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Geochemistry Geophysics Geosystems 3 Article G Volume 5, Number 8 5 August 2004 Q08001, doi:10.1029/2004GC000713 AN ELECTRONIC JOURNAL OF THE EARTH SCIENCES Published by AGU and the Geochemical Society ISSN: 1525-2027 New insights into the burial history of organic carbon on the early Earth Christian J. Bjerrum Geological Institute, University of Copenhagen, Øster Voldgade 10, DK-1350 Copenhagen K, Denmark ([email protected]) Donald E. Canfield Danish Center for Earth system Science (DCESS) and Institute of Biology, University of Southern Denmark, Campusvej 55, DK-5230 Odense, Denmark ([email protected]) [1] The isotope record of organic matter and calcium carbonate is often used to infer the burial history of organic carbon through time. As organic carbon burial is widely held to control long-term oxygen production, the isotope record also relates to the production rates of oxygen on Earth. Current interpretations of the record suggest a long-term consistency in the proportion of total carbon buried as organic carbon ( f ratio), with some important periods of much higher burial proportions. The isotope record is analyzed here with a new carbon isotope mass balance model, which considers submarine hydrothermal weathering of ocean crust as a significant removal pathway of inorganic carbon. With this model the f ratio is considerably reduced if isotopically depleted inorganic carbon is precipitated during hydrothermal weathering and if hydrothermal weathering dominates inorganic carbon removal from the surface environment. In contrast to previous calculations, our analysis of the carbon isotope record shows that organic carbon burial in the Archean accounted for only between 0% and 10% of the total carbon burial. These low burial proportions would have contributed to a slow accumulation of atmospheric oxygen in the Archean. Components: 5711 words, 4 figures. Keywords: Archean; carbon isotopes; mass balance; model; Precambrian. Index Terms: 1010 Geochemistry: Chemical evolution; 1040 Geochemistry: Isotopic composition/chemistry; 4805 Oceanography: Biological and Chemical: Biogeochemical cycles (1615) Received 6 February 2004; Revised 25 April 2004; Accepted 25 June 2004; Published 5 August 2004. Bjerrum, C. J., and D. E. Canfield (2004), New insights into the burial history of organic carbon on the early Earth, Geochem. Geophys. Geosyst., 5, Q08001, doi:10.1029/2004GC000713. 1. Introduction [2] The history of organic carbon burial is usually deduced from steady state carbon isotope mass balance models. In constructing these models it is assumed that inorganic carbon and organic carbon are removed from the surface mixed layer of the oceans into underlying sediments [Hayes, 1983; Hayes et al., 1999; Kump, 1991; Des Marais et al., 1992; Des Marais, 2001]. This type of model, applied to the Precambrian carbon isotope Copyright 2004 by the American Geophysical Union record, has lead to the prevailing view of a long term consistency through time in the proportion of total carbon buried as organic matter [e.g., Schidlowski, 1988; Watanabe et al., 1997; Catling et al., 2001; Kump et al., 2001; Holland, 2002]. This does not, however, mean that the burial flux of organic carbon has remained the same. Higher heat flow in the Archean most likely was accompanied by a higher flux of CO2 from the mantle [Des Marais, 1985, 1997; Sleep and Zahnle, 2001]. Greater CO2 degassing would have been 1 of 9 Geochemistry Geophysics Geosystems 3 G bjerrum and canfield: burial history of organic carbon accompanied by greater total carbon burial. Therefore associated with the long term consistency in burial proportions of organic carbon is the prospect that early Archean O2 production rates were even higher than today, which is paradoxical with evidence for low atmospheric oxygen concentrations [Holland, 2002]. We question, however, an important basic premise of this model. In particular, we relax the assumption that inorganic carbon is only removed from the surface ocean reservoir. This assumption is not strictly valid today, as bicarbonate is extracted from the deep ocean during hydrothermal alteration of ocean crust (ocean crust carbonatization) [Brady and Gislason, 1997; Wallmann, 2001], and most likely it was