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Exchange of oxygen isotopes between water and organic material
Isotope Geoscience, 1 (1983) 357--370 Elsevier Science Publishers B.Y. , Amsterdam - Printed in The Netherlands
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EXCHANGE OF OXYGEN ISOTOPES BETWEEN WATER AND ORGANIC MATERIAL
K.W. WEDEKING and J.M. HAYES
Biogeochemical Laboratories. Departments of Chemistry and of Geology, Indiana Uniuersity , Bloomington , IN 47405 (U.S.A .) (Received April 12, 1983 ; accep ted for pu blication June 14, 1983)
ABSTRACT Wedeking, K.W. and Hayes, J.M. , 1983. Exchange of oxygen isotopes between water and organic material. Isot. GeoscL , 1: 357--370. The exchange of IBO at concentrations near natural abundance between benzoic acid and water and between algal biomass and water has been followed as a function of time and of pH at BODC. Exchange proceeds at conveniently measurable rates at pH 1. 0. The fir st-order rate constant for exchange of 'BO between benzoic acid and water under these conditions is (1.4 ± 0.1) • 10-' min ". An equilibrium isotope effect of 1. 72 ± 0 .09% favoring partitioning of 18 0 into the carboxylic acid is found . The exchange of "0 between wa ter and oxygen-bearing functional groups in algal biomass does not proceed to completion , with only 30-33% of the oxygen in the in ' soluble residue being exchangeable. The first-order rate constant for this exchange is (9 .3 ± 0. 5 ) . 10-' min. - , at 8 0 DC at pH 1.0. If it is assumed that exchange is first order in H+, a half-time of 150 a is calculated for turnover of the exchangeable portion of the oxygen pool at pH 6.0. The existence of a non-exchangeable oxygen pool suggests that isotopic analyses of oxygen in se d im en tary organic matter can be of value in the reconstruction of ancient environments and biochemical pathways.
INTRODUCTION
An extensive literature exists concerning the biogeochemistry of the stable isotopes of carbon, hydrogen and nitrogen. In contrast, the potentially interesting organic geochemistry of oxygen isotopes (i.e. 16 0 and 18 0 ) has been largely ignored, probably for at least two reasons: (1) Until recently (Wedeking and Hayes, 1984), no analytical method has been described that allows highly precise determinations (an uncertainty of 1%0 or better) of the oxygen isotopic compositions of whole cells or of sedimentary organic matter. Previously described methods (Hardcastle and Friedman, 1974 ; Ferhi et al., 1976; Thompson and Gray, 1977 ; Brenninkmeijer and Mook, 1981 ; Hoering and Estep , 1981) have been limited to analyses of pure organic compounds containing only carbon and oxygen.
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(2) Interest in the organic geochemistry of oxygen has been restrained because it is known (e.g., Samuel and Silver, 1965) that organically bound oxygen, in particular the oxygen of carbonyl and carboxyl functional groups, exchanges with water. Thus, it is not surprising that studies of the fractionation of oxygen isotopes within living systems have been limited to that associated with the biosynthesis of cellulose, the oxygen of which is only very slowly exchangeable at physiological pH (Epstein et al., 1977; DeNiro and Epstein, 1979,1981). It is not known whether the oxygen isotopic composition of sedimentary organic matter is related to that of its biological precursors (i.e, carbohydrates, proteins and lipids, or simply, "biomass"). In "the case of the carbon isotopes, the isotopic compositions of humic substances and kerogen formed in marine environments are thought to reflect closely (i.e. within a few per mil) the average isotopic compositions of the living material from which these substances are derived (Degens, 1969; Nissenbaum and Kaplan, 1972). Thus, kerogen contains carbon isotopic information that can be related to biochemical processes. Analyses of the carbon isotopes in ancient kerogens have, for example, led to inferences regarding Precambrian biochemical evolution (Baur, 1983; Hayes, 1983; Schidlowski et al., 1983). Before oxygen isotopic evidence can similarly contribute . to the characterization of ancient biochemistries, the fundamental relationship between the oxygen isotopic composition of kerogen or "proto-kerogen" formed in marine environments and the isotopic composition of biogenically derived precursor material (i.e. planktonic biomass) must be determined. Because a fraction of the oxygen in biomolecules is exchangeable with water, the isotopic composition of oxygen in kerogens formed in marine sediments is likely to be primarily dependent upon both the isotopic composition of the precursor material and the extent of exchange during incorporation into kerogen. If exchange is extensive, biogenically derived isotopic "information" would be overwritten by an abiogenic process. The purpose of this paper is to assess experimentally the extent to which the oxygen in organic matter undergoes isotopic exchange under conditions simulating the initial stages of kerogen formation. The non-hydrolyzable insoluble residue of modern blue-green algae may be representative of the precursor of algal kerogens in ancient sediments (Philp, 1976). The earliest phase of kerogen formation from algal biomass would involve formation of such an insoluble residue by sedimentation and lysis of cells, followed by partial hydrolysis of proteins and glycoproteins, triglycerides, polysaccharides, etc. Hydrolytic decomposition of biopolymers within the sedimentary environment might be catalysed by bacterial enzymes or by clay minerals. Oxygen exchange between the forming residue and water would be likely to occur. In order to estimate quantitatively both the extent of, and the rate of oxygen isotopic exchange that would occur during the formation of algal
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"proto-kerogen" , acid-eatalyzed hydrolytic decompositions of blue-green algal biomass have been carried out in water enriched in 18 0 . By measuring the oxygen isotopic composition of the insoluble material as a function of time , the kinetics of oxygen exchange between the residue and water have been followed and the dependence of the exchange rate upon pH has been studied. To facilitate oxygen exchange, all experiments were carried D out at 80 C. Results show that even after extensive hydrolytic dissolution of the algal organic matter had occurred, a significant fraction of the oxygen within the insoluble residue had not undergone isotopic exchange. In a control experiment, the kinetics of oxygen isotopic exchange between benzoic acid and water have been followed, as done previously (Roberts and Urey, 1939). At equilibrium, the benzoic acid was enriched in 18 0 relative to the water, thus indicating an equilibrium isotope effect in the exchange reaction. The observation of this isotope effect facilitated mathematical treatment of the results obtained in the exchange of algal organic matter and suggests a possible mode of oxygen isotopic fractionation during biochemical reactions. MATERIALS AND METHODS
Mass spectrometry and calculations
Isotopic analyses of CO 2 were conducted using a dual viscous inlet triple collector isotope ratio mass spectrometer (Nuclide Corp., State College, Penn., model 6.00), using well-established procedures (Mook and Grootes, 1973). Isotopic abundances are reported as delta values, in per mil, vs. standard mean ocean water (SMOW): fj x =
[(R x - R SM O W )/R SM O W ]10 3
where R = 18 0 / 16 0 (R SM OW = 2.0052-10- 3 ; Baertschi, 1976). The mass DC spectrometer is calibrated using CO 2 prepared at 25 from the TKL-l and K-2 standards assuming DTKL-l = 26.47, DK-2 = 3.19 (Blattner and Hulstan, 1978), and Qeo -calcite = 1.01025 (Friedman and O'Neil, 1977). The calibration has been fudependently · confirmed by carrying standard water samples through the sample-preparation procedure employed for organic samples (see p. 360). Using a recently des cribed pyrolytic method (Wedeking and Hayes, 1984), the oxygen of organic samples was converted to carbon dioxide for mass spectrometric analysis of 18 0 /160 ratio. The standard deviation of a single isotopic analysis by this technique is ± 0 .6%0, based on a pooled standard deviation of 17 analyses of pure cane and beet sucrose, reported by Wedeking and Hayes (1984) . The procedure involves no significant oxygen blank « 0.2 umol 0) in analysis of organic samples; the carbon dioxide obtained contains oxygen which is derived from only the sample. It has been shown (Wedeking and Hayes, 1984) that isotopic analyses of
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biological materials (lung, brain) give precision not different from that obtained in analyses of nitrogen- and sulfur-free, pure organic material (e.g., sucrose). The oxygen of water samples was converted to carbon dioxide by a similar procedure in which excess carbon was added to the sample prior to pyrolysis. This procedure involves a small oxygen blank (8 blank = -10 ± 1%0, amount = 3 ± 0.5 ,.mlol 0; determined by analyzing sucrose samples with similar carbon addition) . The effect of this blank must be taken into account, particularly in the analysis of isotopically enriched waters (for which the isotopic contrast between the sample and the blank is great) . Application of this procedure (C0 2 production and blank correction) in analysis of a 2.2-mg sample of the NBS-l A water standard yielded a result of -24.2 ± 0.6%0; the accepted value for this standard is -24.3 ± 0.3%0 (Gonfiantini, 1978). In mathematical treatment of kinetic data (see Results and Discussion), interconversion between cS -values and fractional abundances of ISO was frequently necessary. Given that:
r; = IS0x/eSOx + 170x + 160x) = [1 + e 70x/ I SOx) + (1 / 1sR x )r 1
(1)
where F x symbolizes the fractional abundance of ISO in oxygen pool x and R x represents the IS0/160 ratio of that pool, we can write:
(2) This useful approximation is obtained by assuming that the 170 /1S0 ratio is constant for all samples and is the same as that in SMOW, that is, 0.186. An approximate expression relating cS x and F x is obtained by substituting eq. 3 into eq. 2:
Rx
=
[10- 38 x + 1] R SM OW
(3)
The resulting expression is accurate to 0.1%0 over the range of enrichments encountered in this study.
