From FISCHER G, WEFER G (eds), 1999, Use of Proxies in Paleoceanography: Examples from the South Atlantic. Springer-VerlagBerlin Heidelberg, pp 489-512

Reassessing Foraminiferal Stable Isotope Geochemistry: Impact of theOceanic Carbonate System (Experimental Results)

J. Bijma1*, H.J. Spero2 and D.W. Lea3

1Universität Bremen, Fachbereich Geowissenschaften, Postfach 33 04 40,D-28334 Bremen, Germany

2Department of Geology, University of California Davis, Davis,California 95616, U.S.A.

3Department of Geological Sciences and the Marine Science Institute, University ofCalifornia Santa Barbara, Santa Barbara, California 93106, U.S.A.

* corresponding author (e-mail): jbijma@uni-bremen. de

Abstract: Laboratory experiments with living planktic foraminifers show that the δ13C and δ18Ovalues of shell calcite decrease with increasing sea water pH and/or carbonate ion concentration.The effect has been quantified in symbiotic (Orbulina universa) and non-symbiotic (Globigerinabulloides) species and is independent of symbiont activity and temperature. It is concluded that akinetic fractionation process affects both the carbon and oxygen isotopic composition of the shellsimultaneously. At present it cannot be determined definitively whether the relationship is controlledby the pH dependent balance between hydration and hydroxylation of CO2 or by [CO3

2 -] relatedvariations in the calcification rate. However, independent of which factor ultimately controls therelationship between the carbonate chemistry and isotopic fractionation, in the real ocean [CO3

2 -]and pH covary linearly across the relevant pH range. The true relationship between shell isotopiccomposition and the bulk carbonate chemistry is masked by the fact that host respiration andsymbiont activity locally modify the carbonate system. Respiration lowers and photosynthesisincreases ambient pH and [CO3

2 -]. This translates into modified absolute shell values but leaves theslope between the shell isotopic composition and the bulk carbonate chemistry unaffected. A se-cond level of shell isotopic modification is introduced by the incorporation of respired carbon,enriched in 12C, which depletes the shell δ13C value. In symbiont bearing species this depletion ispartially negated by a shell δ13C enrichment in the light. As an alternative to the RUBISCO hypothesis(enrichment via preferential removal of 12CO2), we propose that scavenging of respired CO2 duringphotosynthesis, raises the shell δ13C value. Our results have partly been documented before (Speroet al. 1997) and demonstrate that the carbonate chemistry is undoubtedly a major control on temporalgeochemical variability in the fossil record. For instance, the sea water carbonate system of the pre-Phanerozoic world (Berner 1994; Grotzinger and Kasting 1993) or during glacials (Sanyal et al. 1995)was significantly different from today confounding direct interpretation of foraminiferal stable isotopedata using existing relationships (see companion paper in this volume by Lea et al.).

IntroductionSpero et al. (1997) describe a series of experimentsdemonstrating that planktic foraminiferal δ18O andδ13C are influenced by sea water carbonate chem-istry. In this paper, we describe these experimentsin greater detail, discuss previously unpublished

experiments and speculate on the potential mecha-nisms controlling the observed responses.

Based on a theoretical study of the thermody-namic properties of isotopic substances, Urey (1947)proposed that the 18O content of calcium carbon-

490 Bijma et al.

ate could be used as a paleothermometer for theocean. The method is based on the fact that seawater and calcite differ in their 18O/16O ratios,when in thermodynamic equilibrium, and that thisdifference decreases with increasing temperature.McCrea (1950) demonstrated in the laboratory thatsynthetic calcium carbonate when precipitated inequilibrium with sea water is enriched in 18O rela-tive to sea water but less so with increasing tem-perature and thus behaves as predicted by Urey.This empirical relationship was the first, so called,paleotemperature relationship. Another milestonewere the papers by Epstein et al. (1951, 1953) whowere the first to show that marine invertebratesgrown in the natural environment deposit calciumcarbonate in equilibrium with sea water. Many newpaleotemperature equations have been developedsince this early pioneer work (for a review seeBemis et al. 1998) but the problem in determiningthe "true" paleotemperature is not only one of choos-ing the correct equation but, more importantly, oneof choosing the correct value for δ18O of sea wa-ter (δ18Ow). For the geologic past, this amounts toknowing paleosalinity and the relationshipbetween salinity and δ18Ow (e.g. Berger 1979;Broecker and Van Donk 1970; Fairbanks et al.1992; Rozanski et al. 1993; Schrag et al. 1996;Shackleton 1967, 1974).

Beyond the glacial ice volume effect and theeffect of precipitation/evaporation, so called "vitaleffects" were recognized. Contrary to Emiliani’s(1954) interpretation that the oxygen isotopic dif-ferences among planktic foraminifers reflect differ-ences in depth habitat and hence differencesin precipitation temperature, Duplessy et al. (1970)argued that the differential oxygen isotopic fraction-ation in benthic foraminifers from the same level ina core requires a different explanation and theysuggested that metabolic CO2 depleted in 18O isincorporated into the shells of benthic and possiblyalso planktic foraminifers. The hypothesisthat planktic foraminifers have different18O-fractionation factors was subsequently con-firmed by Shackleton et al. (1973) and later sub-stantiated by numerous other studies (e.g. Curry andMatthews 1981a; Duplessy et al. 1981; Fairbankset al. 1980; Shackleton 1974; Vergnaud-Grazzini

1976; Williams et al. 1979). For instance, Buchardtand Hansen (1977) note that benthic foraminifersbearing symbiotic algae are depleted in 18O rela-tive to symbiont barren species and Erez (1978b)argues that with greater photosynthetic activity,incorporation of light metabolic CO2 results inanomalously light carbon and oxygen isotopic com-positions of hermatypic corals and benthicforaminifers. However, when dealing with differ-ences in δ18O through time, vital effects have gen-erally been assumed to be constant (or zero) andwere therefore neglected.

The interpretation of the carbon isotopic com-position proved to be more complicated than δ18O.Due to the preferential fixation of 12CO2 by primaryproducers and subsequent transport and decompo-sition of organic matter below the euphotic zone,surface waters of the ocean are depleted in 12Cwhereas the deeper ocean is 12C-enriched. Thus,as a consequence of community production andrespiration, δ13C depth profiles are related to theoxygen concentration and inversely correlated withthe nutrient concentration (the stoichiometric pro-portionality constant with PO4 is ca. 0.93 ‰ permol,kg-1: Broecker and Peng 1982). Hence theδ13C of surface dwelling foraminifers is used as asurface water fertility proxy and the difference inδ13C between planktics and benthics (∆δ13C) wasintroduced as a measure of the strength of the bio-logical pump (e.g. Broecker 1971; 1973; 1981).Another source for variation of δ13C in dissolvedinorganic carbon (δ13CΣCO2) is the fractionationbetween carbon in atmospheric CO2 and the totaldissolved CO2 (ΣCO2) in surface ocean water. The13C/12C ratio in atmospheric CO2 is on average 9‰ lower than that in surface ocean ΣCO2 and theextent of fractionation depends on temperature (e.g.Broecker and Maier-Reimer 1992; Charles et al.1993; Lynch-Stieglitz et al. 1995). As a result, highlatitude surface water have a higher 13C/12C ratiothan low latitude surface water. Generally, this ther-modynamic imprint is not very strong due to veryslow isotopic equilibration (Broecker and Peng1982), however, if there is sufficient timefor gas exchange, a significant signal is produced(Charles and Fairbanks 1990). A third source ofvariation is the so called "Mackensen effect".

Reassessing Foraminiferal Stable Isotope Geochemistry 491

Mackensen et al. (1993) hypothesised that growthand reproduction of benthic foraminifers coincideswith the seasonal flux of phytodetritus. Becauseδ13C of freshly accumulated phytodetritus was 3to 4 ‰ lower than surface sediment organic car-bon they argued that δ13C of epibenthic shells,formed during remineralisation of this fluffy layer(locally and temporarily depleting δ13CΣCO2), is de-pleted compared to δ13CΣCO2 found throughout therest of the year.