not true early in Earth history [Walker, 1990]. [3] Thus, on the very early Earth, before significant continental crust formed, nearly all buffering of volcanic CO2 degassing must have been through ocean crust carbonatization (OCC) [Walker, 1990]. Even after significant amounts of continental crust had formed, the presence of ubiquitous hydrothermal carbonates in Archean greenstone belts suggests that ocean crust carbonatization was still an important CO2 sink [Veizer et al., 1989b; Nakamura and Kato, 2002]. Model studies of the early carbon cycle furthermore indicate that through the Archean OCC probably accounted for between 30% and 90% of the inorganic carbon removal [Sleep and Zahnle, 2001]. Ocean crust carbonatization influences carbon isotope mass balance when the carbonates so precipitated are depleted in 13C relative to the surface waters. This occurs, for example, when there is an isotope gradient from the surface to the deep ocean. Today, the decay of 13C-depleted organic matter in the deep ocean results in a 2% surfaceto-deep difference in the isotopic composition of DIC (dissolved inorganic carbon) [Broecker and Peng, 1982]. Therefore any carbonate precipitated in isotopic equilibrium [Alt, 1999] with waters at depth due to OCC will be on average 2% d13C-depleted relative to carbonates precipitated in the surface ocean [Des Marais, 2001]. Other processes discussed below also could have contributed to 13C depletion during OCC. [4] In Archean and early Proterozoic rocks an isotope contrast is observed between carbonates precipitated from surface and deep water. The ancient carbonate rocks can be divided into two broad groups with different isotopic compositions depending on where they were precipitated. Inorganic carbon precipitated from the surface ocean has d13C values around 0%, whereas carbonates of 10.1029/2004GC000713 hydrothermal origin, formed from massive carbonatization, have more 13C-depleted mean values of 1.3% [Des Marais, 2001; Veizer et al., 1989b]. Furthermore, typical micritic siderites in deepwater BIF (banded iron formation) facies have d13C values around 3% to 5% indicating deeper marine waters 13 C-depleted compared to surface water [Beukes et al., 1990; Kaufman et al., 1990]. The observations, then, point to two isotopically distinct burial sinks for inorganic carbon in the Archean, surface-derived carbonates with a d13C of around 0% and deep water-derived carbonates with relatively 13C-depleted values compared to the surface ocean. In what follows we will focus our discussion on OCC as it probably represented the most important deep-water sink for inorganic carbon [Walker, 1990; Sleep and Zahnle, 2001; Nakamura and Kato, 2002]. We recognize that early diagenetic carbonate formation in deep water sediments, fuelled by the decomposition of organic matter, could also remove isotopically depleted DIC, and that this process would also contribute to isotope balance in the same manner as we will explore below. [5] Thus deep ocean carbonate removal likely represented a sink for 13C-depleted carbon in early Earth history. This sink influences isotope mass balance, and importantly, the isotopic composition of surface DIC, independent of organic carbon burial. We develop a new carbon-isotope mass balance model including deep-ocean carbonate precipitation by way of OCC. We then calculate the impact of this on organic carbon burial as deduced from the Precambrian carbon isotope record [Des Marais, 2001; Shields and Veizer, 2002]. 2. Three Sink Carbon Isotope Mass Balance Model [6] Steady state mass balance requires that the flux of carbon into the ocean-atmosphere system (Fin) balances the total flux of inorganic carbon (Fb) and organic carbon (Fbo ) out of the system [e.g., Hayes, 1983; Des Marais, 2001]: Fin Fbo þ Fb ¼ 0: ð1Þ Total inorganic carbon removal is the sum of two district fluxes; Fb = Fb,sc + Fb,oc, where Fb,sc is the burial flux of sedimentary carbonate carbon and Fb,oc is the removal flux of OCC carbon (Figure 1). The carbon removed by OCC can be written as a 2 of 9 Geochemistry Geophysics Geosystems 3 G bjerrum and canfield: burial history of organic carbon 10.1029/2004GC000713 isotopic differences the following isotopic mass balance is obtained: din dsc Db f lDs ð1 f Þ ¼ 0 Figure 1. New carbon isotope