The rise of oxygen and siderite oxidation during the Lomagundi Event Aviv Bachan 1 and Lee R. Kump Department of Geosciences, The Pennsylvania State University, University Park, PA 16802 Edited by Mark H. Thiemens, University of California, San Diego, La Jolla, CA, and approved April 14, 2015 (received for review November 21, 2014) The Paleoproterozoic Lomagundi Event is an interval of 130–250 million years, ca. 2.3–2.1 billion years ago, in which extraordinarily 13 C enriched (>10‰) limestones and dolostones occur globally. The high levels of organic carbon burial implied by the positive δ 13 C values suggest the production of vast quantities of O 2 as well as an alkalinity imbalance demanding extremely low levels of weathering. The oxidation of sulfides has been proposed as a mechanism capable of ameliorating these imbalances: It is a potent sink for O 2 as well as a source of acidity. However, sulfide oxida- tion consumes more O 2 than it can supply CO 2 , leading to insur- mountable imbalances in both carbon and oxygen. In contrast, the oxidation of siderite (FeCO 3 proper, as well as other Fe 2+ -bearing carbonate minerals), produces 4 times more CO 2 than it consumes O 2 and is a common—although often overlooked—constituent of Archean and Early Proterozoic sedimentary successions. Here we propose that following the initial rise of O 2 in the atmosphere, oxidation of siderite provided the necessary carbon for the contin- ued oxidation of sulfides, burial of organic carbon, and, most im- portantly, accumulation of free O 2 . The duration and magnitude of the Lomagundi Event were determined by the size of the preex- isting Archean siderite reservoir, which was consumed through oxidative weathering. Our proposal helps resolve a long-standing conundrum and advances our understanding of the geologic his- tory of atmospheric O 2 . carbon isotopes | oxygen | siderite | carbon cycle | Great Oxidation Event R econstructing the geologic history of atmospheric oxygen is among the foremost scientific challenges of our time (1). The level of atmospheric oxygen (pO 2 ) without doubt played a key role in the evolution of the Earth System (2), exerting a major influence on the biosphere, especially the evolution of metazoans (3). With no direct way of measuring oxygen concentrations in deep geologic time, the stable isotopes of carbon recorded in marine limestones provide key constraints (4). Carbon enters the ocean−atmosphere system through volcanoes and weathering of carbon-bearing sedimentary rocks and can exit in one of two ways: (i ) uptake during photosynthesis and burial of organic carbon leading to O 2 production and (ii ) reaction during weath- ering and formation of CaCO 3 in the ocean. The carbon isotopic record tells us how carbon was partitioned between these two sinks: A δ 13 C value of 0‰ indicates that ∼80% of incoming car- bon was buried as carbonate carbon and 20% as organic carbon. Positive excursions in δ 13 C are unusual and indicate that a larger fraction of carbon was fixed and buried as organic carbon and, with it, a larger amount of O 2 was produced. Following the indications for the first rise of O 2 in the atmo- sphere (5, 6) is the largest and most protracted period of 13 C enrichment in the geologic record, known as the Lomagundi Event (Fig. 1). Limestones and dolostones with extreme carbon isotopic values of + 8‰ to greater than + 15‰ occur globally (6–8), and a duration of between 128 million years (m.y.) and 249 m.y. is suggested by current age constraints (9). The highly ele- vated δ 13 C values indicate the burial of tremendous amounts of organic carbon, and the production of correspondingly vast amounts of O 2 . In fact, the duration and magnitude of the iso- topic excursion in δ 13 C bespeak of O 2 fluxes so large that they challenge our understanding of geochemical cycles. Calculations indicate an integrated production of far larger amounts of O 2 than currently exist—or likely ever existed—in Earth’s atmosphere, implying the concurrent existence of effective O 2 sinks (10). A second, hitherto unrecognized problem exists as well. The elevated δ 13 C values indicate a repartitioning of the incoming carbon in favor of organic carbon burial. However, if the total amount of carbon entering the ocean−atmosphere system remains unchanged, then any increase in the organic carbon burial flux can only be at the expense of the other output flux, that of carbonate carbon. However, the burial of carbonate carbon represents the burial not only of carbon but also of alkalinity, and thus a decrease in its magnitude demands a commensurate decrease in the input of alkalinity from weathering. Critically, δ 13 C values of +10‰ indicate the burial of so large a fraction of organic carbon that a 90% reduction in carbonate burial, and hence weathering, would have been necessary to balance it (see SI Appendix for cal- culation). Assuming that weathering is proportional to pCO 2 to the 0.3 power (11, 12), a 90% reduction in weathering would have entailed a decline from a pCO 2 baseline of 10,000 ppm to single part per million levels. Consequently, in the face of such high levels of organic carbon burial, without an additional source of CO 2 or sink for alkalinity, a near-complete shutdown of weather- ing would have been required to balance the inputs and outputs of carbon. The more plausible alternative is that during Lomagundi times processes that consume O 2 and release CO 2 compensated the inferred imbalances such that pCO 2 levels were bolstered and pO 2 levels moderated. The oxidation of sedimentary sulfides is an attractive option for alleviating the attendant imbalances, as it is a potent sink for oxygen (13) and, in conjunction with acidi- fication of carbonates, a source of CO 2 (14). The oxidation of sulfides following the first rise of oxygen is supported by evidence Significance The evolution of Earth’s oxygen-rich atmosphere occurred in two major steps, the first of which took place approximately 2.4 billion years ago. Following the initial rise of oxygen, car- bon isotope evidence indicates the burial of vast quantities of organic carbon and the production of correspondingly large amounts of oxygen. However, if not accompanied by an ad- ditional supply of carbon, the extreme levels of organic carbon burial imply nonphysically low atmospheric pCO 2 levels. Here we propose that the initial rise in O 2 led to the oxidation of a large preexisting reservoir of siderite (FeCO 3 ), which provided the necessary carbon for the burial of organic matter, pro- duction of further O 2 , and substantial accumulation of oxygen in Earth’s atmosphere for the first time. Author contributions: A.B. and L.R.K. designed research, performed research, and wrote the paper. The authors declare no conflict of interest. This article is a PNAS Direct Submission. 1 To whom correspondence should be addressed. Email: [email protected]. This article contains supporting information online at www.pnas.org/lookup/suppl/doi:10. 1073/pnas.1422319112/-/DCSupplemental. www.pnas.org/cgi/doi/10.1073/pnas.1422319112 PNAS Early Edition | 1 of 6 EARTH, ATMOSPHERIC, AND PLANETARY SCIENCES
Transcript
The rise of oxygen and siderite oxidation during theLomagundi EventAviv Bachan1 and Lee R. Kump
Department of Geosciences, The Pennsylvania State University, University Park, PA 16802
Edited by Mark H. Thiemens, University of California, San Diego, La Jolla, CA, and approved April 14, 2015 (received for review November 21, 2014)
The Paleoproterozoic Lomagundi Event is an interval of 130–250million years, ca. 2.3–2.1 billion years ago, in which extraordinarily13C enriched (>10‰) limestones and dolostones occur globally.The high levels of organic carbon burial implied by the positiveδ13C values suggest the production of vast quantities of O2 aswell as an alkalinity imbalance demanding extremely low levelsof weathering. The oxidation of sulfides has been proposed as amechanism capable of ameliorating these imbalances: It is a potentsink for O2 as well as a source of acidity. However, sulfide oxida-tion consumes more O2 than it can supply CO2, leading to insur-mountable imbalances in both carbon and oxygen. In contrast, theoxidation of siderite (FeCO3 proper, as well as other Fe2+-bearingcarbonate minerals), produces 4 times more CO2 than it consumesO2 and is a common—although often overlooked—constituent ofArchean and Early Proterozoic sedimentary successions. Here wepropose that following the initial rise of O2 in the atmosphere,oxidation of siderite provided the necessary carbon for the contin-ued oxidation of sulfides, burial of organic carbon, and, most im-portantly, accumulation of free O2. The duration and magnitude ofthe Lomagundi Event were determined by the size of the preex-isting Archean siderite reservoir, which was consumed throughoxidative weathering. Our proposal helps resolve a long-standingconundrum and advances our understanding of the geologic his-tory of atmospheric O2.
Reconstructing the geologic history of atmospheric oxygen isamong the foremost scientific challenges of our time (1). The
level of atmospheric oxygen (pO2) without doubt played a keyrole in the evolution of the Earth System (2), exerting a majorinfluence on the biosphere, especially the evolution of metazoans(3). With no direct way of measuring oxygen concentrations indeep geologic time, the stable isotopes of carbon recorded inmarine limestones provide key constraints (4). Carbon enters theocean−atmosphere system through volcanoes and weathering ofcarbon-bearing sedimentary rocks and can exit in one of twoways: (i) uptake during photosynthesis and burial of organiccarbon leading to O2 production and (ii) reaction during weath-ering and formation of CaCO3 in the ocean. The carbon isotopicrecord tells us how carbon was partitioned between these twosinks: A δ13C value of 0‰ indicates that ∼80% of incoming car-bon was buried as carbonate carbon and 20% as organic carbon.Positive excursions in δ13C are unusual and indicate that a largerfraction of carbon was fixed and buried as organic carbon and, withit, a larger amount of O2 was produced.Following the indications for the first rise of O2 in the atmo-
sphere (5, 6) is the largest and most protracted period of 13Cenrichment in the geologic record, known as the LomagundiEvent (Fig. 1). Limestones and dolostones with extreme carbonisotopic values of + 8‰ to greater than + 15‰ occur globally(6–8), and a duration of between 128 million years (m.y.) and249 m.y. is suggested by current age constraints (9). The highly ele-vated δ13C values indicate the burial of tremendous amountsof organic carbon, and the production of correspondingly vastamounts of O2. In fact, the duration and magnitude of the iso-topic excursion in δ13C bespeak of O2 fluxes so large that they
challenge our understanding of geochemical cycles. Calculationsindicate an integrated production of far larger amounts of O2than currently exist—or likely ever existed—in Earth’s atmosphere,implying the concurrent existence of effective O2 sinks (10).A second, hitherto unrecognized problem exists as well. The
elevated δ13C values indicate a repartitioning of the incomingcarbon in favor of organic carbon burial. However, if the totalamount of carbon entering the ocean−atmosphere system remainsunchanged, then any increase in the organic carbon burial flux canonly be at the expense of the other output flux, that of carbonatecarbon. However, the burial of carbonate carbon represents theburial not only of carbon but also of alkalinity, and thus a decreasein its magnitude demands a commensurate decrease in the input ofalkalinity from weathering. Critically, δ13C values of +10‰ indicatethe burial of so large a fraction of organic carbon that a 90%reduction in carbonate burial, and hence weathering, wouldhave been necessary to balance it (see SI Appendix for cal-culation). Assuming that weathering is proportional to pCO2to the 0.3 power (11, 12), a 90% reduction in weathering wouldhave entailed a decline from a pCO2 baseline of 10,000 ppm tosingle part per million levels. Consequently, in the face of such highlevels of organic carbon burial, without an additional source of CO2or sink for alkalinity, a near-complete shutdown of weather-ing would have been required to balance the inputs and outputsof carbon.The more plausible alternative is that during Lomagundi times
processes that consume O2 and release CO2 compensated theinferred imbalances such that pCO2 levels were bolstered andpO2 levels moderated. The oxidation of sedimentary sulfides isan attractive option for alleviating the attendant imbalances, as itis a potent sink for oxygen (13) and, in conjunction with acidi-fication of carbonates, a source of CO2 (14). The oxidation ofsulfides following the first rise of oxygen is supported by evidence
Significance
The evolution of Earth’s oxygen-rich atmosphere occurred intwo major steps, the first of which took place approximately2.4 billion years ago. Following the initial rise of oxygen, car-bon isotope evidence indicates the burial of vast quantities oforganic carbon and the production of correspondingly largeamounts of oxygen. However, if not accompanied by an ad-ditional supply of carbon, the extreme levels of organic carbonburial imply nonphysically low atmospheric pCO2 levels. Herewe propose that the initial rise in O2 led to the oxidation of alarge preexisting reservoir of siderite (FeCO3), which providedthe necessary carbon for the burial of organic matter, pro-duction of further O2, and substantial accumulation of oxygenin Earth’s atmosphere for the first time.
