A Key Role for Green Rust in the Precambrian Oceans and the Genesis of Iron Formations

ARTICLES PUBLISHED ONLINE: 16 JANUARY 2017 | DOI: 10.1038/NGEO2878

A key role for green rust in the Precambrian oceans and the genesis of iron formations

I. Halevy1*†, M. Alesker1†, E. M. Schuster1, R. Popovitz-Biro2 and Y. Feldman2

Iron formations deposited in marine settings during the Precambrian represent large sinks of iron and silica, and have been used to reconstruct environmental conditions at the time of their formation. However, the observed mineralogy in iron formations, which consists of iron oxides, silicates, carbonates and sulﬁdes, is generally thought to have arisen from diagenesis of one or more mineral precursors. Ferric iron hydroxides and ferrous carbonates and silicates have been identiﬁed as prime candidates. Here we investigate the potential role of green rust, a ferrous–ferric hydroxy salt, in the genesis of iron formations. Our laboratory experiments show that green rust readily forms in early seawater-analogue solutions, as predicted by thermodynamic calculations, and that it ages into minerals observed in iron formations. Dynamic models of the iron cycle further indicate that green rust would have precipitated near the iron redoxcline, and it is expected that when the green rust sank it transformed into stable phases within the water column and sediments. We suggest, therefore, that the precipitation and transformation of green rust was a key process in the iron cycle, and that the interaction of green rust with various elements should be included in any consideration of Precambrian biogeochemical cycles.

lthough the chemical, mineralogical and isotopic chemical environments, GR transforms into several of the major composition of iron formations (IFs) has been widely minerals observed in IFs, suggesting a role for GR as a precursor used to constrain the chemistry and oxidation state of the to IF mineral assemblages. A 1–3 Precambrian oceans and atmosphere , many aspects of their genesis remain poorly understood. It is agreed that even in the Thermodynamics and experiments suggest a GR precursor best preserved IFs, observed mineral assemblages reflect diagenetic Minerals observed in IFs are thermodynamically stable (Fig. 1a alteration of one or more mineral precursors1–5. Attention has and Methods), although barriers to nucleation and growth prevent focused on ferric iron (FeIII) oxyhydroxides, thought to have the direct precipitation of some (for example, haematite), or formed by the oxidation of ferrous iron (FeII) near the ocean require high degrees of supersaturation of others (for example, 3 surface, in ‘oases’ of O2-rich seawater generated by cyanobacteria , siderite). Instead, metastable phases (for example, the hydrous by anoxygenic FeII photosynthesis6, or by abiotic light-driven FeII FeIII oxyhydroxide, ferrihydrite) often precipitate from solution and oxidation7. In contrast, the texture of IFs from various ages supports transform into the stable minerals. Unless reductants (for example, primary precipitation of FeII-bearing, nanoparticulate silicates or organic carbon) are available for reduction of the precursors, iron carbonates from anoxic waters, and much later oxidation to ferric may persist in FeIII minerals despite anoxic conditions in the 8,9 4 oxides . The identity of the precursors remains uncertain, as do sediments . In O2-poor seawater-like solutions, thermodynamics the pathways for their transformation into the minerals observed predicts that GR should be a metastable precursor precipitate in IFs, and the dependence of the final assemblage on the presence (Fig. 1b), and that its transformation to the stable minerals should of organic matter. depend on solution pH and oxidation state, and on the aqueous A mineral neglected in the study of IFs is the ferrous– concentrations of dissolved inorganic carbon (DIC), total sulfur ferric hydroxy salt, green rust (GR). A naturally occurring GR, and total silica (Supplementary Fig. 1). Unlike FeIII oxyhydroxide II III x+[ n− × ]x− II fougèrite ([Fe1−x Fex Mgy (OH)2+2y ] x/nA mH2O , with A precursors, approximately two-thirds of the iron in GR is Fe , − − 2− 2− representing OH , Cl , SO4 ,CO3 ), forms today in water- so that organic carbon is not strictly required for its diagenetic logged iron-rich soils, and as a product of steel corrosion transformation into mixed-valence mineral assemblages. This may 10,11 II III in O2-poor environments . Following exposure to air GR explain the occurrence of Fe /Fe assemblages in some organic rapidly oxidizes to FeIII oxyhydroxides12,13. Recently, GR has been carbon-poor IFs17. found to form near the iron redoxcline and persist through To preclude kinetic inhibition of GR precipitation and the water column and sediments in ferruginous Lake Matano constrain its transformation products, we performed precipitation (Indonesia), an Archaean ocean analogue14. Additionally, GR has experiments from Archaean seawater-analogue solutions at been inadvertently formed in experiments conducted in Archaean- room temperature (∼22 ◦C), with geologically plausible pH analogue seawater15,16. Despite these indications that GR forms and concentrations of FeII, DIC, sulfate and silica (Methods). under conditions relevant to Earth’s early oceans, its potential role in Initial precipitates were analysed by X-ray diﬀractometry (XRD), IF genesis is unexplored. Combining thermodynamic calculations, scanning electron microscopy (SEM) and transmission electron laboratory experiments and dynamic models of the early iron microscopy (TEM), and reanalysed after ageing periods of cycle, we find that precipitation of GR was probably a major early several days to four months (at ∼22 and 50 ◦C). In all solutions sink of seawater iron. Furthermore, following exposure to diﬀerent with pH>∼ 7, irrespective of DIC, sulfate and silica content,

