K O Konhauser

N J Planavsky

D S Hardisty

T J Warchola

R Haugaard

S V Lalonde

C A Partin

P B H Oonk

H Tsikos

T W Lyons

A Bekker

Iron Formations: A Global Record of Neoarchaean to Palaeoproterozoic Environmental History

 that will improve the quantitative framework we use to interpret IF deposition. In this work, we also discuss how recent discoveries have provided new insights into the processes underpinning the global rise in atmospheric oxygen and the geochemical evolution of the oceans.

Introduction

Iron formations (IFs) are iron-rich (15-40 wt.% Fe) and siliceous (40-60 wt.% SiO 2 ) chemical sedimentary rocks that precipitated from seawater throughout the Precambrian Eon, the majority having been deposited between 2.80 and 1.85 Ga in the Neoarchaean and Palaeoproterozoic (Figs 1-2). Minimal detrital input during deposition of deep water IF is indicated by low concentrations of Al 2 O 3 (<1 wt.%) and trace elements that are enriched in crustal rocks (Ti,Zr,Th,Hf,and Sc <20 ppm). Most IFs show distinctive layering of variable thickness, from macrobands (meter-thick) to mesobands (centimeter-thick) by which they are typically defined (i.e., banded iron formations, BIFs), to microbands (millimeter and submillimeter layers). Microbands have been linked to episodic hydrothermal input, hypothesised in early studies to represent an annual depositional process (e.g., [START_REF] Trendall | The iron formations of the Precambrian Hamersley Group, Western Australia with special reference to the associated crocidolite[END_REF][START_REF] Morris | Genetic modelling for banded iron-formations of the Hamersley Group, Pilbara Craton, Western Australia[END_REF]. Layering in BIF can be laterally extensive, indicating deposition below wave base. Some IF, known as granular iron formations (GIFs), however, lack this characteristic banding. These are composed of granules of chert and iron oxides or silicates, with early diagenetic chert, carbonate or hematite cement filling pore space (e.g., [START_REF] Simonson | Sedimentology of cherts in the Early Proterozoic Wishart Formation, Quebec-Newfoundland, Canada[END_REF]. GIFs are generally thought to have been deposited in shallow water in proximal settings. IFs of all types can provide important information regarding the composition of past oceans, but only if their physical and chemical depositional environment, basin tectonic setting, and primary versus secondary signals can be teased apart. In this review, we address the current state of knowledge surrounding IF deposition and what it can tell us about the evolving Earth.

IFs are broadly defined as being either Algoma-type or Superior-type based on depositional environment [START_REF] Gross | A classification of iron-formation based on depositional environments[END_REF], although it is more practical to consider these classifications as end members having a spectrum of intermediate varieties (Bekker et al., 2012). Algoma-type IF are interlayered with, or stratigraphically and genetically linked to, submarine-emplaced maficultramafic to felsic volcanic rocks and associated volcaniclastic greywackes and shales in greenstone belts, and in many cases, spatially coupled with volcanogenic massive sulfide (VMS) deposits [START_REF] Bekker | Iron formation: The sedimentary product of a complex interplay among mantle, tectonic, oceanic, and biospheric processes[END_REF]. Algoma-type IFs were apparently formed close to volcanic arcs and spreading centres and produced by exhalative hydrothermal processes related to submarine volcanism (e.g., [START_REF] Goodwin | Structure, stratigraphy, an origin of iron formations, Michipicoten Area, Algoma District, Ontario, Canada[END_REF][START_REF] Gross | A classification of iron-formation based on depositional environments[END_REF][START_REF] Gross | Tectonic systems and the deposition of iron-formation[END_REF]. These IFs contain oxide, silicate, and carbonate facies and commonly grade into sulfidic sediments, which can be enriched in copper, zinc, lead, silver, and gold. The lack of any sedimentological features besides fine banding (i.e., absence of current-, tide-, or wave-generated sedimentary structures) in Algoma-type IF indicates a deep-water depositional environment, likely distal to continental landmasses [START_REF] Bekker | Iron formation: The sedimentary product of a complex interplay among mantle, tectonic, oceanic, and biospheric processes[END_REF].

In this regard, Algoma-type IFs have compositions that reflect the influence of local hydrothermal vent fluids in either partially closed basins or open seawater systems [START_REF] Gourcerol | Depositional setting of Algoma-type banded iron formation[END_REF]. Iron formations that precipitated before 3.0 Ga are generally described as being Algoma-type, though there are also examples that formed at 2.7-2.6 Ga, as well as non-Algoma, shallow-marine jasper/IF sequences in the Palaeoarchaean.

In contrast, Superior-type IFs developed in passive-margin, sediment-dominated settings (i.e., after stable continental shelves evolved) and generally lack direct stratigraphic relations with volcanic rocks. These IFs are regarded as having been deposited in near-shore shelf environments (e.g., [START_REF] Trendall | The significance of iron-formation in the Precambrian stratigraphic record[END_REF] because they are typically interbedded with, or grade into, carbonates and black shales [START_REF] Gross | A classification of iron-formation based on depositional environments[END_REF]. Unlike most Algoma-type IFs, which rarely extend for more than 10 km along strike and typically are not more than 50 m thick, Superior-type IFs can be very laterally extensive, with original aerial extents estimated in some cases to be over 100,000 km 2 [START_REF] Isley | Hydrothermal plumes and the delivery of iron to banded iron formation[END_REF][START_REF] Bekker | Iron formation: The sedimentary product of a complex interplay among mantle, tectonic, oceanic, and biospheric processes[END_REF]. In terms of mass, the largest Superior-type IFs contain over 10 14 tons of iron [START_REF] Isley | Hydrothermal plumes and the delivery of iron to banded iron formation[END_REF], supplying the majority of iron required for the world's economy. These IFs include those of the Hamersley Range in Australia, Minas Gerais in Brazil, Transvaal Supergroup in South Africa, Krivoy Rog in Ukraine, and the Labrador Trough in Canada (see Fig. 1).

As noted above, texturally, IF are divided into two groups -BIF and GIF. BIF are dominant in Eoarchaean to early Palaeoproterozoic successions. They consist mainly of interbanded iron-and silica-rich layers and were generally, but not entirely, deposited in relatively deep-water settings, as they typically lack evidence for wave or storm action [START_REF] Simonson | Was the deposition of large Precambrian iron formations linked to major marine transgressions?[END_REF][START_REF] Trendall | The significance of iron-formation in the Precambrian stratigraphic record[END_REF][START_REF] Krapež | Hydrothermal and resedimented origins of the precursor sediments to banded iron formations: Sedimentological evidence from the early Palaeoproterozoic Brockman Supersequence of Western Australia[END_REF] (Fig. 3). Most BIF commonly are either overlain or underlain by organic matter-rich and sulfidic shales, and, in some cases, are interstratified with them. For example, the well-studied 2.48 Ga Dales Gorge Member of the Brockman Iron Formation in Western Australia is composed of 17 BIF macrobands intercalated with 16 shale ("S") macrobands [START_REF] Trendall | The iron formations of the Precambrian Hamersley Group, Western Australia with special reference to the associated crocidolite[END_REF], a sequence that has been correlated with age-equivalent IFs in South Africa [START_REF] Beukes | Origin and paleoenvironmental significance of major iron formations at the Archean-Paleoproterozoic boundary[END_REF]. These shales are thought to be shallower-water equivalents of BIF (e.g., Beukes and Klein, 1990;Beukes et al., 1990;[START_REF] Beukes | A lithostratigraphic-sedimentological reference profile for the Late Archaean Mozaan Group, Pongola Sequence: Application to sequence stratigraphy and correlation with the Witwatersrand Supergroup[END_REF][START_REF] Bau | Distribution of yttrium and rare-earth elements in the Penge and Kuruman iron-formations, Transvaal Supergroup, South Africa[END_REF]. There are, however, several key examples of shallow-water Archaean BIF that are interbedded with sandstones (e.g., Fralick and Pufahl, 2006).

GIF first appear in the rock record at ca. 2.90 Ga [START_REF] Siahi | Carbonate ooids of the Mesoarchean Pongola Supergroup, South Africa[END_REF], reach their acme at ca. 1.88 Ga, and are then replaced by ironstones in the Phanerozoic record (after ca. 543 Ma) [START_REF] Bekker | Iron formations: Their origins and implications for ancient seawater chemistry[END_REF]. They are characterised by granules that can vary in morphology, size (micrometers to centimetre in diameter), and composition (chert, iron oxides, iron carbonates, and/or iron silicates); the cement is typically chert, carbonate or hematite (Fig. 4). Many of the granules are considered detrital, with some being derived by sedimentary re-working of iron-rich clay, mudstones, arenites, and even stromatolites (e.g., [START_REF] Ojakangas | Tidal deposits in the Early Proterozoic basin of the Lake Superior regionthe Palms and the Pokegama Formations: Evidence for subtidal shelf deposition of Superior type banded iron formation[END_REF][START_REF] Simonson | First discovery of ferruginous chert arenites in the early Precambrian Hamersley Group of Western Australia[END_REF]. Others, such as the ca.

2.32 Ga oolitic GIF of the Timeball Hill Formation in South Africa, are composed of concentric cortices of hematite [START_REF] Dorland | Paleoproterozoic laterites, red beds and ironstones of the Pretoria Group with reference to the history of atmospheric oxygen[END_REF] that were likely precipitated where Fe(II)-rich waters met more oxygenated shallow seawater. The granules are either grain-or cement-supported; in the case of the latter, it appears that the cement precipitated early, protecting the granules from compaction and preserving "floating" textures. The presence of wave-and current-formed sedimentary structures and hummocky cross-stratification indicates that the granules were deposited near-shore in water depths close to, or above, storm and fair-weather wave base (e.g., [START_REF] Dorland | Paleoproterozoic laterites, red beds and ironstones of the Pretoria Group with reference to the history of atmospheric oxygen[END_REF][START_REF] Pufahl | Depositional controls on Palaeoproterozoic iron formation accumulation, Gogebic Range, Lake Superior region, USA[END_REF].

What is the mineralogy of IFs and what do they record about coeval seawater? The bestpreserved successions are remarkably uniform, composed mostly of fine-grained quartz (in the form of chert), magnetite, and hematite, with variable amounts of Fe-rich silicate minerals (e.g., stilpnomelane, minnesotaite, greenalite, and riebeckite), carbonate minerals (siderite, ankerite, calcite, and dolomite), and locally sparse sulfides (pyrite and pyrrhotite). The presence of both ferric and ferrous minerals gives IFs an average oxidation state of Fe 2.4+ [START_REF] Klein | Time distribution, stratigraphy, and sedimentologic setting, and geochemistry of Precambrian iron-formation[END_REF]. [START_REF] James | Sedimentary facies of iron-formation[END_REF] defined four facies of IFs: oxide, silicate, carbonate, and sulfide; they are generally interbedded with variably recrystallised chert [START_REF] Simonson | Origin and evolution of large Precambrian iron formations[END_REF]. The oxide facies consist predominantly of magnetite and/or hematite, whereas carbonate-facies varieties contain siderite or ankerite as major constituents. The mineralogy of silicate-facies IF is more complex and depends to a large extent on the degree of metamorphism. Under relatively low-grade metamorphic conditions (greenschist facies or below), greenalite, minnesotaite, stilpnomelane, chamosite, and ripidolite (Fechlorite), riebeckite, and ferriannite may be present [START_REF] Klein | Some Precambrian banded iron-formations (BIFs) from around the world: Their age, geologic setting, mineralogy, metamorphism, geochemistry, and origin[END_REF]. At higher grades, cummingtonite, grunerite, actinolite, hornblende, pyroxene, garnet, and fayalite can be present. The sulfide-facies IF described by [START_REF] James | Sedimentary facies of iron-formation[END_REF] are now believed to be pyritic carbonaceous shales and ferruginous slates, and therefore do not represent a chemical precipitate [START_REF] Bekker | Iron formation: The sedimentary product of a complex interplay among mantle, tectonic, oceanic, and biospheric processes[END_REF]. Given that the accumulation of sulfide and subsequent pyrite formation requires iron limitation [START_REF] Berner | Sedimentary pyrite formation: An update[END_REF][START_REF] Poulton | Ferruginous conditions: A dominant feature of the ocean through Earth's history[END_REF], [START_REF] Bekker | Iron formation: The sedimentary product of a complex interplay among mantle, tectonic, oceanic, and biospheric processes[END_REF] suggested that sulfide-facies IF should no longer be considered a variety of IF, sensu stricto.

It is generally agreed that none of the minerals in IF are primary in origin in the sense that the original seafloor precipitate was likely a water-rich gel that was not preserved. Instead, the observed minerals reflect significant post-depositional alteration under diagenetic and metamorphic conditions, including, in some cases, post-depositional fluid flow, particularly within oxidised and mineralised zones. The effect of increasing temperature and pressure is manifested by the progressive change in mineralogy through replacement and recrystallisation, increase in grain size, and obliteration of primary textures [START_REF] Klein | Some Precambrian banded iron-formations (BIFs) from around the world: Their age, geologic setting, mineralogy, metamorphism, geochemistry, and origin[END_REF]. Traditionally, iron oxides in IF are interpreted to have formed from an initial Fe(III)-oxyhydroxide phase, e.g., ferrihydrite, Fe(OH) 3 , which precipitated in the photic zone from seawater via the oxidation of dissolved ferrous iron, Fe(II), at concentrations that may have ranged from 0.03 to 0.5 mM; [START_REF] Holland | Oceans -possible source of iron in iron-formations[END_REF][START_REF] Morris | Genetic modelling for banded iron-formations of the Hamersley Group, Pilbara Craton, Western Australia[END_REF]. By contrast, modern seawater typically has iron concentrations in the range of 0.05 to 2.0 nM (de Baar and de [START_REF] De Baar | Distributions, sources and sinks of iron in seawater[END_REF]. However, recent work suggests that in the presence of dissolved silica, the initial water column precipitate may have been a Fe(III)-Si gel (e.g., [START_REF] Percak-Dennett | Iron isotope fractionation during dissimilatory iron reduction under simulated Archean conditions[END_REF].

During IF mineralisation, sinking of Fe(III)-silica particles to the seafloor eventually led to formation of (1) magnetite or iron carbonates when organic remineralisation was coupled to Fe(III) reduction, either during diagenesis or metamorphism; (2) hematite, through dewatering and silica release, assuming insufficient organic carbon for Fe(III) reduction; (3) iron silicates, possibly in the form of a precursor mineral, such as greenalite ((Fe) 3 Si 2 O 5 (OH) 4 ), when silica sorbed onto or was incorporated in Fe(III)-oxyhydroxides reacted with other cationic species within sediment pore waters [START_REF] Morris | Genetic modelling for banded iron-formations of the Hamersley Group, Pilbara Craton, Western Australia[END_REF]Fisher and Knoll, 2009); and (4) precipitation of quartz through dehydration of opaline silica and/or Fe(III)-Si gels. Ferrous iron sorption to those particles may also have given rise to 'green rust'-type deposits that eventually transformed into magnetite or iron silicates [START_REF] Tamaura | The synthesis of green rust II(FeII1-FeII2) and its spontaneous transformation into Fe3O4[END_REF][START_REF] Zegeye | Green rust formation controls nutrient availability in a ferruginous water column[END_REF]Halevy et al., 2017). Recently, greenalite has also been suggested as a primary precipitate that formed in the water column [START_REF] Rasmussen | Precipitation of iron silicate nanoparticles in early Precambrian oceans marks Earth's first iron age[END_REF][START_REF] Rasmussen | Dust to dust: Evidence for the formation of "primary" hematite dust in banded iron formations via oxidation of iron silicate nanoparticles[END_REF][START_REF] Rasmussen | Greenalite precipitation linked to the deposition of banded iron formations downslope from a late Archean carbonate platform[END_REF]; a more detailed discussion of the primary mineralogy of BIF is given below in section 2.3, but we note here the consensus suggests that the primary mineral phase being a silica-associated Fe(III)-oxyhydroxide.

Organic material, whether in the form of pelagic rain or intimately associated with sinking Fe(III)-oxyhydroxide particles, would have fuelled the organic remineralisation reactions mentioned above, and is likely to have been largely consumed by excess Fe(III) oxidants during diagenesis and metamorphism (see section 2.6). Although siderite in some IF has been interpreted to be a primary mineral phase (e.g., Beukes and Klein, 1990), there are strong isotopic and petrographic arguments that point instead to formation linked to the remineralisation of organic matter (e.g., [START_REF] Johnson | Ancient geochemical cycling in the Earth as inferred from Fe isotope studies of banded iron formations from the Transvaal Craton[END_REF]Johnson et al., , 2008;;[START_REF] Higgins | Oxygenation of the ocean and sediments: Consequences for the seafloor carbonate factory[END_REF][START_REF] Pecoits | Petrography and geochemistry of the Dales Gorge banded iron formation: Paragenetic sequence, source and implications for palaeo-ocean chemistry[END_REF][START_REF] Heimann | Fe, C, and O isotope compositions of banded iron formation carbonates demonstrate a major role for dissimilatory iron reduction in ~ 2.5 Ga marine environments[END_REF][START_REF] Johnson | Iron formation carbonates: Paleoceanographic proxy or recorder of microbial diagenesis?[END_REF].