mass balance model. Removal of inorganic carbon (Fb) occurs as sedimentary stromatolithic carbon and inorganic precipitates in the surface reservoir, but also in the deep ocean through ocean crust carbonatization (OCC). The proportion of DIC removed by OCC is given by l. Fob is the flux of organic carbon buried, and Fin is the flux of dissolved inorganic carbon (DIC) into the oceans. In the Archean, significant OCC probably resulted in a greater deep ocean inorganic carbon sink than today. fraction (l) of the total inorganic carbon removed, so Fb = (1 l)Fb + lFb, and the isotopic mass balance of equation (1) may then be written: Fin din Fbo dorg þ ð1 lÞFb dsc þ lFb doc ¼ 0; ð2Þ where doc is the isotopic composition of inorganic carbon removed by OCC and dsc is the isotopic composition of inorganic carbon precipitated and buried in shallow water. din is the isotopic composition of carbon entering the ocean-atmosphere system, and dorg is the isotopic composition of buried organic carbon. Dividing equation (2) through by Fin leads to din f dorg ð1 lÞð1 f Þdsc lð1 f Þdoc ¼ 0; ð3Þ where f = Fob (Fin)1 is the fraction of total carbon buried as organic matter ( f ratio), and following from equation (1), (1 f ) = Fb (Fin)1. [7] In principle, the isotopic difference between the inorganic carbon removed by OCC and the carbonate precipitated from the surface ocean can be constrained from observations, and this difference is given as Ds = doc dsc. Likewise, the isotopic difference between organic carbon and inorganic carbon removed from the surface ocean can be constrained from observations as is given by Db = dorg dsc. [e.g., Hayes, 1983; Schidlowski, 1988; Kump, 1991; Des Marais et al., 1992]. With these ð4Þ The last term on the left-hand side in equation (4) is new compared to previous isotope mass balance equations [e.g., Hayes, 1983; Kump, 1991; Des Marais, 2001]. This term designates the importance of OCC as an inorganic carbon sink (l), whose significance depends on the magnitude of l and the isotopic difference between inorganic carbon removed in the surface and deep ocean (Ds). [8] If l = 0, then all inorganic carbon is removed from the surface reservoir, and equation (4) is identical to previous isotope models [e.g., Hayes, 1983; Kump, 1991; Des Marais, 2001]. In this limit the isotopic composition (dsc) of sedimentary inorganic carbonate removed from the surface ocean is given only by din, Db, and f, and it is independent of the internal processing of carbon in the ocean [Kump, 1991]. This is similarly true if there is no 13 C gradient in the ocean, meaning Ds = 0%. [9] For a constant value of f, and assuming a Ds of 2%, as today, the isotopic composition of DIC removed from the surface ocean (dsc) increases by around 2% as l increases from 0 to 1 (Figure 2a). This increase is independent of organic carbon burial. The increase in dsc results from mass balance considerations and occurs as 13C-depleted doc (compared to dsc) starts to dominate the inorganic sinks (as l increases; Figure 2). Therefore the isotopic composition of carbonate removed from surface oceans is not a unique indicator of the burial proportion of organic carbon, f, as long as l > 0. An even larger effect on dsc is observed, as a function of l, with a larger Ds (Figure 2b). With Ds = 5%, as has been suggested for the Archean and early Proterozoic (see below), increasing l from 0 to 1 increases dsc by around 4%. Such a significant influence of l on dsc shows that previous calculations of the f ratio in the Precambrian are potentially very uncertain, and a reevaluation of the carbon isotope record is thus required. 3. Precambrian Carbon Burial [10] The history of the fraction of carbon buried as organic carbon ( f ) over geologic time can be analyzed by rearranging equation (4): f ¼ ðdin dsc lDs ÞðDb lDs Þ1 : ð5Þ In using this equation, dsc and Db are obtained directly from the isotopic record, Ds is potentially 3 of 9 Geochemistry Geophysics Geosystems 3 G bjerrum and canfield: burial history of organic carbon 10.1029/2004GC000713 floor weathering [Walker, 1990; Nakamura and Kato, 2002]. Furthermore, recent carbon cycle models indicate significant early Archean OCC, with l values ranging between 0.35 and 0.95, and decreasing to 0.25 by the end of the Proterozoic