Author contributions: A.B. and L.R.K. designed research, performed research, and wrotethe paper.
The authors declare no conflict of interest.
This article is a PNAS Direct Submission.1To whom correspondence should be addressed. Email: [email protected].
This article contains supporting information online at www.pnas.org/lookup/suppl/doi:10.1073/pnas.1422319112/-/DCSupplemental.
www.pnas.org/cgi/doi/10.1073/pnas.1422319112 PNAS Early Edition | 1 of 6
including the disappearance of detrital pyrites (15, 16), the ap-pearance of sedimentary evaporites (17), and Cr enrichment iniron-rich sedimentary rocks indicating a highly acidic weath-ering regime (18).Nonetheless, sulfide oxidation alone could not have fully com-
pensated the imbalances resulting from the elevated burial of or-ganic carbon during the Lomagundi Event. Close examination ofEqs. 1–3 reveals that pyrite oxidation coupled to acidification ofcarbonates leads to unavoidable imbalances in carbon and oxygendue to the stoichiometry of the overall reaction, which consumesmore oxygen than it releases carbon. The oxidation of 4 mol ofpyrite requires 15 mol of O2 (Eq. 1), while the associated acid-ification of carbonates can release at most only 8 mol of CO2(Eq. 2). Thus, on the whole, pyrite oxidation consumes 15 mol of O2but produces only 8 mol of CO2 (Eq. 3).
However, the continued oxidation of pyrite requires the burial of15 mol of CO2 as organic carbon (Eq. 4)—7 mol more than canbe supplied by pyrite oxidation coupled to the acidification ofcarbonates.
15CO2 + 15H2O→ 15CH2O+ 15O2 . [4]
Conversely, if the amount of CO2 required by organic carbonburial is assumed to balance the amount that can be supplied bypyrite oxidation coupled to acidification of carbonates (8 mol),not enough oxygen is produced during organic carbon burial(8 mol) to balance the demand of pyrite oxidation (15 mol), suchthat any pyrite oxidation would grind to a halt. Consequently, theoxidation of sulfides cannot be a sustained source of carbon overgeological timescales, even during intervals of highly elevatedoxygen production, and much less so during periods when thisis not the case (14).Here we suggest that siderite oxidation (including siderite proper,
FeCO3, as well as other Fe2+ bearing carbonate minerals) providedthe necessary CO2 during Lomagundi times. Siderite is a majorconstituent of Archean and Early Proterozoic sediments: It is
extremely abundant in banded iron formations, often even moreso than iron oxides (19). Siderite is also found in anomalouslyhigh concentrations in Proterozoic limestones and dolomites(20), where it arises from the replacement of Ca2+ and Mg2+ byFe2+ in the carbonate mineral lattice. Crucially, the oxidation ofsiderite produces 4 times more CO2 than it requires O2. Theoxidation of siderite (Eq. 5) followed by photosynthetic CO2fixation (Eq. 6) is a net source of oxygen (Eq. 7),
4FeCO3 + O2 → 2Fe2O3 + 4CO2 [5]
4CO2 + 4H2O= 4CH2O+ 4O2 [6]
4FeCO3 + 4H2O= 2Fe2O3 + 4CH2O+ 3O2 . [7]
Hence, in principle, the burial of the CO2 evolved from sideriteoxidation as organic carbon can produce oxygen 3 times in excessof the O2 required by siderite oxidation, with the surplus going tothe oxidation of sulfide, oxidation of reduced crustal iron, andthe accumulation of free O2 (21).
Calculations and Numerical ResultsSiderite oxidation would have contributed to transient 13C en-richment of the exogenic reservoir in two principle ways. First, bythe delivery of 13C enriched carbon. Massively bedded Archeanand Proterozoic marine siderites have an average carbon isotopiccomposition of −0.9‰ (19), ∼4‰ more enriched than the av-erage weathering input. Second, siderite oxidation could drive upexogenic δ13C by changing the ratio of organic to carbonate carbonburial fluxes as governed by the stochiometries of siderite, sulfide,and iron silicate oxidation. Consider first the oxidation of sideritecoupled to the burial of organic carbon and oxidation of ironsilicates (Eq. 8):
FeCO3 + 3FeSiO3 +H2O→ 2Fe2O3 + 3SiO2 +CH2O. [8]
Neither O2 nor CO2 appears in the above reaction; it is balancedfor both. The burial of 1 mol of organic carbon consumes 1 molof CO2, which is supplied by the oxidation of 1 mol of siderite.One quarter of the resulting mol of O2 that is produced goes tothe oxidation of siderite, while the other 3/4 mol goes to theoxidation of iron silicates. Siderite oxidation coupled to iron sili-cate oxidation and organic carbon burial can thus drive an increasein the δ13C of the exogenic pool with no imbalances in oxygenor carbon.Consider next the effects of sulfide oxidation (Eq. 3) on the
carbon cycle: The acidity produced by sulfide oxidation is neu-tralized by the release of calcium from limestones and silicates,which releases carbon. We need not distinguish between the directacidification of limestones by sulfuric acid (Eq. 2) and the sulfuricacid weathering of silicates; both reactions equally lead to net re-lease of carbon dioxide: the first by the conversion of carbonate toCO2, and the second by replacing carbonic acid weathering withsulfuric acid weathering and thus allowing volcanic CO2 to go un-consumed. Equally, viewed from the product side, the production ofcalcium sulfate makes available carbon that would otherwise be tiedto the burial of calcium carbonate. Coupled to siderite oxidation, aCO2 and O2 neutral reaction can be written (Eq. 9):
3FeS2 + 6CaCO3 + 7FeCO3 + 13H2O→ 5Fe2O3 + 6CaSO4
+ 13CH2O. [9]
We can use Eqs. 8 and 9 to construct a carbon isotopic mass bal-ance. We assume that some fraction, α, of the siderite is oxidizedaccording to Eq. 8, and the rest, ð1− αÞ, according to Eq. 9. Fromthe resulting isotopic mass balance, an expression can be obtained
Age (Ga)1.52.02.53.0
33S (‰
)0
4
8
12
16
–4
13C
(‰)
–12
–4
4
12
16
8
0
–8
Fig. 1. Archean and Proterozoic carbon and sulfur isotopic data. The Loma-gundi Event refers to the interval of highly positive δ13C values (black band)between 2.3 Ga and 2.0 Ga (9). The preceding collapse in the range of Δ33Svalues (in red and gray) indicates the increase in atmospheric O2 levels fromvanishing Archean levels for the first time (5). Adapted with permission fromMacmillan Publishers Ltd: Nature ref. 35.
2 of 6 | www.pnas.org/cgi/doi/10.1073/pnas.1422319112 Bachan and Kump
for the isotopic composition of limestones as a function of thesiderite oxidation flux (see SI Appendix):
δbcarb½‰=Fwsid
δwsid + e α′
Fwin + α′ Fw
sid[10]
where α′= 137 −
67 α. When the siderite oxidation flux is zero, Eq.
10 gives the long-term isotopic composition of marine carbon-ates of 0‰. As the net flux of siderite oxidation increases, sodoes the isotopic value of marine carbonates (Fig. 2). For themiddle range of δ13C values observed in carbonates belonging tothe Lomagundi Event ð+8‰< δ13C< +10‰Þ, a siderite oxida-tion flux between 13 Tmol/y (1012 mol/y) and 35.5 Tmol/y isrequired (dotted lines in Fig. 2).To further evaluate the time-dependent behaviors of CO2 and
O2, as well as place bounds on the quantities of reactants in-volved in the Lomagundi Event, we use a simple model of thesedimentary and oceanic carbon cycles. The model includes thesedimentary reservoirs of organic and carbonate carbon, sulfateand sulfide sulfur, siderite, and reduced iron. It also includes theoceanic reservoirs for calcium, sulfate, phosphate, and alkalinity,and the ocean−atmosphere reservoirs for carbon and oxygen. Thecarbonate system parameters (pH, pCO2, and the carbonate satu-ration, Ω) are calculated at every time step. Full model descriptionand code are available in SI Appendix. We set the initial modelconditions to postulated pre-Lomagundi conditions: low sulfate(50 μM), low O2 (10
–5 present atmospheric level), and high pCO2(10,500 ppm).We first use the model to explore the impacts of organic
carbon burial without accompanying oxidation of siderite. Wethen add the oxidation of siderite and iron silicates and evaluatethe resultant changes in pCO2 and pO2. Forcing organic carbonburial for 130 m.y. so that a +10‰ positive δ13C excursion arisesresults in a dramatic drop in pCO2 (Fig. 3, solid black line).Without additional CO2 input from siderite oxidation, organiccarbon burial consumes more carbon than can be supplied bysulfide oxidation forcing a precipitous decline in pCO2 to46 ppm. Following the cessation of this forcing, as organic carbonburial ceases to consume carbon, and sulfide oxidation begins to
consume oxygen at a rapid pace, pCO2 rises to 20 × 103 ppm (2×model baseline). The extremely low minimum pCO2 levels thatresult in this model run suggest that this scenario is implausible.Conversely, with increasing input of carbon from siderite oxi-dation, minimum pCO2 levels increase (Fig. 3). With a lowamount of FeCO3 oxidized (813 Emol, 1018 mol), pCO2 is re-duced from the initial 10,800 ppm to 2,800 ppm (red dash-dottedline in Fig. 3). Oxidizing 1,118 Emol of FeCO3 results in a moremoderate pCO2 decline to 7,600 ppm (green dashed line in Fig.3). Increasing the amount of siderite to 1,424 Emol of FeCO3results in pCO2 not being reduced by any significant amount(solid blue line in Fig. 3). In all cases, the modeled pO2 risessubstantially in association with the positive δ13C excursion. Theminimum modeled pO2 is 0.14 atm, or approximately two thirdsthe modern value, lending support to previous suggestions for sub-stantial O2 accumulation in association with the Lomagundi Event(13, 22).We also tabulate the cumulative amounts of reactants con-
sumed (“Consumed” columns in Table 1) and products gener-ated (“Produced” columns in Table 1) during the three modelruns, and present them together with estimates of crustal massesculled from the literature. For siderite, the amounts required fordriving the Lomagundi Event were likely available for oxidationat 2.3 Ga. In the case of carbonate carbon, its weathering in thefirst two runs is actually reduced due to the lower pCO2, and inall cases, our calculated amount of carbonate that was requiredconstitutes only a small fraction of the existing reservoir. Ourcalculated mass of products, in particular of Fe2O3 and Corg, fallwithin the range of values estimated for modern reservoir sizes.Our calculated gypsum production stands at roughly 20% ofthe total modern exogenic sulfur reservoir (which likely existedentirely as sulfide sulfur prior to the Lomagundi event). Our
0 10 20 30 40 500
2
4
6
8
10
12
14
16
Fwsiderite
[Tmols/yr]
13C
[‰]
= 0 = 1
Fig. 2. Effect of siderite oxidation and organic carbon burial on the carbonisotopic composition of the exogenic carbon pool. Calculated according toEq. 10 using a total carbon input flux of 50 Tmol/y. The range of valuesobserved during Lomagundi times, 8‰ < δ13C < 10‰, require the oxidationof 13–35.5 Tmol/y (1012 mol/y) of siderite (dashed lines), depending on therelative proportions of siderite oxidized together with sulfide (α= 0, blueline) versus siderite oxidized together with FeSiO3 (α= 1, green line).