1Department of Earth and Planetary Sciences, Weizmann Institute of Science, Rehovot 76100, Israel. 2Department of Chemical Research Support, Weizmann Institute of Science, Rehovot 76100, Israel. †These authors contributed equally to this work. *e-mail: [email protected]

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a b FeOH2+

1.0 1.0 Fe3+ Fe3+ Schwertmannite

− Fe(OH)4 0.5 0.5

Fe2+ Fe2+

Eh (V) Haematite Eh (V) Ferrihydrite 0.0 0.0

+ FeHCO3 ∑Fe = 10−5 M ∑S = 10−5 M + ∑ −2.7 FeHCO3 −0.5 SiO2 = 10 M −0.5 2− −2 Siderite GRCO3 pCO = 10 atm 0 2 FeCO3 ° Greenalite − T = 25 C Fe(OH)3 0 2 4 6 8 10 12 14 0 2 4 6 8 10 12 14 pH pH

Figure 1 | Thermodynamic stability of iron-bearing minerals in ‘early seawater’. a, Thermodynamically stable, but kinetically inhibited phases. b, Metastable, apparently uninhibited phases, which may act as precursors to formation of the stable phases. The grey region represents the range of Eh and pH encountered in modern marine environments. The Eh of the early, anoxic ocean was lower.

2− CO3 -bearing GR precipitated initially. Ferrihydrite formed at As the values of the required model parameters are highly lower pH. Depending on the solution composition, the GR aged to uncertain, we first explore the sensitivity to these values, and then iron oxyhydroxides, siderite or iron-bearing silicates (Fig.2). The provide a best-guess temporal evolution of the parameter values silicates, which were too scarce or poorly crystalline to detect by to assess the importance of GR over Earth history. Improving XRD, were identified by TEM electron diffraction (ED) patterns constraints on early CO2 and O2 levels and ocean chemistry will consistent with the structures of modulated 2:1 sheet silicates (for reduce uncertainty in such model histories. Due to the strong example, stilpnomelane, minnesotaite). Energy-dispersive X-ray dependence of GR saturation on pH (Fig. 3a), the results are highly spectroscopy coupled to the ED measurements indicated that the sensitive to pCO2 , through its effect on ocean pH (Fig. 3b). Increasing II iron was associated with the sheet silicates. Sample results are in the pO2 leads to decreasing fGR due to more rapid oxidation of Fe Supplementary Information. (Fig. 3c). Higher hydrothermal iron influxes result in a higher FeII concentration in the deep ocean, and in lower fGR (Fig. 3d), if A dynamic model predicts a major GR sink for marine iron siderite precipitates at saturation (see discussion below). Within a II The experiments demonstrate that GR precipitates from Archaean reasonable range of values, fGR is insensitive to rates of Fe photo- and Proterozoic ocean-analogue solutions, and that it ages to oxidation (Supplementary Fig. 3). minerals observed in IFs in the absence of organic matter. However, We explored the effects of assumptions about the sink of the importance of GR as a precursor to IF mineral assemblages iron to primary FeII carbonates (siderite) or silicates (greenalite). depends on the dynamics of the early iron cycle. For example, if the Siderite precipitation in modern anoxic environments typically rate of FeII oxidation is rapid enough that FeIII generation outstrips requires up to ∼50 times supersaturation, possibly due to surface- its precipitation in GR, then FeIII concentrations increase until the controlled kinetics19. A requirement for 10 times siderite saturation II next least soluble phase (ferrihydrite) saturates and precipitates. (sid = 10) results in deep-ocean Fe concentrations controlled Below the iron redoxcline, FeII and DIC or silica concentrations by upwelling of FeII-bearing water, near-surface oxidation of the may be high enough that precipitation of FeII-bearing carbonates upwelled FeII, and precipitation of GR and/or FeIII oxyhydroxides or silicates is a dominant iron sink8,9,18. Thus, the proportion of as the only iron sinks (Fig.3 solid). When no supersaturation GR in the flux to the sediments depends on the relative rates is required (sid = 1, for example, if precipitation surface pre- of FeII supply, FeII oxidation (FeIII production), and precipitation exists) FeII concentrations are co-controlled by deep-ocean siderite and sinking of GR, FeIII oxyhydroxides, and FeII minerals, as precipitation and oxidation of upwelled FeII. Accordingly, the iron affected by the concentrations of FeII, FeIII, DIC and silica, and by sink comprises siderite, GR and/or FeIII oxyhydroxides (Fig.3 ocean pH. dotted). Irrespective of the required siderite supersaturation and To quantify the relative contribution of GR to iron sedimentation the resulting FeII concentrations, and using an upper limit on −2.7 fluxes (fGR), we developed a dynamic model of the early iron silica activity of 10 and an experimentally determined apparent 18 cycle (Methods). Model inputs are pCO2 , pO2 , the hydrothermal and greenalite solubility , greenalite remained undersaturated. riverine influxes, ocean mixing rates, and various thermodynamic With estimated time-dependent parameter values and kinetic constants. The model tracks the aqueous concentrations (Supplementary Information), GR constitutes a large fraction of FeII and FeIII in the surface, intermediate and deep ocean, the of the iron flux to the sediment during the Archaean (Fig. 3e), saturation state of the various precipitates in these environments, irrespective of uncertainty in GR thermochemical properties and their proportion in the sedimentary iron flux. On the basis of the (Supplementary Fig. 4). Before the rise of O2, a hydrothermal experimental results shown here for GR and in other studies for FeIII source of FeII ∼ 2–3 times larger than present is balanced by oxyhydroxides, we assume that the primary precipitates transform oxidation of FeII in the photic zone and formation of GR particles, into the minerals observed in IFs, but the transformation is not as well as by deep-ocean siderite precipitation, if sid =1. Resulting explicitly modelled. FeII concentrations are ∼1–10 µM and <1 nM in the deep and