The main complication associated with the in-terpretation of δ13C records from foraminiferalshells is a prominent vital effect. Already Craig(Revelle and Fairbridge 1957) had shown thatthe carbon isotopes in foraminiferal shells are pre-cipitated out of equilibrium with δ13C of ΣCO2 andpotentially reflect utilization of metabolic CO2. Thelarge range of disequilibrium precipitation was con-firmed by many studies (for a review see e.g.Berger 1979; Duplessy 1978). A milestone forthe investigation of vital effects on stable isotopicfractionation was the establishment of culture tech-niques for planktic foraminifers by the researchgroups of Allen Bé, Roger Anderson, ChristophHemleben and Michael Spindler (e.g. Bé 1979; Bé1980; Bé 1982; Bé and Anderson 1976; Bé et al.1979; Bé et al. 1977; Hemleben and Spindler 1983;Hemleben et al. 1989; Hemleben et al. 1985;Hemleben et al. 1987; Spindler and Hemleben1980). Using laboratory cultures, Erez and Luz(1983) derived the first empirical paleotemperatureequation for Globigerinoides sacculifer and, forinstance, Bouvier-Soumagnac and Duplessy (1985)were able to verify their plankton tow basedtemperature:δ18O relationships. Culture experi-ments were also imperative to investigate the im-pact of algal symbiosis on the shell δ13C composi-tion of benthic (Erez 1978b; Williams et al. 1981b;Zimmermann et al. 1983) and planktic foraminifers(Spero and DeNiro 1987). In contrast to benthicforaminifers and some corals (Erez 1978b; Landet al. 1977), Spero and Deniro concluded thatincreased symbiont photosynthetic activity inOrbulina universa, resulted in a 13C enrichmentof the shell. Their conclusion was supportedby other studies (Cummings and McCarty 1982;Goreau 1977; Weber and Woodhead 1970)

and later verified for other species of plankticforaminifers (Bijma et al. 1998; Spero and Lea1993; Spero et al. 1991). The enrichment was at-tributed to the preferential removal of 12CO2 by theCO2 fixing enzyme ribulose 1,5-biphosphatecarboxylase-oxygenase (RUBISCO) during sym-biont photosynthesis, enriching the calcifying envi-ronment in 13C. Host respiration was shown todeplete the shell isotopic composition (e.g. Bijmaet al. 1998; Spero and Lea 1996). This observationwas explained by the contribution of metabolic CO2,enriched in 12C, to the calcifying environment.

During the past four decades, stable oxygen andcarbon isotope measurements on biogenic calciteand aragonite (for a review see e.g. Wefer andBerger 1991) have become standard tools for re-constructing paleoceanographic and paleoclimaticchange. By assuming that our understanding of themajor parameters controlling stable isotope incor-poration into biogenic calcite is complete, thesegeochemical proxies have been used to reconstructglacial ice volumes (e.g. Fairbanks 1989; Mix 1987),sea surface (SST) and deep water temperatures(e.g. Zahn and Mix 1991), ocean circulation changes(e.g. Boyle 1990; Charles and Fairbanks 1992;Labeyrie et al. 1987) and glacial-interglacial shiftsbetween the terrestrial and oceanic carbon pools(Shackleton 1977). However, the list of factors con-trolling the stable isotopic compositions of shellsis still incomplete. Data from isotopic calibrationstudies have suggested that an unidentified param-eter affects carbon and oxygen isotope ratios inforaminiferal shells (e.g. Duplessy et al. 1981;Fairbanks et al. 1982; Williams et al. 1979). Al-though shell δ13C values may be affected by vitaleffects, δ18O values should be insensitive to physi-ology. Yet, oxygen isotope data from foraminiferscollected in surface plankton tows yield SST recon-structions that are several degrees warmer thanmeasured ocean temperatures (Williams et al.1981a). Similarly, the covariation of carbonand oxygen isotopes in planktic foraminifers (e.g.Berger et al. 1978; Curry and Matthews 1981a;Curry and Matthews 1981b; Kahn and Williams1981) cannot be explained by shifting depthhabitats during ontogeny. In terms of equilibriumfractionation the trends are opposite to those found

492 Bijma et al.

in the water column. If foraminifers were frac-tionating in isotopic equilibrium lightest δ18O val-ues should be associated with heaviest δ13C val-ues. Stable isotope studies on other calcifying or-ganisms such as calcareous algae, hermatypiccorals and echinoderms show similar disequilibria(McConnaughey 1989b; Wefer and Berger 1991)suggesting that these proxies are also equivocal andthat some environmental or biological parameter(s)in addition to those already identified must be af-fecting skeletal stable isotope values. Interestingly,laboratory and field data show that skeletalδ18O and δ13C covary (McConnaughey 1989a;McConnaughey 1989b; Spero et al. 1997; Speroand Lea 1996) suggesting the mechanism that af-fects shell 18O/16O ratios also has an effect on 13C/12C.

In this paper we examine the impact of changesin carbonate chemistry on the stable carbon andoxygen isotope composition of a symbiont bearingand a symbiont barren species of plankticforaminifers and subsequently discuss potentialequilibrium and kinetic fractionation mechanisms.The implications of the observed isotope effectsfor paleo-oceanographic and -climatic reconstruc-tions are discussed in the companion paper by Leaet al. (this volume).

Our initial hypothesis was that glacial symbiontbearing foraminifers should be depleted in 13C rela-tive to their Holocene counterparts because dis-solved CO2 availability (the CO2aq concentration)was reduced in the glacial ocean. Since the δ13Cof algae increases as less CO2aq is available(Farquhar et al. 1982) and [CO2aq] was reducedin the glacial ocean, we predicted that symbiontfractionation and hence 13C enrichment of the cal-cifying environment was reduced during glacialtimes. We found that symbiont bearing foraminifersgrown at lower [CO2aq] are indeed depleted in 13Cbut for a different reason than we initially hypo-thesised.

Material and MethodsDuring the summers of 1993 through 1996 cultureexperiments were conducted at the Wrigley Envi-ronmental Science Center on Santa Catalina Island(California) to investigate the influence of sea

water carbonate chemistry on the δ13C and δ18Oof the shells of the symbiont bearing plankticforaminifer Orbulina universa (d’Orbigny) andthe symbiont barren Globigerina bulloidesd’Orbigny.

Orbulina UniversaOrbulina universa has a tropical to temperatedistribution in the euphotic zone. Each adultindividual is associated with 3,000 to 7,000(Spero and Parker 1985) dinoflagellate symbionts(Gymnodinium beii). The symbionts are distrib-uted in a halo around the calcitic shell (cf. Fig.1a).Orbulina universa, like most spinose species,is carnivorous and feeds primarily on calanoidcopepods at a rate of one to two per day (Spindleret al. 1984). New chambers are added to the ex-isting trochospire at regular intervals. The transi-tion to the adult stage is marked by the formationof a spherical chamber, a feature that is unique inthis species. This terminal chamber is secretedaround the thinly calcified trochospiral juvenileshell. At this stage numerous large pores facilitatethe migration of the symbionts along the spines. Thespherical chamber continues to calcify for a periodof 1-7 days (Spero 1988) before spine resorptionand gametogenetic (GAM) calcification signal im-pending gametogenesis. GAM calcification pro-duces 13-28 % of the shell mass over a period ofseveral hours (Bé 1980; Bé et al. 1983; Spero1986). After gametogenesis the terminal sphereconstitutes between 90 to 95 % of the calcite byweight. Frequently the spiral stage is dissolvedcompletely. Reproduction of this species is tunedto the lunar cycle (Bijma et al. 1990).

