Fractionation between inorganic and organic carbon during ... · an open-marine deltaic setting and...

Earth and Planetary Science Letters 271 (2008) 278–291

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Fractionation between inorganic and organic carbon during the Lomagundi(2.22–2.1 Ga) carbon isotope excursion

A. Bekker a,⁎, C. Holmden b, N.J. Beukes c, F. Kenig d, B. Eglinton b, W.P. Patterson b

a Geophysical Laboratory, Carnegie Institution of Washington, 5251 Broad Branch Rd., N.W., Washington, DC 20015 USAb Saskatchewan Isotope Laboratory, Department of Geological Sciences, University of Saskatchewan, 114 Science Place, Saskatoon, Saskatchewan, S7 N 5E2, Canadac Department of Geology, University of Johannesburg, Auckland Park 2006, South Africad Department of Earth and Environmental Sciences, University of Illinois at Chicago, Chicago, IL 60607, USA

⁎ Corresponding author. Present address: DepartmUniversity of Manitoba, Winnipeg, MB, R3T 2N2 Canad+1 204 474 7623.

E-mail address: [email protected] (A. Bekker)

0012-821X/$ – see front matter © 2008 Elsevier B.V. Adoi:10.1016/j.epsl.2008.04.021

A B S T R A C T
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The Lomagundi (2.22–2.1 G

Received 25 November 2007Received in revised form 7 April 2008Accepted 10 April 2008Available online 26 April 2008

Editor: H. Elderfield

Keywords:Precambriancarbon cycleLomagundi Eventcarbon isotope fractionation between organicand carbonate carbonocean redox state

a) positive carbon isotope excursion in shallow-marine sedimentary carbonateshas been associated with the rise in atmospheric oxygen, but subsequent studies have demonstrated that thecarbon isotope excursion was preceded by the rise in atmospheric oxygen. The amount of oxygen released tothe exosphere during the Lomagundi excursion is constrained by the average global fractionation betweeninorganic and organic carbon, which is poorly characterized. Because dissolved inorganic and organic carbonreservoirs were arguably larger in the Paleoproterozoic ocean, at a time of lower solar luminosity and lowerocean redox state, decoupling between these two variables might be expected.We determined carbon isotope values of carbonate and organic matter in carbonates and shales of theSilverton Formation, South Africa and in the correlative Sengoma Argillite Formation, near the border inBotswana. These units were deposited between 2.22 and 2.06 Ga along the margin of the Kaapvaal Craton inan open-marine deltaic setting and experienced lower greenschist facies metamorphism. The prodelta tooffshore marine shales are overlain by a subtidal carbonate sequence. Carbonates exhibit elevated 13C valuesranging from 8.3 to 11.2‰ vs. VPDB consistent with deposition during the Lomagundi positive excursion. Thetotal organic carbon (TOC) contents range from 0.01 to 0.6% and δ13C values range from −24.8 to −13.9‰.Thus, the isotopic fractionation between organic and carbonate carbon was on average 30.3±2.8‰ (n=32) inthe shallow-marine environment. The underlying Sengoma shales have highly variable TOC contents (0.14 to21.94%) and δ13C values (−33.7 to −20.8‰) with an average of −27.0±3.0‰ (n=50). Considering that theshales were also deposited during the Lomagundi excursion, and taking δ13C values of the overlyingcarbonates as representative of the δ13C value of dissolved inorganic carbon during shale deposition, a carbonisotope fractionation as large as ~37‰ appears to characterize the production of bulk organic matter in thedeeper part of the Pretoria Basin at that time. This enhanced fractionation relative to that observed inshallow-water environments likely reflects heterotrophic (secondary) and chemotrophic productivity at andbelow a pronounced redoxcline, consistent with the euxinic conditions inferred from independent evidencefor the deeper part of the Pretoria Basin. Greater variability in organic carbon vs. carbonate carbon isotopicvalues on the shallow-marine carbonate platform suggests that the carbon cycling was dominated by a largedissolved inorganic carbon reservoir during the Lomagundi excursion.Our study suggests that in contrast to the Late Neoproterozoic and Phanerozoic, when carbon isotopefractionation between carbonate and organic carbon in the open ocean was mostly controlled by primaryproducers, in the Paleoproterozoic redox-stratified ocean heterotrophic and chemotrophic productivityoverprinted a signal of primary productivity below the redoxcline. This strong imprint of heterotrophic andchemotrophic productivity on organic carbon isotope records complicates the reconstruction of spatialpatterns and secular trends in the δ13C values of dissolved inorganic carbon in the Paleoproterozoic seawater.

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1. Introduction

Subsequent to its discovery within carbonates of the Paleoprotero-zoic Lomagundi Group, Zimbabwe (Schidlowski et al., 1975, 1976), thepositive carbon isotope excursion in shallow-marine sedimentarycarbonates, later named the Lomagundi Event, has been observed in anumber of Paleoproterozoic basins worldwide (e.g. Baker and Fallick,

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279A. Bekker et al. / Earth and Planetary Science Letters 271 (2008) 278–291

1989a,b; Karhu, 1993; Karhu and Holland, 1996; Melezhik and Fallick,1996; Bekker et al., 2006a). Although the details of secular carbonisotope variations in Paleoproterozoic seawater are as yet poorlyconstrained, these chemostratigraphic studies clearly demonstratedthat the Lomagundi carbon isotope excursion started before 2.22 Ga,shortly after the Paleoproterozoic glacial epoch, and endedbetween2.11and 2.06 Ga. Accordingly, carbon isotope values of seawater remainedhigh for more than 100 Ma of Paleoproterozoic history (Karhu andHolland, 1996; Melezhik et al., 2007). Atmospheric oxygenation waslinked genetically to the Lomagundi Event by Karhu andHolland (1996).More recently, however, Bekker et al. (2004) provided geochronologicalevidence that the rise in atmospheric oxygen occurred earlier. However,regardless of this genetic relationship, steady-state carbon isotopemassbalance constraints predict that a large amount of oxygen (~12 and 22times the present atmospheric level) was released to the atmosphereover the 100 Ma duration of the Lomagundi Event (Karhu and Holland,1996). Thewide range of values mostly reflects a poor knowledge of theglobal average carbon isotope fractionation between carbonates andorganic carbon for this time interval. Unfortunately, carbonate succes-sions of this age are generally very lean in organic matter, preventingdetermination of carbonate and organic carbon isotope values on thesame samples. An additional complication is the finding of Karhu andHolland (1996) that the carbon isotope record of organic-rich shalesdeposited in the deeper-water settings appears not to record theLomagundi Event. This led Hayes and Waldbauer (2006) to reinterpretthe Lomagundi Event as a period of globally-enhanced methaneproduction below the marine sediment–water interface during theprogressive oxidation of diagenetic environments accompanying therise in atmospheric oxygen; an explanation that strongly contrasts withthe earlier model based on high relative burial rates of organic carbon(e.g. Karhu and Holland,1996). Because the age of these shales is poorlyconstrained, we cannot be entirely sure that they are in factstratigraphically equivalent to the carbonate deposits that record theLomagundi Event in the shallower-water settings.