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a b 200 nm

(030) (300) (110) (120) (210)

(330) (120) (110) (300) (030) NaCl Green rust

5 nm−1 Green rust

GR f NaCl GR GR NaCl NaCl GR GR GR NaCl GR NaCl

10 20 30 40 50 60 70 Minnesotaite α θ ° CuK 2 ( ) (006) (131) (003) S (131) c (110)

Siderite + magnetite (131) (003) (131) (006) 500 nm 2 nm−1 M M g S S S Stilpnomelane M S S S M M S M

20 30 40 50 60 CuKα 2θ (°)

d e −1 50 nm 2 nm 200 nm Siderite

2 nm−1 Magnetite

(220) ED11 (620) (440)

(660) (642) 200 nm

Figure 2 | Green rust and its ageing products. a,b, TEM image and XRD pattern (a), and indexed ED pattern of green rust (b). c–e, XRD pattern of siderite (S) and magnetite (M) (c), SEM image of siderite (d), and TEM image and ED pattern of magnetite (e), which are some of the ageing products in the absence of silica. f,g, TEM image and ED patterns of Fe phyllosilicates, which are some of the ageing products in the presence of silica at room ◦ temperature (f) and at 50 C(g).

III III surface ocean, respectively. Fe concentrations are everywhere The rise of atmospheric O2 results in dominance of Fe <0.01 nM, too low for saturation and precipitation of FeIII oxyhydroxides in the sedimentary iron flux (Fig. 3e), due partly to oxyhydroxides. Realistically, heterogeneity is expected to have more rapid oxidation of FeII, and partly to an assumed decrease in III existed in the relative importance of GR and Fe oxyhydroxides, the concentrations of reduced greenhouse gases (for example, CH4, for example, due to spatio-temporal variability in phototrophic C2H6), which causes an increase in pCO2 via the silicate weathering activity and in deep water upwelling rates. Similarly, variability climate feedback20, and a concomitant decrease in pH21. As solar in seawater pH and DIC concentrations (for example, due to luminosity increases into the Proterozoic, pCO2 decreases and pH generation of alkalinity and DIC during anaerobic remineralization increases, but rapid FeII oxidation prevents GR from contributing of organic matter), or in silica concentrations (for example, due to to the near-surface iron sink. However, FeIII oxyhydroxides formed variable distance from hydrothermal silica sources), would lead to in the photic zone may be transformed into GR, either biotically22 or heterogeneity in the relative importance of primary FeII carbonate abiotically23, when they sink below the iron redoxcline into FeII-rich and silicate precipitation relative to GR or FeIII oxyhydroxide waters. The sedimentary flux of iron-bearing particles may have precipitation. thus contained an appreciable fraction of GR even if photic zone

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a c 1.0 1.0 0.8 0.8 Ω sid: 10 1 0.6 Siderite 0.6 Ferrihydrite 0.4 Green rust 0.4 Fe sink fraction Fe Fe sink fraction Fe 0.2 0.2 0.0 0.0

6.4 6.6 6.8 7.0 7.2 7.4 7.6 7.8 10−5 10−4 10−3 10−2 pH p (atm) O2 b d 1.0 1.0 0.8 0.8 0.6 0.6 0.4 0.4 Fe sink fraction Fe Fe sink fraction Fe 0.2 0.2 0.0 0.0 10−2 10−1 1 2 3 4 5 p (atm) f (× modern) CO2 hyd

e Rise of O2 1.0

0.8

0.6 Proterozoic Archaean 0.4 Fe sink fraction Fe

0.2

0.0 1.5 2.0 2.5 3.0 3.5 Age (Ga)