0 100 200 300
0
5
10
Time (Ma)
13C
[‰]
0 100 200 300
0
5
10
15
20
25
pCO
2 [103 p
pmv]
Time (Ma)0 100 200 300
0
0.1
0.2
0.3
0.4
0.5
pO2 [
atm
]
Time (Ma)
0 100 200 3000
10
20
30
Fw F
eCO
3 [Tm
ol/y
r]
Time (Ma)
A B
C D
Fig. 3. Dynamic model runs simulating the Lomagundi Event. In all modelruns, organic carbon burial is increased such that a large (+10‰) positiveδ13C excursion lasting 130 m.y. is generated (A). With increasing amounts ofsiderite and iron silicate oxidation (B), minimum pCO2 values rise (C), whilepeak pO2 values decline (D). When organic carbon burial is not accompaniedby any siderite oxidation (solid black line), the deficit in CO2 for organiccarbon burial drives pCO2 to the extremely low level of 46 ppm. With a smallamount of siderite oxidized (red dash-dotted line), pCO2 declines to a moremoderate 2,800 ppm; with an intermediate amount (dashed green line),pCO2 declines to 7,600 ppm; and with a high amount (blue solid line), pCO2
does not decline at all. In all cases, O2 rises in association with the δ13C ex-cursion. Minimum modeled pO2 is 0.14 atm, or approximately two thirds themodern value, suggesting substantial oxygen accumulation in associationwith the Lomagundi Event. Overall, the modeling indicates that, when ac-companied by siderite oxidation, the Lomagundi Event can be successfullyaccommodated without it giving rise to nonphysical ocean or atmosphericchemistries.
calculated value is consistent with evidence for moderate seawater sulfate concentrations and gypsum precipitation at 2.1 Ga(23). Thus, the oxidation of a large preexisting sedimentaryreservoir of siderite, following the rise of O2 in the atmosphere, cansuccessfully accommodate the existence of a large, protracted,positive carbon isotope excursion, driven by organic carbon burial,without it resulting in nonphysical ocean or atmospheric chemistriesor violation of global mass balance constraints.
DiscussionOur proposal that siderite oxidation played a key role in theLomagundi Event helps resolve several conundrums related toits timing and duration. It has been pointed out (4) that the orderof oxygenation (as indicated by the disappearance of the massindependent fractionation of sulfur isotopes; Fig. 1) and theorganic carbon burial event (as evidenced by the positive δ13Cvalues 200 m.y. later) is reversed from what would be expected iforganic carbon burial was responsible for the rise of O2. Ourmechanism offers a plausible explanation for the observed orderof events. Geologically slow processes, such as changes in platetectonic regime, secular mantle cooling, or a shift in the locus ofvolcanism leading to changes in volcanic gas composition [amongnumerous proposed mechanisms (24)], were likely responsible fordriving a gradual long-term increase in pO2. We postulate thatsuperimposed on this long-term increase in atmospheric pO2 was apulse of O2 production: Once the threshold for oxidation of sideritewas surpassed, a positive feedback was triggered whereby sideriteoxidation supplied CO2 for organic carbon burial, which in turnsupplied oxygen for further siderite, sulfide, and iron silicate oxi-dation. The siderite was likely oxidized both subaerially, where ex-posed, and in reaction with oxidizing groundwaters that would havepenetrated the continental shelves and intracratonic basins for thefirst time (25). Concomitantly, the burial of siderite would havediminished as a result of a reduction in the inputs of Fe2+ from bothterrestrial weathering and hydrothermal inputs. Delivery of ferrousiron from terrestrial settings would have ceased under a high-O2atmosphere, and hydrothermal inputs of reduced iron would havelikely diminished in an ocean bearing appreciable quantities ofsulfate (26). The oxidation of siderite and burial of organic carbonwould have continued until the siderite reservoir was consumed, onthe timescale of hundreds of million years. Once the large Archeanreservoirs of reduced minerals were exhausted and abundant crustaloxidants were produced, little siderite remained to fuel organiccarbon burial, and O2 production was curtailed.Our model, which shows an initial drop in pCO2 (Fig. 3), is
further consistent with the occurrence of glacial episodes precedingthe Lomagundi Event. It is also consistent with the pattern of δ34Schanges that occur during the Lomagundi interval. Measurementsof both carbonate-associated sulfate (CAS) and evaporite sulfate(27) show an increase in δ34S followed by a protracted decrease,
which coincides with the peak of the δ13C excursion, followed byanother increase (Fig. 4). These trends have been interpreted asreflecting the balance of sulfide burial to sulfide oxidation inresponse to changing O2 levels, during, and in the wake of, theLomagundi Event (27). Our model agrees with and refines thisinterpretation. In our model, there is an initial increase in δ34Sthat is driven by an increase in the fractionation factor associatedwith an increase in sulfate concentrations (28). The δ34S valuesthen decline as large amounts of relatively light sulfate (+7‰)are delivered to the ocean from the oxidation of sulfides, thenrise again as that flux wanes, and finally fall as sulfate concen-trations decline once more. In the model, 14–18% of the total Sreservoir is oxidized to sulfate and sulfate concentrations in theocean reach a peak of 4 mM, which is ∼15% of their modernvalue. Following this modest burst of sulfate production anddeposition during the Lomagundi Event, the δ34 S of sedimentarypyrite indicate that pyrite burial became predominant oncemore, and that dominance was maintained until the second riseof oxygen in the Neoproterozoic (29).Incidentally, the pattern of δ13C and δ34S variation argues
against two otherwise interesting hypotheses that have been putforth in an effort to account for the Lomagundi Event. The firsthypothesis postulates that in response to rising pO2, methano-genic activity in the shallow marine realm led to the creation ofpools of highly 13C enriched carbon, which was then incorporatedinto Lomagundi-aged carbonates (4). However, under this scenario,carbonate-associated sulfate incorporated during the precipitationof these carbonates should be extremely 34S enriched, in contrastto the moderate δ34S values that are observed (27)—although
Table 1. Total amounts of reactants consumed and products generated, in Emol (1018 mol), for the three modelruns that are accompanied by varying levels of siderite oxidation
The ratio between the amount of siderite oxidation accompanied by FeSiO3 oxidation (Eq. 8) versus siderite oxidation accompaniedby FeS2 oxidation (Eq. 9) is given by α. Estimated reservoir sizes given are for the present, except for siderite, which is given at 2.3 Ga.The correspondence between the calculated values and the crustal estimates indicates that, with siderite oxidation, the Lomagundi Eventcan successfully be explained without violation of global mass balance constraints. See SI Appendix for references and additional details.
0
14
21
28
7
35
Age (Myr)2,300 2,250 2,200 2,150 2,100 2,050
Fig. 4. δ34S data from Lomagundi age sediments (27). Light gray boxes areδ34S values measured in sulfate from evaporites (gypsum and anhydrite), anddark gray boxes are δ34S measured in carbonate associated sulfate. Lines arethe same model outputs as in Fig. 3. In the model, the initial increase in δ34Sis driven by an increase in the fractionation factor associated with an in-crease in sulfate concentrations (28). The subsequent decline arises from theinput of light sulfur (+7‰) from sulfide oxidation; the δ34S values then riseagain as that flux wanes, and finally fall as sulfate concentrations declineonce more. Modeled oceanic sulfate levels reach a peak value of 4 mM,which is ∼15% of their modern value.
4 of 6 | www.pnas.org/cgi/doi/10.1073/pnas.1422319112 Bachan and Kump
the controls on incorporation of CAS during carbonate pre-cipitation remain incompletely understood. A stronger casecan be made against the second hypothesis, that the source of13C enrichment in Lomagundi-aged carbonates was the burial of13C-depleted authigenic carbonates in other (unsampled) settings(30). Oxidation of organic carbon via bacterial sulfate reductionduring early diagenesis would have led to the precipitation ofauthigenic carbonates but, at the same time, would have also ledto the precipitation of 34S-depleted pyrite, resulting in correla-tive enrichments in both 13C and 34S in shelf carbonates, incontradiction with the observed inverse correlation.A key remaining question is the mechanism by which the
carbon produced by siderite and pyrite oxidation was directed toorganic productivity. Increased carbon input is a necessary, butnot sufficient, condition for increased organic carbon burial andO2 production. Additional limitations arise from the nutrientrequirements of organisms, chiefly phosphate (31). However, asimple increase in weathering and delivery of phosphate to theocean would not generate a positive δ13C excursion, since theresulting increase in organic carbon burial would be coupled toan identical increase in carbonate carbon burial arising fromincreased delivery of alkalinity. To generate a large positive δ13Cexcursion, the delivery of phosphate from weathering must bedecoupled from the delivery of alkalinity to the ocean. Sulfuricacid weathering is a particularly effective way to mobilize P fromapatite (18), allowing for increased organic carbon burial withoutthe concomitant delivery of alkalinity that would demand car-bonate burial. An additional effective way to supply phosphatefor organic productivity is to reduce the output flux of P relativeto that of organic carbon. More-efficient remineralization ofphosphate allows for the export of more organic carbon per unitphosphate buried. Our model indicates a fivefold increase in C:Pratios through the event (from 106 to 500), which, althoughlarge, is still far less than the values found in much younger blackshales (∼ 4,000). The mechanisms by which the C:P burial ratioincreased (and by inference, why it was kept low before and afterthe Lomagundi Event) are still contested (32, 33). We speculatethat a combination of the two aforementioned mechanisms mayhave been at play: Following the initial increase in pO2, oxidationof sulfides allowed for increased P delivery concomitant withincreased availability of sulfate. The development of euxinic(H2S-rich) bottom waters (34) facilitated mobilization of phosphate
back into the water column. Increased P delivery together withmore efficient P utilization allowed for continued organicproductivity and organic carbon burial. This mechanism wascurtailed following the exhaustion of the siderite reservoirand concomitant drop in sulfate levels (27).A second key question pertains to the size of the Archean
reduced sedimentary reservoirs that were subsequently oxidized.Estimates are uncertain, as they are extrapolated from preservedArchean sedimentary volumes, and thus depend on poorlyknown parameters such as the areal extent of Archean conti-nents and the geologic history of sediment recycling. Nonethe-less, it is reasonable to expect that siderite and sulfide, as well asother reduced crustal minerals, were oxidized following the ini-tial rise of O2. The question of whether the supply of CO2 andacidity from these oxidizing species played a pivotal role indriving O2 production hinges critically on the interpretation ofthe carbon isotopic record. If the highly 13C enriched values arerepresentative of the exogenic carbon pool, and large amounts oforganic carbon were buried during the Lomagundi event, thenmass balance demands that large amounts of siderite and pyritewere oxidized concurrently to compensate for the enhancedpCO2 drawdown and O2 production.We conclude by noting that Garrels and Perry (21), in a paper
predating much of our understanding of Precambrian redoxdynamics, presciently underscored the importance of sideriteoxidation in determining the current redox state of the atmo-sphere: “The role of the oxidation of FeCO3 in creating freeoxygen should be emphasized. The level of oxygen in the presentsteady-state atmosphere would seem to be fortuitous, in that itapparently was controlled by the relative amount of siderite inthe preoxygen rocks.” Once the transfer of carbon from thesiderite to the organic carbon reservoir was complete, a modernoxidizing atmosphere was established, and Earth settled into itslong Mesoproterozoic stasis.
