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Proc Natl Acad Sci U S A. Aug 9, 2005; 102(32): 11131–11136.
Published online Aug 1, 2005. doi:  10.1073/pnas.0504878102
PMCID: PMC1183582

The Paleoproterozoic snowball Earth: A climate disaster triggered by the evolution of oxygenic photosynthesis


Although biomarker, trace element, and isotopic evidence have been used to claim that oxygenic photosynthesis evolved by 2.8 giga-annum before present (Ga) and perhaps as early as 3.7 Ga, a skeptical examination raises considerable doubt about the presence of oxygen producers at these times. Geological features suggestive of oxygen, such as red beds, lateritic paleosols, and the return of sedimentary sulfate deposits after a ≈900-million year hiatus, occur shortly before the ≈2.3–2.2 Ga Makganyene “snowball Earth” (global glaciation). The massive deposition of Mn, which has a high redox potential, practically requires the presence of environmental oxygen after the snowball. New age constraints from the Transvaal Supergroup of South Africa suggest that all three glaciations in the Huronian Supergroup of Canada predate the Snowball event. A simple cyanobacterial growth model incorporating the range of C, Fe, and P fluxes expected during a partial glaciation in an anoxic world with high-Fe oceans indicates that oxygenic photosynthesis could have destroyed a methane greenhouse and triggered a snowball event on timescales as short as 1 million years. As the geological evidence requiring oxygen does not appear during the Pongola glaciation at 2.9 Ga or during the Huronian glaciations, we argue that oxygenic cyanobacteria evolved and radiated shortly before the Makganyene snowball.

Keywords: oxygen, Makganyene glaciation, Huronian glaciations, cyanobacteria

After the rise of life itself, the most radical transformation of Earth's biogeochemical cycles occurred in the transition from an anoxic to an oxic world. This transformation took place in three phases. First, oxygenic photosynthesis evolved and brought into the world locally oxic environments. Second, oxygen became a major component of the atmosphere; some authors (1, 2) have suggested that this period was a protracted phase during which the ocean became euxinic. Finally, the whole ocean–atmosphere system took on its modern oxygen-dominated cast.

Although the timing of and relationship between the three stages have been topics of active research for many decades, there is still a wide divergence of opinion. Evidence from organic biomarkers (35) and arguments concerning trace element mobility (6) and biological productivity (7) have convinced many that O2-generating cyanobacteria and aerobic eukaryotes evolved no later than ≈2.78 giga-annum before present (Ga) and perhaps as long ago as 3.7 Ga. Meanwhile, the developing record of mass-independent fractionation (MIF) of sulfur isotopes in the sedimentary record, as well as some other geochemical tracers, has been interpreted as supporting a protracted atmospheric oxygenation over the period of ≈2.5–2.2 Ga (8).

An early origin for oxygenic photosynthesis demands an explanation of how surface oxidation was muted for perhaps as long as 1,500 million years (My), until cyanobacteria finally surmounted some geochemical, environmental, or ecological barrier and successfully oxidized the planet. Perhaps this scenario is correct, and some abiotic change, such as the long-term escape of hydrogen to space (9), was the direct cause of planetary oxygenation. We suggest, however, that the data are also consistent with scenarios without oxygenic photosynthesis in the Archean. Herein we discuss an alternate hypothesis, one in which the evolution of cyanobacteria destroyed a methane greenhouse and thereby directly and rapidly triggered a planetary-scale glaciation, the ≈2.3–2.2 Ga Makganeyene “snowball Earth.”

Geological Setting

The earliest evidence for glaciation comes from the late Archean and early Proterozoic, which suggests Earth at this time experienced global temperatures not much different from those today. The oldest known midlatitude glaciation, recorded in the Pongola Supergroup diamictite, occurred at 2.9 Ga (10). The period from 2.45 Ga until some point before 2.22 Ga saw a series of three glaciations recorded in the Huronian Supergroup of Canada (11) (Fig. 1). The final glaciation in the Huronian, the Gowganda, is overlain by several kilometers of sediments in the Lorrain, Gordon Lake, and Bar River formations (Fms.). The entire sequence is penetrated by the 2.22 Ga Nipissing diabase (12); the Gowganda Fm. is therefore significantly older than 2.22 Ga.

Fig. 1.
Proposed correlation of the Huronian Supergroup and the upper Transvaal Supergroup. The three Huronian glacial units, penetrated and capped by the Nipissing diabase, predate the Makganyene diamictite in the Transvaal. The uppermost Huronian glacial unit, ...

In its eastern domain, the Transvaal Supergroup of South Africa contains two glacial diamictites, in the Duitschland and Boshoek Fms. The base of the Timeball Hill Fm., which underlies the Boshoek Fm., has a Re-Os date of 2,316 ± 7 My ago (13). The Boshoek Fm. correlates with the Makganyene diamictite in the western domain of the Transvaal Basin, the Griqualand West region. The Makganyene diamictite interfingers with the overlying Ongeluk flood basalts, which are correlative to the Hekpoort volcanics in the eastern domain and have a paleolatitude of 11° ± 5° (14). In its upper few meters, the Makganyene diamictite also contains basaltic andesite clasts, interpreted as being clasts of the Ongeluk volcanics. The low paleolatitude of the Ongeluk volcanics implies that the glaciation recorded in the Makganyene and Boshoek Fms. was planetary in extent: a snowball Earth event (15). Consistent with earlier whole-rock Pb–Pb measurements of the Ongeluk Fm. (16), the Hekpoort Fm. contains detrital zircons as young as 2,225 ± 3 My ago (17), an age nearly identical to that of the Nipissing diabase in the Huronian Supergroup. As the Makganyene glaciation begins some time after 2.32 Ga and ends at 2.22 Ga, the three Huronian glaciations predate the Makganyene snowball.

