SCIENCE ADVANCES | RESEARCH ARTICLE GEOCHEMISTRY Copyright © 2019 The Authors, some Photoferrotrophy, deposition of banded iron rights reserved; exclusive licensee formations, and methane production in Archean oceans American Association for the Advancement of Science. No claim to Katharine J. Thompson1, Paul A. Kenward1, Kohen W. Bauer1, Tyler Warchola2, Tina Gauger3, original U.S. Government Raul Martinez4, Rachel L. Simister1, Céline C. Michiels1, Marc Llirós5, Christopher T. Reinhard6, Works. Distributed Andreas Kappler3, Kurt O. Konhauser2, Sean A. Crowe1* under a Creative Commons Attribution Banded iron formation (BIF) deposition was the likely result of oxidation of ferrous iron in seawater by either NonCommercial oxygenic photosynthesis or iron-dependent anoxygenic photosynthesis—photoferrotrophy. BIF deposition, License 4.0 (CC BY-NC). however, remains enigmatic because the photosynthetic biomass produced during iron oxidation is con- spicuously absent from BIFs. We have addressed this enigma through experiments with photosynthetic bacteria and modeling of biogeochemical cycling in the Archean oceans. Our experiments reveal that, in the presence of silica, photoferrotroph cell surfaces repel iron (oxyhydr)oxides. In silica-rich Precambrian seawater, this repulsion would separate biomass from ferric iron and would lead to large-scale deposition of BIFs lean in organic matter. Excess biomass not deposited with BIF would have deposited in coastal sediments, formed organic-rich shales, and fueled microbial methanogenesis. As a result, the deposition of BIFs by photoferrotrophs would have contributed fluxes of methane to the atmosphere and thus helped to stabilize Earth’s climate under a dim early Sun. INTRODUCTION photosynthesis end up in the atmosphere as methane and the Banded iron formations (BIFs) host the world’s largest iron ore oxidized products are buried as Fe(III) in BIF, can, as we show here, deposits, and they formed predominantly through the deposition of be quantitatively and mechanistically linked to the activity of photo- ferric iron [Fe(III)] from ferruginous oceans during the Archean ferrotrophs in ferruginous Archean oceans. Eon (1, 2). Most models for BIF deposition invoke photosynthesis Observations and models of extant photoferrotrophs demonstrate in the oxidation of ferrous iron from seawater to induce its sub- their capacity to both oxidize Fe(II) at rates sufficient to form even sequent precipitation and deposition as mixed valence iron (oxyhydr) the largest BIFs and support appreciable rates of primary produc- oxides and carbonate phases (3–6) [see (7) and (8) for alternative tion (5, 13, 20, 21). Laboratory experiments to date, paradoxically, models]. BIFs thus likely record the activity of Earth’s early photo- reveal a tight physical association between photosynthetic ferric synthetic biosphere (3). Two modes of photosynthesis have been iron metabolic by-products [e.g., Fe(III) (oxyhydr)oxides] and cel- implicated in Fe(II) oxidation—canonical oxygenic photosynthesis lular biomass that leads to their co-sedimentation (5, 20, 22). In the by the ancestors of modern cyanobacteria (3) and iron-dependent oceans, the deposition of Fe(III), along with this photosynthetic anoxygenic photosynthesis (4, 5), referred to as photoferrotrophy. biomass, would fuel sedimentary respiration that couples oxidation of Photoferrotrophic bacteria grow using light and Fe(II) to fix CO2 organic carbon to microbial Fe(III) mineral reduction and converts into biomass and produce Fe(III) as a metabolic by-product (5, 6, 9)– Fe(III) to dissolved Fe(II) and secondary minerals such as siderite they can do so in the complete absence of oxygen (9). Current liter- and magnetite (22). This sedimentary respiration thus closes the ature suggests that this anoxygenic photosynthesis is likely the evo- iron redox cycle and precludes the preservation of ferric iron in lutionary predecessor of oxygenic photosynthesis (10–12) and that BIF. Current observations and models therefore cannot explain BIF Fe(II) oxidation could thus have been driven by photoferrotrophy deposition through photoferrotrophy in light of this co-sedimentation before the emergence and proliferation of oxygenic photosynthesis of biomass with Fe(III) and ensuing diagenetic reactions (22, 23). (13, 14). Models of nutrient cycling in the Archean ocean implicate By contrast, a lack of Fe(III) reduction would lead to the co-deposition photoferrotrophs as key primary producers (13, 15) before the rise of ferric iron in BIF, with organic matter at relative concentrations of atmospheric oxygen 2.4 to 2.3 billion years (Ga) ago during the close to the 4:1 stoichiometry of Fe(II) oxidation to C fixation in Great Oxidation Event (GOE) (16–18). High methane concentrations photoferrotrophy [e.g., ~2 weight % (wt %) C based on 37 wt % in the Archean atmosphere have also been qualitatively attributed Fe(III) in Dales Gorge BIF (24)]. A compilation of organic matter to C and Fe cycling associated with BIF deposition (19). Such an concentrations and Fe redox states in BIFs, however, reveals that “upside-down” biosphere (19), in which the reduced products of they have very low organic matter concentrations (mean of 0.27 wt %; Fig. 1 and table S5) (24, 25) and contain appreciable Fe(III), with an 1 Departments of Microbiology and Immunology and Earth, Ocean, and Atmospheric average Fe redox state of 2.6 (fig. S1 and table S5) (24, 25). If the iron Sciences, University of British Columbia, Vancouver, British Columbia, Canada. 2Depart- minerals in BIFs were the product of photosynthesis, then the ment of Earth and Atmospheric Sciences, University of Alberta, Edmonton, Alberta T6G 2E3, Canada. 3Center for Applied Geosciences, University of Tuebingen, 72076 fate of the corresponding photosynthetic biomass remains entirely Tuebingen, Germany. 4Institut für Geo- und Umweltnaturwissenschaften, Albert-­ enigmatic. Furthermore, regardless of the mode of Fe(II) oxidation, Ludwigs-Universität, Mineralogie-Geochemie, 79104 Freiburg, Germany. 5Depart- current estimates for the magnitude of hydrothermal Fe(II) fluxes ment of Genetics and Microbiology, BioSciences Faculty, Universitat Autònoma de to the Archean ocean appear to be deficient of the mass fluxes needed Barcelona, Catalunya, Spain. 6School of Earth and Atmospheric Sciences, Georgia Institute of Technology, Atlanta, GA, USA. to sustain BIF deposition (26) and may imply an important role for *Corresponding author. Email: sean.crowe@ubc.ca terrestrial weathering fluxes (27, 28). We thus combined experiments Thompson et al., Sci. Adv. 2019; 5 : eaav2869 27 November 2019 1 of 9 SCIENCE ADVANCES | RESEARCH ARTICLE 10 in media with reduced phosphate concentrations (3 to 6 M), ap- Organic carbon (weight %) proaching those of modern seawater but still somewhat higher than the submicromolar concentrations found in Precambrian seawater 1 (15), show that 50% of strain KB01 cells are associated with Fe(III) (oxyhydr)oxides, leading to co-sedimentation of biomass and Fe(III) 0.1 (Fig. 2A). While phosphate concentrations were very low in the sea- water from which BIFs deposited (15, 33), silica concentrations in Precambrian oceans were high (~1 mM) (15), and this would have 0.01 altered the physical and chemical properties of Fe(III) precipitates Sample type formed in seawater (33). In experiments with 1 mM silica, the con- BIF centration implied for Precambrian seawater (15), strain KB01 cells Shale and other Precambrian rock types did not associate to, or co-sediment