ACKNOWLEDGMENTS. We thank Jim Kasting, Chester Harman, Ying Cui,and Jeff Havig for valuable comments. We thank Noah Planavsky and ananonymous reviewer for constructive reviews. A.B. thanks the CanadianInstitute for Advanced Research for a postdoctoral fellowship given throughthe Earth Systems Evolution Program. L.R.K. acknowledges support from theNational Aeronautics and Space Administration Astrobiology Institute, theUS National Science Foundation Geobiology and Low-Temperature Geo-chemistry Program, and the Canadian Institute for Advanced Research.
1. Kump LR (2008) The rise of atmospheric oxygen. Nature 451(7176):277–278.2. Lovelock JE, Margulis L (1974) Atmospheric homeostasis by and for the biosphere: the
Gaia hypothesis. Tellus 26:2–10.3. Payne JL, et al. (2009) Two-phase increase in the maximum size of life over 3.5 billion
years reflects biological innovation and environmental opportunity. Proc Natl AcadSci USA 106(1):24–27.
4. Hayes JM, Waldbauer JR (2006) The carbon cycle and associated redox processesthrough time. Philos Trans R Soc Lond B Biol Sci 361(1470):931–950.
5. Farquhar J, Bao H, Thiemens M (2000) Atmospheric influence of Earth’s earliest sulfurcycle. Science 289(5480):756–759.
6. Karhu JA, Holland HD (1996) Carbon isotopes and the rise of atmospheric oxygen.Geology 24:867–870.
7. Melezhik VA, et al. (2013) The Palaeoproterozoic perturbation of the global carboncycle: The Lomagundi-Jatuli Isotopic Event. Global Events and the FennoscandianArctic Russia - Drilling Early Earth Project, Frontiers in Earth Sciences, eds Melezhik VA,et al. (Springer, Berlin), Vol. 3, pp 1111–1150.
8. Schidlowski M, Eichmann R, Junge CE (1976) Carbon isotope geochemistry of thePrecambrian Lomagundi carbonate province, Rhodesia. Geochim Cosmochim Acta40:449–455.
9. Martin AP, Condon DJ, Prave AR, Lepland A (2013) A review of temporal constraintsfor the Palaeoproterozoic large, positive carbonate carbon isotope excursion (theLomagundi–Jatuli Event). Earth Sci Rev 127:242–261.
10. Aharon P (2005) Redox stratification and anoxia of the early Precambrian oceans:Implications for carbon isotope excursions and oxidation events. Precambrian Res137:207–222.
11. Walker JC, Hays P, Kasting JF (1981) A negative feedback mechanism for the long-term stabilization of Earth’s surface temperature. J Geophys Res 86(C10):9776–9782.
12. Berner RA, Lasaga AC, Garrels RM (1983) The carbonate-silicate geochemical cycleand its effect on atmospheric carbon dioxide over the past 100 million years. Am J Sci283(7):641–683.
13. Bekker A, Holland H (2012) Oxygen overshoot and recovery during the early Paleo-
proterozoic. Earth Planet Sci Lett 317:295–304.14. Torres MA, West AJ, Li G (2014) Sulphide oxidation and carbonate dissolution as a
source of CO2 over geological timescales. Nature 507(7492):346–349.15. Rasmussen B, Buick R (1999) Redox state of the Archean atmosphere: Evidence from
detrital heavy minerals in ca. 3250–2750 Ma sandstones from the Pilbara Craton,
Australia. Geology 27:115–118.16. Johnson JE, Gerpheide A, Lamb MP, Fischer WW (2014) O2 constraints from Paleo-
proterozoic detrital pyrite and uraninite. Geol Soc Am Bull 126:813–830.17. Melezhik VA, Fallick AE, Rychanchik DV, Kuznetsov AB (2005) Palaeoproterozoic
evaporites in Fennoscandia: Implications for seawater sulphate, the rise of atmo-
spheric oxygen and local amplification of the δ13 C excursion. Terra Nova 17:
141–148.18. Konhauser KO, et al. (2011) Aerobic bacterial pyrite oxidation and acid rock drainage
during the Great Oxidation Event. Nature 478(7369):369–373.19. Ohmoto H, Watanabe Y, Kumazawa K (2004) Evidence from massive siderite beds for
a CO2-rich atmosphere before approximately 1.8 billion years ago. Nature 429(6990):
395–399.20. Veizer J (1978) Secular variations in the composition of sedimentary carbonate rocks,
II. Fe, Mn, Ca, Mg, Si and minor constituents. Precambrian Res 6:381–413.21. Garrels R, Perry E (1974) Cycling of carbon, sulphur and oxygen through geologic
time. Marine Chemistry, The Sea: Ideas and Observations on Progress in the Study of
the Seas, ed Goldberg ED (Wiley, New York) Vol 5, pp 303–336.22. Partin CA, et al. (2013) Large-scale fluctuations in Precambrian atmospheric and
oceanic oxygen levels from the record of U in shales. Earth Planet Sci Lett 369-370:
284–293.23. Reuschel M, et al. (2012) Isotopic evidence for a sizeable seawater sulfate reservoir at
2.1 Ga. Precambrian Res 192:78–88.24. Kasting JF (2013) What caused the rise of atmospheric O2? Chem Geol 362:13–25.
Bachan and Kump PNAS Early Edition | 5 of 6
EART
H,A
TMOSP
HER
IC,
ANDPL
ANET
ARY
SCIENCE
S
25. Rasmussen B, Krapez B, Meier DB (2014) Replacement origin for hematite in 2.5 Gabanded iron formation: Evidence for postdepositional oxidation of iron-bearingminerals. Geol Soc Am Bull 126:438–446.
26. Kump LR, Seyfried WE, Jr (2005) Hydrothermal Fe fluxes during the Precambrian:Effect of low oceanic sulfate concentrations and low hydrostatic pressure on thecomposition of black smokers. Earth Planet Sci Lett 235:654–662.
27. Planavsky NJ, Bekker A, Hofmann A, Owens JD, Lyons TW (2012) Sulfur record ofrising and falling marine oxygen and sulfate levels during the Lomagundi Event.Proc Natl Acad Sci USA 109(45):18300–18305.
28. Habicht KS, Gade M, Thamdrup B, Berg P, Canfield DE (2002) Calibration of sulfatelevels in the Archean ocean. Science 298(5602):2372–2374.
29. Canfield DE, Farquhar J (2009) Animal evolution, bioturbation, and the sulfate con-centration of the oceans. Proc Natl Acad Sci USA 106(20):8123–8127.
30. Schrag DP, Higgins JA, Macdonald FA, Johnston DT (2013) Authigenic carbonate andthe history of the global carbon cycle. Science 339(6119):540–543.
31. Laakso TA, Schrag DP (2014) Regulation of atmospheric oxygen during the Pro-terozoic. Earth Planet Sci Lett 388:81–91.
32. Bjerrum CJ, Canfield DE (2002) Ocean productivity before about 1.9 Gyr ago limitedby phosphorus adsorption onto iron oxides. Nature 417(6885):159–162.
33. Konhauser KO, Lalonde SV, Amskold L, Holland HD (2007) Was there really anArchean phosphate crisis? Science 315(5816):1234.
34. Reinhard CT, Raiswell R, Scott C, Anbar AD, Lyons TW (2009) A late Archean sulfidicsea stimulated by early oxidative weathering of the continents. Science 326(5953):713–716.
35. Lyons TW, Reinhard CT, Planavsky NJ (2014) The rise of oxygen in Earth’s early oceanand atmosphere. Nature 506(7488):307–315.
6 of 6 | www.pnas.org/cgi/doi/10.1073/pnas.1422319112 Bachan and Kump
Table S1: Values of β calculated in accordance with Equation S8. Also shown are the corre-sponding δ13C values (assuming the canonical values of -5 h and 25 h for the carbon inputand photosynthetic fractionation), the RCO2 value, and the pCO2 value required to maintainsteady-state.