Figure 3 | Sensitivity of iron sink to model parameters. a–e, Fraction of precipitates out of the iron sink for variable seawater pH (a), atmospheric pCO2 (b), atmospheric pO2 (c), the hydrothermal circulation enhancement factor (d) and time (e). In a,c,d, parameters other than the varied parameter were at default values (Supplementary Information). In b, pH depended on pCO2 , and other parameters were at default values. In e, time-dependent estimates of all parameters were used (Supplementary Information). Results depend on siderite supersaturation required for precipitation (sid =10 or 1 in solid or dotted, respectively). Sensitivity to the parameters is explained in the Supplementary Information. dynamics favoured formation of FeIII oxyhydroxides. We expect that metals are bioessential nutrients, whereas others have been used to GR continued to shuttle iron to the sediments until the second, constrain the oxidation state of the early oceans and atmosphere. Neoproterozoic, rise of atmospheric O2, when oxygenation of much The mobility and chemistry of these elements, and perhaps others, of the ocean irreversibly decreased seawater FeII concentrations. probably changed radically as the importance of the GR iron shuttle varied with time. Experimental and theoretical work is Implications and tests of a GR iron shuttle required to fully explore the implications for metal cycles, biological A quantitatively important GR shuttle from the surface ocean to productivity and evolution, and for the use of transition metals as the sediments affects the cycles of several other elements. Silica proxies for early ocean chemistry. sorbs strongly to GR and is known to retard GR oxidation24. How might a GR-derived mineral assemblage manifest in the Adsorption of silica onto sinking GR particles may have coupled geologic record? The relatively short timescale of GR transformation the transport of iron and silica to sediments, leading to enrichment makes it difficult to texturally distinguish primary precipitation of of both elements in IFs. Phosphate sorbs strongly to GR, but FeII carbonates or silicates8,9 from their early formation from a GR unlike FeIII oxyhydroxides, which release the phosphate following precursor in the water column or at the sediment/water interface. reductive dissolution, the phosphate sorbed to GR may get Furthermore, transformation by dissolution and reprecipitation sequestered in the FeII-phosphate vivianite25. Did the waning obliterates many of the microtextures associated with GR importance of GR at the Archaean–Proterozoic transition lead precipitation. The low density of GR relative to other iron-bearing to more efficient phosphate recycling, to a positive feedback on precipitates (Supplementary Table 5), however, suggests density productivity, organic matter burial and O2 release, and to the gradient-driven sedimentary textures as possible indicators of a protracted Lomagundi positive C isotope excursion26? Nickel sorbs GR precursor. The occurrence of centimetre-scale flame structures onto GR more effectively than onto FeIII oxyhydroxides14, with from magnetite bands into overlying chert bands in a Mesoarchaean proposed implications for the evolution of methanogens, which IF in the Red Lake Greenstone Belt (northwest Ontario) may be an require Ni. Reductive immobilization of U and Cr by GR and example30. Finally, the different affinity and reactivity of GR towards isomorphic substitution of divalent transition metals (Zn, Cd, Ni, a variety of elements, as well as the occurrence of both FeII and FeIII Co) for FeII in the GR structure are known23,27–29. Some of these in GR, may yield distinctive chemical and isotopic fingerprints,

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oxidation of Fe(OH)2, their products of oxidation and reduction of ferric Additional information Supplementary information is available in the online version of the paper. Reprints and oxyhydroxides: Eh-pH Pourbaix diagrams. C. R. Geosci. 338, 433–446 (2006). 14. Zegeye, A. et al. Green rust formation controls nutrient availability in a permissions information is available online at www.nature.com/reprints. Correspondence and requests for materials should be addressed to I.H. ferruginous water column. Geology 40, 599–602 (2012). 15. Wiesli, R. A., Beard, B. L. & Johnson, C. M. Experimental determination and Fe isotope fractionation between Fe(II), siderite and ‘‘green rust’’ in abiotic Competing ﬁnancial interests systems. Chem. Geol. 211, 343–362 (2004). The authors declare no competing financial interests.