In contrast to the Makganyene Fm., the three Huronian diamictites are unconstrained in latitude. Poles from the Matachewan dyke swarm, at the base of the Huronian sequence, do indicate a latitude of ≈5.5° (18), but ≈2 km of sedimentary deposits separate the base of the Huronian from the first glacial unit (19), which makes it difficult to draw conclusions about the latitude of the glacial units based on these poles. Low latitude poles in the Lorrain Fm. (20, 21), which conformably overlies the Gowganda diamictite, are postdepositional overprints (22).

MIF of Sulfur

The recent discovery of MIF of S isotopes in Archean and early Proterozoic rocks has provided a major constraint on atmospheric O2. Large MIF of S (up to several hundred permil) is produced by photolysis of SO2 to S by light of wavelengths of <200 nm, which would have been unable to penetrate to the lower atmosphere had O2 levels been above a few percent of the present atmospheric level (23). To preserve MIF, multiple atmospheric S species must be maintained as partially isolated reservoirs rather than being homogenized by oxidation of reduced species like hydrogen sulfide and polysulfur, as would occur at greater than ≈10–5 present atmospheric level O2 (24). Mixing and dilution of the atmospherically fractionated component, both in the atmosphere and in the oceans, presumably yields the observed Archean MIF values of up to ≈10‰ (25). An active oceanic S cycle, as would exist at moderately high O2 levels, also would likely prevent the preservation of MIF.

Before the deposition of the ≈2.32 Ga Rooihogite and Timeball Hill Fms. in the Transvaal Supergroup (13, 26) and the McKim Fm. in the Huronian Supergroup, sedimentary sulfides often display MIF (8, 26, 27) over an order of magnitude larger than the largest values observed in modern sulfate aerosols (25, 28). One plausible interpretation of the diminished MIF observed in the Rooihogite, Timeball Hill, and McKim Fms. is a rise in atmospheric O2. At present, only glaciers record the MIF sometimes present in sulfate aerosols; in the marine environment, with riverine input composing ≈90% of sulfate input and aerosols composing ≈10%, MIF is not preserved (25). The small range of MIF could reflect an environment in which atmospheric chemistry began to approach modern conditions, decreasing the magnitude of MIF, but in which aerosols formed the major component of marine sulfate input, allowing its preservation.

The small range of MIF also permits an opposite interpretation. Rather than a decrease in atmospheric fractionation, the diminution could be a product of increased continental input and ocean/atmosphere mixing driven by glacial conditions. The period of 2.5–2.2 Ga was a time of glaciations. Enhanced glacial weathering could have made unfractionated S from igneous sources a larger part of the marine S budget, whereas sharper thermal gradients would have driven homogenization of the S pool. Both effects would have decreased sedimentary MIF. If this interpretation is correct, only the final decrease of MIF after ≈2.09 Ga may reflect the rise of atmospheric O2, and the oxygenation event may have been more rapid than the length of Farquhar and Wing's (8) MIF Stage II (2.45–2.09 Ga) indicates.

Appearance of Local O2

Both the Huronian and the Transvaal Supergroups contain features suggestive of, but not demanding, the first appearance of O2 in the period between the Gowganda and Makganyene glaciations. As noted in ref. 8, MIF in the Huronian rocks examined so far is at diminished values compatible with either increased atmospheric O2 or enhanced glacial weathering and ocean/atmosphere mixing. The Lorrain Fm. contains red beds, as well as a hematitic paleosol at Ville-Marie, Quebec, and the Bar River Fm. contains pseudomorphs after gypsum, consistent with an increase in sulfate levels from oxidative weathering of sulfide minerals on land. In the Transvaal Supergroup, the upper Timeball Hill Fm., like the Lorrain Fm., contains red beds.

Stronger evidence for O2 appears after the Makganyene glaciation. The Ongeluk Fm. is overlain by the Hotazel Fm., which consists predominantly of banded iron formation (BIF). The basal half meter of the Hotazel Fm. contains dropstones (29) (see Supporting Text, which is published as supporting information on the PNAS web site), which suggests it was deposited toward the end of the glacial period. The Hotazel Fm. hosts a massive Mn member (29): a blanket of Mn deposition unmatched by any other known in the world, ≈50 m thick and with an erosional remnant, the Kalahari Mn Field, measuring ≈11 × 50 km in extent (30).

The Kalahari Mn Field indicates the release of large quantities of O2 into the ocean and therefore suggests a highly active postglacial aerobic biosphere. In seawater at circumneutral pH, only NO3 and O2 can chemically oxidize soluble Mn2+ to produce insoluble Mn(IV) (31) (Table 1). Although a Mn2+ oxidizing phototroph also could produce Mn(IV), no such organism has ever been identified. Given the high redox potential of the Mn(IV)/Mn(II) couplet, it is likely that no such organism exists. The unexcited P870 reaction center in purple bacteria has a redox potential too low to accept electrons from Mn(II) (Table 1), but no known photosystem reaction center aside from the P670 reaction center of photosystem II has a higher redox potential. Photosynthetic reduction of Mn(IV) therefore would likely require a two-part photosystem akin to that involved in oxygenic photosynthesis.