The Precambrian ocean might have had large dissolved inorganiccarbon (DIC) and dissolved organic carbon (DOC) reservoirs (Rothmanet al., 2003; Bartley and Kah, 2004); the latter suggestion was used toexplain decoupled organic and carbonate carbon records in the earlyNeoproterozoic (Fike et al., 2006). The DOC reservoir, following Fikeet al. (2006), refers to a mass of dead organic carbon suspended in thePrecambrian ocean that includes colloidal, particulate, and truly‘dissolved’ organic matter. The concept of a high DOC reservoir in thePrecambrian ocean was originally developed by Logan et al. (1995),who argued for a redox-stratified ocean in the Late Precambrian as aresult of slow depositional rates of organic matter before the advent ofmetazoan grazers and fecal pellet production. Logan et al. (1995)inferred that suspended organicmatterwas processed byheterotrophsin awater columnwith a prominent redoxcline. Because the biomass ofheterotrophs is 13C-enriched compared to primary producers, thecontribution of heterotrophs to TOC may be recognized if specificbiomarkers are present or if TOC displays elevated δ13C values. Forexample, the TOC of the Terminal Neoproterozoic deeper-water shalesis 13C-enriched relative to sediments deposited in shallower-watersettings, consistentwith a larger contribution of heterotrophs toTOC inthe deeper-water settings where degradation of primary biomasswould have been more advanced (Logan et al., 1999).

The concept of a large DIC reservoir in the early Precambrianseawater is based on the inferred lower solar luminosity at that timerequiring higher atmospheric CO2 levels to maintain greenhouseconditions (Grotzinger and Kasting, 1993) and finds supportingevidence in carbonate textures such as inorganic precipitates (tufa),synsedimentary cement crusts, and micrites (e.g. Grotzinger, 1994;Kah and Knoll, 1996). The required atmospheric CO2 level at ca.2.3 Ga was estimated at 0.03 to 0.3 bars, which is 100 to 1000 timeslarger than the present atmospheric CO2 level (Grotzinger andKasting, 1993).

Another related issue is the relatively common assumption thatthe isotopic fractionation between organic and carbonate carbon (Δδ)during the Precambrianwas constant (e.g. Eichmann and Schidlowski,1975; Schidlowski and Aharon, 1992). This assumed constantfractionation is used as a test for post-depositional alteration (e.g.Knoll et al., 1986; Kaufman and Knoll, 1995) and for the reconstructionof carbonate carbon isotopic values from the organic carbon isotoperecord (e.g. Calver, 2000; Karlstrom et al., 2000; Walter et al., 2000;Kaufman and Xiao, 2003). In both cases, fractionation between 25 and30‰ is commonly used as a canonical range of values despite theobservation of temporal variations in this parameter in the Neopro-terozoic and Phanerozic rock record (Hayes et al., 1999). Large DIC andDOC reservoirs as well as redox stratification in the Precambrian oceanshould have had a dramatic effect on isotopic fractionation betweencarbonate and organic carbon, possibly leading to the decoupling oftheir records both locally and globally.

To address these issues, we present new δ13C data on both thecarbonate and organic fractions in subtidal carbonates of the 2.22–2.06 Ga Silverton Formation (Pretoria Group, South Africa) and thecorrelative Sengoma Argillite Formation, Segwagwa Group acrossthe border in Botswana (Figs. 1 and 2). These data are compared tothe δ13Corg values in the underlying organic-rich shales of the SengomaArgillite Formation. Buick et al. (1998) found a positive δ13Ccarbexcursion in carbonates of the lower part of the Silverton Formationfrom the northeastern part of the Transvaal Basin (Button, 1973; Fig. 2)within the contact-metamorphic aureole of the 2.06 Ga BushveldComplex, which they correlated to the 2.22–2.1 Ga Lomagundi carbonisotope excursion. Because both the Silverton–Sengoma carbonatesand Sengoma shales were deposited during the Lomagundi Event,δ13Ccarb and δ13Corg datawill be used to test whether the carbonate andorganic carbon records are coupled, and whether enhanced methanecycling alone can account for the massive and prolonged shift inδ13Ccarb values recorded in shallow-marine carbonates between 2.22and 2.1 Ga.

1.1. Regional geology and stratigraphy

The Silverton Formation belongs to the Paleoproterozoic PretoriaGroup (Fig. 1) deposited in the open-marine epicontinental TransvaalBasin in South Africa (Fig. 2). The formation is bracketed in age by theunderlying ca. 2.22 Ga Hekpoort Lava and the overlying or intruding2.06–2.05 Ga Rooiberg Felsite Group and Bushveld Complex, respec-tively (Walraven, 1997; Buick et al., 2001; Dorland, 2004). Furtherindirect age constraints might be inferred from the ca. 2.14–2.12 Ga U–Pb ages of authigenic metamorphic monazite in metashales, metasilt-stones, and quartzites of the Chuniespoort Group and the olderCentral Rand Group of the Witwatersrand Supergroup reflectingrecently recognized tectonic event in the Transvaal Basin (e.g.Rasmussen et al., 2007). This pre-Bushveld compressional event fold-ed the upper part of the Pretoria Group including the Silverton and theoverlying Magaliesberg formations in the far western part of theTransvaal Basin into open upright symmetrical folds (Hartzer, 1995,2000; Bumby et al., 1998). Furthermore, units in the upper part of thePretoria Group, above the Magaliesberg Formation overlying theSilverton Formation, are less mature with respect to the lower part ofthe Pretoria Group deposited on the passive continental marginindicating change to an active tectonic regime and a sedimentprovenance to the north of the Kaapvaal craton (Button, 1973, 1986;Schreiber et al., 1992). These data suggest that the Silverton Formationwas deposited during the 2.22–2.1 Ga Lomagundi carbon isotopeexcursion and most likely shortly before the 2.14–2.12 Ga tectonicevent. We therefore assume a ca. 2.15 Ga age for deposition of theSilverton Formation.

In our study area, in the northwestern part of the Transvaal Basin,the Silverton Formation experienced only lower greenschist faciesmetamorphism. The deltaic to offshore marine Silverton Formation is

Fig. 1. Correlation chart between the Transvaal Supergroup in the Bushveld Basin of southeastern Botswana and in the northwestern part of the Transvaal Basin, South Africa(modified from Key, 1983 and Bekker et al., 2004 with the Rooihoogte–Timeball Hill age from Hannah et al., 2004 and carbon isotope data from this paper). The age of the HekpoortFormation is inferred based on the correlationwith the Ongeluk Andesite of the GriqualandWest Basin and its whole rock Pb–Pb age (Cornell et al., 1996) together with the age of theyoungest inherited zircons reported by Dorland (2004) for the Hekpoort Formation.

280 A. Bekker et al. / Earth and Planetary Science Letters 271 (2008) 278–291

sandwiched between shallow-marine tidally-influenced quartz sand-stones of the Daspoort and Magaliesberg formations (Button, 1973;Fig. 1). The lower contact with the Daspoort Formation is sharp

Fig. 2. Schematic map of the Late Archean and early Paleoproterozoic sedimentary succChuniespoort (dark grey) and Pretoria (light grey) groups in the Transvaal structural basinstructural basin of South Africa taken from the digital 1:1,000,000 geology maps of South Afsub-Kalahari formations (Council for Geoscience, 1997; Key and Ayres, 2000; Haddon, 2001).Kanye (K) structural basins, Botswana. All these structural basins used to belong to the sameSilverton Formationwith high carbon isotope values occur (Button, 1973; Buick et al., 1998). ‘Africa and Botswana, respectively. ‘L’ refers to the location of sequence-stratigraphically corrwith high carbon isotope values.

whereas the upper contact with the Magaliesberg Formation isgradational. The Silverton Formation is divided by basaltic lava intotwo thick units of shales and siltstones with minor chert and

essions of South Africa, showing outcrop and sub-Kalahari sand distribution of theand Ghaap (dark grey) and Postmasburg (light grey) groups in the Griqualand West

rica and Botswana with the additional information from the digital 1:2,500,000 map ofSimilar color patterns are also used to label the correlative units in the Bushveld (B) anddepositional basin. ‘S1’ indicates the location where carbonates in the lower part of theS2’ and ‘Se’ point to the locations of sampled carbonate and black shale sections in Southelative Lucknow Formation in the Griqualand West Basin that also contains carbonates


carbonate layers. It has a thick carbonate sequence at the top in thenorthwestern part of the Transvaal Basin at the border between SouthAfrica and Botswana. The Silverton Formation is known in Botswanaas the Sengoma Argillite Formation (Key, 1983; Figs. 1 and 2). Ourstudy is focused on the carbonate sequence in both South Africa andBotswana and on the immediately underlying black shales of theupper unit in the drill core Strat 2 located near Lobatse in Botswana.The formation is between 500 and 700m thick, whereas the carbonatesequence is up to 167 m at Sengoma Hill in Botswana (Key, 1983).

