Carbon isotopes in biological carbonates: Respiration and photosynthesis

Geochimicaet CosmochimicaActa, Vol. 61, No. 3. DD.61l-622, 1997 Copyright 0 1997Eikvier Science Ltd Printed in the USA. All rightsreserved

Pergamon

0016.7037/97 jl7.00 + .OO

PII SOO16-7037(96)00361-4

Carbon

isotopes in biological

TED A. MC~ONNAUGHEY,

’JIM

carbonates:

Respiration

BLJRDETI-,2JOSEPH F. WHELAN,~

and photosynthesis

and CHARLES K. PAULL~

‘Marine Research, Biosphere 2 Center, Highway 77, PO Box 689, Oracle, Arizona 85623, USA %SGS, Box 25046, MS 963, Denver, Colorado 80225, USA ‘Geology Department and Marine Sciences Curriculum, University of North Carolina, Chapel Hill, North Carolina 27599-33 15, USA (Received January 5, 1996; accepted in revised form October 14, 1996)

Abstract-Respired carbon dioxide is an important constituent in the carbonates of most air breathing animals but is much less important in the carbonates of most aquatic animals. This difference is illustrated using carbon isotope data from freshwater and terrestrial snails, ahermatypic corals, and chemoautotrophic and methanotrophic pelecypods. Literature data from fish otoliths and bird and mammal shell and bone carbonates are also considered. Environmental C02/02 ratios appear to be the major controlling variable. Atmospheric C02/02 ratios are about thirty times lower than in most natural waters, hence air breathing animals absorb less environmental CO2 in the course of obtaining Oz. Tissue CO*, therefore, does not isotopically equilibrate with environmental CO2 as thoroughly in air breathers as in aquatic animals, and this is reflected in skeletal carbonates. Animals having efficient oxygen transport systems, such as vertebrates, also accumulate more respired CO2 in their tissues. Photosynthetic corals calcify mainly during the daytime when photosynthetic CO2 uptake is several times faster than respiratory CO2 release. Photosynthesis, therefore, affects skeletal S’“C more strongly than does respiration. Corals also illustrate how “metabolic” effects on skeletal isotopic composition can be estimated, de&e the uresence of much larger “kinetic” isotope effects. Copyright 0 1997 Elsevier Science itd L 1.

INTRODUCTION

withdraws, 13C-depleted carbon from it. These metabolic effects on the 613C of the internal DIC pool carry over to precipitating carbonates where they are superimposed on sometimes larger “kinetic” isotope fractionations (McConnaughey, 1989b). Kinetic fractionations originate from slower hydration and hydroxylation of CO2 by molecules bearing the heavy isotopes 13Cand 180. These reactions come into play when CO1 (of any origin, not necessarily respired) diffuses across the cell membrane into the confined extracellular region where calcification occurs and reacts, with kinetic fractionations, to form HCO; . Kinetic fractionations are expressed when this HCO; precipitates as CaC03 before re-establishing isotopic equilibrium with cell DIC. Figure 1 schematically illustrates the apparent mechanism of most biological calcification (McConnaughey and Falk, 1991; McConnaughey, 1994; McConnaughey and Whelan, 1996) and shows where metabolic and kinetic isotope fractionations enter the system. This study examines metabolic “C effects in biological carbonates from several perspectives. We critically review the geochemical and physiological literature and discuss the basis for distinguishing “metabolic” from “kinetic” 13C depletions (relative to 6 13Cequilibrium between the carbonate and environmental inorganic carbon ) We then estimate metabolic effects in selected animals. Finally, we present a simple respiratory gas exchange model to examine the mixing of respired and environmental CO, in the tissue DIC pool. This model is applied to air and water breathing animals and extended to examine photosynthetic 13C enrichments in coral skeletons.

The effects of respiration and photosynthesis on the 6’% values of biological carbonates are poorly understood. Especially problematic are divergent results from terrestrial and aquatic animals. Land snails, birds, and mammals produce shell and bone carbonates which isotopically reflect the animal’s diet (e.g., Goodfriend and Magaritz, 1987; Sullivan and Krueger, 1981) and must therefore contain much respired C02. Aquatic invertebrates, in contrast, usually produce carbonates which isotopically resemble ambient dissolved inorganic carbon (DIG), and therefore don’t appear to incorporate much respired CO*. This is fundamental to the use of foraminifera and other calcareous aquatic animals as recorders of ancient carbon cycle processes (e.g., Shackleton and Pisias. 1985; Spero et al., 1991). Fish otoliths appear to be variable with respect to incorporation of respired CO2 (e.g., Kalish, 1991). An additional complication involves photosynthesis, which often appears to influence the isotopic composition of coral and algal carbonates more than does respiration (McConnaughey, 1989a). Such photosynthetic 6 13Cenrichments are paleontologically useful for distinguishing photosynthetic from nonphotosynthetic organisms (Stanley and Swart, 1995) and for deriving records of light levels and cloudiness (McConnaughey, 1989a). These examples all concern so-called “metabolic” isotopic effects that result from changes in the 613C of an internal DIC “pool” (Swart, 1983; McConnaughey, 1989a; Kuile et al., 1989b). Although the isotopic composition of this internal DIC pool has not been measured in any animal, it is widely assumed that respiration adds, and photosynthesis 611

Ted A. McConnaughey et al.

612 Photosynthesis lica; =cq

‘2c

‘PI

faster

Food I

t

I

/

2H+

Ca++

Fig. I Origin of metabolic and kinetic isotope effects in biological carbonates. Metabolic effects arise from changes in the 6r3C of the internal DIC “pool”, by respiration and photosynthesis. Respiration adds, and photosynthesis removes, r3C depleted carbon from this

pool. Kinetic fractionations arise from faster hydration and (especially) hydroxylation of CO, by molecules bearing the light isotopes “C and 160. Kinetic fractionations are expressed when the HCO; resulting from CO2 hydration and hydroxylation precipitates before reestablishing isotopic equilibrium with cell DIC. As illustrated in this model, calcification appears to be driven largely by proton removal from the calcifying fluid, catalyzed by the enzyme Ca’+ ATPase (large dot on calcifying membrane).