Globigerina BulloidesGlobigerina bulloides (Fig. 1b) is a non-symbi-otic species that is typically associated with tem-perate to sub-polar water masses but is also char-acteristic in lower latitudes upwelling environments(Bé and Hutson 1977; Bé and Tolderlund 1971;Naidu and Malmgren 1996a; 1996b). In these re-gions, G. bulloides often dominates the flux to theocean floor (Sautter and Thunell 1989; Sautter andSancetta 1992) and is therefore an important source

Reassessing Foraminiferal Stable Isotope Geochemistry 493

Fig. 1. Orbulina universa withsymbionts (A; scale bar = 300 µm),Globigerina bulloides (B; scale bar= 300 µm), protruding membrane inO. universa (C; scale bar = 300 µm).

494 Bijma et al.

of geochemical information for paleoceanographicreconstructions (Bard et al. 1987; Kallel et al. 1988;Sautter and Thunell 1991). Contrary to most spinosespecies, an important part of its diet consists of algaeas evidenced by the olive green to brownish col-oration of the cytoplasm of freshly collected speci-mens. Recently, a study on the population dynam-ics (Schiebel et al. 1997) demonstrated a lunarperiodicity of reproduction in this species as well.As evidenced by SEM, gametogenic calcificationis much less pronounced than in O. universa, butthe extent of GAM calcite addition has yet to bequantified.

Culture ProtocolThree sets of experiments were designed to inves-tigate the effect of the carbonate chemistry on shellisotopic fractionation. Culture procedures followedstandard protocols (e.g. Bemis et al. 1998;Hemleben et al. 1989). Briefy, O. universa and G.bulloides were collected by scuba divers fromsurface waters of the San Pedro Basin, SouthernCalifornia Bight, U.S.A. and were maintained inlaboratory culture at 17 °C and 22 °C (± 0.2 °C).Specimens were grown in 0.8 µm filtered sea wa-ter (FSW) whose carbonate chemistry was modi-fied. Sea water ΣCO2 and total alkalinity (alkt)were modified via the addition of Na2CO3 and/ortitration with HCl or NaOH. The pH of the waterwas determined potentiometrically whereas alkt andΣCO2 were determined by titration and equilibriumcalculations respectively. Coulometric determi-nations of a few sea water samples confirmedthe accuracy of the calculated ΣCO2 values0(A. Sanyal, pers. comm. 1994). During the experi-ments, Southern California Bight ambient surfacewater pH was 8.13 ± 0.02 (at 22 °C) and ambient(alkt was 2241 ± 19 µeq/kg (n=64). CalculatedΣCO2 was 2010 ± 18 µmol/kg, while[CO3

2 -] was calculated to have varied between153-184 µmol/kg.

The first two sets of experiments were designedto vary [CO3

2 -] at constant alkt or at constantΣCO2. Because [CO3

2 -] and pH covary linearlyacross the relevant pH range in sea water (e.g. 7.9-8.8), these two experiments did not allow us tomechanistically distinguish between pH and

[CO32 -] as a controlling factor. Therefore we de-

signed a third set of experiments at constant pH.The basic difference in the carbonate chemistrybetween the three experimental groups is that atconstant alkt and constant ΣCO2, the [CO3

2 -] de-creases as [CO2aq] increases whereas at constantpH both carbon species covary (Fig. 2).The [CO3

2 -] of the culture water was modified by:1) elevating total alkalinity (alkt) to constant levelsof 2842 ± 80 µeq/kg (n=29) and letting pH andΣCO2 vary (Fig. 3a, b). We chose to run the ex-periments at elevated alkt because removing inor-ganic carbon quantitatively by acidifying the me-dium is much more difficult than adding ΣCO2 oralkalinity. With an alkt >2800 µeq/kg, [CO3

2 -] ex-ceeding 600 µmol/kg could be achieved at ambientΣCO2.2) keeping ΣCO2 constant at 2032 ± 15 µmol/kg;(n = 15) and letting pH and alkt vary (Fig. 3c, d andFig. 4a, b). We chose to work with a slightly el-evated experimental ΣCO2 for the same reason asabove.3) keeping pH constant at 8.15 ± 0.05 and lettingΣCO2 and alkt vary. For [CO3

2 -] above ambient,we increased ΣCO2 by adding Na2CO3 and sub-sequently titrated with HCl to bring the pH back to8.15. For [CO3

2 -] below ambient, however, we re-moved inorganic carbon by acidifying the medium

Fig. 2. Relationship between [CO32 -] and [CO2aq] at con-

stant alkalinity (broken line), constant ΣCO2 (solid line)and constant pH (straight broken line).

Reassessing Foraminiferal Stable Isotope Geochemistry 495

Fig. 3. Effect of CO32 - ion concentration on the δ13C and δ18O values of Orbulina universa shell calcite under con-

stant alkalinity (a and b), constant ΣCO2 (c and d) and constant pH (e and f) conditions. Circles indicate specimensgrown at 22 °C and squares indicate specimens grown at 17 °C. Open symbols are mean values of specimens grownunder high light (HL), closed symbols are specimens grown in the dark. Plotted data are group mean values ± 1 s.d;most groups are composed of between 3 to 10 individual shell analyses. Lines are linear regressions fitted to thedata.

496 Bijma et al.

to a calculated intermediate pH (Wolf-Gladrowpers. comm. 1994) at which the ΣCO2 attained avalue that would yield the desired [CO3

2 -] at pH8.15. After reaching the intermediate pH, the head-space was replaced by a N2 atmosphere to pre-vent uptake of CO2 while raising the pH to 8.15via addition of NaOH.

The specimens were transferred to sealed 125ml acid-cleaned, air-tight Pyrex™ jars containingthe modified culture water and fed an one-day-oldartemia nauplius every third day (total time inculture ≈ 6-8 days). For O. universa, specimenswere maintained in the dark and under non-photoinhibitory Pmax light levels of 400-700 µE m-2

s-1 (1 Einstein = 1 mole photons) to quantify theeffect of symbiont photosynthesis on this relation-ship (Spero and Parker 1985). Following game-togenesis, the empty foraminiferal shells were ar-chived for analysis. In O. universa, individualspherical chambers were used for isotope analy-sis. Globigerina bulloides chambers secreted dur-ing the experiments were amputated from the shellwhorl and pooled according to chamber position(Spero and Lea 1996). The samples were roastedat 375 °C in vacuo and analyzed with a FisonsOptima isotope ratio mass spectrometer using anIsocarb common acid bath autocarbonate systemat 90 °C.

To confirm the stability of the carbonate chem-istry and isotopic composition of the culture waterthroughout each experiment, water samples werecollected at the start and end of each experimentto analyse alkalinity, ΣCO2, δ

13CΣCO2 and δ18Owater.

Cytoplasmic Inorganic Carbon Pool

Fig. 4. Effect of CO32 - ion concentration on the δ13C (a) and δ18O (b) values of Globigerina bulloides chamber

calcite under constant ΣCO2 conditions. Chambers from specific positions in the shell whorl are pooled andanalysed together to eliminate the effect of ontogenetic isotope variability characteristic of this species (Spero andLea 1996) Linear regression analyses are plotted for data from chambers 12 and 13, the last two chambers in theshell whorl.

To explain the stable carbon isotopic signature in avariety of calcifying organisms, a so called inter-nal inorganic carbon pool has been proposed (e.g.Erez 1978a; Erez 1978b; Goreau 1963; Goreau1977; Kuile et al. 1989; Weber and Woodhead1970). To investigate the presence of such a cyto-plasmic inorganic carbon pool in O. universa,pulse-chase experiments were designed. Freshlycollected O. universa were fed and placed in 13C-spiked solution (after a rinse through a 13C-spikedmedium). The specimens were monitored at regu-lar intervals to catch the moment of membraneprotrusion (Fig. 1c) At this stage, the specimenswere rinsed three times in normal FSW, transferredto unspiked culture water and allowed to calcify thespherical chamber (see also Lea et al. 1995).Specimens were kept in the dark or at Pmax lightlevels and fed every third day. After gametogen-esis, the specimens were archived and processedas described above.