Chert is commonly considered to have precipitated from the water column or in sediment porewaters as colloidal silica co-precipitated with Fe(III)-oxyhydroxide particles, but it is also possible that chert formed through the dewatering of Fe(III)-Si gels. Similar to Fe(II), the Archaean ocean had significantly elevated concentrations of dissolved silica, at least as high as at saturation with cristobalite (0.67 mM at 40 o C in seawater), between 0.5 and 1.5 mM for Fe(III)-Si gels, depending on the Fe/Si ratio, or as high as 2.20 mM for amorphous silica [START_REF] Siever | The silica cycle in the Precambrian[END_REF][START_REF] Maliva | Secular change in the Precambrian silica cycle: Insights from chert petrology[END_REF]Konhauser et al., 2007a;[START_REF] Zheng | Abiologic silicon isotope fractionation between aqueous Si and Fe(III)-Si gel in simulated Archean seawater: Implications for Si isotope records in Precambrian sedimentary rocks[END_REF]. In contrast, modern seawater has dissolved silica typically <0.10 mM [START_REF] Tréguer | The silica balance in the world ocean: A reestimate[END_REF], reflecting uptake by silica-secreting organisms.

Alternative interpretations for the chert include formation (1) at the sediment-water interface as a replacement phase of a precursor sediment [START_REF] Krapež | Hydrothermal and resedimented origins of the precursor sediments to banded iron formations: Sedimentological evidence from the early Palaeoproterozoic Brockman Supersequence of Western Australia[END_REF], (2) within the sediment following desorption of silica that was adsorbed to the surface of Fe(III)-oxyhydroxides during the diagenetic reduction of Fe(III) by organic carbon (Fisher and Knoll, 2009), or (3) as a cement phase that infilled the pore spaces within primary iron-rich silicate muds [START_REF] Rasmussen | Precipitation of iron silicate nanoparticles in early Precambrian oceans marks Earth's first iron age[END_REF]. Given the diversity of chert occurrences among different IF, these interpretations are not mutually exclusive.

What Iron Formations Indicate About the Precambrian Environment

Hydrothermal sources of iron to the oceans

The abundance of IFs in Precambrian successions and their absence in Phanerozoic successions was used in early studies to argue for a largely anoxic Precambrian atmosphere and ocean system (e.g., [START_REF] Cloud | Paleoecological significance of banded iron-formation[END_REF][START_REF] Holland | Oceans -possible source of iron in iron-formations[END_REF][START_REF] Holland | The Chemical Evolution of the Atmosphere and Oceans[END_REF]. The accumulation of such large masses of iron (in the form of Superior-type BIF) required the transport of dissolved Fe(II), because Fe(III) is essentially insoluble at circumneutral pH values. Early studies invoked a continental source of iron, where Fe(II) would have been mobile during weathering in the absence of atmospheric O 2 (e.g., [START_REF] James | Sedimentary facies of iron-formation[END_REF][START_REF] Lepp | Origin of Precambrian iron formations[END_REF], as well as the expectation that early continents would have been more mafic than today, and hence richer in iron [START_REF] Condie | Chemical composition and evolution of the upper continental crust: Contrasting results from surface samples and shales[END_REF]. However, detailed studies of IF that followed in the Hamersley Province, Western Australia, suggested that the amount of iron deposited was on the order of 1 x 10 13 g/yr [START_REF] Trendall | The iron formations of the Precambrian Hamersley Group, Western Australia with special reference to the associated crocidolite[END_REF], requiring rivers the size of the modern Amazon to transport orders of magnitude more iron than they do today. Furthermore, IFs typically lack significant detrital input, which argued against a riverine source of the iron [START_REF] Holland | The Chemistry of the Atmosphere and Oceans[END_REF]. [START_REF] Holland | Oceans -possible source of iron in iron-formations[END_REF] had previously suggested that the iron was instead sourced from deep-marine, hydrothermally influenced waters and supplied to the depositional settings via upwelling processes.

With the 1977 discovery of modern seafloor-hydrothermal systems (e.g., [START_REF] Corliss | Chemistry of hydrothermal mounds near Galapagos Rift[END_REF], the observed high flux of dissolved iron derived from slow-spreading mid-ocean ridges (e.g., [START_REF] Saito | Slowspreading submarine ridges in the South Atlantic as a significant oceanic iron source[END_REF], and the recognition that even modern hydrothermal systems may contribute up to 75% of dissolved iron to the iron budget in the deep oceans [START_REF] German | Heat, volume and chemical fluxes from submarine venting: A synthesis of results from the Rainbow hydrothermal field, 36ºN MAR[END_REF][START_REF] Carazzo | The remarkable longevity of submarine plumes: Implications for the hydrothermal input of iron to the deep-ocean[END_REF][START_REF] German | Hydrothermal processes[END_REF], the interpretation that iron was sourced from the deep ocean became increasingly favoured. Moreover, it has been demonstrated that, even in the modern ocean, particulate and dissolved iron can be transported more than 2000 km from the ridge axis in seafloorhydrothermal settings (Fitzsimmons et al., 2017), and that variability in this input and ocean mixing regimes have influenced the extent of iron transport even in the more recent past [START_REF] Horner | Persistence of deeply sourced iron in the Pacific Ocean[END_REF]. Along similar lines, it has also been proposed that dissolved iron could have been supplied directly by hydrothermal plumes to the uppermost levels of the water column. Archaean ridge crests were likely shallower than today [START_REF] Isley | Hydrothermal plumes and the delivery of iron to banded iron formation[END_REF], reflecting higher mantle temperatures, and under such conditions, hydrothermal effluent could have risen buoyantly upwards through the water column to the photic zone -where various oxidative mechanisms could have occurred (see section 2.4). Although modern hydrothermal Fe(II) input concentrations (1-3 nM) are significantly below what would be required to form an IF, this system provides evidence that the long-distance transport of iron and the formation of Superior type IF is mechanistically possible. This would especially be the case under less oxidising conditions, such as those predicted for the Archaean and Palaeoproterozoic, favouring dissolved Fe(II) transport over even longer distances than those observed today. Moreover, the elevated temperatures of water-rock reactions in the Archaean and the low concentration of marine sulfate (which affect the Eh of hydrothermal fluids, as well as limit near-vent precipitation of iron sulfides) suggest that dissolved Fe(II) in hydrothermal effluents may have been orders of magnitude greater than today (e.g., [START_REF] Kump | Hydrothermal Fe fluxes during the Precambrian: Effect of low oceanic sulfate concentrations and low hydrostatic pressure on the composition of black smokers[END_REF]. Indeed, the concentration of dissolved Fe(II) effused from some modern deep-sea vents is as high as 25 mM/kg; a few, in arc-and back-arc settings have as much as 2500 mM [START_REF] German | Hydrothermal Processes[END_REF], highlighting the possibility that Archaean hydrothermal fluids could have been significantly enriched in dissolved Fe(II). Furthermore, the observation that Archaean shales are enriched in iron, relative to average Phanerozoic shales, may also indicate a larger hydrothermal flux of Fe(II) to the early oceans [START_REF] Kump | Iron in Precambrian rocks: Implications for the global oxygen budget of the ancient Earth[END_REF][START_REF] Bekker | Oxygen overshoot and recovery during the early Paleoproterozoic[END_REF], or greater delivery of iron-rich siliciclastic sediment due to the Archaean crust being more mafic (e.g., [START_REF] Dhuime | Emergence of continental crust about 3 billion years ago[END_REF][START_REF] Lee | Twostep rise in atmospheric oxygen linked to the growth of continents[END_REF].

Rare earth elements and yttrium (REE+Y) in hydrothermal fluids have characteristic distributions that provide additional evidence for a hydrothermal Fe(II) source for IFs, based on the assumption that their patterns record the source of iron due high partition coefficients of the REEs onto hydrothermal precipitates (e.g., [START_REF] Koeppenkastrop | Sorption of rare-earth elements from seawater onto synthetic mineral particles: An experimental approach[END_REF]. Modern seawater generally displays heavy REE (HREE) enrichments relative to light REE (LREE), as well as positive La, Gd, and Y anomalies, and strong negative Ce anomalies when normalised to shales (e.g., post-Archaean average shale -PAAS) (e.g., [START_REF] Elderfield | The rare earth elements in seawater[END_REF][START_REF] Taylor | The continental crust: Its composition and evolution: an examination of the geochemical record preserved in sedimentary rocks[END_REF][START_REF] De Baar | Rare earth elements in the Pacific and Atlantic Oceans[END_REF][START_REF] Derry | The chemical evolution of Precambrian seawater: Evidence from REEs in banded iron formations[END_REF][START_REF] Bau | Yttrium and holmium in South Pacific seawater: Vertical distribution and possible fractionation mechanisms[END_REF][START_REF] Alibo | Rare earth elements in seawater: particle association, shalenormalization, and Ce oxidation[END_REF][START_REF] Bolhar | Characterization of early Archaean chemical sediments by trace element signatures[END_REF]Planavsky et al., 2010). These patterns are typically mimicked in the REE+Y patterns of post-Archaean IF [START_REF] Laajoki | Rare-earth elements in Precambrian iron formations in Vaydankyla, South Puolanka area, Finland[END_REF]Fryer, 1977a, b;[START_REF] Graf | Rare earth elements, iron formations and sea water[END_REF][START_REF] Barrett | Two-stage hydrothermal formation of a Lower Proterozoic sediment-hosted massive sulfide deposit, northern Labrador Trough, Quebec[END_REF][START_REF] Klein | Geochemistry and sedimentology of a facies transition from limestone to iron-formation deposition in the early Proterozoic Transvaal Supergroup, South Africa[END_REF]. Iron formations older than 2.5 Ga display positive La, Gd, and Y anomalies as well, but lack regular negative Ce anomalies and instead contain large positive Eu anomalies (e.g., [START_REF] Alexander | Continentally-derived solutes in shallow Archean seawater: Rare earth element and Nd isotope evidence in iron formation from the 2.9 Ga Pongola Supergroup, South Africa[END_REF]. Estimates suggest that a 1:100 or 1:1000 mixture of modern seawater to hydrothermal fluid may account for the observed REE+Y patterns of IF [START_REF] Barrett | Two-stage hydrothermal formation of a Lower Proterozoic sediment-hosted massive sulfide deposit, northern Labrador Trough, Quebec[END_REF][START_REF] Dymek | Chemistry, petrology and origin of banded iron-formation lithologies from the 3800 Ma Isua Supracrustal Belt, West Greenland[END_REF][START_REF] Klein | Geochemistry and sedimentology of a facies transition from limestone to iron-formation deposition in the early Proterozoic Transvaal Supergroup, South Africa[END_REF].

Recent in-situ and fraction-specific REE+Y trends for the ca. 2.47 Ga Kuruman and Griquatown IFs show a marked difference in carbonate versus oxide mineral-hosted REE+Y patterns [START_REF] Oonk | Fraction-specific geochemistry across the Asbestos Hills BIF of the Transvaal Supergroup, South Africa: Implications for the origin of BIF and the history of atmospheric oxygen[END_REF]. The carbonate REE+Y patterns seem to mimic that of modern seawater even closer than do bulk-rock patterns. In addition, the changing steepness of the REE+Y spectra (Yb SN /Ce SN ) could be related to corresponding changes in depth of deposition, as suggested for modern seawater (e.g., [START_REF] Piepgras | The behavior of rare earth elements in seawater: Precise determination of variations in the North Pacific water column[END_REF][START_REF] Alibo | Rare earth elements in seawater: particle association, shalenormalization, and Ce oxidation[END_REF]. The oxide fraction is the most depleted in REE+Y and has a slight MREE enrichment [START_REF] Oonk | Fraction-specific geochemistry across the Asbestos Hills BIF of the Transvaal Supergroup, South Africa: Implications for the origin of BIF and the history of atmospheric oxygen[END_REF], which has been described from modern oxygen-limited ferrous environments (e.g., [START_REF] Haley | Rare earth elements in pore waters of marine sediments[END_REF].

Europium anomalies in REE+Y patterns have been central in tracing Fe sources. Europium enrichment in chemical sedimentary rocks that precipitated from seawater indicates a strong influence of high-temperature hydrothermal fluids on the seawater dissolved REE+Y load (e.g., [START_REF] Klinkhammer | Rare-earth elements in seawater near hydrothermal vents[END_REF][START_REF] Derry | Causes and consequences of mid-Proterozoic anoxia[END_REF]Jacobsen, 1988, 1990;[START_REF] Bau | Distribution of yttrium and rare-earth elements in the Penge and Kuruman iron-formations, Transvaal Supergroup, South Africa[END_REF][START_REF] Viehmann | Geochemistry of the Krivoy Rog Banded Iron Formation, Ukraine, and the impact of peak episodes of increased global magmatic activity on the trace element composition of Precambrian seawater[END_REF]. Positive Eu anomalies in hydrothermal fluids are linked to the breakdown of plagioclase minerals within the volcanic rocks underlying hydrothermal vents, as plagioclase is intrinsically enriched in Eu relative to other REE+Y during igneous crystallisation [START_REF] Schnetzler | Partition coefficients of rare-earth elements between igneous matrix material and rock-forming mineral phenocrysts-II[END_REF][START_REF] Graf | Rare earth elements as hydrothermal tracers during the formation of massive sulfide deposits in volcanic rocks[END_REF][START_REF] Graf | Rare earth elements, iron formations and sea water[END_REF][START_REF] Fryer | Archaean volcanogenic oceans[END_REF]. Furthermore, Eu remains in the soluble Eu 2+ state in hydrothermal fluids because Eu 2+ is more stable at higher temperatures [START_REF] Sverjensky | Europium redox equilibria in aqueous solution[END_REF][START_REF] Bau | Rare-earth element mobility during hydrothermal and metamorphic fluid-rock interaction and the significance of the oxidation state of europium[END_REF]. It is generally assumed that Fe and REE+Y will not be fractionated during transport from spreading ridges or other exhalative centres, owing to the high sorption capacity of Fe(III)oxyhydroxides; therefore, a large positive Eu anomaly (>2) indicates that the Fe in the precursor sediment was hydrothermally derived (e.g., [START_REF] Michard | The REE content of some hydrothermal fluids[END_REF][START_REF] Bau | Rare-earth element mobility during hydrothermal and metamorphic fluid-rock interaction and the significance of the oxidation state of europium[END_REF][START_REF] Danielson | The europium anomalies in banded iron formations and the thermal history of the oceanic crust[END_REF][START_REF] Slack | Suboxic deep seawater in the late Paleoproterozoic: Evidence from hematitic chert and iron formation related to seafloorhydrothermal sulfide deposits, central Arizona, USA[END_REF]Slack et al., , 2009)). However, Eu concentrations can also be elevated by the erosion of Eu-enriched Archaean continental crust [START_REF] Taylor | The continental crust: Its composition and evolution: an examination of the geochemical record preserved in sedimentary rocks[END_REF]Bohlar et al., 2004), and anomalies generated in vent systems may be diluted by mixing of seawater and hydrothermal fluids [START_REF] Olivarez | The europium anomaly of seawater: Implications for fluvial versus hydrothermal REE inputs to the oceans[END_REF]. Nonetheless, secular trends in the magnitude of Eu anomalies in large Superior-type IFs have historically been assumed to record variations in hydrothermal flux (e.g., Fryer, 1977a;[START_REF] Derry | The chemical evolution of Precambrian seawater: Evidence from REEs in banded iron formations[END_REF][START_REF] Danielson | The europium anomalies in banded iron formations and the thermal history of the oceanic crust[END_REF][START_REF] Bau | Rare-earth element systematics of the chemically precipitated component in early Precambrian iron formations and the evolution of the terrestrial atmospherehydrosphere-lithosphere system[END_REF][START_REF] Sreenivas | Emerging views on the evolution of atmospheric oxygen during the Precambrian[END_REF][START_REF] Viehmann | Geochemistry of the Krivoy Rog Banded Iron Formation, Ukraine, and the impact of peak episodes of increased global magmatic activity on the trace element composition of Precambrian seawater[END_REF]. This trend is clearly demonstrated through a compilation of chondrite-normalised Eu anomaly data (Eu CN /Eu CN * ) in IF spanning the Precambrian (Fig. 5). The magnitude of positive Eu anomalies steadily decreases from the Eoarchaean, which reflects both high hydrothermal input and low dissolved oxygen required to stabilise Eu 2+ for transport before the deep oceans became oxygenated [START_REF] Viehmann | Geochemistry of the Krivoy Rog Banded Iron Formation, Ukraine, and the impact of peak episodes of increased global magmatic activity on the trace element composition of Precambrian seawater[END_REF]. The very high peak between 2.7 and 2.5 Ga likely corresponds to a very high mantle input of hydrothermally-derived iron at that time.

Positive Y anomalies are typical of aqueous systems [START_REF] Bau | Rare-earth element systematics of the chemically precipitated component in early Precambrian iron formations and the evolution of the terrestrial atmospherehydrosphere-lithosphere system[END_REF]. Aqueous fluids and their precipitates display Y/Ho ratios significantly higher than the molar ratio of 52 shown by common igneous rocks and epiclastic sediments (Bau, 1996, and references therein), reaching molar ratios up to 90 [START_REF] Zhang | The comparative behaviors of yttrium and lanthanides in the seawater of the North Pacific[END_REF][START_REF] Bau | Sources of rare-earth elements and iron in Paleoproterozoic iron-formations from the Transvaal Supergroup, South Africa: Evidence from neodymium isotopes[END_REF][START_REF] Nozaki | The fractionation between Y and Ho in the marine environment[END_REF]. The Y/Ho ratio in chemical precipitates is also typically depressed by clastic input because crustal rocks have lower Y/Ho ratios [START_REF] Kamber | The geochemistry of late Archaean microbial carbonate: Implications for ocean chemistry and continental erosion history[END_REF][START_REF] Bolhar | Characterization of early Archaean chemical sediments by trace element signatures[END_REF]. Similarly, [START_REF] Viehmann | Geochemistry of the Krivoy Rog Banded Iron Formation, Ukraine, and the impact of peak episodes of increased global magmatic activity on the trace element composition of Precambrian seawater[END_REF] noted the depression of the La anomaly in IF with increased detrital contamination; La anomalies are strongly expressed in pure IF (La SN /La SN * ratios between 1.32 and 1.80), whereas they are small or absent in impure IF (between 0.85 and 1.52). Detrital contamination would also act to decrease the slope of the REE+Y pattern. Typically, HREE are enriched in the oxide facies of IF, as the result of interaction between hydrothermal fluids and mid-ocean ridge basalt (MORB) at spreading centres, which are typically tholeiitic and depleted in LREE [START_REF] Derry | The chemical evolution of Precambrian seawater: Evidence from REEs in banded iron formations[END_REF][START_REF] Bau | Rare-earth element mobility during hydrothermal and metamorphic fluid-rock interaction and the significance of the oxidation state of europium[END_REF].