An isotopic mixing equation describes the combination of respired and environmental carbon incorporated into the precipitating carbonate.: R(6’3CREsp) + (1 - R) (SL3CEN”) = 6’3CSHELL - A

(1)

where R is the fraction of respired carbon in the shell, S’3CREsp and S’3CENv are the isotopic compositions of HCO; derived from respiration and environmental HCO; respectively, and A represents 13Cfractionation between tissue HCO; and the shell. A assumes the equilibrium value ) if the carbonate precipitates for cCalcite-HCO,- (or EAragonite_uco; in isotopic equilibrium with tissue HCO;. This is difficult to verify, but can probably be assumed, if the shell precipitates in oxygen isotope equilibrium with its surroundings, and “kinetic” isotope fractionations (sensu McConnaughey, 1989b) are therefore likely to be small. S13CENvis approximately equal to the 613C of ambient DIC in mildly alkaline waters, or, for air breathers, the S’“C of atmospheric CO* plus the equilibrium fractionation cuco;_co,. Romanek et al. ( 1992) estimated the equilibrium 13Cfractionations for calcite and aragonite relative to HCO; as = +l.O and c~agoni~e_hco~ = +2.7, while fraction&kite-HCOj ations relative to CO2 are c,,lcite_COz = 11.98 - 0.12 (t”C) and Earagonire_co2 = 13.88 - 0.13 (t”C). Zhang et al. (1995) estimated bicarbonate-CO2 equilibrium fractionations as EHCO; -CO2 = 10.78 - .141(t”C). The isotopic composition of respired organic matter will be conserved in the resulting mix of CO, + HCO; At isotopic equilibrium, HCO; will be 13C enriched compared

to CO, by 7- 10%0. depending on temperature (Zhang et al., 1995 ). At pH values typical of invertebrate tissues (usually 7-8; Roos and Boron, 1981), more than 90% of respired CO2 hydrates and ionizes to produce HC03- . 6”CREsp, the isotopic composition of HCO; derived from respiration, should therefore be at most l%a enriched in 13C compared to starting organic material. Further corrections could be added. Respiration discriminates slightly against 13C(Deniro and Epstein, 1978; McConnaughey and McRoy, 1979), and will tend to lower the 613C of tissue HC0.T. Various biochemicals differ by several %Oin isotopic composition (Jacobson et al., 1970; Smith and Benedict. 1974)) and the mix of biochemicals being respired therefore affects the S”C of respired HCO; Diffusive fractionations may become significant when CO2 transfer between the animal and its environment becomes essentially a one-way outward flux (Jost, 1960; Cunningham and Williams, 1980). Zhang et al ( 1995) estimated diffusive fractionations against 13C02 in water to be around 0.8-l .O%Oat the temperatures of interest. Diffusive fractionations will tend to enrich tissue HCO; in 13C. For purposes of calculation, 5 ‘3CREsp,the HCO; derived from respiration, will be assumed to be 0.5%~ heavier than the organic material being respired. An error of rOS%o in this value will not affect estimates of R by more than a few percent for most animals, hence, the correction can usually be ignored. Mollusks were preferred for this analysis because their shells usually precipitate near “0 equilibrium with ambient waters (Epstein et al., 1952), and kinetic isotope effects are therefore likely to be small. Both terrestrial and aquatic snails were available. Symbiont containing methanotrophic and chemoautotrophic bivalves were included because their tissues, and therefore also respired COz, can be extremely depleted in 13C.Corals were included to introduce the complications of strong kinetic isotopic disequilibria, as well as photosynthesis. 1.1. Literature Review 1.1.1. Respiration Craig (1953), Blau et al. (1953), Keith et al. (1964). and Mook ( 1971) observed that carbonate 6 13C and A14C values in aquatic mollusks were usually similar to those of ambient DIC, but also noted cases of 13C depletion. Craig (cited in Revelle and Fairbridge, 1957) first suggested that skeletal incorporation of respired CO1 might be responsible. Swart (1983), Aharon (1991), and Wefer and Berger ( 1991) have more recently addressed the issue of metabolic 13C effects. The present discussion attempts to minimize duplication while expanding on selected aspects of the problem. Pearse ( 1970). Sikes et al. ( 1981), and Kuile and Erez (1987) fed 14C spiked foods to corals, sea urchins, and forams, and observed 14C incorporation into skeletal carbonates. Goreau (1963), Wheeler et al. (1975), Erez (1977, 1978, 1983), Kuile and Erez (1987), and Sikes et al. (1981) spiked culture media with DI14C and 45Ca, and, in most cases, observed slower fixation of 14Cinto skeletal materials.

Carbon isotopes in biological carbonates This suggested a dilution of DIi4C by unlabeled respired CO*, implying that respired COZ contributed significantly to the skeleton. DI’4C/45Ca incorporation ratios greater than one have also been observed however, sometimes in the same organisms (e.g., Wheeler et al., 1975; Erez, 1978), suggesting that respired CO* did not contribute significantly to the carbonate. The discrepant results have not been satisfactorily resolved although Tambutte et al. (1995, 1996) discuss some relevant complications with the 45Ca methods. Foraminifera have been extensively investigated with regard to metabolic effects on skeletal 613C. Ortiz et al. ( 1996)) for example, suggested that the shells of four planktonic species were 13Cdepleted by l-2.4%0 relative to isotopic equilibrium with seawater. Spero et al. ( 1991) and Spero and Lea ( 1993) concluded that respired CO* had minor isotopic effects on the skeletons of symbiotic foraminifera. Spero and Lea (1996) also fed foods of differing 613C to foraminifera and concluded that about 8% of skeletal carbon derived from respiration. Tanaka et al. ( 1986) used 6 13Cand A 14Cdata to produce independent estimates of R (the fraction of shell respired carbon) in barnacles and mollusks in a 14C contaminated estuary. Estimates of R ranged from 23-85%, but 13C and ‘“C estimates showed little correlation. Calculation procedures also appeared to exaggerate both 13Cand 14Cestimates of R. The 14Ccalculation used the expression R = (A 14Cshell - A "Co,cV(A"Ctmue - A “Cn’c), which yields meaningful results only if organic production and calcification utilize different DIC sources. In this case, seaweeds and animal tissues had A14C values as high as +882%0, with great variability, while seawater DIC yielded A’“C values between +79 and +127, with no overlap. Sampling may therefore have missed the full range of A’“C,,, which was actually higher and more variable. The S13C calculation used an expression equivalent to R(6 ‘3C’lssue+ eHcO,_co,) + ( 1 - R)(S”C,,) = 613CsHELL- ~~~~~~~~~~~~ The addition of ~HC03-_C02 to the first term biased estimates of R upward. Equation 1 yields R values between 0 and 25%. Two deep-sea studies largely overcame the limitations imposed by small isotopic differences between food and DIC. Griffin et al. ( 1989) reported that carbonate A’“C values in several deep-sea corals (about -60 to -SO%,,) were similar to ambient DIC, while coral organic carbon was similar to surface particulate matter (about +60%0). Paul1 et al. (1989) reported that the shells of methanotrophic mussels were similar in both 613C and A’“C to ambient DIC and quite different from organic components of the animals or the methane upon which they fed (Figs. 2, 14). That case will be further examined here. Finally, Fritz and Poplowski ( 1974) manipulated the S13C of DIC in aquaria containing aquatic snails. Shell S13C resembled ambient DIC, rather than the snails’ food. In summary, aquatic invertebrates probably don’t fix much respired CO1 into their shells, but the issue hasn’t been completely resolved. Unlike aquatic invertebrates, land snails appear to incorporate considerable respired COZ in their shells. By Eqn. 1, the shells of Helix raised by Deniro and Epstein (1978)

