summarizing in one place, new data for the stable isotopes of Ca, Mg, Sr, Li, Ni, Cr, and Cu and their applica- tion in the interpretation of environmental processes. Although some, such as δ44Ca and δ26Mg, have previously been used to track changes in seawater chemistry throughout geological his- tory, recent studies report their application as geochemical proxies of post-depositional process- es. Similarly, isotopic signatures of strontium, previously used in radioactive isotope chronology, and δ7Li, used in tracing plate subduction and crust/mantle material cycling, found a new appli- cation in studies of weathering patterns. The use of δ53Cr and δ 65Cu isotope signatures, on the other hand, reflects their fractionation under different redox conditions, whereas δ60Ni, due to its adsorption and co-precipitation with sulfide species and Fe-Mn phases, is used in interpreting the contributions of different material sources. And while the isotopic signatures of all these ele- ments indicate certain environmental conditions and processes (e.g. post-depositional process- es, redox conditions, organic matter input, the contribution of sources, etc.), by combining them a more comprehensive insight into the investigated environment can be achieved. biological activity), as well as in deciphering the contributions of different sources. While the isotopic signatures of Li, Sr, Ca, and Mg in sediments as indicators of subduction and continental crust formation or as age and temperature palaeo-proxies have been widely discussed in the literature (NÄGLER et al., 2000; BÖHM et al., 2006; CHAN et al., 2006; HIPPLER et al., 2006; FAN- TLE & TIPPER, 2014; SAENGER & WANG, 2014; TENG, 2017; GUO et al., 2019), their use in the interpretation of post- depositional processes has only recently been more extensively investigated (FANTLE & HIGGINS, 2014; CHANDA & FAN- TLE, 2017; TENG et al., 2017; BRADBURY & TURCHYN, 2018; HU et al., 2019; LI et al., 2019). The same applies to the ap- plication of isotopic variability of the transition metals as a tool for solving specific problems in palaeo- and recent environmen- tal settings (GUEGUEN et al., 2016; BACONNAIS et al., 2018; LITTLE et al., 2018; BRUGGMANN et al., 2019). To date, re- search on the latter is extremely scarce and additional efforts need to be made to produce a more comprehensive database on the isotope variability of these elements in different matrices. Although the existing literature provides a valuable over- view of the current knowledge on non-traditional isotopes, it most often relates to only one, rarely several, isotopes together (CHAN et al., 2006; FANTLE & HIGGINS, 2014; LITTLE et al., 2014; BLÄTTLER et al., 2015; CHANDA & FANTLE, 2017; LITTLE et al., 2017; BACONNAIS et al., 2018; BRADBURY & TUR- CHYN, 2018; CISCATO et al., 2018; BRUGGMANN et al., 2019; HU et el., 2019; LI et al., 2019). However, the intertwining of different processes occurring simultaneously in an environ- ment necessitates the use of larger data sets that include more stable isotopes as well as their total element concentrations. The review presented here thus provides an overview of the latest re- search on the subject, summarizing in one place new data on non- Article history: Manuscript received November 11, 2020 Revised manuscript accepted March 29, 2021 Available online June 30, 2021 Keywords: non-traditional isotopes; Ca and Mg; Sr and Li; transition metals; geological matrices G eo lo gi a C ro at ic a 178 Geologia Croatica 74/2 traditional isotopes (Ca, Mg, Sr, Li, Ni, Cr, and Cu) in different geological environments and highlighting the need for a compre- hensive database that would allow us to properly argue their use in interpreting environmental changes through Earth’s history. The listed elements were selected as those that have found new applications with the development of spectrometry techniques or as representatives of a group of transition metals the isotopic composition of which is nowadays used in the interpretation of numerous processes in palaeo- and recent environments. 2. STABLE ISOTOPES AND NOTATION Nowadays, stable isotope determinations are generally per- formed on multi-collector inductively coupled plasma mass spec- trometry instruments (MC-ICP-MS; e.g. Nu instruments, UK or Neptune, Thermo-Fisher, Bremen, Germany) (e.g. BLÄTTLER et al., 2015; LITTLE et al., 2014; LITTLE et al., 2017; CISCATO et al., 2018; HU et al., 2019; etc.). Detailed guidelines and recom- mended terms for expressing stable isotope ratios, as well as a description of the assessment of international reference materials for isotope-ratio analysis, can be found elsewhere (COPLEN, 2011; BRAND et al., 2014). The following are the elements de- scribed in this review paper, their isotopes, and an explanation of associate isotope composition records. Magnesium has three stable isotopes, 24Mg, 25Mg and 26Mg, with the relative abundances of 78.99%, 10.00% and 11.01%, re- spectively, whereas there are six naturally occurring calcium iso- topes: 40Ca, 42Ca, 43Ca, 44Ca, 46Ca and 48Ca, with abundances of 96.941%, 0.647%, 0.135%, 2.086%, 0.004% and 0.187%, respec- tively. Magnesium stable isotope data are reported using the stand- ard per mil (‰) notation of δ26Mg, i.e., the per mil deviation of the measured 26Mg/24Mg ratios of the unknowns relative to those of the international δ‐zero reference material DSM3. Although the DSM3 reference material is no longer available and has now been replaced by the ERM-AE143, calibrated to the DSM3, re- sults are still reported vs. DSM3: δ26Mg = [(26Mg/24Mg)sample / (26Mg/24Mg)DSM3 – 1] x 1000 (1) For Ca, the isotopic composition is expressed in ‰ relative to the international standard reference material NIST SRM 915a or b, as δ44/40Ca or δ44Ca. δ44Ca = [(44Ca/40Ca)sample / (44Ca/40Ca)NIST SRM 915a/b – 1] x 1000 (2) Strontium is an alkali-earth metal with four stable naturally occurring isotopes; 84Sr, 86Sr, 87Sr, and 88Sr, whereby 87Sr is also partly radiogenic, being produced by beta-decay of the radionu- clide 87Rb with a