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Methods used (that is, when the equilibrium constant for both for sulfate and carbonate GR Thermodynamics. Thermodynamic calculations were carried out using the is calculated using the same chemical potentials). In all of these cases, carbonate Geochemist’s Workbench31. As a basis, we used the thermochemical data in the rather than sulfate GR forms. LLNL database32, which contains a variety of relevant iron-bearing silicates (for example, greenalite, stilpnomelane), carbonates (for example, siderite) and sulfides Experimental synthesis and diagenesis of GR. GR synthesis and ageing was (for example, pyrite). We used a version of the database to which various iron performed in seawater-analogue solutions (see below) with variable 33 sulfate and iron chloride salts were added . We added to the database several concentrations of dissolved inorganic carbon (DIC, ∼25–75 mM) and total SiO2 ferrous–ferric hydroxides, including four varieties of GR (carbonate, sulfate, (0, 100–250 µM). All syntheses were conducted in 1-l glass bottles, at room chloride, hydroxide). We initially took compiled thermochemical properties for the temperature (∼22 ◦C). Bottles were sealed with caps with PTFE-coated silicon hydroxides (including the GR)34, and in preliminary calculations found that only liners, except during purging or diffusive exchange with the atmosphere, when caps carbonate and sulfate GR came close to saturation under the conditions of interest. with a gas inlet and outlet were used. The inlet is fused to a ceramic sinter, which Then, to explore the effects of uncertainty in the thermochemical properties of was lowered into the solution for purging, or raised above the solution during carbonate and sulfate GR, which are less well established than most iron-bearing diffusive exchange with the atmosphere. All solutions were magnetically stirred minerals, we compiled their values from literature sources and tested the sensitivity during preparation. of the thermodynamic calculations to the values of the equilibrium constants used. Doubly distilled water (18.2 m) was purged with N2 gas (99.99%) for 2 h to In calculations showing the thermodynamically stable phases, we excluded remove dissolved O2. The bottles were then sealed and transferred to an anaerobic magnetite and pyrite, which are generally accepted to be of diagenetic origin in glovebox (O2 <1 ppm). Inside the glovebox, salts were added to simulate 1 IFs . We also excluded various high-temperature/pressure minerals (for example, Precambrian seawater (Supplementary Table 1). Stable, O2-inert salts were ferrosilite, fayalite). Inclusion of magnetite and pyrite in calculations of the introduced first (NaCl, MgCl2× 6H2O, MgSO4× 7H2O), and the solution was thermodynamically stable phases results in stability fields of these minerals below purged with argon gas (99.996%) for 1 h to remove residual O2 brought into the the haematite field, but does not affect calculations of the metastable precursors, glovebox with the distilled water or with the salts. FeCl2× 4H2O, pre-purified to from which magnetite and pyrite were excluded. In calculations showing the remove oxidation products43, was then added and the solution purged with argon possible metastable precursors, we excluded thermodynamically stable minerals for an additional 0.5 h. The ceramic sinter was then raised above the liquid surface,