613

Mussel shells

0 f

-

20 t 0

Deep-sea methane seep Methane AA -80

after Paul1 et al. 1989 -80

-40 SW

-20

0

CL,

Fig. 2. 6°C and A’“C of shells and tissues of methanotrophic mussels from an abyssal methane rich brine seep, modified from Paul1 et al. ( 1989). Diagonal scale added to show the approximate percentage of respired carbon in mussel shells, assuming that respired carbon and ambient DIC have isotopic compositions near the respective ends of the scale. Most of the mussel shells analyzed appear to contain less than 10% respired carbon.

contain 80-90% respired Con, assuming -8%0 for the 6 ‘“C of the atmospheric CO? experienced by the snails. Goodfriend and Hood ( 1983) concluded that Jamaican land snails obtained shell carbon 25-40% from plant foods, 30-60% from atmospheric CO*, and up to 33% from ingested limestone, based on 13C and 14C data. Goodfi-iend and Magaritz (1987) concluded that the shells of terrestrial snails responded isotopically to the C/C, mixture in local plant communities, indicating significant incorporation of respired CO*. Among vertebrates, the otoliths of fish show a range (210%0) of apparent metabolic 13C effects (Mulcahy et al., 1979; Radtke et al., 1987; Kalish, 1991). Kalish suggested that otolith 13Cdepletions increase with metabolic rates. The eggshells of birds (Von Schimding et al., 1982; Shaffner and Swart, 1991) and apparently also dinosaurs (Sarkar et al., 1991) are generally several %Odepleted in “C compared to atmospheric CO*, which translates to perhaps 60-80% of eggshell carbon derived from respiration. Respired carbon also makes up most of the carbonate in mammalian bones (Deniro and Epstein, 1978; Sullivan and Krueger, 1981; Shoeninger and Deniro, 1982). Thus, respired COZ apparently contributes little to the shells of most aquatic invertebrates, probably more to the otoliths of some fish, and dominates in the shells and bone carbonates of land snails, birds, and mammals. 1.1.2. Photosynthesis Photosynthesis preferentially uses “C and should leave the internal DIC pool enriched in 13C. As with respiratory effects, this has not been directly demonstrated, but the carbonates of photosynthetic organisms generally have higher 6 13Cthan those of nonphotosynthetic animals having similar

614

Ted A. McConnaughey et al.

degrees of “0 disequilibrium. Skeletal 613C in photosynthetic corals also tends to correlate with light intensity, both over the annual cycle and with depth in the water column, and be higher than in nonphotosynthetic corals having comparable 6’*0 disequilibria (Fairbanks and Dodge, 1979; Swart, 1983; McConnaughey, 1989a; Aharon, 1991). Much has been said of apparent contradictory cases. Erez ( 1978) observed apparent negative correlations between photosynthesis and carbonate 613C in symbiotic forams, but since 6180 was also negatively correlated with photosynthesis, this may have been due to stronger kinetic effects. McConnaughey (1989a) observed negative correlations between illumination and skeletal 613C in a tidepool coral although corals from deeper waters showed the more “normal” positive correlations. The results from the tidepool coral, therefore, suggest that high light or temperatures stressed the tidepool coral during sunny periods and reduced photosynthesis. Jones et al. ( 1986) observed higher skeletal S13C in a nonphotosynthetic gastropod than in a symbiotic giant clam. Subsequent comparison with a symbiont-free encrusting bivalve suggested, however, that algal symbionts didn’t cause skeletal 13Cdepletion in the giant clam (Romanek and Grossman, 1989). Aharon ( 1991) also observed near-equilibrium 13Cprecipitation in other giant clams. The intuitive conclusion that photosynthesis elevates skeletal 13C, therefore, appears to remain valid, and satisfactory explanations have been found for situations which at first appeared to be exceptions. 1.1.3. Kinetic effects