Treatment of organic samples Exchange studies were carried out using reagent grade benzoic acid and a culture of the blue-green alga Spirulina. Dry algal powder was suspended in distilled water and the cells were disrupted using a chilled French pressure cell at 1.24.10 5 kN m ". The homogenized material was lyophilized and aliquots of a single preparation were used in the exchange experiments. A sample of ISO-enriched water was divided into fractions and these were acidified to pH 1.0 and pH 2.0 by addition of aqueous 37% HCI. These water samples were used as the exchange media. The isotopic composition of the pH 1.0 water sample was directly determined by the method described above, duplicate analyses giving 8 = 221.3 ± 0.9%0' The aqueous
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exchange media were assumed to be of identical oxygen isotopic compositions; the amounts of aqueous HCI added to the original water sample were not great enough to affect the total isotopic composition by more than 1%0' The exchange reactions were carried out at 80 ± 1°C, following a procedure described previously (Roberts and Urey, 1939). Samples of benzoic acid and of the algal homogenate were weighed and transferred into reaction vials, 100 ± 2 mg of material being used in all cases . The organic samples were suspended in 8 ml of the appropriate medium which had been . previously heated to 80°C. The vials were tightly capped and shaken, resulting in complete dissolution of the benzoic acid and an initial solubilization of -35% of the algal material (on a dry weight basis). The reaction vials were held at 80°C for 0-174 hr. The supernatant solutions showed no pH change over the reaction periods employed. At appropriate times, vials were removed from the 80°C oven and cooled rapidly to room temperature. The benzoic acid samples precipitated from solution at this time. The tubes were centrifuged and the supernatant solutions discarded. The insoluble algal residues were freeze-dried for 12 hr.; losses by partial hydrolytic dissolution were found to be as great as 80% of the initial dry weight of material. Benzoic acid samples were freeze-dried for only 4 hr. in order to minimize losses by volatilization (as in the procedure of Roberts and Urey, 1939). Separate freeze-drying containers were used to avoid cross-contamination. Oxygen isotopic compositions were determined after samples were free of residual water. RESULTS AND DISCUSSION
Exchange of oxygen between benzoic acid and water Results obtained in the exchange of benzoic acid oxygen with the aqueous medium at pH 1.0 and 80°C are presented in Table 1. Fig. 1 shows the observed oxygen isotopic compositions for benzoic acid (top curve) plotted vs. time in hours. The kinetics of 18 0 exchange between benzoic acid and water have been studied previously by Roberts and Urey (1939), under similar conditions (i.e. 80°C and 0.1 M HCI). These workers found that the kinetics of the exchange reaction are independent of benzoic acid concentration, first order with respect to H+ concentration, and are described by the integrated first-order rate expression:
(4) where k is the first-order rate constant in min. -I; and F j • F t and F w are the fractional abundances of 18 0 in the benzoic acid initially, in the benzoic acid at time t and in the water, respectively. The equation was derived as-
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- - - -- - - - - -(H20 )- - - - - - - - 200
oL..-..L-....L........L-L.---l_L.-..L-....L........L-L.---I.---....J
o
~O
50
t,hr Fig. 1. Isotopic compositions of organic oxygen vs. time for materials exchanged with I·O-enriched water. All experiments carried out at BOce . {, '" 0 of water = +221.3 ± 0.9%0 vs. SMOW in all experiments ( rJ = benzoic acid, pH 1.0; 0 = algal homogenate, pH 1.0 ; I> = algal homogenate pH 2.0). Solid curues calculated assuming first-order exchange kinetics, as explained in text. TABLE I Exchange of oxygen isotopes between benzoic acid and water at 80 ce * ' Time (hr.)
o'"OSMOW* ' (%0)
F(* 3)
0.0 5 .0 30.0 103.0( 00)
+23 .2 +133 .2 +227.4 +242.0
2.0467 2.2662 2.4541 2.4B32
( x 10 3 )
*, pH 1.0 ; s '"0 of water = +221.3 ± 0.9%0 vs. SMOW. *, Observed isotopic composition of benzoic acid oxygen. *3 Frac tional abundance of ISO in the benzoic acid calculated from observed value of {,'·OSMOW·
suming that the 18 0 is initially concentrated in the water, that the benzoic acid oxygen atoms are equivalent, and that the water represents an infinite exchange pool (i.e, F w is constant). According to eq. 4, the exchange is first order in (Fw-Ft) , that is, the reaction will be complete when the 18 0 abundance of the benzoic acid is the same as that of the water. The rate constant given by eq. 4 is proportional to the H+ concentration and was previously found to be (1.42 ± 0.04)-10- 3 min. -1 (Roberts and Urey, 1939) for 0.1 M HCI at 80°C, giving a half-time for the exchange reaction of 8.14 hr .
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The results summarized in Table I and plotted in Fig. 1 (top curve) show that the oxygen isotopic composition of the benzoic acid approaches an equilibrium value that is heavier than the water (i.e. the acid is enriched in 18 0 ). An equilibrium isotope effect is, therefore, present. Such an effect was not noted by Roberts and Urey (1939), who studied the reaction only under conditions far from equilibrium; that is, for times less than the halftime of the reaction and with large initial differences in the 18 0 concentrations of the benzoic acid and the water. Oxygen isotopic exchange reactions of organic compounds have been comprehensively reviewed by Samuel and Silver (1965); such an isotope effect, involving the exchange of a carboxylic compound with water, was not reported. No such reports have subsequently appeared in the literature. Given the isotopic compositions of the H 20 and of the benzoic acid at 103 hr. (Table I) the equilibrium fractionation factor for the exchange reaction: H 2180 +l.C6HsC1602H "'" H/ 60 +.!.C6HsCI802H 2 2
(5)
can be calculated. The fractionation factor, abcnzoic acid-water = R acid / R water ' calculated according to Friedman and O'Neil (1977), is found to be 1.0172 ± 0.0009 for the reaction at 80°C.
A modified rate expression for benzoic acid oxygen exchange. Since an implicit assumption in the derivation of eq. 4 is that the oxygen isotopic composition of the benzoic acid will be equal to that of the water at equilibrium, a modified expression, reflecting the approach of the 18 0 abundance in the benzoic acid to its equilibrium value, is needed. The modified form of eq. 4 should be first order in the difference (Fe-Ft), where Fe represents the final (equilibrium) fractional abundance of 18 0 in the benzoic acid: (6)
In eq. 6, k, t, F i and F t are as in eq. 4 but F w (the fractional abundance of 18 0 in the water) has been replaced by Fe. Note that Fe can be calculated knowing both the isotopic composition of the water and the equilibrium isotope effect for the exchange reaction. For the data of Table I, taking Fe = 2.4832.10- 3 , a plot of-log (Fe-FI) vs. time allows calculation of the first-order rate constant from the slope of the line obtained. The rate constant found is (1.4 ± 0.1)'10- 3 min. -I, in close agreement with the previously reported value (Roberts and Urey, 1939). Substitution of the calculated rate constant into eq. 6 and solution for F t allows calculation of 0 180 for the benzoic acid as a function of time. As seen in Fig. 1, the curve obtained falls close to the points corresponding to benzoic acid.