Table 1.
Midpoint potentials of relevant redox couplets

One plausible interpretation of the sequence of events leading up to the Paleoproterozoic snowball Earth is shown in Fig. 1 and is as follows.

  1. Three glaciogenic units were deposited in the Huronian. The extent of the glaciations is not constrained, but they generally lack the lithographic features associated with snowball Earths, such as a sharp transition from a diamictite to a cap carbonate (32). Although the Espanola carbonate (33) could be a cap for the Bruce glaciation, without paleomagnetic data or additional lithographic features, its presence alone is insufficient to conclude that the Bruce glaciation was a snowball event.
  2. Some of the earliest continental red beds were deposited in the Firstbrook member of the Gowganda Fm. and in the Lorrain and Bar River Fms. in Canada, as well as in the upper Timeball Hill Fm. in South Africa. The basal Timeball Hill Fm. has recently been dated at 2,316 ± 7 My ago (13). In our proposed correlation, all of the red bed-bearing units were deposited after the last Huronian glaciation and before the Makganyene glaciation. The formation of the red beds could involve local O2, although it does not demand it (34). Syngenetic pyrite from the basal Timeball Hall Fm. shows only slight MIF of S (26), consistent with the initiation of planetary oxygenation or enhanced glacial activity.
  3. The low-latitude glaciation that formed the Makganyene and Boshoek diamictites (14) was initiated when the production of O2 triggered the collapse of a methane greenhouse. Although Pavlov et al. (35) proposed a similar trigger, they assumed a delay of at least ≈400 My between the onset of oxygenic photosynthesis and its surficial expression. We suggest a more immediate linkage.
  4. The Nipissing diabase intruded into the Huronian sequence (36), and the Ongeluk and Hekpoort volcanics were deposited (16).
  5. The Hotazel Fm., which includes BIF and Mn members, was deposited in oxygenated waters in the aftermath of the snowball.
  6. The upper siderite facies of the Hotazel Fm. and perhaps also part of the overlying Mooidrai dolomite were deposited as cap carbonates in the process of removing CO2 from a thick postsnowball greenhouse (29).

Timescale for Methane Greenhouse Collapse

For cyanobacteria to be directly responsible for triggering a planetary glaciation, they must have been able to produce enough O2 to destroy the methane greenhouse before the carbonate–silicate weathering cycle could compensate. On sufficiently long timescales, global cooling would slow silicate weathering and carbonate precipitation, thereby allowing CO2 to build up in the atmosphere (37). The response time of the carbonate–silicate weathering cycle is generally estimated at ≈1 My (38), although the time to replace the greenhouse capacity lost in a methane greenhouse collapse may be longer. To estimate the timescale for destruction of the methane greenhouse, we constructed a steady-state ocean biogeochemistry model based on the assumptions that biological productivity was controlled by P and N and that Fe(II) oxidation was the main inorganic O2 sink (see Supporting Text).

The critical P/Fe flux ratio for net oxidation of the surface ocean increases with the burial rate of C (see Fig. 2a, which is published as supporting information on the PNAS web site). For a P flux similar to today's value of ≈8 × 1010 mol/y (39) and a burial fraction of 2 × 10–2, similar to modern anoxic environments, the model results indicate that the critical value is ≈1/50. The current ratio of P flux to hydrothermal Fe flux is ≈1/2, whereas the ratio of P flux to hydrothermal and terrigenous Fe flux is ≈1/64 (1). Rare earth element patterns of BIFs (40) and the stability of terrestrial ferrous sulfides in an anoxic atmosphere suggest that the main source of reactive Fe in the Archean was hydrothermal, so the P/Fe ratio may have been closer to 1/2 than 1/64.

When the P/Fe ratio falls below the critical value, essentially all O2 produced is captured by Fe2+. Above the critical P/Fe value, there is a net release of O2 (Fig. 2a). Other dissolved O2 sinks, such as Mn2+, H2, and CH4, will consume some O2, but hydrothermal fluxes of these reductants are up to an order of magnitude less than the flux of Fe2+ (41). At today's levels, these are capable of consuming ≈7 × 1010 mol of O2 per y, equivalent to 4 × 10–10 bar/y. Although the flux of H2S from vent fluids in today's high sulfate oceans is comparable to that of Fe2+, in an ocean where Fe2+ is the main electron donor, H2S levels would be lower. A formaldehyde rainout flux of ≈1012 mol/y (24) would be the largest O2 sink, capable of adsorbing ≈6 × 10–9 bar of O2 per y, but would be nullified by H2O2 rainout once even a small amount of O2 accumulated in the atmosphere. O2 production in excess of ≈6 × 10–9 bar/y (1 bar = 100 kPa) would therefore be sufficient to initiate CH4 oxidation, and, once begun, net O2 production in excess of ≈4 × 10–10 bar/y would suffice to continue it. At the rates predicted by our model, destruction of a 1-mbar methane greenhouse, if it occurred at all, would likely occur within a few My, a timescale comparable with the carbonate–silicate weathering cycle. Therefore, either cyanobacteria did not evolve until shortly before the oxygenation event, or the nutrient flux did not reach sufficiently high levels at any point after the evolution of cyanobacteria until then.