Fig. 3. Measured partial section of the Silverton Formation, South Africa near the border withshown as well as carbon isotope fractionation between carbonate and organic carbon (Δδ). I

Two partial (~30 and 50 m thick) sections of the carbonatesequence were sampled in South Africa 1 to 2 km from the borderwith Botswana (Figs. 2 and 3). Carbonates, interlayered with shales,are partially dolomitized, laminated, and rarely display the featuresof a shallow-water depositional environment such as wave ripples,domal stromatolites, silicified nodules, soft-sediment deformationstructures, water and gas escape structures, or small-scale cross-bedding. Another partial (46 m thick) section of the carbonatesequence was sampled at Sengoma Hill, Botswana (Fig. 4). This

Botswana with stratigraphic variations in carbonate and organic carbon isotope valuesnsets shown scatter plots for carbonate and organic carbon isotope values vs. Δδ values.


carbonate section displays wave ripple cross-laminations and flatlaminations. In the lower part of this section sheet cracks (Fig. 5) areextensively developed and filled with recrystallized carbonate.These likely developed during early diagenesis as methane and CO2produced during remineralization of organic matter escaped fromsediments.

Fig. 4. Measured partial carbonate section of the Sengoma Argillite Formation, Botswana in tcarbon isotope values shown as well as carbon isotope fractionation between carbonate andvalues vs. Δδ values.

The drillcore Strat 2 (see Key, 1983) stored at the GeologicalSurvey of Botswana was also logged and sampled (Fig. 6). It consistsof two upward-shallowing cycles with organic-rich pyriferousshales in the lower parts of the cycles deposited in a prodeltaenvironment above wave base and hematite-rich quartz sandstonesat the top of the cycles deposited in a delta plain environment. Veins

he type section of Sengoma Hills with stratigraphic variations in carbonate and organicorganic carbon (Δδ). Insets shown scatter plots for carbonate and organic carbon isotope

Fig. 5. Sheet cracks in carbonates of the Sengoma Argillite Formation at Sengoma Hills.


with carbonates, quartz, migrated pyrobitumen, and remobilizedsulfides occur locally as, for example, in the 107–115 m depth interval.

2. Methods

The least altered (i.e. lacking veins, discoloration, weathering rinds,and silicification) and finest-grained portions of polished thicksections of carbonates from the Silverton Formationwere microdrilledwith 2 mm in diameter diamond drill bits and reacted at 70°C using aKiel III carbonate device directly coupled to a Thermo Scientific 253 gasisotope ratio mass-spectrometer in the Saskatchewan Isotope Labora-tory, University of Saskatchewan. Each set of 6 samples was bracketedby an internal calcite standard calibrated against the (NIST) NBS-19standard. Conversion to the VPDB scale was performed using thevalues −2.20‰ and 1.95‰ for δ18O and δ13C, respectively, for (NIST)NBS-19. The external uncertainty (±1σ) based on NBS-19 analyses is0.10‰ for δ18O and 0.05‰ for δ13C. Carbonates from the SengomaFormation were microdrilled and reacted at 70 °C using a GasBenchcarbonate device connected to a Thermo Scientific 252 instrument atthe Geophysical Laboratory, Carnegie Institution of Washington. Eachset of 6–7 samples was bracketed by the Geological Survey of Finlanddolomite standard (Tytyri dolomite; Karhu, 1993) and conversion tothe VPDB scale was performed using the values 0.78‰ and −7.07‰ forδ18O and δ13C, respectively, for this internal standard calibrated againstthe (NIST) NBS-19 and NBS-20 standards. The external uncertainty(±1σ) is based on the reproducibility of the Tyryri dolomite, which isbetter than 0.2‰ for δ18O and δ13C. Major and minor element analysesfor both sets of samples were performed by ICP-AES using Thermo Irisinstrument. The reproducibility of the analyses is ±5% (1σ).

Samples for δ13Corg and TOC analyses were prepared by decarbona-ting whole rock powders with 1 N HCl, followed by a thorough rinsingin distilledwater and drying in an oven at 50 °C for 24 h. δ13Corg analyseswere conducted on the Thermo Scientific Delta Plus XL instrumentcoupled to a Carlo Erba NC2500 Elemental Analyzer at the Environ-mental Isotope Geochemistry Laboratory, University of Illinois atChicago. Samples marked with an asterisk in the Supplementarymaterials, Table 1 were prepared and analyzed at the GeophysicalLaboratory, Carnegie Institution of Washington. Aliquots of powderedsamples between 1 and 15 mg in weight, depending on TOC content,were decarbonatedwith 6 NHCl inmuffled silver boats, dried overnightin a hood and, subsequently, in an oven at 50 °C, and analyzed on theThermo Scientific Delta Plus XL instrument coupled to a Carlo ErbaNC2500 Elemental Analyzer. All stable isotope values are reportedin per mil vs. VPDB and USGS-40, USGS-41, caffeine UIC, and acetanilide(Costech Analytical Technologies) standards were used to monitorexternal and internal uncertainty that was better than 0.2‰. Totalorganic carbon abundances were calculated based on carbon contentmeasured with Elemental Analyzer and loss during decarbonation.

2.1. Geochemical data and preservation of primary δ13C values

Carbonate samples analyzed are dolostones, limestones, and marlswith xenotopic mosaics of anhedral, coarsely-crystalline to verycoarsely-crystalline grains with pseudospar, and rarely contain quartzor carbonate veins. Carbonates have a narrow range of δ13C valuesbetween +7.6 and +11.2‰ VPDB (Supplementary materials, Table 1).δ18O values show a much larger range from −16.7 to −5.2‰ VPDB. Acorrelation between δ13Ccarb and δ18Ocarb values, which has often beenattributed to metamorphic decarbonation reactions (e.g. Valley, 1986;Baumgartner and Valley, 2001), is not present in our sample set. Thecarbonates have variable TOC contents ranging from 0.01 to 0.60% anda large range in δ13Corg values from −24.8 to −13.9‰. There is nocorrelation between TOC content and δ13Corg values as would beexpected from the breakdown and volatilization of organic moleculesby diagenetic and metamorphic processes (Hayes et al., 1983). Carbonisotope exchange between graphite and carbonate accounts for 13C-enrichment in graphite associated with marble, however graphiteforms at temperatures above 500–600 °C (Valley, 2001), well abovethe lower greenschist facies metamorphism experienced by carbo-nates of the Silverton and Sengoma Argillite formations. Δδ values arealso highly variable, ranging between 23.9 and 33.6‰, although theaverage (30.3‰) is close to 30‰, that is typical for Phanerozoic andPrecambrian unmetamorphosed carbonates (Eichmann and Schi-dlowski, 1975; Knoll et al., 1986; Hayes et al., 1999). Mn and Srcontents and Mn/Sr ratios are also widely used to evaluate diageneticeffects in carbonate rocks (e.g. Kaufman et al., 1993). Carbonate sam-ples analyzed in this study have Mn contents ranging from 27 to1505 ppm (most samples contain less than 670 ppm), Sr contents varyfrom 26 to 119 ppm, and Mn/Sr ratios are between 0.5 and 26.5(Supplementary materials, Table 1). These data are consistent withcalcite or dolomite as the precursor mineralogy, rather than aragonite,and indicate only a minor to moderate amount of post-depositionalalteration with respect to other Paleoproterozoic carbonates world-wide (e.g. Bekker et al., 2001, 2003a,b, 2005, 2006a; Melezhik et al.,1999; Veizer et al., 1992).