with, their Fe(III) oxyhydroxide Modern marine sediments by-­products (Fig. 2A). Instead, almost all of the cells remained sus- Modern OMZ sediments pended (94 ± 6%). Likewise, experiments conducted with strain Theoretical 4:1 Fe:C KoFox demonstrated that cell-mineral association was diminished Fig. 1. The organic matter concentrations in BIFs, other Precambrian sedi- relative to the silica-free experiments, with 72 ± 7% of strain KoFox mentary rocks, typical modern marine sediments, and oxygen minimum zone cells remaining suspended (Fig. 2A and fig. S5A). We also grew both (OMZ) sediments. The solid lines represent the group means, while the dotted lines delineate one SD above and below the mean. The black dashed line that strains KB01 and KoFox in their standard growth medium containing spans the figure represents the theoretical organic carbon concentration that 4 mM phosphate and no silica, which led to little cell-mineral asso- would be expected in BIFs deposited by photoferrotrophs (2.04 wt % C), assuming ciation for strain KB01 (Fig. 2A) and modest association for strain quantitative co-sedimentation of biomass with ferric Fe. References for this figure KoFox (fig. S5A). Detailed electron microscopy revealed that the can be found in table S5. surfaces of KB01 cells were entirely free from Fe(III) oxyhydroxide precipitates (Fig. 3), whereas strain KoFox tended to form multicellular using modern photosynthetic bacteria with revised theoretical aggregates that variably associated with the mineral precipitates estimates of Fe(II) fluxes to the Archean ocean and box modeling of (fig. S3A). Strain KB01, and to a somewhat lesser extent strain KoFox, coupled carbon and iron cycling to show that photosynthetic Fe(II) thus avoids encrustation with Fe in low-P, high-Si waters, as well as in oxidation could have sustained large-scale, organic-poor BIF deposi- standard growth media with high phosphate, remaining suspended tion in ocean upwelling systems over hundreds of millions of years. despite Fe(III) oxyhydroxide precipitation. We also subjected Synechococcus cells to Fe-Si–rich growth con- ditions in a similar fashion (Fig. 2A and fig. S5A). Synechococcus RESULTS AND DISCUSSION associated to a greater extent with the Fe(III) oxyhydroxide products Separation of biomass and Fe(III) of Fe(II) oxidation than did the photoferrotrophs under these con- To explore the fate of biomass and Fe(III) during photosynthetic ditions. Such an association was expected, despite the presence of silica, Fe(II) oxidation, we conducted experiments using Chlorobium given that oxygen effuses from Synechococcus cells, reacts rapidly phaeoferrooxidans strain KB01–a photoferrotrophic bacterium that with Fe(II), and causes precipitation of Fe(III) (oxyhydr)oxides on was isolated from the water column of ferruginous Kabuno Bay, a the cell surfaces—an association likely maintained due to binding sub-basin of Lake Kivu in East Africa (21). We also conducted with organic ligands, as observed in many previous environmental experiments with Chlorobium ferrooxidans strain KoFox, a sediment-­ and laboratory studies (32, 34, 35). Photoferrotrophs adapted to pelagic dwelling photoferrotroph that tends to associate with Fe(III) (oxyhydr) lifestyles thus appear most capable of avoiding co-sedimentation with oxides without becoming encrusted (23, 29), as well as a model Fe(III) and remain buoyant with the potential for separation of marine oxyphototrophic cyanobacterium of the genus Synechococcus. biomass from Fe(III) at larger scales in the environment. Strain KB01 is the only known photoferrotroph from a pelagic environment (21), and pelagic photoferrotrophs, including those Mechanisms of cell-mineral separation that would have populated Precambrian ferruginous oceans, face a Surface charge often influences cell-mineral association, or lack special challenge—they must maintain their position within sunlit thereof, and as solution P:Fe or Si:Fe ratios increased, the surface waters despite the precipitation of heavy Fe(III) oxyhydroxide meta- charge on the Fe(III) (oxyhydr)oxides formed became more nega- bolic by-products that have a tendency to adhere to cell surfaces and tive (Fig. 2B) due to the incorporation of P or Si anions into the cause their rapid sedimentation. Most bacteria have negatively charged oxyhydroxide structure (33). Conversely, Fe(III) oxyhydroxide cell surfaces at neutral pH (30), while Fe(III) (oxyhydr)oxides typi- particles were positively charged under low P or Si conditions, cally have a positive surface charge at the same pH (31). This gener- which is consistent with observations of positively charged Fe(III) ally leads to a strong attraction between cell surfaces and Fe(III) precipitates in many environments (31). Strain KB01 cells had (oxyhydr)oxides that often manifests as encrustation of bacterial strongly negative surface charges (fig. S4), whereas cells of strain cells and formation of Fe(III) oxyhydroxide cell aggregates (6, 32). KoFox had near neutral surface charges (fig. S4), similar to some micro­ To test the association between photoferrotrophs and their Fe(III) aerophillic Fe(II) oxidizers (36). This difference in surface charge metabolic by-products, we grew strain KB01 until late exponential between strains KB01 and KoFox reflects different cell surface growth phase, gently inverted the cell-mineral suspension to resuspend chemistries (table S2), notably an abundance of anionic surface sedimented Fe(III) (oxyhydr)oxides and cells, allowed the heavy functional groups on strain KoFox with different acid-base behavior Fe(III) (oxyhydr)oxides to resettle, and determined the percentage than strain KB01 (Supplementary Materials and fig. S3). Benthic of cells that remained suspended (Fig. 2A). Experiments conducted microorganisms, such as strain KoFox, commonly produce surface Thompson et al., Sci. Adv. 2019; 5 : eaav2869 27 November 2019 2 of 9 SCIENCE ADVANCES | RESEARCH ARTICLE 20 Interact. ener. (kT) 100 30 0.6 15 P:Fe 10 20 Si:Fe P:Fe 0.5 5 Si:Fe Interaction energy (kT) Zeta potential (mV) 75 0 10 0.4 −5 Percentage 0 1 2 3 4 5 Distance (nm) 0 0.3 50 −10 0.2 AB: KB- LW: KB- EL: KB- −20 0.1 TOT: KB- 25 −30 0.0 0 −40 −0.1 0 1 2 3 4 5 6 7 8 3.0 3.5 4.0 4.5 5.0 :C :D :A SW C D 1 2 KB :B :C Si:Fe and P:Fe ratio : 2: KB KB Distance (nm) KB KB Cy KF A B C Fig. 2. Cell surface characteristics for strain KB01 and the relationship between ferric iron surface charge and medium anions. Fractions of planktonic (blue) versus sedimented (red) cells (A) for photoferrotrophic strain KB01 under varying geochemical conditions: 400 M Fe(II) with low P (3 M) [unmarked]; 400 M Fe(II), low P (3 M), with 0.6 mM Si [A]; 400 M Fe(II), low P (3 M), with 1.0 mM Si [B]; 400 M Fe(II), low P (3 M), with 1.5 mM Si [C]; and 10 mM Fe(II) with 4.4 mM P [D]. C. ferrooxidans and Synechococcus are also shown under the [C] conditions. Rhodobacter strain SW2 is shown under the [D] conditions. The zeta potential (B), in millivolts, of the ferric iron precipitates is depicted with an increasing ratio of Si or P to Fe, where the error bars are all within the data points. Last, for the extended DVLO modeling for strain KB01 (C), the main graph depicts the interaction energies of the three forces (AB, Lewis acid base; LW, Lifshitz–van der Waals; EL, electrostatic) and the total (TOT) for those forces from 3 to 5 nm, while the inset depicts the forces from 0 to 5 nm. 100 nm that develop between these surfaces (Supplementary Materials) (36, 38). DVLO modeling uses measurements of physical surface