If cyanobacteria were present during the Huronian glaciations, the increased P flux into the oceans generated by glacial weathering (42) should have triggered the oxygenation event. Instead, the oxygenation event seems to correlate with the later Makganyene glaciation; this finding suggests the evolution of cyanobacteria occurred in the interval between the Huronian glaciations and the Makganyene glaciation.

Whether N limitation could have delayed the destruction of a methane greenhouse depends on the Fe demand of the N2 fixers and the ability of cyanobacteria to capture Fe before it reacted with O2 and sank beneath the photic zone as a ferric precipitate. With anoxic deep waters providing a large source of Fe to the photic zone, life for an early Proterozoic diazotroph would have been easier than for modern diazotrophs that must subsist off of Fe transported from the continents. C/Fe ratios in N2-fixing populations of Trichodesmium range from <2,000 to 50,000 (43). If early cyanobacteria were inefficient at both capturing and using Fe (e.g., C/Fe = 2,000; capture efficiency = 1%), then N limitation could have protected the methane greenhouse, but under more optimistic assumptions (either a higher capture efficiency or a higher C/Fe ratio), it would not have done so (Fig. 2b). The geologic record, not computational models, must ultimately decide.

Implications of the Possible Late Evolution of Cyanobacteria

The interpretation presented here suggests that planetary oxygenation began in the interval between the end of the Huronian glaciations and the onset of the Makganyene glaciation and that the Paleoproterozoic snowball Earth was the direct result of a radical change in the biosphere. In the Archean and earliest Proterozoic oceans, life may have been fueled predominantly by Fe, with Fe(II) used as the electron donor for photosynthesis and Fe(III) as the main electron acceptor for respiration. The sediments therefore would be moderately oxidizing and the surface waters reduced (34). Because the redox potential for Mn2+ oxidation is much higher than that of Fe2+, Mn would have to be removed from the oceans in reduced form. The carbonates precipitated at this time contain up to ≈2% Mn (30), indicating that Mn2+ reached shallow waters and coprecipitated with Ca+2; oxidized Mn is extremely rare. The atmosphere was likely reducing. Astrophysical models predict the Sun was substantially dimmer than today, but a CH4 greenhouse (44) produced by methanogens living in a reduced upper ocean would have kept the planet warm enough to allow for the presence of liquid water without leading to massive siderite precipitation (45).

This world would have been overthrown when cyanobacteria capable of oxygenic photosynthesis evolved, which molecular phylogenies indicate occurred later than the main bacterial radiation (46). The surface waters became oxidizing and more productive. Reduced C accumulated in the sediments; methanogenesis moved from the surface ocean to the deep ocean, where it was isolated from the atmosphere. Methanotrophy in the ocean and photochemistry in the atmosphere used O2 to transform atmospheric CH4 to CO2, a less effective greenhouse gas. The rise of O2 thus might have triggered a glacial interval without the ≈400 My delay assumed by others (35).

Phosphate flux into the oceans correlates with increased continental weathering during glacial intervals (42), so increased continental weathering produced by the glaciations may have increased nutrient availability above the threshold for net O2 release into the atmosphere. Because of the relatively low global temperature, it would have taken only a trace of OH radicals from the oxygenic bloom produced when cyanobacteria appeared in a high-P ocean to damage the CH4 greenhouse enough to bring average global temperatures to <0°C. During the at least ≈35 My (29, 47) it took to build up a sufficient CO2 greenhouse to escape from the snowball, hydrothermal fluxes would have built up a rich supply of nutrients in the oceans. When the planet finally warmed, the oceans were ripe for a cyanobacterial bloom. The O2 produced by the bloom cleared out tens of My worth of accumulated reductants and thus produced the Kalahari Mn field (29).

Biomarker Counterevidence for Archean O2

Despite the parsimony of cyanobacterial evolution occurring within a few My before the onset of the Paleoproterozoic snowball Earth, some organic biomarker evidence and indirect sedimentological and geochemical arguments have been used to suggest that the origin of cyanobacteria dates to far earlier times: at least 2.78 Ga and maybe as long ago as 3.7 Ga.

The critical piece of evidence placing the origin of cyanobacteria and locally oxic environments in the Archean is the discovery in bitumens from rocks as old as 2.78 Ga of organic biomarkers apparently derived from lipids used by cyanobacteria and eukaryotes in their cell membranes. Although Brocks et al. (48) concluded that the bitumens were likely syngenetic, they could not exclude the possibility that they were postdepositional contaminants. Even if the biomarkers are as old as their host rocks, however, the uniformitarian extrapolation of modern lipid distributions to the Archean should be viewed cautiously.

Hopanes. Among modern organisms, 2-methyl-bacteriohopanepolyol is produced predominantly by cyanobacteria and in trace quantities by methylotrophs like Methylobacterium organophilum (3). Hopanes derived from 2-methyl-bacteriohopanepolyol can be preserved in sedimentary rocks, where they have been used as tracers for cyanobacteria (3). Hopanol synthesis has traditionally been assumed, based on the understood modern distribution of hopanols, to occur only in aerobic organisms. Fischer et al. (79) recently demonstrated, however, that Geobacter sulfurreducens can synthesize diverse hopanols, although not 2-methyl-hopanols, when grown under strictly anaerobic conditions. Thus, fossil hopanes do not necessarily imply the presence of O2-producing organisms, and nothing about 2-methyl-hopanols suggests that they are any different in this respect. Archean 2-methyl-hopanes also might have been produced by ancestral cyanobacteria that predated oxygenic photosynthesis.