half-life of 48.8 x109 years (SEMENISHCHEV et al., 2020). Contrary to the 87Sr/86Sr ratio that gradually in- creases over time, the proportion between 84Sr, 86Sr, and 88Sr is constant in nature, but can change in processes where natural fractionation of isotopes occurs (SEMENISHCHEV et al., 2020). For Sr, the stable isotopic composition is expressed in ‰ relative to the international standard reference material SRM987 (NIST): δ88Sr = [(88Sr/86Srsample) / (88Sr/86SrSRM987) − 1] x 1000 (3) Unlike the δ88Sr ratio, expressed in per mil, the relative pro- portion of radiogenic to stable Sr isotopes, i.e. 87Sr/86Sr isotopic ratio, has commonly been reported in terms of a four- to six-place decimal notation instead of δ-notation. Lithium was one of the few elements produced during the Big Bang nucleosynthesis. It is composed of two stable isotopes 6Li and 7Li, with the heavier one being far more abundant (92.41%). Its isotopic composition is expressed in ‰ relative to the international standard reference material NIST L-SVEC or IRMM-016 (LI et al., 2010) as δ7Li values: δ7Li = [(7Li/6Li)sample / (7Li/6Li)standard – 1] x 1000 (4) Naturally occurring nickel is composed of five stable iso- topes: 58Ni, 60Ni, 61Ni, 62Ni, and 64Ni, with 58Ni being the most abundant (68.077% natural abundance), and 26 radioisotopes. Its isotopic composition is expressed in ‰ relative to the interna- tional standard reference material NIST SRM 986 as δ60Ni val- ues: δ60Ni = [(60Ni/58Ni)sample / (60Ni/58Ni)SRM 986 – 1] x 1000 (5) Naturally occurring chromium is composed of four stable isotopes: 50Cr, 52Cr, 53Cr, and 54Cr, with 52Cr being the most abun- dant (83.789% natural abundance). Its isotopic composition is ex- pressed in ‰ relative to the international standard reference ma- terial NIST SRM 979 as δ53Cr values: δ53Cr = [(53Cr/52Cr)sample / (53Cr/52Cr)SRM 979 – 1] x 1000 (6) Copper has two stable isotopes, 63Cu (69.17%) and 65Cu (30.83%), along with 27 radioisotopes. Natural mass-dependent variations in 65Cu/63Cu span 15‰, and are expressed relative to the NIST SRM 976 standard as δ65Cu values: δ65Cu = [(65Cu/63Cu)sample / (65Cu/63Cu)SRM 976 – 1] x 1000 (7) 3. NON-TRADITIONAL ISOTOPES AS GEOCHEMICAL PROXIES OF POST-DEPOSITIONAL PROCESSES 3.1. Weathering The radiogenic Sr isotope ratio (87Sr/86Sr) has long been known as a reliable tool in geochronological studies as well as in prove- nance investigation. Namely, the radiogenic 87Sr/86Sr ratio gene- rally varies with age and Rb/Sr ratios result in the 87Sr/86Sr ratio being an excellent tracer for the source(s) of Sr (ANDREWS et al., 2016; SEMENISHCHEV et al., 2020) (Table 1). However, with recent advances in mass-spectrometry, small variations in stable Sr isotope abundances (reported as δ88Sr) can be accurately quantified (FIETZKE & EISENHAUER, 2006; OHNO & HI- RATA, 2007) and are increasingly used to gain further insight into the Sr cycle and its primary controls (ANDREWS et al., 2016). So far, the fractionation of the stable isotopes of Sr has been documented in terrestrial, marine, and biological processes; wherein the lighter δ86Sr isotopes are preferentially incorporated into biogenic carbonates (+0.14‰ to +0.27‰, KRABBENHÖF et al., 2010), while inorganic Ca-carbonates and seawater are en- riched in the heavy Sr isotope and therefore display more positive δ88Sr values (+0.25‰ to +0.37‰ in carbonates (HALICZ et al., 2008); 0.310(8)‰ as an average value for seawater (KRABBEN- HÖF et al., 2010); 0.407 ± 0.012‰ in seawater from the Pacific and Atlantic oceans (WAKAKI et al., 2017); or 0.35 ± 0.06‰ for G eologia C roatica 179Fiket et al.: Non-traditional stable isotope signatures in geological matrices as a tool for interpreting environmental changes – a review the Mediterranean Sea (HALICZ et al., 2008; Table 1), analogous to Ca isotopes. Still, the Sr isotopic fractionation in terrestrial and marine systems is relatively small, with δ88Sr values varying by ∼2.5‰ (FIETZKE & EISENHAUER, 2006; OHNO & HIRATA, 2007; KRABBENHÖF et al., 2010; ANDREWS et al., 2016; WAKAKI et al., 2017; SEMENISHCHEV et al., 2020). Compared to other elements and isotopes, Sr is unique in that its isotopic records contain information on both the source of Sr (87Sr/86Sr) and the mass-fractionating processes that it has wit- nessed (δ88Sr). The combined application of these data can thus lead to a more comprehensive understanding of the weathering cycle of Sr. The latter was also suggested by MÍKOVÁ (2012) in Table 1. Strontium and lithium isotope composition in different types of samples. Isotope δ Sample type Reference δ88Sr +0.310(8)‰ seawater, average KRABBENHÖF et al., 2010 +0.407‰ ± 0.012‰ Pacific and Atlantic seawater WAKAKI et al., 2017 +0.39‰ ± 0.02‰ deep Pacific seawater SCHER et al., 2013 +0.14‰ to +0.27‰ marine biogenic carbonates KRABBENHÖF et al., 2010 +0.27‰ to +0.37‰ marine inorganic carbonates HALICZ et al., 2008 +0.35‰ ± 0.06‰ Mediterranean Sea HALICZ et al., 2008 −0.17‰ ± 0.06‰ terra rossa soil and speleothem calcite HALICZ et al., 2008 +0.26‰ ± 0.1‰ aragonite HALICZ et al., 2008 +0.22‰ ± 0.07‰ corals HALICZ et al., 2008 +0.35‰ ± 0.06‰ seawater HALICZ et al., 2008 -0.14‰ to -0.20‰ terrestrial carbonates HALICZ et al., 2008 +0.253‰ to 0.361‰ hydrothermal fluids KRABBENHÖF et al., 2010 +0.28‰ ± 0.09‰ gneiss DE SOUZA et al., 2010 +0.15‰ to +0.31‰ granite DE SOUZA et al., 2010 +0.29‰ to +0.31‰ soil DE SOUZA et al., 2010 +0.24‰ to +0.42‰ rivers KRABBENHÖF et al., 2010 +0.27‰ to +0.33‰ river sediments ANDREWS et al., 2016 ~0.267‰ carbonate fraction of soils and sediments ANDREWS et al., 2016 ~0.263‰ silicate fraction of soils and sediments ANDREWS et al., 2016 +0.223‰ to +0.369‰ soil exchangeable fraction ANDREWS et al., 2016 -0.11‰ to +0.42‰ plants OESER & BLANCKENBURG, 2020 δ7Li +31‰ ocean seawater PENNISTON-DORLAND et al., 2017 +1‰ to +44‰ river water PENNISTON-DORLAND et al., 2017 +17‰ to +35‰ lake water PENNISTON-DORLAND et al., 2017 +6‰ to +29‰ groundwater PENNISTON-DORLAND et al., 2017 +5‰ to +11‰ hydrothermal fluids PENNISTON-DORLAND et al., 2017 -3‰ to +26‰ geothermal waters PENNISTON-DORLAND et al., 2017 31‰ seawater CHAN et al., 2006 -2‰ to 5‰ clays CHAN et al., 2006 -4.3‰ to +14.5‰ marine sediments CHAN et al., 2006 -1.5‰ to +5% clay rich detrital sediments CHAN et al., 2006 +0.5‰ to +3.6‰ Laki basalts (Hawaii) PISTINER & HENDERSON, 2003 +2.2‰ to +4.7‰ Hawaii soil PISTINER & HENDERSON, 2003 +9.8‰ rainwater (unfiltered) PISTINER & HENDERSON, 2003 +18.1‰ rainwater (filtered) PISTINER & HENDERSON, 2003 +2.4‰ to +4.6‰ basalt PISTINER & HENDERSON, 2003 +8.8‰ to +9.1‰ granite PISTINER & HENDERSON, 2003 +3.1‰ to +12.0‰ Sao Miguel rivers bedload POGGE VON STRANDMANN et al., 2010 +3.5‰ to +9.5‰ Sao Miguel rivers suspended load POGGE VON STRANDMANN et al., 2010 +18.2‰ to +36.2‰ Sao Miguel rivers dissolved load POGGE VON STRANDMANN et al., 2010 +5.9‰ to +9.8‰ Sao Miguel rivers dissolved load in rivers influenced by hydrothermal input POGGE VON STRANDMANN et al., 2010 +6.9‰ Sao Miguel hot spring POGGE VON STRANDMANN et al., 2010 +32.8‰ Sao Miguel rivers rainwater POGGE VON STRANDMANN et al., 2010 +2.8‰ to +8.9‰ Iceland rivers bedload POGGE VON STRANDMANN et al., 2006 −1.3‰ to 7.5‰ Iceland rivers suspended load POGGE VON STRANDMANN et al., 2006 +17.0‰ to +43.7‰ Iceland rivers dissolved load POGGE VON STRANDMANN et al., 2006 +10.9‰ Iceland hot spring POGGE VON STRANDMANN et al., 2006 +22.8‰ Hraunfossar, Iceland groundwater POGGE VON STRANDMANN et al., 2006 +33.3‰ Ice (Iceland) POGGE VON STRANDMANN et al., 2006 + 0.6‰ ± 0.6‰ continental crust SAUZÉAT et al., 2015 0‰ ± 2‰ continental crust TENG et al., 2004 +6‰ to +32‰ carbonates POGGE VON STRANDMANN et al., 2019 6.1‰ ± 1.3‰ seawater calcite POGGE VON STRANDMANN et al., 2019 9.6‰ ± 0.6‰ seawater aragonite POGGE VON STRANDMANN et al., 2019 G eo lo gi a C ro at ic a 180 Geologia Croatica 74/2 her review on tracing weathering processes and flow pathways in surface and ground waters in Antarctica and terrains with a re- treating glacier using strontium isotopic composition. A similar approach was also adopted by ANDREWS et al. (2016) using ra- diogenic Sr isotope ratios (87Sr/86Sr), stable Sr isotope ratios (δ88Sr), and major ion data to identify river cation sources and their biogeochemical cycling within the Cleddau and Hollyford River catchments in the Milford Sound region of Fiordland, New Zealand. Their results suggested that the soil water δ88Sr values are ∼0.30‰ higher than the bedrock δ88Sr values, while mixing calculations showed that the plant-fractionated soil water pool contributes ∼27% of the riverine Sr. Also, by combining the δ88Sr values and Ca/Sr and 87Sr/86Sr ratios, three major contributions to the Sr riverine budget in the Milford Sound region were iden- tified; i) silicate (with δ88Sr amounting to ~0.263 ‰); ii) carbona te weathering (with δ88Sr amounting to ~0.264‰); iii) soil water input (with δ88Sr between 0.223‰ and 0.369‰). Similarly, con- tinental waters were found to be enriched with 88Sr compared to the rocks in their drainage basins (e.g. KRABBENHÖF et al., 2010; MOYNIER et al., 2010). To explain this phenomenon, SHALEV et al. (2017) suggested that during precipitation of con- tinental carbonates (i.e. carbonates precipitated from the surface, pedogenic, or groundwater) a fractionation of strontium isotopes occurs that contributes to the enrichment of 88Sr in rivers. How- ever, this would also require that a surprisingly large proportion (∼40%) of the originally weathered Sr to co-precipitate with con- tinental carbonates, suggesting that some other mechanisms, such as Sr uptake in plants and/or incongruent weathering of silicates, play a significant role in the Sr cycle (SHALEV et al., 2017). OESER & BLACKENBURG (2020) concluded that the release of Sr during weathering is isotopically congruent and that, despite a strong biological fractionation (-0.40‰ to +0.34‰), the plant production – degradation cycle is isotopically neutral. The poten- tial fractionation in the bioavailable Sr pool was explained by the loss of Sr from the system by organic solids enriched in the light 86Sr isotope. The understanding of stable Sr isotope fractionation during weathering is however, still very limited (HALICZ et al., 2008; DE SOUZA et al., 2010; SHALEV ET AL., 2017) (Table 1), and necessitates further investigation. In addition to Sr, Li isotopes have also been reported as con- tributing to a better understanding of weathering processes ( Fig. 1). Namely, the significant fractionation of Li isotopes occurs during silicate weathering, from over 26‰ and up to 35‰ in river and surface environments (PISTINER & HENDERSON, 2003), with the preferential incorporation of 6Li in neo-formed clay min- erals (e.g. PISTINER & HENDERSON, 2003; POGGE VON STRANDMANN et al., 2010; SAUZÉAT et al., 2015), and the mobilization of heavier 7Li into solution. On continents, the δ7Li value is thus controlled by the leaching rates of silicate source rocks and the amounts of neo-formed clays. Consequently, areas with weathering-limited regimes (e.g. mountainous or glaciated regions), as well as those with transport-limited regimes (e.g. areas with thick soil sequences), are usually characterized by high δ7Li signatures, except when underlain by shales. In rivers, the δ7Li isotopic signature is driven by the lithology of the drainage area, where the primary lithology determines the secondary mine- ralogy (e.g. POGGE VON STRANDMANN et al., 2006; 2010; SAUZÉAT et al., 2015), the corresponding weathering intensity, the concentration of river suspended load (POGGE VON STRANDMANN et al., 2006;), as well as the subsurface resi- dence time of infiltrating meteoric waters (PISTINER & HEN- DERSON, 2003; SAUZÉAT et al., 2015) (Table 1). All the latter conditions show significantly greater variability of δ7Li in rivers. For example, in river waters the δ7Li is +1‰ to +44‰ (PENNIS- TON-DORLAND et al., 2017), -1.3‰ to +9.5‰ in the suspended load (POGGE VON STRANDMANN et al., 2006; 2010), and 17.0‰ to 43.7‰ in the dissolved load (POGGE VON STRAND- MANN et al., 2006; 2010), while in silicate rocks the