Reassessing Foraminiferal Stable Isotope Geochemistry 497

A spike stock-solution was prepared by dissolv-ing 7 mg of 99.1 % Na2

13CO3 in 100 ml deionizedwater. From this stock-solution 35 ml were addedto 1 l FSW (2028 µMol ΣCO2; δ

13CΣCO2 ≈ 2 ‰).Initial and final samples of the 13C-spiked culturemedium were measured, indicating that the aver-age spike was 77.8 ‰ ± 5.2 ‰ (n=11). The pulse-and chase-solution for each experiment were takenfrom the same batch of modified sea water andtherefore had identical carbonate chemistry. The[CO3

2 -] varied between 153 and 517 µmol/kg at anaverage constant alkalinity of 2563 ± 371 µeq/kg.

ResultsWater samples collected at the start and end ofeach experiment showed that the carbonate chem-istry (as determined by pH, alkt and equilibriumcalculations) remained constant throughout theexperiments. The isotope chemistry (δ13CSCO2 andδ18Owater) of the culture medium remained fairlyconstant as well. Analyses show that δ18Owater wasconstant at -0.23 ± 0.05 ‰ (vs. VSMOW) (n=100).The δ13CSCO2 depended on the amount of Na2CO3that was added to modify the carbonate chemistryin each experiment. Analysis of initial and finalwater show that δ13CSCO2 varied on average 0.16± 0.10 ‰ (n=30).

All foraminiferal δ13C data have been standard-ized to a δ13CΣCO2 = 2.00 ‰ (ambient sea waterδ13CΣCO2 = 1.90 ± 0.08 ‰, n=18) by adding or sub-tracting the difference between 2.00 ‰ and themeasured experimental water δ13CΣCO2 to eachshell value. Here δ18O (and similar for δ13C) =[(18O/16Osample/

18O/16Ostandard) - 1] x 1000. All iso-tope values are relative to VPDB unless noted. Toconvert δ18Ow values from the VSMOW to VPDBscale, the most recent correction of -0.27 ‰ (Hut1987) was used.

Constant Total Alkalinity ExperimentResults demonstrate that O. universa shell δ13Cand δ18O values decrease approximately 3.9 ‰and 1.5 ‰ respectively as carbonate increasesfrom 41 to 642 µmol/kg (pH increases from 7.38to 8.83) (Fig. 3a, b). Shell δ13C values do not in-crease further when sea water [CO3

2 -] drops be-

low ca. 100 µmol/kg. Comparison of foraminifersgrown under high light (HL = maximum symbiontphotosynthesis) with specimens maintained in thedark (no symbiont photosynthetic activity) indicatesthat HL shells are enriched in 13C by 1.1 ‰ anddepleted in 18O by 0.3 ‰ relative to dark speci-mens. These offsets are seen across the full rangeof [CO3

2 -]. Although the effect of symbiontphotosynthesis on shell δ13C and δ18O was docu-mented previously (Spero and DeNiro 1987; Speroand Lea 1996; Spero and Williams 1988), the factthat the δ13C/[CO3

2 -] and δ18O/[CO32 -] slopes

from HL and dark experiments are ca. 0.0065 and0.002 ‰/µmol kg-1 respectively (Table 1) indicatesthat symbiont photosynthesis does not affect thestable isotope:carbonate ion relationship.

Six experiments were carried out at a 5°C lowerculture temperature. Although, the temperaturedecrease does not seem to significantly affect theδ13C/[CO3

2 -] nor the δ18O/[CO32 -] relationship,

additional experiments are needed to furtherbolster this conclusion.

Constant Total Inorganic CarbonExperimentIn a second series of experiments, we manipulatedthe carbonate concentration between 75 - 774 µmol/kg (pH increase from 7.87 to 8.97) by varyingalkt and maintaining ΣCO2 at ambient values(Fig 3c, d). Orbulina universa shell δ13C and δ18Odecreased as [CO3

2 -] increased, with similar slopesas under constant alkt: 0.006 and 0.002 ‰/µmol kg-

1 respectively (Table 1). Again, comparison offoraminifers grown under HL with specimensmaintained in the dark (no symbiont photosyntheticactivity) indicates that HL shells are enriched in 13Cby 1.1 ‰ and depleted in 18O by0.4 ‰ relative to dark specimens and that theseoffsets are constant across the full range of[CO3

2 -].We also cultured the non-symbiotic species,

G. bulloides, across a [CO32 -] range of 103-436

µmol/kg at constant ΣCO2. Because G. bulloidesdisplays a large chamber-to-chamber ontogeneticeffect for both δ13C and δ18O (Spero and Lea1996), we amputated chambers secreted during the

498 Bijma et al.

experiments and pooled them according to theirdiscrete positions in the shell whorl for stable iso-tope measurement (Fig. 4a, b). Regression analy-ses on data from the 11th, 12th and 13thchamber groups yield similar δ13C/[CO3

2 -] andδ18O/[CO3

2 -] slopes demonstrating that ontogeneticeffects have little or no affect on this relationship.However, the average G. bulloides δ13C/[CO3

2 -]and δ18O/[CO3

2 -] slopes of 0.013 and 0.0045 ‰/µmol kg-1 are twice that of O. universa (Table 1).

Constant pH ExperimentResults demonstrate that shell δ13C and δ18O val-ues of O. universa grown in the dark are approxi-mately constant (2.0 ‰ and -1.75 ‰ respectively)as carbonate increases from ambient to 780 µmol[CO3

2 -]/kg (ΣCO2 increases from 748 to 9,273µmol/kg; alkt increases from 875 to 10,073 µeq/kg)(Fig. 3e, f). If anything, there is an insignificant butdistinct trend towards slightly higher values for δ13C

and δ18O with increasing [CO32 -] in the dark. Com-

parison of foraminifers grown under HL (maximumsymbiont photosynthesis) with specimens main-tained in the dark indicates that HL shells are en-riched in 13C by ca. 1.5 ‰ at ambient [CO3

2 -], butthat this enrichment decreases as [CO3

2 -] increases.Because this offset is not constant across the fullrange of [CO3

2 -], symbiont photosynthesis doeshave an effect on the δ13C/[CO3

2 -] relationshipunder constant pH. On the contrary, the δ18O val-ues in the light are slightly depleted compared tothose in the dark and the offset (≈ 0.1 ‰) is moreor less constant across the full range of [CO3

2 -].

Table 1. Experiment regression slopes.

Species Experiment Regression Slope

δ13

32C

CO[ ]−δ18

32O

CO[ ]−

Orbulina universa Constant Alkalinity High Light 0.006 0.002

Dark 0.007 0.002

Constant ΣCO2 High Light 0.006 0.002

High Light* 0.006 0.0012

Dark 0.006 0.002

Globigerina bulloides Constant ΣCO2 12th Chamber 0.014 0.005

13th Chamber 0.012 0.004

11th Chamber* 0.013 0.005

12th Chamber* 0.012 0.004

13th Chamber* 0.013 0.004

* Data from a feeding experiment not described here

Inorganic Carbon Pool (Pulse-Chase)ExperimentAfter a mean incubation time of two days in the"pulse" solution (δ13CΣCO2 = 77.8 ‰ ± 5.2 ‰), thespecimens were transferred, upon membrane pro-trusion, to a "chase" solution with the same carbon-

Reassessing Foraminiferal Stable Isotope Geochemistry 499

ate chemistry as the "pulse" solution, fed as usualand left to calcify the adult chamber and undergogametogenesis. It can be calculated that the smallamount of the "pulse" solution that is enclosed bythe membrane and inevitably transferred to thechase culture medium does not affect the shell δ13Ceven if all of this "enclosed" carbon is used forcalcification (cf. Lea et al. 1995). The δ13C val-ues of the spiked specimens were slightly higherthan those of the control group (Fig. 5a-d) but massbalance calculations demonstrate that less than 1% of the carbon used for calcification originatesfrom the spike and hence that, if present at all, theinorganic carbon pool is insignificant.