perturbed pCO2 to baseline pCO2. We also include the estimated reduction in pCO2 from a48
baseline of 10,000 ppm.49
As can be seen in the values in Table S1, the values of β decline precipitously as the values50
of forg increase. At high values of forg (and δ13C), β is dramatically reduced suggesting that51
non-physically low levels of weathering are required.52
S2 Volcanic carbon inputs and organic carbon burial53
Could a long-lived, but temporary increase in carbon inputs potentially prevent atmospheric54
CO2 levels from crashing, while still allowing increased organic matter burial? To estimate the55
required increase in carbon input the following calculation might be carried out: carbon isotope56
mass balance requires that forg be equal to 0.6 for δ13Ccarb to reach + 10h (if epsilon is 25h).57
Then, assuming the total influx is composed of the initial influx (x), 20% of which is buried58
as organic matter, and an additional influx (y), 100% of which is buried as organic matter, the59
So a doubling of the total carbon input is implied, i.e.:62
y
y + x= 2 (S14)
However, this calculation, similarly to the calculation presented in the opening to Section S1,63
fails to take into account the constraints imposed by alkalinity. Using the framework presented64
in S1 we can carry out a more refined calculation and show the difficulty with this proposed65
solution. We include an additional carbon flux to the mass balance so that it becomes:66
Fvolc + Fextra +(Fworg + Fw
carb
)β
(Fwcarb + Fw
sil)β=
1
1− forg(S15)
The magnitude of Fextra should be large enough so that no reduction in weathering occurs (β ≥67
1), and it must also honor the isotopic constraints which suggest that 60% of incoming carbon68
was buried as organic carbon (forg = 0.6). Using the same values for the fluxes as in Section69
S1 (Fvolc = 5; Fwcarb = 36; Fworg = 9; Fwsil = 4, all in Tmol/yr), substituting β = 1 and forg = 0.6,70
and then solving for Fextra yields:71
5 + Fextra + 9 + 36
36 + 4=
1
1− 0.6(S16)
50 + Fextra
40= 2.5 (S17)
5
Fextra = 50 Tmol/yr (S18)
So, the results indicate that 50 Tmol/yr of extra carbon are required, an amount that is equal to72
the total baseline input flux of carbon (Fvolc + Fwcarb + Fworg). Thus, this calculation indicates73
that a doubling of the total input flux of carbon is required, exactly as indicated by the simple74
calculation above, but now the difficulty becomes apparent: the extra carbon has to be supplied75
in “volcanic” form, that is as CO2 and not HCO3– or CO3
2–. The reason is that carbon in the form76
of carbonate alkalinity (HCO3– + CO3
2–), such as supplied by the weathering of calcium silicates77
and carbonates, is already tied to burial of carbonates and thus is unavailable for organic carbon78
burial (or at least without causing an imbalance in alkalinity and a decline in pCO2). An increase79
in the organic carbon weathering flux is also out of the question since it would simply undo the80
13C enrichment effected by the elevated organic carbon burial. Consequently, an eleven-fold81
increase in volcanism is required (from 5 to 55 Tmol/yr), and not simply a two fold increase82
in total carbon input. This required increase in volcanic input is very large, and, even if such83
an increase were geologically plausible, it would only solve half the problem: an increase in84
volcanic CO2 input would resolve the conundrum of CO2 deficit, but not answer the difficulty85
of O2 accumulation. Thus, other sources of CO2 which also consume O2, such as siderite and86
sulfide oxidation, must be seriously considered.87
S3 Organic carbon burial accompanied by pyrite oxidation88
An increase in organic carbon burial, unaccompanied by increased carbon input, leads to non-89
physically low levels of weathering because of the resulting imbalance in alkalinity. Pyrite90
oxidation can ameliorate this imbalance by supplying acidity, thus bolstering pCO2 levels. As91
stated in the main text, both the acidification of limestones by sulfuric acid and the sulfuric acid92
6
weathering of silicates equally lead to the release of carbon dioxide. The former acts by direct93
conversion of carbonate to CO2, while the latter by replacing carbonic acid weathering with94
sulfuric acid weathering and thus allowing volcanic CO2 to go un-consumed. Consequently,95
the resulting CO2 is available for organic carbon burial without any charge balance constraints,96
in contrast to carbonate alkalinity ([HCO3–],[CO3
2–]). It is, nonetheless, important to note that97
the supply of carbon from the sulfuric acid weathering of silicates should not exceed the flux of98
volcanic carbon that would otherwise exit as CaCO3 (4 Tmol/yr). Acidification of Ca-silicates99
in excess of this quantity would lead to an increase in CaCO3 burial and a decrease in the δ13C100
of the ocean-atmosphere system.101
To constrain the required fluxes we augment our previous carbon cycle model with a sulfate102
box, with one input flux of sulfate from pyrite oxidation and one output flux of CaSO4 (Figure103
S2):
C
Ca
5
9
4
40-X
10+X
36
SO4
X
X
Figure S2: Basic carbon cycle model for the ocean-atmosphere carbon pool with a sulfate box.X is the variable pyrite oxidation flux. Modified from Kump and Arthur (1). Units in Tmol/yr.
104
7
We start with the three mass balance equations, one for each element:105
C : Fvolc + Fworg + Fw
carb = Fborg + Fb
carb (S19)
Ca : Fwsil + Fw
carb = Fbcarb + Fb
sulf (S20)
S : Foxpyr = Fb
sulf (S21)
Rearranging we get:106
Fwsil + Fw
carb − Foxpyr = Fb
carb (S22)
Dividing by Fbcarb107
Fvolc + Fworg + Fw
carb
Fbcarb
=1
1− f(S23)
And substituting back into the expression for carbon:108
Fvolc + Fworg + Fw
carb
Fwsil + Fw
carb − Foxpyr
=1
1− f(S24)
Then isolating the oxidation flux of pyrite gives:109
Foxpyr = (Fw
sil + Fwcarb)− (1− f)(Fvolc + Fw
org + Fwcarb) (S25)
We can then calculate the necessary pyrite oxidation fluxes required to balance the burial of110
organic carbon by removal of calcium as gypsum (Table S2):111
Table S2: Pyrite oxidation flux required to balance elevated organic carbon burial, in Tmol/yr.The O2 imbalance is the extra O2 required for pyrite oxidation beyond that which is producedby organic carbon burial, in Tmol/yr.
8
As an example of how pyrite oxidation makes CO2 available for organic carbon burial, consider112
the end-member case where f = 1, i.e. all the carbon coming in is buried as organic carbon113
and none as carbonate carbon. In this case carbonic acid weathering of silicates and carbonates114
must be zero. For pCO2 to remain elevated, carbonic acid weathering must be replaced by115
sulfuric acid weathering. Thus, a SO4= flux of 40 Tmol/yr, with 4 coming from sulfuric acid116
weathering of Ca-silicates and 36 from acidification of carbonates completely compensates for117
missing carbonic acid weathering. Now, 40 Tmol/yr of Ca that would otherwise exit as CaCO3118
get buried as gypsum (or accumulate in the ocean), making available for organic carbon burial119
40 Tmol/yr of carbon. The only remaining problem is that pyrite oxidation requires 15/8*40 =120
75 Tmol/yr of O2, which is 35 Tmol/yr more than organic carbon burial can supply.121
It is pertinent to note that the oxygen deficit would be even more severe if sulfide oxidation and122
carbonate dissolution occurred without a concomitant increase in organic carbon burial. In the123
framework of the above example 75 Tmol/yr of O2 would be consumed and not just 35 Tmol/yr.124
Consequently, the oxygen imbalance which arises during sulfide oxidation argues against recent125
claims for sulfide oxidation as a long-term source of carbon (2). And, if sulfide oxidation did126
make a significant contribution of carbon during Himalayan uplift, as argued by Torres et al.127
(2), it would have driven a substantial drop in pO2 during the Cenozoic, something for which128
there is no evidence.129
S4 Siderite Oxidation and δ13C130
Siderite oxidation can be called upon as a source of carbon over geologic timescales since, in131
contrast to sulfide oxidation, it produces more CO2 than it consumes O2 (Equation S26) and132
so can supply CO2 in excess of the amount that is required by organic carbon burial to keep133
the process going. However, in the context of the Lomagundi Event, the oxidation of siderite134
9
must have also been coupled to processes that produce acidity (sulfide oxidation) or consume135
O2 (iron silicate oxidation) as siderite oxidation alone together with the simple (carbonic acid)136
weathering silicates could not have generated the Lomagundi Event. The reason is that the137
alkalinity generated during the consumption of the excess carbon by silicate weathering would138
have negated the 13C enrichment effected by organic carbon burial.139
To appreciate this difficulty, and more generally, some of the intricacies of siderite oxidation,140
consider the oxidation of siderite coupled to weathering of silicates and the burial of organic141
carbon and carbonate carbon:142
4FeCO3 + O2 → 2Fe2O3 + 4CO2 (S26)143
CaSiO3 + CO2 → CaCO3 + SiO2 / ∗ x (S27)144
CO2 + H2O→ CH2O + O2 / ∗ y (S28)
One could combine these equations in two ways: first, as an overall CO2-neutral reaction (x +145
y = 4), and second, as an O2-neutral reaction (y = 1). Together with the constraints from the146
carbon isotope record (forg = y/(y + x) = 0.6) one can obtain the desired stoichiometries. In147
the O2 neutral case y = 1 and x = 2/3, and in the CO2 neutral case x = 1.6 and y = 2.4. Hence,148
the CO2-neutral reaction leads to net release of O2 at a rate of 1.4 moles of O2 per 4 moles of149
siderite oxidized, while the O2 neutral reaction leads to net release of CO2, at a rate of 2 1/3150
moles per 4 moles of siderite oxidized.151
These imbalances cannot be maintained over timescales of hundreds of m.y. of years over which152
the Lomagundi Event occurs. In the CO2 neutral reaction one is left with with a large excess of153
O2, and in the O2 neutral reaction case one is left with a large excess of CO2. The duration of154
the event is such that if the organic carbon burial were not fully compensated in terms of both155
CO2 and O2 physically non-permissible atmospheric compositions would quickly arise (i.e. pO2156
10
or pCO2 1 atm). Thus, CO2 or O2 must not appear in the overall reaction for oxidation of157
siderite together with the weathering of calcium silicates; nor for that matter should they appear158
in any other reaction which is postulated to have taken place over the duration of the Lomagundi159
Event. Below we give the appropriate reactions for siderite oxidation coupled to calcium silicate160
weathering, siderite oxidation coupled to Fe-silicate weathering, and siderite oxidation coupled161
to Fe-sulfide oxidation and acidification of carbonates.162
What are the biggest differences between the three reactions? Examine the first reaction above165
(Equation S29): the 1 mol of oxygen produced during the burial of organic matter is balanced166
by the oxidation of 4 moles of siderite which produces 4 moles of CO2. One mole of the 4167
moles of CO2 which is produced goes to the burial of organic carbon while the other 3 to the168
weathering of silicates. The overall result is an forg ratio of 0.25. Thus, this reaction is nearly169
isotopically transparent. The increase in the burial of organic carbon in this case is balanced by170
an increase in the burial of carbonate carbon at a ratio that is very near the long term average171
δ13C of the exogenic cycle. Such a reaction would have contributed little to 13C enrichment172