values are ~0‰ in the continental crust (TENG et al., 2004; SAUZÉAT et al., 2015) and ~3–5‰ in basalts (PISTINER & HENDERSON, 2003; SAUZÉAT et al., 2015). Marine sediments are, on the other hand, enriched in heavier 7Li, with δ7Li values ranging from 4.3‰ to +14.5‰ (CHAN et al., 2006), while carbonates display an even broader range and generally correspond to δ7Li values between +6‰ and +32‰ (POGGE VON STRANDMANN et Figure 1. Schematic representation of the distribution of δ88Sr and δ7Li in different environmental compartments. G eologia C roatica 181Fiket et al.: Non-traditional stable isotope signatures in geological matrices as a tool for interpreting environmental changes – a review al., 2019). In the ocean, lithium acts as a conservative element with a residence time of about one million years and a δ7Li iso- topic signature of +31‰ (PENNISTON-DORLAND et al., 2017). The steady-state is maintained by river inputs during con- tinental weathering (+23‰), inputs by groundwater (+15‰) and hydrothermal fluids (+7‰), and the preferential removal of 6Li from seawater by clay minerals (PENNISTON-DORLAND et al., 2017; and references therein); the combination of which also explains the heavier δ7Li signature compared to its inputs. The prevalent incorporation of Li into silicate minerals com- pared to other mineral phases results in the Li isotopic signature being a promising proxy for evaluating changes related to the continental silicate weathering patterns (PISTINER & HEN- DERSON, 2003; CHAN et al., 2006). In contrast to radiogenic isotopic ratios used to monitor chemical weathering, such as 87Sr/86Sr, the Li isotopic composition of river water is not highly sensitive to those of the bedrock, as it mainly reflects the weat- hering intensity and formation of clay minerals. Thus, by combin- ing the Li and Sr isotopic signatures of different geological envi- ronments (Fig. 1) with mineralogical data, a more complete insight into the weathering processes as well as into the cycling of these elements in the environment can be achieved. 3.2. Diagenesis After deposition, sediments are subjected to diagenetic alteration, which can affect their geochemical fingerprint. The extent to which the primary sediment was altered can provide crucial in- formation needed for the reconstruction of environmental condi- tions under which it was deposited (BLÄTTLER et al., 2015; CHANDA & FANTLE, 2017; HIGGINS et al., 2018). To success- fully tackle this topic, the following authors used a multi-parame- ter concept. FANTLE & HIGGINS (2014) combined the traditional ap- proach, using elemental ratios and stable isotopes of C and O, with the non-traditional, i.e. isotopic composition of Ca and Mg, to assess the extent of diagenetic alteration of geochemical prox- ies in shallow marine carbonates from ODP (Ocean Drilling Pro- gram) Site 1196A. The results of their study suggest that the vari- ability in δ44Ca (from 0.60‰ to 1.131‰) and δ26Mg (from -3.91‰ to -2.59‰) values, along with the variations in other geochemical proxies (C and O isotopes), can be used to elucidate diagenetic alterations, namely limestone diagenesis, and dolomitization. These authors also highlighted the importance of using several parameters when deciphering diagenetic changes in the sedimen- tary record. A similar multi-proxy approach has been implemented by HIGGINS et al. (2018) in their study of early-marine diagenetic alterations and their effect on the geochemical proxies of shallow- water carbonate sediments. The authors combined measurements of the isotopic composition of Ca and Mg with other geochemical proxies (δ13C, δ18O, Sr/Ca, Mg/Ca) and the mineralogical compo- sition in sediments and pore fluids at ten sites in the Bahamas and the Eucla Shelf. Their study showed that variations in the isotopic composition of Ca (-1.56‰ to -0.35‰) and Mg (-3.46‰ to -2.19‰) in the stratigraphic record were influenced by changes in the car- bonate mineralogy and the type of diagenetic alteration, fluid- buffered vs. sediment-buffered diagenesis (Table 2). In the majori ty of the Bahamian sediments, fluid-buffered diagenetic processes combined with the transformation of aragonite and high-Mg calcite to low-Mg calcite and dolomite resulted in the coherent stratigraphic variability of δ44Ca values and carbonate- specific geochemical proxies, such as δ13C, δ18O, and Sr/Ca as well as in the trace element records. Based on previous findings (BLÄTTLER et al., 2015; CHANDA & FANTLE, 2017), AHM et al. (2018) presented a nu- merical model that estimated the extent and type (fluid- vs. sedi- ment-buffered) of early-diagenetic processes in carbonate sedi- ments and their influence on other carbonate-specific geochemical proxies. Through tracking the elemental and isotopic composition of carbonate-bound geochemical proxies (Ca, Mg, C, O, Sr/Ca), the model proposed by these authors was able to estimate the extent of diagenetic alteration, describe the diagenetic fluid and provide more insight into the variation of geochemical proxies in ancient carbonate deposits. Recently, LI et al. (2019) combined the δ26Mg isotopic sig- nature of dolomite (ranging from -2.28‰ to -1.78‰), with trace element analysis, the isotopic composition of C, and O, and mine- ralogical analysis to determine the dolomite formation depth in massive Cretaceous dolostones. The authors highlighted the im- portance of the dolomitization depth and mechanism on the Mg-C isotope response in the sedimentary record. Studies of modern dolomitization processes (FANTLE & HIGGINS, 2014; BLÄTTLER et al., 2015; CHANDA & FAN- TLE, 2017) show that during early diagenesis δ26Mg values in dolomites are controlled by porewater chemistry. Thereby, the evolutionary patterns of porewater chemistry differ in different types of early diagenetic systems due to changes in the hydro- logical conditions and result in stratigraphic variations in the composition of the Mg isotope in carbonates (AHM et al., 2018; HIGGINS et al., 2018). On the other hand, Mesozoic-Palaeozoic dolostones formed by massive