Biogeochemical significance. The mechanism of oxygen isotopic exchange between carboxylic acids and water is thought to involve hydration of the
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acid to give a symmetrical intermediate (i.e the ortho-acid) , which can dehydrate to give the exchanged product (Bender et al., 1956). A similar mechanism, involving a symmetrical hydrated intermediate, is responsible for the exchange of aldehyde and ketone oxygen with water (Byrn and Calvin, 1966). The direction, if not the magnitude, of the isotope effect observed here may be representative of the effect that would be associated with carbonyl oxygen exchange at 80°C. At lower temperatures, equilibrium fractionations commonly increase. The direct observation of an isotopic fractionation in the exchange of oxygen in benzoic acid suggests that biosynthesized organic molecules would be enriched in 18 0 relative to cell water if oxygen within intermediates or products had undergone exchange via carbonyl or carboxyl hydration. Interestingly, DeNiro and Epstein (1981) have found that the cellulose synthesized by aquatic plankton is uniformly enriched in 18 0 , relative to cell water, by an average of 27%0' Thus, the results obtained here support these workers' suggestion that hydration of carbonyl functional groups of intermediates during the biosynthesis of cellulose is responsible for the observed fractionation. It seems likely that a similar mechanism of isotopic fractionation would play a role in determining the oxygen isotopic comp osit ions of biosynthesized products such as proteins and fatty acids.
Exchange of oxygen in algal biomass Results obtained in the hydrolytic decompositions of algal organic matter at 80°C and at both pH 1.0 and pH 2.0 are presented in Table II. The oxygen isotopic compositions of the aqueous media were unchanged from experiments involving benzoic acid . The table giyes the measured total organic oxygen isotopic compositions of the insoluble residues and the corresponding times of reaction; the isotopic compositions are shown plotted vs. reaction time in Fig . 1 (middle and lower curves, corresponding to reaction at pH 1.0 and pH 2.0, respectively ). Also given in Table II are the percentages of starting material recovered, calculated on a dry weight basis.
Qualitative interpretation of results. Fig. 1 (lower two curves) shows that the 18 0 abundances of the insoluble residues increased relatively smoothly with time, with the rate of isotopic enrichment being greater at lower pH. A parallel decrease in the percent recovery of starting material occurred at pH 1.0, while recovery of material was relatively constant at pH 2.0 (Table II). Thus, both dissolution of the material and 18 0 enrichment are acid-catalyzed processes. This suggests that dissolution of material over time (besides that fraction initially soluble) proceeded by acid-catalyzed hydrolytic decomposition of biopolymers. Oxygen isotopic exchange between water and the organic residue would accompany hydrolysis, thus accounting for, at least in part, the dependence of 18 0 enrichment rates upon pH. Of course, oxygen exchange could occur without hydrolytic
365 TABLE II Exchange of oxygen isotopes between algal homogenate and water at 80°C* I Tim e (h r .)
°(%0) tat 18 0 SMOW *'
0ex 18 0 SMOW*' (%0)
Mat er ial rec overy*4
+20.3 +45 .0 +49 .6 +66.3 +79 .3 +85.9
+ 20.3 +103.8 +119.3 +175.7 +219.6 +242.0
+66 +56 + 47 +4 1 +33 +26
+26.5 +28.9 +35.4
+ 41.2 + 49.4 + 71.3
+74 +64 +63
( %)
pH 1.0 : 0.0 5.0 10.0 20. 0 40 .0 10 3.0 pH 2.0 : 10 .0 40 .0 103.0
* ' 0" 0 of water = +221.3 ! 0.9%0 vs. SMOW. *, Ob served isotopic compositi on of total organic oxygen within insolu ble algal res idues. *3Iso top ic com p ositio n of exchangeable fr ac t ion of oxygen with in algal residue. Calcu late d as ex p lain ed in te xt. *4 Calcu lated relative to amoun t o f starti ng material, on a dry weight basis .
dissolution of material, bu t in this case exch ange would still be favored at lo wer pH. The observed enrichments of 18 0 with time are probably caused by exchange of oxygen isotopes between the 180-enriched water and the insoluble residue. Po ssible alternative mechanisms include: (1 ) dissolution of organic matter containing oxygen that is depleted in 18 0 relative to the average isotopic composition of the material ; (2) loss of material togeth er with a kinetic fra ctionation of oxygen isot opes in the residue; and (3) oxidation of the material by O 2 in equilibrium with water. Reactions were carr ied out in air -tight containers with minimal head space, thus (3) may be ignored. Kinetic effects during hydrolytic dissolution of material may be responsible for a portion of the observed 18 0 enrichments with time . Such effects, if present, are difficult to distinguish from exchange, and have not been separately treated. Two lines of evidence indicate that selective losses of 180-depleted material did not account for a significant fraction of the observed enrichments: (1) at pH 2.0 , enrichments of up to 25%0 were observed without loss of material (relative to recovery at zero time), thus at this pH , exchange alone had occurred ; and (2 ) oxygen eleme n tal abundances of the residues showed no monotonic variations with time at either pH , with averages of 23 ± 5 and 24 ± 1 wt. % O 2 (± 1a) being found for the residues obtained at pH 1.0 and at pH 2.0 , respectively . Thus progressive loss of an isotopically distinct, single component (i.e.
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protein, lipid or carbohydrate) is not indicated, since each of these components have, on average, distinctly different oxygen elemental abundances. It appears that solubilization of "average" organic material had occurred. It is, then, reasonable to assume that oxygen isotopic exchange is principally responsible for the observed variations in 18 0 abundances of the algal residues. In the exchange of algal oxygen at pH 1.0, an asymptotic isotopic composition was essentially reached after 103 hr. (Fig. 1), suggesting attainment of equilibrium in the reaction. The isotopic composition after this apparent equilibration is significantly depleted in 18 0 relative to the water. The equilibrium difference of 135%0 (H 20 - residue) is too large to have been caused by an isotope effect; exchange of oxygen between organic compounds and water would result in enrichment of the heavy isotope within the organic compound, as opposed to the observed depletion. Thus, it must be concluded that a significant fraction of the oxygen within the algal residue had not undergone isotopic exchange after 103 hr. at pH 1.0. This conclusion, in conjunction with the apparent attainment of equilibrium, requires the existence of at least two pools of organic oxygen with greatly differing exchange rates. A kinetic model that assumes the existence of both exchangeable and completely non-exchangeable pools of organic oxygen within the insoluble algal residue adequately fits the observed data, as shown below. In this model, the relative sizes of the two pools are constant. Non-exchangeable oxygen might exist as non-hydrolyzable amides and esters (e.g., in proteins and triglycerides), as well as hydroxyl oxygens of polysaccharides. The exchangeable oxygen pool probably consists of ketones and as free carboxylic acid and aldehyde groups that, in part, may have been present as ester, amide and acetal groups prior to hydrolytic exchange.