Fig. 5. Effect of pre-culture labeling (δ13CΣCO2 = 77.8 ‰) on the δ13C and δ18O values (under constant alkalinityconditions) as a function of CO3

2 - ion concentration (squares). For comparison, the data of the constant alkalinityexperiment (Fig. 2 a,b) are also plotted (circles). Open symbols (a, b) are mean values of specimens grown underhigh light, (HL); closed symbols (c, d) are specimens grown in the dark. Plotted data are group mean values ± 1 s.d.The lines are linear regressions fitted to the data.

DiscussionBecause the sea water carbonate chemistry sys-tem was modified in fundamentally different ways(varying proportions of ΣCO2 and alkt) yet pro-duced indistinguishable regression slopes, the ob-served stable isotope:carbonate ion relationship isreal and not an artifact of our experiments. In ad-dition, identical δ13C and δ18O values of pre- andpost-gametogenic shells (Spero et al. 1997) dem-onstrate that the relationships are a function ofcalcification during normal growth and not due toinorganic precipitation on the surface of the emptypost-gametogenic shell.

500 Bijma et al.

Fig. 6. Mass balance calculations (Mook et al. 1974;Zhang et al. 1995) show the redistribution of isotopesbetween the dissolved carbon species as a function ofpH.

The experiments carried out with the symbiontbarren G. bulloides as well as the dark experi-ments with O. universa clearly demonstrate thatwhatever controls the stable isotope:carbonatechemistry relationship is independent of symbiontactivity. Experiments carried out at a 5°C lowertemperature further suggest that this relationshipis also independent of temperature. In addition, anunexpected stable oxygen isotope:[CO3

2 -] relation-ship (also independent of symbiont activity and tem-perature) was found that cannot be explainedin terms of known "vital effects". These observa-tions force us to reassess previously proposedfractionation mechanisms.

In the following sections we will first discussinorganic equilibrium and kinetic fractionationmechanisms. We will then discuss the influence ofrespiration and photosynthesis and finally proposea combination of mechanisms that control the iso-topic composition of symbiont barren and symbi-ont bearing foraminifers.

Equilibrium and Kinetic FractionationMass balance calculations (Mook et al. 1974;Zhang et al. 1995) show that the δ13C:carbonatechemistry relationship documented here cannot beexplained in terms of equilibrium fractionation, i.e.by a simple redistribution of isotopes between thedissolved carbon species (Fig. 6). The influence oftemperature (17 vs. 22 °C) on the isotopic redistri-bution between the carbon species is only discern-ible in the δ13C composition of CO2. The δ18O sig-nal of the carbon species is insensitive to changesin the carbonate chemistry because the water res-ervoir is huge compared to the ΣCO2 reservoir.

With regard to kinetic fractionation, the earlyinorganic precipitation experiments of McCrea(1950) are of particular interest. He demonstratedthat the 18O/16O ratio of rapidly precipitated CaCO3decreases with increasing percentage CO3

2 - in so-lution. McCrea proposed that the oxygen isotopeequilibrium fractionation between the sum of thedissolved carbonate species and water is linearlyrelated to the equilibrium between the carbonatespecies, a conclusion that was, more than 40 yearslater, confirmed by Usdowski and Hoefs (1993).Comparison of our experimental O. universa data

with those of McCrea show remarkable similarityin the slopes of the relationships (Fig. 7). We notethat the larger G. bulloides δ18O:[CO3

2 -] relation-ship demonstrates that additional fractionationmechanisms are operative. McCrea proposed thatthe δ18O:[CO3

2 -] relationship for inorganic precipi-tates could be a function of calcification rate. Thisconclusion is not supported by our observationsbecause: 1) Above ambient [CO3

2 -], the averageshell weight, in both the dark and HL, is fairly con-stant (albeit variable) with increasing [CO3

2 -]. Be-cause the life span of O. universa was similarbetween and within treatments, the average calci-fication rate must have been relatively constant inthe [CO3

2 -] range above ambient as well. On thecontrary, below ambient [CO3

2 -], the calcificationrates in the dark and in HL decrease significantlywith decreasing [CO3

2 -] (Fig. 8). Thus, if it is as-sumed that changes in the calcification rate con-trol the δ18O fractionation we would expect a clearbreak in the shell δ18O:[CO3

2 -] relationship at am-bient [CO3

2 -]. On the contrary, the shell δ18O:[CO3

2 -] relationship decreases linear across the full[CO3

2 -] range. 2) The difference in shell weight,and by extension the difference in calcification rate,between the HL and dark experiments (cf. Lea etal. 1995) is larger than the difference in shell weight

Reassessing Foraminiferal Stable Isotope Geochemistry 501

between ambient and high [CO32 -] (Fig. 8). Con-

sequently, if the fractionation is controlled by thecalcification rate, the light/dark effect on δ18Oshould be larger than the [CO3

2 -] effect. On thecontrary, the light/dark shift in δ18O is on the aver-age -0.35 ‰ for O. universa and constant acrossthe full range of the experimental [CO3

2 -] while thedepletion is approximately 1.5 ‰ over a [CO3

2 -]range of ca. 600 and 700 µmol/kg at constantalkt and constant ΣCO2, respectively. Shells ofG. bulloides are depleted by 1.4 ‰ across a[CO3

2 -] range of only 333 µmol/kg (at constantΣCO2). Again, this suggests that the calcificationrate does not control 18O incorporation. 3) Finally,the final shell weight of O. universa, ranging be-tween 10 to 70 µg, is very variable between andwithin treatments (Fig. 8). Thus, if the calcificationrate is the controlling factor, the correlation be-tween the oxygen isotopic composition of the shelland the percentage carbonate ion in solution (Fig.7) should be very low. At this stage we thereforehave to conclude that the δ18O:[CO3

2 -] relationshipfor inorganic and biological precipitates is probablynot a function of calcification rate as proposed byMcCrea. However, because we find it difficult tobelieve that the rate at which a calcifying organ-ism accretes CaCO3 is not a fundamental factorfor isotope fractionation, additional experiments willhave to be carried out to confirm this hypothesis.It should also be noted that although the similarityin life span of O. universa between and withintreatments can be used to roughly quantify calcifi-cation rate (because spherical chamber formationis a continuous process) it cannot be applied to G.bulloides because chambers are formed discretely.

More importantly, the similarity between ourinvestigation and the inorganic precipitate study ofMcCrea implies a kinetic mechanism which wouldaffect all calcifying organisms and inorganicprecipitates. Evidence for a common kinetic frac-tionation mechanism exists in corals and other in-vertebrate groups. For instance, McConnaughey(1989b) demonstrated that the carbon and oxygenisotopic composition of symbiotic and non-symbiotic coral skeletons can vary significantlyacross skeletal surfaces that were presumablysecreted synchronously. In the non-photosyntheticspecies, Tubastrea, the values covary with a posi-

Fig. 7. Comparison of high light (open circles) and dark(closed circles) O. universa δ18O data from the constantalkalinity experiment and high light G. bulloides δ18Odata (open triangle) from the constant ΣCO2 experimentwith inorganic precipitate results from McCrea (1950).Oxygen isotope data are plotted vs. percentage CO3

2 - insolution where: % CO3

2 - = [CO32 -] / ([CO3

2 -] + [HCO3- ]).

Note that the slopes of the O. universa data are indis-tinguishable from that of the McCrea data set but thatthe slope of the G. bulloides data is much steeper.