during the Lomagundi Event, and even then mostly via the relatively heavy δ13C value of the173
siderite carbon. Contrast Equation S29 with Equations S30 and S31. In the latter two reactions174
the burial of organic carbon is not accompanied by the offsetting burial of carbonates making175
them much more effective in driving 13C enrichment.176
Thus, there are more and less efficient ways to drive 13C enrichment via siderite oxidation cou-177
pled to organic carbon burial. When inverting the δ13C record, one is given a degree of freedom178
11
in choosing the varying proportions of the three reactions. By choosing a large proportion of179
siderite oxidation coupled to non-Fe-silicate weathering (Equation S29) one could explain the180
Lomagundi Event in a way that involves an amount of reactants that exceeds those which were181
likely available for oxidation or leads to the production of more products than can be accounted182
for.183
Does such a worst-case-scenario calculation does invalidate the hypothesis? No, it is the best184
case scenario which is the test: if it can be shown that even under the most propitious circum-185
stances the hypothesis fails, then it can be confidently rejected. Hence, in our tables and model186
runs below we focus on the most effective ways to drive the Lomagundi Event and show that187
these do not violate mass balance constraints. We do not claim that back reactions (sulfide188
burial, oxidation of organic carbon) or the weathering of non-Fe-silicates did not occur, but we189
do maintain that these must have been minor in relation to the forward reactions, as demanded190
by constraints imposed by the global mass balance on the one hand and the δ13C (and δ34S)191
record on the other.192
S5 Derivation of Equation 10193
We can use Equations S30 and S31 to construct a carbon isotopic mass balance. We assume194
that some fraction, α, of the siderite is oxidized together with Fe-silicates according to Equation195
S30, and the rest, (1 − α), reacts with sulfides and carbonates according to Equation S31. The196
first reaction (Equation S30) implies that the burial of organic carbon occurs at a 1:1 ratio with197
siderite oxidation, such that a certain portion of organic carbon burial is proportional to the198
siderite oxidation flux associated with iron silicate oxidation. Following the same logic, the199
stoichiometry of the second reaction is such that for every mol of siderite oxidized, 6/7 mol200
of carbonate are acidified and 13/7 mol of organic carbon are buried, inducing the following201
12
isotopic mass balance:202
(1− α)Fwsid δ
wsid +
6
7(1− α)Fw
sid δwcarb =
13
7(1− α)Fw
sid δborg (S32)
Adding the above expressions to the mass balance of the pre-Lomagundi carbon cycle,203
Fw,0in δw,0in = Fb,0
carb δbcarb + Fb,0
org δborg (S33)
gives:204
Fw,0in δw,0in +αFw
sid δwsid + (1− α)Fw
sid δwsid +
6
7(1− α)Fw
sid δwcarb
= Fb,0carb δ
bcarb + Fb,0
org δborg + αFw
sid δborg +
13
7(1− α)Fw
sid δborg
We assume that the isotopic composition of all the organic carbon burial is fractionated by a205
constant amount relative to seawater, so given by (δbcarb − ε), where ε is the photosynthetic206
fractionation. Collecting terms gives the following mass balance equation:207
Fw,0in δw,0in + Fw
sid
[δwsid +
6
7(1− α) δwcarb
]= Fb,0
carb δbcarb+
[Fb,0org +
(13
7− 6
7α
)Fwsid
](δbcarb − ε)
(S34)
Rearranging so that it is given for δ13C of the ocean-atmosphere system:208
δbcarb =Fw,0in δw,0in + Fw
sid
[δwsid + 6
7(1− α) δwcarb
]+ ε
[Fb,0org + (13
7− 6
7α) Fw
sid
]Fb,0org + Fb,0
carb + (137− 6
7α) Fw
sid
(S35)
Note that:209
Fw,0in = Fb,0
org + Fb,0carb (S36)
and210
Fw,0in δw,0in = Fb,0
org(δcarb − ε) + Fb,0carbδ
bcarb (S37)
so that assuming δw,0carb = 0 h gives:211
Fwin δin + Fb
org ε = 0 (S38)
13
Thus:212
δbcarb =Fw,0in δw,0in + εFb,0
org + Fwsid
[δwsid + (13
7− 6
7α) ε]
Fwin + (13
7− 6
7α) Fw
sid
(S39)
and finally:213
δbcarb =Fwsid
[δwsid + ε
(137− 6
7α)](
Fwin + (13
7− 6
7α) Fw
sid
) (S40)
To make the expression more compact we define an alpha prime:214
α′ =13
7− 6
7α (S41)
So the final expression is:215
δbcarb =Fwsid [δwsid + ε α′]
Fwin + α′ Fw
sid
(S42)
Equation 10 in the main text (or Equation S42 above) has a Michaelis-Menten form (as is borne216
out in Figure 2: as the value of Fwsid increases, δbcarb asymptotically approaches the value of217
δwsidα′ + ε, which is approximately 24 h. When the siderite flux is zero, the carbon isotopic218
composition of the ocean-atmosphere system returns to its long-term steady-state value of 0h.219
S6 Global mass balance220
Using Equation S42 it is further possible to calculate the siderite oxidation flux required by a221
Gaussian shaped positive excursion of up to +10 h of a given duration (Figure S3). The total222
siderite flux is then given by the area under the Fwsid curve.223
To calculate an alpha, we use a constrained optimization algorithm (fmincon, Matlab (3)), such224
that a minimum amount of siderite is used, coupled to the constraint that the total integrated225
oxidation of pyrite not exceed the total exogenic sulfur pool of 534 Emol of sulfur. The logic226
behind this choice is that the total amount of exogenic sulfur, which is currently partitioned227
14
0 100 200 3000
2
4
6
8
10
δ1
3C
ca
rb
Time [Ma]0 100 200 300
0
10
20
30
40
50
Fw sid
[T
mo
l/yr]
Time [Ma]
100
130
190
250
B.A. B.A. B.A.
Figure S3: A. Gaussian shaped positive δbcarb excursions up to +10 h with durations of 100,130, 190, and 250 m.y. B. The resulting siderite oxidation flux according to Equation S42, withthe values of alpha set to keep the total sulfate production below 534 Emol. Area under curvesgiven in Table S3. Different colored curves in both plots correspond to different durations.
between the ocean, continental sulfide, and continental sulfate, likely existed entirely as con-228
tinental sulfide prior to the Lomagundi event, and was thus available for oxidation during the229
event. We do not claim that all the sulfur was oxidized and precipitated as gypsum. Rather,230
we utilize this constraint as an upper theoretical bound on the extent of siderite oxidation via231
Equation S31.232
The results are given in Table S3 for four durations (100, 130, 190, and 250 m.y.) and three233
input fluxes (25, 50, and 75 Tmol/yr). The code to produce the figure and table is given in234
the supplementary files. We also include the cumulative amounts of reactants consumed and235
products produced during our dynamic model runs, which are discussed in Section S7. The236
three runs (Model 1–3) given are the same model runs that are presented in Figure S7. Positive237
values for CaCO3 in the first two runs indicate that carbonate dissolution was lower in the238
perturbed state than in steady-state. This is due to the lower than baseline pCO2 which occurs239
Table S3: Top: total amounts in Emol (1018) of reactants consumed and products generated duringa Gaussian shaped δ13C excursion of up to + 10 h, obtained via two different calculations. First, byutilizing Equation 10 (Figure S3), together with the constraint that the amount of pyrite oxidized togetherwith siderite (Reaction 9) did not generate sulfate in excess of the modern CaSO4 + FeS2 reservoirs (534Emol); siderite oxidation accompanied by iron silicate oxidation (Reaction 8) was presumed to make upthe remainder. Second, using the dynamic model runs, in which case runs with non-physical atmosphericchemistries (pO2 1 atm and pCO2 0 atm) were rejected. The three runs (Model 1–3) are the samemodel runs presented in Figure S7. Positive values for CaCO3 in the first two runs indicate that carbonatedissolution was lower in those runs than in steady-state. In all cases the fraction of siderite accompaniedby FeSiO3 oxidation (Reaction 8) is given by α. Bottom: estimated reservoir sizes. All are given for thepresent, except for siderite which is given for 2.2 Ga. Higher Fe2O3 estimate includes oxidized iron incrystalline silicate rocks.
To compare the results of our calculations to measured values, we compiled estimates of rele-241
vant sedimentary reservoirs sizes. First, we give some comments on the estimates culled from242
the literature which are presented at the bottom of Table S3. We then discuss the degree of243
correspondence between our calculations and the estimates. The most detailed inventory of the244
16
sizes of the crustal reservoirs of CaCO3, Corg, FeS2, and CaSO4 can be found in Holser et al.245
(6), which also summarizes previous estimates from Holser and Kaplan (10), Li (11), Garrels246
and Perry (4), Schidlowski and Eichmann (12), Nielsen (13), and Garrels and Lerman (14). The247
reported values fall in the range of 84 – 294 Emol S for the sulfide reservoir; 81 – 240 Emol S248
for the sulfate reservoir; 3505 – 6460 Emol for CaCO3; and 930 – 1300 Emol for Corg. Hayes249
and Waldbauer (7) give an updated and in-depth discussion, which, in addition to the values250
given by Holser et al., summarizes more recent values for sedimentary compilations given by251
Wedepohl (15), Hunt (16), Des Marais (17), Berner (18), and Arvidson et al. (19), which fall252
within the range of 2800 – 6500 Emol of CaCO3 and 675 – 1300 Emol of Corg. They also cite253
mass-age data on carbonates from Wilkinson and Walker (20), which suggest a somewhat larger254
carbonate reservoir size (7900 – 9600 Emol), and thus a correspondingly larger Corg reservoir255
size (1400 – 1700).256
Estimates for crustal iron are given by Yaroshevsky (8), who summarizes previous results from257
Vinogradov (21) and Ronov et al. (22): the sedimentary shell is estimated to contain 393.75258
Emol of oxidized iron and 1018.5 Emol of reduced iron, with another 888 Emol of oxidized259
iron and 1867 Emol of reduced iron in the upper crust (granitic-metamorphic shell), totaling260
1280 Emol of oxidized iron and 2886 of reduced iron. For oxidized iron, Garrels and Perry261
estimate the excess oxidized iron in sedimentary rocks at 263 Emol Fe2O3. Sleep (5) gives a262
lower estimate for sedimentary oxidized iron: 50 Emol of sedimentary Fe2O3, requiring 100263
Emol of Fe and 25 Emol of O2 to have been produced, but a higher estimate for oxidized crustal264
iron (including hard rocks) of 4000 Emol of Fe2O3, requiring 8000 Emol Fe and 2000 Emol265
O2, though he does acknowledge that the uncertainties in composition of the lower continental266
crust could lead to the lower, but still very large, estimate of 2000 Emol Fe2O3, requiring 1000267
Emol of O2 to have been produced. Hayes and Waldbauer (7) cite Ronov and Yaroshevsky268
(23) for an estimate of 1020 O2 equivalents, or 2040 Emol of Fe2O3, and a higher estimate269
17
from Goldschmidt of 1860 Emol O2 equivalents, or 3720 Emol Fe2O3—near the estimate of270
Sleep. Altogether, the estimates converge on oxidized crustal iron as being the most substantial271
reservoir of O2 equivalents.272
For siderite, other than the anecdotal descriptions of its relative abundance given in Ohmoto273
et al. (24), the only quantitative estimate is by Ronov et al. (9) based on observations of the274
Russian platform. At 2.3 Ga, they estimate that Jasperlites (which, as stated in the text, are275
considered to be altered siderites) to constitute 14% of the sedimentary shell of 25,000x1020 gr,276
or 3.0x1021 mol FeCO3. Incorporating models for sediment recycling results in an even larger277
estimate of 22%, as discussed by Garrels and Mackenzie (25). In fact, the large increase in278
Jasperlites followed by their total disappearance in the phanerozoic is perhaps one of the most279
