Ostracod shell chemistry as proxy for paleoenvironmental change

Quaternary International xxx (2013) 1e21

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Ostracod shell chemistry as proxy for paleoenvironmental change Nicole Börner a, *, Bart De Baere b, Qichao Yang c, Klaus Peter Jochum c, Peter Frenzel d, Meinrat O. Andreae c, Antje Schwalb a a

Institut für Geosysteme und Bioindikation, Technische Universität Braunschweig, Langer Kamp 19c, 38106 Braunschweig, Germany Department of Earth, Ocean and Atmospheric Sciences, University of British Columbia, Vancouver, Canada c Biogeochemistry Department, Max Planck Institute for Chemistry, Mainz, Germany d Institut für Geowissenschaften, Friedrich-Schiller-Universität Jena, Jena, Germany b

a r t i c l e i n f o

a b s t r a c t

Article history: Available online xxx

The application of ostracod shell chemistry as a paleoenvironmental tool has grown within the past decades. Most studies have investigated Mg and Sr in ostracod shells as proxies for temperature and salinity, and the use of a wide range of trace elements as prospective paleoenvironmental indicators has yet to be developed for lacustrine systems. Only a few preliminary studies have used trace metals in paleolimnological studies such as Cd, Ba and Zn as paleonutrient indicators, or Mn, Fe and U as redox and oxygenation indicators. This paper reviews the state of the art of geochemical analyses in microfossils such as ostracods, foraminifera, and corals, and provides insights in new trace element proxies with the goal to promote the use of trace elements in ostracod shells as paleoenvironmental proxies. In paleoceanography, foraminifera and corals are most prominently used to reconstruct past climate conditions. Well-established proxies such as d18O, Mg/Ca and Sr/Ca provide information about changes in sea surface temperatures. In addition, a great number of new proxies have been developed recently, such as radiogenic isotopes and redox sensitive trace elements. In paleolimnology, ostracod shell chemistry is widely used to assess paleohydrological changes. Reconstruction of temperature and salinity changes in lake environments is often achieved by oxygen isotopes as well as Mg/Ca and Sr/Ca ratios, but depending on the hydrological and geological settings of the lake system, local calibrations are needed to assess which proxy is suited to reflect which processes. New proxies need to be tested by novel techniques that recently have become available. Compared to conventional instrumentation used in ostracod shell chemistry, methods such as Laser Ablation ICP-MS and NanoSIMS allow single shell analysis and provide high-resolution data. The potential of ostracods in paleolimnology is not yet fully assessed, but can be developed by learning from paleoceanographic studies. Ó 2013 Elsevier Ltd and INQUA. All rights reserved.

1. Introduction

In recent decades ostracods, have been increasingly used as microfossil indicators in lake systems, while the majority of geochemical approaches in paleoclimate research using biogenic carbonates focus on marine foraminifera. In order to interpret coastal, brackish, or freshwater environments, ostracods are more suitable, as they tolerate a wide salinity range. Ostracod assemblages have initially been used as a salinity proxy (De Deckker and Geddes, 1980; De Deckker, 1981, 1982; Forester, 1983, 1986; Smith, 1993; Curry, 1997), as they allow establishing conductivity transfer functions (Mezquita et al., 2005; Mischke et al., 2007). In the marine environment, Mg/Ca in planktonic foraminifera and Sr/Ca in corals are used as recorders of sea surface temperatures (SST), Mg/Ca in benthic foraminifera and ostracods as proxies of bottom water temperature, and the coupling of d18O and Mg/Ca can be used to adjust for the temperature-dependency of d18O by

Biological remains are particularly well suited for the reconstruction of paleoenvironments and paleohydrology because they are abundant and very diverse, comprising hundreds of taxa including diatoms, chrysophytes, charophytes, ostracods, corals, mollusks and foraminifera (Gasse et al., 1987). In lacustrine sediments, the most abundant calcareous organism remains are shells from ostracods (Holmes, 2003). These provide a discrete source of biogenic carbonate, an extremely valuable Quaternary paleoenvironmental indicator.

* Corresponding author. E-mail address: [email protected] (N. Börner). 1040-6182/$ e see front matter Ó 2013 Elsevier Ltd and INQUA. All rights reserved. http://dx.doi.org/10.1016/j.quaint.2013.09.041

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isolating the temperature signal from the d18Owater record, which then can be used to reconstruct local changes in evaporation and precipitation (Rosenthal and Linsley, 2006). Primary production and decay of organic matter can be assessed using carbon isotopes (Kahn, 1979; Gasse et al., 1987; Basak et al., 2009). In coastal environments the freshwater influx can be deduced using O-, C- or Srisotopes (Janz and Vennemann, 2005). In lacustrine environments, d18O, d13C, Mg/Ca and Sr/Ca ratios in ostracod shells are useful tools for reconstructing paleotemperatures, paleosalinity, paleoproductivity as well as methanogenesis (Schwalb, 2003; Mischke et al., 2008b; Decrouy et al., 2012; Schwalb et al., 2013). Following recent developments in foraminifera studies, many other trace metals become important in studies of ostracod shell chemistry, as trace metals play an essential role in many geochemical cycles (Anbar and Knoll, 2002; Algeo et al., 2012). The ability to track chemical changes in the oceans, continental waters and the atmosphere through time depends on a thorough understanding of the processes that influence the trace element uptake by sediment and biogenic carbonate, and on the development of suitable trace element proxies for paleoceanographic and paleolimnological research (Canfield, 1998; Saito and Sigman, 2003; Algeo et al., 2012). The use of redox sensitive trace elements in paleoceanographic studies has been extensively investigated (Ricketts et al., 2001; Tribovillard et al., 2006; Boiteau et al., 2012). Because of their conservative behavior and long residence times in oxic seawater, molybdenum and uranium (800 ka for Mo and 450 ka for U, respectively) were found to be most useful in marine environments (Algeo et al., 2012). The MneFe redox cycling within the water column represents an important additional proxy (Algeo et al., 2012). Patterns in redox-sensitive trace element enrichment can thus be used in paleoceanographic studies to evaluate deepwater restriction, deepwater residence time and changes in deepwater chemical composition (Algeo and Lyons, 2006; Algeo and Tribovillard, 2009; Algeo et al., 2012). The investigation of the controls of co-variation, such as the co-variation of U and Mo, is a relatively new area of research (Tribovillard et al., 2012). New proxies need to be tested by novel techniques that recently have become available. Compared to conventional instrumentation used in ostracod shell chemistry, methods such as Laser Ablation ICP-MS and NanoSIMS allow single shell analysis and provide highresolution data. The potential of ostracods in paleolimnology is not yet fully explored, but can be developed by learning from paleoceanographic studies. This study will give an overview of stable isotope and trace element proxies in paleoceanographic and paleolimnological research. A wide range of proxies have found application in foraminiferal, coral or molluscan studies. Prominent proxies such as stable oxygen and carbon isotopes as well as Mg/Ca and Sr/Ca ratios have found wide application in ostracod shell chemistry. On the other hand, the use of radiogenic isotopes or trace elements such as barium, cadmium, or boron has not yet found their way into ostracod research. But these parameters may become available for ostracod research, if they show similar behavior in marine and lacustrine environments. Thus, the development of the use of trace elements in ostracod shells may significantly broaden the range of existing paleoenvironmental proxies. 2. History During the last three decades trace element and stable isotope geochemistry of ostracods has become a prominent tool in paleoenvironmental reconstruction. First attempts to develop quantitative tools using oxygen isotopes in biogenic carbonates had already been undertaken in the middle of the twentieth century by Urey

(1947), who suggested that oxygen isotopes reflect changes in both temperature and ice volume. His students Epstein and Mayeda (1953) and Emiliani (1955, 1966) developed the first paleotemperature reconstructions using foraminifera from deep-sea cores to calculate changes in ocean temperature. Shackleton et al. (1973) described the fractionation of oxygen isotopes in foraminifera calcite and their relation to ice volume and temperature. Comparisons of stable isotope based reconstructions with those from other proxies confirmed the outstanding paleoenvironmental information derived from geochemical data (Fairbanks et al., 1980; Gasse et al., 1987; Lewis and Anderson, 1992; Fritz et al., 1994). Until the 1980s, species assemblages and stable carbon and oxygen isotopes were the major paleoclimatic tools, recently complemented by new proxies in foraminifera. Besides stable oxygen and carbon isotopes, Mg/Ca ratios (e.g., Nürnberg et al., 1996) and Sr/Ca ratios as well as calcium isotopes (d44Ca) are used to reconstruct paleotemperatures (Kisakürek et al., 2011). Proxies have been developed to reconstruct past salinity changes in marine and marginal marine (Carpenter et al., 1991; Corrège and De Deckker, 1997; Gillikin et al., 2006) as well as lacustrine environments (Chivas et al., 1985; De Deckker et al., 1988a; Engstrom and Nelson, 1991). Other authors have attempted to reconstruct past ocean circulation using Cd/Ca ratios and radiogenic isotopes (Vance et al., 2004; Colin et al., 2010; Makou et al., 2010; Van de Flierdt et al., 2010). Especially for the marine environment, additional proxies have been developed, such as the reconstruction of the carbon cycle including proxies for ocean productivity (231Pa/230Th, U concentration) (Tribovillard et al., 2006; Boiteau et al., 2012), nutrient utilization (Cd/Ca, d15N, d30Si) (Boyle, 1981; Perga, 2010; Versteegh et al., 2011), alkalinity (Ba/Ca) (Boyle, 1981; Chivas et al., 1983), pH (d11B), carbonate ion concentration (Eggins et al., 2003; Elmore et al., 2012) and atmospheric CO2 (d11B, d13C) (Mii et al., 2001). Stable radiogenic isotopes (87Sr, 187Os, 143Nd) can be used to assess fluxes from continent to ocean (Vance and Burton, 1999; Henderson, 2002; Kober et al., 2007). 3. Ostracods as paleoenvironmental proxies Ostracods are bivalved micro-crustaceans and can be found in most aquatic environments, e.g., in continental, estuarine, marine and hypersaline waters (Holmes, 1992). They are generally 0.2e 1.0 mm long and weigh between 20 and 200 mg (Rosenthal and Linsley, 2006). Ostracods have up to 8 molting stages until they become adults, each time secreting an exoskeleton (carapace) of low-magnesium calcite using Ca2þ and HCO 3 ions from ambient water. Prior to molting, large amounts of calcium and phosphate are stored in the outer epidermis to have calcium quickly available for calcifying the new carapace (Keyser and Walter, 2004). The procuticle, the calcified part, is between 20 and 45 mm thick (Sylvester-Bradley and Benson, 1971) and covered by a continuous cuticular integument, consisting of a thin chitinous and organic membrane (Bennett et al., 2011). Keyser and Walter (2004) summarized that the carapace consists of 80e90% calcite, 2e15% organic material (chitin plus protein) and various minor and trace elements (Delorme, 1970). Shell calcification takes place very quickly, typically within a few hours to several days, in geochemical equilibrium with the ambient water. After molting, the components used to form the shell are taken directly from the host waters without storage prior to molting (Turpen and Angell, 1971). Thus, the chemistry of ostracod shells reflects the chemical conditions of the ambient water (e.g., temperature, salinity, dissolved ion composition, hydrology) at the time of molting and new valve growth (Chivas et al., 1985; Holmes, 1992; Smith, 1993; Smith and Horne, 2002; Mischke et al., 2007; Mischke and Holmes, 2008) and can be used as proxy for