properties to calculate interfacial forces as a function of distance between surfaces. Our model results revealed that the negative surface charge on strain KB01 led to electrostatic repulsion between its surface and negatively charged Fe(III) (oxyhydr)oxides (Fig. 2C). The more neutral charge on strain KoFox, conversely, led to a weak electro- static attraction (fig. S5C). Lewis acid-base and Lifshitz–van der 100 nm Waals forces were much weaker than the electrostatic forces, but we note that the acid-base properties of strain KB01 lead to repulsion A B of Fe(III) (oxyhydr)oxides, even without electrostatic effects. Surface Fig. 3. Scanning electron microscopy and transmission electron microscopy charge, therefore, controls cell-mineral association, and ultimately image of strain KB01. Scanning electron photomicrograph (A) and transmission like-charged surfaces cause the physical separation of strain KB01 electron photomicrograph (B) of C. phaeoferrooxidans strain KB01 revealing various and, by extension, other pelagic photoferrotrophs from their Fe(III) internal and external cell structures as well as a lack of encrustation (A and B) and a oxyhydroxide by-products. These findings have important implica- rare association with Fe(III) precipitates (A). tions for BIF deposition and the coupling of carbon and iron cycles in the Archean Eon. We note that while strain KB01 is a derived layers rich in anionic functional groups to facilitate attachment to member of the phylum Chlorobi, its cell surface chemistry is typical solid substrates (37), and these anionic surface functional groups for Gram-negative bacteria (30, 39), and our results are thus likely tend to strongly bind Fe(III), effectively neutralizing surface charge extensible to its ancestors and most other Gram-negative bacteria, (fig. S4). The role of Fe(III) in controlling surface charge on strain with the exception of benthic organisms, like strain KoFox, that have KoFox was confirmed by rinsing KoFox cells with reducing agents unusual modifications to their cell surfaces. Stem group, pelagic (sodium dithionite) that liberated Fe(III) and caused a shift in the photoferrotrophs in Archean oceans, therefore, most likely had cell surface charge to more negative values (fig. S4). These same rinses surface chemistries and interacted with Fe(III) (oxyhydr)oxides much had little effect on cells of KB01 (fig. S4), confirming the role of like strain KB01. surface chemistry and Fe(III) binding in controlling association of cells to Fe(III) (oxyhydr)oxides. The direct role of cell surface Revised Precambrian Fe budgets charges in dictating mineral association and co-sedimentation is Before proceeding to model coupled carbon and iron cycling, we also evident from the strong inverse relationship (fig. S5B) between sought to reconcile the possible material fluxes of Fe(II) to the mineral surface charges and cell-mineral separation. oceans and those needed to sustain BIF deposition (26, 40). These To assess the biophysical mechanisms that control cell-mineral material fluxes serve as boundary conditions for our models and associations, we determined interfacial properties of the cell and help tether our results to the geologic record. Peak BIF deposition Fe(III) oxyhydroxide surfaces and conducted extended Derjaguin-­ rates imply an Fe burial flux of 45 mol m−2 year−1 over areas greater than Landau-Verwey-Overbeek (DVLO) modeling to quantify the forces 1011 m2, or ~4.5 Tmol Fe year−1 (5), and at steady state, Fe(II) must Thompson et al., Sci. Adv. 2019; 5 : eaav2869 27 November 2019 3 of 9 SCIENCE ADVANCES | RESEARCH ARTICLE be actively resupplied to the oceans at this rate through a com- Archean oceans, taking into consideration the low oxygen atmo- bination of hydrothermal venting and continental and seafloor sphere (45) and smaller continent sizes (43), resulting in a total weathering. Previous estimates of modern hydrothermal Fe fluxes Fe(II) weathering flux of 5 Tmol year−1 to the oceans (Supplementary of ~2 Tmol year−1 (Supplementary Materials), on their own, are in- Materials). Summing continental weathering and hydrothermal Fe(II) sufficient to support steady-state BIF deposition (table S3) (1, 26, 40). fluxes yields an upper possible global Fe(II) flux of ~45 Tmol year−1 These estimates were based on the product of hydrothermal fluid to the oceans 2.5 Ga (table S3). This is more than sufficient to flow and Fe(II) concentrations measured in circulating on- and deposit BIF at 4.5 Tmol year−1 and is supplied mostly through pre- off-axis fluids (41). Global Fe fluxes, however, would have been viously unconsidered off-axis hydrothermal venting. These revised much different in the Archean Eon because of enhanced hydrothermal Fe(II) fluxes place upper boundaries on the magnitude of global activity (40, 42), smaller continents (43), reduced seawater sulfate coupled carbon and iron cycling, which we explore below using (42, 44), weathering of Earth’s crust at low oxygen (18, 45, 46), and biogeochemical box models. pervasively anoxic oceans. Hydrothermal fluid fluxes to the oceans have been calculated Modeling Archean marine iron and carbon cycles from a variety of geochemical and physical data resulting in a wide We assume that biological production in our model is driven by range of estimates for these fluxes in the literature. A recent compi- photoferrotrophy, and while oxygenic photosynthesis could have been lation from a number of sources (47), however, provides a synthesis active from the Mesoarchean (46, 50), its contribution to primary of both on- and off-axis fluid flow as well as a modeled distribution production may have been small before the GOE due to competi- of these values that yields averages (modes) of 5 × 1013 and 1.5 × tion with photoferrotrophs (15, 51). Nevertheless, discriminating 1016 kg year−1, respectively (47). These values are greater than between modes of photosynthesis would have no real effect on our values considered in previous estimates of modern hydrothermal model outcomes because oxygen produced would react with Fe(II), Fe(II) budgets, implying that modern Fe(II) fluxes to the oceans yielding the same 4:1 stoichiometry between iron oxidation and were underestimated (41). Archean hydrothermal fluid flow was, carbon fixation as in photoferrotrophy. Biomass degradation in our furthermore, likely higher than today due to greater heat loss from model was first channeled through heterotrophic Fe(III) reduction the Archean lithosphere (26, 48). Following previous work (26), we and then, given the near absence of sulfate in the Archean ocean thus scaled the updated estimates for modern fluid flow by the ratio (44), microbial methanogenesis (22, 44). We implicitly assumed of past to modern lithospheric heat loss (Supplementary Materials). that both Fe(III) reduction and methanogenesis are preceded by the This yields estimates for hydrothermal fluid flow at 2.5 Ga of 1 ± 0.6 breakdown of organic matter through hydrolysis, glycolysis, and × 1014 and 4 ± 0.6 × 1016 kg year−1 for on- and off-axis venting, respec- fermentation. Rates of organic matter breakdown were parameterized tively (table S3). Crucially, these revised estimates for fluid flow increase on the basis of observations from modern anoxic marine basins the possible Fe(II) fluxes to both the modern and Archean oceans. (52, 53). However, we note that, unlike today, the biological pump The concentration of Fe(II) in modern on-axis hydrothermal fluids in the Archean oceans would have operated without ballasting from is ~6 