McConnaughey ( 1989b) concluded that kinetic discrimination against 13C and I80 during the conversion of CO* to HCO; caused most of the heavy isotope depletion in many biological carbonates. Aragonites appear to be more likely than calcites to exhibit strong kinetic effects, but both minerals may exhibit these effects (e.g., Carpenter and Lohman, 1995). The kinetic effects are expressed when calcification occurs within thin, alkaline, Ca*+ rich solutions separated from adjacent cells by CO* permeable membranes. CO1 diffuses back and forth between the cells and the calcifying solution, and some of the CO* reacts (with kinetic fractionations) and precipitates rapidly as CaC03. This kinetic environment was simulated in the laboratory, resulting in carbonates which were depleted in “0 and 13C by as much as 6 and 22%0, respectively, compared to isotopic equilibrium, or about one-third more than in corals or most other disequilibrium biological carbonates. Isotopic equilibration in the biological system apparently occurs by two mechanisms, both of which tend to produce linear correlations between 5I*O and S 13C,extending toward the point of isotopic equilibrium for both isotopes (McConnaughey, 1989b). Spero and Lea (1996) recently discovered that the S13C and 6”O of foraminiferal calcite decrease as the pH of the culture nedium increases, CO;- concentrations increase, and molecular CO2 concentrations decrease. 13C and “0 depletions are proportional, yielding slopes (As 13C/AS l8O) similar to those previously observed, for example, in nonphotosynthetic corals, suggesting “kinetic” origins. Some-

thing in the external pH-CO? system, therefore, affects the magnitude of kinetic effects, possibly by influencing the rate of COZ flux across the calcifying membrane. 2. METHODS Terrestrial snails were collected from the Spring Mountain and sheep ranges of southern Nevada, primarily during the spring of 1993. Aquatic snails were collected in the Spring Mountains and Ash Meadows, along with water samples. for analysis of S l8O. Saxon Sharpe (Desert Research Institute, Reno. NV. USA) guided some field collections and identified most of the snails. Most of the aquatic systems have been extensively monitored for temperature, isotopes ( 6180, 613C, A ?), and major ion chemistry for some time (Winograd and Pearson, 1976; Perfect et al., 1995). Soil carbonates and soil gas samples from southern Nevada were collected for isotopic analysis during the same period as the snails, although not from the same locations (McConnaughey and Whelan, 1996). and were used as a general guide to 613C values in the plant communities. C. K. Paul1 and H. Cavanaugh provided shells from mussels previously collected at a deep-sea methane-rich brine seep in the Gulf of Mexico. M. Arthur provided a shell from the clam Calyptogena sp., collected at a methane seep at a depth of 601 m on the Peruvian continental shelf. Mussel and clam shells were embedded in resin, sectioned, and sampled in transects running from the outside to the inside of the shell. G. T. Shen provided sample splits of hermatypic corals (Puvona clavus and P. gigantea) from San Cristobal and Champion Islands, Galipagos. Ecuador, collected during 1989. These coral data are presented along with selected data from the same islands, previously described by McConnaughey ( 1989). Molluskan carbonates were crushed and roasted at 380°C in vacua for 1 h. Carbonates were reacted with 100% H3P04 to generate CO? for mass spectrometric measurement of shell 613C and S’*O. Two acid dissolution procedures were used. High temperature (75°C) dissolution using an automated “Kiel” carbonate extraction device connected directly to a Finnigan MAT 251 mass spectrometer was used for the isotopic sections through mussel and clam shells and corals, and for many of the analyses of snails. Low temperature (25°C) dissolution was used for the remainder of the snails. with isotopic analysis performed on a Finnigan MAT 252 mass spectrometer. Frequent analysis of isotopic standards by both methods allowed all results to be expressed in the conventional manners, as if dissolution always occurred at 25°C. Water samples were analyzed for 6 I80 by equilibrating 75 PL of water with about 30 pmol CO, at 25°C (Kishima and Sakai, 1980). CO, from CO,-water equilibrations were analyzed on a Finnigan MAT 252 mass spectrometer. Precision on replicate isotopic analyses was better than 0.1%0 for 6”O and 0.07 for 613C.Data quality was maintained, in part, through frequent comparisons against internationally recognized isotopic standards and frequent analysis of secondary laboratory standards. Results are expressed relative to V-PDB except for water 6”O values, which are expressed relative to V-SMOW. Equilibrium 13Cfractionations for calcite and aragonite a~ estimated using the equations of Romanek et al.( 1992). Equilibrium CaC03-H20 I80 fractionation factors are meanwhile estimated as (6, - S,) = 21.90 - 3.162(t + 31.06)05 for calcite (after O’Neil et al., 1969; see Wefer and Berger, 1991) and (S, - &) = 4.75 - .23t for aragonite (Grossman and Ku, 1986). In this notation, 5, is the 6 I80 of CO2 released from CaCO? by H3P0, reaction at 25”C, and &, is the 6’“O of CO2 equilibrated with water at 25°C. 6, and 6, are both expressed on the PDB-CO2 isotopic scale in applying these equations, correcting 6, values measured on the V-SMOW scale by -0.26%0 (Coplen et al., 1983). 3. RESULTS AND DISCUSSION

3.1. Isotopic Mixtures The shell of a methanotrophic mussel from an abyssal brine seep in the Gulf of Mexico ranged in 6 13C from -2

615

Carbon isotopes in biological carbonates

Methanotrophic

,

4-

mussel

Isotopic equilibrium with seawater

Outside

Aragonite Inside

oz3-

,,_$

5

.

Calcite

’

Hinge

&

I

I

I

I

I

I

I

I

10

9

8

7

6

5

4

3

2-

II 2

I 1

0

% Respired CO2 (for calcite) L

-10

I

I

-9

-8

-7

-6

-5

-4

-3

-2

-1

0

1

2

3

s'3ceJ Fig. 3. Cross plot of shell “0 vs. “C for a methanotrophic mussel from the Gulf of Mexico. Microsampling track proceeded from the outside to the inside of the shell, approximately one-third of the way between the umbo and the shell margin. Scale for fraction respired carbon in shell based on Eqn. 1, as described in text.