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environmental variables (Chivas et al., 1986a; Dean and Schwalb, 2002; Ito et al., 2003; Alvarez Zarikian et al., 2005; Anadón et al., 2006; Mischke et al., 2010b; Marco-Barba et al., 2013). 4. Ostracod shell chemistry For the use of ostracod shell chemistry as paleoenvironmental proxy and for appropriate data analysis it is crucial to know the ecology and phenology of the target species (Pérez et al., 2013). Phenological variation often occurs with respect to the hatching and molting of juveniles, females and males (Xia et al., 1997a), and insight in timing and location of calcification with respect to the preferred microhabitat is invaluable (Van der Meeren et al., 2011). The major driving factors of variations in ostracod valve chemistry are spatiotemporal variability in the physicochemical environment of the host water, life history, habitat and biocalcification (Van der Meeren et al., 2011). Ostracod shells have been used as source material for geochemical analysis of stable isotope and trace element composition in paleolimnological reconstruction of lake hydrochemistry and climate as they provide insight into past water balance and solute evolution of lakes (Lister, 1988; Chivas et al., 1993; Haskell et al., 1996; Bridgwater et al., 1999; Yu and Ito, 1999; Gouramanis et al., 2010). A review about proxies and their interpretation is summarized in Table 1. However, the challenge remains to adequately clean samples prior to analysis, because surface coatings, added after formation of the carbonate, often contain significant quantities of the trace metal of interest (Holmes, 1996; Ito, 2002; Keatings et al., 2006). In addition, studies of trace element and stable isotope composition of ostracod shells from pre-Pleistocene sediments are relatively sparse. Extending ostracod shell chemistry to longer time scales is complicated, as diagenesis becomes a greater problem and water chemistry is not well known (Mucci and Morse, 1983; De Deckker et al., 1999; Dwyer et al., 2002; Bennett et al., 2011).

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4.1. Paleotemperature and paleosalinity Information about global ice volume, salinity and temperature are provided by oxygen isotopes in biogenic carbonates (Chivas et al., 1986a, 1986b; Talbot, 1990). Isotopically depleted water is preferentially stored in polar ice sheets, whereas the ocean is characterized by more positive d18O values (Shackleton et al., 1973). Hence, more positive oxygen isotope values in fossil specimens from the marine realm are interpreted as indicative of glacial conditions; more negative d18O as interglacial conditions. Temperature affects the fractionation of oxygen isotopes during formation of carbonate. An increase in temperature by 1  C results in a corresponding decrease in d18Ovalve of 0.2& (Craig, 1965; Chivas et al., 1986a; Bennett et al., 2011). Oxygen isotope values in ostracod valves are also affected by salinity changes. Marco-Barba et al. (2012) reported a significant relationship between ostracod and water d18O for Cyprideis torosa, except at high Total Dissolved Solids (TDS >20 g L1) where d18Ovalve values were lower than expected. Evaporation and hence higher salinity results in higher d18Owater values and thus higher d18Ovalve (Chivas et al., 1986a; Wrozyna et al., 2010; Pérez et al., 2013). A strong correlation between TDS and Mg/Ca or Sr/Ca ratios has been shown in numerous studies. Van der Meeren et al. (2011) suggested that Mg/Ca might have the highest potential as salinity proxy. Both Mg/Cavalve and Sr/ Cavalve correlate strongly with the respective Mg/Cawater and Sr/ Cawater (Hu et al., 2008). Other studies described a relationship between Sr/Ca ratios and salinity, as high Sr/Ca ratios are indicative for lacustrine environments and a rapid decrease in Sr/Ca coincides with the transition from lacustrine to marine environments (Torgersen et al., 1988; De Deckker et al., 1988a). The uptake of Sr and Mg is primarily a function of their concentration in ambient water and water salinity; in the case of Mg it is also related to water temperature (Carpenter and Lohmann, 1992; Hu et al., 2008; Ito and Forester, 2009). Teeter and Quick (1990) reported a negative

Table 1 Summary of paleoenvironmental proxies and their interpretation in ostracod research. Proxy & indicated environmental parameter

Species

Reference

Location of statement

Candona candida, Cytherissa lacustris, Herpetocypris brevicaudata/chevreuxi Eucypris mareotica, Fabaeformiscandona danielopoli, Limnocythere inopinata Several species Candona candida, Cytherissa lacustris

Hammarlund, 1999

Oxygen isotope records, Conclusion Section 5.

Decrouy et al., 2012 Lister, 1988

Candona subtriangulata

Lewis and Anderson, 1992

Candona sp., Cytherissa lacustris Candona negIecta, Leucocythere mirabilis, Limnocythere sanctipatricii Candona candida, C. levanderi, C. marchica, Cytherissa lacustris, Limnocythere sanctipatricii Pseudocandona marchica Candona neglecta Candona neglecta, Ilyocypris gibba/bradyi, Prionocypris zenkeri Leucocythere mirabilis, Limnocytherina sanctipatricii, Candona rawsoni, C. subtriangulata

Von Grafenstein et al., 1992 Schwalb et al., 1994

Cytheridella ilosvayi Cyprideis torosa Eucypris inflata, Limnocythere inopinata Limnocythere sappaensis Limnocythere ceriotuberosa Candona neglecta

Alvarez Zarikian et al., 2005 Marco-Barba et al., 2012 Lister et al., 1991 Schwalb et al., 1999 Cohen et al., 2000 Ricketts et al., 2001

d18O Temperature

Temperature & d18Owater (d18Owater reflects d18Oprecipitation and thus air temperature or source of inflow)

d18Owater (reflecting shifts in water source, such as input of melt-, groundwater or river water) Temperature, P/E balance & water source

Mischke et al., 2008b

Von Grafenstein et al., 1994, 1996

Section 4. Oxygen Isotopes, Conclusion Climatic interpretation and discussion e e

Von Grafenstein et al., 2000 Anadón et al., 2006 Holmes et al., 2010

Water temperature effects; Section 3.3 Sections 4.1 & 4.2 Concluding remarks Section 4.5

Schwalb et al., 2013

Section 5.1, Conclusion

Last et al., 1994

Stable isotopes in ostracodes Section 7.2, Conclusion Section 4.5 Conclusion Sections 4. & 5. Stable Isotopes Section 4.2.1 (continued on next page)

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Table 1 (continued ) Proxy & indicated environmental parameter

P/E balance & Salinity

P/E balance & water source

Species

Reference

Location of statement

Several species Several species

Schwalb, 2003 Janz and Vennemann, 2005

Eucypris lutzae, Ilyocypris binocularis, Strandesia risgoviensis Fabaeformiscandona breuili, F. spelaea, Pseudocandona compressa Cyprideis torosa Eucypris inflate Eucypris mareotica, Fabaeformiscandona rawsoni, Leucocythere sp. Several species Limnocythere inopinata Cytheridella ilosvayi, Limnocythere opesta, Pseudocandona sp. Several species Australocypris, Mytilocypris

Tütken et al., 2006

Summary and conclusions Section 5.2, Fig. 7, Conclusion Section 5.2

Anadón et al., 2008

Geochemistry: discussion

Anadón and Gabàs, 2009 Zhu et al., 2009 Mischke et al., 2010b

Discussion Section 4.2 Section 4.3

Wrozyna et al., 2010 Van der Meeren et al., 2011 Escobar et al., 2012

Section 5. Section 5.2 Sections 4.1 & 7.

Pérez et al., 2013 Chivas et al., 1993

Section 5.2, Conclusion Oxygen Isotopes, Conclusions Discussion Section 6.2 Discussion of lake evolution

Cyprideis torosa Several species Eucypris mareotica

Marco-Barba et al., 2013 Gouramanis et al., 2010 Mischke et al., 2010a

Cyprideis torosa Candona negIecta, Leucocythere mirabilis, Limnocythere sanctipatricii Candona candida, Cytherissa lacustris, Herpetocypris brevicaudata/chevreuxi Limnocythere ceriotuberosa Eucypris mareotica, Fabaeformiscandona danielopoli, Limnocythere inopinata Cyprideis torosa Several species Several species Candona neglecta Ilyocypris binocularis Eucypris mareotica, Fabaeformiscandona rawsoni, Leucocythere sp. Eucypris mareotica Candona candida, Cytherissa lacustris

Gasse et al., 1987 Schwalb et al., 1994

Biogenic carbonates

Hammarlund, 1999 Cohen et al., 2000 Mischke et al., 2008b

Carbon isotope record, Conclusion Stable Isotopes Section 5.

Anadón and Gabàs, 2009 Wrozyna et al., 2010 Decrouy et al., 2012 Ricketts et al., 2001 Tütken et al., 2006 Mischke et al., 2010b

Discussion Section 5. Section 4. Section 4.2.2 Section 5.8 Section 4.3

Mischke et al., 2010a Lister, 1988 Schwalb et al., 1999 Schwalb, 2003 Alvarez Zarikian et al., 2005 Janz and Vennemann, 2005 Anadón et al., 2006 Anadón et al., 2008

Discussion of lake evolution Carbon isotopes, Conclusion Sections 4. & 5. Summary and conclusions Conclusion Section 5.3, Conclusion Concluding remarks Geochemistry: discussion

Gouramanis et al., 2010 Holmes et al., 2010

Section 6.2 Section 4.4

Escobar et al., 2012

Sections 4.2 & 7.

Marco-Barba et al., 2012 Marco-Barba et al., 2013 Pérez et al., 2013 Schwalb et al., 2013

Section 4.4 Discussion Section 5.2, Conclusion Sections 5.2 & 5.3

Chivas et al., 1983 Holmes et al., 1992 Holmes et al., 1995 Corrège and De Deckker, 1997 De Deckker et al., 1999 Cohen et al., 2000 Janz and Vennemann, 2005 Gouramanis and De Deckker, 2010 Elmore et al., 2012 Chivas et al., 1986b Gasse et al., 1987 Hu et al., 2008 De Deckker et al., 2011 De Deckker et al., 1988a De Deckker and Williams, 1993 Decrouy et al., 2012

Conclusions Section 5.2 Section 5.1, Conclusions Section 3.3 Section 3., Discussion Mg/Ca and Sr/Ca results Section 5.4, Fig. 7 Discussion Sections 3.2 & 3.3 Conclusion Biogenic carbonates Section 4.3, Conclusions Section 8. Fig. 3, Discussion; Results

d13C Productivity

Productivity & CO2 exchange

Productivity & carbon source (possible carbon sources are organic matter decay, CO2 exchange, methanogenesis)

Mg/Ca Temperature

Temperature & Salinity (reflecting P/E balance, lake level, groundwater inflow) Temperature & Mg/Cawater (indicative for lake level changes)

Limnocythere sappaensis Several species Cytheridella ilosvayi Several species Candona neglecta Fabaeformiscandona breuili, F. spelaea, Pseudocandona compressa Several species Candona neglecta, Ilyocypris gibba/bradyi, Prionocypris zenkeri Cytheridella ilosvayi, Limnocythere opesta, Limnocythere sp., Pseudocandona sp. Cyprideis torosa Cyprideis torosa Several species Limnocytherina sanctipatricii, Leucocythere mirabilis Mytilocypris henricae Ilyocypris bradyi Cypretta brevisaepta Bythocypris, Krithe Cyprideis australiensis Limnocythere ceriotuberosa Several species Australocypris, Diacypris, Mytilocypris, Reticypris Krithe Several species Cyprideis torosa Ilyocypris microspinata, Lineocypris jiangsuensis Diacypris spp., Reticypris spp. Cyprideis Cyprideis Several species

Section 4.