mmol kg−1 (41), but this depends on the chemistry of seawater fecal pellets. Biological production and nutrient cycling in the modern that circulates through seafloor basalts. Notably, high concentrations oceans can be divided into three broad oceanographic provinces— of sulfate (28 mM) in modern seawater lead to production of hydro- open ocean, coastal zone, and upwelling regions (52). In detail, the gen sulfide in anoxic hydrothermal fluids, and this hydrogen sulfide distribution and biological activity of these provinces depend on reacts with Fe(II) to form iron-sulfide minerals that limit the con- ocean circulation patterns, continental configurations, and nutrient centration of Fe(II) in effluxing fluids (42). Sulfate concentrations supply, but in the absence of robust constraints on these parameters in Archean seawater were much lower than today due, in part, to for the Archean oceans, we assumed that the relative contributions the lack of oxidative pyrite weathering on the continents (44). Models of similar such provinces to biological production scale with conti- of hydrothermal fluid chemistry, therefore, predict that Fe(II) con- nental area and were otherwise similar to the modern. We thus dis- centrations in high-temperature hydrothermal fluids were 10-fold tributed total biological production across these three provinces higher when seawater was sulfate free versus the modern 28 mM according to their relative productivities in the modern ocean (52) (42), implying that high-temperature vent fluids in the Archean and scaled continental area from 0 to 100% of the modern (Fig. 4 Eon could have had up to 60 mmol kg−1 Fe(II). Combining such and Supplementary Materials). a high-temperature hydrothermal Fe(II) concentration (42) with Our results revealed that upwelling provinces support Fe(II) revised on-axis hydrothermal fluid flows results in Fe(II) fluxes of oxidation and Fe(III) oxyhydroxide sedimentation at rates sufficient 8 ± 3 Tmol year−1 at 2.5 Ga (table S3), and these are similar to, but to deposit even the largest BIFs, such as those of the Hamersley Basin smaller than, other recent estimates (49). The effect of seawater sulfate in Western Australia (Fig. 5, Table 1, and Supplementary Materials). concentrations on the Fe(II) concentration in lower-temperature, Notably, Fe(III) (oxyhydr)oxides, with a relatively high density off-axis hydrothermal fluids appears less pronounced, and so we (3.8 g cm−1) and a tendency to aggregate to larger particle sizes (54), conservatively consider off-axis hydrothermal Fe(II) concentrations had an average settling velocity of 2 × 104 m year−1. This led to their equivalent to today (~0.75 mmol kg−1) (42). This results in possible deposition within a maximum of 40 km from their locus of initial off-axis Fe(II) fluxes of up to 30 ± 5 Tmol year−1 (table S3). Com- precipitation, assuming a 150-m-deep water column and current bining these new estimates for on- and off-axis Fe(II) fluxes, we velocities less than 650 m year−1 (fig. S6 and table S4) (52, 55). Further- redefine upper possible limits on hydrothermal Fe(II) delivery to more, this localized deposition occurred despite the strong horizontal the oceans at 2.5 Ga as ~40 Tmol year−1 (table S3). current velocities that are characteristic of upwelling provinces (52, 55). In addition to the Fe(II) fluxes from hydrothermal venting, In stark contrast, the average settling velocity of biomass not associated modern continental weathering also contributes Fe to the oceans. with Fe(III) (oxyhydr)oxides was 0.35 m year−1, and biomass can We thus assessed the potential delivery flux of continental Fe(II) to thus be transported at distances more than 6000 km—greater than Thompson et al., Sci. Adv. 2019; 5 : eaav2869 27 November 2019 4 of 9 SCIENCE ADVANCES | RESEARCH ARTICLE A B hv 35 20 30 Fe dep. (mol m−2 year −1) CH4/PP (Tmol year −1) 4Fe2+ + CO2 + 11H2O 15 CH2O H2 CH4 25 8H+ + 4Fe(OH)3 + [CH2O] 20 Fe deposition 10 Fe2+ 15 Global PP BIF Shale 10 A 5 5 Upper Upper Upper coastal upwelling open ocean 0 0 0 25 50 75 100 0 25 50 75 100 Iron Iron Iron % of modern continents % Cells associated with Fe oxidation oxidation oxidation C D 60 25 Fe dep. (mol m−2 year −1) 50 CH4/PP (Tmol year −1) Deep/intermediate 20 40 ocean Sediments 15 Iron reduction/ 30 Fe2+ methanogenesis 10 Carbon 20 Burial B 10 5 Fig. 4. Model of an Archean coastal upwelling zone. A schematic (A) depicting 0 0 the cycling of iron and carbon in the model and the boxes (B) that were used to 0 25 50 75 100 0 250 500 750 1000 create the model structure, accompanied by arrows demonstrating the fluxes of [Fe] µM Upwelling (m year −1) various biochemical parameters between the boxes. Fig. 5. Iron and carbon box model sensitivity results. (A) Model sensitivity results for varying continent size. (B) Percentage of cells associated with Fe(III) (oxyhydr) the width of the modern Pacific Ocean (fig. S6 and table S4). Sub- oxides. (C) Deep ocean Fe(II). (D) Varying upwelling rates in the upwelling provinces. Iron deposition rates in the upwelling provinces are depicted on the y axis, while stantial biomass can thus be exported from productive upwelling global rates of primary production and methane production are both shown on areas and broadly distributed between oceanic provinces before its the x axis in Tmol/year. deposition. Notably, the separation of biomass from Fe(III) (oxyhydr) oxides led to little pelagic and sedimentary Fe recycling within the upwelling province itself. The small fraction of biomass deposited material (Supplementary Materials and fig. S7). At low sedimentation in upwelling regions fueled diagenesis, resulting in both Fe(II) recy- rates characteristic of deep-water environments, organic carbon cling to the water column and the authigenesis of reduced Fe phases concentrations are similar to those of the organic carbon–rich such as siderite and magnetite. Siderite can be an abundant component shales that deposited throughout the Archean Eon (45). in many BIF (24, 25), and an example carbon isotope mass balance The globally integrated rates of these processes varied depending suggests that 30% of the siderite in BIFs can be diagenetic, with the on continental area, biomass association to Fe(III) (oxyhydr)oxides, balance likely representing a primary precipitate (Supplementary Fe(II) concentrations in the deep ocean, and rates of upwelling Materials). In our benchmark scenario (Supplementary Materials), (Fig. 5). These variables can be constrained to likely ranges based on where 15% of cells are associated with their Fe(III) by-products, material fluxes recorded in sedimentary rocks (1, 2, 40), bounds for organic carbon deposition rates could have supported the conversion global cycling (1, 28, 40), and tempered analogies to modern systems of 10% of the total Fe(III) deposited to diagenetic siderite. When the (21, 56)—model outputs within these ranges are delineated by the diagenetic siderite is combined with primary nondiagenetic siderite, shaded area in Fig. 5. Deep ocean Fe(II) concentrations are con- as inferred from the isotopic composition of siderite, the total siderite strained to <70 M, similar to earlier estimates of 50 M based on content of BIFs in our benchmark scenario is as high as 30%. This siderite and calcite solubility (28), because higher Fe(II) concentra- combination of primary and diagenetic siderite and remaining tions led to global Fe(II) oxidation rates that exceed the revised Fe(III) (oxyhydr)oxides yielded an average redox state of Fe in BIF maximum global