to -7%0 (Fig. 3). Cavanaugh et al. ( 1981) and Paul1 et al. (1989) previously reported values of -4.8%0 and -2.7 to -8.1%0 for mussels from these seeps. Shell 6180 values ( +2.6 to +3.8%0) are similar to, and generally slightly higher than, values for the surrounding hard ground carbonates (Paul1 et al.: 1992) or for expected oxygen isotopic equilibrium with ambient waters (4.5”C, aI80 = -0.1%0 SMOW; Paul1 et al., 1991; ijstlund et al., 1987). Kinetic effects on shell 613C values are therefore likely to be small. Methane, with a 613C value near -83%0, provided the main food source for the seep community (Paul1 et al., 1992; Martens et al., 1991; Cary et al., 1989). Mussel tissues averaged around -70%0 (Paul1 et al., 1991; Fig. 2). Measured seawater DIC 613C values ranged from -0.43 to +0.42%0, with +0.28%0 as a suggested value (Paul1 et al., 1991). Sediment porewaters and carbonates had substantially lower 613C values (Paul1 et al., 1991, 1992), hence the mussels may have been exposed to a 13C depleted seawater-brine mixture. Nevertheless, applying Eqn. 1 with 6’3CREsp = -70%0 and 613CENVDIC = +0.28%0 yields estimates of shell respired carbon of 5 - 12% (equilibrium calcite) or 7 14% (equilibrium aragonite) (Fig. 3). These numbers are similar to estimates based on shell A 14C (Fig. 2). The calcitic shell 613C for the Peruvian Calyptogena ranged from + 1.2 to -0.5%~ first increasing and then decreasing somewhat over the life of the animal. Seawater 613C(DIC) appears to lie between 0 and 1%0 (Kroopnick, 1985), hence shell S13C values are probably depleted compared to calcite 13Cequilibrium by about 2%0. By assuming a 613C for respired carbon of -35%0 (e.g., Kulm et al., 1986; Saino and Ohta, 1989; Rau et al., 1990; Rio et al., 1992; Masuzawa et al, 1995), it would appear that respired carbon makes up less than 7% of shell carbon. Kinetic contributions to 13C depletion are again small, as calyptogenids generally precipitate their shells near “0 equilibrium with seawater. In this case, shell 6’*0 ( +2.6 to +1.9%0) would be in equi-

librium with seawater (6 “0 = +0.3 V-SMOW) at a temperature of 6-9 “C. If shell 6 13C is plotted against tissue S13C for mollusks from various deep-sea seep and vent environments (Fig. 4), the slope of the apparent correlation (about 0.1) suggests that about 10% shell carbon derives from respiration. These organisms are particularly sensitive to respired CO2 due to the large isotopic difference between animal tissues and ambient DIC although the uncertainty is probably still several percent.

5

1

Abyssal vent and seep animals

D

Typical

_+=*

benthic molluscs

__0-*-

__,-

__**

0

??

mussels

0 Calyptogena b Solemya ??

-10 ’ -80

1 -70

I -60

I -50

I -40

trochid

I -30

I -20

-10

B3C (%,) Tissue organic carbon Fig. 4. Relationship between shell 13Cand tissue 13Cfor various deep-sea mollusks. (Data from Paul1 et al., 1989, 1991, 1992; Cavanaugh et al., 1981: Rio et al., 1992; Kulm et al., 1986 ).

? 616

Ted A. McConnaughey et al.

(vs. a,ir)

(vs. DIC)

I

Snails

Aquatic non-pulmonates

1

Algal m yhonate;

-15

-10

(vs. DIC)

~

(vs. DE), 1

-5

&13C Disequilibrium

0

,

1

+5

(“I_,)

Fig. 5. 13Cdepletions in terrestrial and aquatic snails from Nevada. Disequilibria for aquatic carbonates are calculated with respect to ambient aquatic DIC, and disequilibria for land snails calculated with respect to atmospheric CO*. Both methods of calculation are shown for air breathing, pulmonate aquatic snails.

The aragonitic shells of pseudobranch aquatic snails from springs in southern

(water breathing) Nevada ranged in

6 “C from - 1.9 to - 10.5%0. Dissolved inorganic carbon in these waters ranged from 623C = -4.6 to -9.7%0 (Winograd and Pearson, 1976; Perfect et al., 1995). Calculated 13C disequilibria for these snails, relative to aragonite precipitated in isotopic equilibrium with local DIC, averages - 1.3%0 (Fig. 5). By applying Eqn. 1, respired carbon appears to contribute about 6% on average to their shells. This calculation assumes a 613C value for snail food of -28%0, based on Riggs’ (1984) analysis of algae from nearby springs. Riggs also reported snail shell S13C values of - 1.9 to -2.95%0 for nearby springs, compared to calculated equilibrium values of - 1.9 to -2.3%0. Summarizing the results from aquatic mollusks, it would appear that respired carbon generally contributes less than 10% of shell carbon. The shells of land snails from southern Nevada ranged in S13C from -5.0 to -10.2%0, with an average around -7%0. This is about 9- 10%0 depleted in 13Ccompared to aragonite precipitated in isotopic equilibrium with atmospheric COZ (Fig. 5)) or, alternatively, perhaps 4-5%0 enriched compared to aragonite precipitated in isotopic equilibrium with respired CO? (assuming S 13CREsp= -22%0, based on isotopic measurements made on soil gases at comparable elevations.) Upon applying Eqn. 1, respired CO* appears to contribute 54-90% (average 69%) of shell carbon. Land snails may encounter a local “COa enriched atmosphere due to soil respiration, in which case the above numbers overstate the importance of the snail respiration to shell carbon. The apparent difference between the terrestrial snails and aquatic mollusks is nevertheless striking. The shells of pulmonate (air breathing) aquatic snails averaged -5.8%0 13Cdepleted compared to equilibrium with aquatic DIC or -9.8%0 depleted compared to atmospheric

CO*. The similarity of the latter number to that observed in air breathing land snails suggests that breathing air, rather than an aquatic habitat, controls the 13Cfractionation in these snails. The slightly larger 13C depletion observed in the aquatic pulmonate snails, compared to land snails, may reflect foods which are isotopically lighter or exposure to slightly lighter ambient CO,. The mollusk shells discussed so far are thought to precipitate near 13C equilibrium with the internal DIC pool. Any depression in shell 613C, compared to calculated aragonite equilibrium with ambient DIC, therefore, reflects the influence of respiration on the internal DIC pool. Respiratory 13C effects are harder to evaluate in organisms displaying substantial kinetic isotopic disequilibrium. In coral skeletons (Fig. 6)) kinetic isotope fractionations apparently cause most of the 13C depletion, as well as all of the “0 depletion, compared to aragonite precipitated in isotopic equilibrium with seawater (McConnaughey, 1989a,b). Kinetic effects are often expressed unevenly within the skeletons of corals and other organisms. The nonphotosynthetic coral Tubastrea sp. spans a range of about 3.3%0 in S’*O. while the photosynthetic corals Pavona clavus and Pavona gigantea span about 3.9%0 (Fig. 6). The variable

Gakpagas

corals

+5 Photosynthetic

Eq

(hermatypic)

44 ,’I’ .:.?? <’ ,.

corals slow

fast growth 0

growth rp

-

,,j’ ...” ,1’ ..:.