remarkable features of Ronov et al’s data, as already pointed out by Garrels and Mackenzie (25).280
Garrels and Perry (4) give a value of 350 Emol of sedimentary FeCO3, though this estimate is281
based on the amounts required to balance oxidized sedimentary iron rather than by rock data.282
Nonetheless, their logic holds, and the much larger estimates for oxidized iron which include283
iron in crystalline silicate rocks correspond very well to the total estimated the mass of siderite284
available for oxidation at 2.3 Ga.285
How did such large amounts of oxidized iron accumulate in the crust? During the Archean iron286
was likely delivered in reduced form from the weathering of silicate rocks and the dissolution287
of pyrite and siderite in sedimentary rocks. It likely exited the ocean in equally reduced form288
as pyrite and siderite. Ferric to total iron ratios in shales are not much different from mantle289
values until the Great Oxidation Event (38) and it is only following during it that the ferric iron290
content of shales rises substantially. During the Lomagundi Event pO2 rose to significant levels291
for the first time and reduced iron delivered from weathering became oxidized on land for the292
first time. The weathered iron accumulated as oxidized iron in shales and as redbeds on the con-293
18
tinents (redbeds make their first appearance following the Great Oxidation Event). In addition,294
diagenetic alteration of reduced iron as oxidizing fluids flowed through sedimentary basins for295
the first time likely also led to substantial accumulation of iron oxides (39). Oxidation of iron296
in mid-ocean ridge basalts by oceanic sulfate (which likely rose from very low Archean levels297
for the first time) likely also contributed to growth of the oxidized iron reservoir. The iron in298
oceanic basalts undergoing subduction would have been particularly likely to be incorporated299
into crystalline rocks of the continental crust. During the remainder of the Proterozoic, fol-300
lowing the Lomagundi Event, weathering would have delivered both oxidized iron and reduced301
iron. The fraction of iron that in reduced form was likely oxidized subaerially. Some of that iron302
accumulated as redbeds while some of it was delivered to the marine realm where it was likely303
reduced and exited as pyrite. This situation likely persisted until oxidation of the deep ocean,304
much later, during the Phanerozoic, which led to an additional loci of iron oxide deposition in305
deep sea.306
Comparing the values computed according to Equation 10, and the estimated sizes of the crustal307
reservoirs, in particular of organic carbon, oxidized iron, and sulfate, given in Table S3, shows308
that the lower estimates (corresponding to a lower duration of the Lomagundi Event and/or309
lower estimates for total carbon input) match reasonably well to the existing crustal reservoirs.310
Moreover, the larger estimates for organic carbon, though they exceed the estimates for the311
current reservoirs, do not invalidate the conclusions, as material could have been lost since the312
Lomagundi Event. For instance, subduction of organic carbon or its oxidation subsequently313
to the Lomagundi Event could have led to a smaller fraction of the produced organic carbon314
being preserved. In particular, the hypothesis that subduction of organic carbon as well as its315
oxidation to methane were accelerated during and immediately following the Lomagundi Event316
is an attractive one, as it would explain the unidirectional and permanent nature of the Earth317
surface oxidation that occurred in association with it.318
19
Many workers have highlighted the fact that Earth has an excessively oxidized atmosphere319
and crust. In particular, Lovelock (27, 28) pointed out that while Mars and Venus are at the320
equilibrium redox potential appropriate to their stations in the solar system, that of Earth is far321
more oxidized than its position would suggest. The implication is that Earth’s unique features,322
plate tectonics and life, are likely responsible for its current redox state (see also Hayes and323
Waldbauer, 7). This process was very likely aided by a methane “hydrogen balloon” (Lovelock324
and Lodge (27), Catling et al. (29)) which transported hydrogen to the upper atmosphere. We325
suggest that the processes of hydrogen loss to space and to the mantle may have been pulsed326
as well, and tightly coupled to episodes of organic carbon production (and burial). Thus, the327
oxidized products accumulated in the crust and atmosphere, while the reducing power was328
transferred to organic carbon, and then subsequently subducted into the mantle and lost to space.329
It is worth emphasizing the main difference between our own interpretation of the carbon cycle330
and that of Hayes and Waldbauer, and others before them, is that while they postulate that331
the accumulation of oxidants (Fe3+, SO2−4 , O2) and reductants (Corg, FeS2) occurred gradually332
throughout the Geozoic, we acknowledge the possibility that large portions of these reservoirs333
could have been accumulated, and destroyed, in shorter periods of geologic time (on the order334
of tens to hundreds of millions of years).335
S7 Model Description336
The model includes mass boxes for the oceanic concentrations of carbon, calcium, sulfur, phos-337
phate, oxygen, and alkalinity, as well as isotopic mass for carbon, calcium and sulfate (Figure338
S4). The sedimentary reservoirs included in the model are organic carbon, carbonate carbon,339
sulfide sulfur, sulfate sulfur, reduced iron, and siderite. A carbonate system solver, modified340
from Emerson and Hedges (30) and Zeebe and Wolf-Gladrow (31), is used to calculate pCO2,341
20
pH, and the carbonate saturation state (Ω) from values of alkalinity and DIC at every model342
timestep. Complete list of constitutive equations and flux relations is given in Section S8. The343
full code is supplied in supplemental files.344
SO4
O2
ALK
p
CO2
pH
C sys
!44Ca
!13C
!34S
FeCO3
FeO
Figure S4: Sketch of model setup. Oceanic and atmospheric reservoirs in blue, sedimentaryreservoirs in brown. Burial fluxes from the ocean in blue, weathering fluxes in brown, subduc-tion fluxes and volcanic fluxes in red. Csys stands for carbonate system solver.
The first numerical experiment we perform is a simple sanity test: we force the model with345
weathering fluxes of siderite, sulfide, and carbonate, and burial of organic carbon and sulfate,346
in stochiometric relations according to Equation 9 in the main text, such that a +10 permil347
excursion results. This perturbation is a “Goldilocks” solution with the reactants and products348
exactly balanced so should incur no changes in pO2 or pCO2. Results are shown in Figure349
S5. In Table S4 we give the values produced by the model, calculated in two ways: firstly,350
by integrating the relevant fluxes with respect to time, and secondly by subtracting the initial351
21
Duration FeS2 CaCO3 FeSiO3 FeCO3 Fe2O3 CaSO4 CH2O αModel Int 100 my -249.44 -498.89 0 -582.03 415.74 498.89 1080.92 0Model Sub 100 my -249.44 -498.88 0 -582.03 415.74 498.88 1080.92 0Calc 100 my -249.06 -498.11 0 -581.13 415.09 498.11 1079.24 0
Table S4: Table comparing model output to analytical calculations. First row (Model Int) isobtained by integrating the time-varying fluxes in the model. The second row is obtained bysubtracting initial and final sedimentary reservoir sizes. The third row (Calc) are the same valuesgiven in Table S3 (5th row: 50 Tmol/yr, 100 m.y.) but without rounding. Nearly identical solu-tions between the first and second rows indicate that the model conserves mass. Nearly identicalsolutions of the semi-analytical calculation and numerical model indicate that the model is ac-curate.
and final sedimentary reservoir masses. The differences between the first row (integration) and352
second row (subtraction) are in the second decimal point, indicating that the model preserves353
mass. The differences between the model and the calculation arise due to truncation error, as354
well as the maximal δ13C in the model, which is 10.001h, as opposed to an exact 10h in the355
calculation. The differences are on the order of 0.15% of the analytical solution, indicating that356
the model is reasonably accurate.357
The next numerical experiment we perform is to incorporate parameterizations for the weath-358
ering and burial fluxes, and force a pulse of organic carbon burial. We first force the model359
without any pyrite oxidation, and we then add two different parameterizations of the pyrite360
oxidation flux. We show the model outputs for δ13C, pCO2, and pO2 in Figure S6.361
The model results are that under the scenario of pyrite oxidation according to the Williamson362
and Rimstidt parameterization (green dashed line) pCO2 falls to 46.6 ppm. Under the scenario363
where the pyrite and gypsum weathering fluxes are set to the modern ones with modification364
for the reservoir size (blue solid line), pCO2 falls the least, but still reaches very low values365
of 108.3 ppm. The decline in pCO2 is driven by the burial of organic carbon: the removal of366
DIC from ocean water, with little accompanying alkalinity, drives the carbonate system from367
neutrality and towards a zone of higher pH and lower pCO2. The weathering fluxes, which are368
22
0 5 10
x 107
54
54
54
54
54
CC
0 5 10
x 107
2.08
2.1
2.12
x 10−6
pO
2
0 5 10
x 107
1.5
2
2.5
3x 10
21
Mse
d
Co
rg
0 5 10
x 107
2.8
3
3.2
3.4
3.6x 10
21
Mse
d
Fe
CO
3
0
5
10
15
δC
arb
1.0822
1.0822
1.0822
1.0822
x 104
pC
O2
5.5
5.6
5.7
5.8
5.9
6
x 1021
Mse
d
Ca
CO
3
0
0.5
1
1.5
2
x 1013
Fw F
eC
O3
A. B.
C D
Figure S5: Model outputs for a pulse of carbon burial coupled to carbonate carbon acidification,sulfide oxidation, and siderite oxidation in accordance with Equation 9 in the main text, suchthat a +10 permil excursion results. This perturbation does not result in changes in pO2 or pCO2
since it follows a stoichiometrically balanced reaction for both species. A. The concentration ofinorganic carbon in the ocean (CC) and its isotopic composition (δCarb). B. Atmospheric oxygenand carbon dioxide concentrations (pO2, pCO2) C. Mass of sedimentary organic carbon andcarbonate carbon reservoirs (Msed
Corg and MsedCcarb). D. Mass of siderite and the siderite oxidation
flux (MsedFeCO3 and FW
FeCO3).
23
0 100 200 300
0
2
4
6
8
10
Time (Ma)
δ1
3C
0 100 200 30010
−10
10−5
100
105
Time (Ma)pO
2 [a
tm]
0 100 200 30010
0
105
Time (Ma)
pC
O2 [ppm
v]
A. B.
C.
Figure S6: Model outputs for a pulse of carbon burial unaccompanied by increased carboninput. All simulations show a drop in pCO2 to extremely low levels. Under the scenario of nopyrite oxidation (red dash-dotted line) pCO2 falls to 3.76 ppm, close to the theoretical valuecalculated in Table S1. Under the scenario of pyrite oxidation according to the Williamsonand Rimstidt parameterization (green dashed line) pCO2 falls to 46.6 ppm. Under the scenariowhere the pyrite and gypsum weathering fluxes are set to the modern ones with modificationfor the reservoir size (blue solid line), pCO2 falls the least, but still reaches very low valuesof 108.3 ppm. The inclusion of pyrite oxidation impacts the pO2 response as well. Withoutpyrite oxidation, oxygen accumulates and remains high, whereas with pyrite oxidation, pO2returns to baseline after the perturbation. The peak values for pO2 are 0.46 atm (2.1 PAL) underthe Williamson and Rimstidt parameterization, and 3x10−5 (3x baseline) under the alternativeparameterization. Note log scale on y axes in B. and C.