Quantitative treatment of results. Calculation of the relative sizes of the exchangeable and non-exchangeable oxygen pools can be accomplished using the mass-balance equation: (7)
where 8 tot, 8 ex and 8 nex represent the isotopic compositions of the total algal oxygen pool, the exchangeable and the non-exchangeable oxygen pools, respectively; and where f ex is the fraction of the total oxygen residing in exchangeable functional groups within the insoluble residue. It is assumed that f ex remains constant. Eq. 7 would be exact if 8 -values were replaced by the corresponding fractional abundances, but the approximation as written is accurate to 0.1%0 over the range of isotopic compositions encountered here. Calculation of the fraction of exchangeable oxygen within the algal residue is accomplished as follows. The value of 8tot = +85.9% 0 of Table II, at a time of 103 hr., gives the total isotopic composition of the algal ma-
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terial after complete equilibration of the exchangeable pool with water. By making the reasonable assumption that the non-exchangeable oxygen pool has an isotopic composition always given by the total isotopic composition at time zero (i.e. that both pools had the same initial isotopic composition), only the equilibrium value of [j ex is needed to calculate the relative pool sizes. The value of [j ex after complete exchange is probably close to that of the water (i.e. +221%0 vs. SMOW). An isotope effect like that encountered in the benzoic acid study might, however, intervene. This possibility gives rise to some uncertainty in the calculated value of feX' Assuming the equilibrium value of [j ex to be the same as the water gives a value of 33% for the relative percentage of exchangeable oxygen, according to eq. 7. On the other hand, assuming the exchangeable pool to be comprised entirely of "carboxyl-like" oxygen, implies that [j ex would approach an equilibrium value of +242%0 vs. SMOW, as observed directly in the case of benzoic acid exchange. In this case, a value of 30%, only slightly different from 33%, is obtained for the relative size of the exchangeable pool. Given the observed total isotopic compositions of Table II, the isotopic composition of the exchangeable oxygen pool can be calculated at each time of reaction, by use of eq. 7. This has been carried out by assuming that the exchangeable pool is "carboxyl-like", that is, f ex = 0.33 and that {j ex = 242%0 at equilibrium; the values obtained are given in Table II, and are symbolized by [j ex 180SMOW' The data summarized in Table n can be used to calculate rate constants for the exchange of algal oxygen, assuming first-order kinetics, according to eq. 6. The procedure used is identical to that described for the case of benzoic acid exchange; values of -log(Fe - F t ) are calculated and plotted vs. time. Linear plots were found for the algal oxygen exchange at both pH 1.0 and pH 2.0, thus verifying the occurrence of first-order kinetics. A rate constant of (9.3 ± 0.5).10- 4 min."! is calculated for the exchange at pH 1.0, while at pH 2.0 a rate constant of (3.7 ± 0.8).10- 5 min.T ' is found. Thus, the exchange of algal oxygen is close to being first order in H+ concentration. If the calculated rate constants are substituted into eq. 6, solving for F t allows calculation of the isotopic composition of the exchangeable oxygen pool as a function of time at both pH 1.0 and pH 2.0. Substitution of calculated isotopic compositions into eq. 7 then allows calculation of the total isotopic compositions with time. Curves calculated in this way fall close to the observed points, as seen in Fig. 1 (middle and lower solid curves). Since the isotopic composition of exchangeable oxygen in the algal organic matter has been found to follow eq. 6, it would seem that the exchangeable oxygen pool is kinetically homogeneous. That is, exchangeable oxygen pools with greatly differing exchange rates are apparently not present. The observed rate constant for algal exchange at 80 and pH 1.0 is within a factor of 1.5 of that of benzoic acid under identical GC
368
conditions, and is close to the exchange rates of the amino acids serine, lysine and leucine recently found, under similar conditions, by Murphy and Clay (1981). Since the exchange of the algal material involved an insoluble organic oxygen pool, a direct comparison of the observed exchange rates with those of compounds studied in solution may not be informative. Thus, an unequivocal interpretation of the observed exchange rates in terms of the functional group(s) involved is not possible. However, the observed rates of exchange and the dependence upon pH are consistent with a mechanism involving hydration of carboxyl or carbonyl functional groups.
Geochemical significance of algal exchange studies. The minimal time required for isotopic equilibration of exchangeable oxygen in algal biomass under conditions less acidic than those employed in this study can be crudely estimated by assuming that the rate constant for the exchange of algal organic matter is first order with H+ concentration. For example, at pH 6.0 and 80cC, a rate constant of 9.10- 9 min.-I is calculated, given the rate at pH 1.0. This corresponds to a half-time for exchange of ~ 150 a. This result strongly suggests that freshly sedimented marine organic matter would have an isotopic composition near that of the living material, even if the organic material were exposed to water having a greatly differing oxygen isotopic composition. In addition, it can be expected that brief laboratory exposures of biological samples to water would result in virtually no oxygen exchange. On the other hand, on a geological time scale, complete equilibration of the exchangeable pool would be expected. It has been found that a significant fraction (i.e. ~70%) of the oxygen in non-hydrolyzable algal biomass does not undergo isotopic exchange with water under conditions simulating the initial phases of kerogen formation. This implies that insoluble non-hydrolyzable residues of plankton, or "proto-kerogen" formed in aqueous settings, would have oxygen isotopic compositions that would be, in part, dependent upon the original isotopic composition of the living biomass from which the proto-kerogen was derived. The process of oxygen exchange between organic oxygen and water during hydrolytic decomposition of algal material would furthermore result in a partial dependence of the oxygen isotopic composition of the proto-kerogen residue upon the isotopic composition of water in the sedimentary environment. Thus, the oxygen isotopic compositions of recently formed algal kerogens may be relatable both to biochemical processes involving fractionation of oxygen isotopes and to environmental conditions that prevailed during proto-kerogen formation. CONCLUSIONS cC,
(1) In the exchange of oxygen between benzoic acid and water at 80 the equilibrium isotopic composition of the benzoic acid is not equal to
369
that of the water, but is enriched in 18 0 by - 1 7%0' This indicates the occurrence of an equilibrium isotope effect during the exchange reaction, favoring concentration of 18 0 in the benzoic acid. (2) Direct observation of an equilibrium isotope effect favoring enrichment of 18 0 within carboxyl oxygen during hydrolytic exchange suggests that hydration of carboxyl or carbonyl oxygens of intermediates in bioch emical reactions might lead to enrichment of 18 0 in biomolecules, relative to cell water. (3) Results of experiments involving the exchange of oxygen in bluegreen algal biomass with water during hydrolytic decomposition are consistent with the existence of both exchangeable and essentially non-exchangeable oxygen within the insoluble residue. After equilibration of the exchangeable oxygen pool, an isotopic mass balance indicated that 67-70% of the oxygen within the algal residue had not undergone exchange with water. (4) Calculated first-order rate constants for exchange of algal oxygen at pH 1.0 and pH 2.0 are not inconsistent with a mechanism of exchange involving hydration of carboxyl or carbonyl functional groups within the material. (5) The observed dependence of the exchange rate of algal oxygen upon pH suggests that near pH 7 .0, in the absence of other catalysts, exchange would proceed with a half-time measured in hundreds of years. (6) The existence of an essentially non-exchangeable pool of oxygen within algal biomass indicates that "proto-kerogen " , formed in natural environments by hydrolytic de composition of algal biomass, would have an oxygen isotopic composition dependent upon both (1) the isotopic composition of the source material (i.e. living biomass) and (2) the isotopic composition of the water within the sedimentary environment. ACKNOWLEDGEMENTS
The authors are grateful to Dr. Howard Gest and Mr. Jeffrey Favinger for their assistance in preparation of the Spirulina homogenate. They appreciate the technical assistance of Mr. Stephen Studley. This work was supported by Grant NGR 15-003-118 from the National Aeronautics and Space Administration.
REFERENCES Baur, M.E. , 1983. Diversity in Precambrian microbial communities. J. GeoI. Soc., 140: 5-12. Baertschi, P., 1976. Absolute oxygen-18 content of standard mean ocean water. Earth Planet. Sci. Lett. , 31(3) : 341-344 . Bender, M.L., Stone, R .R. and Dewey, R.S., 1956. Kinetics of isotopic exchange between substituted benzoic acids and water. J . Am. Chern. Soc ., 78: 319-321.