Figure 2. The isotopic composition of carbonate carbon precipitated in the surface ocean (dsc) contoured as a function of l, the relative importance of deep ocean DIC removal, and f, the fraction of total carbon buried as organic matter. (a) Contoured using equation (4) with Db = 30 %, the Phanerozoic mean (equation (2)) and with Ds = 2 %, as today. (b) The same as in Figure 2a, but with Ds = 5 %. Increasing l significantly decreases the f ratio associated with a given isotopic composition of the surface carbonates (Figure 2b). constrained from observations, and l is constrained from independent carbon cycle modeling. As in previous f ratio calculations, we need not consider what happens to the carbon after it is buried [Hayes, 1983; Schidlowski, 1988; Des Marais et al., 1992; Hayes et al., 1999]. In other words, we assume that the isotopic composition of volcanic and weathered lithosphere carbon (din) remains constant over time, with a value of 5% [Haggerty, 1999]. To obtain values for dca and Db, the isotope database record of Des Marais [2001] and Shields and Veizer [2002] was filtered with a Butterworth filter with a half power point frequency cutoff of (0.3Gyr)1, after the d13C record was first averaged in 10 Myr intervals and linearly interpolated when no data were present in a time interval (Figures 3b and 4b). [11] Archean submarine basalts are extensively carbonatized indicating probable significant sea- Figure 3. Modeled Archean-Proterozoic carbon cycle (3.5 Giga-annum before present (Ga) to 0.7 Ga). (a) Conservatively calculated decrease in l the fraction of DIC removed though ocean crust carbonatization (OCC), with methane as an important greenhouse gas in the early Precambrian [Sleep and Zahnle, 2001]. (b) The isotopic composition of inorganic carbon (top, dsc, from Shields and Veizer [2002]) and organic carbon (bottom, from Des Marais [2001]) are shown through the Precambrian. Solid curves are the filtered d13C record. (c) Different possible values of Ds, the isotope difference between carbonate removed from the surface and deep ocean reservoirs. The solid and dash-dot curves are extrapolations of Ds estimated from isotope composition of Archean rock (see text). The dashed line is the present mean ocean Ds. (d) Calculated fraction of total carbon buried as organic matter ( f ) from equation (5), where Db is the difference between the lower and upper curve in Figure 3a, din = 5%, and l is from Figure 3b. Ds is a free parameter, where the three curves are as in Figure 3c (dashed, solid, and dash-dot). The fine dotted line is for l = 0, as in equation (4). In all cases the calculated f ratio is lower than previously interpreted (thin dashed line, f 0.2). 4 of 9 Geochemistry Geophysics Geosystems 3 G bjerrum and canfield: burial history of organic carbon 10.1029/2004GC000713 [Sleep and Zahnle, 2001]. The lower estimates of l were calculated assuming that methane was an important Archean greenhouse gas, and that consequently, atmospheric CO2 concentrations were relatively low (Figure 3a) [e.g., Pavlov et al., 2000; Sleep and Zahnle, 2001]. These models assume a constant and modern continental crust size. The mass of the continental crust, and its availability to remove CO2 by weathering, will influence the evolution of l. For this reason we have recalculated from the equations of Sleep and Zahnle [2001] the evolution of l for the methane greenhouse scenario (with low atmospheric CO2) and assuming continental growth. To do this, the silicate weathering expression of Sleep and Zahnle [2001] was multiplied by the normalized continental crust volume reconstruction of Kamber et al. [2003] where the crust reached 70% of its present volume by 2 Ga (Figure 4a). (The free parameters of Sleep and Zahnle [2001] were set to Fridge = 2.4 1012 mol/yr, a = 0.9, and thydro = 3 106 year.) In this example very high l values of near 1 are observed in the early Archean, decreasing to around 0.39 at the end of the Precambrian. The model results all point to significant Archean OCC and the prospect of an important second sink for DIC from the ocean-atmosphere system. If the isotopic composition of the OCC sink differs from the surface reservoir sink there will be, as we have seen above, an error in the calculation of the f ratio if the OCC sink is