Fe(II) supplies and recycling combined (table S3). of 2.7 and is very similar to that of many siderite-bearing Neoarchean Even at these relatively low Fe(II) concentrations (20 to 70 M), BIFs (fig. S1 and table S5). upwelling provinces could have supported area-specific rates of BIF Biomass exported from upwelling provinces augmented the deposition of up to 54 mol m−2 year−1 over areas greater than 1011 m2 biomass produced in the other provinces and drove extensive pelagic (Table 1), as needed to deposit the largest of the Archean BIFs (5). and sedimentary Fe recycling through heterotrophic Fe(III) reduc- Notably, with decreased cell-mineral association, increased continent tion. As decoupling of biomass from Fe(III) (oxyhydr)oxides in up- size, increased deep-water Fe(II) concentrations, and increased up- welling regions led to excess biomass over Fe(III) in both other welling rates, Fe(II) consumption could have eventually outpaced provinces, it fueled methanogenesis once Fe(III) had been entirely even upper limits for Fe(II) supply, and this would have led to reduced or was buried as Fe-poor, organic carbon–bearing coastal depletion of deep ocean Fe(II). Within the most likely parameter and deep-sea sediments (Fig. 4). The deposition of biomass in these space, rates of global primary production through photoferrotrophy sediments would have led to organic carbon concentrations of be- would have been less than 1% of modern primary production tween 0.5 and 5 wt % depending on sedimentation rates of detrital (Table 1) and, in the absence of oxygen and sulfate, most of this Thompson et al., Sci. Adv. 2019; 5 : eaav2869 27 November 2019 5 of 9 SCIENCE ADVANCES | RESEARCH ARTICLE Table 1. Model results for both the carbon cycle (primary production, carbon burial, carbon remineralization, and methane production) and the iron cycle (iron recycling and iron deposition) at 2.5 Ga. Low, Middle, and High represent the model outputs from the low, middle, and high parts of the ranges depicted by the gray boxes in Fig. 5 and fig. S8. For reference, modern rates of primary production (PP) range from 2760 to 3510 Tmol C year−1 (60). Fe C to CH4 Fe deposition Global Fe C PP C burial C remineralization C to Fe reduction Global CH4 production recycling (upwelling deposition province) Tmol % of % of Tmol Tmol Tmol Tmol % of PP % of PP C/year remineralization remineralization Fe/year Fe/year Fe/year CH4/year 2.5 Ga Low 4.5 14.8 85.2 72.6 27.4 11.1 1.0 17.9 0.5 2.5 Ga Middle 14.8 14.7 85.3 49.1 50.9 24.8 9.9 59.1 3.2 2.5 Ga High 22.5 14.6 85.4 41.0 59.0 31.4 17.5 89.9 5.7 biomass would have been remineralized through Fe(III) reduction be delivered to the atmosphere (58). On the basis of new solutions and methanogenesis in the open ocean, with the remainder buried (49) to photochemical models (59), this biospheric methane flux in sediments (Table 1). Global primary production is limited by (3.2 Tmol year−1) would support a 10 ppmv (parts per million by P availability over geological time and marine primary production volume) methane atmosphere under our benchmark scenario, with scales with P concentrations in the oceans (57). Reconstruction of a possible range of between 1 and 20 ppmv across the likely model P concentrations in the Archean oceans implied 0.04 to 0.13 M parameter space (Table 1). While these methane concentrations alone (15), which is ~1 to 4% of the modern and sufficient to support our are unlikely to support a warm climate, positive feedbacks with other estimates for primary production. Notably, up to 45% of all remineralized concurrent modes of photosynthesis, like H2-based anoxygenic carbon was channeled through methanogenesis, and the fraction of photosynthesis, markedly increase both biospheric methane fluxes remineralization through methanogenesis increased with decreased to the atmosphere and atmospheric methane concentrations (49). cell-mineral association, increased continent size, increased deep-­ Our observations of extant photoferrotrophs, revised global Fe budgets, water Fe(II) concentrations, and increased upwelling rates (Fig. 5 and models of coupled C and Fe cycling thus support an upside-­ and Table 1). Variations in all of these processes thus likely influ- down Archean biosphere, similar to that originally proposed by enced global rates of iron deposition and methane production and Walker (19), in which the deposition of BIFs leads to a methane-rich could have contributed to the development of differing BIF facies. atmosphere. Our results further imply that cell-Fe(III) separation The model prediction that methanogenesis would have played a augments marine methane production through the breakdown of key role in carbon degradation in Archean ferruginous oceans is also photoferrotrophic biomass and mechanistically ties biospheric methane supported by other Earth system models (49) and observations from fluxes to the deposition of BIF. The role this plays in atmospheric modern ferruginous basins. In Kabuno Bay, East Africa, for example, chemistry and greenhouse warming should be tested through further carbon produced through photoferrotrophy is remineralized through biogeochemical-photochemical modeling efforts. methanogenesis (21), leading to an export of ferric iron from the euphotic zone even under extremely quiescent physical conditions. Likewise, despite abundant Fe(III) (oxyhydr)oxides, as much as 50% METHODS of the carbon degradation in ferruginous Lake Matano is channeled Photoferrotrophic strains and cyanobacteria were grown in basal through methanogenesis, rather than Fe(III) reduction (56). Empirical media (21) and in alternative media (Supplementary Materials and observations thus point generally to carbon and iron cycling under table S1) until late exponential phase. Subsamples of both Fe(II)/ ferruginous conditions that ultimately leads to methane production Fe(III) and pigments were used to track the growth kinetics and to even in the absence of the large-scale advective processes that oper- assess the in vitro cellular association to Fe(III). The surface properties ate in the oceans. When integrated into a box model of the Arche- of each strain and their metabolic Fe(III) by-products were assessed by an marine iron and carbon cycles, our observations of cell-Fe(III) measuring their surface potential using the Particle Metrix ZetaView. oxyhydroxide separation during photoferrotrophy quantitatively The acid-base properties of cell surfaces were determined through titra- describe an Archean biosphere that: (i) accounts for the Fe(III) depo- tions. Cell surface contact angles were measured for DVLO modeling sition rates needed to form BIFs; (ii) explains the fate of biomass in three different liquids—water, glycerol, and diiodomethane (36, 38). and its absence from BIFs; and (iii) provides a strong source of Electron microscopies [scanning electron microscopy (SEM) and methane to the biosphere. transmission electron microscopy (TEM)] were used to image This model demonstrates that, within the geologically con- cell-mineral interactions (Supplementary Materials). The particle sizes strained parameter space, global rates of methanogenesis can easily of the Fe(III) (oxyhydr)oxides produced by the photoferrotrophic reach 3.2 Tmol year−1 (Table 1), and with the near absence of oxygen strains were measured using a Mastersizer 2000 particle size analyzer, and sulfate to fuel methane oxidation, most of this methane would and these were then used to model the impact of horizontal ocean Thompson et al., Sci. Adv. 2019; 5 : eaav2869 27 November 2019 6 of 9 SCIENCE ADVANCES | RESEARCH ARTICLE current velocities on the settling time of both Fe(III) particles and 14. P. Kharecha, J. Kasting, J. Siefert, A coupled atmosphere–ecosystem model of the early Archean Earth. Geobiology 3, 53–76 (2005). cellular biomass (Supplementary Materials, fig. S6, and table S4). 