Non-photosynthetic _,o -

(ahermatypic) coral

#,’ ,,,;. ,’

, ,: -4

I -3

1 -2

I -1

J 0

8”OCL) Fig. 6. Skeletal 613C and S180 of photosynthetic (Pavona spp. ) and nonphotosynthetic (Tubastrea sp. ) corals from San Cristobal and Champion Islands, Galhpagos, Equador. The square marked “Eq” in the upper right of the figure represents estimated mean annual isotopic equilibrium for aragonite with local seawater. “Kinetic” isotope fractionations alone would hypothetically produce the dashed line passing through 13Cand I80 equilibrium with seawater. Data from the nonphotosynthetic coral are offset from this kinetic line by about - 1. 5%0in 6 “C (arrow labeled ‘‘R” ) , due to skeletal incorporation of respired CO*. For a particular value of skeletal 6”0, photosynthesis elevates skeletal 613C in the photosynthetic corals, compared to the nonphotosynthetic coral.

Carbon isotopes in biological carbonates

strength of the kinetic effect apparently reflects partial isotopic equilibration between DIC in the extracellular fluids where calcification occurs and in adjacent cells. As long as this equilibration occurs mainly through CO* exchange across the boundary membrane (Fig. 1) , carbon and oxygen isotopic disequilibria should be proportional, and the isotopitally heavy end of the 613C - 6l*O correlation line should approach isotopic equilibrium with the internal DIC pool ( McConnaughey. 1989b). The 6 13C - 6 I80 correlation line is fairly well defined in the case of the nonphotosynthetic coral and even reaches estimated “0 equilibrium with ambient seawater. The approximately 1.5%0 depression in skeletal S13C, compared to calculated aragonite equilibrium with seawater DIC, therefore, provides an estimate of the S 13Cdepression in the tissue DIC pool, compared to ambient seawater. This respiratory effect is marked “R” in Fig. 6. This 1.5%~respiratory effect will be assumed to apply to the whole coral. The skeletal 6 ‘jC enrichments in the photosynthetic corals, caused by preferential withdrawal of “C from the internal DIC pool during photosynthesis, can then be estimated by comparing portions of the photosynthetic and nonphotosynthetic corals having similar 6’*0 values. Slow growing parts of the photosynthetic corals are about 3%0 13Cenriched, while some of the fast growing parts reach ‘“C enrichments of as much as I I%o (Fig. 6). Strong relationships between photosynthetic 13C enrichments in the skeleton and skeletal growth rates should not, however, be expected since some corals grow fastest in parts of the skeleton which are only minimally photosynthetic (Fang et al., 1983). Several lessons, therefore, spring from the corals. First, as with other aquatic animals, respiratory additions of isotopically light CO2 to the internal DIC pool cause relatively small (about 1.5%~) depressions of skeletal 6 13C. Second, photosynthetic withdrawals of isotopically light CO* from the internal DIC pool can cause much larger (up to 11%0) elevations of skeletal b 13C.Finally, kinetic effects can dwarf the effects of both respiration and photosynthesis. All of the coral materials analyzed here were depleted in 13Ccompared to aragonite equilibrium, and even the most photosynthetically 13C enriched coral skeletons were several %0 depleted in ‘“C compared to seawater DIC.

3.2. Respiratory Gas Exchange Model A simple respiratory gas exchange model offers some insight into the above numbers. The premise is that a passive influx of COz from ambient waters (or the atmosphere) dilutes the CO:, produced internally by respiration and isotopitally equilibrates the internal DIC “pool” with environmental DIC. The question is, how large is this passive CO2 influx compared to respiratory CO2 generation, and what controls its magnitude? Respiration consumes O2 and produces CO* in approximately I : I stoichiometry, depending mainly on the composition of the material being respired: CHZO + OZ -+ CO2 + HZ0

(2)

617

Both O2 uptake and CO* loss occur through passive diffusion across respiratory epithelia. Fluxes (J, in mol m-’ s-‘) of O2 (subscript “0”) and CO2 (subscript “C”) may be parameterized by epithelial permeabilities (P, in m s _’) to 02 and CO1 times the gas concentration differences (in mol mm”) between blood (subscript “b” ) and ambient waters (subscripted ‘$w” ) : Jo = PO ([021w‘[0?lh)

(3)

JC = PC ([C02lw-[CO?lh)

(4)

The net outward CO2 flux- Jc therefore consists of the difference between passive inward and outward fluxes. The inward flux of environmental COZ, E = PC [CO,]w dilutes the CO? produced internally by respiration, R = Jo = PO ( [021w[O&). Skeletal incorporation of environmental CO? and, therefore, 13Cequilibration between internal DIC and environmental COz should therefore depend on ratio E/R. It follows that: E/R = (PcIP,).[CO,],I[O~];(l

- [O,],/[O,]w,m’

(5)

Basic principles of gas solubility and diffusion permit the terms of this equation to be estimated. Permeabilities of respiratory epithelia depend on the resistivities (r = I/P, in s m-‘) of component layers to gas diffusion. The two main types of resistive element in a respiratory epithelium are the lipid rich cell membranes and aqueous boundary layers (including cell interiors). The sum of these resistivities determines membrane permeability (P = l/Cr,).