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Table 1 (continued ) Proxy & indicated environmental parameter Salinity

Mg/Cawater (but no correlation to salinity)

Mg/Cawater & Salinity (reflecting changes in water source and P/E balance)

Sr/Ca Salinity (signal can be biased by aragonite precipitation)

Sr/Cawater (but no correlation to salinity)

Sr/Cawater & Salinity (information about P/E balance and water source)

Temperature & Sr/Cawater (lake level changes) Carbonate mineralogy (aragonite vs. calcite) Ba/Ca Temperature U/Ca Oxygenation (vertical mixing, organic matter decay) 87 Sr/86Sr Water source (continental weathering, water mass mixing, transition from marine to lacustrine and v.v., freshwater inflow)

Nd/144Nd Water mass tracer Li/Ca Temperature Fe, Mn Oxygenation

Species

Reference

Location of statement

Cyprideis torosa Cyprideis americana

Marco-Barba et al., 2012 Teeter and Quick, 1990

Candona rawsoni Candona rawsoni Candona rawsoni Candona rawsoni Limnocythere inopinata Australocypris, Mytilocypris Cyprideis australiensis Candona neglecta Cyprideis ruggierii, Ilyocypris cf. gibba, Loxoconcha minima Cypridopsis vidua Candona patzcuaro, Heterocypris punctata Eucypris mareotica, Fabaeformiscandona danielopoli, Limnocythere inopinata Cyprideis torosa

Engstrom and Nelson, 1991 Fritz et al., 1994 Haskell et al., 1996 Yu and Ito, 1999 Van der Meeren et al., 2011 Chivas et al., 1986a De Deckker et al., 1999 Ricketts et al., 2001 Anadón et al., 2002

Section 4.2, Conclusions Observation and Discussion, Conclusions Salinity calibration Results Section 4.2 Geochemical analysis Section 5.1 Conclusions Section 3., Discussion Section 4.2.3 Section 6.1

Ito and Forester, 2009 Bridgwater et al., 1999 Mischke et al., 2008b

Summary and Conclusions Sections 4.3.1 & 5. Section 5.

Anadón and Gabàs, 2009

Discussion

Australocypris robusta, Mytilocypris praenuncia and several other species Cyprideis Cyprideis, Ilyocypris

Chivas et al., 1985; 1986a

Conclusion

De Deckker et al., 1988a,b McCulloch et al., 1989

Candona rawsoni Ilyocypris bradyi Australocypris, Mytilocypris Limnocythere ceriotuberosa Candona neglecta Ilyocypris microspinata, Lineocypris jiangsuensis Cyprideis australiensis Cyprideis ruggierii, Ilyocypris cf. gibba, Loxoconcha minima Cyprideis torosa Cypridopsis vidua Australocypris, Diacypris, Mytilocypris, Reticypris Several species Cyprideis torosa Cyprideis torosa Cyprideis Cypretta brevisaepta Candona patzcuaro, Heterocypris punctata Fabaeformiscandona breuili, F. spelaea, Pseudocandona compressa Eucypris mareotica, Fabaeformiscandona danielopoli, Limnocythere inopinata Candona neglecta, Ilyocypris gibba/bradyi, Prionocypris zenkeri Diacypris spp., Reticypris spp. Cyprideis torosa Cyprideis Several species Candona rawsoni

Engstrom and Nelson, 1991 Holmes et al., 1992 Chivas et al., 1993 Cohen et al., 2000 Ricketts et al., 2001 Hu et al., 2008 De Deckker et al., 1999 Anadón et al., 2002

Fig. 3, Discussion Comparison of 87Sr/86Sr and Sr/Ca ratios Salinity calibration Section 5.2 Paleosalinity, Conclusions Mg/Ca and Sr/Ca results Section 4.2.3 Section 4.3, Conclusions Section 3., Discussion Section 6.1

Anadón and Gabàs, 2009 Ito and Forester, 2009 Gouramanis and De Deckker, 2010 Gouramanis et al., 2010 Marco-Barba et al., 2012 Gasse et al., 1987 De Deckker et al., 1988a,b Holmes et al., 1995 Bridgwater et al., 1999 Anadón et al., 2008

Discussion Summary and Conclusions Discussion Section 6.2 Section 4.3, Conclusions Biogenic carbonates Fig. 3, Discussion; Results Section 5.1, Conclusions Section 4.3.1 & 5. Geochemistry: discussion

Mischke et al., 2008b

Section 5.

Holmes et al., 2010

Section 4.3

De Deckker et al., 2011 Marco-Barba et al., 2013 De Deckker and Williams, 1993 Decrouy et al., 2012 Haskell et al., 1996

Section 8. Discussion e Section 4. Section 4.2

Mytilocypris henricae

Chivas et al., 1983

Conclusions

Candona neglecta

Ricketts et al., 2001

Section 4.2.4

Cyprideis, Ilyocypris

McCulloch et al., 1989

Several species Cyprideis torosa Cypretta brevisaepta Several species Candona neglecta, Ilyocypris gibba/bradyi, Prionocypris zenkeri

Janz and Vennemann, 2005 Vasiliev et al., 2006 Holmes et al., 2007 Kober et al., 2007 Holmes et al., 2010

Paleoenvironmental implications Section 5.1, Conclusion Sections 5. & 6. Section 3.5 Section 4.3 Section 4.3

Several species

Janz and Vennemann, 2005

Section 5.1, Conclusion

Eucypris inflata

Zhu et al., 2009

Section 4.3

Cyprideis torosa

Gasse et al., 1987

Biogenic carbonates

143

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correlation between the Mg concentration in shells of living Cyprideis americana and salinity, with little or no temperature dependence of Mg. No temperature effect was also observed by MarcoBarba et al. (2012) as well as no influence of Mg/Cawater on the Mg/Ca ratio in ostracods. In addition, the authors found that in waters with Mg/Ca ratios 6 all Cyprideis torosa shells showed the same Mg/Ca content over the whole salinity range, indicating a lack of correlation between salinity and ostracod Mg/Ca ratios in lowmagnesium waters. The temperature dependence can easily be masked by small changes in the Mg/Ca ratio of the ambient water (De Deckker et al., 1999). Factors influencing the Mg/Ca and Sr/Ca ratios in ostracod shells are vital offsets, as well as the influence of other soluble ions that may change the complexing of the divalent cations in the host waters and hence partitioning of Sr, Mg, and Ca into the shells (see Section 5.). Van der Meeren et al. (2011) assessed how valve chemistry is affected by regional patterns and seasonal trends in solute evolution, the physicochemical stability of the host water, and the time and place of biocalcification. The authors found, that the correlation of Sr/Cavalve to Sr/Cawater was more significant than that of Mg/ Cavalve to Mg/Cawater, and suggested a lesser influence by calcification rate and temperature on the Sr incorporation. A local reduction in dissolved Sr can be caused by aragonite precipitation (Haskell et al., 1996; Hu et al., 2008; Mischke et al., 2008b), and reduced Mg uptake by ostracods can occur in waters with high Mg-content (De Deckker et al., 1999; Gouramanis and De Deckker, 2010), both masking the temperature-dependent signal. An overview of interpretations of geochemical parameters is given in the publications by Miller et al. (1991), Griffiths and Holmes (2000), and Zachos et al. (2001), based on non-marine ostracods from closedbasin lakes where hydrochemical changes occur as a function of precipitation and evaporation as well as groundwater input. According to Zhu et al. (2009) another possible tracer for past temperatures is provided by Li/Ca ratios in ostracod shells. They found a negative correlation between the Li/Ca ratio and temperature in Eucypris inflata from Lake Qinghai, Tibet. A temperature dependency of Li/Ca ratios was also described by Hall and Chan (2004b) and Lear et al. (2010) for several foraminifera species. Hall and Chan (2004b) showed that the conservative behavior of lithium in the ocean (1.5 Ma residence time) results in constant Li/ Ca ratios of water on centennial to millennial scales, indicating that the Li/Ca ratios of carbonates are not mainly controlled by solubility. Tomascak et al. (2003) calculated a residence time of 28 ka for lithium in Mono Lake, California, but Zhu et al. (2009) suggested that lithium in larger lakes, like Lake Qinghai, has longer residence time than in Mono Lake. Thus, in large lakes, the Li/Ca ratios of carbonates are primarily controlled by temperature and not biased by the Li/Ca ratios of the ambient water. 4.2. Carbon cycle The d13C content of carbonate depends on the isotopic composition of the Total Dissolved Inorganic Carbon (DIC) from which it precipitates, influenced by the rate of local production of CO2 and the rate of isotopic exchange with the atmospheric reservoir (Hammarlund, 1999; Alvarez Zarikian et al., 2005; Tütken et al., 2006; Mischke et al., 2010b). A temperature dependency of the carbon isotope composition of the carbonate precipitated in ostracod shells could be excluded in many studies (Durazzi, 1977; Marco-Barba et al., 2012). Additional factors that impact d13C are biological productivity, pH, organic matter decay, and bacterial processes (De Deckker and Forester, 1988; Schwalb et al., 1999; Liu et al., 2008; Escobar et al., 2012). Freshwater or groundwater inflow has been documented to shift the d13C signal toward lighter values (Janz and Vennemann, 2005; Gouramanis et al., 2010; Holmes et al.,

2010). Methanogenesis and methane oxidation also influence the d13C signal in ostracods (Schwalb, 2003; Escobar et al., 2012; Schwalb et al., 2013). Methane oxidation close to the sedimentwater-interface results in 13C-depleted DIC that is used for incorporation into ostracod shells. 4.3. New chemical proxies in ostracod shell chemistry Very few studies exist that address the potential suitability of proxies other than d18O, d13C, Mg/Ca and Sr/Ca. In an early study, Chivas et al. (1983) tested the suitability of Ba/Ca ratios in ostracods. Their experiments revealed a positive correlation between ostracod Ba/Ca ratios and temperature and no relationship to the Ba/Ca concentrations of the ambient water. Palacios-Fest et al. (2003) tried to track anthropogenic influence by measuring Co, Ni, Cu, Zn, Cd, Pb, and REEs in ostracod shells, which may reflect levels of urban and industrial pollution. The authors found that most of the studied heavy metals as well as REEs show increased values in ostracod valves from polluted waters and are thus suitable to track anthropogenic influence. Contradictory results were obtained for Cd and Pb as their concentration in polluted and unpolluted conditions did not show any differences. Additionally, manganese and iron (Gasse et al., 1987; Holmes, 1997) as well as uranium (Ricketts et al., 2001) to calcium ratios were tested as proxies for redox and oxygenation cycles. The authors reported increased element concentrations in reducing environments and related changes in Fe/Ca, Mn/Ca, and U/Ca to shifts in bottom water oxygenation. Ricketts et al. (2001) suggested that vertical water mass mixing and organic matter decay were possible controlling factors of changes in oxygenation. In the marine and marginal marine environment, ostracod valve chemistry was used to assess the extent to which a marine basin is isolated from the open ocean using Nd isotopes as the isotopic composition of this element, 2010 reflects different sources of carbonate, depending on surface/bedrock composition (Janz and Vennemann, 2005). The 87Sr/86Sr ratios trace transitions from marine to non-marine environments and vice versa as well as freshwater inflow (McCulloch and Deckker, 1989; McCulloch et al., 1989; Vasiliev et al., 2006). Strontium isotopes have been applied to limnological studies as they give information about the mixing of waters from different water sources, e.g., river inflow, groundwater inflow, or marine influence (Holmes et al., 2007; Kober et al., 2007; Holmes et al., 2010). 5. Limitations of ostracod shell chemistry as environmental proxy Various theoretical and practical limitations to the use of stable isotopes and trace elements in ostracod shells as paleoenvironmental proxies have been described (Engstrom and Nelson, 1991; Holmes et al., 1995; Ito and Forester, 2009), including nonequilibrium fractionation of stable isotopes and trace elements in ostracod shells, diagenesis and alteration, and the effect of different pre-treatment methods on the shell composition. 5.1. Non-equilibrium fractionation and vital effects Vital effects are very important for calculating past environmental conditions, e.g., temperature. Vital effects describe speciesor genus-specific offsets from the expected equilibrium. Disequilibrium fractionation in biogenic carbonates has been reported in many papers. Most studies, however, focus on disequilibrium in planktonic foraminifera (Moberly, 1968; Fairbanks et al., 1980; Duplessy et al., 1981; Erez and Honjo, 1981; Kahn and Williams, 1981; Kozdon et al., 2009; Kisakürek et al., 2011) (see Section 7).