15. C. Jones, S. Nomosatryo, S. A. Crowe, C. J. Bjerrum, D. E. Canfield, Iron oxides, divalent The box model of the Archean marine carbon and iron cycles was cations, silica, and the early earth phosphorus crisis. Geology 43, 135–138 (2015). designed with three boxes (each representing an oceanic province) 16. G. Luo, S. Ono, N. J. Beukes, D. T. Wang, S. Xie, R. E. Summons, Rapid oxygenation to capture variability in fluxes through each box (Supplementary of Earth’s atmosphere 2.33 billion years ago. Sci. Adv. 2, e1600134 (2016). Materials, Figs. 4 and 5, fig. S8, and Table 1). 17. P. Philippot, J. N. Ávila, B. A. Killingsworth, S. Tessalina, F. Baton, T. Caquineau, E. Muller, E. Pecoits, P. Cartigny, S. V. Lalonde, T. R. Ireland, C. Thomazo, M. J. van Kranendonk, V. Busigny, Globally asynchronous sulphur isotope signals require re-definition SUPPLEMENTARY MATERIALS of the Great Oxidation Event. Nat. Commun. 9, 2245 (2018). Supplementary material for this article is available at http://advances.sciencemag.org/cgi/ 18. H. D. Holland, The oxygenation of the atmosphere and oceans. Philos. Trans. R. Soc. 361, content/full/5/11/eaav2869/DC1 903–915 (2006). Supplementary Materials and Methods 19. J. C. Walker, Was the Archaean biosphere upside down? Nature 329, 710–712 (1987). Section S1. Cell surface features and acid-base chemistry 20. A. Kappler, C. Pasquero, K. O. Konhauser, D. K. Newman, Deposition of banded iron Section S2. Cell-iron surface interaction and extended DVLO modeling formations by anoxygenic phototrophic Fe(II)-oxidizing bacteria. Geology 33, 865–868 Section S3. Iron concentration and supply (2005). Section S4. Physical separation of ferric iron oxyhydroxides and cellular biomass in an ocean 21. M. Llirós, T. García-Armisen, F. Darchambeau, C. Morana, X. Triadó-Margarit, Ö. Inceoğlu, setting C. M. Borrego, S. Bouillon, P. Servais, A. V. Borges, J.-P. Descy, D. E. Canfield, S. A. Crowe, Section S5. Box model of Archean marine carbon and iron cycles Pelagic photoferrotrophy and iron cycling in a modern ferruginous basin. Sci. Rep. 5, Section S6. Organic carbon burial and diagenesis 13003 (2015). Table S1. Range of concentrations in the growth media used throughout experiments. 22. K. O. Konhauser, D. K. Newman, A. Kappler, The potential significance of microbial Fe(III) Table S2. Cell surface characteristics and cell-mineral interaction modeling. reduction during deposition of Precambrian banded iron formations. Geobiology 3, Table S3. Modern and Archean Fe fluxes. 167–177 (2005). Table S4. Different scenarios of the physical separation model, with each case using a different 23. T. Gauger, J. M. Byrne, K. O. Konhauser, M. Obst, S. Crowe, A. Kappler, Influence of water velocity. organics and silica on Fe (II) oxidation rates and cell–mineral aggregate formation by Table S5. Data compilations for Fig. 1 and fig. S1. the green-sulfur Fe (II)-oxidizing bacterium Chlorobium ferrooxidans KoFox–Implications Fig. S1. The redox state of iron in BIF through time where the red bars indicate the siderite-rich BIFs. for Fe (II) oxidation in ancient oceans. Earth Planet. Sci. Lett. 443, 81–89 (2016). Fig. S2. Growth curve for C. phaeoferrooxidans strain KB01. 24. A. J. Kaufman, J. Hayes, C. Klein, Primary and diagenetic controls of isotopic compositions Fig. S3. Additional SEM and TEM images of strains KB01 and KoFox under two conditions. of iron-formation carbonates. Geochim. Cosmochim. Acta 54, 3461–3473 (1990). Fig. S4. Surface charge of strains KB01 and KoFox. 25. C. Klein, N. J. Beukes, Geochemistry and sedimentology of a facies transition Fig. S5. Additional cell surface characteristics for strain KoFox. from limestone to iron-formation deposition in the early Proterozoic Transvaal Fig. S6. Modeling the settling velocity of carbon and iron using a range of horizontal current Supergroup, South Africa. Econ. Geol. 84, 1733–1774 (1989). velocities. 26. R. P. Lowell, S. M. Keller, High-temperature seafloor hydrothermal circulation over Fig. S7. Modeled weight % organic carbon in the coastal and open ocean sediments. geologic time and Archean banded iron formations. Geophys. Res. Lett. 30, (2003). Fig. S8. Iron and carbon box model sensitivity results. 27. D. E. Canfield, A new model for Proterozoic ocean chemistry. Nature 396, 450 (1998). References (61–103) 28. H. D. Holland, The Chemical Evolution of the Atmosphere and Oceans (Princeton Univ. View/request a protocol for this paper from Bio-protocol. Press, 1984). 29. S. Heising, L. Richter, W. Ludwig, B. Schink, Chlorobium ferrooxidans sp nov., a phototrophic green sulfur bacterium that oxidizes ferrous iron in coculture with a “Geospirillum” sp strain. Arch. Microbiol. 172, 116–124 (1999). REFERENCES AND NOTES 30. J. B. Fein, C. J. Daughney, N. Yee, T. A. Davis, A chemical equilibrium model for metal 1. A. E. Isley, D. H. Abbott, Plume-related mafic volcanism and the deposition of banded adsorption onto bacterial surfaces. Geochim. Cosmochim. Acta 61, 3319–3328 (1997). iron formation. J. Geophys. Res. Solid Earth 104, 15461–15477 (1999). 31. D. A. Sverjensky, N. Sahai, Theoretical prediction of single-site surface-protonation 2. C. Klein, Some Precambrian banded iron-formations (BIFs) from around the world: Their equilibrium constants for oxides and silicates in water. Geochim. Cosmochim. Acta 60, age, geologic setting, mineralogy, metamorphism, geochemistry, and origins. 3773–3797 (1996). Am. Mineral. 90, 1473–1499 (2005). 32. K. O. Konhauser, Diversity of bacterial iron mineralization. Earth Sci. Rev. 43, 91–121 3. P. Cloud, Paleoecological significance of banded iron formation. Econ. Geol. 68, (1998). 1135–1143 (1973). 33. K. O. Konhauser, S. V. Lalonde, L. Amskold, H. D. Holland, Was there really an Archean 4. R. Garrels, E. Perry, F. Mackenzie, Genesis of Precambrian iron-formations phosphate crisis? Science 315, 1234 (2007). and the development of atmospheric oxygen. Econ. Geol. 68, 1173–1179 (1973). 34. V. R. Phoenix, D. G. Adams, K. O. Konhauser, Cyanobacterial viability during hydrothermal 5. K. O. Konhauser, T. Hamade, R. Raiswell, R. C. Morris, F. G. Ferris, G. Southam, biomineralisation. Chem. Geol. 169, 329–338 (2000). D. E. Canfield, Could bacteria have formed the Precambrian banded iron formations? 35. E. D. Swanner, Physiology, Fe (II) oxidation, and Fe mineral formation by a marine Geology 30, 1079–1082 (2002). planktonic cyanobacterium grown under ferruginous conditions. Front. Earth Sci. 3, 60 6. A. Kappler, D. K. Newman, Formation of Fe(III)-minerals by Fe(II)-oxidizing (2015). photoautotrophic bacteria. Geochim. Cosmochim. Acta 68, 1217–1226 (2004). 36. G. Saini, C. Chan, Near-neutral surface charge and hydrophilicity prevent mineral 7. B. Rasmussen, B. Krapež, D. B. Meier, Replacement origin for hematite in 2.5 Ga banded encrustation of Fe-oxidizing micro-organisms. Geobiology 11, 191–200 (2013). iron formation: Evidence for postdepositional oxidation of iron-bearing minerals. Bulletin 37. O. Stukalov, A. Korenevsky, T. J. Beveridge, J. R. Dutcher, Use of atomic force microscopy 126, 438–446 (2014). and transmission electron microscopy for correlative studies of bacterial capsules. 8. N. J. Tosca, S. Guggenheim, P. K. Pufahl, An authigenic origin for Precambrian greenalite: Appl. Environ. Microbiol. 