Gas solubility in lipids determines membrane resistivities (e.g., Forster et al., 1969; Osada and Nakagawa, 1992). Membrane (erythrocyte) resistivity to CO1 diffusion is about 280 s m-’ (Gutknecht et al., 1977). O2 is about 10% as soluble as CO* in lipids, so membrane resistivity to O? should then be about IO times greater, or about 2800 s m-‘. Aqueous boundary layers resist gas diffusion as given by r = z/D, where z is the thickness of the boundary layer and D is the molecular diffusivity of the gas in water. Molecular diffusivities of gases in water are about 2-3 x lo-” m’ S _I, depending mainly on temperature. With boundary layer dimensional scales on the order of z = 10 -‘m, the resistivity of an aqueous boundary layer to gas diffusion is about r = (IO-’ m)/(2-3 X low9 m2 s-‘) = 3-5 X 10’ s m-‘. CO?, in effect, diffuses faster than O2 in water, however, because it hydrates and ionizes to yield DIG concentrations which are over 100 times higher than that of molecular COZ (at pH values typical of seawater or the blood of marine invertebrates.) COn ionization can be considered instantaneous at respiratory epithelia, due to catalysis by the enzyme carbonic anhydrase. The resistivity of a respiratory aqueous boundary layer to CO? diffusion is therefore on the order of I % of the resistivity of the same layer to OZ diffusion. For the purposes of calculation, the resistivities of aqueous boundary layers to CO1 and O? diffusion may be estimated at 40 and 4000 s rn-.’ respectively. The permeability of a lipid rich membrane plus one associated aqueous boundary layer to COZ diffusion is therefore

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0.2

0.4

0.6

0.6

1.0

Blood Oxygen I Ambient Oxygen

Fig. 7. Respired COz as a fraction of total CO* in the tissues and precipitating carbonates of a calcareous animal, calculated from respiratory gas exchange model. The two input parameters in the model are the C02/02 ratios in ambient air or water and the blood to ambient O2 ratio.

Pc = l/(280 + 40) = 3.1 X 10m3 m s-l. Likewise, its oxygen permeability is PO = I/( 2800 + 4000) = 1.5 X 10e4 m s-‘. The ratio of CO2 to O2 penneabilities is therefore PC/PO = 20. Additional lipid and aqueous layers don’t greatly change these permeability ratios. For the second term in Eqn. 5, atmospheric C02/02 ratios may be calculated from partial pressures (3.5 x 1O-4 for CO?, and 2.09 X 10 -’ for 0,)) yielding a concentration ratio of 1.7 X 10e3 in air. Surface waters near partial pressure equilibrium with the atmosphere will have C02/02 ratios equal to the atmospheric ratio times the ratio of gas solubilities. C02/02 solubility ratios in seawater range from 34 at 0°C to 29 at 24”C, with variations depending mainly on temperature and salinity (Broecker, 1974; Kennish, 1989). Seawater C02/02 concentration ratios are therefore about 5.0 x lo-*. Photosynthesis, respiration, and other processes may, of course, change gas ratios in specific environments. The ratio of blood O2 to ambient O2 concentrations is often about 0.9 in invertebrates, but often below 0.4 (and quite variable) in the venous blood of aquatic and air breathing vertebrates (Dejours, 1975; Gordon, 1977). Efficient circulatory systems and oxygen transport proteins in blood presumably enable organisms such as vertebrates to tolerate reduced blood O2 concentrations. The above numbers provide the basis for estimating E/R, the ratio of environmental to respired COz within an animal. For aquatic invertebrates, E/R = 20.0.05*(1 - .9)-l = 10, which is to say that -90% of the CO2 inside the animal derives from the water and -10% derives from respiration (Fig. 7). Fishes might incorporate more respired CO2 into their otoliths due to lower ratios of blood to environmental 02, accounting for their apparently stronger 13C depletions (e.g., Kalish, 1991). If land snails have blood to environmental O2 ratios similar to those of typical aquatic invertebrates, their carbonates might contain about 70% respired C02, which is similar to what was estimated from isotope

data. Birds and mammals should contain more respired CO1 due to reduced blood to environmental O2 ratios. The model, therefore, yields results which are close to those deduced from isotopic evidence, subject, of course, to the caveats that it incorporated L‘soft’’ physiological data, and applies to “typical” environmental conditions. Blood CO2 levels appear to offer a more direct approach to this question. The COz partial pressure in human venous blood, for example, is about 100 times higher than that of ambient air (Gordon, 1977); hence, >99% of blood CO2 would appear to derive from respiration. The pH sensitivity of CO, partial pressure data complicates the picture, however, and some carbon may also be excreted by nonrespiratory pathways. The above approach is therefore preferred and has the additional benefits that it retains the connection to oxygen fluxes and does not require data on blood and environmental CO? levels, which tend to be less accurately known than O2 levels. An implicit assumption in the gas exchange model is that respiratory gas exchange is sufficient to meet oxygen requirements, but not much more. Occasional low oxygen conditions or bouts of activity might, of course, force organisms to have larger and more permeable gills than would be required for routine Oz uptake. Spare respiratory capacity would tend to equilibrate CO2 between tissues and environment beyond the levels predicted here. Bivalve mollusks (and various other animals) also feed with their gills; hence, gill size may exceed respiratory requirements. Any such condition would help to flush respired CO1 from the tissues. Environmental C02/02 ratios also vary, affecting the CO2 uptake term E. Finally, the fluids from which calcification occur may exchange water with the environment directly ( McConnaughey, 1989b), providing an additional nonrespiratory mode for ‘3C/“C equilibration between calcifying fluids and the environment. These considerations may be significant, but the respiratory gas exchange model in its simple form offers plausible explanations for the respiratory 6’“C effects observed in a wide variety of aquatic and terrestrial animals. Temperature, better physiological data, assays for carbonic anhydrase, and, of course, isotopic measurements for tissue DIC are obvious places to seek refinements. It is also interesting to note that because environmental COz/02 ratios largely control shell ‘“C fractionations in this model, shell S13C may sometimes find applications as recorders of environmental C02/02 ratios. 3.3. Photosynthesis Photosynthetic 13C enrichment in the internal DIC pool, like respiratory 13C depletion, is constrained by the permeabilities of membranes and tissues to CO2 exchange with ambient waters. Several factors can, however, lead to larger photosynthetic 13C enrichments than respiratory r3C depletions. First, photosynthetic organisms typically calcify mainly during the daytime (e.g., Goreau, 1963; Barnes and Crossland, 1978). Respiration may also be suppressed in the dark (Ktthl et al., 1995). Carbonate isotopic composition there-