Sources of REE+Y and the fidelity of their signals have been investigated using the 138 La-138 Ce and 147 Sm-143 Nd isotope systems (e.g., [START_REF] Tanaka | Combined La-Ce and Sm-Nd isotope systematics in petrogenetic studies[END_REF][START_REF] Derry | The chemical evolution of Precambrian seawater: Evidence from REEs in banded iron formations[END_REF][START_REF] Shimizu | La-Ce and Sm-Nd systematics of siliceous sedimentary rocks: A clue to marine environment in their deposition[END_REF][START_REF] Hayashi | Origin of negative Ce anomalies in Barberton sedimentary rocks, deduced from La-Ce and Sm-Nd isotope systematics[END_REF][START_REF] Viehmann | Geochemistry of the Krivoy Rog Banded Iron Formation, Ukraine, and the impact of peak episodes of increased global magmatic activity on the trace element composition of Precambrian seawater[END_REF]Haugaard et al., 2016). Both Ce and Nd have short residence times in the modern ocean, 90-165 and 1000-1500 years, respectively [START_REF] Amakawa | Precise determination of variations in the 138 Ce/ 142 Ce ratios of marine ferromanganese nodules[END_REF], and in the Archaean oceans these elements likely had a similar residence times on the order of 10 3 years, suggesting that Ce and Nd isotope compositions would be spatially heterogeneous. In the Archaean and Proterozoic oceans, ε Nd (t) values (which describe the deviation of the 143 Nd/ 144 Nd ratio measured in a sample relative to the 143 Nd/ 144 Nd ratio in a chondritic uniform reservoir at time t and in parts per 10 4 ) are estimated to have been +1 to +2 for deep waters dominated by hydrothermal sources, whereas values down to -3 would have been typical of shallow-waters dominated by terrestrial sources [START_REF] Miller | Sources of Precambrian chemical and clastic sediments[END_REF]Jacobsen andPimentel-Klose, 1988a, 1988b;[START_REF] Alexander | Continentally-derived solutes in shallow Archean seawater: Rare earth element and Nd isotope evidence in iron formation from the 2.9 Ga Pongola Supergroup, South Africa[END_REF]. The relative contribution of continentally-sourced versus mantle-sourced REE+Y, discussed in more detail in the next section, has been estimated using Ce and Nd isotope compositions (Piepgras andWasserburg, 1980, Bau et al., 1997;[START_REF] Shimizu | Sources of iron formations in the Archean Isua and Malene supracrustals, West Greenland: Evidence from La-Ce and Sm-Nd isotopic data and REE abundances[END_REF]1991;[START_REF] Alexander | Continentally-derived solutes in shallow Archean seawater: Rare earth element and Nd isotope evidence in iron formation from the 2.9 Ga Pongola Supergroup, South Africa[END_REF][START_REF] Heubeck | An early ecosystem of Archean tidal microbial mats (Moodies Group, South Africa, ca. 3.2 Ga)[END_REF][START_REF] Viehmann | Geochemistry of the Krivoy Rog Banded Iron Formation, Ukraine, and the impact of peak episodes of increased global magmatic activity on the trace element composition of Precambrian seawater[END_REF]. In some cases, 138 La-138 Ce systematics show that Ce anomalies in IFs are a feature produced by recent hydrothermal events and are not intrinsic to ancient seawater [START_REF] Hayashi | Origin of negative Ce anomalies in Barberton sedimentary rocks, deduced from La-Ce and Sm-Nd isotope systematics[END_REF]. Similarly, young 147 Sm-143 Nd isotope ages indicate post-depositional REE+Y mobility during metamorphic events (e.g., [START_REF] Shimizu | Sources of iron formations in the Archean Isua and Malene supracrustals, West Greenland: Evidence from La-Ce and Sm-Nd isotopic data and REE abundances[END_REF]. Such results highlight the need to consider the possibility of post-depositional REE+Y mobility in some IFs, particularly those that have been subjected to metamorphic and supergene alteration.

Building on the premise of a hydrothermal iron source, [START_REF] Isley | Plume-related mafic volcanism and the deposition of banded iron formation[END_REF] suggested a direct link between mantle plume activity between 3.8 and 1.8 Ga and global IF deposition. It is now widely accepted that the deposition of large, economically important IF coincided in time with mantle plume breakout events, as recorded by the secular distribution of large igneous provinces (LIPs), dike swarms, and submarine-emplaced mafic volcanic rocks (e.g., [START_REF] Isley | Plume-related mafic volcanism and the deposition of banded iron formation[END_REF][START_REF] Rasmussen | Deposition of 1.88-billion-year-old iron formations as a consequence of rapid crustal growth[END_REF], as well as that of major mantle plume events [START_REF] Viehmann | Geochemistry of the Krivoy Rog Banded Iron Formation, Ukraine, and the impact of peak episodes of increased global magmatic activity on the trace element composition of Precambrian seawater[END_REF]. LIPs are linked to relatively short-lived (<10-20 Myr) igneous events with magma having been produced in the mantle, resulting in relatively rapid intrusion and eruption of high volumes of mafic to ultramafic magma [START_REF] Coffin | Large igneous provinces: Crustal structure, dimensions, and external consequences[END_REF][START_REF] Ernst | Large mafic magmatic events through time and links to mantleplume heads[END_REF]. Greater oceanic spreading rates, increased submarine and subaerial volcanic activity, high sea-level, greenhouse conditions, and an enhanced production of VMS are all consequences of mantle plume breakout events (e.g., [START_REF] Condie | Precambrian superplumes and supercontinents: A record in black shales, carbon isotopes and paleoclimates?[END_REF][START_REF] Barley | Emplacement of a large igneous province as a possible cause of banded iron formation 2.45 billion years ago[END_REF][START_REF] Barley | Late Archean to early Paleoproterozoic global tectonics, environmental change and the rise of atmospheric oxygen[END_REF][START_REF] Bekker | Iron formation: The sedimentary product of a complex interplay among mantle, tectonic, oceanic, and biospheric processes[END_REF][START_REF] Bekker | Iron formations: Their origins and implications for ancient seawater chemistry[END_REF].

Continental sources of iron to the oceans

Compilations of Nd isotope data from Archaean to Palaeoproterozoic IFs highlight the fact that most IFs have ε Nd (t) values that are intermediate between those of the depleted mantle and ancient continental crust [START_REF] Alexander | Neodymium isotopes in Archean seawater and implications for the marine Nd cycle in Earth's early oceans[END_REF][START_REF] Viehmann | Geochemistry of the Krivoy Rog Banded Iron Formation, Ukraine, and the impact of peak episodes of increased global magmatic activity on the trace element composition of Precambrian seawater[END_REF]. Assuming that the sources of REE+Y closely track the sources of Fe(II), Nd isotope data for IFs suggest a significant contribution from "continental" iron [START_REF] Haugaard | Petrology and geochemistry of the ~2.9 Ga Itilliarsuk banded iron formation and associated supracrustal rocks, West Greenland: Source characteristics and depositional environment[END_REF]Li et al., 2015). Based on Nd-Fe isotope variations in the ca. 2.5 Ga Brockman Iron Formation (Fig. 6), Li et al. (2015) modelled the mixing between a hydrothermal Fe(II) source (high δ 56 Fe, ε Nd (t)) and a continental Fe(II) source (low δ 56 Fe, ε Nd (t)).

The high δ 56 Fe values of the hydrothermal source are interpreted to reflect partial oxidation of hydrothermally-derived Fe(II), whereas the low δ 56 Fe values of the continental source are attributed to microbial dissimilatory iron reduction (DIR) that released dissolved Fe(II) from pore fluids in continental margin sediments that was transported to the deep basin. Basin-wide transport of DIRgenerated dissolved Fe(II) has been referred to as a "benthic iron shuttle" [START_REF] Anderson | Sources and mechanisms for the enrichment of highly reactive iron in euxinic Black Sea sediments[END_REF], and has been inferred to have occurred in Archaean and Proterozoic basins [START_REF] Severmann | Modern iron isotope perspective on the benthic iron shuttle and the redox evolution of ancient oceans[END_REF].

A benthic iron shuttle predicts that proximal sediments should have relatively high δ 56 Fe values, reflecting the residual iron pool after export of low-δ 56 Fe dissolved Fe(II). Support for this model comes from the Kuruman Iron Formation, where near-shore samples have consistently higher δ 56 Fe values than more distal samples [START_REF] Heimann | Fe, C, and O isotope compositions of banded iron formation carbonates demonstrate a major role for dissimilatory iron reduction in ~ 2.5 Ga marine environments[END_REF]. In terms of the Brockman Iron Formation, which is correlative with the Kuruman Iron Formation [START_REF] Beukes | Origin and paleoenvironmental significance of major iron formations at the Archean-Paleoproterozoic boundary[END_REF], the observed co-variations in δ 56 Fe -ε Nd (t) values (Fig. 6) lends itself to differing interpretations, even amongst the co-authors. On the one hand, the scatter in the data might indicate that REEs and Fe isotopes are decoupled, and hence, given the range of potential Fe and REE pathways and variable basin geometry, the relations in Fig. 6 do not neccessarily support a mixture of hydrothermal and benthic Fe sources. On the other hand, some scatter is expected given the range in δ 56 Fe values for the hydrothermal component (recording differing extents of oxidation), and the range in ε Nd (t) values for the continental endmember, and perhaps the focus of interpretation is the fact that the observed data do not lie near the "L-shaped" trend that is required by a strictly hydrothermal model, such as that proposed by [START_REF] Rouxel | Iron isotope constraints on the Archean and Paleoproterozoic ocean redox state[END_REF], who argued that the negative δ 56 Fe values of IFs reflect a residual iron component after extensive oxidation.

The broader database of Nd isotope compositions reported from the Hamersley Basin (Jacobsen and Pimentel-Klose, 1998a;[START_REF] Alibert | Rare element and neodymium isotopic compositions of the banded iron-formations and associated shales from Hamersley, Western Australia[END_REF]Li et al., 2015)

Plausible primary iron mineralogy

Iron formations are defined by their unusual mineralogy, where silica and a wide range of iron-rich (but aluminium-poor) minerals constitute the majority. The notion that some of the Fe(III) mineral phases (hematite, magnetite) are the primary precursor phase for IF was suggested over four decades ago. [START_REF] Ayres | Genesis of iron-bearing minerals in banded iron formation mesobands in the Dales Gorge Member, Hamersley Group, Western Australia[END_REF] documented the presence of possible original granular textures, including microspheres of 5-15 µm in diameter, in several of the hematite-rich mesobands of the Dales Gorge Member. Also present within the hematite layers are chert globules containing hematite inclusions. [START_REF] Trendall | The iron formations of the Precambrian Hamersley Group, Western Australia with special reference to the associated crocidolite[END_REF] reported "jaspery platelets" of hematite (2-10 µm in diameter) within chert that they argued was paragenetically the oldest of the hematite grains observed in the Dales Gorge Member. [START_REF] Ahn | Hematite nanospheres of possible colloidal origin from a Precambrian banded iron formation[END_REF] observed 120-200 nm hematite nanospheres in the underlying Marra Mamba Formation, which they attributed to structural ordering that accompanied dehydration of colloidal Fe(III) oxyhydroxide particles. These spheroids are similar to hematite spheroids found in the Quaternary hydrothermal fields of Milos Island in Greece [START_REF] Fru | Fossilized iron bacteria reveal a pathway to the biological origin of banded iron formation[END_REF] and in modern iron oxide deposits of the Red Sea (Taitel-Goldman and Singer, 2002), and may record early diagenetic transformations of original Fe(III)-oxyhydroxide particles. In their review of IF genesis, [START_REF] Beukes | Origin and paleoenvironmental significance of major iron formations at the Archean-Paleoproterozoic boundary[END_REF] stressed that the earliest oxide minerals in Superior-type IFs was fine-grained hematite dispersed throughout the chert. An early paragenesis for hematite is supported by the recent study of [START_REF] Sun | Primary hematite in Neoarchean to Paleoproterozoic oceans[END_REF], who conducted highresolution transmission electron microscopy (TEM) and selected area electron diffraction (SAED)

analyses to ascertain the timing and paragenesis of hematite within fine mineral structures in IF from the 2.73 Ga Abitibi greenstone belt in northeastern Canada and the Kuruman Iron Formation.

Those authors interpreted the 3-5 nm ultrafine hematite crystals in the hematite bands, as well as submicrometer (sub)euhedral hematite crystals randomly distributed in the chert matrix of transitional zones between iron oxide-and chert-rich bands, to have formed directly by dehydration from a precursor Fe(III)-oxyhydroxide phase.

Although the case for primary hematite discussed above is strong in some examples, hematite in IFs is typically secondary in origin, linked with iron enrichment leading to ore genesis, even in sections distal to iron ore bodies. Iron formations that contain hematite and martite-goethite are some of the largest known ore bodies. Their origin remains controversial, but hydrothermal processes are well documented [START_REF] Taylor | Genesis of highgrade hematite orebodies of the Hamersley Province, Western Australia[END_REF][START_REF] Thorne | Oxygen isotope compositions of iron oxides from high-grade BIF-hosted iron ore deposits of the central Hamersley Province, Australia: Constraints on the evolution of hydrothermal fluids[END_REF]. Ore hematite is typically prismatic or platy, and commonly occurs with bladed carbonate minerals, indicating late-stage carbonate recrystallisation, in contrast to finer-grained hematite that characterises occurrences distal to ore. Interpretation of platy hematite as depositionally or diagenetically early is suspect, however, because such a morphology is the typical form of ore-generation hematite. Indeed, geochronological studies of xenotime and monazite paragenesis within hematite and martite-goethite ore bodies that formed during hydrothermal alteration of IF indicate that iron-ore formation was a long-lived and multi-stage process [START_REF] Rasmussen | Prolonged history of episodic fluid flow in giant hematite ore bodies: Evidence from in situ U-Pb geochronology of hydrothermal xenotime[END_REF]. Within the iron ore bodies of the Hamersley Province, hematite-carbonate alteration occurred at temperatures of ~250°C under high fluid pressures [START_REF] Taylor | Genesis of highgrade hematite orebodies of the Hamersley Province, Western Australia[END_REF].

Silica concentrations in Precambrian seawater are estimated to have been high (~1 to 2 mM), assuming control by the solubility of amorphous silica and sorption to silicate minerals [START_REF] Maliva | Secular change in the Precambrian silica cycle: Insights from chert petrology[END_REF], raising the possibility that the primary Fe(III) minerals were enriched in SiO 2 . [START_REF] Percak-Dennett | Iron isotope fractionation during dissimilatory iron reduction under simulated Archean conditions[END_REF] and [START_REF] Wu | Stable iron isotope fractionation between aqueous Fe(II) and model Archean ocean Fe-Si coprecipitates and implications for iron isotope variations in the ancient rock record[END_REF] studied dissolved Fe(II) oxidation in the presence of dissolved silica, which produced an Fe(III)-Si gel that was homogenous on the nm scale, indicating the presence of a single phase in which Si and Fe are tightly bonded and distinct from, for example, Si sorbed to ferrihydrite, as has been previously described from terrestrial hot springs (e.g., [START_REF] Lalonde | Chemical reactivity of microbe and mineral surfaces in hydrous ferric oxide depositing hydrothermal springs[END_REF]. Previous spectroscopic studies indicate that co-precipitation of Fe(III) and Si from a mixed solution tends to form an intimately bonded Fe(III)-Si phase, rather than separate Fe(III) and Si phases [START_REF] Doelsch | Speciation and crystal chemistry of iron(III) chloride hydrolyzed in the presence of SiO 4 ligands. 1. An Fe K-Edge EXAFS study[END_REF][START_REF] Doelsch | Speciation and crystal chemistry of Fe(III) chloride hydrolyzed in the presence of SiO 4 ligands. 2. Characterization of SiFe aggregates by FTIR and 29 Si Solid-State NMR[END_REF][START_REF] Doelsch | Chemistry and structure of colloids obtained by hydrolysis of Fe(III) in the presence of SiO 4 ligands[END_REF]. Identifying the specific primary Fe(III) precipitate in IFs is important, because when exposed to dissolved Fe(II), Fe(III)-Si gels are resistant to phase transformations that may occur in dissolved Fe(II)-ferrihydrite experiments (e.g., [START_REF] Johnson | Experimental constraints on Fe isotope fractionation during magnetite and Fe carbonate formation coupled to dissimilatory hydrous ferric oxide reduction[END_REF], and uptake of Fe(II) into the solid is inhibited under abiological conditions [START_REF] Zheng | Abiologic silicon isotope fractionation between aqueous Si and Fe(III)-Si gel in simulated Archean seawater: Implications for Si isotope records in Precambrian sedimentary rocks[END_REF]. [START_REF] Delstanche | Silicon isotopic fractionation during adsorption of aqueous monosilicic acid onto iron oxide[END_REF]. Recent experimental studies, however, document large Si isotope fractionations between dissolved Si and Fe(III)-Si gel, significantly larger than those produced by sorption of dissolved Si to Fe(III)-oxyhydroxides [START_REF] Zheng | Abiologic silicon isotope fractionation between aqueous Si and Fe(III)-Si gel in simulated Archean seawater: Implications for Si isotope records in Precambrian sedimentary rocks[END_REF], thus providing an explanation for the highly negative δ 30 Si values observed in some IF cherts (Fig. 7). Such observations strongly support the idea that the primary Fe(III)

precipitate in IFs was an Fe(III)-Si gel. Moreover, [START_REF] Reddy | Silicon isotope fractionation during microbial reduction of Fe(III)-Si gels under Archean seawater conditions and implications for iron formation genesis[END_REF] found that even larger Si isotope fractionations occur upon microbial reduction of Fe(III)-Si gels, which may explain the tendency of magnetite-rich IFs to have lower δ 30 Si values than hematite-rich IFs (Fig. 7).