(magnetite, pyrite, haematite, siderite, greenalite) and several metastable minerals and NaHCO3 (99.7%, ACS) and Na2SiO3× 5H2O were added to simulate waters thought to be of diagenetic or metamorphic origin (for example, minnesotaite). with variable concentrations of DIC (24.2, 48.3 or 72.5 mM), and total SiO2 (0, 100, All calculations were carried out at a temperature of 25 ◦C, a pressure of 1 atm 150, 200 or 250 µM), respectively. Solution pH was adjusted to 7–8 by addition of (1.013 bar) and a total Fe concentration of 10−5 M. Phase diagrams in the main text NaOH (99.99%, ACS). We note that when the resulting solutions were left to stand, (Fig.1) were calculated with parameter values aimed at simulating an Archaean no precipitate formed over the course of several hours. Argon was flowed −2 −5 ocean. CO2 fugacity was 10 atm (ref. 35), total S concentration was 10 M through the headspace for 5 min and the bottles were then taken out of −2.7 (ref. 36), and total SiO2 concentration was 10 M (ref. 37). We tested the effect of the glovebox. uncertainty in the values of these parameters on the expected stable and precursor Partial oxidation of the FeII to FeIII was achieved by diffusive replacement of the mineral assemblage (Supplementary Fig. 1). For the stable mineral assemblage, an argon headspace by air. The bottles were exposed to air for ∼1.5 h, without increase in pCO2 and in total SiO2 concentrations leads to growth of the stability agitation or mixing, at which time the solution obtained a pale greenish-blue fields of siderite and greenalite, respectively, at the expense of haematite colour, but remained clear. The bottles were then returned to the glovebox and (Supplementary Fig. 1a,e). An increase in total S has no effect (Supplementary sealed. Precipitation of GR was complete within a few hours, but the first analysis Fig. 1c), as pyrite precipitation was disallowed (see above). was performed the following day to allow crystallinity to develop. Identical bottles As mentioned above, we tested the sensitivity also to uncertainty in the were left to age in the glovebox, wrapped in aluminium foil to prevent thermochemical properties of carbonate and sulfate GR. Drissi et al.38 used Eh/pH photochemical reactions in the solution and solid. All synthesis products were aged ◦ measurements during Fe(OH)2 oxidation to carbonate GR to calculate the at room temperature (∼22 C), whereas additional products from repeats of the 2− ◦ chemical potential of carbonate GR, given the chemical potentials of CO3 ,H2O SiO2-bearing experiments were also aged at 50 C. The pH of the ageing solutions −1 and Fe(OH)2. They reported a mean value of −852,870 cal mol for anhydrous was not buffered or measured, but the persistence of GR (Supplementary Table 2) carbonate GR, but no estimate of the error. Using 2σ uncertainty bounds on indicates that pH did not drop below ∼7. their Eh/pH measurements, and the chemical potentials reported in their study, we calculated the equilibrium constant for the carbonate GR Characterization of solids. Solids synthesized in the experiments were separated dissolution reaction, from solutions by centrifuge. Their morphology and elemental composition were determined by scanning electron microscopy (SEM, Leo Supra 55VP with EDS + + + − detector, Oxford INCA), and transmission electron microscopy (TEM, Philips FeIIFeIII(OH) (CO )+13H =4Fe2 +2Fe3 +HCO +12H O 4 2 12 3 3 2 CM-120 operating at 120 kV with EDS detector, EDAX-Phoenix Microanalyzer). Mineralogy was determined by X-ray diffraction (XRD, Rigaku TTRAX III and The equilibrium constant is logK =39.1±0.4 kJ mol−1. The uncertainty of Bruker D2 Phaser, using Cu Kα radiation), and by electron diffraction (ED) in 0.4 kJ mol−1 on the Gibbs free energy of reaction comes from the 2σ uncertainty the TEM. bounds on the Eh/pH measurements. This leads to negligible uncertainty in the For TEM analysis, samples were diluted by suspension in de-aerated stability field of carbonate GR (Supplementary Fig. 2a), and has no discernible ethanol/H2O and mounted on carbon/collodion-coated 400 mesh-sized Cu grids. bearing on the outcome of the thermodynamic calculations and dynamic model in For SEM analysis, samples were mounted on Al, carbon tape-coated SEM stubs (no this study. Much larger uncertainty comes not from the Eh/pH measurements, but sputter coating). For XRD analysis in the Bruker D2 Phaser, samples were mounted from the choice of chemical potentials of the species participating in the Fe(OH)2 on airtight sample holders, also by dripping a suspension of the solid in de-aerated oxidation reaction and in GR formation. Using, instead of the potentials in ethanol/H2O. For analysis in the Rigaku TTRAX III, samples were mixed with Drissi et al.38, the values recommended in the CRC Handbook of Chemistry and glycerol to prevent oxidation and mounted on sample holders. All sample Physics39 yields logK =33.2±0.4 kJ mol−1, significantly expanding the stability preparation was done in the anaerobic glovebox, after which samples were field of carbonate GR at the expense of ferrihydrite and aqueous FeII species transferred in an underpressurized desiccator to the analytical facilities. (Supplementary Fig. 2c). Using the more updated chemical potentials for Fe2+ and XRD patterns were automatically compared with the International Center for Fe3+ from Parker and Khodakovskii40 and the rest of the chemical potentials from Diffraction Data (ICDD) database. Lattice parameters (d-spacing) derived from Drissi et al.38 yields logK =45.5±0.4 kJ mol−1, significantly diminishing the ED patterns were manually compared with d-spacings from the ICDD database stability field of carbonate GR (Supplementary Fig. 2b). However, we can rule out and from the RRUFF database44. The average oxidation state of iron in the samples this high value of logK on the basis of our precipitation experiments, which yielded was not determined. The results of the analyses are shown in Supplementary GR at pH ∼8. This would not be possible if the equilibrium constant were as large Table 2. as 1045.5. We similarly explored the sensitivity to the thermochemical properties of Dynamic iron cycle model. The conceptual model is shown in Supplementary sulfate GR. Refait and Génin41 reported a chemical potential of anhydrous sulfate Fig. 5, and the parameter values used are listed in Supplementary Tables 3 and 4. GR of −904,100 ± 350 cal mol−1, and Olowe et al.42 reported a value of Parameter choices for a time-dependent calculation over the Archaean and part of −902,890 cal mol−1. Using the chemical potentials in these studies, logK = 29.4 to the Proterozoic are discussed in the Supplementary Information. 30.1. Using the chemical potentials in the CRC Handbook of Chemistry and The ocean is divided into three boxes representing the surface (0–100 m), Physics, logK = 23.5 to 24.1. Using the chemical potentials for Fe2+ and Fe3+ from intermediate (100–1,000 m), and deep (1,000–4,000 m) ocean. The volume of each Parker and Khodakovskii40 and the rest of the chemical potentials from Refait and box is calculated under the assumption that oceans cover 85% of Earth’s surface, Génin41, logK = 35.8 to 36.4. Given the greater stability of carbonate GR, this large which accounts for a smaller land area in the Archaean and early Proterozoic. uncertainty does not affect the results when consistent equilibrium constants are Results are insensitive to the choice of fractional land area.