Fig. 8. Shell weights of O. universa plotted as a func-tion of CO3

2 - ion concentration. Chamber mass data forG. bulloides were inconclusive and are therefore notshown. The shaded area represents the range of ambi-ent [CO3

2 -]. Circles, triangles and squares represent theconstant alkalinity, ΣCO2 and pH experiment, respec-tively. Diamonds indicate the pulse-chase experiments.Open symbols are mean values of specimens grownunder high light (HL), closed symbols are specimensgrown in the dark. Linear regressions were fitted to databelow and above ambient [CO3

2 -]; HL data (solid line)and dark data (broken line).

502 Bijma et al.

tive δ18O/δ13C slope of 0.29. The foraminiferalδ18O/δ13C slopes derived from the experimentaldata are virtually identical to Tubastrea, with val-ues ranging between 0.29 and 0.33 (Fig. 9). Otherahermatypic corals, calcareous algae and inverte-brates such as cidaroid urchins show similaroxygen and carbon isotope covariance(McConnaughey 1989b; Wefer and Berger 1991)although the δ18O/δ13C slopes can differ from theexperimental range reported here. For instance,in symbiont-bearing organisms such as the coralPavona, the slope can approach unity or evenchange sign due to the additional 13C-enrichingeffect of symbiont photosynthesis (McConnaughey1989b). Aside from organisms that display a largesymbiont effect, the slope relationship similaritiesamong different protozoans, invertebrate groupsand some calcifying algae suggest that a commonkinetic mechanism is responsible for the observedcovariance. Moreover, the linearity between δ18Oand δ 13C suggests that whatever controls thefractionation of oxygen also affects carbon.

McConnaughey (1989a) has proposed that thefractionation observed in Tubastrea occurs duringCO2 hydration and hydroxylation to form HCO3

-.

The rates of hydration and hydroxylation are pHdependent and are the rate-limiting stepsresponsible for the kinetic discrimination againstthe heavier 18O and 13C isotopes in the calcifyingmicroenvironment adjacent to the skeleton.McConnaughey further argues that CO2 hydrationand hydroxylation reactions may exhibit differentkinetic isotope effects and that the balance be-tween these two reactions changes with pH. Sub-sequent isotopic equilibration between HCO3

- andCO3

2 - is effectively instantaneous because theprotonation of CO3

2 - and the deprotonation ofHCO3

- are extremely fast. With regard to δ18O,McConnaughey’s prediction is in agreement withresults from inorganic precipitation experiments(Usdowski et al. 1991).

pH or Carbonate Ion ControlOur experiments allow us to mechanistically dis-tinguish between pH and [CO3

2 -] as controllingfactors for the observed fractionation. The fact thatunder constant pH, δ13C and δ18O of shells kept inthe dark do not vary with increasing [CO3

2 -]suggest that pH, as proposed by McConnaughey(1989a), controls the stable isotope fractionationand not [CO3

2 -]. This conclusion is supported by ex-periments that were recently carried out with G.sacculifer at constant pH (Bijma, Spero and Leaunpubl. results). Fig. 10a and b demonstrate that,at constant alkalinity, below pH ≈ 8 the fractionationis approximately constant (the CΣCO2 experi-ments suggest a similar trend). Apparently, the par-titioning of the hydration and hydroxylation does notchange very much in that pH range. One couldargue that, although insignificant, the dark constantpH experiments show a trend towards higher val-ues for both δ13C and δ18O at higher [CO3

2 -]. Atpresent we have no explanation for this observa-tion. It can neither be justified by a redistributionof isotopes between the dissolved carbon speciesas pH is more or less constant (Fig. 11c) nor by achange in the relative proportion of light metabolicCO2 incorporation (because the δ18O differencebetween respired and bulk CO2 is negligible as willbe demonstrated later).

An intriguing question is why the shell stableisotope composition responds linear to [CO3

2 -]

Fig. 9. Comparision of foraminiferal carbon and oxygenisotope data from all experiments with data from the non-symbiotic coral, Tubastrea spp (McConnaughey 1989b).The similarity of the positive covariance between coral(aragonite) and foraminiferal (calcite) stable isotopessuggest that similar mechanisms are responsible for theobserved variations. In this plot, [CO3

2 -] and pH increasetowards the origin.

Reassessing Foraminiferal Stable Isotope Geochemistry 503

(Figs. 3,4) but not to pH (Fig. 10)? Intuitively onewould conclude that kinetic fractionation is appar-ently controlled by the [CO3

2 -]. However, in a non-equilibrium situation where the kinetic fractionationis dependent on the balance between CO2hydration and hydroxylation as suggested byMcConnaughey (cf. Fig. 6 in McConnaughey1989a) and both reactions exhibit different but con-stant kinetic isotope effects, the pH-δ13CΣCO2 (cf.Fig. 6) and the pH-δ18OΣCO2 relationships are notconstant but apparently decrease with increasingpH (Fig. 10a-d). Because the pH/[CO3

2 -] relation-ship in the constant alkalinity and the constantΣCO2 experiments is also non-linear (Fig. 11a, b),the [CO3

2 -]:stable isotope relationships in these ex-

periments turn out to be linear. On the other hand,as argued before, we believe that the calcificationrate (for instance, the time available for isotopicequilibration between CO3

2 - and solid CaCO3 be-fore subsequent layers prevent isotopic exchange)should play an important role and is likely to dependon the saturation state (Ω) and hence on the[CO3

2 -]. Independent of which factor ultimatelycontrols the relationship between the carbonatechemistry and isotopic fractionation, the depend-ence may be expressed as a function of pH or[CO3

2 -]. For obvious reasons we prefer the linearrelationship with [CO3

2 -]. It should also be noted,that in the real ocean [CO3

2 -] and pH covary lin-early across the relevant pH range.

Fig. 10. Effect of pH on the δ13C and δ18O values of Orbulina universa shell calcite under constant alkalinity(a and b) and constant ΣCO2 (c and d). Circles indicate specimens grown at 22°C and squares indicate specimensgrown at 17°C. Open symbols are mean values of specimens grown under high light (HL), closed symbols arespecimens grown in the dark. Plotted data are group mean values ± 1 s.d. Lines are linear regressions fitted to thedata.

504 Bijma et al.

Impact of Respired CO2

Because isotopes are conserved in metabolism, theextent of 13C depletion of respired CO2 must re-flect the δ13C of the carbon source and the carbonbudget of the consumer. Thus, the δ13C of respiredCO2 depends on the δ13C value of the food sourceand the δ13Corg of the foraminifer. The followingmass balance calculation may be considered:

δ13Csource = [r•δ13Cforam] + [(1-r)•δ13Cresp]

where r is the fraction of carbon retained by theforaminifer. Rearranged:

δ13Cresp = (δ13Csource - [r•δ13Cforam])/(1-r)

For symbiont bearing species it was demonstratedthat a substantial fraction (denoted s) of the host'scarbon is derived from the symbionts, probablytranslocated during the night when the symbiontsare retracted in the shell (Bemis pers. comm.). Thusfor O. universa the equation must be expanded:

δ 1 3C r e s p= ( [ s • δ 1 3C s y m b i o n t ] + [ 1 - s • δ 1 3C f o o d ]- [r•δ13Cforam])/(1-r)

δ13C of the artemia, the host and the symbiontsis a function of their biochemical composition, es-pecially of the lipid content. Corrected for thecarapace, that is discarded after digestion, theδ13C value a Great Salt Lake artemia nauplii(GSL) is -15 ‰. The δ13Corg of O. universa andG. bulloides fed GSL artemia is -17.4 ‰ and-18.5 ‰ respectively (Uhle unpubl. results). Basedon compound specific isotope analysis (Table 2) andthe assumption that lipids are 5 ‰ depleted com-pared to the average δ13C of symbiont organicmatter, a δ13Csymbiont of -24 ‰ can be estimated(Schouten pers. comm.). If we further assume that56 % of the carbon taken up by O. universa origi-nates from the GSL-artemia and that the rest isderived from the symbionts (Bemis pers comm) andthat only one tenth of the carbon taken up is re-tained, the δ13C of the respired CO2 is -19.1 ‰ and-14.6 ‰ for O. universa and G. bulloides respec-tively (Table 3). Bemis (unpubl data) determinedthat the δ13Corg of O. universa fed artemia nauplii

of a San Francisco Bay (SFB) strain was -18.6 ‰,while the artemia were on the average -20.4 ‰(corrected for the carapace). The δ13C of the re-spired CO2 under these conditions is -22.4 ‰(Table 3). Unfortunately, we have no data for G.bulloides fed SFB artemia.