The paragenesis of magnetite in IFs is quite variable. Formation through dewatering of primary Fe(OH) 2 and Fe(OH) 3 precipitates has been proposed as the origin of disseminated finegrained magnetite [START_REF] Klein | Some Precambrian banded iron-formations (BIFs) from around the world: Their age, geologic setting, mineralogy, metamorphism, geochemistry, and origin[END_REF], but most workers consider magnetite to have formed through diagenetic and metamorphic reactions and there is in cases compelling petrographic evidence for a metamorphic origin (see below). In the case of diagenesis, Y-L. Li et al. (2013) proposed a biological model to account for its formation via a three-stage sequence, beginning with magnetite formation by DIR within an initial ferrihydrite-rich sediment, coupled to the oxidation of lysed phytoplankton biomass. DIR was followed by magnetite crystal aging, and ultimately pressure-temperature-induced abiotic growth of the pre-existing biogenic magnetite during metamorphism.

In support of the notion that magnetite nucleated during diagenesis, [START_REF] Li | Mineral ecophysiological data provide growing evidence for microbial activity in banded-iron formations[END_REF] conducted detailed crystallochemical analysis of magnetite from the Dales Gorge Member and reported a lattice constant and Fe 2+ /Fe 3+ stoichiometry very similar to those produced by DIR bacteria including Geobacter, Shewanella, and Thermoanaerobacter. Interestingly, [START_REF] Halama | Evaluation of siderite and magnetite formation in BIFs by pressure-temperature experiments of Fe(III) minerals and microbial biomass[END_REF] experimentally simulated the metamorphism of precursor and diagenetic iron-rich minerals (ferrihydrite, goethite, hematite) under low-grade metamorphic conditions (170°C, 1.2 kbar) for 14 days, and showed that thermochemical magnetite formation was limited to samples containing ferrihydrite and glucose, or goethite and glucose. No magnetite was formed from Fe(III) minerals when microbial biomass was present as the carbon and electron sources for thermochemical Fe(III) reduction, suggesting that biomass-derived organic molecules were bound to the mineral surfaces and prevented, or at least inhibited, solid-state conversion of ferric iron phases to magnetite.

Petrographic evidence certainly implies a secondary origin for magnetite in IFs [START_REF] Ewers | Studies of the Dales Gorge Member of the Brockman Iron Formation, Western Australia[END_REF]. Magnetite is a common replacement or overgrowth of early hematite and siderite, and of most rhombic, iron-rich carbonate. For instance, [START_REF] Krapež | Hydrothermal and resedimented origins of the precursor sediments to banded iron formations: Sedimentological evidence from the early Palaeoproterozoic Brockman Supersequence of Western Australia[END_REF] reported several forms of diagenetic and metamorphic magnetite, including (1) disseminated grains within, but obscuring, sedimentary laminae; (2) laminated beds that clearly truncate sedimentary layering; (3) layerdiscordant veins; and (4) infilling of cleavage. [START_REF] Ayres | Genesis of iron-bearing minerals in banded iron formation mesobands in the Dales Gorge Member, Hamersley Group, Western Australia[END_REF] reported that magnetite euhedra predate rhombic carbonate, but post-date matrix chert. In contrast, laminated bands of magnetite can post-date burial stylolites and destroy primary layering; [START_REF] Krapež | Hydrothermal and resedimented origins of the precursor sediments to banded iron formations: Sedimentological evidence from the early Palaeoproterozoic Brockman Supersequence of Western Australia[END_REF] referred to such laminated bands as iron proto-ore. Discordant veins, laminae, and cleavage fillings of magnetite cut these bands. [START_REF] Beukes | Origin and paleoenvironmental significance of major iron formations at the Archean-Paleoproterozoic boundary[END_REF] documented the occurrence of magnetite as a replacement of hematite-chert pods, and proposed formation through the addition of Fe(II) aq to ferrihydrite.

Finally, magnetite is a common ore mineral, showing preserved transformations from proto-ore to ore magnetite, ore hematite to ore magnetite, and ore magnetite to ore hematite.

The genetic relations between hematite and magnetite in IFs are well traced using O and Fe isotope compositions (Fig. 8). Bulk mineral separates of hematite, magnetite, and quartz from the Dales Gorge Member have average δ 18 O values of ~ -2‰, ~ +3‰, and ~ +21‰, respectively, which produces quartz-hematite and quartz-magnetite 18 O/ 16 O fractionations of ~ +23‰ and +18‰, respectively, indicating that on average, hematite formed at a lower temperature than magnetite [START_REF] Becker | Oxygen isotope study of a Precambrian banded ironformation, Hamersley Range, Western Australia[END_REF]. In situ oxygen isotope analysis, however, indicates significantly larger ranges in δ 18 O values, which correlate with petrographic relationships and chemical compositions (W. Li et al., 2013a). Fine-grained and dispersed hematite, considered to reflect the earliest generation, have δ 18 O values as low as -9‰ as determined by in situ analysis, indicating formation temperatures as low as ~25 o C (Fig. 8). The earliest generation of magnetite in the Dales Gorge Member contains minimal Si content [START_REF] Huberty | Silician magnetite from the Dales Gorge Member of the Brockman Iron Formation, Hamersley Group, Western Australia[END_REF], and in situ analysis shows that this magnetite has δ 18 O values of ca. -6 to +4‰, indicating formation temperatures between ~50 o C and ~100 o C (Fig. 8). This is consistent with an early diagenetic origin for the low-Si magnetite.

Significantly higher δ 18 O values observed for Si-bearing magnetite, as well as for coarse, euhedral magnetite, indicate higher formation temperatures up to ~250 o C (Fig. 8). As discussed in W. Li et al. (2013a), the very low diffusivity of oxygen in hematite makes this mineral relatively resistant to post-formation oxygen isotope exchange with fluids or quartz, whereas magnetite and quartz will more readily exchange with each other, or with a fluid, at elevated temperatures.

In contrast to the systematic changes in δ 18 O values with increasing formation temperature for hematite and magnetite, δ 56 Fe values should remain relatively unchanged due to the low solubility of iron in aqueous solutions at hydrothermal temperatures [START_REF] Chou | Solubility of magnetite in supercritical chloride solutions[END_REF][START_REF] Simon | Magnetite solubility and iron transport in magmatic-hydrothermal environments[END_REF], slow diffusion of Fe in hematite and magnetite [START_REF] Hallstrom | Modelling of iron diffusion in the iron oxides magnetite and hematite with variable stoichiometry[END_REF], and lack of a large external reservoir for equilibration, in contrast to quartz for oxygen isotope exchange. The spread in δ 56 Fe values for hematite and magnetite, as determined by in situ methods, is therefore likely to record original values upon increase in temperature (W. Li et al., 2013a), allowing inferences to be made regarding the sources of iron (e.g., hydrothermal versus DIR; Fig.

8).

Iron-bearing carbonates, which we broadly refer to as "siderite" are the most common carbonate minerals in IF, and include carbonates along the magnesite-siderite (MgCO 3 -FeCO 3 ) and dolomite-ankerite (Ca 0.5 Mg 0.5 CO 3 -Ca 0.5 Fe 0.5 CO 3 ) solid-solution series; the former having an average composition of Mg 0.2 Fe 0.8 CO 3 [START_REF] Klein | Some Precambrian banded iron-formations (BIFs) from around the world: Their age, geologic setting, mineralogy, metamorphism, geochemistry, and origin[END_REF]. Siderite typically occurs as minute, single rhomb-shaped crystals or massive layers of microcrystalline aggregates (James, 1951;[START_REF] Trendall | The iron formations of the Precambrian Hamersley Group, Western Australia with special reference to the associated crocidolite[END_REF]Ayers, 1972). In some beds, siderite may constitute >50% of the rock, and such occurrences are often classified as "siderite facies IF". Small siderite nodules are common in some layers, where they display evidence of differential compaction, thus indicating that the nodules probably formed prior to deep burial. The texturally earliest siderite is present as minute (<20 µm)

spheroids, which occur in some IFs and intercalated mudstones from the Hamersley Province [START_REF] Ayres | Genesis of iron-bearing minerals in banded iron formation mesobands in the Dales Gorge Member, Hamersley Group, Western Australia[END_REF]. The siderite appears to have replaced or overgrown an earlier phase (e.g., ferric oxyhydroxides), which occurs in the core of some crystals. Consistent with this suggestion, it has been shown that in the 1.88 Ga Gunflint Iron Formation in Ontario (Canada) some siderite rhombohedra contain spheres in their cores and a clear progression exists from siderite spheres to siderite rhombohedra without any trace of spheres remaining [START_REF] Laberge | A model for the biological precipitation of Precambrian iron-formations-A: Geological evidence[END_REF][START_REF] Carrigan | Petrological and stable isotope studies of carbonate and sulfide minerals from the Gunflint Formation, Ontario: Evidence for the origin of early Proterozoic iron-formation[END_REF]. This progression was also found in the Dales Gorge Member, where siderite spheres and rhombohedra occur near each other. Such features suggest that some siderite formed during the diagenetic (biological or non-biological), and subsequent metamorphic, oxidation of microbial biomass to CO 2 /bicarbonate coupled to the reduction of Fe(III), followed by reaction of the product Fe 2+ aq with sediment pore-water bicarbonate (Johnson et al., 2008;[START_REF] Heimann | Fe, C, and O isotope compositions of banded iron formation carbonates demonstrate a major role for dissimilatory iron reduction in ~ 2.5 Ga marine environments[END_REF][START_REF] Köhler | Biological carbon precursor to diagenetic siderite with spherical structures in iron formations[END_REF][START_REF] Halama | Evaluation of siderite and magnetite formation in BIFs by pressure-temperature experiments of Fe(III) minerals and microbial biomass[END_REF]. Importantly, in most cases, the C, O, Fe, and Sr isotope compositions of IF carbonates clearly preclude formation in equilibrium with seawater [START_REF] Johnson | Iron formation carbonates: Paleoceanographic proxy or recorder of microbial diagenesis?[END_REF]. With that said, although petrographically most siderite (and ankerite) exhibit recrystallised and, therefore, apparently secondary textures, their geochemistry does indicate that some potentially "primary" signals can be retained in IF carbonates. Iron isotopes of extracted carbonate fraction remained unaffected by low-grade metamorphism [START_REF] Frost | Preservation of Fe isotope heterogeneities during diagenesis and metamorphism of banded iron formation[END_REF] and carbonate-specific REE+Y patterns, both in situ [START_REF] Alibert | Rare earth elements in Hamersley BIF minerals[END_REF] and extracted as bulk [START_REF] Oonk | Fraction-specific geochemistry across the Asbestos Hills BIF of the Transvaal Supergroup, South Africa: Implications for the origin of BIF and the history of atmospheric oxygen[END_REF], mimic REE+Y aspects of present-day seawater.

Despite the consensus amongst workers that the precursor phase during the sedimentation of IF was most likely Fe(III)-oxyhydroxides (e.g., [START_REF] Konhauser | Could bacteria have formed the Precambrian banded iron formations?[END_REF][START_REF] Kappler | Deposition of banded iron formations by anoxygenic phototrophic Fe(II)-oxidizing bacteria[END_REF][START_REF] Posth | Biogenic Fe(III) minerals: From formation to diagenesis and preservation in the rock record[END_REF][START_REF] Sun | Primary hematite in Neoarchean to Paleoproterozoic oceans[END_REF] or an Fe(III)-Si gel (e.g., [START_REF] Percak-Dennett | Iron isotope fractionation during dissimilatory iron reduction under simulated Archean conditions[END_REF][START_REF] Zheng | Abiologic silicon isotope fractionation between aqueous Si and Fe(III)-Si gel in simulated Archean seawater: Implications for Si isotope records in Precambrian sedimentary rocks[END_REF], there is renewed interest in possible alternative phases that may have influenced or been important for BIF sedimentation. These alternative authigenic phases include dominantly Fe(II) or mixed-valence phases, such as greenalite and various "green-rust" phases.

The possible significance of greenalite versus other Fe(II) phases has been discussed

previously [START_REF] Eugster | The depositional environments of Precambrian banded ironformations[END_REF]. Greenalite has been identified in chert layers from several Palaeoproterozoic IFs, including the Brockman Iron Formation in Western Australia (Kaufman et al., 1990;[START_REF] Rasmussen | Precipitation of iron silicate nanoparticles in early Precambrian oceans marks Earth's first iron age[END_REF][START_REF] Alibert | Rare earth elements in Hamersley BIF minerals[END_REF] and the Transvaal Supergroup in South Africa [START_REF] Rasmussen | Greenalite precipitation linked to the deposition of banded iron formations downslope from a late Archean carbonate platform[END_REF]. The identification of fine-grained greenalite nanoparticles in laminated cherts [START_REF] Rasmussen | Dust to dust: Evidence for the formation of "primary" hematite dust in banded iron formations via oxidation of iron silicate nanoparticles[END_REF][START_REF] Rasmussen | Greenalite precipitation linked to the deposition of banded iron formations downslope from a late Archean carbonate platform[END_REF] has been taken as evidence for water column precipitation of this mineral, and by extrapolation, as a critical precursor to hematite and magnetite. However, at least in the case of the nanoparticles examined by [START_REF] Rasmussen | Greenalite precipitation linked to the deposition of banded iron formations downslope from a late Archean carbonate platform[END_REF], the succession examined more likely represents distal deep-water deposits not characteristic of IF deposited on a shelf. Moreover, ascribing a precursor phase containing significant ferrous iron requires a diagenetic mechanism to convert the Fe(II) on a layer-by-layer basis to Fe(III). [START_REF] Rasmussen | Replacement origin for hematite in 2.5 Ga banded iron formation: Evidence for postdepositional oxidation of iron-bearing minerals[END_REF] have argued that this oxidation, in the case of the Dales Gorge Member, occurred in association with a major collisional event at ca. 2.2 Ga when oxygenated and heated groundwaters (>200-400 o C) migrated along sedimentary layers of pre-existing BIF sediment [START_REF] Rasmussen | Isotopic dating of the migration of low-grade metamorphic front during orogenesis[END_REF]. Such an interpretation, however, is in conflict with the very low δ 18 O values measured for hematite in the Dales Gorge Member (Fig. 8), which requires a very low temperature of formation. Late-stage oxidation also conflicts with the paragenetic sequence of oxides outlined by [START_REF] Huberty | Silician magnetite from the Dales Gorge Member of the Brockman Iron Formation, Hamersley Group, Western Australia[END_REF] for the Dales Gorge Member, which relies on the temperature-dependence of Si substitution. Although it is clear that oxygenated groundwaters may produce hematite mineralisation in ore zones [START_REF] Mclellan | Fluid flow in extensional environments; numerical modelling with an application to Hamersley iron ores[END_REF], the evidence does not support this mechanism for all hematite in the Dales Gorge Member.

Finally, the laboratory experiments by [START_REF] Tosca | An authigenic origin for Precambrian greenalite: Implications for iron formation and the chemistry of ancient seawater[END_REF], who examined the solubility of freshly precipitated greenalite suggest that its precipitation requires a pH of 7.75-8.3, values that are at odds with those predicted for Archaean seawater (e.g., [START_REF] Grotzinger | New constraints on Precambrian ocean composition[END_REF]Blätter et al., 2016;[START_REF] Halevy | The geologic history of seawater pH[END_REF]. However, it is important to note that porewater pH in systems dominated by DIR will be in the range expected for greenalite precipitation, and that the petrographic evidence outlined above could also be consistent with early diagenetic precipitation of greenalite.

Carbonated green-rust has been observed in modern ferruginous and sulfate-poor Lake Matano in Indonesia [START_REF] Zegeye | Green rust formation controls nutrient availability in a ferruginous water column[END_REF], an environment proposed to be an analogue for the Archaean ocean [START_REF] Crowe | Photoferrotrophs thrive in an Archean ocean analogue[END_REF], and it is possible that green-rust was a precursor to magnetite in IFs. In Lake Matano, nanoparticle clusters of ferrihydrite or nanogoethite form at the redoxcline;

while at 10-15 m below the redoxcline, small hexagonal plates that are characteristic of green-rust or fougerite were observed. [START_REF] Zegeye | Green rust formation controls nutrient availability in a ferruginous water column[END_REF] attributed the formation of the green-rust phase to one of several possible pathways, including: (1) direct precipitation from a supersaturated water column, (2) as a microbial by-product of either photoferrotrophy or DIR, or (3) due to the interaction between aqueous Fe 2+ and the surface of ferrihydrite below the redoxcline. Regardless of the mechanism, [START_REF] Zegeye | Green rust formation controls nutrient availability in a ferruginous water column[END_REF] demonstrated that green-rust can be found within a naturally stratified water column, and that it may act as an important transitory phase between the redoxcline and the export of iron phases to the sediment. Thermodynamic calculations and precipitation experiments by Halevy et al. (2017) support carbonated "green-rust" as a possible phase given the appropriate Eh-pH conditions. Those authors further suggested that green-rust may have been an

important factor in Archaean IF deposition before the Great Oxidation Event (GOE) -the time in Earth's history ca. 2.45 Ga when the atmosphere first became oxygenated (see [START_REF] Lyons | The rise of oxygen in Earth's early ocean and atmosphere[END_REF] and the subsequent dominance of ferrihydrite as the precursor phase to IF. It is unknown, however, whether such a conclusion may change in the presence of high concentrations of dissolved Si.