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Ostracod shells calcify with a constant species-specific offset from isotopic equilibrium with the host water, as shown by results from oxygen isotope studies by Xia et al. (1997b), Von Grafenstein et al. (1999), Li and Liu (2010), Kalm and Sohar (2010) and Decrouy et al. (2011). Candona rawsoni shows a clear and consistent temperature dependence of oxygen isotope fractionation during its biological calcification (Xia et al., 1997b). At 25  C, the optimum molting and calcification temperature, C. rawsoni specimens reached maturity faster than at 15  C and showed an offset of þ2& in d18O values. In inorganic carbonate a temperature difference of 10  C results also in a difference of 2&, suggesting that d18O in ostracod shells is well defined by d18O and the temperature of the host water. An oxygen isotope fractionation similar for all species of Candoninae was reported by Decrouy et al. (2011) with an offset in d18O of more than þ3& relative to equilibrium values for inorganic calcite. In Cytheroidea the oxygen isotope fractionation is less discriminative against 18O, resulting in an enrichment in d18O of 1.7e2.3&. Von Grafenstein et al. (1999) reported offsets from isotopic equilibrium to be greater than 2& for Candona species, 1.5& for Cytherissa lacustris, and around 0.8& for Limnocythere inopinata and Darwinula stevensoni. Li and Liu (2010) calculated a fractionation factor of 1.0311 at 13.5  C in Eucypris mareotica, similar to that of inorganic calcite (1.0308 at 13.5  C). They concluded that in high pH and high salinity waters, such as in Lake Qinghai, the vital offsets in E. mareotica are small enough to be negligible. They also found that fractionation factors move farther away from the equilibrium fractionation as temperature increases. Factors affecting oxygen isotope fractionation are the temperature dependent rate of calcification as well as incomplete calcification caused by environmental stress (Xia et al., 1997b). Erez and Luz (1983) suggested that non-equilibrium fractionation in planktonic foraminifera occurs only in the earlier stages of shell formation when metabolic activity is more intense (Moberly, 1968) and the more rapid calcification is characterized by less discrimination against 16O, resulting in lower d18O for not completely calcified shells. Similar observations were published by De Villiers et al. (1995) for coralline aragonite, describing stronger discrimination against 16O and thus more incorporation of 18O at slow growth rates. Opposite results were found by Xia et al. (1997b) with respect to ostracods, showing less discrimination against 16O at slower calcification. In addition, Keatings et al. (2002b) reported that oxygen isotope fractionation into ostracod shell calcite is not only a function of temperature but is also dependent on the pH of the ambient water. Vital effects affect not only d18O values, but also a wide range of other isotopes and trace elements. The d13C values of different genera were reported to be generally in or near isotopic equilibrium with the host waters (Keatings et al., 2002b; Decrouy et al., 2011). Disequilibrium incorporation into ostracod calcite was observed for Sr/Ca and Mg/Ca (Xia et al., 1997b; Wansard et al., 1998; Von Grafenstein et al., 1999). In high-Mg waters a large physiological energy is required to exclude Mg and Sr during shell calcification (Reddy and Wang, 1980). Hence, an upper limit to the amount of Mg and Sr that ostracods will incorporate into their shells exist as a higher Mg/Ca ratio in the solution leads to lower partition coefficient of Mg (Morse and Bender, 1990). De Deckker et al. (1999) and Gouramanis and De Deckker (2010) reported a reduced uptake of magnesium in high-Mg waters, and an additional influence of high water temperatures on the Mg uptake is given by Marco-Barba et al. (2012). Vital effects and Sr and Mg partition coefficients vary between species (Von Grafenstein et al., 1999) and even within individuals of same species (Wansard et al., 1998). Morishita et al. (2007) studied the distributions of Mg and Sr in the marine ostracod Neonesidea oligodentata and found a banded structure of chemical

7

distributions. The three different bandings with variations in Mg and Sr were described by the authors as a high-Mg and high-Sr outer band, a heterogenous low-Mg and low-Sr middle band, and a high-Mg and low-Sr inner band. They suggested that differences in calcification rates account for the different Mg and Sr distributions. Systematic differences in element distributions between right and left valve, female and male were not observed (Morishita et al., 2007; Marco-Barba et al., 2012). Another problem that should be accounted for is the seasonality of shell formation. Little is known about seasonality for most ostracod species, for example whether shells form during a restricted season or throughout the year, and if seasonality contributes to the variability in the biogeochemistry of fossil assemblages. If molting and shell calcification of A-1 instars (pre-adults) to adults occurs at different times of the year, the resulting intraannual noise may overwhelm long-term paleoclimatic signals. Freshwater populations of Limnocythere inopinata disappear during winter and hatch in early spring, whereas saline water populations overwinter as adults or late juvenile instars (Yin et al., 2001; Van der Meeren et al., 2011). Another example is given in a study by Xia et al. (1997a) investigating the effects of seasonality on the stable isotope variability in C. rawsoni. This species hatches in spring, but adults may calcify in late summer or the following spring depending on water temperature. Seasonal changes in water temperature and variations in water chemistry result in high variability in d18O, d13C, Mg/Ca and Sr/Ca. If C. rawsoni molts over a wide range of temperatures, the mean isotope and trace element signatures of the population as a whole may give an integrated seasonal signal for both temperature and water chemistry. 5.2. Pre-treatment For trace element analysis the removal of contaminants comprising (1) adhering aluminosilicates, (2) carbonate minerals, (3) chemical precipitates, and (4) organic material (Keatings et al., 2006) is essential. Stable isotope studies need to remove organic material (Keatings et al., 2006), as such contaminants can mask the original biogenic signal. Most studies recommend that the valves only be cleaned manually with a fine brush, needle and deionized water whenever possible, as chemical pre-treatment methods have the potential to alter the stable isotope and trace metal composition of the ostracod valves (Holmes, 1996). Only a few studies exist that have investigated the effects of different pre-treatment methods on the ostracod shell composition (Jin et al., 2006; Keatings et al., 2006; Li et al., 2007; Mischke et al., 2008a). More extensive work has been done on the effect of pre-treatment on isotopic composition of other biogenic and inorganic carbonates (Sarkar et al., 1990; Boiseau and Juillet-Leclerc, 1997; Grottoli et al., 2005; Serrano et al., 2008). Other carbonates (such as foraminifera) differ structurally and in some cases mineralogically from ostracod calcite and those studies do not necessarily provide valid models for how ostracod calcite will behave under similar treatments (Keatings et al., 2006). A comprehensive overview of different cleaning techniques and their effects is given by Keatings et al. (2006). In this study, differences in valve composition following pre-treatment were measured by comparing treated and untreated valves from the same carapaces, as the composition of both valves of a single carapace is statistically identical. The removal of organic material for stable isotope studies is usually accomplished by soaking in hydrogen peroxide (Curtis and Hodell, 1993; Von Grafenstein et al., 1996) or sodium hypochlorite (Durazzi, 1977; Grossman et al., 1986; Scotchman, 1989), vacuum roasting or plasma ashing (Boomer, 1993; Andrews et al., 1995). In order to remove adhering aluminosilicate and ferromanganese minerals from the shell surface, as is essential for trace element analysis, reducing or complexing

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solutions are used. With respect to the use of hydrogen peroxide, Ito (2002) suggested that it insufficiently removes organic contamination and that the resulting organic acids can partially dissolve the carbonate (Keatings et al., 2006). Boiseau and JuilletLeclerc (1997) found no exchange between hydrogen peroxide and coral aragonite when they exposed it to 16O enriched hydrogen peroxide. In addition, Keatings et al. (2006) stated that hydrogen peroxide cleaning is a suitable method for oxygen isotope analysis. On the other hand, Xia et al. (1997b) found that cleaning for 15 min in hot (80  C) 5% H2O2 resulted in small (0.1e0.3&) nonsystematic shifts in oxygen isotope composition. Keatings et al. (2006) concluded that pre-treating should only be done if absolutely necessary and with full awareness of any resulting signal bias. They suggested that no pre-treatment should be carried out unless a visual inspection revealed substantial organic or mineral contamination of valves. Furthermore, Keatings et al. (2006) suggested to employ both hydrogen peroxide and plasma ashing for oxygen isotopes analyses, and only plasma ashing when measuring carbon isotopes. Sodium hypochlorite pretreatment is recommended for the analysis of Mg/Ca and Sr/Ca. The most effective method for the removal of the chitin sheath from living or recently dead ostracod valves may be freeze drying of the valves, which loosens chitin for easy removal, leaving only the mineralized portion of the carapace behind, but this approach needs further testing. Vacuum roasting, hydrogen peroxide and sodium hypochlorite caused a reduction in the mean d13C values and increased deviations in d18O values by 0.8&. Sodium hypochlorite has less impact to Mg/Ca and Sr/Ca ratios than hydrogen peroxide, and hydroxylamine hydrochlorite caused no significant changes in Mg/Ca and Sr/Ca ratios. Both techniques led to a reduction in the shell weight, suggesting that Mg, Sr and Ca are removed in equal proportions (Keatings et al., 2006). In conclusion, no single pre-treatment method was found to be free from any effects on the ostracod valves. Even when using the most suitable pre-treatment method, sample drying can also significantly alter the isotopic signature of the ostracod shells. Mischke et al. (2008a) tested the effects of drying the sieve residues either from tap water, deionized water, or ethanol rinsing. The stable isotope values of shells dried from water (deionized and tap water, respectively) were lower for both oxygen and carbon as a result of calcite crystals precipitated on the shell surfaces during the drying process. In contrast, samples dried from ethanol showed a smooth and clean surface. The observed isotopic depletion of water-dried shells in comparison to ethanol-dried shells is attributed to the precipitation of inorganic calcite on the ostracod shell surfaces. Deionized water is very mildly acidic and may cause carbonate dissolution and later precipitation during drying, while tap water is a source of additional carbonate. These effects were also reported in foraminifera by Sperling et al. (2002). And as the precipitated crystals are not easily assessable by visual inspection of shells prior to analysis, water drying should be generally avoided (Mischke et al., 2008a).