not considered. Figure 4. Modeled Archean-Proterozoic carbon cycle with continental growth. (a) Recalculated decrease in l, the fraction of DIC removed though ocean crust carbonatization (OCC) with methane as an important greenhouse gas and continental growth up to 2 Ga (calculation after Sleep and Zahnle [2001]). (b) As in Figure 3b, solid curves are the filtered d13C record. (c) Different possible values of Ds, the isotope difference between carbonate removed from the surface and deep ocean reservoirs, are presented, where the solid and dash-dot curves are extrapolations of Ds estimated from isotope composition of Archean rocks (see text). The dashed line is the present mean ocean Ds. (d) Calculated fraction of total carbon buried as organic matter (f) from equation (5), where Db is the difference between the lower and upper curves in Figure 4a, din = 5 %, and l is from Figure 4b. Ds is a free parameter, where the three curves in Figure 4c (dashed, solid, and dash-dot) are used to calculate three possible f ratio curves. Fine dotted line is for l = 0, as in equation (4). The solid line represents our best estimate for the f ratios in the Archean and early Proterozoic for the given l. [12] To ascertain the isotopic composition of the OCC sink, we must evaluate Ds. As mentioned above, Archean carbonatized rocks of various ages, which are considered part of the regional lower temperature carbonatization halo, have a mean d13C value of 1.3 ± 1.6%, and an end-member seawater(?) value of +2.2% [Veizer et al., 1989b]. The regional halo is sedimentologically and geochemically distinct from carbonate in sepentinized shear zone and conduit rocks with d13C = 4 ± 2%, which probably had strong affinity to a deepseated mantle source. Veizer et al. [1989a] also analyzed sedimentary carbonates in Archean greenstone belts and found mean d13C value of +1.5 ± 1.5% [Veizer et al., 1989a]. The carbonates formed from OCC, and the sedimentary carbonates, are isotopically different at the 99.9% level. These values imply therefore a Ds 2.8%. The diagenetic alteration of Precambrian carbonates studied by Veizer et al. [1989a, 1989b] is thought to be small on the basis of trace element and strontium isotope considerations. However, progressive recrystallization during hydrothermal 5 of 9 Geochemistry Geophysics Geosystems 3 G bjerrum and canfield: burial history of organic carbon circulation has led to continues loss of Sr and a small increase in the d13C values in OCC carbonates. The sedimentary carbonates have also apparently experienced a small decrease in d13C values upon alteration. Thus the estimated Ds is conservative and originally could have been larger (i.e., more negative). [13] Typical micritic siderites in deep-water BIF (banded iron formation) facies are also isotopically depleted with d13C values around 5%, ranging up to around 3% [Beukes et al., 1990; Kaufman et al., 1990]. In contrast, contemporaneous shallow-marine carbonates have d13C values mostly around 0%, suggesting Ds values of between 3 to 5% at this time, approximating the Ds value indicated from averages of carbonatized Archean rocks. [14] Various processes could have contributed to a large Ds in the Archean and early Proterozoic. A large Ds could have resulted from an equally large vertical isotope gradient in marine DIC. The factors establishing such a large vertical isotope gradient are not well understood but could include reduced ocean circulation and/or a more efficient biological isotope-organic pump. These would act to increase the accumulation of carbon oxidation products in the deep ocean reservoir leading to a larger vertical isotope DIC gradient [Broecker and Peng, 1982]. For example, today, DIC in deep waters of the stratified Black Sea and Kyllaren Fjord are, respectively, depleted by 6% and 15% relative to surface waters [Fry et al., 1991; Smittenberg et al., 2004]. The oxidation of 13C-depleted organic matter from chemoautotrophic carbon fixation at the chemocline [Hayes et al., 1999], or from the oxidation of methane across the chemocline [Hayes, 1994], could also help establish a larger vertical isotope gradient in DIC. The processes acting to establish a large vertical isotope gradient in DIC in the Archean would, however, need to have been quite pronounced. This is