The organic matter in the middle part of the Silverton Formationcarbonate section, between 15 and 32 m from the base (Fig. 3), is 13C-enriched, with δ13C values of −18.8 to −13.9‰, compared to backgroundvalues from −22.5 to −19.1‰. Although this stratigraphic interval has, onthe whole, the lowest TOC contents in the section, there are someindividual samples from outside of this interval with even lower TOCcontent that do not show a tendency towards 13C-enrichment comparedto neighboring samples (see Supplementary materials, Table 1). Inaddition, it is difficult to interpret this interval as having been subjectedto a higher degree of post-depositional alteration since the δ13Ccarb and

Fig. 6. Partial section of the Sengoma Argillite Formation, Botswana from the drill core Strat 2 with stratigraphic variations in organic carbon isotope values in two upward-coarseningcycles shown. Inset shows scatter plot for organic carbon isotope values vs. TOC content. Two samples that were strongly altered are shown with empty circles.


δ18Ovalues, and traceelementabundancesdonot appear themselves tobeunusual. Alternatively, the observed 13C-enrichment through this strati-graphic interval might reflect high phytoplankton growth rates in theupwelling area where nutrients were delivered from the deep ocean. Asimilar interpretationwas offered to explain the presence of 13C-enrichedorganic matter in stromatolitic phosphorites and carbonates of the ca.2.0 Ga Jhamarkotra Formation, Aravalli Supergroup, Rajasthan, India(Banerjee et al., 1986). In our case, the relatively low TOC levels in thisstratigraphic interval seem to be in conflict with the high productivitymodel unless organic matter was reworked from the high productivityarea and transported to this depositional site. Regardless of the underlyingcause of the 13C-enriched organic matter in this part of the section, it isimportant to note that there is no correlation between δ13Corg values andTOC contents in this section (see Supplementary materials, Table 1) and

even if these data were removed from the graph of δ13Corg vs. Δδ on theinset in Fig. 3, a significant correlation with a similar slope would still beevident, suggesting that data from this 13C-enriched interval most likelyreflect primary values.

δ13C values of carbonates in the sampled sections of the SilvertonFormation are on average 1‰ more positive than carbonates sampledat Sengoma Hills (9.6 vs. 8.6‰; Figs. 3 and 4). Because these sectionsare relatively close to each other and are partial sections of muchthicker carbonate sequence in both areas, we relate this discrepancy toa slight difference in the stratigraphic position of the sections withinthe carbonate sequence rather than a shelf gradient in δ13CDIC.

The underlying shales of the Sengoma Argillite Formation havehigher TOC content, ranging from 0.14 to 21.94 wt.%, and highlyvariable δ13Corg values from −33.7 to −20.8‰ with the majority of


values and the average value (−27.0‰) lower than −25‰ (Fig. 6;Supplementary materials, Table 1). The basal part of the section hasthe lowest δ13C values. The overlying strata become progressively 13C-enriched upsection reaching a maximum at 219 m. At this point thetrend reverses with δ13C values declining to 205 m. From 205 m untilto the top of the section there are no variations in δ13Corg values. Twoshale samples at the top of the section between 107 and 115 m depthcontain veins with carbonates, quartz, migrated pyrobitumen, andremobilized sulfides. Based on their elevated δ13Corg values, thesesamples appear to be more significantly altered than the sampleswithout veins (Fig. 6).

In summary, both the petrographic features and geochemical datasuggest that carbonate and organic carbon fractions in the carbonatesand shales were not significantly altered following deposition.Therefore, we conclude that δ13Ccarb values reflect the δ13C value ofDIC in contemporaneous seawater in the Pretoria Basin. δ13Corg valuescorrespond to the composition of organic matter deposited in thisbasin andmay also reflect contemporaneous seawater DIC δ13C values,offset by some fractionation. Whether or not this was indeed the casedepends on the relative contribution of organic matter sourced fromprimary (photosynthetic) and secondary (e.g. methanotrophic andsulfur-oxidizing) pathways that indirectly reflects redox stratificationand CO2, CH4, and S cycling in the basin.

3. Discussion

3.1. Do high δ13Ccarb values of the Silverton–Sengoma carbonates reflectlocal or global conditions?

Accepting that the Silverton–Sengoma δ13Ccarb values closely ap-proximate seawater δ13CDIC values in the Pretoria epicontinental sea,the extent to which local scale carbon cycling processes may haveoverprinted the global-scale signature of the surface ocean remainsunclear. Processes on both scales have been inferred to interpret thePhanerozoic carbonate carbon isotope records (e.g. Kump and Arthur,1999; Melchin and Holmden, 2006). Modern (e.g. Florida Bay, BahamaBanks, and Shark Bay) and Phanerozoic carbonates show a large rangein δ13C values that might reflect vital effects of carbonate-secretingorganisms, local environmental conditions (e.g. hypersalinity, biolo-gical productivity, terrestrial runoff, and diagenetic mineralization oforganic matter), and shelf gradients in epicontinental basins with therestricted connection to the open ocean. Local scale effects were citedby Melezhik et al. (1999) as an explanation for very positive δ13Ccarbvalues (N+6 to +8‰) recorded by stromatolitic carbonates that mayhave been deposited during the Lomagundi Event in potentiallyhypersaline settings. Because we are interested in comparing frac-tionations between organic and inorganic carbon in sediments fromshallow and deeper parts of the Pretoria basin, and we lack directconstraints on carbon isotope values of DIC in the deeper part of thebasin, the following discussion is focused on whether the Silverton–Sengoma δ13Ccarb values reflect local or global signals. If global signalsare implied, δ13Ccarb values in the shallow part of the basin can be usedas a rough estimate for the carbon isotope values of seawater DIC inthe deeper part of the Pretoria basin.

Several models have been proposed to explain positive δ13Ccarbexcursions in modern and Phanerozoic shallow-marine carbonatesthat are relevant to this discussion. These include i) weathering ofcarbonates during sea level lowstands, and the impact of this on theδ13C value of terrestrial runoff (Kump and Arthur, 1999; Melchin andHolmden, 2006); ii) enhanced productivity in restricted settings (e.g.Swart and Eberli, 2005); and iii) 13C-enrichment in evaporating brines(Stiller et al., 1985; but see Lazar and Erez, 1990 for a different view).With respect to the carbonate weathering model, the Silverton andSengoma Argillite formations do not directly overlie an extensivecarbonate platform, although the Late Archean Campbellrand–Mal-mani subgroups containing carbonates with δ13Ccarb values close to

0‰ occur significantly below these units. Enhanced cyanobacterialproductivity has been inferred to explain 13C-enrichment beyondbackground levels during the Lomagundi Event in shallow-marinesettings (e.g. Melezhik et al., 1999) with the Bahama Banks and SharkBay used as modern analogues. High productivity on the BahamaBanks is maintained by red and green calcareous algae in shallow,restricted settings that are not analogous to an open and extensivecarbonate platform in the Pretoria Basin where stromatolites reflect-ing cyanobacterial activity are rare due to the deeper-water de-positional environment. Similarly, regardless of the validity of theevaporative model for the origin of extreme 13C-enrichments, there isno evidence for hypersaline conditions during the deposition of theSilverton or Sengoma Argillite formations in the Transvaal Basin.On this basis, we conclude that none of these models appears tosatisfactorily explain the 13C-enrichment in the carbonates of theseunits, and we suggest that δ13Ccarb values in the nearshore settingsmainly reflect the global, rather than local signals.