24
set to be proportional to pCO20.3, respond by declining, thus lowering the input of carbon and369
alkalinity from weathering. The carbonate burial flux, which is proportional to the carbonate370
saturation state, declines as well, even as pH increases, because of the reduction in the input of371
Ca2+. The system thus approaches a new equilibrium, one in which the inputs of alkalinity are372
equal to the much reduced outputs of alkalinity. When sulfide oxidation is added to the model373
weathering fluxes, the contribution of sulfate helps mitigate the imbalance in alkalinity, thus374
resulting in higher pCO2 levels. The oxidation of pyrite also leads to a more realistic behavior375
of O2, whereby instead of accumulating in the atmosphere and remaining constant after the376
perturbation, O2 returns to steady-state after the perturbation (compare dash-dotted curve to the377
dashed and solid lines in Figure S6.B).378
Since the parameterization of the pyrite oxidation flux exerts such a strong control on the pCO2379
and pO2 response of the model during the positive excursion, a brief description of the avail-380
able choices is in order. The first, simpler yet probably less appropriate parameterization, is381
one that takes the current modern estimates and scales them to the size of the Late Archean -382
Early Proterozoic sulfate and sulfide sedimentary reservoirs, which were likely much smaller383
and much larger, respectively. Thus, if the Phanerozoic estimates for the sulfide and sulfate384
sedimentary reservoirs are 294 and 240 Emol (a ratio of 0.55), and associated fluxes are 0.93385
and 0.76 Tmol S /yr (using the values from Garrels and Perry, 4), then assuming a 0.99 ratio386
of sedimentary reservoir masses in favor of sulfide gives fluxes of 1.72 and 0.01 Tmol S /yr for387
the Late Archean - Early Proterozoic. We allow the sulfide oxidation flux to scale as the square388
root of the ratio of pO2 to its initial value, and linearly with the size of the remaining pyrite389
reservoir:390
FwFeS2 = Fw,i
FeS2 ∗[pO2
pO2,i
]0.5∗
[Msed
FeS2
MsedFeS2,i
](S43)
The alternative formulation takes into account the kinetics of pyrite oxidation as experimentally391
25
constrained by Williamson and Rimstidt (32) (see also Bolton et al., 33):392
RFeS2 =10−8.19 [ pO2 ∗KO2
H ]0.5
[H+]0.11∗ 31536000; (S44)
where RFeS2 is the rate at which pyrite is oxidized with units of [ molm2 yr
] (the factor of 31536000393
converts from seconds to years). It is observed to scale with the square root of the dissolved394
aqueous O2 concentration (which is given by atmospheric O2 multiplied by Henry’s constant for395
oxygen, 0.00126 [mol/L/Atm at 25C]), and to be weakly inversely proportional (0.11 power) to396
the concentration of protons in the weathering solution, which we calculate as pH of pristine397
rainwater in equilibrium with atmospheric CO2, which is given as the roots of a cubic equation398
with the appropriate rate constants: kH = 10−1.47, k1 = 10−6.35, k2 = 10−10.33, kw = 10−14.400
Once the oxidation rate is known, a scaling relationship between the calculated rate of pyrite401
oxidation under modern pCO2 and pO2 conditions (RmodFeS2
), and the estimated modern pyrite402
oxidation flux (Fw,modFeS2 ) and reservoir size (M sed,mod
FeS2 ) can be obtained:403
KoxPyr =
Fw,modFeS2
RmodFeS2∗Msed,mod
FeS2
(S46)
This constant (KoxPyr) can then by used to calculate the pyrite oxidation flux under different404
boundary conditions of reservoir size and oxidation rate (as a function of pO2 and pCO2):405
FwFeS2 = Kox
Pyr ∗ RFeS2 ∗MsedFeS2; (S47)
The result of scaling to a higher pCO2, lower pO2, and larger reservoir size, is that pyrite406
oxidation is calculated to be substantially smaller in the low-O2 Archean (approximately 1x109407
Tmol/yr versus 1x1011 Tmol/yr in the present). The initial magnitude of the sulfide oxidation408
flux under the different parameterizations (4.47x109 or 1.67x1012 Tmol/yr) makes a difference409
26
for how high pO2 needs to go up to increase the flux such that it balances the accumulation of410
alkalinity due to organic carbon burial. The higher the initial value, the less pO2 has to increase411
in order to balance the input and outputs of carbon and alkalinity. Nonetheless, because of the412
inherent imbalances associated with pyrite oxidation, which lead to a shortage of carbon, pCO2413
falls to low levels, irrespective of the parameterization used.414
Next, we perturb the model with a Gaussian shaped excess organic carbon burial flux, and allow415
for siderite and iron silicate oxidation, in addition to pyrite oxidation. The results are shown416
in Figure S7. With increasing siderite contribution the pO2 peak increases, the pCO2 minimum417
rises, and the pCO2 maximum becomes more pronounced. In the first case (red dash-dotted418
line) with 813 Emol of FeCO3 consumed pCO2 falls to 2800 ppm and then rises 13,600 ppm.419
In the second case (green dashed line) with 1118 Emol of FeCO3 consumed pCO2 falls to 7600420
ppm and then rises 16,500 ppm. In the third case (blue solid line) with 1424 Emol of FeCO3421
consumed pCO2 is not reduced at all and then rises 21,150 ppm. Perhaps counter-intuitively,422
the increasing amount of siderite causes the pCO2 peak associated with the declining limb of423
the δ13C excursion to be smaller. This is because the oxidation of siderite (and iron silicates)424
diverts O2 from sulfide oxidation, and it is the sulfide oxidation flux which drives the increase425
in pCO2 through its effects on alkalinity and thus carbonate burial. A higher sulfide oxidation426
flux leads to a smaller carbonate burial flux because of the sulfide oxidation’s contribution of427
acidity, and hence a larger imbalance which then is required to correct itself, leading to a pCO2428
overshoot.429
We show the changes in [SO4] and δ34S of the oceanic sulfate box associated with each of430
these three model runs in Figure S8 and Figure S9. The resulting trends in δ34S, as stated431
in the main text, arise mainly due to two effects. The sharp rise and decline in δ34S are a432
result of increased fractionation associated with increased availability of sulfate. We use the433
27
0 100 200 300
0
2
4
6
8
10
δ1
3C
[‰
]
0 100 200 300−0.1
0
0.1
0.2
0.3
0.4
pO
2 [a
tm]
0 100 200 3000
5
10
15
20
25
pC
O2 [1
03 p
pm
v]
Time (Ma)
Figure S7: Model outputs for a pulse of carbon burial accompanied by siderite, pyrite, andiron silicate oxidation. With increasing siderite contribution the pO2 peak increases, the pCO2minimum rises, and the pCO2 maximum becomes more pronounced. In the first case with a lowamount of siderite oxidized (red dash-dotted line) pCO2 falls to 2800 ppm and then rises 13600ppm, and pO2 rises to 0.14 Atm. In the second case with an intermediate amount of sideriteoxidized (green dashed line) pCO2 falls to 7600 ppm and then rises 16500 ppm, and O2 rises to0.23 Atm. In the third case with a large amount of siderite oxidized (blue solid line) pCO2 isnot reduced at all and then rises 21150 ppm, while pO2 rises to 0.34 Atm.
28
0 100 200 3000
0.5
1
1.5
2
2.5
3
3.5
4
[SO
4] [m
M]
Time (Ma)
0 100 200 3005
10
15
20
25
30
δ3
4S
[‰
]
Time (Ma)
Figure S8: Variation in the [SO4] and δ34S of the oceanic sulfate box, same model runs as inFigure S7 above.
parameterization suggested by Habicht et al. (36), with a fractionation factor of 1.029 above434
a threshold of 190 µM and a linear decrease to 1.000 as sulfate decreases towards zero. Since435
all model runs include a rise in SO4 above 190 µM there is little variation in the response with436
varying pyrite oxidation rates. In contrast, the drop in δ34S in the middle is a result of the influx437
of light sulfide to the sulfate box from pyrite oxidation, and some variation is apparent: lower438
δ34S result from increased sulfate input. As shown in Figure S9 and in the main text these439
results fit the data from Planavsky et al. (37) quite well, lending support to our interpretation of440
the events which occurred during Lomagundi times.441
29
0
142128
7
35
Age (Myr)2,300 2,250 2,200 2,150 2,100 2,050
Figure S9: Plot of modeled δ34S variation overlaying data from Planavsky et al. (37). In themodel the sharp rise and fall in δ34S are a result of increased fractionation associated withincreased availability of sulfate, while the drop in δ34S in the middle is a result of the influxof light sulfide to the sulfate box from pyrite oxidation. Lines are the same model runs as inFigures S7 and S8.
30
S8 Model Equations442
31
Table S5: Steady-state values for model fluxes and reservoirsDescription Name Value UnitDissolved inorganic carbon DIC 54.0 · 10−3 mol/kgDissolved calcium Ca 0.13 · 10−3 mol/kgAlkalinity ALK 58.5 · 10−3 eq/kgDissolved phosphate PO4 0.25 · 10−6 mol/kg
Partial pressure of CO2 pCO2 10822 ppmvCalcite Saturation Ω 1.45 -pH ph 8.03 -
Burial of organic carbon Fborg 10 · 1012 mol/yrBurial of CaCO3 Fbcarb 40 · 1012 mol/yrBurial of phosphate Fbp 9.3 · 1010 mol/yrC:P burial ratio CP 106 -
δ13C of volcanic flux δvolc -5 permilδ13C of carbonate weathering δwcarb 0 permilδ13C of organic carbon weathering δworg -25 permil
δ13C of carbonate burial δ 0 permilδ13C of organic carbon weathering δborg -25 permilPhotosynthetic fractionation ε 25 permilFlux values after DePaolo (40) and Kump and Arthur (1). Magnitude of photosyn-thetic fractionation from Hayes et al. (41). Concentrations of carbon, calcium, andphosphate converted to masses using an ocean volume of 1.32 · 1021 L and salinity of1.035 kg/L.
32
Table S6: Isotopic mass equations for oceanic reservoirs
dMC
dt= Fw
Cvolc+ Fw
CaCO3+ Fw
Corg− Fb
Corg− Fb
Ccarb(S48)
dMCa
dt= Fw
Casil+ Fw
CaCO3+ Fw
CaSO4− Fb
CaCO3− Fb
CaSO4(S49)
dMPO4
dt= Fw
PO4− Fb
PO4(S50)
dMS
dt= Fw
Svolc+ Fw
CaSO4+ Fw
Pyr − FbCaSO4
− FbPyr (S51)
dO2
dt= Fb
Corg+ 1.875 Fb
Pyr − FwCorg− 1.875 Fw
Pyr − fredoxvolcCFwCvolc
(S52)
− 1.875 fredoxvolcSFwSvolc
(S53)
dδCdt
=[FwCvolc
(δCvolc− δC) + Fw
CaCO3(δwCaCO3
− δC)
+ Fworg(δ
worg − δC)− (Fb
org)(−εC)] 1
MC
(S54)
dδCa
dt=[FwCasil
(δwCasil− δCa) + Fw
CaCO3(δwCacarb
− δCa)
+ FwCaSO4
(δwCasulf− δCa)− (Fb
CaCO3+ Fb
CaSO4)(−εCa)
] 1
MCa
(S55)
dδSdt
=[FwSvolc
(δSvolc − δS) + FwCaSO4
(δwSulf − δS) + FwPyr(δ
wPyr − δS)
− FbPyr(−εS)
] 1
MS
(S56)
33
Table S7: Mass and isotopic mass equations for sedimentary reservoirs