Critically, green-rust phases are of mixed valence (e.g., [START_REF] Rickard | Chemistry of iron sulfides[END_REF], and accordingly, require a source of Fe(III). For instance, [START_REF] Etique | Magnetite as a precursor for green rust through the hydrogenotrophic activity of the iron-reducing bacteria Shewanella putrefaciens[END_REF] demonstrated that green rust might be the product of a bacterially mediated reaction involving Fe(III)-reducing bacteria and magnetite. In experiments conducted by [START_REF] Li | The formation of magnetite in the early Archean oceans[END_REF], the authors reacted natural ferrihydrite-encrusted bacterial mats with dissolved Fe(II) and showed that hexagonal platelets of a siliceous green-rust phase formed, which then agglomerated into aggregates. The latter dissolved within hours and reprecipitated as ultrafine crystals of magnetite tens of nm in size, similar to the structure made of magnetite nanoparticles produced in previous mineral synthetic experiments (e.g., [START_REF] Guan | A simple one-pot synthesis of single-crystalline magnetite hollow spheres from a single iron precursor[END_REF]. An argument against a green-rust precursor to BIF magnetite, however, is found in the Fe isotope composition of the Dales Gorge Member (Fig. 8), where the positive δ 56 Fe values cannot be explained using the 56 Fe/ 54 Fe fractionation of ca. -0.2‰ measured between Fe(II) aq and green-rust [START_REF] Wiesli | Experimental determination of Fe isotope fractionation between aqueous Fe(II), siderite and "green rust" in abiotic systems[END_REF]. Nevertheless, the possibility of green-rust as a precursor precipitate for IFs remains a subject of interest.

Available oxidants in seawater for Fe(II) oxidation

The mineralogy and Fe isotope composition of IF dictates that some oxidation of Fe(II) was necessary for precipitation, yet which mechanism(s) were dominant remains uncertain. Prior to the rise of atmospheric oxygen and the development of a protective ozone layer, Earth's surface would have been subjected to high levels of ultraviolet radiation. Ocean waters at this time were anoxic and could have supported high concentrations of dissolved Fe(II). Under such conditions, dissolved ferrous iron species, such as Fe 2+ or Fe(OH) + , absorb radiation in the 200-400 nm range, leading to the photochemical formation of dissolved ferric iron [reaction 1], which in turn, hydrolyses to form Fe(III)-oxyhydroxides at circumneutral pH [START_REF] Cairns-Smith | Precambrian solution photochemistry, inverse segregation, and banded iron formations[END_REF][START_REF] Braterman | Photooxidation of hydrated Fe 2+significance for banded iron formations[END_REF][START_REF] Anbar | The photochemistry of manganese and the origin of banded iron formations[END_REF].

[1] 2Fe 2+ (aq) + 2H + + hv à 2Fe 3+ (aq) + H 2 , where hv represents ultraviolet radiation However, experiments focused on determining the specific rates of Fe(II) photochemical oxidation did not simulate the complex, disequilibrium water chemistry characteristic of an ocean where Fe(II)-and Si-rich hydrothermal waters mixed with ambient Si-saturated seawater that also contained high concentrations of HCO 3 -. Indeed, in fluids having high dissolved Fe(II), Si(OH) 4 , and HCO 3 -, the oxidation effects of either UV-C (200-280 nm wavelength) or UV-A (320-400 nm wavelength) were found to be negligible compared to the precipitation of ferrous silicates and ferrous carbonates (e.g., Konhauser et al., 2007b). The point here is that although UV photooxidation does occur, in an ocean approaching saturation or supersaturation with respect to ferrous carbonate and silicate phases, UV photo-oxidation may have been relatively insignificant, particularly at anything but the shallowest water depths. Importantly, iron-rich waters, near saturation with ferrous iron phases, are likely essential for driving IF deposition, instead of simply background marine sedimentation.

Renewed interest in UV photo-oxidation of Fe(II) has focused on the possibility of UVderived oxidants, such as hydrogen peroxide (H 2 O 2 ), causing Fe(III) precipitation [START_REF] Pecoits | Atmospheric hydrogen peroxide and Eoarchean iron formations[END_REF]. Based on mass-balance calculations and consideration of the amount of hydrogen peroxide potentially generated by UV photo-oxidation of water, it was determined that the amount of hydrogen peroxide derived from atmospheric reactions is orders of magnitude too small to account for the Fe(III) sedimentation required to form IF [START_REF] Pecoits | Atmospheric hydrogen peroxide and Eoarchean iron formations[END_REF]. Studies of Fe and O isotopes have attempted to constrain the potential contribution of UV photo-oxidation of Fe(II) for IF deposition; but, results proved to be non-unique isotope signatures associated with UV photooxidation, thus hindering such an approach [START_REF] Nie | Iron and oxygen isotope fractionation during iron UV photo-oxidation: Implications for early Earth and Mars[END_REF]. Moreover, the recent UV photooxidation experiments of Nie et al. ( 2017) utilised a relatively simple, borate-buffered solution without silica that does not correspond to Archaean and Palaeoproterozoic seawater compositions.

Accordingly, although it remains possible that UV photo-oxidation contributed to IF precipitation, this mechanism seems unlikely to have provided the major source of Fe(III) for the deposition of large IFs.

As an alternative to the abiological model, the presence of ferric iron minerals in IFs has also been ascribed to the metabolic activity of planktonic bacteria in oceanic photic zones (Fig. 9).

Although a biological role in Fe(III)-oxyhydroxide precipitation was suggested around a century ago (e.g., [START_REF] Leith | The Mesabi iron-bearing district of Minnesota[END_REF][START_REF] Gruner | The origin of sedimentary iron-formations: The Biwabik Formation of the Mesabi Range[END_REF], the importance of bacteria in this process began to receive greater acceptance with the discovery of microfossils in the ca. 1.88 Ga IFs in the Animikie Basin of the Lake Superior region (e.g., Tyler and Barhoorn, 1954;[START_REF] Barghoorn | Microorganisms from Gunflint Chert[END_REF][START_REF] Cloud | Significance of Gunflint (Precambrian) microflora[END_REF]. Based on the assumption that the microfossils were cyanobacteria, or their predecessors, [START_REF] Cloud | Paleoecological significance of banded iron-formation[END_REF] proposed that these primitive O 2 -producing photosynthetic bacteria lacked suitably advanced oxygen-mediating enzymes, and, consequently, required Fe(II) as an oxygen acceptor.

Therefore, these microorganisms would have flourished when Fe(II) (and other nutrients) were made episodically available, allowing for the indirect oxidation and precipitation of Fe(III)oxyhydroxides. By contrast, the photosynthetic population would have declined in number when Fe(II) availability was limited. Recent analyses of the microfossils within the 1.88 Ga Gunflint Iron Formation appear to confirm that cyanobacteria were a significant component of shallow-marine communities at that time [START_REF] Lepot | Iron minerals within specific microfossil morphospecies of the 1.88 Ga Gunflint Formation[END_REF]. What makes this model particularly intriguing, however, is experimental work demonstrating that Fe(II) concentrations in the range of 10s to 100s of µM would lead to oxidative stress in cyanobacteria, resulting from reactions between Fe(II) and oxygen that produce reactive oxygen species [START_REF] Swanner | Modulation of oxygen production in Archaean oceans by episodes of Fe (II) toxicity[END_REF]. These contradictory findings may suggest that primitive cyanobacteria had, at least by 1.9 Ga, some form of protection mechanism against Fe(II) concentrations becoming too high in proximal environments. In this regard, the study of [START_REF] Lepot | Iron minerals within specific microfossil morphospecies of the 1.88 Ga Gunflint Formation[END_REF] proposes that intracellular ferric iron biomineralisation might have offered protection against Fe(II) toxicity by reducing proximal iron concentrations.

Other bacteria may have directly utilised low O 2 concentrations available in surface waters of the Archaean oceans for Fe(II) oxidation. For example, [START_REF] Holm | The 13C/12C ratios of siderite and organic matter of a modern metalliferous hydrothermal sediment and their implications for banded iron formations[END_REF] speculated that oxidation of Fe(II) by chemolithoautotrophic species, such as Gallionella ferruginea, would have been kinetically favoured in an ocean having limited free oxygen because abiological rates of Fe(II) oxidation at circumneutral pH are slow under microaerobic conditions (e.g., [START_REF] Søgaard | Conditions and rates of biotic and abiotic iron precipitation in selected Danish freshwater plants and microscopic analysis of precipitate morphology[END_REF].

Interestingly, some of the ca. 1.88 Ga Gunflint-type microfossils from the Animikie Basin have been reinterpreted as chemolithoautotrophic Fe(II)-oxidisers [START_REF] Golubic | Early cyanobacterial fossil record: Preservation, palaeoenvironments and identification[END_REF][START_REF] Planavsky | Iron-oxidizing microbial ecosystems thrived in late Paleoproterozoic redox-stratified oceans[END_REF]. It is important to note that some Gunflint fossil groupings (e.g., Gunflintia) may encompass multiple types of bacteria, including those with different metabolisms [START_REF] Shapiro | Hematite-coated microfossils: Primary ecological fingerprint or taphonomic oddity of the Paleoproterozoic?[END_REF]. The prospect of an early community of Fe(II)-oxidisers has been bolstered by the recognition of the importance of these bacteria in modern Fe-cycling in Chesapeake Bay [START_REF] Field | Planktonic marine iron-oxidizers drive iron mineralization under low oxygen conditions[END_REF]. These authors demonstrated that the Fe(II)-oxidisers were associated with low oxygen concentrations and active Fe redox cycling in the oxic-anoxic transition zone (<3 µM O 2 , <0.2 µM H 2 S). Although cyanobacteria were also detected in this transition zone, oxygen concentrations were too low to support significant rates of abiotic Fe(II) oxidation. Therefore, [START_REF] Field | Planktonic marine iron-oxidizers drive iron mineralization under low oxygen conditions[END_REF] speculated that cyanobacteria may be providing oxygen for microaerophilic Fe(II) oxidation through a symbiotic relationship; at high Fe(II) levels, cyanobacteria would gain protection against Fe(II) toxicity. The implication of recognising microaerophilic Fe(II)-oxidisers in modern ferruginous environments for IF deposition has also been discussed in recent reviews by [START_REF] Chan | The role of microaerophilic Fe-oxidizing microorganisms in producing banded iron formations[END_REF] and [START_REF] Koeksoy | Using modern ferruginous habitats to interpret Precambrian banded iron formation deposition[END_REF].

A different biological model was proposed by [START_REF] Garrels | Genesis of Precambrian iron-formations and development of atmospheric oxygen[END_REF] and [START_REF] Hartman | The evolution of photosynthesis and microbial mats: A speculation on the banded iron formations[END_REF], who both suggested photoferrotrophy as a metabolism supporting Fe(III) production. During this form of anoxygenic photosynthesis, light, not O 2 , may have coupled the carbon and iron cycles, with the bacteria using Fe(II) rather than H 2 O as an electron donor, and producing Fe(III) instead of O 2 as the metabolic waste [reaction 2].

[2] 4Fe 2+ + HCO 3

-+ 10H 2 O à 4Fe(OH) 3 + CH 2 O + 7H +
Since then, several experimental studies have confirmed that various purple and green bacteria can use Fe(II) as a reductant for CO 2 fixation (e.g., [START_REF] Widdel | Ferrous iron oxidation by anoxygenic phototrophic bacteria[END_REF][START_REF] Heising | Chlorobium ferrooxidans sp. nov., a phototrophic green sulfur bacterium that oxidizes ferrous iron in co-culture with a "Geospirillum" sp. strain[END_REF][START_REF] Straub | Rhodovulum iodosum sp. nov., and Rhodovulum religious sp. nov., two new marine phototrophic ferrous-iron-oxidizing purple bacteria[END_REF]. The ferric iron minerals these strains produce are also consistent with the most likely precursor IF minerals, Fe(III)-oxyhydroxides [START_REF] Kappler | Formation of Fe (III) minerals by Fe(II) oxidizing photoautotrophic bacteria[END_REF].

Photoferrotrophs have also been identified in Kabuno Bay, Democratic Republic of Congo, a basin characterised by ferruginous conditions and by observed rates of Fe(II) oxidation that appear to exceed those required for IF deposition [START_REF] Llirós | Pelagic photoferrotrophy and iron cycling in a modern ferruginous basin[END_REF]. In particular, photoferrotrophs produce amorphous to poorly crystalline Fe(III)-oxyhydroxide minerals, which carry a net positive charge.

Such biogenic minerals are expected to bind to organic carbon (cells) and silica, with the net effect being the deposition of Fe-Si-C aggregates onto the sea floor [START_REF] Posth | Alternating Si and Fe deposition caused by temperature fluctuations in Precambrian oceans[END_REF][START_REF] Posth | Size, density and composition of cellmineral aggregates formed during anoxygenic phototrophic Fe(II) oxidation: Impact on modern and ancient environments[END_REF]. One potential issue with the suggestion of photoferrotrophy in the Archaean is that a lack of an early ozone layer would have allowed high incoming UV radiation in the photic zone where photoferrotrophs are presumed to have resided. A solution to this issue lies in the work of [START_REF] Gauger | Protection of phototrophic iron (II)-oxidizing bacteria from UV irradiation by biogenic iron (III) minerals: Implications for early Archean banded iron formation[END_REF][START_REF] Gauger | Protection of nitrate-reducing Fe(II)-oxidizing bacteria from UV radiation by biogenic Fe(III) minerals[END_REF], who provided evidence that minerals produced by photoferrotrophs, as well as other Fe(II)-oxidising bacteria, can protect the cells from UV radiation. These Fe(III) "shields" consist of nanometer-sized grains of ferrihydrite loosely attached to the cell surfaces.

The geobiology of iron formations

Evidence in the rock record for the evolution of oxygenic photosynthesis

To determine the significance of different biological oxidative mechanisms that may have driven IF deposition it is essential to constrain when oxygenic photosynthesis and anoxygenic photosynthesis evolved. Although the timing of cyanobacterial and photoferrotroph evolution is still debated, with some molecular clock and phylogenomic analyses suggesting that the crown group of oxygenic cyanobacteria post-dates the Proterozoic rise in oxygen [START_REF] Shih | Crown group Oxyphotobacteria postdate the rise of oxygen[END_REF], many indicators exist in the rock record pointing to an earlier appearance in the Archaean for both. In the case of cyanobacteria, these indicators include:

(1) Stromatolitic assemblages in the 2.72 Ga Tumbiana Formation, Western Australia, are presumed to have been constructed by photoautotrophs that utilised oxygenic photosynthesis [START_REF] Buick | The antiquity of oxygenic photosynthesis -evidence from stromatolites in sulfatedeficient Archean lakes[END_REF].

This view is supported by the earliest recognised fossil assemblage of filamentous and coccoidal cell colonies from the ca. 2.6 Ga Campbellrand Group, South Africa, possibly including oscillatoriacean cyanobacterial genera such as Phormidium and Lyngbya [START_REF] Altermann | Microfossils from the Neoarchean Campbell Group, Griqualand West Sequence of the Transvaal Supergroup, and their paleoenvironmental and evolutionary implications[END_REF]. [START_REF] Dodd | Evidence for early life in Earth's oldest hydrothermal vent precipitates[END_REF] have recently identified microscale hematite-encrusted filaments within the ca. 3.8 Ga Nuvvuagittuq Supracrustal Belt in northern Québec, which they suggest may be ancient microaerophilic Fe(II)-oxidisers. If true, this implies a metabolism that relied on the presence of oxygen, and, by extension, cyanobacteria. Putative stromatolites have also been identified in the ca. 3.7 Ga Isua Greenstone Belt, southwest Greenland [START_REF] Nutman | Rapid emergence of life shown by discovery of 3,700-million-year-old microbial structures[END_REF], which the authors ascribed to an autotrophic carbon-fixing metabolism. However, the construction of stromatolite-like structures by the anoxygenic phototroph Rhodopseudomonas palustris [START_REF] Bosak | A possible role for anoxygenic photosynthetic microbes in the formation of ancient stromatolites[END_REF] challenges the notion that stromatolites uniquely mark cyanobacterial presence in the geologic record and that microfossil metabolism can be inferred from morphology alone.

(2) When microbial mats grow on siliciclastic sediment (e.g., in tidal flats, lagoons, sabkhas) they locally leave distinctive wrinkle and roll-up structures, laminated textures, and micro-impressions collectively referred to as mat-related structures (MRS). The filamentous bacteria, typically cyanobacteria or chemotrophic sulfide-oxidising bacteria, such as Beggiatoa, that comprise the mats, trap and bind sediments with their exopolysaccharide (EPS) secretions as they grow upward;

this biostabilisation method allows them to withstand erosion by waves and currents. MRS are one of the oldest forms of fossil evidence, having been found in rocks as old as the 3.2 Ga Moodies Group, South Africa [START_REF] Noffke | A new window into early Archean life: Microbial mats in Earth's oldest siliciclastic tidal deposit (3.2 Ga Moodies Group, South Africa)[END_REF][START_REF] Heubeck | An early ecosystem of Archean tidal microbial mats (Moodies Group, South Africa, ca. 3.2 Ga)[END_REF] and the 2.9 Ga Pongola Supergroup, South Africa [START_REF] Noffke | Earth's earliest microbial mats in a siliciclastic marine environment (2.9 Ga Mozaan Group, South Africa)[END_REF]. In the case of the Moodies Group, recent stratigraphic and depositional facies analysis have documented the association of three principal mat morphotypes representing different environmental settings: (1) planar-type in coastal floodplains, (2) wavy-type in intertidal zones, and (3) tufted-type in upper inter-to supratidal zones. All mat types suggest formation by phototrophic biota, but the tufted morphology implies an intricate level of coordinated growth commonly known from cyanobacterial mats in modern environments [START_REF] Homann | Morphological adaptations of 3.22 Ga-old tufted microbial mats to Archean coastal habitats (Moodies Group, Barberton Greenstone Belt, South Africa)[END_REF].