Steranes. Produced by eukaryotes for use in the cell membrane, sterols are preserved in the rock record as steranes. Brocks et al. (4, 5) recovered steranes, along with 2-methyl-bacteriohopanepolyol, from 2.78 Ga Pilbara Craton shales. Because there is no known anaerobic sterol synthesis pathway, they used their discovery to argue for the presence of O2.

Cholesterol biosynthesis in modern organisms is a long biochemical pathway that employs the following four O2-dependent enzymes: (i) squalene epoxidase, (ii) lanosterol 14-α-methyl demethylase cytochrome P450, (iii) sterol-4-α-methyl oxidase, and (iv) lathosterol oxidase (49). These O2-dependent enzymes perform reactions that, although not currently known to occur biochemically in anaerobic organisms, could feasibly occur without O2. Moreover, the substitution of an O2-dependent enzyme for an anaerobic step in a biosynthetic pathway appears to be a common evolutionary occurrence. Raymond and Blankenship (50) found that, of the 473 O2-dependent enzymatically catalyzed reactions in the BioCyc database (www.biocyc.org), 20 have at least one O2-independent counterpart that performs the same reaction. For instance, AcsF catalyzes an O2-dependent cyclization step in the synthesis of chlorophyll and bacteriochlorophyll, a pathway that must have existed before the evolution of oxygenic photosynthesis. The O2-independent enzyme BchE performs the same reaction as AcsF but uses vitamin B12 in place of O2 (50). The assumption that sterol synthesis is always O2-dependent and always has been therefore merits close inspection.

Indeed, the Hamersley bitumens include their own cautionary message about the application of uniformitarian assumptions to fossil lipids. Dinosterane is generally accepted to be characteristic of dinoflagellates and is interpreted as a dinoflagellate biomarker in Phanerozoic rocks (51). Yet even though an Archean origin for dinoflagellates seems implausible, because it would indicate the Archean origin of modern eukaryotic group not known in the fossil record until at least the Paleozoic (52), Brocks et al. (5) found dinosterane. They interpreted the molecule as being produced by eukaryotes of unknown affinities, although an alternative explanation is that these modern-style putative Archean biomarkers are contaminants.

Indirect Counterevidence for Archean O2

Although organic biomarkers may be difficult to interpret, they are a significant improvement on several other geological tracers that have been used to argue for the presence of cyanobacteria and environmental O2, including microfossils, stromatolites, BIFs, and assorted isotopic fractionations.

Microfossils. Before 2 Ga, when diversified assemblages with affinities to major groups of cyanobacteria first appear in the fossil record, the microfossil record is murky (53). Some have interpreted filamentous forms in earlier rocks as cyanobacterial remains (5456), but Brasier et al. (57) recently questioned the biogenic nature of these objects. Moreover, cyanobacteria cannot be identified solely by a filamentous form. Many nonoxygenic bacteria are also filamentous, including some mat-forming green nonsulfur and purple sulfur bacteria (58, 59) and a methanogenic archeon (60). The wide variety of filamentous prokaryotes highlights a problem in identifying fossil microbes lacking clear evidence of cell differentiation based on morphology: Any given form has probably arisen many times in Earth history, both in extant and extinct organisms.

Stromatolites and BIFs. Two types of late Archean rock Fms. have often been interpreted as indicating cyanobacterial activity: stromatolites and BIFs. Des Marais (7) argued that large stromatolite reefs indicate the presence of cyanobacteria and therefore a locally aerobic environment, but large reefs also can form under anaerobic conditions. Populations of anaerobic methane oxidizers, for instance, have built massive reefs at methane seeps in the Black Sea (61). In addition, the Archean and Paleoproterozoic oceans were likely more supersaturated with respect to calcite and aragonite than the modern oceans (62), which would have facilitated the precipitation of large reefs even without biological participation. Indeed, abiotic processes may have played a major role in the formation of many Precambrian stromatolites (63). Moreover, although the deposition of ferric iron in BIFs has traditionally been taken to imply the presence of free O2 (40, 64, 65), BIFs also could have formed in a O2-free environment, either by photooxidation (66) or by Fe(II)-oxidizing phototrophic bacteria (67, 68).

Isotopic Evidence. Rosing and Frei (6) argued based on isotopic evidence that >3.7 Ga metashale from West Greenland preserves signs of an aerobic ecosystem. They found organic C with δ13C values of –25.6‰ in the same sediments as Pb with isotopic ratios indicating that the samples were originally enriched in U with respect to Th and interpreted this finding as reflecting an environment in which U could be cycled between its reduced, insoluble U(IV) form and its oxidized, soluble U(VI) form. They concluded the light C was produced by oxygenic phototrophs and that biogenic O2 had permitted the redox cycling of U. However, all biological C fixation pathways and some abiotic mechanisms can produce light C, and, just like the Fe(III) in BIFs, U(VI) can form in the absence of O2. At circumneutral pH values, the midpoint potential of the U(VI)/U(IV) couplet is ≈0 V, similar to the Fe(III)/Fe(II) couplet and considerably less oxidizing than Mn(IV)/Mn(II), NO3/N2, or O2/H2O (see Table 1).