dolomitization have been reported as potential archives of coeval seawater Mg isotope signals (HU et al., 2019). HU et al. (2019) studied the relationships between post-depositional processes and Mg isotope variabilities in mas- sive dolomites by complementing magnesium isotope analyses with the analysis of C-O isotopic composition, rare earth element (REE) concentrations and the petrographic features of dolo stones. Their results indicated that different types of syndiagenetic dolo- mites deposited in different sedimentary settings have uniform δ26Mg values of -2.06‰ ± 0.20‰, implying that the chemistry of the porewater was buffered by coeval seawater during dolomiti- zation, irrespective of sedimentary settings. Furthermore, their results suggested that post-depositional processes, including diagenetic and hydrothermal alterations, did not induce notable changes in the Mg isotope signals in the dolomite. 3.3. Authigenic mineral formation Authigenic formation of clay minerals is another sedimentary component that significantly affects the isotope fractionation in the pore fluids. BRADBURY & TURCHYN (2018) used calcium isotopes to evaluate the influence of sediment type on carbonate precipi- tation, dissolution, and recrystallization. The analysis was per- formed on sediments from ODP sites 1081 and 1086, off the coast of West Africa, two geographically close sites with a different share of carbonates and organic matter. The authors combined Ca isotope composition of sediments (δ44Ca = –0.42‰ – –0.21‰, reported relative to Bulk Silicate Earth) with Sr and Ca concen- trations in pore fluids (δ44Ca = –0.08 – +1.04‰) to model the rates of carbonate dissolution and precipitation. Their results suggested the occurrence of Ca isotope fractionation during precipitation of carbonate. G eo lo gi a C ro at ic a 182 Geologia Croatica 74/2 BLÄTTLER et al. (2015) on the other hand, used Mg and Ca isotopes as geochemical tracers for authigenic carbonate forma- tion in marine sediments of Miocene age. Similarly to FANTLE & HIGGINS (2014), the authors combined measurements of δ26Mg and δ44Ca (delta values of Ca are reported relative to mo- dern seawater), with those of δ13C and δ18O, performing a multi- isotope study on authigenic dolomites occurring as beds and nodu- les within the stratigraphic sequence of the Monterey Formation, offshore California. Their data suggested that the obtained iso- topic signature of Mg and Ca point to dolomite precipitation and carbonate recrystallization, respectively. There, the variability of δ26Mg (from -2.86‰ to -0.52‰) and δ44Ca values (from -1.10‰ to -0.35‰) with stratigraphic depth, and anti-correlation of Mg and Ca isotopes, was related to early-diagenetic processes in the sedimentary pore fluid. The large variation of δ13C values in do- lomite on the other hand, was attributed to its formation in zones of bacterial sulfate reduction or methanogenesis. KIMMIG & HOLMDEN (2017) investigated the potential of Mg isotopes, combined with other geochemical proxies (Sr/Ca and Ca/Mg ratios, δ 13C, mineralogy) to identify changes in car- bonate polymorph mineralogy in a carbonate mud deposit marked by the Hirnatian glaciation. Their results showed that in the in- vestigated section, during the glaciation event, a 2–3‰ increase in sedimentary δ26Mg values occurred. The changes in the abun- dance of dolomite in the host limestone were considered respon- sible for the observed shifts in δ26Mg, which was also in agree- ment with other geochemical proxies and indicators of sea-level changes. According to the trends observed for the δ26Mg and Ca/Mg ratios, changes in carbonate mineralogy (e.g. variations in the abundance of aragonite) induced the following changes in the δ26Mg isotopic composition of limestone: -4.45 ± 0.82‰ before glaciation, -2.54 ± 0.23‰ during glaciation and -3.05 ± 0.27‰ after glaciation. This exemplifies the potential of a multi-proxy approach in the identification of changes in carbonate poly- morphs in sedimentary records. 3.4. Adsorption-desorption processes While carbonates preferentially incorporate light Mg and Ca iso- topes (FANTLE & HIGGINS, 2014; BLÄTTLER et al., 2015; KIMMIG & HOLMDEN, 2017; HIGGINS et al., 2018), limestone diagenesis and dolomitization cause their shift towards lower or higher values, respectively (Table 2). In contrast to carbonates, clay minerals display heavier δ26Mg values (-1.82‰ to +0.18‰; WIMPENNY et al., 2014), but also to some extent variable ones, depending on the share of the exchangeable and the structurally incorporated Mg. Similar to magnesium, the calcium isotope composition of pore waters was also found to be influenced by the adsorption/desorption processes of Ca2+ from the surfaces of clay minerals (OCKERT et al., 2013). OCKERT et al. (2013) in- vestigated the influence of cation exchange between different clay minerals (montmorillonite, illite, and kaolinite) and marine sedi- ment, with Ca in artificial seawater. The authors found that in the marine porewater environment the light Ca preferentially ad- sorbs onto clay minerals with the fractionation between the ad- sorbed and the dissolved Ca2+ and δ44Ca varying between +0.09‰ and -2.76‰. Their results also suggested a mineral spe- cific fractionation of Ca, where kaolinite (-1.2‰ and -2.76‰ dur- ing adsorption and desorption, respectively) and illite (-0.82‰ and -1.2‰ during adsorption and desorption, respectively), showed more significant fractionation compared to montmoril- lonite. Also, the presence of ammonium, originating from or- ganic matter decomposition, was considered to cause desorption of Ca2+ from the clay mineral surfaces into the pore water. This topic was further studied by BRAZIER et al. (2019) who experi- mentally showed that the intensity of Ca isotopic fractionation during adsorption onto clay minerals is controlled by their struc- tural characteristics, including layer charge, specific surface area, and interlayer space. The results obtained by these authors high- lighted the importance of adsorption-desorption processes in- volving mineral particles in the assessment of Ca isotopic signa- tures in natural environments. The more so as, in addition to carbonates, silicate minerals are the second dominant source of Ca and Mg in rivers, and therefore seawater, and discerning the key controls of Ca and Mg isotopic composition during the al- teration of primary silicates and formation of alteration phases such as clay minerals can help in understanding how these ele- ments behave during weathering. 