(3) Oxygenic photosynthesis at 2.72 to 2.59 Ga has been inferred from the presence of extremely isotopically depleted kerogen within carbonates and shales in the Hamersley Province of Western Australia, the Kaapvaal Craton in South Africa, and the Superior Craton in Canada. Organic carbon δ 13 C values in these metasediments are as low as -60‰ [START_REF] Hayes | Geochemical evidence bearing on the origin of aerobiosis, a speculative hypothesis[END_REF][START_REF] Eigenbrode | Late Archean rise of aerobic microbial ecosystems[END_REF]. The most 13 C-depleted values have been ascribed to the assimilation of methane by chemolithoautotrophic, methanotrophic bacteria that in the modern world utilise electron acceptors such as O 2 , SO 4 2-, or NO 3 -. However, sulfate can also form via photolysis of SO 2 of volcanic origin, and it might be possible to have methane oxidation via Fe 3+ (e.g., [START_REF] Konhauser | The potential significance of microbial Fe(III) reduction during deposition of Precambrian banded iron formations[END_REF][START_REF] Beal | Manganese-and iron-dependent marine methane oxidation[END_REF][START_REF] Crowe | The methane cycle in ferruginous Lake Matano[END_REF]. As outlined above, Fe(II) oxidation does not require oxygenic photosynthesis.

(4) Nitrogen isotope compositions of kerogens in minimally altered shales from the Campbellrand-Malmani carbonate platform in South Africa and broadly correlative sedimentary succession in Western Australia (Hamersley Group) show a significant rise in δ 15 N values between 2.67 and 2.50

Ga [START_REF] Godfrey | The cycling and redox state of nitrogen in the Archaean ocean[END_REF][START_REF] Garvin | Isotopic evidence for an aerobic nitrogen cycle in the latest Archean[END_REF]. Exceptionally high δ 15 N values are also reported for the 2.72 Ga Tumbiana Formation [START_REF] Thomazo | Extreme 15 N-enrichments in 2.72-Gyr-old sediments: Evidence for a turning point in the nitrogen cycle[END_REF]. This positive shift has been interpreted as evidence for the onset of nitrification-denitrification reactions in the surface oceans (e.g., [START_REF] Beaumont | Nitrogen isotope ratios of kerogens in Precambrian cherts: A record of the evolution of atmosphere chemistry?[END_REF]; importantly, nitrification is typically thought to require oxygen.

However, the emerging view is that it is possible to have anaerobic nitrification via Fe 3+ (e.g., [START_REF] Yang | Nitrogen loss from soil through anaerobic ammonium oxidation coupled to iron reduction[END_REF][START_REF] Busigny | Nitrogen cycle in the Late Archean ferruginous ocean[END_REF], thus making implications of the N isotope record more ambiguous than considered previously. (5) Over the past decade, redox-sensitive trace element data generated from paleosols, shales, carbonates, and IFs suggests an early Archaean rise of oxygenic photosynthesisers (see Robbins et al., 2016 for review). Amongst these data, the most widely cited are from shales in the Campbellrand-Malmani and Hamersley groups that contain high concentrations of Re and Mo.

Strong Re and Mo enrichments and coupled Mo-Fe isotope data were interpreted as evidence for oxidative weathering of sulfide-bearing minerals on the continents, and thus the evolution of oxygenic photosynthesis as early as ca. 2.6 Ga [START_REF] Anbar | A whiff of oxygen before the Great Oxidation Event?[END_REF][START_REF] Kendall | Pervasive oxygenation along late Archaean ocean margins[END_REF][START_REF] Czaja | Evidence for free oxygen in the Neoarchean ocean based on coupled ironmolybdenum isotope fractionation[END_REF]Stüeken et al., 2015b). Re-Os isotopes from the same succession investigated by [START_REF] Anbar | A whiff of oxygen before the Great Oxidation Event?[END_REF] further indicate that the observed Re enrichment is a primary, terrestrial signal, and not the result of post-depositional alteration [START_REF] Kendall | Transient episodes of mild environmental oxygenation and oxidative continental weathering during the late Archean[END_REF]. Additional support for the appearance of oxygenic photosynthesis by 2.7 Ga comes from sulfide concentrations in marginal marine sediments [START_REF] Stüeken | Contributions to late Archaean sulphur cycling by life on land[END_REF]. Also, continuously increasing d 98 Mo values in the black shales, carbonates, and oxide-facies IFs of the Marra Mamba and Wittenoom formations (2.6-2.5 Ga) in the Hamersley

Basin have been tied to the sorption of light Mo isotopes onto Mn-oxides [START_REF] Kurzweil | Manganese oxide shuttling in pre-GOE oceans-evidence from molybdenum and iron isotopes[END_REF].

This pattern is also observed in the IF of the 2.95 Ga Sinqeni Formation, Pongola Supergroup, where the Mo isotope signature is consistent with the interaction with Mn(IV) oxides (Planavsky et al., 2014a). These Mo isotope signatures likely required an oxygenated water column and delivery of Mn(IV) oxides to the sediment pile, given rapid Mn(IV) oxide reduction via Fe 2+ (e.g., [START_REF] Busigny | Iron isotopes in an Archean ocean analogue[END_REF]. This implies that these signatures cannot be easily linked to anoxygenic photosynthetic Mn(II) oxidation more than a half of billion years before the GOE. Similarly, Cr isotopes from the 2.98 Ga Nsuze paleosol, also in the Pongola Supergroup, have been used as evidence for significant oxidative weathering on land, and thus appreciable levels of atmospheric oxygen, even at that time [START_REF] Crowe | Atmospheric oxygenation three billion years ago[END_REF]. Frei et al. ( 2016) combined U/Th ratios and Cr isotopes to argue for oxidative weathering as far back in time as the 3.8-3.7 Ga Isua Supracrustal Belt, but at very low atmosphere oxygen contents (i.e., <<10 -5 present atmospheric levels, PAL; Fig. 10).

In It is also important to note that oxygen evolution in the photic zone may have been decoupled from atmospheric evolution. (6) Several weight percent concentrations of organic carbon in sedimentary rocks provide one of the most straightforward and compelling arguments for the Archaean evolution of oxygenic photosynthesis, although this record becomes sparse for strata older than ~2.8 Ga. A wide range of microbial metabolisms can, of course, generate organic carbon, but only a handful of these metabolisms can realistically produce organic carbon-rich sedimentary rocks [START_REF] Buick | When did oxygenic photosynthesis evolve?[END_REF]Scott et al., 2011;[START_REF] Lyons | The rise of oxygen in Earth's early ocean and atmosphere[END_REF][START_REF] Krissansen-Totton | A statistical analysis of the carbon isotope record from the Archean to Phanerozoic and implications for the rise of oxygen[END_REF]. Since photoferrotrophy produces a particulate Fe(III)-oxyhydroxide as well as organic carbon, this process is more likely, as outlined below, to produce organic carbon-poor, iron-rich rocks like IF, because much of the organic matter would be oxidised via DIR, which, if this process was widespread, would limit the use of organic carbon content in IFs as a direct indicator of organic carbon deposition. Photosynthetic sulfide and H 2 oxidisers can be important primary producers and could in theory produce organic matter-rich rocks like black shales. However, the reductant for these microbial metabolisms, unlike for oxygenic photosynthesis, must be sourced from a hydrothermal system or produced via microbial degradation of organic matter. Sulfide and H 2 , unlike Fe(II), are not thought to have been present at significant levels in the Archaean water column [START_REF] Lyons | The rise of oxygen in Earth's early ocean and atmosphere[END_REF].

Evidence in the rock record for the evolution of photoferrotrophy

Compared to oxygenic photosynthesis, there is less actual physical and chemical evidence for the existence of Fe(II)-oxidising phototrophs in the Archaean. An argument in favour of photoferrotrophy having played a prominent role in IF deposition is that much of the evidence for pre-GOE surface-ocean oxygenation remains contested and proxies that specifically focus on shallow-ocean redox conditions, such as carbonate iodine [START_REF] Hardisty | An iodine record of Paleoproterozoic surface ocean oxygenation[END_REF] and REEs (Planavsky et al., 2010a), point to the dominance of IF and carbonate deposition under broadly anoxic conditions in the oceans until the GOE. Supporting this, molecular phylogenetic analysis of many enzymes involved in (bacterio)-chlorophyll biosynthesis suggests that anoxygenic photosynthetic lineages are likely more deeply rooted than oxygenic cyanobacterial lineages [START_REF] Xiong | Photosynthesis: What color was its origin?[END_REF].

Numerical models have been used to suggest that anoxygenic photosynthesisers would control Fe(II) oxidation even if the surface layer was weakly oxygenated [START_REF] Kappler | Deposition of banded iron formations by anoxygenic phototrophic Fe(II)-oxidizing bacteria[END_REF]. This model is based on the simple idea that a layer of photoferrotrophs could thrive beneath an oxygenated surface layer even at low levels of photosynthetically active light. The model employs Fe(II) oxidation rates from modern photoferrotrophs under a range of reasonable mixing rates in the ocean. It is also likely that photoferrotrophs flourishing at depth could have exhausted dissolved phosphorous in upwelling waters, thus outcompeting cyanobacteria for a key nutrient, and by extension limiting their activity in overlying waters [START_REF] Jones | Iron oxides, divalent cations, silica, and the early Earth phosphorus crisis[END_REF]. [START_REF] Kappler | Deposition of banded iron formations by anoxygenic phototrophic Fe(II)-oxidizing bacteria[END_REF] presented a model for determining the thickness of a marine layer with photoferrotrophic bacteria required to fully oxidise a hydrothermal source of Fe(II) at a variety of Fe(II)-oxidation rates in which eddy-diffusion was the predominant mode of Fe(II) transport. This essentially assumes that diapycnal diffusivity was the main transport term. At a conservative, experiment-based, depth-integrated, phototrophic Fe(II)-oxidation rate of 1.4 × 10 -5 M/day (corresponding to a depth of 100 m), an estimated Fe(II) concentration of 0.5 mM, and the modern global mean eddy-diffusion rate of 0.1 cm 2 /s, it was estimated that a photoferrotrophic bacterial layer 17.6 m thick was required to completely oxidise the Fe(II) input. This result suggests that such bacteria could potentially have oxidised all of the upwelling Fe(II) in Precambrian oceans.

A potential limitation with the approach of [START_REF] Kappler | Deposition of banded iron formations by anoxygenic phototrophic Fe(II)-oxidizing bacteria[END_REF] is that in many settings advection can be the dominant transport process (e.g., upwelling zones). As a simple test to assess the required thickness of a photoferrotrophic bacterial layer to oxidise both upwelling and diffusive Not surprisingly, the thickness of the photosynthetic Fe(II)-oxidising bacterial layer increases when advection is considered as opposed to solely eddy-diffusion transport. The thickness of the photoferrotrophic bacterial layer is quite sensitive to upwelling rates, especially at low rates of Fe(II) oxidation. At the conservatively estimated Fe(II)-oxidation rate of 1.4 × 10 -5 M/day and an Fe(II) concentration of 0.1 mM, a 34 m-or >100 m-thick layer of Fe(II)-oxidising bacteria is required to fully oxidise the Fe(II) input at upwelling rates of 0.5 and 5 m/day, respectively (Fig. 11). These estimates are significantly larger than the 17.6 m thickness calculated with eddydiffusion alone by [START_REF] Kappler | Deposition of banded iron formations by anoxygenic phototrophic Fe(II)-oxidizing bacteria[END_REF]. Furthermore, when considering the expected range of upwelling conditions and reasonable rates of photoferrotrophic Fe(II) oxidation, this result indicates a wide scope in the extent of Fe(II) oxidation in Archaean water masses. In fact, given that upwelling is highly variable on a seasonal time scale (with shifts over an order of magnitude in some settings), a variable extent of oxidation is likely to characterise any Archaean setting where upwelling was an important transport mechanism.

Available reductants and diagenesis of iron formations

Oxidation of Fe(II), either indirectly via oxygenic photosynthesis, or directly via anoxygenic photoferrotrophy, would produce biomass that settled to the seafloor together with the Fe(III) minerals (e.g., [START_REF] Konhauser | The potential significance of microbial Fe(III) reduction during deposition of Precambrian banded iron formations[END_REF][START_REF] Li | Mineral ecophysiological data provide growing evidence for microbial activity in banded-iron formations[END_REF]Posth et al., 2013a, b). If, as today, the organic carbon was oxidised during burial by either diagenesis or metamorphism, the relevant question is which terminal electron acceptor was present at the seafloor during times of IF deposition, and at what concentrations? Despite the possibility of some oxic surface waters being generated by cyanobacterial activity as early as 3.8 Ga (see discussion above), deep waters, and by extension the seafloor, remained anoxic, unlike today where sediment pore-waters can have dissolved oxygen at depths of several millimeters, and locally even greater depths (Fig. 12A). In the absence of O 2 , the fermentation products in the bottom waters and/or shallow sediments would have been oxidised via some other form of anaerobic respiratory process. In terms of such pathways, the paucity of O 2 would also have meant minimal nitrate and sulfate availability, the former being evident from the lack of positive δ 15 N values before ca. 2.7 Ga [START_REF] Godfrey | The cycling and redox state of nitrogen in the Archaean ocean[END_REF][START_REF] Thomazo | Extreme 15 N-enrichments in 2.72-Gyr-old sediments: Evidence for a turning point in the nitrogen cycle[END_REF], and the latter being evident from negligible sulfur isotopic fractionations in the Archaean sedimentary sulfide record (e.g., [START_REF] Strauss | Sulphur isotopes and the early Archaean sulphur cycle[END_REF]. The supply of MnO 2 was likely also not significant until after the GOE, due to the oxidation of dissolved Mn(II) requiring the presence of free dissolved oxygen and an aerobic microbial catalyst (e.g., [START_REF] Tebo | Geomicrobiology of manganese(II) oxidation[END_REF], although the plausibility of a Mn(II)-based photosynthesis pre-dating the rise of oxygen has been hypothesised [START_REF] Johnson | Manganese mineralogy and diagenesis in the sedimentary rock record[END_REF], but not confirmed in modern cultures.

By contrast, the ferric minerals in IF would have represented a favourable electron acceptor for the oxidation of organic matter [START_REF] Walker | Suboxic diagenesis in banded iron formations[END_REF][START_REF] Nealson | Iron reduction by bacteria: A potential role in the genesis of banded iron formations[END_REF] through DIR by various bacteria and archaea (Fig. 12B). Significantly, coupling the reduction of some Fe(III) minerals to the oxidation of organic matter not only explains the low content of organic carbon in IFs (<0.5 wt.%; [START_REF] Gole | Banded iron-formation through much of Precambrian time[END_REF], but it also explains highly negative δ 13 C values of the early diagenetic Fe(II)-rich carbonates [START_REF] Perry | Geology and stable isotope geochemistry of Biwabik Iron Formation, northern Minnesota[END_REF][START_REF] Walker | Suboxic diagenesis in banded iron formations[END_REF][START_REF] Baur | Millimeter-scale variations of stable isotope abundances in carbonates from banded iron-formations in the Hamersley Group of Western Australia[END_REF][START_REF] Heimann | Fe, C, and O isotope compositions of banded iron formation carbonates demonstrate a major role for dissimilatory iron reduction in ~ 2.5 Ga marine environments[END_REF][START_REF] Craddock | Iron and carbon isotope evidence for microbial iron respiration throughout the Archean[END_REF], the small-scale heterogeneity in δ 56 Fe values (e.g., [START_REF] Steinhofel | Deciphering formation processes of banded iron formations from the Transvaal and the Hamersley successions by combined Si and Fe isotope analysis using UV femtosecond laser ablation[END_REF]W. Li et al., 2013a), the presence of Fe(II)-bearing minerals in IF (e.g., magnetite, siderite, greenalite -recall section 2.3), and the general lack of microfossils preserved in the Fe-rich layers lacking silicification. Moreover, the reduction of ferric iron may also have been coupled to the anaerobic oxidation of methane [START_REF] Konhauser | The potential significance of microbial Fe(III) reduction during deposition of Precambrian banded iron formations[END_REF], an observation supported by the reaction's likely occurrence in modern marine sediments [START_REF] Riedinger | An inorganic geochemical argument for coupled anaerobic oxidation of methane and iron reduction in marine sediments[END_REF] and lakes [START_REF] Crowe | The methane cycle in ferruginous Lake Matano[END_REF], as well as in culture experiments [START_REF] Ettwig | Archaea catalyze iron-dependent anaerobic oxidation of methane[END_REF][START_REF] Bray | Shifting microbial communities sustain multiyear iron reduction and methanogenesis in ferruginous sediment incubations[END_REF]).