Carbon isotope values in carbonates of the Sengoma Argillite andSilverton formations are indeed similar to those previously reportedfrom these formations in other parts of the extensive Transvaal Basin(Master et al., 1993; Buick et al., 1998; Swart, 1999; Bekker et al., 2001)and in the sequence-stratigraphically correlative Lucknow Formation inthe GriqualandWest Basin (Coetzee et al., 2006; see Fig. 2 for location),and broadly correlative Paleoproterozoic carbonates deposited duringthe Lomagundi Eventworldwide (Bekker et al., 2006a).We donot implythat small-scale shelf gradients were entirely lacking in the Paleoprote-rozoic epeiric basins; ratherwe suggest that high atmospheric CO2 levelspredicted by climatic models likely caused them to be damped by thehigher seawater DIC content and higher exchange rates betweenatmospheric CO2 and seawater DIC (cf. Panchuk et al., 2005).

3.2. Δδ in shallow-marine carbonate environment: implication for anenhanced methane cycle

Δδ is an important parameter for characterizing carbon cyclingduring the 2.22–2.1 Ga Lomagundi δ13Ccarb excursion (Karhu andHolland, 1996). Whereas the δ13CDIC values are generally assumed tobe homogeneous in open-marine settings and faithfully recorded byδ13Ccarb values, the carbon isotope record of marine TOC can bestrongly influenced by local environmental conditions such as redoxstate, nutrient content, and aqueous CO2 levels, in addition to theseawater δ13CDIC value. It is therefore difficult to constrain an averageorganic carbon isotopic value on the global scale for any specific timeinterval. δ13Corg values or their stratigraphic trends in environmentswith high TOC contents could be strongly influenced by secondaryproductivity occurring in the anoxic water mass of the stratified watercolumn (e.g. methanotrophs or S-oxidizing bacteria; Hayes et al.,1999), especially in the Precambrian ocean where dissolved oxygenlevels were likely much lower (Slack et al., 2007). Intense biogeo-chemical carbon cycling in redox-stratified water columns thuscomplicates the interpretation of δ13Corg and Δδ values. Althoughorganic-rich sediments that were strongly influenced by secondaryproductivity are unlikely to provide genuine records of global sea-water δ13C values, sections with high TOC contents naturally attractbiogeochemical studies and, therefore, bias the secular record ofglobal average δ13Corg values. Shallow-water marine environmentsare better settings to look for faithful records of seawater δ13C trends,particularly for organic carbon isotopes, because secondary produc-tivity there is confined to anoxic porewaters below the water–sediment interface, and direct comparison with carbonate carbonisotope values is possible in carbonate successions. There is relativelylittle data on Δδ values from carbonates deposited during theLomagundi Event because most of these carbonate successions appearto have been deposited in shallow-marine and relatively well-oxygenated waters and are, thus, extremely lean in organic carbon(≪1 mg C/g sample) with δ13C values of refractory organic matter


strongly affected by nearly complete loss of organic carbon (Bekkeret al., 2006a). Carbonates of the Silverton Formation therefore providea rare opportunity to measure Δδ values in shallow-marine carbonateenvironments during the Lomagundi Event.

Assuming that δ13Ccarb values genuinely reflect seawater δ13CDICvalues and that δ13Corg values represent average carbon isotope valuesof organic matter, fractionation between these two reservoirs on theshallow-marine carbonate platform was close to 30‰. Until now, theLomagundi excursion has only been shown to be present in thecarbonate fraction. Since older and younger Late Archean and earlyPaleoproterozoic carbonate successions show similar Δδ values (e.g.Eichmann and Schidlowski, 1975; Schidlowski, 1988; Des Marais et al.,1992), the secular trend of δ13Corg values from shallow-marinecarbonate successions spanning the duration of the LomagundiEvent should record a positive δ13Corg excursion and, thus, couplingbetween the inorganic and organic carbon isotope records. If thisinterpretation is found to be correct on a global scale by further studies,then the hypothesis of an enhanced microbial degradation of organicmatter associated with methanogenesis below the sediment–waterinterface in shallow diagenetic settings, and a decoupling of organicand carbonate carbon records during the progressive build-up ofmolecular oxygen in the oceans and atmosphere (Hayes andWaldbauer, 2006) is not a valid explanation for the origin of highδ13C values of carbonates deposited during the Lomagundi Event. Byanalogywith Phanerozoic, and younger Precambrian examples, aΔδ of~30‰ suggests that the organic matter was mostly fixed by organismsutilizing a pathway of oxygenic photosynthesis in thewater column, orperhaps at the sediment–water interface. The 13C-enrichment in bothinorganic and organic carbon on the shallow-marine carbonateplatform during the Lomagundi Event, therefore, provides strongevidence that thewhole marine DIC reservoir was affected rather thanjust shallow porewaters below the sediment–water interface aspredicted in the model of Hayes and Waldbauer (2006).

3.3. Chemical and redox stratification in the Silverton Basin

Carbonates in the upper part of the Silverton–Sengoma Argilliteformations and the underlying black shales of the same units weredeposited during the 2.22–2.1 Ga Lomagundi Event based on theirgeochronologic constraints discussed in ‘Regional Geology andStratigraphy’ section and positive carbon isotope values of carbonatesin the lower and upper parts of the Silverton Formation bracketing thestudied shales (Buick et al., 1998; this paper). Walther's Law dictatesthat in offshore settings, shales, that are interlayered with carbonatesin the South African section (see Fig. 3), and conformable with anoverlying carbonate sequence, are deeper-water time equivalentsof shallow-water carbonates found higher in the succession. As dis-cussed above, shales from the Strat 2 drill core in Botswana contain upto 22% TOC and are more organic-rich than the contemporaneousshallow-marine carbonates of the Silverton and Sengoma Argilliteformations. The shales, therefore, suggest high organic carbon burialin offshore settings and, potentially, a strong biological pump to thedeeper-water settings (cf. Hotinski et al., 2004) during the LomagundiEvent. In contrast to the associated organically lean carbonates, theTOC in the offshore shales displays significantly lower δ13C values. Theapparent depth gradient in δ13Corg values between shallow-watercarbonates and deeper-water shales is about 6‰, much larger than thewater column gradient in the modern ocean (≤2‰; Kroopnick, 1985),even though the difference in depth of deposition for these carbonatesand shales in the epicontinental basin was unlikely more than 100 m.Shore-to-basin gradients of about 4‰ have been documented acrossseveral epeiric platforms during the Hirnantian carbon isotopeexcursion (e.g. Melchin and Holmden, 2005). Although evidence fora seawater shelf gradient in δ13CDIC has not yet been convincinglydemonstrated in Paleoproterozoic basins, we recognize that it is achallenging problem to amass the necessary evidence on account of

the lack of in situ carbonate production on the deeper shelf, and thepotential overprinting of deeper shelf TOC records by secondaryproductivity, as discussed above. Furthermore, Hotinski et al. (2004)concluded based on box modeling that significant water columngradients would be unlikely to develop in the high CO2 atmospherepredicted by Precambrian climatic models (cf. Kasting, 1993) asvigorous CO2 exchange between the atmosphere and ocean wouldlimit the impact of biological pump.