A strong negative δ13C excursion in organic C at ≈2.7 Ga has been interpreted as evidence for the evolution of aerobic methanotrophy (69). Such light C suggests repeated fractionation: first in the production of CH4 subsequently oxidized to CO2, then in re-reduction by a primary producer; similar fractionations are observed today in environments with methanotrophs. But although CH4 oxidation was once thought to require O2, geochemical measurements (70) and recent microbiological work (71, 72) have demonstrated that CH4 oxidation also can occur with alternative electron acceptors, so O2 is not needed to explain the isotopic excursions. Although the anaerobic methane oxidizing bacteria studied today rely on sulfate, which is unlikely to have been available in a high-Fe Archean ocean, thermodynamics permits a variety of electron acceptors, including Fe(III), to be used for CH4 oxidation.

S isotope data indicate local spikes in sulfate concentration starting ≈3.45 Ga (73, 74). Canfield et al. (75) argued that these spikes were produced in high-O2 environments. But given the low redox potential of the sulfide/sulfate couplet, local sulfate spikes can be explained by scenarios that do not involve O2. Moreover, sedimentary sulfate deposits, which disappear in the rock record after ≈3.2 Ga (75), do not reappear until after the Huronian glaciations, which suggests that high sulfate conditions were rare.

None of these indirect lines of evidence necessitate oxygenic photosynthesis. The case for cyanobacteria and locally oxic environments existing before the disappearance of MIF of S isotopes and the massive deposition of Mn in the Kalahari Mn Field rests, at the moment, solely on steranes.

Future Directions

Our model for a Paleoproterozoic origin of cyanobacteria is testable by several methods. It suggests that sterols in Archean rocks, if they are original, were synthesized by anaerobic reactions. It therefore should be possible to find or engineer enzymes capable of synthesizing cholesterol under anaerobic, biochemically plausible conditions. In addition, a continuous biomarker record that stretches back from time periods when O2 is definitely present into the Archean might reveal transitions in community composition. Current work with samples from recent drilling programs targeting the late Archean and the Paleoproterozic has begun this task. An intensive search for biomarkers with definite relationships to metabolism, such as those derived from the pigment molecules of phototrophic bacteria, also would produce a more convincing record. A search for cyanobacterial or eukaryotic fossils that predate 2.0 Ga yet have affinities to modern groups would complement the geochemical approaches.

With respect to the record of MIF of S, the timeline we propose for a rapid oxygenation scenario suggests that decreased fractionation during the interval of the Huronian glaciations may be a byproduct of enhanced glacial weathering and ocean/atmosphere mixing. If this scenario is correct, a similar phenomenon should have occurred in association with the Pongola glaciation at ≈2.9 Ga. Investigation of the Huronian deposits where low-MIF S is found should reveal these deposits to be sedimentologically immature, reflecting glacial input. Additionally, the syngenicity of the MIF values should be tested through techniques such as the paleomagnetic search for significantly postdepositional formation of sulfides.

Finally, our model predicts that the Makganyene snowball Earth was a singular event. Convincing paleomagnetic evidence (including positive syn-sedimentary field tests) that demonstrated the Huronian glaciations were low-latitude would contradict our model and instead support a protracted episode of planetary oxygenation with multiple snowball events not directly triggered by a singular event, the evolution of cyanobacteria.


Because of the importance of the evolution of cyanobacteria and the planetary oxygenation event in Earth history, it is particularly useful to consider multiple working hypotheses about these events. We propose a model that takes a skeptical attitude toward the evidence for Archean cyanobacteria and a protracted early Proterozoic planetary oxygenation. In our alternative scenario, an evolutionary accident, the genesis of oxygenic photosynthesis, triggered one of the world's worst climate disasters, the Paleoproterozoic snowball Earth. Intensive investigation of the time period of the Paleoproterozoic glaciations may reveal whether a novel biological trait is capable of radically altering the world and nearly bringing an end to life on Earth.

Supplementary Material

Supporting Information:


We thank R. Adler, N. Beukes, R. Blankenship, J. Brocks, H. Dorland, A. Kappler, J. Kasting, A. Maloof, D. Newman, S. Ono, A. Sessions, D. Sumner, T. Raub, B. Weiss, and three anonymous reviewers for advice and discussion; R. Tada for help with fieldwork in the Huronian; A. Pretorius and Assmang Limited for access to Nchwaning Mine; and P. Hoffman for communicating this manuscript. This work was supported in part by the Agouron Institute and by a National Aeronautics and Space Administration Astrobiology Institute cooperative agreement with the University of Washington. R.E.K. was supported by a National Science Foundation Graduate Research Fellowship and a Moore Foundation Fellowship.


Author contributions: R.E.K. and J.L.K. designed research; R.E.K., I.A.H., and C.Z.N. performed research; R.E.K., J.L.K., and I.A.H. analyzed data; and R.E.K. and J.L.K. wrote the paper.

Abbreviations: Ga, giga-annum before present; My, million years; BIF, banded iron formation; Fm., formation; MIF, mass-independent fractionation.


Wing, B. A., Brabson, E., Farquhar, J., Kaufman, A. J., Rumble, D., III, & Bekker, A. (2002) Geochim. Cosmochim. Acta 68, A840 (abstr.).

Falkowski, P. G., Follows, M. & Fennel, K. (2003) EOS Trans. Am. Geophys. Union 84, Suppl., abstr. U52C-01.