4. NON-TRADITIONAL ISOTOPES AS INDICATORS OF ENVIRONMENTAL CONDITIONS 4.1. Redox conditions The chromium isotope system functions as an atmospheric redox proxy because oxidative weathering of crustal Cr(III)-bearing minerals results in the release of 53Cr-enriched mobile Cr(VI) to the solution (GILLEAUDEAU et al., 2016). Cr(VI) is then car- ried to the oceans via rivers, thus imparting a positively fractiona- ted δ53Cr signal to modern seawater (~0.5 ± 0.1‰) (BONNAND et al., 2013) compared to crustal values (–0.123 ± 0.102 ‰) (SCHO- ENBERG et al., 2008). In marine environments, chromium has a tendency to co-precipitate with different mineral phases and readily accumulates in sediments in reducing conditions (e.g. GILLEAUDEAU et al., 2016; Fig. 2), which makes δ53Cr a valuable redox proxy not only in palaeo-environmental studi es but also in recent marine settings (BONNAND et al., 2013; SCHEIDERICH et al., 2015; GILLEAUDEAU et al., 2016). According to SCHEIDERICH et al. (2015), modern surface seawater is heterogeneous in δ53Cr, with isotopic heterogeneity confined to the surface water (from 0.41‰ to 1.51‰ in the Atlantic ocean, and from 0.91‰ to 1.43‰ in the Pacific ocean; SCHEIDERICH et al., 2015), while the deep-water masses in the Atlantic and Pacific displayed similar δ53Cr values (~0.5‰, SCHEIDERICH et al., 2015); Table 3). In contrast, the study of BRUGGMANN et al. (2019) on chromium isotope cycling in the water column and sediments of the Peruvian continental margin, a modern oxygen minimum zone (OMZ) in an open marine set- ting, yielded somewhat different values. Their results show δ53Cr ranging from 0.02 ± 0.16‰ to 0.59 ± 0.11‰ and from 0.31 ± 0.07‰ to 0.92 ± 0.12‰ in seawater and sediments, respectively. The sediments deposited in permanently anoxic waters showed δ53Cr at 0.77 ± 0.19‰, while those deposited in oxic bottom waters were characterized by δ53Cr at 0.46 ± 0.19‰, suggesting that sediment Cr concentrations and δ53Cr values are influenced by water column redox (e.g. reductive dissolution and transport of Fe-oxides) and/or early diagenetic processes (e.g. redistribution of Cr during phosphogenesis). Although both mentioned studies emphasize the importance of redox conditions for Cr fractiona- tion in marine environments, little is known about redox cycling of Cr within the water column, changes across the sediment- - water interface, and the exact mechanisms by which Cr is incor- porated into marine sediments, which remain poorly constrained ( SCHEIDERICH et al., 2015; BRUGGMANN et al., 2019). All of this is important as most of the studies conducted have neglected the internal fractionation processes in the marine Cr cycle, such G eologia C roatica 183Fiket et al.: Non-traditional stable isotope signatures in geological matrices as a tool for interpreting environmental changes – a review as the export of lighter Cr(III) with organic matter into the sedi- ment (SEMENIUK et al., 2016). The isotopic composition of Ni can also be used (Fig. 2) in elucidating the redox processes in sediments. Namely, its co-pre- cipitation with Fe–Mn oxides represents an important output from the dissolved pool in an oxygenated environment (e.g., GALL et al., 2013; CAMERON & VANCE, 2014; TENG et al., 2017; Table 3). In reducing conditions, however, Ni shows a prefe- Figure 2. Schematic representation of the distribution of δ53Cr, δ60Ni and δ65Cu in different environmental compartments. Table 2. Magnesium and calcium isotope composition in different types of samples. Isotope δ Sample type Reference δ26Mg -2.86‰ to -0.52‰ bulk dolomite BLÄTTLER et al., 2015 –2.54‰ ± 0.23‰ calcite KIMMIG & HOLMDEN, 2017 +0.18‰ illite WIMPENNY et al., 2014 -0.30‰ montmorillonite WIMPENNY et al., 2014 -1.82‰ kaolinite WIMPENNY et al., 2014 -3.46‰ to -2.19‰ Great Bahama Bank HIGGINS et al., 2018 -3.02‰ to -2.64‰ Little Bahama Bank HIGGINS et al., 2018 -3.13‰ to -3.07‰ bank-top sediment HIGGINS et al., 2018 -3.26‰ to -2.62‰ Eucla Shelf HIGGINS et al., 2018 -0.72‰ to -0.39‰ pore water CHANDA & FANTLE, 2017 -5.00‰ to -2.23‰ bulk carbonates CHANDA & FANTLE, 2017 -3.91‰ to -2.59‰ bulk carbonate FANTLE & HIGGINS, 2014 -3.60‰ ± 0.25‰ limestone, average FANTLE & HIGGINS, 2014 -2.68‰ ± 0.07‰ dolostone, average FANTLE & HIGGINS, 2014 -2.28‰ to -1.78‰ dolostone LI et al., 2019 -2.29‰ ± 0.10‰ dolostone GALY et al., 2002 -2.06‰ ± 0.20‰ syndiagenetic dolomite HU et al., 2019 -0.22‰ ± 0.10‰ UCC LI et al., 2010 -0.63‰ to 0.64‰ Iceland rivers POGGE VON STRANDMANN et al., 2008 0.85‰ hydrothermal springs POGGE VON STRANDMANN et al., 2008 -0.83‰ ice POGGE VON STRANDMANN et al., 2008 δ44Ca -1.10‰ to -0.35‰ bulk dolomite BLÄTTLER et al., 2015 -1.56‰ to -0.40‰ Great Bahama Bank HIGGINS et al., 2018 -0.68‰ to -0.35‰ Little Bahama Bank HIGGINS et al., 2018 -1.49‰ to -1.20‰ bank-top sediment HIGGINS et al., 2018 -1.24‰ to -0.88‰ Eucla Shelf HIGGINS et al., 2018 0.60‰ to 1.131‰ bulk carbonate FANTLE & HIGGINS, 2014 0.97‰ ± 0.24‰ limestone, average FANTLE & HIGGINS, 2014 1.03‰ ± 0.15‰ dolostone, average FANTLE & HIGGINS, 2014 -0.08‰ to 1.04‰ pore fluid Site 1081 BRADBURY & TURCHYN, 2018 -0.42‰ to -0.21‰ sediment Site 1081 BRADBURY & TURCHYN, 2018 0.11‰ to 0.78‰ pore fluid Site 1086 BRADBURY & TURCHYN, 2018 -0.42‰ to -0.25‰ sediment Site 1086 BRADBURY & TURCHYN, 2018 G eo lo gi a C ro at ic a 184 Geologia Croatica 74/2 rence towards dissolved sulfide (e.g., CISCATO et al., 2018) and is either scavenged by reactive sulfide species or sulfide precipi- tation. While the Fe-Mn oxides show preference to isotopically heavier isotopes (+0.9‰ to +2.5‰, GALL et al., 2013), sulfide phases prefer the isotopically lighter Ni (+1.3 ± 0.4‰, CISCATO et al., 2018). 