74, 5457–5465 (2008). Implications for iron formation and the chemistry of ancient seawater. Bulletin 128, 38. C. J. Van Oss, Interfacial forces in aqueous media (CRC press, 2006). 511–530 (2016). 39. N. Yee, J. Fein, Cd adsorption onto bacterial surfaces: A universal adsorption edge? 9. F. Widdel, S. Schnell, S. Heising, A. Ehrenreich, B. Assmus, B. Schink, Ferrous iron oxidation Geochim. Cosmochim. Acta 65, 2037–2042 (2001). by anoxygenic phototrophic bacteria. Nature 362, 834–836 (1993). 40. A. E. Isley, Hydrothermal plumes and the delivery of iron to banded iron formation. 10. W. F. Martin, D. A. Bryant, J. T. Beatty, A physiological perspective on the origin J. Geol. 103, 169–185 (1995). and evolution of photosynthesis. FEMS Microbiol. Rev., (2017). 41. S. Poulton, R. Raiswell, The low-temperature geochemical cycle of iron: From continental 11. W. W. Fischer, J. Hemp, J. E. Johnson, Evolution of oxygenic photosynthesis. fluxes to marine sediment deposition. Am. J. Sci. 302, 774–805 (2002). Annu. Rev. Earth Planet. Sci. 44, 647–683 (2016). 42. L. R. Kump, W. E. Seyfried, Hydrothermal Fe fluxes during the Precambrian: Effect of low 12. R. E. Blankenship, M. T. Madigan, C. E. Bauer, Anoxygenic Photosynthetic Bacteria (Springer oceanic sulfate concentrations and low hydrostatic pressure on the composition of Science & Business Media, 2006), vol. 2. black smokers. Earth Planet. Sci. Lett. 235, 654–662 (2005). 13. D. E. Canfield, M. T. Rosing, C. Bjerrum, Early anaerobic metabolisms. 43. J. Korenaga, Archean Geodynamics and the Thermal Evolution of Earth (Geophysical Philos. Trans. R. Soc. B. Biol. Sci. 361, 1819–1834 (2006). Monograph Series, American Geophysical Union, 2006), pp. 7–32. Thompson et al., Sci. Adv. 2019; 5 : eaav2869 27 November 2019 7 of 9 SCIENCE ADVANCES | RESEARCH ARTICLE 44. S. A. Crowe, G. Paris, S. Katsev, C. Jones, S.-T. Kim, A. L. Zerkle, S. Nomosatryo, D. A. Fowle, 73. D. G. Brown, P. R. Jaffé, Effects of nonionic surfactants on the cell surface hydrophobicity J. F. Adkins, A. L. Sessions, J. Farquhar, D. E. Canfield, Sulfate was a trace constituent and apparent Hamaker constant of a Sphingomonas sp. Environ. Sci. Technol. 40, 195–201 of Archean seawater. Science 346, 735–739 (2014). (2006). 45. T. W. Lyons, C. T. Reinhard, N. J. Planavsky, The rise of oxygen in Earth/’s early ocean 74. W. Norde, J. Lyklema, Protein adsorption and bacterial adhesion to solid surfaces: and atmosphere. Nature 506, 307–315 (2014). A colloid-chemical approach. Colloids Surf. 38, 1–13 (1989). 46. S. A. Crowe, L. N. Døssing, N. J. Beukes, M. Bau, S. J. Kruger, R. Frei, D. E. Canfield, 75. J. N. Ryan, P. M. Gschwend, Effects of ionic strength and flow rate on colloid release: Atmospheric oxygenation three billion years ago. Nature 501, 535–538 (2013). Relating kinetics to intersurface potential energy. J. Colloid Interface Sci. 164, 21–34 47. I. Halevy, A. Bachan, The geologic history of seawater pH. Science 355, 1069–1071 (2017). (1994). 48. N. H. Sleep, K. Zahnle, P. Neuhoff, Initiation of clement surface conditions on the earliest Earth. 76. S. W. Poulton, D. E. Canfield, Ferruginous conditions: A dominant feature of the ocean Proc. Natl. Acad. Sci. U.S.A. 98, 3666–3672 (2001). through Earth’s history. Elements 7, 107–112 (2011). 49. K. Ozaki, E. Tajika, P. K. Hong, Y. Nakagawa, C. T. Reinhard, Effects of primitive 77. S. M. McLennan, S. Taylor, Continental freeboard, sedimentation rates and growth photosynthesis on Earth’s early climate system. Nat. Geosci. 11, 55–59 (2017). of continental crust. Nature 306, 169–172 (1983). 50. N. J. Planavsky, D. Asael, A. Hofmann, C. T. Reinhard, S. V. Lalonde, A. Knudsen, X. Wang, 78. D. Lowe, Major events in the geological development of the Precambrian Earth, in F. Ossa Ossa, E. Pecoits, A. J. B. Smith, N. J. Beukes, A. Bekker, T. M. Johnson, The Proterozoic Biosphere: A Multidisciplinary Study (Cambridge Univ. Press, 1992), K. O. Konhauser, T. W. Lyons, O. J. Rouxel, Evidence for oxygenic photosynthesis half pp. 67–75. a billion years before the Great Oxidation Event. Nat. Geosci. 7, 283–286 (2014). 79. K. C. Condie, Chemical composition and evolution of the upper continental crust: 51. S. L. Olson, L. R. Kump, J. F. Kasting, Quantifying the areal extent and dissolved oxygen Contrasting results from surface samples and shales. Chem. Geol. 104, 1–37 concentrations of Archean oxygen oases. Chem. Geol. 362, 35–43 (2013). (1993). 52. J. H. Martin, G. A. Knauer, D. M. Karl, W. W. Broenkow, VERTEX: Carbon cycling 80. C. Woldringh, J. Binnerts, A. Mans, Variation in Escherichia coli buoyant density measured in the northeast Pacific. Deep Sea Res. Part A Oceanogr. Res. Pap. 34, 267–285 (1987). in Percoll gradients. J. Bacteriol. 148, 58–63 (1981). 53. A. H. Devol, H. E. Hartnett, Role of the oxygen-deficient zone in transfer of organic carbon 81. K. D. Stolzenbach, K. A. Newman, C. S. Wong, Aggregation of fine particles at to the deep ocean. Limnol. Oceanogr. 46, 1684–1690 (2001). the sediment-water interface. J. Geophys. Res. Oceans 97, 17889–17898 (1992). 54. J. F. Banfield, S. A. Welch, H. Zhang, T. T. Ebert, R. L. Penn, Aggregation-based crystal growth 82. W. S. Reeburgh, Oceanic methane biogeochemistry. Chem. Rev. 107, 486–513 (2007). and microstructure development in natural iron oxyhydroxide biomineralization products. 83. S. Crowe, J. A. Maresca, C. Jones, A. Sturm, C. Henny, D. A. Fowle, R. P. Cox, E. F. Delong, Science 289, 751–754 (2000). D. E. Canfield, Deep-water anoxygenic photosythesis in a ferruginous chemocline. 55. D. Kadko, Upwelling and primary production during the US GEOTRACES East Pacific Geobiology 12, 322–339 (2014). Zonal Transect. Global Biogeochem. Cycles 31, 218–232 (2017). 84. S. A. Crowe, C. Jones, S. Katsev, C. Magen, A. H. O’Neill, A. Sturm, D. E. Canfield, 56. S. Crowe, S. Katsev, K. Leslie, A. Sturm, C. Magen, S. Nomosatryo, M. A. Pack, J. D. Kessler, G. D. Haffner, A. Mucci, B. Sundby, D. A. Fowle, Photoferrotrophs thrive W. S. Reeburgh, J. A. Roberts, L. González, G. D. Haffner, A. Mucci, B. Sundby, D. A. Fowle, in an Archean Ocean analogue. Proc. Natl. Acad. Sci. U.S.A. 105, 15938–15943 The methane cycle in ferruginous Lake Matano. Geobiology 9, 61–78 (2011). (2008). 57. C. J. Bjerrum, D. E. Canfield, Ocean productivity before about 1.9 Gyr ago limited by 85. L. Kuntz, T. Laakso, D. Schrag, S. Crowe, Modeling the carbon cycle in Lake Matano. phosphorus adsorption onto iron oxides. Nature 417, 159 (2002). Geobiology 13, 454–461 (2015). 58. D. C. Catling, K. J. Zahnle, C. P. McKay, Biogenic methane, hydrogen escape, 86. P. R. Craddock, N. Dauphas, Iron and carbon isotope evidence for microbial iron and the irreversible oxidation of early Earth. Science 293, 839–843 (2001). respiration throughout the Archean. Earth Planet. Sci. Lett. 303, 121–132 (2011). 