Therefore, accepting that TOC of the Sengoma Argillite Formationwith its low δ13C values does not reflect a global ocean water columngradient in δ13CDIC, two alternatives must be considered: i) either theDIC in offshore settings was locally affected by enhanced mineraliza-tion of organic matter, or ii) inputs from secondary productivitysignificantly contributed to the TOC of the offshore shales. The BlackSea is a good example for the first scenario in that it has a significantcarbon isotope water column gradient in the DIC from +0.8‰ near thesurface to −1.4‰ at 100 m depth and to about −6.3‰ at the bottom inthe deep sulfidic part of the basin (Fry et al., 1991; Schouten et al.,2004). Therefore, the carbon isotope value of the DIC reservoir can bestrongly affected in stratified basins where the exchange with theocean is restricted at depth by sills. In contrast, areas with high organicproductivity associated with upwelling zones on continental marginsare vertically well-mixed and at most 1 to 2‰ lower in δ13CDIC valuesthan the global surface ocean (e.g. Peru Upwelling Zone; Bidigareet al., 1997). Although mineralization of organic matter and methanecycling may have been important in the Pretoria basin, and bearing inmind the fact that sedimentary records of offshore δ13CDIC values arelacking or compromised, there are several lines of indirect evidencesuggesting that seawater δ13CDIC values in offshore settings can beapproximated by the sedimentary records of carbon in nearshoresettings. First, as discussed above, it seems rather odd that a large,open-marine, tidally-influenced epicontinental basin under highPaleoproterozoic atmospheric CO2 levels would develop long-termconditions above the wave base under which δ13CDIC values could beso significantly 12C-enriched offshore. Low δ13C values similar to thoseobserved in the thick black shale section in the drill core Strat 2 alsooccur in shales with lower TOC (mostly less than 1 wt.%) collectedfrom the outcrop sections in the more proximal parts of the basin(Watanabe et al., 1997; Coetzee, 2001). Second, if a carbon isotopelateral gradient occurred in the Pretoria basin, as observed now in theBlack Sea, one would expect to find carbon isotope values ofcarbonates to be highly variable and depth-controlled, with lowerδ13C values at the base of the upward-shallowing cycles immediatelyabove shales (see Fig. 3) and increasingly higher δ13C values at theirtops. In contrast, our data do not show significant variations in carbonisotope values of carbonates with sea level changes and only highlypositive carbon isotope values were found in this and previous studies(Buick et al., 1998; Swart, 1999). We therefore infer that δ13C values ofDIC in the Pretoria basin were not significantly affected on the basinscale by mineralization of organic matter in sediments with high TOCcontent and use the δ13C values of Silverton–Sengoma carbonates as aproxy for carbon isotope values of DIC in settings where Sengomashales accumulated.

It has been suggested that Δδ values larger than 32‰ in the LateNeoproterozoic interglacial time intervals might be related to theactivity of methane-oxidizing or sulfide-oxidizing bacteria incorpor-ating light carbon either from methane or mineralized organic matter(Hayes et al., 1999). Returning to the Black Sea as a potential modernanalogue for the Precambrian redox-stratified ocean, significantheterotrophic (secondary) and chemotrophic (e.g. sulfide-oxidizingand thiosulfate-disproportionating bacteria) production of organicmatter in the water column there contributes ~45% of the total carbonflux to the sediments; the remainder is from primary productivity inthe surface ocean above the redoxcline (Karl and Knauer, 1991). Bulkδ13Corg values for the organic carbon in the bottom sediments of theBlack Sea are less than 4‰ lower in δ13Corg value than phytoplankton


in the surface waters (Fry et al., 1991; Hayes et al., 1999). This dif-ference is consistent with the ~6‰ on the average lowering of δ13Corgvalues between nearshore and offshore settings in the Pretoria basin,suggesting that a significant portion of the organic matter in thedeeper-water shales originated from heterotrophic and chemotrophicmetabolism. This would also imply a stable redoxcline in this part ofthe basin (Fig. 7). The redoxcline with the pronounced underlyingsulfidic conditions was indeed inferred independently for the Pretoriabasin on the basis of Fe speciation and δ34S data (Shen and Bekker,2006; Bekker et al., 2006b; Scott et al., 2008). Sedimentary structuresin the black shale section in the drill core Strat 2 (see Fig. 6), such aswave ripples with internal laminations, climbing ripples, and soft-sediment deformation structures, however, suggest that the redox-cline developed in relatively shallow waters, above wave base, incontrast to the Black Sea analogue. It is plausible that deposition ofthese organic-rich shales reflects intense upwelling from the deepocean, bringing nutrients to fuel primary productivity in the basin. Thedramatic difference in carbon isotope fractionation between TOC andDIC in the Pretoria basin (~36‰ on the average) and in the BlackSea (~27‰ maximum) likely reflects a shallower redoxcline, reducedamounts of seawater sulfate and other oxidants and, consequently,extensive methanogenesis in the deeper anoxic part of the Pretoriabasin with methanotrophy at the redoxcline.

A related question, therefore, is whether other shallow-waterenvironments spanning the Lomagundi Event show evidence for apronounced redoxcline, sulfidic conditions, and significant 13C-depletion in the preserved organic matter. Evidence from thePaleoproterozoic Francevillian Series of Gabon suggests that thismay have been the case. Carbonates in the transgressive sequencewith δ13Ccarb values ranging from +2.6 to +6.3‰ are interlayered withorganic-rich shales with δ13Corg values as low as −46.2‰ (Gauthier-Lafaye and Weber, 1989; 2003). The overlying volcanics were recentlydated at 2083 ±6 Ma (Gauthier-Lafaye, 2006) establishing that theorganic-rich shales with low δ13Corg values were deposited during theLomagundi Event. Based on the speciation of Fe in the black shales(Scott et al., 2008), and basin analysis by Gauthier-Lafaye and Weber(2003), the intercalated shales were likely deposited under euxinicconditions in the epicontinental basin influenced, perhaps, byupwelling from the deep ocean. If these findings from two Paleo-proterozoic basins may be taken as evidence for extensive euxinia inshallow-water epicontinental seas during the Lomagundi Event, thequestion arises as to the redox state of the deep ocean during that

Fig. 7. Model for biogeochemical cycling of carbon in the shallow-water and deeper-water eSengoma Argillite formations (modified from Bekker and Kaufman, 2007). Note the sulfidic oepicontinental basin overlain by the oxygenated upper part of thewater column and underlaiof thewater columnwas dominated by primary productivity whereas secondary (heterotrophto TOC in the deeper-water environments (cf. Coetzee et al., 2006; Bekker and Kaufman, 20

time. Giant sedimentary Mn-deposits and associated minor ironformations deposited during the Lomagundi Event (e.g. Gauthier-Lafaye and Weber, 2003; Bekker et al., 2003b) strongly suggest thatthe deep ocean was neither oxic nor euxinic but rather suboxic toanoxic at that time.