Respiring organisms consume 16O16O 10 to20 ‰ more rapidly than 18O16O. With respect to

Fig. 11. Relationship between [CO32 -] and pH at constant

alkalinity (a), constant ΣCO2 (b) and, constant pH (c).

Reassessing Foraminiferal Stable Isotope Geochemistry 505

the metabolic fractionation during respiration afractionation factor of 1.016 has been proposed foroxygen (Lane and Dole 1956). Others (Kiddon etal. 1993), measured a respiratory isotope effectbetween 14 ‰ to 26 ‰ for seven representativeunicellular marine organisms and estimate that theaverage fractionation of the dominant marinerespirers is about 20 ± 3 ‰. A generalized respi-ration equation is given by:

CH2O + O2 → CO2 + H2O

If we assume a fractionation of 20 ‰ during res-piration and that the δ18O of O2 is -0.45 ‰ vs.VSMOW (equivalent to H2O) and if we furtherassume that the δ18O of the combined host/symbi-ont organic matter is 28 ‰ (Swart 1983) andthat CO2 and H2O maintain their characteristicfractionation of 1.0412 at 25 °C (Friedman andO’Neil 1977), respiration may produce a pool of18O-depleted CO2 and cell H2O. However, due tothe large cell water reservoir and the large waterfluxes across the membrane (according toMcConnaughey (1989a) 6 orders of magnitudelarger than photosynthetic or respiration fluxes) theδ18O of the cell water quickly returns to -0.45 ‰.Thus, although respired CO2 can be 13 ‰ depletedcompared to δ18O of bulk CO2, the fast isotopeexchange with cell water may completely cancelthis offset. However, once converted to CO3

2 - apart of the signal may be preserved due to the slowexchange of oxygen isotopes between dissolvedCO3

2 - and H2O (McConnaughey 1989a).In summary, incorporation of respired CO2 can

significantly lower shell δ13C but does probablynot affect δ18O. For example, Spero and Lea(1996) demonstrated that G. bulloides, when fedSFB artemia is ca. 0.5 ‰ depleted in 13C comparedto specimens fed GSL artemia but that shell δ18Owas unaffected. Thus, incorporation of respiredCO2 lowers shell δ13C but the effect is relativelysmall compared to the ca. 3 times larger photosyn-thetic effect that has been determined for O.universa and G. sacculifer. However, the calcu-lations by Spero and Lea (1996) are based on theisotopic difference of the food source. In the fol-lowing we will use the calculated isotopic differ-ence in the respired CO2 (Table 3). If all respiredCO2 would be incorporated into the shell (100 %efficiency), the shell δ13C difference of specimensfed SFB or GSL artemia should be 3.3 ‰, i.e. re-flect the full difference in the carbon isotopic com-position of respired CO2. Since the difference inshell δ13C of G. sacculifer fed SFB and GSLartemia, is on the average only 0.55 ‰ (Bijma,Spero and Lea unpubl. results) it can be calculatedthat the impact of respiration on the shell isotopicvalue is ca. 17 %. Consequently, with an efficiencyof 17 %, respiration can lower the shell isotopic

Table 2. Stable carbon isotope values of dominant freesterols. Ergosta-5,24(28)-dienol was not used for thedetermination of the average δ13C. We believe that themuch heavier isotope signature of this compound origi-nates from heterotrophic "swarmers".

Table 3. Calculated δ13C of respired CO2 for a symbiontbearing and a symbiont barren species fed two differ-ent strains of artemia (GSL = Great Salt Lake; SFB = SanFrancisco Bay).

artemia strain δ13Cresp (‰)

O. universa G. bulloides

GSL (-15 ‰) -19.1 -14.6

SFB (-20.4 ‰) -22.4 n.a.

"free sterols" δ13C

cholest-5,22-dienol -29.5 ‰ ± 0.2 ‰

cholesterol (=cholest-5-enol) -28.7 ‰ ± 0.2 ‰

ergosta-5,22-dienol -29.1 ‰ ± 0.1 ‰

ergosta-5,24(28)-dienol (-24.9 ‰ ± 0.1 ‰)

average ca. -29 ‰

506 Bijma et al.

we found a positive shift of 1.1 ‰ for O. universaand others have shown that compared to specimensgrown in the dark, δ13C of O. universa and G.sacculifer kept above saturation light levels, isenriched by 1.7 ‰ (Spero 1992) and 1.4 ‰ (Speroand Lea 1993), respectively. Consequently, if wetake the kinetic effect on δ13C and δ18O (inducedby local changes in the carbonate chemistry) intoaccount, we have to invoke an additionalmechanism to explain a concomitant enrichment inδ13C between 2.2 to 2.8 ‰ for O. universa andeven 3.4 ‰ for G. sacculifer.

RUBISCO has always been the first choice toexplain shell enrichment in the light. However, thismechanism requires that sufficient CO2 is availableto enrich the residual CO2 pool or, if HCO3

- is thecarbon source, that not all CO2 (once convertedfrom HCO3

-) if fixed by RUBISCO and that somediffuses to the site of calcification. Numericalexperiments (Wolf-Gladrow et al. 1999) have dem-onstrated that the symbionts are CO2 limited, evenat low photosynthetic rates, and therefore useHCO3

- as their main carbon source. These experi-ments have also demonstrated that a steep CO2gradient exists between the shell surface and thesymbiont halo, probably acting as an efficient bar-rier against CO2 diffusion towards the shell. As theHCO3

- pool is very large and hardly affected by pho-tosynthetic carbon fixation, and the δ13C of HCO3

-

is very close to the δ13C of ΣCO2, we argue thatRUBISCO is probably not responsible for theHL:dark shift. This conclusion is supported by ex-periments carried out at constant pH (Fig. 3e). AsCO2 availability increases, shells get lighter insteadof heavier, suggesting that enzyme mediated 13C-enrichment of the ambient environment plays asubordinate role, if at all.

As an alternative hypothesis to explain the ob-served positive shift between the HL and darkexperiments we propose scavenging of respired12CO2 by the symbionts. In the previous section wehave shown that respired CO2 can lower shell δ13Cby 3.2 to 3.8 ‰ (depending on the food source),enough to explain the required enrichment in thelight of up to 3.4 ‰. The decrease in shell δ13Cwith increasing [CO2aq] at constant pH can nowbe explained by a decrease in the scavenging effi-ciency of respired CO2 as bulk CO2 increases. Thus

Symbiont EffectOur experimental results have separated the influ-ence of symbiont photosynthesis from the influenceof the carbonate chemistry on the shell isotopiccomposition calcification. The observed 13C-en-richment and 18O-depletion in HL O. universarelative to specimens maintained in the dark areclearly a function of symbiont photosynthesisand have been reported before. With respect to13C-enrichment, a mechanism has been proposedwhereby RUBISCO preferentially removes 12CO2during photosynthesis and subsequently enrichesthe calcifying environment with 13C (Spero 1992;Spero and DeNiro 1987; Spero and Lea 1993;Spero et al. 1991; Spero and Williams 1989).