visualizing techniques are generally used for Quaternary material (Bennett et al., 2011). Chivas et al. (1986b) stated that partial dissolution of valves in mildly acidic waters has no effect on the Sr/Ca ratios, but leaching of Mg occurred in under-calcified Cyprideis valves collected from saline lakes in Australia. These observations were confirmed by laboratory tests with Krithe and Loxoconcha valves, which showed 12% reduced Mg/Ca ratios but original Sr/Ca ratios after 40% of the valve material was removed by dissolution in deionized water (Dwyer et al., 2002). Similarly, De Deckker et al. (1999) found that Sr/Ca ratios of partially dissolved valves are very similar to unaltered shells, but Mg/Ca ratios are significantly lower in altered material. These results suggest that Sr is homogeneously distributed, whereas Mg is not. All authors suggested to use only pristine valves in geochemical studies, but this is not always possible (Corrège, 1993). There tends to be a relationship between the physical preservation of valves in the sediment and the presence or absence of the enveloping chitinous cuticle. Removal or damage of this cuticle tends to result in poorly preserved valves. Similar findings were also reported for foraminifera by Lorens et al. (1977). In contrast, an investigation into the effects of preservation on Mg/Ca, Sr/Ca, d18O and d13C values in late Quaternary non-marine ostracod valves by Keatings et al. (2002a) yielded no evidence that the preservation state affected valve geochemistry. Diagenetic overprinting on low-Mg calcite brachiopod shells is well documented and shows that diagenesis tends to lead to lower carbon and oxygen isotope values (Banner and Kaufman, 1994; Mii et al., 2001). Diagenetic alteration by later meteoric water also tends to lead to lower carbon and oxygen isotope compositions compared to alteration by seawater (Immenhauser et al., 2002). Another important factor is burial depth, as increasing temperatures with depth lead to lower d18O, which was shown in studies of benthic foraminifera bulk carbonate (Schrag et al., 1995). Bennett et al. (2011) published a detailed study about the diagenesis of fossil ostracods, assessing the ultrastructure of the calcite crystals of the shells for different diagenetic stages. The authors described six diagenetic stages: replacement of original calcite by (1) neomorphic calcite, (2) framboidal and euhedral pyrite, (3) ferroan calcite, (4) ferroan dolomite, (5) siderite growth, (6) sphalerite and barite mineralization. The values of d13C and d18O tend to become more negative with increasing diagenesis, although the effect is often not measureable for carbon isotopes. Nevertheless, the ostracod neomorphic calcite (stage 1) preserves a seawater signal, even though there may have been some alteration soon after shallow burial. As the recrystallization process is currently poorly understood, it has been suggested that the original crystals may not have completely dissolved but instead just grown in size (Bennett et al., 2011). For diagenetically altered planktonic foraminifera a similar ultrastructure to that of neomorphic calcite in ostracods was described and interpreted to have grown by dissolution and recrystallization of the test at shallow burial depth of less than 300 m (Pearson et al., 2001).

5.3. Diagenesis 6. State of the art in foraminiferal geochemistry A number of studies have suggested that post-mortem diagenesis may significantly affect the chemical composition of nonmarine ostracod valves (Mucci and Morse, 1983; Chivas et al., 1986b), but the question, under which circumstances diagenesis can alter the shell chemistry, has received little attention in the literature (Keatings et al., 2002a). Diagenetically altered specimens feature pitted, dissolved or re-crystallized surfaces. A visual assessment of the state of preservation is possible using light microscopy (Dwyer et al., 2002; Keatings et al., 2002a). The degree of surface recrystallization can be identified using scanning electron microscopy imagery (Mischke et al., 2008a). Both of these

6.1. Paleothermometry and paleosalinity For paleoenvironmental reconstruction, especially in the marine environment, both faunal and geochemical studies have concentrated on foraminifera (Table 2). Sea surface temperature reconstructions using oxygen isotopes were originally developed using foraminifera-based paleoproxies (Emiliani, 1955, 1966; Zachos et al., 2001; Sagawa et al., 2013). In addition to the effect of temperature, the d18O values in foraminifera tests are also influenced by the seawater d18O composition. The value of

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d18Oseawater reflects the hydrological balance between evaporation and precipitation (E/P) and surface salinity (Corrège and De Deckker, 1997; Guildersen and Pak, 2005). On longer time scales, d18Oseawater is also a function of the extent of continental ice sheets (Shackleton et al., 1973; Miller et al., 1991; Mulitza et al., 2004).

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Foraminiferal Mg/Ca was proposed to be a proxy for temperature, as Mg/Ca in marine carbonates varies with latitude, suggesting a temperature-dependence (Rosenthal et al., 1997b; Lea, 2003; Lear et al., 2010). The application of Mg/Ca on longer time scales is limited by a lack of knowledge about past seawater composition.

Table 2 Summary of established paleoenvironmental proxies in foraminiferal, coral and molluscan geochemical research (selected papers). Proxy & indicated environmental parameter

Foraminifera

Coral

Mollusk

Temperature

Emiliani, 1955, 1966; Shackleton et al., 1973; Sagawa et al., 2013

Emiliani et al., 1978; Hart and Cohen, 1996

Temperature & salinity

Corrège and De Deckker, 1997; Guildersen and Pak, 2005 Zachos et al., 2001; Mulitza et al., 2004; Lear et al., 2010 Miller et al., 1991

Von Grafenstein et al., 1992; Brand et al., 1993; Von Grafenstein et al., 1994; Von Grafenstein et al., 2000; Wurster and Patterson, 2001; Dutton et al., 2002; Gillikin et al., 2005; Schöne et al., 2005; Gillikin et al., 2006; Elliot et al., 2009; Batenburg et al., 2011 Carroll et al., 2009

Smith et al., 1997; Linsley et al., 2004

Anadón et al., 2008

18

d O

Temperature & ice volume Ice volume Temperature & d18Owater (meltwater and freshwater input) Vital effects (growth rate) Salinity & P/E balance d13C Productivity & nutrient distribution

Nutrient distribution and CO2 exchange Productivity & carbon source Mg/Ca Temperature

Temperature & salinity Temperature & saturation state Vital effects (growth rate)

Tripati et al., 2001; Apolinarska, 2013 Gasse et al., 1987 Lea, 1995; Rosenthal et al., 1997b; Bickert and Mackensen, 2004; Mackensen and Licari, 2004

Anadón et al., 2008; Taft et al., 2013 Rosenthal et al., 1997a; Gagan et al., 2004; Sadekov et al., 2008; Hathorne et al., 2009; Sadekov et al., 2009 Nürnberg et al., 1996 Lear et al., 2010

Hart and Cohen, 1996; Sinclair et al., 1998; Fallon et al., 1999; Montagna et al., 2007

Productivity & freshwater input Temperature Salinity U/Ca Temperature Temperature & freshwater input Redox changes & Oxygenation U/Cawater d11B, B/Ca pH

Lazareth et al., 2003; Anadón et al., 2008; Carroll et al., 2009; Elliot et al., 2009; Batenburg et al., 2011 Dodd, 1965; Dodd and Crisp, 1982 Dutton et al., 2002; Carré et al., 2006; Schöne et al., 2011; Schöne et al., 2013

Hart and Cohen, 1996; Sinclair et al., 1998; Fallon et al., 1999; Marshall and McCulloch, 2002; Gagan et al., 2004; Linsley et al., 2004; Cohen et al., 2006; Montagna et al., 2007

Dodd, 1965; Lazareth et al., 2003; Carroll et al., 2009

Kisakürek et al., 2011 Rosenthal et al., 1997a Dutton et al., 2002; Gillikin et al., 2005; Carré et al., 2006; Schöne et al., 2011; Schöne et al., 2013 Brand et al., 1993; Anadón et al., 2002; Anadón et al., 2008 Dodd and Crisp, 1982

Sr/Cawater Salinity Ba/Ca Alkalinity & nutrient distribution Productivity

Gasse et al., 1987; Elorza and GarcìaGarmilla, 1996; Wurster and Patterson, 2001; Dutton et al., 2002; Schöne et al., 2005; Gillikin et al., 2006; Carroll et al., 2009; Apolinarska, 2013

Makou et al., 2010

Sr/Ca Temperature

Temperature and Salinity Saturation state Vital effects (growth rate)

Emiliani et al., 1978

Lea and Spero, 1994; Lea, 1995 Fallon et al., 1999

Lazareth et al., 2003; Elliot et al., 2009; Batenburg et al., 2011 Carroll et al., 2009

Montagna et al., 2007 Gillikin et al., 2006, 2008 Russell et al., 1996; Yu et al., 2008

Sinclair et al., 1998; Fallon et al., 1999 Montagna et al., 2007

Boiteau et al., 2012 Russell et al., 1994 Sanyal et al., 1996; Lemarchand et al., 2000; Pearson and Palmer, 2000; Sanyal et al., 2000

Hemming et al., 1998

(continued on next page)

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Table 2 (continued ) Proxy & indicated environmental parameter pH & Temperature Temperature

87

Sr/86Sr Sr-source & continental weathering

Sr-source & water masses mixing (freshwater input, ocean circulation)

Temperature & continental weathering Nd/144Nd Ocean circulation & continental weathering

Foraminifera

Coral

Mollusk

Yu et al., 2007; Hathorne et al., 2009 Hart and Cohen, 1996; Sinclair et al., 1998; Fallon et al., 1999; Montagna et al., 2007 Dia et al., 1992

Fan et al., 2010

Van de Flierdt et al., 2006; Colin et al., 2010; Copard et al., 2010; Van de Flierdt et al., 2010; López Correa et al., 2012 Rüggeberg et al., 2008

Whittaker and Kyser, 1993; Tripati et al., 2001; Anadón et al., 2002

Palmer and Elderfield, 1985; Vance and Burton, 1999; Burton and Vance, 2000; Pomiès et al., 2002; Vance et al., 2004; Klevenz et al., 2008; Roberts et al., 2012

Sholkovitz and Shen, 1995

Whittaker and Kyser, 1993

Marriott et al., 2004 Hall and Chan, 2004a,b; Lear et al., 2010 Lear and Rosenthal, 2006

Case et al., 2010

Hess et al., 1986; Hodell et al., 1989; Martin and Macdougall, 1991; Miller et al., 1991; Dia et al., 1992 Shijie, 1996

143

Li/Ca, Mg/Li (corals) Temperature Temperature & saturation state Seawater saturation F/Ca Temperature Cd/Ca Productivity and water mass mixing

Nutrient distribution and exchange with CO2 Mn/Ca Productivity Productivity & circulation Nutrient distribution, freshwater input Oxygenation 44 d Ca Temperature

Hart and Cohen, 1996 Van Geen et al., 1992; Lea, 1995; Rosenthal et al., 1997b; Burton and Vance, 2000; Makou et al., 2010 Rosenthal et al., 1997a

Wyndham et al., 2004

Carroll et al., 2009

Klinkhammer et al., 2009 Lazareth et al., 2003 Brand et al., 1993 De La Rocha and DePaolo, 2000; Nägler et al., 2000; Kisakürek et al., 2011

d15N Productivity, anthropogenic influence REEs Alkalinity Water mass tracer (riverine input, volcanic activity, aeolian influx) Productivity V/Ca Redox changes P/Ca Nutrient distribution 238 Pu/239Pu Anthropogenic influence Al, Fe, Cu, Cr, Hg, V, Zn Anthropogenic influence