370 Blattner, P. and Hulston, J .R., 1978. Proportional variations of geochemical 6'·0 scales - an interlaboratory comparison. Geochim. Cosmochim. Acta, 42 : 59-62. Brenninkmeijer, C.A.M. and Mook, W.G ., 1981. A batch process for direct conversion of organic oxygen and water to carbon dioxide for oxygen-Ld/oxygen-Lf analysis. Int. J. Appl, Radiat. Isot., 32(3) : 137-141. Byrn, M. and Calvin, M., 1966. Oxygen-18 exchange reactions of aldehydes and ketones. J. Am. Chern. Soc., 88(9): 1916-1922. Degens, E.T., 1969. Biogeochemistry of stable carbon isotopes. In: G. Eglinton and M.T .J . Murphy (Editors), Organic Geochemistry, Methods and Results. Springer, New York, N.Y., pp. 304-329. DeNiro, M.J. and Epstein, S ., 1979. Relationship between the oxygen isotope ratios of terrestrial plant cellulose, carbon dioxide, and water. Science, 204: 51-53. DeNiro, M.J. and Epstein, S., 1981. Isotopic composition of cellulose from aquatic organisms . Geochim. Cosmochirn, Acta, 45: 1885-1894. Epstein, S., Thompson, P. and Yapp, C.J., 1977. Oxygen and hydrogen isotopic ratios in plant cellulose. Science, 198: 1209-1215. Ferhi, A.M., Letolle, R. and Lerman, J.C., 1976. Oxygen isotope ratios of organic matter: analyses of natural compositions. In: E.R. Klein and P.D. Klein (Editors), Proc. 2nd Int. Conf. on Stable Isotopes, pp. 716-724. Friedman, 1. and O'Neil, J.R ., 1977 . Compilation of stable isotope fractionation factors of geochemical interest. U.S. Geol. Surv., Prof. Pap. 440-KK. Gonfiantini, R., 1978. Standards for stable isotope measurements in natural compounds. Nature (London), 271 : 534-536 . Hardcastle, K.G . and Friedman, 1., 1974. A method for oxygen isotope analyses of organic material. Geophys. Res. Lett., 1(4) : 165-167 . Hayes, J .M., 1983 . Geochemical evidence bearing on the origin of aerobiosis, a speculative hypothesis. In : J.W. Schopf (Editor), Earth's Earliest Biosphere, Its Origin and Evolution. Princeton University Press, Princeton, N.J. (in press). Hoering, T.C. and Estep, M.L.F. , 1981. Evaluation of a method for measuring the isotopic composition of oxygen in organic matter. Carnegie Inst, Washington, Yearb., 80-81: 410-414. Mook, W.G. and Grootes, P.M., 1973. The measuring procedure and corrections for the high-precision mass-spectrometric analysis of isotopic abundance ratios, especially referring to carbon, oxygen and nitrogen. Int. J. Mass Spectrom. Ion Phys., 12: 273298. Murphy, R.C . and Clay, K.L ., 1981. Synthesis and back exchange of 1·0 labeled amino acids for use as internal standards with mass spectrometry. Biochemistry, 20 : 849· 855. Nissenbaum, A. and Kaplan, 1.R., 1972. Chemical and isotopic evidence for the in situ origin of marine humic substances. Limnol. Oceanogr., 17(4): 570-582. Philp, R.P., 1976. Possible origin for insoluble organic (kerogen) debris in sediments from insoluble cell-wall materials of algae and bacteria. Nature (London), 262: 134136 . Roberts, 1. and Urey , H.C., 1939. Kinetics of the exchange of oxygen between benzoic acid and water. J. Am. Chern. Soc., 61: 258(}-2584. Samuel, D. and Silver, B.L., 1965. Oxygen isotope exchange reactions of organic compounds. In: V. Gold (Editor), Adv. Phys. Org, Chern., 3: 123-186. Schidlowski, M., Hayes, J.M. and Kaplan, 1.R., 1983. Isotopic inferences of ancient biochemistries: carbon, sulfur, hydrogen and nitrogen. In : J.W . Schopf (Editor), Earth 's Earliest Biosphere, Its Origin and Evolu tion. Princeton University Press, Princeton, N.J. (in press). Thompson, P. and Gray, J ., 1977 . Determination of oxygen-18/oxygen-16 ratios in compounds containing carbon, hydrogen, and oxygen. Int. J. Appl. Radiat. Isot., 28(4): 411-415. Wedeking, K.W. and Hayes, J.M. , 1984. Analysis of oxygen isotopes in organic material. Anal. Chern. (submitted).
357
EXCHANGE OF OXYGEN ISOTOPES BETWEEN WATER AND ORGANIC MATERIAL
K.W. WEDEKING and J.M. HAYES
Biogeochemical Laboratories. Departments of Chemistry and of Geology, Indiana Uniuersity , Bloomington , IN 47405 (U.S.A .) (Received April 12, 1983 ; accep ted for pu blication June 14, 1983)
ABSTRACT Wedeking, K.W. and Hayes, J.M. , 1983. Exchange of oxygen isotopes between water and organic material. Isot. GeoscL , 1: 357--370. The exchange of IBO at concentrations near natural abundance between benzoic acid and water and between algal biomass and water has been followed as a function of time and of pH at BODC. Exchange proceeds at conveniently measurable rates at pH 1. 0. The fir st-order rate constant for exchange of 'BO between benzoic acid and water under these conditions is (1.4 ± 0.1) • 10-' min ". An equilibrium isotope effect of 1. 72 ± 0 .09% favoring partitioning of 18 0 into the carboxylic acid is found . The exchange of "0 between wa ter and oxygen-bearing functional groups in algal biomass does not proceed to completion , with only 30-33% of the oxygen in the in ' soluble residue being exchangeable. The first-order rate constant for this exchange is (9 .3 ± 0. 5 ) . 10-' min. - , at 8 0 DC at pH 1.0. If it is assumed that exchange is first order in H+, a half-time of 150 a is calculated for turnover of the exchangeable portion of the oxygen pool at pH 6.0. The existence of a non-exchangeable oxygen pool suggests that isotopic analyses of oxygen in se d im en tary organic matter can be of value in the reconstruction of ancient environments and biochemical pathways.
INTRODUCTION
An extensive literature exists concerning the biogeochemistry of the stable isotopes of carbon, hydrogen and nitrogen. In contrast, the potentially interesting organic geochemistry of oxygen isotopes (i.e. 16 0 and 18 0 ) has been largely ignored, probably for at least two reasons: (1) Until recently (Wedeking and Hayes, 1984), no analytical method has been described that allows highly precise determinations (an uncertainty of 1%0 or better) of the oxygen isotopic compositions of whole cells or of sedimentary organic matter. Previously described methods (Hardcastle and Friedman, 1974 ; Ferhi et al., 1976; Thompson and Gray, 1977 ; Brenninkmeijer and Mook, 1981 ; Hoering and Estep , 1981) have been limited to analyses of pure organic compounds containing only carbon and oxygen.
0167-6695/83 /$03.00
© 1983 Elsevier Science Publishers B.Y.
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(2) Interest in the organic geochemistry of oxygen has been restrained because it is known (e.g., Samuel and Silver, 1965) that organically bound oxygen, in particular the oxygen of carbonyl and carboxyl functional groups, exchanges with water. Thus, it is not surprising that studies of the fractionation of oxygen isotopes within living systems have been limited to that associated with the biosynthesis of cellulose, the oxygen of which is only very slowly exchangeable at physiological pH (Epstein et al., 1977; DeNiro and Epstein, 1979,1981). It is not known whether the oxygen isotopic composition of sedimentary organic matter is related to that of its biological precursors (i.e, carbohydrates, proteins and lipids, or simply, "biomass"). In "the case of the carbon isotopes, the isotopic compositions of humic substances and kerogen formed in marine environments are thought to reflect closely (i.e. within a few per mil) the average isotopic compositions of the living material from which these substances are derived (Degens, 1969; Nissenbaum and Kaplan, 1972). Thus, kerogen contains carbon isotopic information that can be related to biochemical processes. Analyses of the carbon isotopes in ancient kerogens have, for example, led to inferences regarding Precambrian biochemical evolution (Baur, 1983; Hayes, 1983; Schidlowski et al., 1983). Before oxygen isotopic evidence can similarly contribute . to the characterization of ancient biochemistries, the fundamental relationship between the oxygen isotopic composition of kerogen or "proto-kerogen" formed in marine environments and the isotopic composition of biogenically derived precursor material (i.e. planktonic biomass) must be determined. Because a fraction of the oxygen in biomolecules is exchangeable with water, the isotopic composition of oxygen in kerogens formed in marine sediments is likely to be primarily dependent upon both the isotopic composition of the precursor material and the extent of exchange during incorporation into kerogen. If exchange is extensive, biogenically derived isotopic "information" would be overwritten by an abiogenic process. The purpose of this paper is to assess experimentally the extent to which the oxygen in organic matter undergoes isotopic exchange under conditions simulating the initial stages of kerogen formation. The non-hydrolyzable insoluble residue of modern blue-green algae may be representative of the precursor of algal kerogens in ancient sediments (Philp, 1976). The earliest phase of kerogen formation from algal biomass would involve formation of such an insoluble residue by sedimentation and lysis of cells, followed by partial hydrolysis of proteins and glycoproteins, triglycerides, polysaccharides, etc. Hydrolytic decomposition of biopolymers within the sedimentary environment might be catalysed by bacterial enzymes or by clay minerals. Oxygen exchange between the forming residue and water would be likely to occur. In order to estimate quantitatively both the extent of, and the rate of oxygen isotopic exchange that would occur during the formation of algal