An estimate of the organic carbon content (C org ) that may have been initially deposited with Fe(III) precipitates in IFs may be made by calculating the electron equivalents of the organic C tied in Fe-carbonates, based on their δ 13 C values, as well as Fe(II) quantities required to make magnetite (W. Li et al., 2013a). Although measured C org contents of the large Transvaal BIFs, correlative with the Hamersley BIFs, are quite low (~0.01 wt.%), estimated initial C org contents are much higher, ~5

wt. % (Fig. 13), based on the C isotope and magnetite mass balances. Using C isotope data, CO 2 content, and Fe(II) and Fe(III) data from Beukes et al. (1990), Beukes and Klein (1990), and [START_REF] Kaufman | Geochemical and mineralogic effects of contact metamorphism on banded iron-formation: An example from the Transvaal Basin, South Africa[END_REF], the average initial C org based on C isotope composition is 1.28 wt.% for oxidefacies IF and 3.24 wt. % for siderite-facies IF. For oxide-facies IF, based on Fe(II) abundances in magnetite, and assuming an equivalent electron balance was contained in organic carbon, an additional 3.42 wt.% C org is calculated, for a total of 4.70 wt.% initial C org in oxide-facies IF.

Although this calculation assumes that magnetite was produced solely by reduction of hematite, stable isotope results discussed below support that interpretation. Such quantities of C org are similar to those in Neoarchaean black shales, and highlight the possibility that significant organic carbon was buried with Fe(III) precipitates during IF deposition.

The possibility of DIR producing the large Fe(II) inventory in IFs is supported by molecular phylogeny, specifically from the observation that many deeply-branching Archaea (i.e., some of the oldest purported species) can use H 2 to reduce Fe(III) to support chemolithoautotrophic growth [START_REF] Vargas | Microbiological evidence for Fe(III) reduction on early Earth[END_REF]. Fe(III)-reducing bacteria can even use quinone moieties as electron shuttles between solid-phase iron minerals and H 2 , thereby alleviating the need for direct contact between the cell and mineral surface [START_REF] Lovley | Reduction of humic substances and Fe(III) by hyperthermophilic microorganisms[END_REF]. Moreover, DIR has been shown to be broadly distributed amongst several known Proteobacteria genera, suggesting that this form of metabolism became widespread early in the course of evolution [START_REF] Barns | Microbial diversity in ocean, surface and subsurface environments[END_REF].

Finding clear evidence of Fe(III) reduction in the mineralogical record of IFs, however, has been difficult, although a few possibilities exist. For instance, the spheroidal forms of siderite found in several IFs are similar to those of the siderite grains formed experimentally when ferrihydrite is reacted with organic carbon at pressure and temperature conditions commensurate with IF diagenesis [START_REF] Köhler | Biological carbon precursor to diagenetic siderite with spherical structures in iron formations[END_REF]. Likewise, the large magnetite grains found in many IFs can be replicated experimentally through a three-stage sequence, beginning with DIR of an initial ferric iron-rich sediment coupled to the oxidation of dead phytoplankton biomass, followed by magnetite crystal aging, and ultimately by pressure-and temperature-induced abiotic alteration of the biogenic magnetite during metamorphism (Y.-L. Li et al., 2013). Such transformations, if they occurred at temperatures of >100-150 o C, may explain magnetite that has positive δ 18 O values (Fig. 8); inorganic reaction between Fe(III)-oxyhydroxides and organic matter is another possible explanation. However, as noted above, magnetite that has δ 18 O values less than zero requires formation at lower temperature (Fig. 8), and, in such cases, direct, microbially-catalysed reduction seems the best explanation. The strongest proof of microbial Fe(III) reduction, therefore, seems to lie in the isotopic compositions of IF minerals.

Although much of the discussion on microbial Fe(III) reduction in IFs has focused on processes within the sediment, it is possible that Fe(III) reduction occurred in the water column, or exported Fe(II) to the water column, which would have affected the efficiency of Fe and C recycling prior to final burial. In an attempt to test this possibility, [START_REF] Konhauser | The potential significance of microbial Fe(III) reduction during deposition of Precambrian banded iron formations[END_REF] modelled the ancient Fe cycle based simply on conservative experimental rates of photosynthetic Fe(II) oxidation in the photic zone. They showed that under ideal growth conditions, as much as 70% of the biologically produced Fe(III) could have been recycled back into the water column via fermentation and organic carbon oxidation coupled to DIR. By comparing the potential size of biomass generated phototrophically with the reducing equivalents required for Fe(III) reduction and magnetite formation, they also hypothesised that another anaerobic metabolic pathway might have been utilised in the surface sediment to oxidise organic carbon. Based on the premise that the deep ocean waters were anoxic, this role could have been fulfilled by a consortium of fermenters and methanogens [START_REF] Konhauser | The potential significance of microbial Fe(III) reduction during deposition of Precambrian banded iron formations[END_REF].

Iron-rich carbonates (siderite and ankerite) in IFs have δ 13 C values that are significantly lower than those expected for precipitation from seawater (as represented by limestone and dolostone) (Fig. 14). δ 13 C values as low as -16‰ [START_REF] Higgins | Oxygenation of the ocean and sediments: Consequences for the seafloor carbonate factory[END_REF] have been traditionally explained (e.g., [START_REF] Walker | Suboxic diagenesis in banded iron formations[END_REF] as resulting from the mixing of dissolved inorganic carbon (DIC) from seawater and remineralised organic carbon via DIR [reaction 3], as opposed to further carbon inputs derived from Fe(III)-induced methane oxidation [reaction 4]. However, the latter possibility has yet to be tested with simple numerical models that account for the depositional setting of IFs and realistic Archaean seawater chemistry.

[3] 4Fe(OH

) 3 + CH 2 O + 3HCO 3 -à 4FeCO 3 + 3OH -+ 7H 2 O [4] 4Fe(OH) 3 + 0.5CH 4 + 3.5HCO 3 -à 4FeCO 3 + 3.5OH -+ 7H 2 O
Using the equations in [START_REF] Konhauser | The potential significance of microbial Fe(III) reduction during deposition of Precambrian banded iron formations[END_REF], only one mole of carbonate is produced for four moles of Fe(III) reduced, requiring an additional three moles of carbonate, presumably from a seawater source; this is equivalent to a C org :C sw ratio of 1:3 (Fig. 14). In case of the equation 4, C meth :C sw ratio is 1:7. These ratios are consistent with the fact that IF carbonates are rarely pure siderite, but instead consist of magnesite-siderite and dolomite-ankerite solid solutions and hence require a seawater source of Ca and Mg. As discussed in [START_REF] Heimann | Fe, C, and O isotope compositions of banded iron formation carbonates demonstrate a major role for dissimilatory iron reduction in ~ 2.5 Ga marine environments[END_REF], alternative proportions of external carbonate relative to organic carbon can also be envisioned [reactions 5 and 6]:

[5] 4Fe(OH) 3 + CH 2 O + 2HCO 3 -→ 4FeCO 3 + Fe 2+ aq + 4OH -+ 6H 2 O [6] 4Fe(OH) 3 + CH 2 O + HCO 3 -→ 4FeCO 3 + 2Fe 2+ aq + 5OH -+ 5H 2 O
where reactions 5 and 6 have C org :C sw ratios of 1:2 and 1:1, respectively. As the C org :C sw ratio decreases, greater quantities of Fe 2+ aq are produced, which would be expected to produce increasing quantities of magnetite through the reaction of Fe 2+ aq and unreduced Fe(III)-oxyhydroxides. moles reduced for every carbon mole oxidised. The model is tuned to have organic mineralisation generally describing conditions of anoxic remineralisation [START_REF] Boudreau | Diagenetic Models and Their Implementation: Modelling Transport and Reactions in Aquatic Sediments[END_REF].

[A]

The general solution for this expression was obtained assuming steady state, = 0, and the With this approach, it is straightforward to illustrate the main processes that will control porewater DIC carbon isotope values. Importantly, decreasing the sedimentation rate and initial utilisable organic carbon concentrations, and increasing the size of the marine DIC reservoir will result in heavier δ 13 C DIC values. This result is relevant for Archaean IF because there is nearly unanimous agreement that (i) the size of the DIC reservoir was significantly larger in the Archaean than in the modern ocean, (ii) most IF precursor sediments were deposited in offshore environments with limited siliciclastic input and low sedimentation rates, and (iii) proportionally more Fe(III) existed than organic carbon buried with IF precursor sediments relative to typical modern continental margin sediments [START_REF] Ridgwell | The role of the global carbonate cycle in the regulation and evolution of the Earth system[END_REF][START_REF] Bekker | Iron formation: The sedimentary product of a complex interplay among mantle, tectonic, oceanic, and biospheric processes[END_REF].

The effects of sedimentation rate, concentration of initial reactive organic carbon, and size of the marine DIC reservoir on pore-water profiles are illustrated here by comparing conditions likely to be characteristic of IF precursor sediments with those typical of modern continental-margin sediments. In this case, we use a site in the Santa Barbara Basin as representative of the continental margin sediment column. This site is noteworthy for its role in the development of early diagenetic modelling [START_REF] Berner | An idealized model of dissolved sulfate distribution in recent sediments[END_REF], but it is also a fair representation of modern marginal settings with high organic carbon loading and high rates of sulfate reduction. If we consider the modern DIC reservoir, as well as sedimentation rates and organic carbon loads characteristic of continental margin settings, δ 13 C values of pore-water DIC quickly (within upper 20 cm) become highly depleted (δ 13 C DIC < -20‰). However, with sedimentation rates and initial metabolisable organic carbon concentrations being an order of magnitude less than those typical of continental margin sediments, and a DIC reservoir an order of magnitude larger than the modern one, the pore-water DIC δ 13 C values would stabilise at values heavier than -6‰ (Fig. 15). This magnitude of offset, relative to modern marginal sediments, seems likely for IF precursor sediments. In addition, although IF sedimentation was likely highly pulsed and variable [START_REF] Bekker | Iron formation: The sedimentary product of a complex interplay among mantle, tectonic, oceanic, and biospheric processes[END_REF]Haugaard et al., 2016a), slow net accumulation rates were probably linked to a low siliciclastic flux that allowed for accumulation of pure chemical sediments. Lastly, as mentioned above, it is generally accepted that the DIC reservoir was much larger in the Archaean relative to today; in this context, an estimated ten-fold greater size compared to that of the modern reservoir seems reasonable [START_REF] Hotinski | The effectiveness of the Paleoproterozoic biological pump: A δ 13 C gradient from platform carbonates of the Pethei Group (Great Slave Lake Supergroup, NWT)[END_REF][START_REF] Ridgwell | The role of the global carbonate cycle in the regulation and evolution of the Earth system[END_REF][START_REF] Higgins | Oxygenation of the ocean and sediments: Consequences for the seafloor carbonate factory[END_REF], although a commensurate reduction in organic carbon reservoir Detailed stable Fe isotope studies of hematite, magnetite, and siderite in IFs, including in situ or small-sample isotopic analysis, also demonstrate that these minerals did not form in Fe isotope equilibrium with each other or with seawater (Fig. 16). This result potentially provides further support for in situ reduction of Fe(III) oxyhydroxides by DIR in the sediment prior to lithification, consistent with the C and O isotope data discussed above. The Fe isotope composition of Archaean seawater was likely temporally and spatially variable, reflecting, in part, variable extents of Fe(II)

oxidation (e.g., [START_REF] Rouxel | Iron isotope constraints on the Archean and Paleoproterozoic ocean redox state[END_REF]. IF minerals that formed in equilibrium with seawater should have δ 56 Fe values offset from the seawater composition based on specific mineral-fluid Fe isotope fractionation factors. In contrast, the δ 56 Fe values for coexisting oxides and carbonates, from the micron scale to bulk-sample analysis of gram-size quantities, show that Fe isotope compositions scatter closely about a 1:1 line, with some samples trending toward, but not reaching, Fe isotope equilibrium. The 1:1 line can be considered a common source of Fe for specific magnetite-hematite and magnetite-siderite pairs (e.g., W. Li et al., 2013a), taken to reflect authigenic formation of magnetite and siderite through reduction of Fe(III)-oxyhydroxide. Samples that display high δ 56 Fe values for the initial Fe(III) precipitate are interpreted to reflect a hydrothermal source of Fe based on Fe-Nd isotope relations (Fig. 6), whereas samples that have low initial δ 56 Fe values may reflect a microbial source of Fe(II) that was shuttled from a continental shelf environment (Fig. 6), assuming

Iron formations as tracers of seawater redox

Iron formations are among the most widely used lithologies to investigate the composition of the ancient oceans because their precursor minerals, such as Fe(III)-oxyhydroxides, are likely to retain the chemical signature of seawater (e.g., [START_REF] Bau | Distribution of yttrium and rare-earth elements in the Penge and Kuruman iron-formations, Transvaal Supergroup, South Africa[END_REF]Jacobsen and Pimental-Klose, 1988b;[START_REF] Bjerrum | Ocean productivity before about 1.9 Ga ago limited by phosphorus adsorption onto iron oxides[END_REF][START_REF] Bolhar | Characterization of early Archaean chemical sediments by trace element signatures[END_REF][START_REF] Robbins | Limited Zn and Ni mobility during simulated iron formation diagenesis[END_REF][START_REF] Robbins | Trace elements at the intersection of marine biological and geochemical evolution[END_REF]. Moreover, many IFs contain low concentrations of crustal-sourced elements, such as Al, Ti, Zr, Th, Hf, and Sc, implying an authigenic origin and the preservation of geochemical signatures that are uncontaminated by continental detritus.

Manganese, the most abundant redox-sensitive element in IF after Fe [START_REF] Klein | Some Precambrian banded iron-formations (BIFs) from around the world: Their age, geologic setting, mineralogy, metamorphism, geochemistry, and origin[END_REF], is potentially a very important candidate to use for tracking seawater redox changes across periods of IF deposition. To oxidise Mn(II), the oxidation state of the basin needs to be significantly higher than that required for Fe(II) oxidation. In the pre-GOE BIFs of the Transvaal Supergroup, Mn is only present in its divalent form in carbonate minerals and its concentration increases significantly stratigraphically upwards from the Kuruman to the Griquatown and Koegas IFs [START_REF] Kurzweil | Manganese oxide shuttling in pre-GOE oceans-evidence from molybdenum and iron isotopes[END_REF][START_REF] Oonk | Fraction-specific geochemistry across the Asbestos Hills BIF of the Transvaal Supergroup, South Africa: Implications for the origin of BIF and the history of atmospheric oxygen[END_REF]. This pattern indicates a relative Mn enrichment in the basin with time, either passively or due to effective reductive cycling of Mn(IV) oxyhydroxides during IF deposition.

However, at present, there is no conclusive evidence in the Transvaal IF for either the primary oxidative mechanism for episodic Mn sequestration (i.e., aerobic versus photoferrotrophic), nor for the relative role of Mn(II) incorporation into carbonates directly as Mn(II) from porefluid or indirectly through diagenetic bacterial reduction of Mn(IV) oxyhydroxides.

In addition to constraining the secular trend in the magnitude of Eu anomaly in Superior-type IF (discussed above), REE+Y studies have also focused on Ce anomalies as tracers of water column redox behaviour. These studies have drawn from a sound understanding of REE+Y cycling in modern anoxic basins. In general, oxygenated marine settings display a strong negative Ce anomaly when normalised to a shale composite (Ce (SN) ), because oxidation of Ce(III) to Ce(IV) greatly reduces Ce solubility, resulting in its preferential removal onto Mn(IV)-Fe(III)-oxyhydroxides, organic matter, and clay particles [START_REF] Byrne | Marine chemistry and geochemistry of the lanthanides[END_REF][START_REF] Bau | Oxidative scavenging of cerium on hydrous Fe oxide: Evidence from the distribution of rare earth elements and yttrium between Fe oxides and Mn oxides in hydrogenetic ferromanganese crusts[END_REF]. This process leaves the residual water column depleted in Ce, thus creating the characteristic negative Ce anomaly. In contrast, suboxic and anoxic waters lack significant negative Ce (SN) anomalies due to reductive dissolution of settling Mn(IV)-Fe(III)-oxyhydroxide particles [START_REF] German | Redox cycling of rare earth elements in the suboxic zone of the Black Sea[END_REF][START_REF] Byrne | Marine chemistry and geochemistry of the lanthanides[END_REF]. Similarly, depletion of light REE+Y develops in oxygenated waters owing to preferential removal of LREE versus HREE onto Mn(IV)-Fe(III)-oxyhydroxides and other particle-reactive surfaces. As a result, the ratio of light to heavy REEs markedly increases across redox boundaries due to reductive dissolution of Mn(IV)-Fe(III)-oxyhydroxides [START_REF] German | Redox cycling of rare earth elements in the suboxic zone of the Black Sea[END_REF][START_REF] Byrne | Marine chemistry and geochemistry of the lanthanides[END_REF].

In many Archaean and early Palaeoproterozoic IFs there are no Ce anomalies (Planavsky et al., 2010a), and thus no deviation from trivalent Ce behaviour (e.g., Fryer, 1977a;[START_REF] Bau | Rare-earth element systematics of the chemically precipitated component in early Precambrian iron formations and the evolution of the terrestrial atmospherehydrosphere-lithosphere system[END_REF][START_REF] Bau | Sources of rare-earth elements and iron in Paleoproterozoic iron-formations from the Transvaal Supergroup, South Africa: Evidence from neodymium isotopes[END_REF][START_REF] Alexander | Continentally-derived solutes in shallow Archean seawater: Rare earth element and Nd isotope evidence in iron formation from the 2.9 Ga Pongola Supergroup, South Africa[END_REF]Frei et al., 2008). This suggests that the water column from which Fe(III) oxyhydroxides precipitated was relatively anoxic, or at least below the Eh of the Ce 4+ /Ce 3+ redox couple [START_REF] Bau | Distribution of yttrium and rare-earth elements in the Penge and Kuruman iron-formations, Transvaal Supergroup, South Africa[END_REF]. It is also important to consider the rates of Ce 3+ oxidation and potential inhibition mechanisms, such as the presence of Fe 2+ and Mn 2+ (e.g., [START_REF] Moffett | The relationship between cerium and manganese oxidation in the marine environment[END_REF]. Several studies have reported cases of Ce anomalies in Archaean IF (e.g., [START_REF] Kato | Rare earth elements in Precambrian banded iron formations: Secular changes of Ce and Eu anomalies and evolution of atmospheric oxygen[END_REF], but many of these cases can be linked to analytical artefacts or the analysis of samples that have experienced supergene alteration (for details see [START_REF] Braun | Cerium anomalies in lateritic profiles[END_REF][START_REF] Valeton | Supergene alteration since the Upper Cretaceous on alkaline igneous and metasomatic rocks of the Poços de Caldas ring complex, Minas Gerais, Brazil[END_REF].