because higher marine DIC concentrations would have been associated with higher atmospheric CO2 concentrations, and this would decrease the importance of the organic pump relative to DIC mixing [Kump and Arthur, 1999]. Other aspects of the carbon cycle, admittedly speculative, may potentially have contributed to the removal of 13Cdepleted DIC during OCC. [15] Thus Rothman et al. [2003] proposed that Precambrian anoxic waters may have accumulated very high concentrations of dissolved organic carbon (DOC) compared to present-day oxic marine bottom water. They argued that the dissolved 10.1029/2004GC000713 fermentation products of organic matter settling into an anoxic deep ocean would degrade slowly allowing DOC to accumulate to higher concentrations. We speculate that elevated concentrations of DOC would circulate through ocean crust where microbial fermentation, and high temperature decarboxylation of organic compounds with magnetite, would result in the liberation of CO2, CH4, and other organic decarboxylation products [McCollom and Seewald, 2003]. Elevated CO2 could precipitate as calcite or siderite depending on fluid chemistry and represent a sink for carbon. The prospect would have existed, therefore, for the introduction of 13C-depleted DIC into altered ocean crust, although the magnitude of this process is difficult to constrain. [16] The direct extraction of mantle CO2 into carbonatized basalts during OCC is expected to be minimal, as in more recent ocean crust such carbonate generally have precipitated in equilibrium with seawater DIC [Alt, 1999; Alt and Teagle, 2003]. By contrast, in recent basalts organic carbon oxidation can contribute 13C-depleted carbon to hydrothermal carbonates when these are intercalated with sediments [Alt and Teagle, 2003]. A similar process could have introduced 13C-depleted carbon into ancient OCC carbonates. However, this still amounts to the introduction of 13C-depleted carbonate into the crust and it would influence the isotope systematics of the ocean system in the same manner as outlined above. [17] We obviously have an incomplete understanding of what controlled Ds early in Earth history. Processes such as deep ocean siderite precipitation, and in particular OCC, acted as probably significant removal pathways of DIC. Also, direct isotope observations suggests that the deep carbonates were isotopically distinct from surface water carbonates with a likely isotope difference, Ds, of 3 to 5% in the Archean and early Proterozoic oceans. This isotope difference would have generated a sink with the possibility of influencing the isotopic composition of DIC in the surface reservoir independently of organic carbon burial. In what follows we explore the evolution of the f ratio through the Precambrian with a variety of Ds values ranging from today’s value of 2%, to 5%, and finally to an absolute maximum value of 7.5% (Figures 3c and 4c). [18] We begin with the most conservative scenario yielding the smallest isotope effect on the f ratio. This is the scenario assuming a methane-rich atmosphere and constant present-day continental 6 of 9 Geochemistry Geophysics Geosystems 3 G bjerrum and canfield: burial history of organic carbon size, giving the lowest values for l in the Archean (Figure 3b). In this case f ratios straddle 0.07 for most of the Archean with a Ds of 5%, are as low as 0.05 with Ds of 7.5%, and are around 0.1 for a Ds of 2% (Figure 3d). If we assume continental growth through early Earth history much higher Archean values for l are generated, yielding much lower f ratios (Figure 4d). In this case f ratios are around zero in the Archean for Ds < 5%, and are still substantially reduced, compared to the case without OCC, with modern Ds = 2%. Importantly, for all l and Ds values explored, the f ratios are much lower during the Archean and early Proterozoic than at present ( f 0.2), implying low proportions of total carbon buried as organic carbon. [19] There is a pronounced positive excursion in the f ratio between 2.4 and 2.0 Ga, as has been previously noted [Des Marais et al., 1992; Knoll et al., 1986; Karhu and Holland, 1996]. Our results, however, reduce the magnitude of the excursion compared to what is calculated in the absence of deep ocean carbonate removal (l = 0; upper dashed curve, Figures 3d and 4d) by amounts depending on Ds and the evolution of l. Relatively high f ratios for the Neoproterozoic are also preserved with the absolute value depending, again, on the model scenario. 