359
"proto-kerogen" , acid-eatalyzed hydrolytic decompositions of blue-green algal biomass have been carried out in water enriched in 18 0 . By measuring the oxygen isotopic composition of the insoluble material as a function of time , the kinetics of oxygen exchange between the residue and water have been followed and the dependence of the exchange rate upon pH has been studied. To facilitate oxygen exchange, all experiments were carried D out at 80 C. Results show that even after extensive hydrolytic dissolution of the algal organic matter had occurred, a significant fraction of the oxygen within the insoluble residue had not undergone isotopic exchange. In a control experiment, the kinetics of oxygen isotopic exchange between benzoic acid and water have been followed, as done previously (Roberts and Urey, 1939). At equilibrium, the benzoic acid was enriched in 18 0 relative to the water, thus indicating an equilibrium isotope effect in the exchange reaction. The observation of this isotope effect facilitated mathematical treatment of the results obtained in the exchange of algal organic matter and suggests a possible mode of oxygen isotopic fractionation during biochemical reactions. MATERIALS AND METHODS
Mass spectrometry and calculations
Isotopic analyses of CO 2 were conducted using a dual viscous inlet triple collector isotope ratio mass spectrometer (Nuclide Corp., State College, Penn., model 6.00), using well-established procedures (Mook and Grootes, 1973). Isotopic abundances are reported as delta values, in per mil, vs. standard mean ocean water (SMOW): fj x =
[(R x - R SM O W )/R SM O W ]10 3
where R = 18 0 / 16 0 (R SM OW = 2.0052-10- 3 ; Baertschi, 1976). The mass DC spectrometer is calibrated using CO 2 prepared at 25 from the TKL-l and K-2 standards assuming DTKL-l = 26.47, DK-2 = 3.19 (Blattner and Hulstan, 1978), and Qeo -calcite = 1.01025 (Friedman and O'Neil, 1977). The calibration has been fudependently · confirmed by carrying standard water samples through the sample-preparation procedure employed for organic samples (see p. 360). Using a recently des cribed pyrolytic method (Wedeking and Hayes, 1984), the oxygen of organic samples was converted to carbon dioxide for mass spectrometric analysis of 18 0 /160 ratio. The standard deviation of a single isotopic analysis by this technique is ± 0 .6%0, based on a pooled standard deviation of 17 analyses of pure cane and beet sucrose, reported by Wedeking and Hayes (1984) . The procedure involves no significant oxygen blank « 0.2 umol 0) in analysis of organic samples; the carbon dioxide obtained contains oxygen which is derived from only the sample. It has been shown (Wedeking and Hayes, 1984) that isotopic analyses of
360
biological materials (lung, brain) give precision not different from that obtained in analyses of nitrogen- and sulfur-free, pure organic material (e.g., sucrose). The oxygen of water samples was converted to carbon dioxide by a similar procedure in which excess carbon was added to the sample prior to pyrolysis. This procedure involves a small oxygen blank (8 blank = -10 ± 1%0, amount = 3 ± 0.5 ,.mlol 0; determined by analyzing sucrose samples with similar carbon addition) . The effect of this blank must be taken into account, particularly in the analysis of isotopically enriched waters (for which the isotopic contrast between the sample and the blank is great) . Application of this procedure (C0 2 production and blank correction) in analysis of a 2.2-mg sample of the NBS-l A water standard yielded a result of -24.2 ± 0.6%0; the accepted value for this standard is -24.3 ± 0.3%0 (Gonfiantini, 1978). In mathematical treatment of kinetic data (see Results and Discussion), interconversion between cS -values and fractional abundances of ISO was frequently necessary. Given that:
r; = IS0x/eSOx + 170x + 160x) = [1 + e 70x/ I SOx) + (1 / 1sR x )r 1
(1)
where F x symbolizes the fractional abundance of ISO in oxygen pool x and R x represents the IS0/160 ratio of that pool, we can write:
(2) This useful approximation is obtained by assuming that the 170 /1S0 ratio is constant for all samples and is the same as that in SMOW, that is, 0.186. An approximate expression relating cS x and F x is obtained by substituting eq. 3 into eq. 2:
Rx
=
[10- 38 x + 1] R SM OW
(3)
The resulting expression is accurate to 0.1%0 over the range of enrichments encountered in this study.
Treatment of organic samples Exchange studies were carried out using reagent grade benzoic acid and a culture of the blue-green alga Spirulina. Dry algal powder was suspended in distilled water and the cells were disrupted using a chilled French pressure cell at 1.24.10 5 kN m ". The homogenized material was lyophilized and aliquots of a single preparation were used in the exchange experiments. A sample of ISO-enriched water was divided into fractions and these were acidified to pH 1.0 and pH 2.0 by addition of aqueous 37% HCI. These water samples were used as the exchange media. The isotopic composition of the pH 1.0 water sample was directly determined by the method described above, duplicate analyses giving 8 = 221.3 ± 0.9%0' The aqueous
361
exchange media were assumed to be of identical oxygen isotopic compositions; the amounts of aqueous HCI added to the original water sample were not great enough to affect the total isotopic composition by more than 1%0' The exchange reactions were carried out at 80 ± 1°C, following a procedure described previously (Roberts and Urey, 1939). Samples of benzoic acid and of the algal homogenate were weighed and transferred into reaction vials, 100 ± 2 mg of material being used in all cases . The organic samples were suspended in 8 ml of the appropriate medium which had been . previously heated to 80°C. The vials were tightly capped and shaken, resulting in complete dissolution of the benzoic acid and an initial solubilization of -35% of the algal material (on a dry weight basis). The reaction vials were held at 80°C for 0-174 hr. The supernatant solutions showed no pH change over the reaction periods employed. At appropriate times, vials were removed from the 80°C oven and cooled rapidly to room temperature. The benzoic acid samples precipitated from solution at this time. The tubes were centrifuged and the supernatant solutions discarded. The insoluble algal residues were freeze-dried for 12 hr.; losses by partial hydrolytic dissolution were found to be as great as 80% of the initial dry weight of material. Benzoic acid samples were freeze-dried for only 4 hr. in order to minimize losses by volatilization (as in the procedure of Roberts and Urey, 1939). Separate freeze-drying containers were used to avoid cross-contamination. Oxygen isotopic compositions were determined after samples were free of residual water. RESULTS AND DISCUSSION
Exchange of oxygen between benzoic acid and water Results obtained in the exchange of benzoic acid oxygen with the aqueous medium at pH 1.0 and 80°C are presented in Table 1. Fig. 1 shows the observed oxygen isotopic compositions for benzoic acid (top curve) plotted vs. time in hours. The kinetics of 18 0 exchange between benzoic acid and water have been studied previously by Roberts and Urey (1939), under similar conditions (i.e. 80°C and 0.1 M HCI). These workers found that the kinetics of the exchange reaction are independent of benzoic acid concentration, first order with respect to H+ concentration, and are described by the integrated first-order rate expression:
(4) where k is the first-order rate constant in min. -I; and F j • F t and F w are the fractional abundances of 18 0 in the benzoic acid initially, in the benzoic acid at time t and in the water, respectively. The equation was derived as-
362
- - - -- - - - - -(H20 )- - - - - - - - 200
oL..-..L-....L........L-L.---l_L.-..L-....L........L-L.---I.---....J
o
~O
50
t,hr Fig. 1. Isotopic compositions of organic oxygen vs. time for materials exchanged with I·O-enriched water. All experiments carried out at BOce . {, '" 0 of water = +221.3 ± 0.9%0 vs. SMOW in all experiments ( rJ = benzoic acid, pH 1.0; 0 = algal homogenate, pH 1.0 ; I> = algal homogenate pH 2.0). Solid curues calculated assuming first-order exchange kinetics, as explained in text. TABLE I Exchange of oxygen isotopes between benzoic acid and water at 80 ce * ' Time (hr.)
o'"OSMOW* ' (%0)
F(* 3)
0.0 5 .0 30.0 103.0( 00)
+23 .2 +133 .2 +227.4 +242.0
2.0467 2.2662 2.4541 2.4B32
( x 10 3 )
*, pH 1.0 ; s '"0 of water = +221.3 ± 0.9%0 vs. SMOW. *, Observed isotopic composition of benzoic acid oxygen. *3 Frac tional abundance of ISO in the benzoic acid calculated from observed value of {,'·OSMOW·
suming that the 18 0 is initially concentrated in the water, that the benzoic acid oxygen atoms are equivalent, and that the water represents an infinite exchange pool (i.e, F w is constant). According to eq. 4, the exchange is first order in (Fw-Ft) , that is, the reaction will be complete when the 18 0 abundance of the benzoic acid is the same as that of the water. The rate constant given by eq. 4 is proportional to the H+ concentration and was previously found to be (1.42 ± 0.04)-10- 3 min. -1 (Roberts and Urey, 1939) for 0.1 M HCI at 80°C, giving a half-time for the exchange reaction of 8.14 hr .