Versteegh et al., 2011 Roberts et al., 2012 Sholkovitz and Shen, 1995

Whittaker and Kyser, 1993

Wyndham et al., 2004 Hastings et al., 1996

However, because Mg as well as Ca have long residence times (13 Ma for Mg, 1 Ma for Ca) (Rosenthal and Linsley, 2006), the composition of seawater cannot have changed significantly during the Pleistocene and thus provides a good basis for sea surface temperature reconstructions at least at this time scale (Gagan et al., 2004; Sadekov et al., 2008; Sadekov et al., 2009). Nevertheless, inter-species variability in Mg/Ca ratios is correlated with foraminifera calcification depth, as shallow mixed-layer dwellers like Globigerinoides ruber and Globigerinoides sacculifer have high Mg/ Ca, while deep dwellers (e.g., Globigerinoides tumida, Globigerinoides dutertrei) are relatively low in Mg/Ca (Rosenthal and Boyle, 1993). Large compositional variability within individual chambers and within the test as a whole also occurs and depends on the vertical migration during adult life; shallow-dwelling species have a more homogenous composition than deep-dwelling species

LaVigne et al., 2008 Imai and Sakanoue, 1973 Bastidas and García, 1999

(Rosenthal and Linsley, 2006; Allison and Austin, 2008; Hathorne et al., 2009). Laser ablation ICP-MS microanalysis of the foraminifera Globigerinoides ruber revealed large variations of Mg/Ca composition within and between individual tests and suggested a source of significant uncertainty in Mg/Ca thermometry, but mean molar values show a strong exponential correlation with mean annual sea surface temperatures (Sadekov et al., 2008; Hathorne et al., 2009; Sadekov et al., 2009). Mg/Ca ratios decrease with increasing calcification depth and a large interspecific variability among benthic species suggests that using a single species calibration may yield the best precision. The effect of salinity on the Mg/Ca composition of foraminiferal calcite is currently poorly constrained (Sadekov et al., 2008), but Lear et al. (2010) stated that Mg/Ca may be a salinity-independent paleothermometer. By combining Mg/Ca signatures with d18O values it is possible to

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separate bottom water temperature and ice volume effects in the d18O record. An additional, but yet largely untested tool for paleotemperatures is provided by foraminiferal Ca isotopes, as d44Ca is more robust to diagenesis than Mg/Ca. De La Rocha and DePaolo (2000) found that periods of lower global temperature result in low d44Ca concentrations of foraminifera tests. A correspondence between d44Ca and glacial/interglacial SST fluctuations was also reported by Nägler et al. (2000). In contrary, Kisakürek et al. (2011) found this relationship in benthic but not in planktonic species and also just for temperatures below 24  C. Nevertheless, the potential of calcium isotopes as recorders of sea surface temperatures was reported by all authors. A relatively new proxy is the benthic foraminifera Li/Ca ratio, which reflects changes in both seawater saturation state and temperature (Burton and Vance, 2000; Marriott et al., 2004; Hall and Chan, 2004b; Lear and Rosenthal, 2006). For this proxy, coupled Mg/Ca and Li/Ca may be used to reconstruct both temperature and seawater saturation state (Lear et al., 2010). A decrease in the Li/Ca ratio reflects a lower calcification rate caused by decreased seawater carbonate ion concentration (Hall and Chan, 2004b; Lear and Rosenthal, 2006). The correlation between Li/Ca and temperature reveals a remarkable linear fit for temperatures above 5  C (Lear et al., 2010). Furthermore, because it is possible that the Mg/ Ca values might have been affected by non-temperature-related effects, since modern foraminifera living in poorly saturated waters have lower Mg/Ca than expected from global temperature calibrations (Lear et al., 2010), combining the Mg/Ca with the Li/Ca signature may reveal significant information. Temperature dependence was also reported for U/Ca ratios in foraminifera tests (Russell et al., 1996; Yu et al., 2008). Planktonic foraminifera studied by Yu et al. (2008) showed strong correlation between U/Ca and temperature, allowing the reconstruction of sea surface temperatures by combined U/Ca and Mg/Ca analyses. Russell et al. (1994, 1996) stated that variations in U/Ca ratios in foraminifera tests are a function of temperature effects as well as the seawater uranium content, and Boiteau et al. (2012) interpreted the latter relationship as sedimentary redox changes caused by organic carbon flux and bottom water oxygen distribution. For paleosalinity reconstructions, independent geochemical proxies for salinity still have to be developed. To assess paleosalinity, it is possible to use an independent temperature proxy, such as Mg/Ca, to correct the d18Ocarbonate signal for temperature, so that residual d18O variations reflect changing salinity (Corrège and De Deckker, 1997; Guildersen and Pak, 2005). In addition, foraminifera assemblages can provide information about salinity changes (Lear et al., 2010). 6.2. Paleoproductivity The reconstruction of paleoproductivity is based on d13C or Cd/ Ca ratios. The carbon isotope signatures of foraminifera are a function of the nutrient cycling and the exchange of CO2 with the atmosphere (Rosenthal et al., 1997a; Makou et al., 2010). The d13C signal in foraminifera tests reflects the ocean surface water productivity, but is further influenced by water mass distributions and the respiration of organic matter (Bickert and Mackensen, 2004; Mackensen and Licari, 2004). A proposed relationship of benthic foraminiferal d13C values to the bottom water nutrient content was rejected by Lea (1995). To extract the atmospheric signature on the carbon isotopic composition of seawater, coupled cadmium and d13C data can be used (Boyle, 1988; Rosenthal et al., 1997b). Cd/Ca ratios of shallowdwelling foraminifera species reflect the vertical distribution of nutrients and allows the calculation of changes in total dissolved

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inorganic carbon (Lea, 1995; Makou et al., 2010). Burton and Vance (2000) considered that cadmium behaves like the nutrient PO4 and thus indicates fluctuations in the nutrient distribution. Van Geen et al. (1992) used Cd/Ca ratios to reconstruct changes in upwelling intensity in a coastal environment. Rosenthal et al. (1997a) reported the correlation between the Cd/Ca ratio of the foraminifera test and the Cd concentration of the ambient water, which reflects surface water nutrient concentration and hence productivity or deepwater circulation, respectively. 6.3. New paleoenvironmental proxies in foraminifera research Recently, a wide range of radiogenic isotopes and trace elements in foraminifera were investigated for their suitability as paleoenvironmental proxies. Rosenthal et al. (1997b) tested F/Ca ratios, but could not find a depth, temperature, or salinity dependence. F/ Ca ratios of benthic foraminifera are generally constant throughout different calcification depths and thus differing temperatures, but they show significant differences between species, governed primarily by biological processes (Rosenthal et al., 1997b). Another focus was on the reconstruction of ocean circulation based on information about water mass distribution from proxies that mimic nutrients (d13C, Cd/Ca) or rates of flow (radiogenic isotopes) (Vance and Burton, 1999). For the differentiation of water masses that have indistinguishable nutrient signals, radiogenic isotopes, e.g., Nd, Pb, or Hf, can be used since they exhibit spatial variability due to spatial variability in their continental sources (Vance and Burton, 1999). Nd isotopes allow identifying different sources of neodymium and thus variations in ocean circulation and continental input (Palmer, 1985; Vance and Burton, 1999; Burton and Vance, 2000). Klevenz et al. (2008) used Nd isotopes to identify varying contributions of northern- vs. southern-sourced deepwater at the studied location and thus changes in deepwater hydrography. Paleo-oceanic circulation was reconstructed by Pomiès et al. (2002), and Vance et al. (2004) used Nd isotopic signatures in planktonic foraminifera to assess continental weathering, material flux from the continent to the ocean, and ocean circulation. Information about ocean circulation and flow rates are given by two insoluble products of uranium decay: 231Pa (Protactinium) and 230 Th (Thorium) (Yu et al., 1996; Walter et al., 1997; Henderson et al., 1999). Uranium is present in constant proportion to salinity in seawater, thorium is insoluble and removed quickly to the sea floor, whereas protactinium is more soluble and can be advected by circulation. The Pa/Th ratio hence reflects Pa transport pathways, therefore water mass distribution and flow rates (Walter et al., 1997). Sedimentary fluxes of biogenic barite (BaSO4) as well as Th, Pa, and Be have been used to assess past productivity (Kumar et al., 1995). Biological productivity also prefers the uptake of light isotopes of trace metals such as Zn, Fe, and Mo (Beard et al., 1999; Maréchal et al., 2000). The d15N values reflect the degree of nitrate utilization in water; the higher d15N is, the more completely nitrate has been used (Henderson, 2002). Alkalinity and pH can be reconstructed using foraminiferal Ba/ Ca and d11B, respectively. Lea (1995) reported a remarkably linear fit between the Ba/Ca ratio of foraminifera tests and alkalinity. In addition, periods of higher Ba/Ca seem to be associated with glacial intervals, thus the authors suggested variations in the nutrient composition as additional controlling factor. Hall and Chan (2004a) used Ba/Ca ratios as a proxy for water mass distribution and nutrient input, and Lea and Spero (1994) reconstructed nutrient and alkalinity distributions using Ba/Ca ratios. Hönisch et al. (2011) reported that Ba/Ca in foraminifera tests is related to the Ba/Ca concentration of the host waters and that environmental parameters including pH, temperature, salinity, and symbiont

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photosynthesis have no effect on the Ba incorporation into planktonic foraminifera. The use of boron isotopes or B/Ca ratios to reconstruct changes in pH was shown by numerous authors (Sanyal et al., 1996; Lemarchand et al., 2000; Pearson and Palmer, 2000; Sanyal et al., 2000). Boron occurs as two species in seawater whose relative concentration is pH dependent: B(OH)3 and the 20& isotopically lighter B(OH)-4. However, only B(OH)-4 is incorporated into marine carbonate, thus the d11B in carbonates changes with B speciation and therefore with pH (Lea, 1995; Pearson and Palmer, 2000). A small biological effect on the boron incorporation was stated by Sanyal et al. (2000), but its influence on the correlation between the d11B composition of the foraminifera test and seawater pH is negligible. On longer time scales, however, changes in the marine boron isotope budget may mask the pH dependency of the d11B signal (Lemarchand et al., 2000). Hathorne et al. (2009) associated variations in the B/Ca concentration with the vertical migration of the foraminifera and corresponding changes in pH and temperature. A temperature effect on the boron incorporation was also found by Yu et al. (2007). Roberts et al. (2012) investigated the suitability of rare earth element (REE) ions for paleoenvironmental studies. All studied REE concentrations were positively correlated, e.g., neodymium, cerium, and lanthanum. Neodymium isotopes are suitable for reconstructing source and flow direction of water masses (Vance and Burton, 1999; Pomiès et al., 2002; Klevenz et al., 2008). The REE3þ ions precipitate between the inner layers of foraminiferal calcite and are mostly associated with coatings rather than incorporated into the calcite structure. The adsorption and complexation of REEs may be related to the oxidation of organic matter between inner calcite layers of foraminifera, causing carbonate dissolution and providing carbonate ions, which increase REE3þ complexation and adsorption (Tang and Johannesson, 2005; Johnstone et al., 2010). The REEs are also strongly coupled to Fe and Mn cycling because of their strong affinity for ferromanganese oxides (Sholkovitz et al., 1992). High REE concentrations are consistent with increased alkalinity and MnCO3 precipitation (Roberts et al., 2012), while Tang and Johannesson (2005) observed increased REE adsorption with higher pH, as more carbonate ions were available for complexation. The mobility of Mn and Fe (also U and Ce) ions is controlled by the oxygen concentration in the pore waters (Froelich et al., 1979; Boiteau et al., 2012), which allows reconstructing changes in the pore water environment. High Fe and Mn concentrations measured in planktonic foraminifera coincided with periods of high U enrichment, consistent with oxide dissolution under reducing conditions and re-mobilization of redoxsensitive ions (Roberts et al., 2012). Under sub-oxic pore water conditions, authigenic MnCO3 precipitates between the inner calcite layers of planktonic foraminifera, preventing remobilization of the REE ions under more reducing conditions (Roberts et al., 2012). 7. Limitations in foraminiferal geochemistry In foraminifera, similar limitations are present as in ostracod shell chemistry. Significant intratest and intertest variations, for example in Mg/Ca ratios, in various planktonic foraminifera species have been reported (Brown and Elderfield, 1996; Barker et al., 2005). But to what extent this compositional heterogeneity influences the reproducibility and the achievable precision and accuracy of seawater thermometry has not been rigorously assessed. In an interlaboratory comparison study of Mg/Ca and Sr/Ca ratios in three carbonate reference materials, Greaves et al. (2008) found many uncertainties and showed that interlaboratory variability is dominated by inconsistencies among instrument calibrations. They showed that repeatability of Mg/Ca determinations increased with