1. Canfield, D. E. (1998) Nature 396 450–453.
2. Anbar, A. D. & Knoll, A. H. (2002) Science 297 1137–1142. [PubMed]
3. Summons, R. E., Jahnke, L. L., Hope, J. M. & Logan, G. A. (1999) Nature 400 554–557. [PubMed]
4. Brocks, J. J., Logan, G. A., Buick, R. & Summons, R. E. (1999) Science 285 1033–1036. [PubMed]
5. Brocks, J. J., Buick, R., Summons, R. E. & Logan, G. A. (2003) Geochim. Cosmochim. Acta 67 4321–4335.
6. Rosing, M. T. & Frei, R. (2004) Earth Planet. Sci. Lett. 217 237–244.
7. Des Marais, D. J. (2000) Science 289 1703–1705. [PubMed]
8. Farquhar, J. & Wing, B. A. (2003) Earth Planet. Sci. Lett. 213 1–13.
9. Catling, D. C., Zahnle, K. J. & McKay, C. P. (2001) Science 293 839–843. [PubMed]
10. Nhelko, N. (2004) D.Phil. thesis (Rand Afrikaans University, Johannesburg).
11. Young, G. M., von Brunn, V., Gold, D. J. C. & Minter, W. E. L. (1998) J. Geol. 106 523–538.
12. Corfu, F. & Andrews, A. J. (1986) Can. J. Earth Sci. 23 107–109.
13. Hannah, J. L., Bekker, A., Stein, H. J., Markey, R. J. & Holland, H. D. (2004) Earth Planet. Sci. Lett. 225 43–52.
14. Evans, D. A., Beukes, N. J. & Kirschvink, J. L. (1997) Nature 386 262–266.
15. Tajika, E. (2003) Earth Planet. Sci. Lett. 214 443–453.
16. Cornell, D. H., Schutte, S. S. & Eglington, B. L. (1996) Precambrian Res. 79 101–123.
17. Dorland, H. (2004) D.Phil. thesis (Rand Afrikaans University, Johannesburg).
18. Bates, M. P. & Halls, H. C. (1990) Can. J. Earth Sci. 27 200–211.
19. Young, G. M., Long, D. G. F., Fedo, C. M. & Nesbitt, H. W. (2001) Sediment. Geol. 141 233–254.
20. Williams, G. E. & Schmidt, P. W. (1997) Earth Planet. Sci. Lett. 153 157–169.
21. Schmidt, P. & Williams, G. (1999) Earth Planet. Sci. Lett. 172 273–285.
22. Hilburn, I. A., Kirschvink, J. L., Tajika, E., Tada, R., Hamano, Y. & Yamamoto, S. (2005) Earth Planet. Sci. Lett. 232 315–332.
23. Farquhar, J., Savarino, J., Airieau, S. & Thiemens, M. H. (2001) J. Geophys. Res. Planets 106 32829–32839.
24. Pavlov, A. A. & Kasting, J. F. (2002) Astrobiology 2 27–41. [PubMed]
25. Ono, S., Eigenbrode, J. L., Pavlov, A. A., Kharecha, P., Rumble, D., Kasting, J. F. & Freeman, K. H. (2003) Earth Planet. Sci. Lett. 213 15–30.
26. Bekker, A., Holland, H. D., Wang, P.-L., Rumble, D., III, Stein, H. J., Hannah, J. L., Coetzee, L. L. & Beukes, N. J. (2004) Nature 427 117–120. [PubMed]
27. Farquhar, J., Bao, H. & Thiemens, M. (2000) Science 289 756–758. [PubMed]
28. Savarino, J., Romero, A., Cole-Dai, J., Bekki, S. & Thiemens, M. H. (2003) Geophys. Res. Lett. 30 2131.
29. Kirschvink, J. L., Gaidos, E. J., Bertani, L. E., Beukes, N. J., Gutzmer, J., Maepa, L. N. & Steinberger, R. E. (2000) Proc. Natl. Acad. Sci. USA 97 1400–1405. [PMC free article] [PubMed]
30. Cairncross, B., Beukes, N. & Gutzmer, J. (1997) The Manganese Adventure: The South African Manganese Fields (Associated Ore & Metal Corp. Ltd., Johannesburg).
31. Gaidos, E. J., Nealson, K. H. & Kirschvink, J. L. (1999) Science 284 1631–1633. [PubMed]
32. Hoffman, P. F. & Schrag, D. P. (2002) Terra Nova 14 129–155.
33. Bernstein, L. & Young, G. M. (1990) Can. J. Earth Sci. 27 539–551.
34. Walker, J. C. G. (1987) Nature 329 710–712. [PubMed]
35. Pavlov, A. A., Brown, L. L. & Kasting, J. F. (2001) J. Geophys. Res. Planets 106 23267–23287.
36. Noble, S. R. & Lightfoot, P. C. (1992) Can. J. Earth Sci. 29 1424–1429.
37. Walker, J. C. G., Hays, P. B. & Kasting, J. F. (1981) J. Geophys. Res. Oceans Atmos. 86 9776–9782.
38. Berner, R. A., Lasaga, A. C. & Garrels, R. M. (1983) Am. J. Sci. 283 641–683.
39. Bjerrum, C. J. & Canfield, D. E. (2002) Nature 417 159–162. [PubMed]
40. Klein, C. & Beukes, N. J. (1992) in The Proterozoic Biosphere, ed. Klein, C. (Cambridge Univ. Press, Cambridge, U.K.), pp. 147–152.