363
The results summarized in Table I and plotted in Fig. 1 (top curve) show that the oxygen isotopic composition of the benzoic acid approaches an equilibrium value that is heavier than the water (i.e. the acid is enriched in 18 0 ). An equilibrium isotope effect is, therefore, present. Such an effect was not noted by Roberts and Urey (1939), who studied the reaction only under conditions far from equilibrium; that is, for times less than the halftime of the reaction and with large initial differences in the 18 0 concentrations of the benzoic acid and the water. Oxygen isotopic exchange reactions of organic compounds have been comprehensively reviewed by Samuel and Silver (1965); such an isotope effect, involving the exchange of a carboxylic compound with water, was not reported. No such reports have subsequently appeared in the literature. Given the isotopic compositions of the H 20 and of the benzoic acid at 103 hr. (Table I) the equilibrium fractionation factor for the exchange reaction: H 2180 +l.C6HsC1602H "'" H/ 60 +.!.C6HsCI802H 2 2
(5)
can be calculated. The fractionation factor, abcnzoic acid-water = R acid / R water ' calculated according to Friedman and O'Neil (1977), is found to be 1.0172 ± 0.0009 for the reaction at 80°C.
A modified rate expression for benzoic acid oxygen exchange. Since an implicit assumption in the derivation of eq. 4 is that the oxygen isotopic composition of the benzoic acid will be equal to that of the water at equilibrium, a modified expression, reflecting the approach of the 18 0 abundance in the benzoic acid to its equilibrium value, is needed. The modified form of eq. 4 should be first order in the difference (Fe-Ft), where Fe represents the final (equilibrium) fractional abundance of 18 0 in the benzoic acid: (6)
In eq. 6, k, t, F i and F t are as in eq. 4 but F w (the fractional abundance of 18 0 in the water) has been replaced by Fe. Note that Fe can be calculated knowing both the isotopic composition of the water and the equilibrium isotope effect for the exchange reaction. For the data of Table I, taking Fe = 2.4832.10- 3 , a plot of-log (Fe-FI) vs. time allows calculation of the first-order rate constant from the slope of the line obtained. The rate constant found is (1.4 ± 0.1)'10- 3 min. -I, in close agreement with the previously reported value (Roberts and Urey, 1939). Substitution of the calculated rate constant into eq. 6 and solution for F t allows calculation of 0 180 for the benzoic acid as a function of time. As seen in Fig. 1, the curve obtained falls close to the points corresponding to benzoic acid.
Biogeochemical significance. The mechanism of oxygen isotopic exchange between carboxylic acids and water is thought to involve hydration of the
364
acid to give a symmetrical intermediate (i.e the ortho-acid) , which can dehydrate to give the exchanged product (Bender et al., 1956). A similar mechanism, involving a symmetrical hydrated intermediate, is responsible for the exchange of aldehyde and ketone oxygen with water (Byrn and Calvin, 1966). The direction, if not the magnitude, of the isotope effect observed here may be representative of the effect that would be associated with carbonyl oxygen exchange at 80°C. At lower temperatures, equilibrium fractionations commonly increase. The direct observation of an isotopic fractionation in the exchange of oxygen in benzoic acid suggests that biosynthesized organic molecules would be enriched in 18 0 relative to cell water if oxygen within intermediates or products had undergone exchange via carbonyl or carboxyl hydration. Interestingly, DeNiro and Epstein (1981) have found that the cellulose synthesized by aquatic plankton is uniformly enriched in 18 0 , relative to cell water, by an average of 27%0' Thus, the results obtained here support these workers' suggestion that hydration of carbonyl functional groups of intermediates during the biosynthesis of cellulose is responsible for the observed fractionation. It seems likely that a similar mechanism of isotopic fractionation would play a role in determining the oxygen isotopic comp osit ions of biosynthesized products such as proteins and fatty acids.
Exchange of oxygen in algal biomass Results obtained in the hydrolytic decompositions of algal organic matter at 80°C and at both pH 1.0 and pH 2.0 are presented in Table II. The oxygen isotopic compositions of the aqueous media were unchanged from experiments involving benzoic acid . The table giyes the measured total organic oxygen isotopic compositions of the insoluble residues and the corresponding times of reaction; the isotopic compositions are shown plotted vs. reaction time in Fig . 1 (middle and lower curves, corresponding to reaction at pH 1.0 and pH 2.0, respectively ). Also given in Table II are the percentages of starting material recovered, calculated on a dry weight basis.
Qualitative interpretation of results. Fig. 1 (lower two curves) shows that the 18 0 abundances of the insoluble residues increased relatively smoothly with time, with the rate of isotopic enrichment being greater at lower pH. A parallel decrease in the percent recovery of starting material occurred at pH 1.0, while recovery of material was relatively constant at pH 2.0 (Table II). Thus, both dissolution of the material and 18 0 enrichment are acid-catalyzed processes. This suggests that dissolution of material over time (besides that fraction initially soluble) proceeded by acid-catalyzed hydrolytic decomposition of biopolymers. Oxygen isotopic exchange between water and the organic residue would accompany hydrolysis, thus accounting for, at least in part, the dependence of 18 0 enrichment rates upon pH. Of course, oxygen exchange could occur without hydrolytic
365 TABLE II Exchange of oxygen isotopes between algal homogenate and water at 80°C* I Tim e (h r .)
°(%0) tat 18 0 SMOW *'
0ex 18 0 SMOW*' (%0)
Mat er ial rec overy*4
+20.3 +45 .0 +49 .6 +66.3 +79 .3 +85.9
+ 20.3 +103.8 +119.3 +175.7 +219.6 +242.0
+66 +56 + 47 +4 1 +33 +26
+26.5 +28.9 +35.4
+ 41.2 + 49.4 + 71.3
+74 +64 +63
( %)
pH 1.0 : 0.0 5.0 10.0 20. 0 40 .0 10 3.0 pH 2.0 : 10 .0 40 .0 103.0
* ' 0" 0 of water = +221.3 ! 0.9%0 vs. SMOW. *, Ob served isotopic compositi on of total organic oxygen within insolu ble algal res idues. *3Iso top ic com p ositio n of exchangeable fr ac t ion of oxygen with in algal residue. Calcu late d as ex p lain ed in te xt. *4 Calcu lated relative to amoun t o f starti ng material, on a dry weight basis .
dissolution of material, bu t in this case exch ange would still be favored at lo wer pH. The observed enrichments of 18 0 with time are probably caused by exchange of oxygen isotopes between the 180-enriched water and the insoluble residue. Po ssible alternative mechanisms include: (1 ) dissolution of organic matter containing oxygen that is depleted in 18 0 relative to the average isotopic composition of the material ; (2) loss of material togeth er with a kinetic fra ctionation of oxygen isot opes in the residue; and (3) oxidation of the material by O 2 in equilibrium with water. Reactions were carr ied out in air -tight containers with minimal head space, thus (3) may be ignored. Kinetic effects during hydrolytic dissolution of material may be responsible for a portion of the observed 18 0 enrichments with time . Such effects, if present, are difficult to distinguish from exchange, and have not been separately treated. Two lines of evidence indicate that selective losses of 180-depleted material did not account for a significant fraction of the observed enrichments: (1) at pH 2.0 , enrichments of up to 25%0 were observed without loss of material (relative to recovery at zero time), thus at this pH , exchange alone had occurred ; and (2 ) oxygen eleme n tal abundances of the residues showed no monotonic variations with time at either pH , with averages of 23 ± 5 and 24 ± 1 wt. % O 2 (± 1a) being found for the residues obtained at pH 1.0 and at pH 2.0 , respectively . Thus progressive loss of an isotopically distinct, single component (i.e.
366
protein, lipid or carbohydrate) is not indicated, since each of these components have, on average, distinctly different oxygen elemental abundances. It appears that solubilization of "average" organic material had occurred. It is, then, reasonable to assume that oxygen isotopic exchange is principally responsible for the observed variations in 18 0 abundances of the algal residues. In the exchange of algal oxygen at pH 1.0, an asymptotic isotopic composition was essentially reached after 103 hr. (Fig. 1), suggesting attainment of equilibrium in the reaction. The isotopic composition after this apparent equilibration is significantly depleted in 18 0 relative to the water. The equilibrium difference of 135%0 (H 20 - residue) is too large to have been caused by an isotope effect; exchange of oxygen between organic compounds and water would result in enrichment of the heavy isotope within the organic compound, as opposed to the observed depletion. Thus, it must be concluded that a significant fraction of the oxygen within the algal residue had not undergone isotopic exchange after 103 hr. at pH 1.0. This conclusion, in conjunction with the apparent attainment of equilibrium, requires the existence of at least two pools of organic oxygen with greatly differing exchange rates. A kinetic model that assumes the existence of both exchangeable and completely non-exchangeable pools of organic oxygen within the insoluble algal residue adequately fits the observed data, as shown below. In this model, the relative sizes of the two pools are constant. Non-exchangeable oxygen might exist as non-hydrolyzable amides and esters (e.g., in proteins and triglycerides), as well as hydroxyl oxygens of polysaccharides. The exchangeable oxygen pool probably consists of ketones and as free carboxylic acid and aldehyde groups that, in part, may have been present as ester, amide and acetal groups prior to hydrolytic exchange.