59. A. A. Pavlov, L. L. Brown, J. F. Kasting, UV shielding of NH3 and O2 by organic hazes 87. N. J. Beukes, C. Klein, Geochemistry and sedimentology of a facies transition—from in the Archean atmosphere. J. Geophys. Res. Planet 106, 23267–23287 (2001). microbanded to granular iron-formation—in the early Proterozoic Transvaal Supergroup, 60. D. Antoine, J. M. André, A. Morel, Oceanic primary production: 2. Estimation at global South Africa. Precambrian Res. 47, 99–139 (1990). scale from satellite (coastal zone color scanner) chlorophyll. Global Biogeochem. Cycles 88. M. Baur, J. Hayes, S. Studley, M. Walter, Millimeter-scale variations of stable isotope 10, 57–69 (1996). abundances in carbonates from banded iron-formations in the Hamersley Group 61. F. Hegler, N. R. Posth, J. Jiang, A. Kappler, Physiology of phototrophic iron(II)-oxidizing of Western Australia. Econ. Geol. 80, 270–282 (1985). bacteria: Implications for modern and ancient environments. FEMS Microbiol. Ecol. 66, 89. C. Klein, E. A. Ladeira, Geochemistry and mineralogy of Neoproterozoic banded 250–260 (2008). iron-formations and some selected, siliceous manganese formations from the Urucum 62. E. Viollier, P. Inglett, K. Hunter, A. Roychoudhury, P. Van Cappellen, The ferrozine method District, Mato Grosso do Sul, Brazil. Econ. Geol. 99, 1233–1244 (2004). revisited: Fe (II)/Fe (III) determination in natural waters. Appl. Geochem. 15, 785–790 90. R. F. Dymek, C. Klein, Chemistry, petrology and origin of banded iron-formation (2000). lithologies from the 3800 Ma Isua supracrustal belt, West Greenland. Precambr. Res. 39, 63. N.-U. Frigaard, K. L. Larsen, R. P. Cox, Spectrochromatography of photosynthetic 247–302 (1988). pigments as a fingerprinting technique for microbial phototrophs. FEMS Microbiol. Ecol. 91. C. Klein, N. J. Beukes, Sedimentology and geochemistry of the glaciogenic late 20, 69–77 (1996). Proterozoic Rapitan iron-formation in Canada. Econ. Geol. 88, 542–565 (1993). 64. A. Korenevsky, T. J. Beveridge, The surface physicochemistry and adhesiveness 92. C. Klein, E. Ladeira, Geochemistry and petrology of some Proterozoic banded of Shewanella are affected by their surface polysaccharides. Microbiology 153, iron-formations of the Quadrilátero Ferrífero, Minas Gerais, Brazil. Econ. Geol. 95, 405–427 1872–1883 (2007). (2000). 65. D. Suter, C. Siffert, B. Sulzberger, W. Stumm, Catalytic dissoluion of iron (III)(hydr) oxides by oxalic acid in the presence of Fe (II). Naturwissenschaften 75, 571–573 (1988). 93. H. D. Holland, M. Schidlowski, Mineral Deposits and the Evolution of the Biosphere 66. D. R. Lovley, E. J. Phillips, Availability of ferric iron for microbial reduction in bottom (Springer, 1982). sediments of the freshwater tidal Potomac River. Appl. Environ. Microbiol. 52, 751–757 94. A. R. Cabral, R. A. Creaser, T. Nägler, B. Lehmanna, A. R. Voegelin, B. Belyatsky, J. Pašava, (1986). A. A. Seabra Gomes Jr., H. Galbiatti, M. E. Böttcher, P. Escheri, Trace-element and multi- 67. E. E. Roden, J. M. Zachara, Microbial reduction of crystalline iron (III) oxides: Influence isotope geochemistry of Late-Archean black shales in the Carajás iron-ore district, Brazil. of oxide surface area and potential for cell growth. Environ. Sci. Technol. 30, 1618–1628 Chem. Geol. 362, 91–104 (2013). (1996). 95. Y. Watanabe, H. Naraoka, D. J. Wronkiewicz, K. C. Condie, H. Ohmoto, Carbon, nitrogen, 68. R. E. Martinez, D. S. Smith, K. Pedersen, F. G. Ferris, Surface chemical heterogeneity and sulfur geochemistry of Archean and Proterozoic shales from the Kaapvaal Craton, of bacteriogenic iron oxides from a subterranean environment. Environ. Sci. Technol. 37, South Africa. Geochim. Cosmochim. Acta 61, 3441–3459 (1997). 5671–5677 (2003). 96. D. McKirdy, Organic geochemistry in Precambrian research. Precambrian Res. 1, 75–137 69. D. S. Smith, F. G. Ferris, Proton binding by hydrous ferric oxide and aluminum oxide (1974). surfaces interpreted using fully optimized continuous p K a spectra. Environ. Sci. Technol. 97. J. J. Middelburg, J. Nieuwenhuize, P. van Breugel, Black carbon in marine sediments. 35, 4637–4642 (2001). Mar. Chem. 65, 245–252 (1999). 70. R. E. Martinez, D. S. Smith, E. Kulczycki, F. G. Ferris, Determination of intrinsic bacterial 98. A. L. Paropkari, C. P. Babu, A. Mascarenhas, New evidence for enhanced preservation surface acidity constants using a Donnan shell model and a continuous pK of organic carbon in contact with oxygen minimum zone on the western continental a distribution method. J. Colloid Interface Sci. 253, 130–139 (2002). slope of India. Mar. Geol. 111, 7–13 (1993). 71. J. Cowen, Morphological study of marine bacterial capsules: Implications for marine 99. J. Niggemann, T. G. Ferdelman, B. A. Lomstein, J. Kallmeyer, C. J. Schubert, aggregates. Mar. Biol. 114, 85–95 (1992). How depositional conditions control input, composition, and degradation of organic 72. C. Van Oss, Hydrophobicity of biosurfaces—Origin, quantitative determination matter in sediments from the Chilean coastal upwelling region. Geochim. Cosmochim. Acta and interaction energies. Colloids Surf. B Biointerfaces 5, 91–110 (1995). 71, 1513–1527 (2007). Thompson et al., Sci. Adv. 2019; 5 : eaav2869 27 November 2019 8 of 9 SCIENCE ADVANCES | RESEARCH ARTICLE 100. C. Klein, E. A. Ladeira, Petrography and geochemistry of the least altered banded observations in Kabuno Bay. K.J.T. performed photoferrotroph cell settling experiments. T.W. iron-formation of the Archean Carajás Formation. northern Brazil. Econ. Geol. 97, 643–651 performed cyanobacteria settling experiments. K.J.T. and P.A.K. conducted electron (2002). microscopy imaging of strain KB01. T.G. and A.K. conducted electron microscopy imaging of 101. C. Klein, Mineralogy and petrology of the metamorphosed Wabush Iron Formation, strain KoFox. K.J.T., P.A.K., R.L.S., and R.M. characterized cell surface chemistries. K.J.T. southwestern Labrador. J. Petrol. 7, 246–305 (1966). conducted DVLO modeling. K.J.T., S.A.C., K.W.B., and C.C.M. constructed the biogeochemical 102. N. R. Posth, S. Huelin, K. O. Konhauser, A. Kappler, Size, density and composition models. K.J.T. and S.A.C. wrote the manuscript with detailed input from P.A.K., K.O.K., C.T.R., of cell–mineral aggregates formed during anoxygenic phototrophic Fe (II) oxidation: and A.K. and contributions from all other authors. Competing interests: The authors declare Impact on modern and ancient environments. Geochim. Cosmochim. Acta 74, 3476–3493 that they have no competing interest. Data and materials availability: All data needed to (2010). evaluate the conclusions in the paper are present in the paper and/or the Supplementary 103. D. E. Canfield, Sulfate reduction and oxic respiration in marine sediments: Implications Materials. Additional data related to this paper may be requested from the authors. for organic carbon preservation in euxinic environments. Deep Sea Res. Part A Oceanogr. Correspondence and requests for materials should be addressed to S.A.C. (sean.crowe@ubc.ca). Res. Pap. 36, 121–138 (1989). Submitted 2 September 2018 Acknowledgments: Anonymous reviewers provided suggestions and criticism. N. Yee, Accepted 18 September 2019 C. Chan, and J. Smit provided insight into cell surface chemistry. Imaging was conducted at the Published 27 November 2019 UBC Bioimaging Facility, UBC Pharmaceutical Sciences Imaging Facility, and the SFU 4D LABS 10.1126/sciadv.aav2869 Imaging Facility. Funding: This work was supported by NSERC Discovery Grant 04867 to S.A.C., NSERC Postgraduate Scholarship–Doctoral Program, and Canada Research Chairs Program. Citation: K. J. Thompson, P. A. Kenward, K. W. Bauer, T. Warchola, T. Gauger, R. Martinez, R. L. Simister, A.K. was supported by the European Research Council under the European Union’s Seventh C. C. Michiels, M. Llirós, C. T. Reinhard, A. Kappler, K. O. Konhauser, S. A. Crowe, Photoferrotrophy, Framework Program (FP/2007–2013)/ERC Grant agreement no. 307320-MICROFOX. Author deposition of banded iron formations, and methane production in Archean oceans. Sci. Adv. 5, contributions: S.A.C. conceived, designed, and directed the study. S.A.C. and M.L. made eaav2869 (2019). Thompson et al., Sci. Adv. 2019; 5 : eaav2869 27 November 2019 9 of 9