4.2. Organic matter In addition to co-precipitation with Fe-Mn-oxyhydroxides or sulfide phases, depending on the redox state, two other important processes control the marine geochemistry of Ni and its isotope fractionation; biological cycling via cell uptake and the burial of organics (CAMERON et al., 2009; CAMERON & VANCE, 2014). However, sedimentary outputs that control the heavy δ60Ni content of the ocean are at odds with the lighter Ni stable isotope composition of the inputs, by 1.3‰ to 1.7‰ (CISCATO et al., 2018), a feature common to many other transition metals (e.g. LITTLE et al., 2014) (Fig. 2). To find a possible explanation for this imbalance, CISCATO et al. (2018) investigated upwelling margin sediments and found that the organic-rich sediments be- neath upwelling zones are also an important output flux of Ni from the oceans. However, they found that the δ60Ni in sediments (+0.94‰ to +1.33‰) was almost identical to the δ60Ni of modern seawater (+1.19‰ to +1.47‰) (CAMERON & VANCE, 2014). De- spite this similarity, the approach they adopted, by analyzing the Ni isotopic signature of both fractions, the organic-sulfides (with δ60Ni from +0.86‰ to +1.83‰) and the HF-digestible fraction (with δ60Ni from +0.86‰ and +1.37‰), has the potential to quan- tify isotope fractionations associated with biological uptake as well as to record the δ60Ni of the contemporary seawater. Even though Cr is not considered bio-essential, organic mat- ter can also significantly influence the Cr cycle (e.g. SEMENIUK Table 3. Chromium, nickel and copper isotope composition in different types of samples. Isotope δ Sample type Reference δ53Cr −0.26‰ to −0.04‰ continental crust SCHOENBERG et al., 2008 ~0.5 ± 0.1‰ modern seawater BONNAND et al., 2013 +0.412‰ to +0.664‰ Argentine seawater BONNAND et al., 2013 +0.65‰ to +0.76‰ Bahamian and Yucatan ooid BONNAND et al., 2013 +0.41‰ to +1.51‰ Atlantic Ocean water SCHEIDERICH et al., 2015 +0.61‰ to +1.43‰ Pacific Ocean water SCHEIDERICH et al., 2015 +0.99‰ to +1.55‰ Arctic seawater SCHEIDERICH et al., 2015 +0.02‰ to +0.59‰ seawater BRUGGMANN et al., 2019 +0.31‰ to +0.92‰ recent sediments BRUGGMANN et al., 2019 0.46 ± 0.19‰ sediments deposited in oxic conditions BRUGGMANN et al., 2019 0.77 ± 0.19‰ sediments deposited in anoxic conditions BRUGGMANN et al., 2019 δ60Ni +0.9‰ to +2.5‰ Fe–Mn crusts GALL et al., 2013 +1.79‰ ± 0.21‰ North Pacific Fe-Mn crust LI et al., 2019 +1.73‰ ± 0.21‰ South Pacific Fe-Mn crust LI et al., 2019 +1.3‰ ± 0.7 ‰ Atlantic Ocean Fe-Mn crust GALL et al., 2013 +1.7‰ ± 0.8‰ Pacific Ocean Fe-Mn crust GALL et al., 2013 +1.6‰ ± 0.3‰ Indian Ocean Fe-Mn crust GALL et al., 2013 +0.94‰ to +1.33‰ sediment CISCATO et al., 2018 +0.86‰ to +1.83‰ organic matter and associated pyrite fraction CISCATO et al., 2018 +0.86‰ and +1.37‰ HF-dissolvable sediment fraction CISCATO et al., 2018 +1.3 ± 0.4‰ sulfide phases CISCATO et al., 2018 +0.15‰ ± 0.24‰ terrestrial samples (mantle and crust) CAMERON et al., 2009 + 1.19‰ to 1.47‰ deep ocean water CAMERON and VANCE, 2014 +1.44‰ ± 0.15‰ seawater CAMERON and VANCE, 2014 +0.29‰ to +1.34‰ river water CAMERON and VANCE, 2014 +0.14‰ ± 0.23‰ river sediments CAMERON and VANCE, 2014 δ65Cu +0.66‰ ± 0.07‰ South Atlantic deep water LITTLE et al., 2018 +0.65‰ ± 0.07‰ North Atlantic deep water BOYLE et al., 2012 +0.65‰ ± 0.08‰ Indian and Pacific Ocean deep water TAKANO et al., 2014 +0.70‰ ± 0.11‰ Tasman Sea deep water THOMPSON et al., 2014 +0.51‰ ± 0.20‰ Mediterranean Sea BACONNAIS et al., 2018 +0.08‰ ± 0.17‰ lithogenic fraction MOYNIER et al., 2017 +0.31‰ ± 0.11‰ bioauthigenic fraction LITTLE et al., 2017 +0.44‰ ± 0.23‰ Fe-Mn crusts and nodules LITTLE et al., 2014 +0.68‰ river water LITTLE et al., 2018 +0.11‰ ± 0.09‰ lithogenic part of particulate phase LITTLE et al., 2018 +0.40‰ ± 0.10‰ labile fraction of particulate phase LITTLE et al., 2018 +0.10‰ to +0.35‰ marine particles MARÉCHAL et al., 1999 +0.03‰ to +0.52‰ marine particles THOMPSON et al., 2014 +0.08‰ ± 0.17‰ lithogenic fraction, average MOYNIER et al., 2017 G eologia C roatica 185Fiket et al.: Non-traditional stable isotope signatures in geological matrices as a tool for interpreting environmental changes – a review et al., 2016) and Cr(III) can be removed from seawater by phyto- plankton via adsorption or incorporation, both of which may cause isotope fractionations induced by the initial reduction of Cr(VI) to Cr(III) (SEMENIUK et al., 2016). The stable isotopic signature of copper also has the potential to provide significant information on the biogeochemical cycling of this element in the environment (MOYNIER et al., 2017). Similar to chromium, redox transformations between Cu(I) and Cu(II) species are the main processes in natural systems result- ing in Cu isotope fractionation (LITTLE et al., 2014; 2017), e.g., reduced and precipitated Cu(I) species are known to be lighter by 2‰ to 5‰ relative to dissolved Cu(II) species (RYAN et al., 2014). The fractionation between organically complexed and free inorganic species, ranging from +0.1‰ and +0.8‰ (RYAN et al., 2014), suggests that complexation by soluble ligands in natural waters further induces significant isotope fractionation of Cu. In seawater, 99.9% of dissolved Cu is organically complexed, whereby complexation of Cu with OM favours retention of the heavy isotope in solution (VANCE et al., 2008). The organic mat- ter (OM) is thus considered to be a key variable controlling Cu speciation in both terrestrial and aquatic environments. 