Quantitative treatment of results. Calculation of the relative sizes of the exchangeable and non-exchangeable oxygen pools can be accomplished using the mass-balance equation: (7)
where 8 tot, 8 ex and 8 nex represent the isotopic compositions of the total algal oxygen pool, the exchangeable and the non-exchangeable oxygen pools, respectively; and where f ex is the fraction of the total oxygen residing in exchangeable functional groups within the insoluble residue. It is assumed that f ex remains constant. Eq. 7 would be exact if 8 -values were replaced by the corresponding fractional abundances, but the approximation as written is accurate to 0.1%0 over the range of isotopic compositions encountered here. Calculation of the fraction of exchangeable oxygen within the algal residue is accomplished as follows. The value of 8tot = +85.9% 0 of Table II, at a time of 103 hr., gives the total isotopic composition of the algal ma-
367
terial after complete equilibration of the exchangeable pool with water. By making the reasonable assumption that the non-exchangeable oxygen pool has an isotopic composition always given by the total isotopic composition at time zero (i.e. that both pools had the same initial isotopic composition), only the equilibrium value of [j ex is needed to calculate the relative pool sizes. The value of [j ex after complete exchange is probably close to that of the water (i.e. +221%0 vs. SMOW). An isotope effect like that encountered in the benzoic acid study might, however, intervene. This possibility gives rise to some uncertainty in the calculated value of feX' Assuming the equilibrium value of [j ex to be the same as the water gives a value of 33% for the relative percentage of exchangeable oxygen, according to eq. 7. On the other hand, assuming the exchangeable pool to be comprised entirely of "carboxyl-like" oxygen, implies that [j ex would approach an equilibrium value of +242%0 vs. SMOW, as observed directly in the case of benzoic acid exchange. In this case, a value of 30%, only slightly different from 33%, is obtained for the relative size of the exchangeable pool. Given the observed total isotopic compositions of Table II, the isotopic composition of the exchangeable oxygen pool can be calculated at each time of reaction, by use of eq. 7. This has been carried out by assuming that the exchangeable pool is "carboxyl-like", that is, f ex = 0.33 and that {j ex = 242%0 at equilibrium; the values obtained are given in Table II, and are symbolized by [j ex 180SMOW' The data summarized in Table n can be used to calculate rate constants for the exchange of algal oxygen, assuming first-order kinetics, according to eq. 6. The procedure used is identical to that described for the case of benzoic acid exchange; values of -log(Fe - F t ) are calculated and plotted vs. time. Linear plots were found for the algal oxygen exchange at both pH 1.0 and pH 2.0, thus verifying the occurrence of first-order kinetics. A rate constant of (9.3 ± 0.5).10- 4 min."! is calculated for the exchange at pH 1.0, while at pH 2.0 a rate constant of (3.7 ± 0.8).10- 5 min.T ' is found. Thus, the exchange of algal oxygen is close to being first order in H+ concentration. If the calculated rate constants are substituted into eq. 6, solving for F t allows calculation of the isotopic composition of the exchangeable oxygen pool as a function of time at both pH 1.0 and pH 2.0. Substitution of calculated isotopic compositions into eq. 7 then allows calculation of the total isotopic compositions with time. Curves calculated in this way fall close to the observed points, as seen in Fig. 1 (middle and lower solid curves). Since the isotopic composition of exchangeable oxygen in the algal organic matter has been found to follow eq. 6, it would seem that the exchangeable oxygen pool is kinetically homogeneous. That is, exchangeable oxygen pools with greatly differing exchange rates are apparently not present. The observed rate constant for algal exchange at 80 and pH 1.0 is within a factor of 1.5 of that of benzoic acid under identical GC
368
conditions, and is close to the exchange rates of the amino acids serine, lysine and leucine recently found, under similar conditions, by Murphy and Clay (1981). Since the exchange of the algal material involved an insoluble organic oxygen pool, a direct comparison of the observed exchange rates with those of compounds studied in solution may not be informative. Thus, an unequivocal interpretation of the observed exchange rates in terms of the functional group(s) involved is not possible. However, the observed rates of exchange and the dependence upon pH are consistent with a mechanism involving hydration of carboxyl or carbonyl functional groups.
Geochemical significance of algal exchange studies. The minimal time required for isotopic equilibration of exchangeable oxygen in algal biomass under conditions less acidic than those employed in this study can be crudely estimated by assuming that the rate constant for the exchange of algal organic matter is first order with H+ concentration. For example, at pH 6.0 and 80cC, a rate constant of 9.10- 9 min.-I is calculated, given the rate at pH 1.0. This corresponds to a half-time for exchange of ~ 150 a. This result strongly suggests that freshly sedimented marine organic matter would have an isotopic composition near that of the living material, even if the organic material were exposed to water having a greatly differing oxygen isotopic composition. In addition, it can be expected that brief laboratory exposures of biological samples to water would result in virtually no oxygen exchange. On the other hand, on a geological time scale, complete equilibration of the exchangeable pool would be expected. It has been found that a significant fraction (i.e. ~70%) of the oxygen in non-hydrolyzable algal biomass does not undergo isotopic exchange with water under conditions simulating the initial phases of kerogen formation. This implies that insoluble non-hydrolyzable residues of plankton, or "proto-kerogen" formed in aqueous settings, would have oxygen isotopic compositions that would be, in part, dependent upon the original isotopic composition of the living biomass from which the proto-kerogen was derived. The process of oxygen exchange between organic oxygen and water during hydrolytic decomposition of algal material would furthermore result in a partial dependence of the oxygen isotopic composition of the proto-kerogen residue upon the isotopic composition of water in the sedimentary environment. Thus, the oxygen isotopic compositions of recently formed algal kerogens may be relatable both to biochemical processes involving fractionation of oxygen isotopes and to environmental conditions that prevailed during proto-kerogen formation. CONCLUSIONS cC,
(1) In the exchange of oxygen between benzoic acid and water at 80 the equilibrium isotopic composition of the benzoic acid is not equal to
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that of the water, but is enriched in 18 0 by - 1 7%0' This indicates the occurrence of an equilibrium isotope effect during the exchange reaction, favoring concentration of 18 0 in the benzoic acid. (2) Direct observation of an equilibrium isotope effect favoring enrichment of 18 0 within carboxyl oxygen during hydrolytic exchange suggests that hydration of carboxyl or carbonyl oxygens of intermediates in bioch emical reactions might lead to enrichment of 18 0 in biomolecules, relative to cell water. (3) Results of experiments involving the exchange of oxygen in bluegreen algal biomass with water during hydrolytic decomposition are consistent with the existence of both exchangeable and essentially non-exchangeable oxygen within the insoluble residue. After equilibration of the exchangeable oxygen pool, an isotopic mass balance indicated that 67-70% of the oxygen within the algal residue had not undergone exchange with water. (4) Calculated first-order rate constants for exchange of algal oxygen at pH 1.0 and pH 2.0 are not inconsistent with a mechanism of exchange involving hydration of carboxyl or carbonyl functional groups within the material. (5) The observed dependence of the exchange rate of algal oxygen upon pH suggests that near pH 7 .0, in the absence of other catalysts, exchange would proceed with a half-time measured in hundreds of years. (6) The existence of an essentially non-exchangeable pool of oxygen within algal biomass indicates that "proto-kerogen " , formed in natural environments by hydrolytic de composition of algal biomass, would have an oxygen isotopic composition dependent upon both (1) the isotopic composition of the source material (i.e. living biomass) and (2) the isotopic composition of the water within the sedimentary environment. ACKNOWLEDGEMENTS
The authors are grateful to Dr. Howard Gest and Mr. Jeffrey Favinger for their assistance in preparation of the Spirulina homogenate. They appreciate the technical assistance of Mr. Stephen Studley. This work was supported by Grant NGR 15-003-118 from the National Aeronautics and Space Administration.
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