41. Elderfield, H. & Schultz, A. (1996) Annu. Rev. Earth Planet. Sci. 24 191–224.
42. Föllmi, K. B. (1995) Geology 23 503–506.
43. Kustka, A., Carpenter, E. J. & Sanudo-Wilhelmy, S. A. (2002) Res. Microbiol. 153 255–262. [PubMed]
44. Pavlov, A. A., Kasting, J. F., Brown, L. L., Rages, K. A. & Freedman, R. (2000) J. Geophys. Res. Planets 105 11981–11990. [PubMed]
45. Rye, R., Kuo, P. H. & Holland, H. D. (1995) Nature 378 603–605. [PubMed]
46. Olsen, G. J., Woese, C. R. & Overbeek, R. (1994) J. Bacteriol. 176 1–6. [PMC free article] [PubMed]
47. Pierrehumbert, R. T. (2004) Nature 429 646–649. [PubMed]
48. Brocks, J. J., Buick, R., Logan, G. A. & Summons, R. E. (2003) Geochim. Cosmochim. Acta 67 4289–4319.
49. Risley, J. M. (2002) J. Chem. Ed. 79 377–384.
50. Raymond, J. & Blankenship, R. E. (2004) Geobiology 2 199–203.
51. Moldowan, J. M. & Jacobson, S. R. (2000) Int. Geol. Rev. 42 805–812.
52. Falkowski, P. G., Katz, M. E., Knoll, A. H., Quigg, A., Raven, J. A., Schofield, O. & Taylor, F. J. R. (2004) Science 305 354–360. [PubMed]
53. Amard, B. & Bertrand-Sarfati, J. (1997) Precambrian. Res. 81 197–221.
54. Awramik, S. M., Schopf, J. W. & Walter, M. R. (1983) Precambrian. Res. 20 357–374.
55. Schopf, J. W. & Packer, B. M. (1987) Science 237 70–73. [PubMed]
56. Schopf, J. W. (1993) Science 260 640–646. [PubMed]
57. Brasier, M. D., Green, O. R., Jephcoat, A. P., Kleppe, A. K., Van Kranendonk, M. J., Lindsay, J. F., Steele, A. & Grassineau, N. V. (2002) Nature 416 76–81. [PubMed]
58. Pierson, B. K. (1994) in Early Life on Earth, Nobel Symposium, ed. Bengtson, S. (Columbia Univ. Press, New York), No. 84, pp. 161–180.
59. D'Amelio, E. D., Cohen, Y. & Des Marais, D. J. (1987) Arch. Microbiol. 147 213–220. [PubMed]
60. Joulian, C., Patel, B. K. C., Ollivier, B., Garcia, J. L. & Roger, P. A. (2000) Int. J. Syst. Evol. Microbiol. 50 525–528. [PubMed]
61. Michaelis, W., Seifert, R., Nauhaus, K., Treude, T., Thiel, V., Blumenberg, M., Knittel, K., Gieseke, A., Peterknecht, K., Pape, T., et al. (2002) Science 297 1013–1015. [PubMed]
62. Grotzinger, J. P. & Kasting, J. F. (1993) J. Geol. 101 235–243. [PubMed]
63. Grotzinger, J. P. & Knoll, A. H. (1999) Annu. Rev. Earth Planet. Sci. 27 313–358. [PubMed]
64. Cloud, P. (1973) Econ. Geol. 68 1135–1143.
65. Holland, H. D. (1973) Econ. Geol. 68 1169–1172.
66. Cairns-Smith, A. G. (1978) Nature 276 807–808.
67. Widdel, F., Schnell, S., Heising, S., Ehrenreich, A., Assmus, B. & Schink, B. (1993) Nature 362 834–836.
68. Kappler, A. & Newman, D. K. (2004) Geochim. Cosmochim. Acta 68 1217–1226.
69. Hayes, J. M. (1994) in Early Life on Earth, ed. Bengtson, S. (Columbia Univ. Press, New York), Vol. 84, pp. 220–236.
70. Barnes, R. O. & Goldberg, E. D. (1976) Geology 4 297–300.
71. Boetius, A., Ravenschlag, K., Schubert, C. J., Rickert, D., Widdel, F., Gieseke, A., Amann, R., Jorgensen, B. B., Witte, U. & Pfannkuche, O. (2000) Nature 407 623–626. [PubMed]
72. Orphan, V. J., House, C. H., Hinrichs, K. U., McKeegan, K. D. & DeLong, E. F. (2002) Proc. Natl. Acad. Sci. USA 99 7663–7668. [PMC free article] [PubMed]
73. Canfield, D. E., Habicht, K. S. & Thamdrup, B. (2000) Science 288 658–661. [PubMed]
74. Shen, Y. A., Buick, R. & Canfield, D. E. (2001) Nature 410 77–81. [PubMed]
75. Huston, D. L. & Logan, G. A. (2004) Earth Planet. Sci. Lett. 220 41–55.
76. Drever, J. I. (1997) The Geochemistry of Natural Waters: Surface and Groundwater Environments (Prentice–Hall, Englewood Cliffs, NJ).
77. Langmuir, D. (1978) Geochim. Cosmochim. Acta 42 547–569.
78. White, D. (2000) The Physiology and Biochemistry of Prokaryotes (Oxford Univ. Press, New York).
79. Fischer, W. W., Summons, R. E. & Pearson, A. (2005) Geobiology 3 33–40.

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