decreasing Mg/Ca (0.78% at Mg/Ca ¼ 5.56 mmol/mol (CMSI 1767), 0.82% at Mg/Ca ¼ 3.76 mmol/mol (ECRM 752-1) and 1.15% at Mg/ Ca ¼ 0.79 mmol/mol (BAMRS3), respectively) and that interlaboratory reproducibilities were noticeably worse than intralaboratory precision. Nevertheless, the uncertainty decreases when increasing numbers of foraminifera tests are analyzed per sample. If at least 20 tests are analyzed, the uncertainty reduces to less than 1  C (Barker et al., 2003; Anand and Elderfield, 2005). Mg/Ca variability is also influenced by a range of site-specific biological and environmental factors: the amplitude of seasonal and inter-annual temperature changes, differences in sedimentation rate and depth of bioturbation, also redox conditions in the water column as well as in the pore waters. The intratest Mg/Ca variability results from the presence of calcite layers featuring varying Mg/Ca composition, and high- and low-Mg/Ca layers within different chambers of the same foraminifera test (Sadekov et al., 2008). Differences occur also between morphotypes. For example, G. ruber pyramidalis shows consistently lower Mg/Ca composition than G. ruber ruber (Sadekov et al., 2008) because of differing seasonal growth preferences or habitat depth. Other limitations arise by diagenetic alteration as well as pre-treatment methods (Boyle, 1981; Boyle, 1983). A decrease in carbonate saturation level with depth leads to alteration of Mg/Ca in foraminifera by postdepositional dissolution on the sea floor (Dekens et al., 2002). A detailed study on diagenesis in foraminifera and their effect on stable isotope and trace element composition is given by Sexton et al. (2006). Diagenesis causes a reduction in Mg, possibly by removal of a Mg-rich zone that is present within the tests of many planktonic species (Rosenthal and Boyle, 1993). It was also found that benthic foraminifera are generally more resistant to dissolution than planktonic ones (Izuka, 1988; Lear et al., 2000). 8. Geochemical proxies in other micro- and macrofossils In the marine environment a variety of organisms exist that have potential as paleoclimate archives, e.g., corals and mollusks (see Table 2). Corals are widely used to assess past sea surface temperature (SST), especially as they allow resolving intra-annual to weekly variations (Wurster and Patterson, 2001; Montagna et al., 2007). Isotope and trace element studies in mollusk shell carbonates have been used to reconstruct paleoenvironments (Rosenthal and Katz, 1989; Anadón et al., 2002; Schöne et al., 2011), not only in the marine realm but also in continental waters. In addition, biogenic silica from diatoms or sponges has been used for paleothermometry (Leng and Marshall, 2004; Ellwood et al., 2006; Baines et al., 2011). Jochum et al. (2012a) assessed paleotemperatures by the Mg/Ca and d18O composition in spicules from the deep-sea sponge Monorhaphis chuni. Besides their use for paleothermometry, silicon isotopes and germanium composition were also measured in sponges as well as in diatoms to reconstruct inorganic germanium and silicon concentrations in seawater (Ellwood et al., 2006; Hendry and Robinson, 2012). 8.1. Trace elements in corals as a paleoenvironmental proxy Cold water corals have been used as geochemical archives in deep sea research (Hart and Cohen, 1996; Smith et al., 1997) since they inhabit a wide variety of depths and geographic distributions, including ocean basins and high-latitude regions (Case et al., 2010). Their aragonite skeletons record changing deep-water composition and past circulation patterns. Oxygen stable isotope values can be used to reconstruct mean annual temperatures (Emiliani et al., 1978; Hart and Cohen, 1996; Linsley et al., 2004), while intraannual variability is overprinted by biological fractionation (Smith et al., 1997). Coralline Sr/Ca is a useful proxy for changes in coral

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growth environment offering intra-annual resolution, hence both seasonal and inter-annual variability can be assessed (Marshall and McCulloch, 2002; Gagan et al., 2004; Linsley et al., 2004). The Mg/ Ca, Sr/Ca, U/Ca, B/Ca and F/Ca ratios show seasonal cycles consistent with sea surface temperatures (Hart and Cohen, 1996; Sinclair et al., 1998; Linsley et al., 2000; Montagna et al., 2007), but Mg and B also show intra-annual fluctuations that could not be explained by temperature alone (Fallon et al., 1999). The linear correlation for U/ Ca and Sr/Ca to SST seems to be limited to temperatures above 18  C, as in times of extreme cold, an increased elemental incorporation was observed (Fallon et al., 1999). Marshall and McCulloch (2002) reported that thermal stress at extremely warm or cold temperatures can result in lower SST estimates, because the biological control on Sr/Ca fractionation breaks down. Hart and Cohen (1996) limited the use as paleothermometer to the surface water genus Porites, because deep-sea corals seem more affected by vital effects. Cohen et al. (2006) also reported non-equilibrium fractionation of strontium caused by vital effects. In contrast to the temperature dependence that dominates the Mg/Ca and Sr/Ca ratios, Hemming et al. (1998) attributed changes in the B/Ca signature of Porites to changes in seawater pH or productivity and upwelling. Investigations by Montagna et al. (2006) and LaVigne et al. (2008) revealed a linear correlation between seawater dissolved inorganic phosphate (DIP) concentrations and P/Ca ratios in corals, allowing the reconstruction of past seawater nutrient concentrations. Paleoproductivity can also be assessed using Ba/Ca ratios in corals and can thus give information about seasonal upwelling, e.g., that induced by winds (Lea et al., 1989; Fallon et al., 1999), as high Ba concentrations occur in deep, cold and nutrient-rich waters. In contrast, Montagna et al. (2007) found a relationship between Ba/ Ca ratios and sea surface temperatures, but the Ba/Ca ratios may also be sensitive to river runoff. Other possible trace element proxies, e.g., Cd, Pb, Mn, Zn, and V, were studied by Shen and Boyle (1988), but only Pb and Cd showed clear temporal variability associated with industrialization, and in the case of Cd with natural perturbation in ocean circulation. Lewis et al. (2007) studied land-use changes in a river catchment using Ba, Y, and Mn in Porites. Ba and Y are attributed to soil erosion caused by livestock. Yttrium concentrations rise with agricultural intensification as they show a positive correlation with cattle numbers, and thus indicate increasing sediment flux due to landuse changes and prolonged droughts. The authors stated that Y is more reliable than Ba/Ca ratios, as Ba/Ca seems to be influenced by additional factors. Mn is also clearly related to changes in catchment land-use and related erosion (Lewis et al., 2007). Other studies trying to trace anthropogenic impact used a wide range of heavy metals (Al, Fe, Cu, Cr, Hg, V and Zn) as well as Pu isotopes (Imai and Sakanoue, 1973; Bastidas and García, 1999). REEs have been studied in coral aragonite as they could provide information about river water discharge, rainfall and weathering (Scherer and Seitz, 1980; Sholkovitz and Shen, 1995; Wyndham et al., 2004). Concentration of REEs are low in surface waters and increase with depth, thus they provide information about the upwelling of sub-surface seawater. Additionally, suboxic and anoxic seawater and pore waters are characterized by high REE concentrations and Ce anomalies, allowing the reconstruction of past redox conditions. And in coastal corals, REEs also provide information about river water discharge and weathering, as water from terrigenous sources are quite different in REE/Ca from the local seawater (Sholkovitz and Shen, 1995), and about biologic activity and productivity (Wyndham et al., 2004). However, a number of potential complicating factors can limit the use of coralline aragonite. Corals precipitate their aragonitic skeleton out of equilibrium with seawater, resulting in a decrease of Sr, U, d18O and d13C, and an enrichment in Mg relative to abiotic

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aragonite (Bar-Matthews et al., 1993; De Villiers et al., 1995; Hendy et al., 2007; Meibom et al., 2007). Biological processes that fractionate trace elements in corals originate from two different types of aragonite. Sinclair et al. (2006) reported Mg-rich opaque centers of calcification and U-rich large translucent crystals, with the Mgrich material strongly depleted in d18O and d13C. In addition the centers of calcification show enriched trace element concentrations (Meibom et al., 2006). Varying proportions of these crystal types account for most of the variability in d18O and d13C. The disequilibrium offset of, for example, skeletal Sr/Ca from seawater Sr/Ca can vary between corals of the same species, thus only relative changes in SST over time can be reconstructed accurately. Additionally, intra-axis variability occurs, since significantly higher Sr/ Ca ratios are associated with slower skeletal extension rates (Hart and Cohen, 1996; Adkins et al., 2003; Cohen et al., 2006; Meibom et al., 2007). It has been shown that sampling along the coral maximum growth axis is best suited, because off axis sections have higher Sr/Ca ratios, apparently due to smaller polyp size and lower skeletal density (Cohen et al., 2001; López Correa et al., 2010). Cohen et al. (2006) suggested that a combination of temperaturedependent partitioning and changes in the saturation state of the calcifying fluid can account for varying Sr/Ca and Mg/Ca ratios in corals. Additional limitations are related to skeletal dissolution and secondary aragonite overgrowth affecting trace element ratios. Secondary aragonite infillings and overgrowth also influence d18O and d13C signatures, both resulting in cooler temperature reconstructions (Hendy et al., 2007). Contrary to the prevailing opinion that diagenesis occurs increasingly with increasing age, Hendy et al. (2007) reported that postdepositional artifacts can occur in very recent coral skeletons, e.g., a decade old, whereas century-old skeletons have been found to be nearly intact. 8.2. Mollusk geochemistry Mollusk shells, especially bivalves and gastropods, are mainly used to assess paleotemperatures and paleosalinities in marine or brackish water environments. The advantage on using mollusks instead of foraminifera is the preservation of seasonal and even daily variability (Dutton et al., 2002; Schöne et al., 2005). Stable isotope and trace element profiles along the external growth line provide high-resolution records and the possibility to assign calendar years to geochemical proxies, if the growth rate can be calculated. The d18O values have been shown to provide information about sea surface temperatures (Gillikin et al., 2005; Schöne et al., 2005; Carroll et al., 2009; Batenburg et al., 2011) as the shell is precipitated in isotopic equilibrium with ambient water (Elliot et al., 2009). In lake systems the d18O signature can be used to reconstruct temperatures, but it also reflects the oxygen isotopic composition of the lake water (Brand et al., 1993; Von Grafenstein et al., 1994; Wurster and Patterson, 2001). In this way, isotope changes in precipitation resulting from variations in air temperature can be traced (Von Grafenstein et al., 1992; Von Grafenstein et al., 2000). A relationship between d18O in the mollusk shell and salinity could not be found by Dutton et al. (2002). In contrast, Carroll et al. (2009) reported the possible influence of seasonal changes in salinity, occurring after the spring melt. In coastal environments d18O signals can be biased by lower d18O values of meteoric water, suggesting warmer temperatures at lower salinities. But there are also limitations to the use of d18O as differences in growth rate can bias the temperature signal (Tripati et al., 2001; Elliot et al., 2009; Apolinarska, 2013). Carbon isotope signatures in bivalves and gastropods are a reliable proxy for paleoproductivity. Fluctuations in the molluscan d13C value are interpreted as changes in the carbon isotope