Combining these observations, the ocean during the LomagundiEvent was likely redox-stratified, with biological activity coupledbetween oxic surface waters, euxinic intermediate waters, andsuboxic to anoxic bottom waters (Fig. 7). Contrary to the generallyheld view (e.g. Melezhik et al., 2005; Hayes andWaldbauer, 2006) thatthe Lomagundi Event was not accompanied by the deposition oforganic-rich shales consistent with a high relative burial rate oforganic carbon, several basins in West and South Africa and Brazilcontain organic-rich shales. Therefore, in spite of the fact that thesignature of the Lomagundi excursion might be overprinted bysecondary productivity in these deeper-water shales, their high TOCand low δ13Corg values suggest that significant amounts of organiccarbonwith light δ13Corg values were removed from the oceans duringthis time, helping to explain high δ13C values recorded in shallow-water carbonates.

Assuming that our data for the Pretoria basin are typical for globallyaveraged δ13Ccarb and δ13Corg values of carbonates and organic matter,respectively, from the shallow-water environments, shales of theSengoma Argillite Formation deposited in the deeper-water part of thebasin provide some indication of the decrease in δ13C values of organiccarbon below the redoxcline during the Lomagundi Event. It is thereforeimpossible to estimate a global average fractionation between organicand inorganic carbon during the Lomagundi Event from our dataset.However, because primary producers dominate the ocean organiccarbon budget now (Hedges and Keil, 1995) and were likely similarlyimportant during the Lomagundi Event and assuming that shallow-water environments are more likely to reflect primary productivity, wecan bracket this parameter between 30 and 45‰; likely closer to 30‰.

3.4. Size of DIC and DOC reservoirs during the Lomagundi Event

Our data provide a new perspective on carbon cycling during theLomagundi Event and allow for a test of the possible effects of largeDIC and DOC reservoirs on the operation of carbon cycle. It is generallyaccepted that except for relatively short but important time intervalsin the Late Archean, Paleoproterozoic, and Neoproterozoic, the relativeproportion of carbonate and organic matter burial was stable in the

nvironments of the Pretoria Basin, South Africa during deposition of the Silverton andxygen-minimum zone impinging on the shelf in the shallow-marine tidally-influencedn by the anoxic or suboxic deeper-waters. Organic matter in the shallow oxygenated partic) and chemotrophic productivity at and below the redoxcline contributed significantly07).

Fig. 8. a–c: Schematically shown fields for the Cenozoic carbon cycle and two end-members based on the carbon cycle dominated by either large DOC or large DIC reservoirs on thescatter diagrams of δ13Ccarb vs. δ13Corg (a); δ13Ccarb vs. Δδ (b); and δ13Corg vs. Δδ (c). Note different slopes and intercepts predicted by these states of the carbon cycle. d–f: Scatterdiagrams for carbon isotope data of the Silverton and Sengoma Argillite formations. Note lack of correlation between δ13Ccarb and δ13Corg (d) and between δ13Ccarb vs. Δδ (e) butsignificant correlation between δ13Corg and Δδ (f).


Precambrian, and similar to that observed today based on the carbonisotope records of carbonates and organic matter (e.g. Schidlowski,1988; Des Marais et al., 1992). A slightly different treatment of thecarbon cycle involving ocean floor carbonatization, suggests that therelative burial rate of organic matter could have gradually increasedover time if submarine hydrothermal weathering of the ocean crust

was more important on the early Earth. Nevertheless, even in thismodel the range of this gradual change is rather limited (Bjerrum andCanfield, 2004). While the endogenic carbon flux to surface environ-ments, and the ratio of organic carbon to inorganic carbon burialfluxes did not change significantly since the Late Archean, DOC andDIC reservoirs could have been larger in the past due to weaker solar


luminosity and greenhouse requirements, and less efficient burial oforganic carbon (Grotzinger and Kasting, 1993; Kasting, 1993; Rothmanet al., 2003; Bartley and Kah, 2004).

Under the present conditions, the residence time of DIC is on theorder of 105 yr whereas that of DOC is at least several orders ofmagnitude shorter (Holland,1984).As a result, the smallerDOC reservoirresponds faster to environmental perturbations such as changes in pCO2or in the ocean redox state, whereas the larger DIC reservoir respondsmuchmore slowly. This relationship is illustrated schematically in Fig. 8ausing a Cenozoic record that reflects a carbon cycle evolving qua-sistatically in a succession of steady states (Rothman et al., 2003). Notethat for the last ca. 50 Ma average carbon isotope values of TOCconsistently increased whereas those of carbonates were steady ordecreasing. This results in the overall decrease in Δδ values through theCenozoic (see Fig. 3 in Hayes et al., 1999). Laboratory experiments withphytoplankton show (see Hayes et al., 1999 for review) that as pCO2increases δ13C values of photosynthetically-produced organic matterbecomemore negative due to an enhanced biological fractionation. Theincreased isotopic discrimination between fixed organic carbon and DICcauses a corresponding positive shift in carbon isotope values of theseawater DIC pool that is reflected in higher δ13Ccarb values. Based onexperiments with phytoplankton cultures, this relationship conflictswith the conventional treatment of the ocean carbon isotope massbalance (e.g. DesMarais et al.,1992; Schidlowski and Aharon,1992), thatassumesΔδ valueswere invariant over geological time. This assumptionallows us to directly relate secular variations in δ13Ccarb values tochanges in the relative burial rate of organic carbon over time. The linearrelationship betweenΔδ and δ13Ccarb valueswas recently emphasized inthe model of Rothman et al.'s (2003) for the Neoproterozoic carboncycling, based on large variations in the carbon isotope fractionationbetween organic and inorganic carbon since the Neoproterozoic (Hayeset al., 1999). The latter model assumes that carbon isotope variations incarbonates reflect changes inΔδ values driven by environmental factorssuch as pCO2 or ocean redox changes rather than reflecting relativeburial rate of organic carbon since the former changes faster than thelatter. Considering that the size and residence times of DIC and DOCreservoirs could have been different in the Precambrian than today, it isuseful to discuss two end-members with the surface ocean carbon cycledominated by either large DIC or large DOC reservoirs (Fig. 8a–c). If thesize and residence time of these reservoirs are dramatically different,one of themwill be relatively inert to perturbations that would affect itscarbon isotope value on a time scale shorter than its residence time,whereas the other with a significantly shorter residence time willrespond more quickly. Over time scales that are much longer than theresidence times of both reservoirs, their isotopic values will follow acurved path as has been shown for Phanerozoic seawater C and Sreservoirs (Kump and Garrels, 1986). These different scenarios aredepicted schematically in Fig. 8a–c showing C isotope responses to aperturbation in the carbon cycle whose duration is the intermediatebetween the DIC and DOC residence times. Using a similar approach,Rothman et al. (2003) interpreted δ13Corg and δ13Ccarb values recon-structed from the Neoproterozoic carbonate successions recordingpositive carbon isotope excursions as reflecting a dynamic systemwith a large DOC reservoir. In this case, carbon isotope values of car-bonates are highly sensitive to external forcing whereas carbon isotopevaluesof organicmatter are less so (see Fig. 8a–c). The oppositewouldbetrue in case of a large DIC reservoir. Note that the Cenozoic carbonisotope data reflect an intermediate state of the carbon cycle betweenthese two end-member models.