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The mass budgets of FeII and FeIII in the surface ocean are: factor between 0.1 and 100 in the sensitivity analysis. The value of this scaling factor in the time-dependent simulations is discussed in the section describing d II = × II − + × II − these simulations. Fe s JRiv Fe Riv kph kO2 Fe s Js−i dt Mineral precipitation fluxes JGR, JFH and JSid are related to the degree × II − II − × − of saturation: Fe s Fe i 4 JGR JSid (C −C ) = sat × d J fmin III II τ Fe = kph +kO × Fe −Js−i dt s 2 s where C is the concentration (FeII or FeIII), C is the concentration at saturation of × III − III − × − sat Fe s Fe i 2 JGR JFH the mineral of interest, τ is a timescale for return to saturation (chosen to be 1 h),

and fmin is a unitless factor, which increases sigmoidally from zero when C

flux of water out of near-axis hydrothermal vents, JHyd, was taken to be ∼2 times logk0 =21.56−1,545/T the present-day flux of 3×1013 l yr−1 (ref. 45), divided by the volume of the deep ocean box. According to the dependence of the hydrothermal flux on the heat flux where I is the ionic strength (0.7, similar to modern seawater), [OH−] is the out of the oceanic crust through time51, this is equivalent to the hydrothermal water II hydroxide concentration (in M), [O2] is the concentration of dissolved oxygen flux ∼2,800 million years ago. The concentration of Fe in hydrothermal fluids, = × II (in M), [O2] KH pO2 , and Fe Hyd, was taken to be 80 mM, as indicated by thermodynamics of sulfate-poor hydrothermal solutions52. − × 1 − 1 3 1,500 ( T 298.15 ) KH =1.3×10 ×e We did not consider the fate of the mineral particles after their precipitation from seawater (that is, sinking, deposition, and early diagenesis), as we aim to −1 is Henry’s law constant for O2 (M atm ). We account for microaerophilic assess the fraction of iron exiting the ocean via a GR precursor. The exact timing of microbial oxidation of FeII, which can be many times faster than abiotic oxidation, transformation into the thermodynamically stable minerals, whether still in the using information from culture experiments48. The total FeII oxidation rate water column or in the sediments, is a separate and important problem, which was constant (abiotic + microbial) is: not addressed here. On the basis of the experiments conducted in this study for GR, and in other studies for FeIII oxyhydroxides, we assume that these primary kab = O2 precipitates transform into the stable minerals observed in IFs on relatively short kO2 1−fbio timescales, under the appropriate chemical conditions (for example, presence of organic matter and FeIII reducers for the transformation of ferrihydrite into II II where fbio if the fraction of microbial Fe oxidation out of the total oxidation as a Fe -bearing or mixed-valence minerals). function of pO2 (in atm). At the upper boundary of the experimental results: Overall in the ocean at the model steady state, the hydrothermal and riverine II source of Fe is balanced by oxidation in the surface ocean (by O2 and light-driven − − × 3 3.446 log[pO2 ] II III fbio =1.074×10 ×e processes), and by precipitation of Fe -bearing minerals (GR or siderite). The Fe source (FeII oxidation) is balanced by precipitation of FeIII-bearing minerals (FeIII

We cap fbio at 0.95, higher than the maximal experimental value (0.88). This choice oxyhydroxides or GR). and the fit to the upper (rather than lower) boundary of the experimental results, Except for the pH sensitivity analysis, in which ocean pH was prescribed maximizes fbio and, therefore, the total oxidation rate constant. This conservative independently, ocean pH was calculated from charge balance (for all other choice leads to underestimation of the importance of GR as a function of calculations), given pCO2 , and the concentration of conservative ions in seawater: atmospheric pO2 . The base rate constant for photo-oxidation of FeII was derived from the results + − − 2− [H ]−[OH ]−[HCO3 ]−2×[CO3 ]+... of published photo-oxidation experiments49, in which a photo-oxidation flux of II −2 −1 200 mg Fe cm yr for a modern solar flux and down to a depth of 100 m in a + + 2+ 2+ 2+ − ...+[Na ]+[K ]+2×[Mg ]+2×[Ca ]+2×[Fe ]−[Cl ] water column with 100 µM FeII was calculated. This translates into a rate constant ∼1.1×10−7 s−1. As light-driven oxidation occurs only in the surface ocean box, −[ −]− ×[ 2−]= this rate constant must be normalized by the depth of the surface ocean, Br 2 SO4 0 which is here identical to that in the above study, making for a normalization factor + + 2+ − − of unity. For lack of better constraints, the concentrations of Na ,K , Mg , Cl and Br Anoxygenic FeII photosynthesis can be many times more rapid than abiotic were taken to be identical to the present-day ocean (469, 10.2, 52.8, 546 and oxidation6. Conversely, an iron redoxcline located several metres to tens of metres 0.84 mM, respectively). The concentration of Ca2+ was calculated using the II below the ocean surface can result in less radiation available for light-driven Fe prescribed value of pCO2 and assuming CaCO3 saturation. The concentration of free 2+ − + oxidation, both biological and abiotic. To account for possibly higher and lower Fe was taken from the dynamic model. Substituting [OH ] = Kw/[H ], [ −]= × × + [ 2−]= × × × + 2 rates of photo-oxidation, we multiplied the base rate constant by a unitless scaling HCO3 pCO2 KH KA1/[H ], CO3 pCO2 KH KA1 KA2/[H ] ,