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composition of the dissolved inorganic carbon (DIC) and reflect water mixing and primary productivity (Gasse et al., 1987; Elorza and Garcìa-Garmilla, 1996; Wurster and Patterson, 2001; Gillikin et al., 2006; Carroll et al., 2009). Dutton et al. (2002) stated that maxima in d13C coincide with visible growth lines and can be attributed to periods of high productivity. In addition, the d13C curves agree very well with intra-annual fluctuations in chlorophyll a, and thus are correlated to primary productivity (Schöne et al., 2005). Only few studies reported a temperature dependence of Mg/Ca (Lazareth et al., 2003; Batenburg et al., 2011) and Sr/Ca ratios (Dodd, 1965). According to Gillikin et al. (2006), many trace elements that are commonly used in foraminiferal or coral research, such as Sr, Mg and U, cannot be used as environmental proxies in bivalves (Gillikin et al., 2006). For Mg/Ca paleothermometry, contradictory results led to the suggestion that Mg/Ca incorporation in the bivalve shell is mainly controlled by biological processes (Carré et al., 2006; Elliot et al., 2009). The salinity dependence of Sr/Ca and Mg/Ca ratios was studied by Dodd and Crisp (1982) with the result that these ratios can be used as salinity proxy, but only if the salinity does not exceed 8&. Although Sr/Ca ratios are useful paleotemperature proxies in corals and sponges, they seem to be mainly controlled by growth rate in aragonitic bivalves (Dutton et al., 2002; Carré et al., 2006). Gillikin et al. (2005) tested whether the strong correlation between growth rate and temperature allows indirect temperature predictions from bivalve Sr/Ca ratios. Using two different species they found a significant correlation between growth rate and Sr/Ca ratio for Saxidomus giganteus but no correlation for Mercenaria mercenaria. They suggested that variability in Sr/Ca ratios is mainly caused by biological processes and not under thermodynamic control. Later Carroll et al. (2009) published that the primary driver of bivalve growth is mainly food supply, rather than temperature. Carré et al. (2006) also showed that environmental parameters have no significant influence on the incorporation of Mn, Mg, Ba and Sr into the bivalve shell, and that crystal growth rate accounts for most if the trace metal variability (e.g., 74% variance in Sr). Temperature and salinity had only a minor influence on trace element incorporation, and there is also no correlation between the element/calcium ratios of the bivalve shell and seawater (Carré et al., 2006). Elliot et al. (2009) found no correlation between Mg and Sr incorporation and growth rate in Tridacna gigas bivalves and suggested varying metabolic rates as the cause for varying shell Mg/ Ca. And also Shirai et al. (2008) stated that biological processes account for micro-scale elemental distributions, as the elemental composition was significantly associated with sublayer types in Bathymodiolus platifrons shells. Schöne et al. (2013) also found large differences between Sr/Ca and Mg/Ca ratios from simultaneously deposited regions in shells of Arctica islandica and related them to different crystal fabrics, or to the processes that control their formation. Only at the annual growth lines were Sr and Mg deposited in equilibrium with the ambient environment. In contrast, strontium isotopes seem to be a reliable proxy to trace different water masses. Sr isotopes reflect the Sr isotope ratio of the ambient water and allow to separate different strontium sources. Water mass exchange between open ocean and marginal marine environments (Whittaker and Kyser, 1993), freshwater input, and water exchange between separated ocean basins (Tripati et al., 2001) could be reconstructed. In continental environments, saline influxes into lake systems (Anadón et al., 2002) as well as differing hydrochemistry during two lake phases, caused by differences in the relative contributions of water sources (Fan et al., 2010), were assessed. Promising results were reported also for bivalve Ba/Ca ratios, as this proxy is highly reproducible between specimens. Ba/Ca shell

profiles show a flat background signal interrupted by sharp Ba/Ca peaks (Gillikin et al., 2006; Schöne et al., 2013). Given a direct relation between the background and the Ba/Cawater, the Ba/Ca ratios of bivalve shells can be used as proxies for changes in salinity due to the inverse relation between Ba/Cawater and salinity (Gillikin et al., 2006). A temperature dependence of the Barium incorporation into the bivalve shell was not observed. Elliot et al. (2009) found that the Ba/Ca peaks correlate very well with chlorophyll peaks in time and amplitude, indicating the association of these peaks with seasonal increases of local primary productivity. In contradiction, Gillikin et al. (2008) reported that Ba/Ca peaks start around 40 days after the crash of a bloom and thus the blooms could not be the cause. Lazareth et al. (2003) suggested the possibility to use Ba and Mn as paleoproductivity proxies, since Ba and Mn concentrations were directly related to nutrient input during plankton blooms and increased run-off of freshwater. Carré et al. (2006) found no such correlation, as periodicity was lacking in their Ba and Mn profiles, and thus rejected the relation to periodic plankton blooms. Caroll et al. (2009) reported that up to 60% of the annual variability in Ba/Ca ratios could be explained by river discharge, but only at sites close to rivers, as this effect became negligible at farther distances. The possible application of Mn/Ca ratios as indicator for primary production was also suggested by results of Carroll et al. (2009). Complicating factors that may limit the use of trace elemental ratios and stable isotopes in bivalves are similar to those that affect all micro- and macrofossils (Elorza and Garcìa-Garmilla, 1996; Wurster and Patterson, 2001). Finding an appropriate pretreatment prior to analysis is essential also for bivalves. Krause-Nehring et al. (2011) tested the effects of different chemical pretreatment methods with respect to changes in elemental composition, the efficiency to remove organic matter, and the possibility to cause alteration, and found that no pretreatment is without any side effects. They suggested avoiding chemical treatment prior to analysis whenever possible. 9. Techniques in paleoceanography and paleolimnology The most routine analysis in paleoceanography as well as paleolimnology is batch dissolution. Techniques for geochemical analysis of trace elements in foraminifera, mollusks or corals that are becoming more widely used are Laser Ablation Inductively Coupled Plasma Mass Spectrometry (LA-ICP-MS) (Sinclair et al., 1998; Eggins et al., 2003; Mertz-Kraus et al., 2009; Sadekov et al., 2009; Hetzinger et al., 2011; Jochum et al., 2012b), electron probe microanalysis (EPMA) (Cohen et al., 2001; Morishita et al., 2007; Pena et al., 2008; Hathorne et al., 2009), and Nano Secondary Ion Mass Spectrometry (NanoSIMS) (Cohen et al., 2006; Kozdon et al., 2009; Fehrenbacher and Martin, 2010; Jochum et al., 2012a; Hoppe et al., 2013). These methods allow rapid analysis of solidstate samples as well as simultaneous measurement of a wide range of trace elements with high sensitivity (Montagna et al., 2007; Hetzinger et al., 2011). In foraminifera research, laser ablation ICP-MS was used to remove diagenetically modified surface calcites by pre-ablation prior to each analysis (Eggins et al., 2003; Pena et al., 2005). NanoSIMS analysis in corals revealed variations in Mg associated with microstructure zoning (Meibom et al., 2007). A few studies have applied these methods to the analysis of ostracod shells (Morishita et al., 2007). Jochum et al. (2012b) carried out laser ablation ICP-MS on Leucocytherella sinensis valves and summarized the detection limits of a large number of trace elements. With respect to the proxies in foraminifera, elements such as Nd and other REEs are detectable and can thus also be tested as proxies in lacustrine environments. Laser ablation analysis and NanoSIMS can also be used to obtain trace element distribution

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11. Future challenges Summarizing all existing techniques shows that ostracod shell chemistry provides an important paleoenvironmental proxy. However, environmental variables such as salinity, alkalinity, continental weathering, and atmospheric circulation cannot be assessed entirely by a single proxy. More work is required to better understand existing proxies and their interactions. Factors controlling shell chemistry of various genera are very complex, and underlying chemical and biological processes of most of these new proxies are not well understood, even for living ostracods. Future studies should aim at better constraining the sensitivity of different proxies for different ostracod species. In order to assess paleoenvironmental signals, paired trace elemental and stable isotope proxies should be standard. Acknowledgments Fig. 1. U/Th mapping on a 600 mm long ostracod shell (Leucocytherella sinensis) from Tibet determined by laser ablation ICP-MS at the Max Planck Institute for Chemistry, Mainz, using a spot size of 12 mm.

patterns in ostracod valves, e.g., U/Th mapping using LA-ICP-MS (Fig. 1). For example, Sohn and Kornicker (1969) suggested that calcium is evenly distributed through the ostracod shell, while phosphorous is concentrated near the shell margins. But the distribution of most elements is still unclear. Detailed mapping of ostracod shells would thus also give new insights in possible biological controls on the element incorporation in the shells. 10. Conclusions Shell chemistry in ostracods and other micro- and macrofossils such as foraminifera and corals show many similarities even though the extent of research in the latter groups is clearly much greater. Techniques for geochemical analysis such as laser ablation ICP-MS, NanoSIMS, and electron probe microanalysis (EPMA), which are state of the art methods in paleoceanography, provide great potential for application in ostracod research, as they offer high-resolution data and thus the ability to show elemental variation associated with shell microstructures (Shirai et al., 2008). An important question is, which proxies can be transferred from the marine to the lacustrine environment and hence from foraminifera or coral chemistry to ostracods. If similar processes control the element/calcium partitioning in ostracods and, for example, foraminifera, it may be possible to transfer the fundamental basis of the proxies from marine to lacustrine environments. In addition, the assessment of elemental detection limits will determine what is even possible technically. Transferring existing proxies from marine to lacustrine environments is complicated as the factors that control environmental conditions differ. Differences also occur between open and closed lake basins, respectively. Thus the applicability of a new proxy has to be tested case-by-case. Li/Ca ratios as recorder of carbonate saturation and temperature, as well as paleonutrient proxies such as Cd/Ca may be used in closed lake basins. Closed lake systems are primarily influenced by precipitation-to-evaporation ratios and thus climate. In open lake systems the water balance is additionally dependent on inflowing and outflowing water masses. The potential of redox sensitive trace elements providing information about bottom water oxygenation should be exploited in all marine and continental waters.

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