The Silverton–Sengoma carbonate data (Fig. 8d–f) may be eval-uated against these models. First, we agree with Rothman et al. (2003)and Bekker et al. (2003b) that the conventional steady-state model,and in particular its reliance on organic carbon burial as a way tochange seawater δ13C values, cannot explain the extremely highδ13Ccarb values recorded by the Neoproterozoic and Lomagundi Events.Similarly to the Neoproterozoic carbon cycle, the carbon isotope

background for the Lomagundi excursion is +6 to +8‰, but highercarbon isotope values are not uncommon and even +28‰ values wereobserved in a relatively thick carbonate section deposited in an open-marine setting (Bekker et al., 2003b). The Silverton–Sengoma data ploton, or slightly above, the best-fit line on a plot of δ13Ccarb vs. Δδ forNeoproterozoic carbonate successions recording positive carbonisotope excursions and, arguably, a dynamic state of carbon cycle(Rothman et al., 2003). The Silverton–Sengoma data (Fig. 8d–f) showno correlation between δ13Ccarb and either Δδ or δ13Corg, but there is asignificant correlation between δ13Corg and Δδ. The trend is consistentwith the carbon cycle dominated by a large DIC reservoir during theLomagundi Event. Carbonates are indeed abundant in the Paleoproter-ozoic sedimentary successions deposited during the Lomagundicarbon isotope excursion, and the long duration of the LomagundiEvent (N100 Ma) is consistent with the expected stability of carbonisotope values in carbonates at the time when the DIC reservoir waslarge. The DOC reservoir may have been larger than now as well, butdue to a much larger size of the DIC reservoir, its effect on the carboncycle is not apparent. How could the carbon cycle remain in a dynamicstate regardless of how large the DIC reservoir was for the N100 Maduration of the Lomagundi Event? We speculate that extensive intra-cratonic basins and epeiric seas,with restricted connections to a redox-stratified global ocean, developed during the protracted breakup of aPaleoproterozoic supercontinent (Bekker et al., 2006a). The paleogeo-graphyof the period, similar to that in theNeoproterozoic, favored highorganic carbon burial rates, and this in turn caused a global-scale shiftin the carbon cycle to higher δ13C values. Numerical simulations ofcarbon cycling study the effect of a flux perturbation by applying aninstantaneous change from which the system recovers from, or apersistent forcing that drives the system into a new steady state (e.g.Kump and Garrels, 1986; Kump and Arthur, 1999). It is suggested herethat under conditions of high organic carbon burial, carbon cyclingoperates in a dynamic rather than steady state due to perturbationsoccurring on a time scale intermediate between the residence times ofDIC andDOC reservoirs. The carbon cycle could havebeen shifting fromone non-steady dynamic state to another, never reaching a steady stateuntil the breakup of the Paleoproterozoic supercontinent between 2.1and 2.0 Ga that changed the locus of organic matter deposition tocontinental margins (Bekker et al., 2003b).

In conclusion, ourdata support theexistenceof a largeDIC reservoir inthe oceans during the Lomagundi Event. The DIC reservoir likelydecreased gradually as solar luminosity increased over time and,consequently, the carbon cycle shifted to one that was dominated by alarge DOC reservoir sometime before the Late Neoproterozoic. At the endof the Precambrian, the DOC reservoir was drastically reduced as deepocean oxygen levels increased (Logan et al., 1995; Rothman et al., 2003;Canfield et al., 2006; Fike et al., 2006)marking the transition toamodern-style ocean carbon cycle. The DOC reservoir likely increased episodicallyin the Phanerozoic at the timeswhen the oceanwas stratified and anoxic.Shelf gradients likely waxed and waned under Phanerozic relatively lowpCO2 levels as large epeiric seas developed in association with thetectonically-driven sea level changes (e.g. Melchin and Holmden, 2006).

3.5. Implications ofΔδ as a diagenetic filter and of δ13Corg for reconstructionof seawater composition

A number of workers have relied on carbon isotope values of TOCas a tool for evaluating the degree of post-depositional carbonatealteration and to infer seawater δ13CDIC values in Precambriansiliciclastic successions where carbonates are absent (e.g. Knoll et al.,1986; Kaufman and Knoll, 1995; Calver, 2000; Karlstrom et al., 2000;Walter et al., 2000; Kaufman and Xiao, 2003). While this approach hasbeen generally successfully applied to Phanerozoic carbonate andsiliciclastic successions, the results of this study suggest that it canlead to erroneous conclusions when applied to deeper-water depositsin the Precambrian, where the primary photosynthetic signature is


susceptible to being overprinted by heterotrophic and chemolitho-trophic productivity in redox-stratified oceans.

Our study suggests that carbon isotope values of TOC during theLomagundi Event were strongly influenced by environmental factorschanging water column redox conditions and relative contribution ofsecondary productivity to TOC. As a result, changes in sea level andocean circulation affecting position of the depositional site withrespect to the redoxcline in the water column and upwelling currentsmay have influenced carbon isotope values of TOC in deeper-watersettings to such an extent that they no longer behave as reliableindicators of DIC isotope trends in seawater. Since the oceanwas likelyredox-stratified during most of the Precambrian, these environmentalcontrols might be equally applicable to other Precambrian succes-sions. This situation is rather different from that in the modernoxidized ocean where δ13C values of TOC are relatively homogeneousand are generally coupled with δ13CDIC values.

4. Conclusions

1. During the Paleoproterozoic Lomagundi carbon isotope excursion,fractionation between carbonate and organic carbon on theshallow-marine carbonate platform was close to 30‰, but indeeper-water suboxic to anoxic and, locally, sulfidic settings it waslarger, consistent with redox stratification in the ocean;

2. There is increasing evidence for deposition of organic-rich shalesduring the Lomagundi excursion in deeper-water settings;

3. The largemagnitudeof the Lomagundi carbon isotopeexcursion likelyreflects the large isotope fractionation between bulk organic carbonand carbonate carbon, due to contributions from secondary produc-tivity in addition to high relative burial rates of organic carbon. It isdifficult at present to separate these two variables on the global scale;

4. Carbonate and organic carbon isotope data suggest that the carboncycle was in a dynamic state dominated by a large DIC reservoirduring the Lomagundi Event, however the DOC reservoir couldhave also been larger than at present, but its influencewas dwarfedby a much larger DIC reservoir;

5. In redox-stratified Precambrian oceans, carbon isotope values ofTOC in deeper-water facies were likely strongly influenced bycontributions from heterotrophic and chemotrophic productivityand, therefore, estimates of post-depositional carbonate alterationand carbon isotope values of DIC made on basis of carbon isotopevalues of organic carbon should be cautioned.

Acknowledgements

This study was supported by the NASA-Ames Research Center,NASA, and PRF/ACS grants to H.D. Holland and, at a later stage, by NSFgrant EAR-05-45484, NASA Astrobiology Institute award No.NNA04CC09A, and NSERC Discovery Grant to AB. Read Mapeo(University of Botswana) and El-El Coetzee (University of Johannes-burg) guided in the field, Geological Survey of Botswana providedaccess to drill core material and permitted sampling, Marilyn Fogelprovided access to mass-spectrometers and Seth Newsome patientlyhelped with analyses at the Geophysical Laboratory. Lee Kump and ananonymous reviewer provided valuable comments.

Appendix A. Supplementary data

Supplementary data associated with this article can be found, inthe online version, at doi:10.1016/j.epsl.2008.04.021.

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Fractionation between inorganic and organic carbon during the Lomagundi (2.22–2.1 Ga) carbon is.....IntroductionRegional geology and stratigraphy

MethodsGeochemical data and preservation of primary δ13C values

DiscussionDo high δ13Ccarb values of the Silverton–Sengoma carbonates reflect local or global conditions?Δδ in shallow-marine carbonate environment: implication for an enhanced methane cycleChemical and redox stratification in the Silverton BasinSize of DIC and DOC reservoirs during the Lomagundi EventImplications of Δδ as a diagenetic filter and of δ13Corg for reconstruction of seawater composi.....

ConclusionsAcknowledgementsSupplementary dataReferences

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