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2+ = × + 2 × × × [Ca ] KSP [H ] /(pCO2 KH KA1 KA2) into the above charge balance 41. Refait, Ph. & Génin, J. M. R. The transformation of chloride-containing green + − 2− yields a quartic equation for [H ], which has only one real, positive root. rust into sulphated green rust two by oxidation in mixed Cl and SO4 aqueous media. Corros. Sci. 36, 55–65 (1994). Code availability. The code used to calculate the fractions of green rust, 42. Olowe, A. A. & Génin, J. M. R. in Corrosion Science and Engineering: ferrihydrite and siderite in the flux of iron to the sediments is available on request. Proceedings of an International Symposium in honour of Marcel Pourbaix’s 85th Birthday Vol. 158 (eds Rapp, R. A., Gocken, N. A. & Pourbaix, A.) RT 297 Data availability. Representative electron microscope images, electron diﬀraction (Rapports Techniques CEBELCOR, 1989). patterns and X-ray diﬀraction patterns used to identify the primary precipitates 43. Gayer, K. H. & Wootner, L. The hydrolysis of ferrous chloride at 25◦. J. Am. and ageing products are available in the Supplementary Information. Chem. Soc. 78, 3944–3946 (1956). 44. Downs, R. T. The RRUFF Project: An Integrated Study of the Chemistry, References Crystallography, Raman and Infrared Spectroscopy of Minerals 31. Bethke, C. Geochemical and Biogeochemical Reaction Modeling (Cambridge (19th General Meeting of the International Mineralogical Press, 2008). Association, 2006). 32. Johnson, J. W., Oelkers, E. H. & Hegelson, H. C. SUPCRT92: a software 45. Elderfield, H. & Schultz, A. Mid-ocean ridge hydrothermal fluxes and the package for calculating the standard molal thermodynamic properties of chemical composition of the ocean. Annu. Rev. Earth Planet. Sci. 24, minerals, gases, aqueous species, and reactions from 1 to 5000 bar and 0 to 191–224 (1996). 1000 ◦C. Comput. Geosci. 18, 899–947 (1992). 46. Ganachaud, A. & Wunsch, C. Improved estimates of global ocean circulation, 33. Tosca, N. J. et al. Geochemical modeling of evaporation processes on Mars: heat transport and mixing from hydrographic data. Nature 408, insight from the sedimentary record at Meridiani Planum. Earth Planet. Sci. 453–457 (2000). Lett. 240, 122–148 (2005). 47. Millero, F. J., Sotolongo, S. & Izaguirre, M. The oxidation kinetics of Fe(II) in 34. Rickard, D. & Luther, G. W. III. Chemistry of iron sulfides. Chem. Rev. 107, seawater. Geochim. Cosmochim. Acta 51, 793–801 (1987). 514–562 (2007). 48. Druschel, G. K., Emerson, D., Sutka, R., Suchecki, P. & Luther, G. W. III

35. von Paris, P. et al. Warming the early Earth—CO2 reconsidered. Planet. Space Low-oxygen and chemical kinetic constraints on the geochemical niche of Sci. 56, 1244–1259 (2008). neutrophilic iron(II) oxidizing microorganisms. Geochim. Cosmochim. Acta 72, 36. Crowe, S. A. et al. Sulfate was a trace constituent of Archean seawater. Science 3358–3370 (2008). 346, 735–739. 49. Anbar, A. D. & Holland, H. D. The photochemistry of manganese and the 37. Siever, R. The silica cycle in the Precambrian. Geochim. Cosmochim. Acta 65, origin of banded iron formations. Geochim. Cosmochim. Acta 56, 3265–3273 (1992). 2595–2603 (1992). 38. Drissi, S. H., Refait, Ph., Abdelmoula, M. & Génin, J. M. R. The preparation 50. Hotinski, R. M., Kump, L. R. & Najjar, R. G. Opening Pandora’s Box: the and thermodynamic properties of Fe(II)-Fe(III) hydroxide-carbonate (green impact of open system modeling on interpretations of anoxia. rust 1); Pourbaix diagram of iron in carbonate-containing aqueous media. Paleoceanography 15, 267–279 (2000). Corros. Sci. 37, 2015–2041 (1995). 51. Sleep, N. H. & Zahnle, K. Carbon dioxide cycling and implications for climate 39. Haynes, W. M. CRC Handbook of Chemistry and Physics on ancient Earth. J. Geophys. Res. 106, 1373–1399 (2001). 97 edn (CRC Press, 2016). 52. Kump, L. R. Jr & Seyfried, W. E. Hydrothermal Fe fluxes during the 40. Parker, V. B. & Khodakovskii, I. L. Thermodynamic properties of the aqueous Precambrian: eﬀect of low oceanic sulfate concentrations and low hydrostatic ions (2+ and 3+) of iron and the key compounds of iron. J. Phys. Chem. Ref. pressure on the composition of black smokers. Earth Planet. Sci. Lett. 235, Data 24, 1699–1745 (1995). 654–662 (2005).

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