619

Carbon isotopes in biological carbonates fore reflects periods of maximum photosynthetic influence, rather than periods of strong respiratory influence. Second, daytime gross photosynthesis (P) is often several times faster than respiration (R). Nothing is known about the applicable values for the corals shown in Fig. 6, but Davies ( 1984) reported 24 h average P/R of 1.9 and daytime PIR of about 4.5 in a different coral species. Fisher et al. (1985) similarly estimated average 24 h P/R ratios of l-3 for a symbiotic giant clam and daytime P/R of about 4. If respiration reduces the 6’C of the internal DIC pool in the nonphotosynthetic coral by 1.5%0, then net photosynthesis potentially raises pool 613C at least four times as much, or more than 6%0. This crude estimate accounts for most of the apparent elevation of skeletal S13C in photosynthetic corals. Finally, in the case of corals, gas exchange rates may be lower during the day because the coral polyps tend to close their mouths during the daytime. There are at least three possible explanations for this: ( 1) corals feed on plankton mainly at night; (2) symbionts produce oxygen during the day, so the coral doesn’t need to ventilate as much; and (3) the coral stimulates symbiont photosynthesis by discharging protons from calcification into the coelenteron, thereby increasing coelenteron CO2 levels ( McConnaughey, 1994; McConnaughey and Whelan, 1996). Restricting water fluxes through the coelenteron during the day increases the efficiency of this physiology. A carbon isotope balance model sheds further light on photosynthetic “C enrichments. In this model, isotopic composition of an internal DIC pool is determined by the rates of photosynthesis (P) , respiration (R), skeletogenesis (S), carbon uptake from solution (I), carbon efflux to solution (O), and the isotopic fractionations associated with these processes:

I-O=P-R+S

6 Skeleton=

6 Pod

-A

For an organism which depends primarily on photosynthesis for organic carbon, SReap- SPhoto- 6Tiaaues. (Eliminating predation greatly simplifies the mass and isotope balances, but this approximation limits the approach to selected shallow water corals.) The skeleton-pool 13Cfractionation A has the same kinetic origin in the photosynthetic coral as the nonphotosynthetic coral and might, therefore, have a value around 10%0 for rapidly growing corals. During the daytime when most calcification occurs, P, R, and S in corals may have magnitudes of about 4, I, and 2 pmol rn-’ s-’ (e.g., Davies, 1984). I and 0 are unmeasured, but, from the above isotope balance, appear to have values around 13 and 8 pmol m -’ s-’ respectively (Fig. 8). The two largest terms in the carbon budget have therefore never been measured. Photosynthesis enriches the internal DIC pool by about 7%0 compared to ambient DIC, but the kinetic 13Cfractionation during calcification is larger so that the skeleton ends up 3%0 depleted in 13Ccompared to ambient seawater. The lack of experimentally determined quantitative mass

,5 ,

Coral daytime carbon budget

Efflux 7.9

CaCO, 2

-5 -

Photosynthetic fractionation

-10

-

-22.8

-1.5 -

“bO ,Tlssue 3

Food

UMOI m2 s-l

-20 Fig. 8. Daytime carbon isotope balance model for a fully autotrophic coral. Seawater DIC ( 613C = + I%o) is split into three product streams, tissue (assumed 613C = -15%0, 3 j~&ol mm2s-‘),CaC03 (assumed SL3C = -2%~ 2 umol me2 SK’). and DIC efflux (calcuiated 613C = +7.8%0, 7. 9 ‘km01 m-* s-l). An additional input of 13Cdepleted food carbon would allow isotope balances to be constructed with less ‘T enrichment in the calculated DIC pool.

and isotope balances in calcareous organisms may hide deficiencies in such models. Internal DIC probably behaves as multiple semi-independent pools subject to different physiological fluxes and kinetics; photosynthetic 13Cfractionations probably depend on internal CO, levels; active circulation probably influences DIC uptake and efflux; and kinetic fractionations during calcification may be more complicated. It appears, however, that models of this type shed some light on metabolic 13Cfractionations. The relatively large size of metabolic 13C fractionations in calcareous photoautotrophs permits a number of paleontological applications. For example, McConnaughey ( 1989a) suggested that corals can sometimes be used as cloudiness recorders, and Stanley and Swart (1995) determined that tropical corals had already developed photosynthetic symbioses by the Triassic. 4. SUMMARY

Skeletal incorporation of respired CO? typically reduces shell S13C by <2%0 in aquatic invertebrates, and the effects are sometimes imperceptible. Respired CO2 makes much larger contributions to the carbonates of air breathing organisms. The difference apparently lies in the strength of the passive equilibrating CO2 flux from the environment into the animal (stronger in water) and. secondarily, with the animal’s tolerance for reduced gas exchange, which causes respired CO* to accumulate in the tissues (most pronounced in vertebrates). In reef corals, photosynthesis produces stronger “metabolic” effects than does respiration because calcification occurs mainly during periods of active photosynthesis when

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the carbon fluxes associated with photosynthesis are larger than those associated with respiration. The internal DIC pool in photosynthetic corals apparently becomes about 7%u enriched in 13C,compared to seawater. This enrichment is usually masked in the skeleton, however, due to strong kinetic isotope fractionations. The skeleton therefore precipitates several permil depleted in 13Ccompared to aragonite equilibrium with ambient DIC. The shells of aquatic invertebrates usually record the 613C of DIC fairly accurately, while fish otoliths are less likely to do so. Even when substantial kinetic isotopic disequilibria exist, as in corals, it may be possible to retrieve information on photosynthesis, and, therefore, light levels and cloudiness. Terrestrial carbonates (land snails, bird eggs, and mammalian bone carbonates) will not record the S13C of atmospheric COz, but are likely to record the C3/C4 mixture of their diets.

Acknowledgmenrs-Saxon Sharpe of Desert Research Institute, Reno, Nevada, guided the collections of most snails, identified them, and commented on their respiratory habits. Glen Shen, Mike Arthur, and Helen Cavanaugh provided coral and mollusc shells for isotopic analyses. Craig Johnson, Richard Forester, Howard Spero, Scott Carpenter, David Des Marais, and an anonymous reviewer contributed much to whatever intelligibility the manuscript now has.

Editorial

handling:

D. .I. Des Marais

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