5. NON-TRADITIONAL ISOTOPES AS A TOOL FOR DECIPHERING MULTIPLE SOURCES CONTRIBUTION By combining data on partitioning and the isotopic fractionation of transition metals between different phases in both terrestrial and aquatic systems, (e.g. between oxides, sulfides or organic matter, or dissolved vs. particulate; Table 3; Fig. 2), the contribu- tion of the different sources can be differentiated. LITTLE et al. (2018) studied Cu isotope distributions in the dissolved and par- ticulate phase in the South Atlantic. Their observations point to the existence of two pools of Cu isotopes in the particulate phase, a refractory pool with a lithogenic δ65Cu signature (at about 0‰) and a labile pool, associated with organic matter, at about +0.4‰ (Table 3). According to their data, the labile pool is isotopically lighter compared to the homogeneous deep ocean dissolved pool at about +0.7‰ (VANCE et al., 2008; RYAN et al., 2014). Although the ocean interior is considered to have a very ho- mogeneous Cu isotope composition at about +0.65‰ (+0.66 ± 0.07‰ at depths >200 m in the South Atlantic; LITTLE et al., 2018); +0.65 ± 0.07‰, >200 m in the North Atlantic (BOYLE et al., 2012); +0.65 ± 0.08‰, >200 m in the Indian and Pacific Oceans (TAKANO et al., 2014); +0.70 ± 0.11‰, >200 m in the Tasman Sea (THOMPSON et al., 2014); and +0.51‰ ± 0.20‰, in the Mediterranean Sea (BACONNAIS et al., 2018), devia- tions towards lower δ65Cu values have been reported for the up- per water column (between +0.41‰ and 0.49‰), and towards higher δ65Cu values in the deeper parts of the water column (be- tween +0.75‰ and 0.84‰) (MOYNIER et al., 2017) and refere- nce therein). This variation in isotopic composition along the depth profiles can be attributed to biological activity, aerosol depo sition, and local supply (LITTLE et al., 2018). Despite this variability, the Cu isotope composition of seawater is similar to the riverine input value of +0.68‰ and to the calculated com- bined riverine and atmospheric Cu source of +0.63‰, which is substantially higher compared to the upper continental crust (UCC) at +0.08 ± 0.17‰ (MOYNIER et al., 2017). Similar to Ni, only two major sinks have been identified so far for Cu, Fe-Mn crusts and nodules and organic-rich sediments, and both of them display a preference for lighter Cu isotopes, at +0.44‰ (LITTLE et al., 2014) and +0.28‰ (LITTLE et al., 2017), respectively. The discrepancy between the light isotope signature of sinks and the heavy isotope signature of the major inputs by rivers requires identification of the additional light source(s) or heavy sink(s). The role of ferromanganese crusts as archives for deep-water Ni isotope compositions was studied by GUEGUEN et al. (2016) and GALL et al. (2013). GUEGUEN et al. (2016) showed that de- spite different growth rates, textures and geochemical patterns, Fe-Mn crusts from both the North and South Pacific Oceans have had homogenous Ni isotope compositions over the last ∼17 Ma, yielding average δ60Ni values of 1.79 ± 0.21 ‰ and 1.73 ±0.21 ‰, respectively. Results of GALL et al. (2013) suggested that the heavy Ni isotopic signature of crusts (+0.9‰ to +2.5‰) reflects the input to the ocean from continental weathering and possibly also from hydrothermal fluids. Despite significant variation in the Ni isotopic composition of surface scrapings of Fe-Mn crusts, they observed low δ60Ni variability in different layers of crust from the central Pacific Ocean that started growing over 70 Ma, suggesting that oceanic sources and sinks largely remained in steady-state over the Cenozoic. Still, the δ60Ni values reported by GALL et al. (2013) are both heavier and lighter than the modern seawater value determined by CAMERON & VANCE (2014), at 1.44 ± 0.15‰. Moreover, nickel isotope depth profile s in water columns of the Pacific, Atlantic, and Southern Oceans studied by CAMERON & VANCE (2014) showed no variations, suggesting that Ni isotopes are homogenous in deep waters. At steady-state, the Ni isotope composition of seawater is thus controlled by the relative fluxes of Ni inputs to the ocean (e.g. rivers, atmospheric deposits, hydrothermal sources) and Ni uptake (e.g. authigenic sinks, organic matter burial) (GUEGUEN et al., 2016). Since the riverine flux and the scavenging of Ni by Fe-Mn oxides are, re- spectively, the main input and output fluxes for this element in the ocean (GALL et al., 2013; CAMERON & VANCE, 2014), GUEGUEN et al. (2016) proposed that the modern marine Ni isotope mass balance is controlled by the isotopic composition of these fluxes. Although the mechanisms that control the Cu and Ni isotope fractionation in seawater are still not well understood, their δ65Cu and δ60Ni isotopic signatures have the potential to distinguish the contribution of different phases to the sedi- ments and when combined with seawater records could provide some new insights into their cycling in both modern and ancient environments. However, to further understand these processes it is necessary to gather a larger dataset of δ65Cu and δ60Ni encom- passing several sources and sinks, as well as different oceanic water masses and contrasting biogeochemical domains. 6. CONCLUSION From the above, it is evident that the development of new tech- niques not only provided answers to existing questions but also opened new topics and created new challenges. However, it is also increasingly clear that a complete and proper interpretation of the various processes during the geological past requires a multi- proxy approach. The above-listed elements and their isotope sig- natures have the potential to explain a number of processes in both recent and palaeo-environments (e.g. redox conditions, the influence of organic matter, recrystallization, diagenesis, weathe- ring, contribution of sources, etc.). Combining such information with additional geochemical and mineralogical data will cer- tainly ensure their more complete and correct interpretation. The lack of data is currently a limiting factor and calls for new studies that will expand our knowledge on non-traditional isotope signatures and provide new insights into many yet-unre- solved geological problems. G eo lo gi a C ro at ic a 186 Geologia Croatica 74/2 ACKNOWLEDGEMENT The authors would like to acknowledge the valuable contributions of the anonymous reviewers and editors. Funding: This research received no external funding. Conflicts of Interest: The authors declare no conflict of in- terest. 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