4. Perspective [20] There is a long-standing opinion that the f ratio through Precambrian was generally similar to today [Schidlowski, 1988; Catling et al., 2001; Watanabe et al., 1997; Holland, 2002]. This is not supported by our results [see also Des Marais et al., 1992]. However, significant removal of isotopically depleted DIC from deeper ocean waters can have a large influence on the f ratio, and it is likely that the f ratio was less than half the present ratio during nearly all of the Archean, and could have been as low as zero for some of the time. Low f ratios mean less organic carbon burial, and imply relatively lower rates of oxygen release to the surface environment [Holland, 1984; Lenton and Watson, 2000]. Low rates of organic carbon burial could have acted alone, or more likely together, with high fluxes of mantle-derived reduced gases [Kump et al., 2001; Holland, 2002] to explain the low concentrations of O2 in the Archean atmosphere [Holland, 1999; Canfield et al., 2000; Farquhar et al., 2000]. While including OCC into our model reveals lower Archean f ratios than previously recognized, 10.1029/2004GC000713 marine phosphorus limitation could have been one of the factors contributing to low f ratios [Bjerrum and Canfield, 2002]. [21] As pointed out by Des Marais [2000] and Towe [1990], even modest organic carbon burial depends on the activities of oxygen-producing organisms and a carbon cycle involving aerobic respiration. Thus, while our analysis involves no assumptions about the presence of oxygenproducing or oxygen-utilizing organisms, even f ratios as high as 0.05 would apparently require their presence [Des Marais, 2001]. By contrast, our less conservative f ratio results, with values straddling zero (Figure 4d), make no predictions about the presence of cyanobacteria in the early Archean environment, and such low f ratios could conceivably be produced independent of oxygenic photosynthesis. [22] We view this contribution as a starting point. Clearly the consideration of deep ocean DIC precipitation, and in particular OCC, in global carbon cycle modeling has important consequences for our interpretation of f ratios in the past. This becomes extremely important in the early Precambrian when OCC was likely most significant. Our model results depend critically on Ds (the isotope difference between OCC carbon and shallow water derived-inorganic carbon), and the history of l (the proportion of total carbon removed by OCC). These values will be better constrained with focused efforts to better identify the magnitudes of OCC preserved in ancient ocean crust, and the magnitude of past differences between OCC and shallow sedimentary carbonate. Likewise such studies will help identify the processes responsible for the isotopic composition of OCC carbonates. Notation OCC ocean crust carbonatization. f fraction of total carbon buried as organic carbon. d13C isotopic composition of carbon reported relative to the PDB standard. d13C = (Rsample/RPDB 1) 1000, where R = 13 12 C/ C is the isotopic abundance ratios. 13 d d C in abbreviated form. DIC dissolved inorganic carbon. l fraction of total inorganic carbon removed by OCC. 7 of 9 Geochemistry Geophysics Geosystems 3 G bjerrum and canfield: burial history of organic carbon Ds isotopic difference between inorganic carbon removed by OCC and carbonate precipitated from the surface ocean water, %. Ga giga-annum before present. Acknowledgments [23] We thank D. J. Des Marais for allowing us use his d13C database, N. H. Sleep for sharing unpublished model results on ocean crust carbonatization though time, and T. M. McCollom and M. T. Rosing for insight in decarboxylation, serpentinization, and hydrothermal systems. We furthermore thank two reviewers for very helpful comments on an earlier version of the manuscript and two journal reviewers for helpful suggestive comments. This work was funded by the Danish National Research Foundation (Danmarks Grundforskningsfond) and Danish National Science Foundation (SNF). References Alt, J. C. 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