Because it is generally assumed thatphotosynthesis per se does not affect the δ18Oof the photosynthate nor that of the residualH2O-CO2-HCO3

- pool (Swart 1983), 18O-depletionin symbiotic systems were tentatively explainedas the result of light-enhanced calcificationrates (Bouvier-Soumagnac et al. 1986; Spero 1992;Spero and Lea 1993). However, as discussed be-fore, our data suggest that calcification rate maynot explain the oxygen isotope fractionation ob-served in this study. In the rest of this section wepropose an alternative mechanism based on localpH changes induced by carbon fixation of thesymbionts.

Since the symbiotic algae increase the pH of theforaminiferal microenvironment to as much as 8.8during photosynthesis (Jørgensen et al. 1985),the local [CO3

2 -] will be elevated as well (Wolf-Gladrow et al.1999). The observed differences inδ18O between HL and dark range from0.35 ‰ for O. universa (Spero 1992) to 0.63 ‰for G. sacculifer (Spero and Lea 1993). If weassume that this difference is solely due to thesymbiont induced shift in the ambient carbonatechemistry, and further that the δ18O/[CO3

2 -] and aδ13C/[CO3

2 -] slopes are -0.002 ‰ and -0.0065 ‰respectively (Table 1), a simultaneous shift of1.1 ‰ and 2.0 ‰ towards more negative δ13Cvalues at HL can be calculated for O. universaand G. sacculifer, respectively. On the contrary,

composition by 3.2 to 3.8 ‰, respectively, when fedGSL and SFB artemia.

Reassessing Foraminiferal Stable Isotope Geochemistry 507

scavenging of light respired CO2 can explain theobserved difference between the HL and dark re-sults and the trend noticed in the constant pH ex-periments. In addition, effective scavenging of lightrespired CO2 could also explain the very negativevalue of the symbiont organic matter (-24 ‰) eventhough HCO3

- is the main carbon source.The proposed mechanism of CO2 scavenging

can operate only if metabolic CO2 diffuses freelythrough the membrane of the host and is not storedin an internal inorganic carbon pool. The pulse-chase experiments demonstrate that such a pool isnot present in O. universa (Fig. 6).

Proposed Fractionation MechanismsBased on our experiments we propose that 3fractionation mechanisms can be recognized:1) pH or [CO3

2 -] dependent kinetic fractionationof both carbon and oxygen stable isotopes.2) Incorporation of respired CO2 depleted in 13C.In symbiont bearing species in the light.3) Enrichment of the calcifying environment with13C (compared to the dark) due to scavenging oflight respired CO2 (RUBISCO related effects areprobably of minor importance).

shell at higher pH. Hence, because the δ13C of re-spired CO2 is significantly depleted compared tobulk CO2, the fractional change of metabolic ver-sus bulk CO2 incorporation into the shell will rein-force the depletion. In addition, a higher light- com-pared to dark-respiration rate of the symbionts(Weger et al. 1989) may deplete the calcifyingenvironment even more under HL. The enrichmentof shell δ13C is dictated by very effective scaveng-ing of light respired CO2 (mechanism 3). A declinein the scavenging efficiency with increasing[CO2aq] could explain the trend in the constant pHexperiments (more respired CO2 escapes photo-synthetic fixation and gets incorporated into theshell).

The question may be raised why scavengingof respired CO2 dominates the impact of kineticfractionation (mechanism 1) on δ13C signal but noton the δ18O signal? The reason is twofold. First,theoretical considerations require that the impactof kinetic fractionation on 13C/12C is half that of18O/16O (McCrea 1950). Second, the relative de-pletion of respired CO2 with respect to 13C is largewhereas the depletion in 18O may be lost quicklydue to the large water reservoir compared to ΣCO2and the fast exchange reactions.

Decoupling of the δ13C/[CO32 -] Relationship

at High [CO2aq] in the LightA change in the scavenging efficiency may alsoexplain the "plateau" or the decrease in δ13C to-wards lower [CO3

2 -] in the constant alkalinity ex-periments (Fig. 3a). At higher [CO2aq] (i.e. lower[CO3

2 -]), scavenging of metabolic CO2 may be-come less effective and consequently an increas-ing portion of light respired CO2 is incorporated intothe shell. Again, the opposing affect of RUBISCOat high [CO2aq], i.e. the enrichment of the ambi-ent ΣCO2 with 13C apparently plays a subordinaterole. As noted before is the photosyntheticassimilation of carbon strongly CO2 limited (Wolf-Gladrow et al. 1999). Apparently, even at the maxi-mum PCO2 reached in the constant alkalinity ex-periments (ca. 1900 µatm.) the symbionts are stillCO2 limited. In seperate experiments (Spero, Bijmaand Lea unpubl results), PCO2 was varied between1400 µatm and more than 8000 µatm. The results

Light/dark Shift of Shell Stable IsotopesIn all experiments (Calkt, CΣCO2 and CpH),the δ18O value of shells grown in the light isdepleted compared to shells kept in the dark. Sucha decrease in shell δ18O with increasing irradiancehas been documented before for G. sacculiferand O. universa (Spero and Lea 1993). Becausemetabolic processes do not directly affect theδ18O signal of the shell, we argue that an increaseof pH or [CO3

2 -] in the calcifying environment, ei-ther due to bulk carbonate chemistry changes orinduced by photosynthesis, depletes shells in 18Ovia mechanism 1.

The enrichment of δ13C in the light observed inall experiments (Calkt, CΣCO2 and CpH) must beexplained in terms of competing processes. Firstis the depletion of shell δ13C kinetically driven bythe photosynthetically raised pH (mechanism 1).Secondly, Zeebe et al. (1999) have demonstratedthat more respired CO2 gets incorporated into the

508 Bijma et al.

at Bremen University, contribution 203) and theProgram for the Advancement of Special ResearchProjects at the Alfred Wegener Institute,Germany (JB) and by the US National ScienceFoundation (HJS and DWL). Data are availableunder www.pangaea.de/Projects/SFB261.

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(not shown here) support the hypothesis that thedecrease in shell δ13C in the light should be ex-plained in terms of a reduction in the scavengingefficiency and that RUBISCO is of secondaryimportance. The reason that the δ18O signal underHL is not decoupled from [CO3

2 -] at high PCO2(Fig. 3b) is because δ18O of metabolic CO2 andbulk CO2 are probably not much different. Conse-quently, changes in the scavenging efficiency do noteffect shell δ18O.

Conclusions1) Shifts in pH (and for natural sea water in[CO3

2 -]) have a significant impact on foraminiferalshell δ13C and δ18O.2) Both symbiotic (Orbulina universa) and non-symbiotic (Globigerina bulloides) species displaythis effect. Therefore, we hypothesize that thiseffect is common among other species.3) The effect is species dependent and based on acombination of kinetic and metabolic fractionationprocesses.4) A pH/carbonate isotope effect has importantimplications for the interpretation of stable isotopedata in the fossil record. Shells calcifying during coldclimate periods record isotopic signatures that can-not be interpreted directly using present-day rela-tionships (cf. Lea et al. this volume)5) Because the carbonate ion effect appears to bea major factor influencing the foraminiferal isotoperecord, calibration of the response in each of themajor paleoceanographic species is a necessarynext step towards fully understanding and utilizingthe oceanic isotopic record.

Acknowledgments

We thank the staff of the Wrigley EnvironmentalScience Center and E. Kincaid, C. Hamilton, J.Dailey, E. Komsky, T. Mashiotta, M. Uhle, A.Sanyal, D. Chan, E. Mochon, M. Cramer and V.Hubner for their help in the field. Thanks go to S.Schouten for the compound specific isotope analy-sis and to W. Berger and an anonymous reviewerfor many helpful suggestions. This researchwas funded by the Deutsche Forschungs-gemeinschaft (Sonderforschungsbereich 261

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