Additional work is needed to verify the existence of extensive Ce(III) oxidation in the Archaean. In this regard, a novel and potentially valuable tool for moving forward may be the use of Ce stable isotopes, which in combination with Ce anomalies, have recently been shown to provide a more nuanced understanding of redox conditions in modern hot springs and ferromanganese crusts, and hence, may be favourable for studies of strata pre-and contemporaneous to the GOE [START_REF] Nakada | Cerium stable isotope ratios in ferromanganese deposits and their potential as a paleo-redox proxy[END_REF]. Additionally, it is important to note that Ce anomalies need not develop solely within the water column of a basin. For instance, Ce anomalies can develop in groundwater systems [START_REF] Johanneson | Do Archean chemical sediments record ancient seawater rare earth element patterns?[END_REF].

There also appear to be differences in trivalent REE+Y behaviour in IFs before and after the permanent rise of atmospheric oxygen (Fig. 17). Late Palaeoproterozoic IFs show significant ranges in light-to-heavy REE (Pr/Yb (SN) ) and Y/Ho ratios, which likely reflect variable fractionation of REE+Y by Mn(IV)-Fe(III)-oxyhydroxide precipitation and dissolution (Planavsky et al., 2010a).

This interpretation implies deposition of late Palaeoproterozoic IF at ca. 1.88 Ga in basins having varied redox conditions and a strong redoxcline separating an upper oxic water column from deeper waters that were suboxic to anoxic [START_REF] Planavsky | Iron-oxidizing microbial ecosystems thrived in late Paleoproterozoic redox-stratified oceans[END_REF]. These ranges are similar to those seen in modern anoxic basins. In contrast, most Archaean and early Palaeoproterozoic IFs deposited before the rise of atmospheric oxygen are characterised by light REE+Y depletion (Planavsky et al., 2010a). This consistent depletion in LREE suggests the lack of a discrete redoxcline during this time period and points toward O 2 -independent Fe(II) oxidation mechanisms.

Neoarchaean-Palaeoproterozoic Iron Formations

The Neoarchaean to middle Palaeoproterozoic (~2.8-1.9 Ga) marks one of the most important periods in Earth history, with several major interlinked environmental and biological evolutionary events. These include, amongst others, the oxygenation of the atmosphere -the GOE [START_REF] Bekker | Dating the rise of atmospheric oxygen[END_REF][START_REF] Konhauser | Aerobic bacterial pyrite oxidation and acid rock drainage during the Great Oxidation Event[END_REF][START_REF] Gumsley | Timing and tempo of the Great Oxidation Event[END_REF]. The GOE represents a transition from an atmosphere that was essentially devoid of free oxygen (O 2 << 10 -5 PAL) to one with O 2 concentrations >10 -5 PAL, where estimates of oxygen levels are based on modelling massindependent S isotope (S-MIF) variations [START_REF] Pavlov | Mass-independent fractionation of sulfur isotopes in Archean sediments: Strong evidence for an anoxic Archean atmosphere[END_REF]. In the rock record, this event is manifested by a number of changes (see [START_REF] Farquhar | Geological constraints on the origin of oxygenic photosynthesis[END_REF][START_REF] Farquhar | Geological and geochemical constraints on Earth's early atmosphere[END_REF], for reviews), including (1) loss of easily oxidisable detrital uraninite, pyrite, and siderite from fluvial siliciclastic sediments at ca. 2.4 Ga (e.g., [START_REF] Rasmussen | Redox state of the Archean atmosphere: Evidence from detrital heavy metals in ca. 3250-2750 Ma sandstones from the Pilbara Craton, Australia[END_REF][START_REF] England | Paleoenvironmental significance of rounded pyrite in siliciclastic sequences of the Late Archean Witwatersrand Basin: Oxygendeficient atmosphere or hydrothermal alteration?[END_REF][START_REF] Hofmann | Multiple sulphur and iron isotope composition of detrital pyrite in Archaean sedimentary rocks: A new tool for provenance analysis[END_REF][START_REF] Johnson | O 2 constraints from Paleoproterozoic detrital pyrite and uraninite[END_REF]; (2) loss of iron from ancient soil horizons (paleosols) older than ca. 2.4 Ga because of greater iron solubility under reducing conditions (e.g., [START_REF] Rye | Paleosols and the evolution of atmospheric oxygen: A critical review[END_REF]; (3) the appearance of red beds after ca. 2.3 Ga (e.g., [START_REF] Roscoe | Huronian rocks and uraniferous conglomerates in the Canadian Shield[END_REF][START_REF] Chandler | Proterozoic redbed sequences of Canada[END_REF][START_REF] Melezhik | Emergence of the aerobic biosphere during the Archean-Proterozoic transition: Challenges of future research today[END_REF][START_REF] Bankole | Origin of red beds in the Paleoproterozoic Franceville Basin, Gabon, and implications for sandstone-hosted uranium mineralization[END_REF]; (4) the loss of S-MIF in sulfide and sulfate minerals in sedimentary rocks deposited after 2.3 Ga [START_REF] Farquhar | Atmospheric influence of Earth's earliest sulfur cycle[END_REF][START_REF] Bekker | Dating the rise of atmospheric oxygen[END_REF][START_REF] Papineau | Multiple sulfur isotopes from Paleoproterozoic Huronian interglacial sediments and the rise of atmospheric oxygen[END_REF][START_REF] Partridge | Pyrite paragenesis and multiple sulfur isotope distribution in late Archean and early Paleoproterozoic Hamersley Basin sediments[END_REF][START_REF] Guo | Reconstructing Earth's surface oxidation across the Archean-Proterozoic transition[END_REF][START_REF] Williford | Constraining atmospheric oxygen and seawater sulfate concentrations during Paleoproterozoic glaciation: In situ sulfur three-isotope microanalysis of pyrite from the Turee Creek Group, Western Australia[END_REF][START_REF] Luo | Rapid oxygenation of Earth's atmosphere 2.33 billion years ago[END_REF][START_REF] Gumsley | Timing and tempo of the Great Oxidation Event[END_REF]; and ( 5) the increase in Cr and U contents in IF at 2.45 Ga that records the onset of oxidative continental weathering [START_REF] Konhauser | Aerobic bacterial pyrite oxidation and acid rock drainage during the Great Oxidation Event[END_REF]Partin et al., 2013a).

Iron formations deposited before the GOE

Iron formations deposited between 2.60 and 2.45 Ga are the most laterally extensive and economically valuable of any in Earth history. The 2.60 Ga Marra Mamba Formation of the Hamersley Province in Western Australia is one of the thickest IFs in the geologic record. It was deposited in a deep-water basinal setting adjacent to a passive margin during the sea-level highstand of the Marra Mamba basin, and has an average outcrop thickness of about 210 m [START_REF] Trendall | The iron formations of the Precambrian Hamersley Group, Western Australia with special reference to the associated crocidolite[END_REF][START_REF] Krapež | Hydrothermal and resedimented origins of the precursor sediments to banded iron formations: Sedimentological evidence from the early Palaeoproterozoic Brockman Supersequence of Western Australia[END_REF]. The Marra Mamba IF carries a pronounced positive Eu anomaly [START_REF] Alibert | Rare element and neodymium isotopic compositions of the banded iron-formations and associated shales from Hamersley, Western Australia[END_REF], suggesting a strong hydrothermal imprint on the REE systematics during its deposition, although, as noted above (Fig. 6), Nd isotope data suggest a mixture of continental and hydrothermal REE+Y sources. Further, the Marra Mamba IF is dominated by magnetite, ankerite, and fibrous silicate [START_REF] Baur | Millimeter-scale variations of stable isotope abundances in carbonates from banded iron-formations in the Hamersley Group of Western Australia[END_REF], which suggests excess organic carbon was present during diagenesis. Support for this interpretation is found in negative δ 13 C values that increase with increasing magnetite abundance [START_REF] Baur | Millimeter-scale variations of stable isotope abundances in carbonates from banded iron-formations in the Hamersley Group of Western Australia[END_REF].

Major IF deposition occurred between 2.50-2.40 Ga in association with a series of mantle plume breakout events (e.g., [START_REF] Heaman | Global mafic volcanism at 2.45 Ga: Remnants of an ancient large igneous province?[END_REF] and was immediately preceded by the final stage in the supercontinent assembly of the Gawler Craton (South Australia), North China Craton, East Antarctica Craton, and in the central Indian tectonic zone of India [START_REF] Barley | Late Archean to early Paleoproterozoic global tectonics, environmental change and the rise of atmospheric oxygen[END_REF]. Deposition of these IFs occurred on a reactivated continental margin [START_REF] Krapež | Hydrothermal and resedimented origins of the precursor sediments to banded iron formations: Sedimentological evidence from the early Palaeoproterozoic Brockman Supersequence of Western Australia[END_REF], followed by a rise in sea level. [START_REF] Isley | Plume-related mafic volcanism and the deposition of banded iron formation[END_REF] estimated that during those 100 m.y., perhaps as much as 60% of the global volume of preserved IF was deposited.

The largest IFs by mass are associated with the Brockman Iron Formation in the Hamersley Group of the Pilbara Craton in Western Australia [START_REF] Isley | Hydrothermal plumes and the delivery of iron to banded iron formation[END_REF][START_REF] Trendall | Precambrian iron-formation[END_REF]. The Brockman Iron Formation is divided into four lithostratigraphic units, namely the lowermost Dales Gorge Member (BIF), the Whaleback Shale Member, the Joffre Member (BIF), and the uppermost Yandicoogina Shale Member (Fig. 18). Overlying the Brockman Iron Formation are two more BIF units, the Weeli Wolli Formation and the Boolgeeda Iron Formation. The best age estimates for the deposition of the Dales Gorge and Joffre members are 2494-2464 Ma (Trendall et al., 2004) and 2454±3 Ma [START_REF] Pickard | SHRIMP U-Pb zircon ages of tuffaceous mudrocks in the Brockman Iron Formation of the Hamersley Range, Western Australia[END_REF], respectively. The consensus view is that the Brockman Iron Formation was deposited on a large, stable, and clastic-starved continental platform, which was only influenced by episodic inputs of fine-grained tuffaceous detritus (e.g., [START_REF] Ewers | Studies of the Dales Gorge Member of the Brockman Iron Formation, Western Australia[END_REF][START_REF] Gross | Tectonic systems and the deposition of iron-formation[END_REF]. [START_REF] Barley | Emplacement of a large igneous province as a possible cause of banded iron formation 2.45 billion years ago[END_REF] suggested that deposition of the Hamersley IFs was possibly linked to major submarine magmatic plume activity in the form of a large igneous province (LIP). [START_REF] Morris | Genetic modelling for banded iron-formations of the Hamersley Group, Pilbara Craton, Western Australia[END_REF] also proposed that the depositional environment for the Hamersley BIFs included a steady source of silica and iron, and a water depth shallow enough to form the large carbonate platforms.

From a geochemical standpoint, the Brockman Iron Formation portrays characteristics common to pre-GOE IFs. Typical seawater REE+Y patterns are displayed by the Dales Gorge Member, including low Pr/Yb SN (0.08-0.3), low Sm/Yb SN (0.1-0.4), positive La anomalies, positive Eu anomalies, and superchondritic Y/Ho ratios [START_REF] Pecoits | Petrography and geochemistry of the Dales Gorge banded iron formation: Paragenetic sequence, source and implications for palaeo-ocean chemistry[END_REF]. However, the Dales Gorge Member lacks a true negative Ce anomaly, suggesting low-oxygen conditions at the time of deposition. The Dales Gorge Member is also composed of relatively pure chemical precipitates, as supported by low Al, Ti, Th, Zr, Hf, and Sc concentrations [START_REF] Pecoits | Petrography and geochemistry of the Dales Gorge banded iron formation: Paragenetic sequence, source and implications for palaeo-ocean chemistry[END_REF]. Based on bulksample analysis, [START_REF] Becker | Oxygen isotope study of a Precambrian banded ironformation, Hamersley Range, Western Australia[END_REF] noted isotopic exchange between minerals within the Dales Gorge Member at burial metamorphic temperatures of 270 to 310 o C, and suggested that oxygen isotopes in quartz and carbonate may have equilibrated completely. Hematite, on the other hand, underwent negligible change in oxygen isotopes. Detailed in situ oxygen isotope analysis (W. Li et al., 2013a), as discussed above (Fig. 8), documents multiple generations of oxides and suggests a primary precipitation temperature as low as ~25 o C.

The low abundance of Al 2 O 3 and TiO 2 within the Joffre Member also suggests limited continental epiclastic input. However, the Joffre Member contains many stilpnomelane granules of compositions that appear to record a mixture of volcanic (pyroclastic) and chemical components, thus making it a poor chemical proxy for atmospheric conditions at the time of deposition (Haugaard et al., 2016b). Compared to the underlying Dales Gorge Member, the Joffre Member has significantly higher REE concentrations [START_REF] Pecoits | Petrography and geochemistry of the Dales Gorge banded iron formation: Paragenetic sequence, source and implications for palaeo-ocean chemistry[END_REF]Haugaard et al., 2016b (Johnson et al., 2008b;[START_REF] Pecoits | Petrography and geochemistry of the Dales Gorge banded iron formation: Paragenetic sequence, source and implications for palaeo-ocean chemistry[END_REF]W Li et al., 2013a;Haugaard et al., 2016b).

The Weeli Wolli Formation displays a marked increase in Cr concentrations, interpreted by [START_REF] Konhauser | Aerobic bacterial pyrite oxidation and acid rock drainage during the Great Oxidation Event[END_REF] as a remnant signature of bacterial pyrite oxidation and associated acid rock drainage. Aside from this work, few geochemical analyses have been carried out on this formation.

A lack of geochemical data is also apparent for the Boolgeeda Iron Formation, which was deposited at ~2.45 Ga and represents the terminal IF deposit in the Hamersley Basin. High-resolution chemostratigraphic profiling of both the Weeli Wolli and the Boolgeeda IF are currently being perused, and it is anticipated that these data will help elucidate the transition from anoxic to oxic conditions on early Earth, and potentially develop geochemical ties to time-correlative basins worldwide.

The major IFs of South Africa are developed in the Griqualand West Basin of the 2-6-2.4 Ga Transvaal Supergroup, and comprise the Kuruman and Griquatown formations and BIFs of the Koegas Subgroup (Fig. 19). [START_REF] Beukes | Origin and paleoenvironmental significance of major iron formations at the Archean-Paleoproterozoic boundary[END_REF] presented detailed correlations between the Hamersley and Transvaal BIFs, and suggest that these sequences reflect deposition in the same basin, as originally proposed by [START_REF] Button | Transvaal and Hamersley basins-review of basin development and mineral deposits[END_REF] and then [START_REF] Cheney | Sequence stratigraphy and plate tectonic significance of the Transvaal succession of southern Africa and its equivalent in Western Australia[END_REF], although it cannot be ruled out that they were two separate, but coeval basins. The contact-metamorphosed Penge BIF of the eastern Transvaal Basin is considered to be the lithostratigraphic equivalent of the Kuruman BIF and is not discussed further here. In the Griqualand West Basin, the BIFs constitute the upper Ghaap Group of the lower Transvaal Supergroup (Asbestos Hills Subgroup), whereas carbonate rocks of the Campbellrand Subgroup dominate the lower Ghaap Group [START_REF] Beukes | Origin and paleoenvironmental significance of major iron formations at the Archean-Paleoproterozoic boundary[END_REF]). The regional transition from the stromatolitic Gamohaan Formation upward into the Kuruman BIF is continuous and indicates a progressive deepening of the basin [START_REF] Klein | Geochemistry and sedimentology of a facies transition from limestone to iron-formation deposition in the early Proterozoic Transvaal Supergroup, South Africa[END_REF][START_REF] Sumner | Carbonate precipitation and oxygen stratification in Late Archean seawater as deduced from facies and stratigraphy of the Gamohaan and Frisco formations, Transvaal Supergroup, South Africa[END_REF]. The rhythmically micro-banded Kuruman BIF was deposited largely in an open-shelf environment below wave base and contains multiple cycles of shallowing-upward sequences [START_REF] Beukes | Sedimentology of the Kuruman and Griquatown iron-formations, Transvaal Supergroup, Griqualand West, South Africa[END_REF]. A SHRIMP U/Pb zircon age, determined on stilpnomelane-rich tuffaceous mudrocks, constrains the age of the top part of the Kuruman BIF at ca. 2.46 Ga [START_REF] Pickard | SHRIMP U-Pb zircon ages for the Palaeoproterozoic Kuruman Iron Formation, Northern Cape Province, South Africa: Evidence for simultaneous BIF deposition on Kaapvaal and Pilbara Cratons[END_REF], thus making it broadly contemporaneous with the Brockman IF in Australia.

The predominantly clastic-textured, granular Griquatown IF is interpreted to have formed in a shallow water, storm-dominated depositional environment [START_REF] Beukes | Sedimentology of the Kuruman and Griquatown iron-formations, Transvaal Supergroup, Griqualand West, South Africa[END_REF]Beukes and Klein, 1990). The transition from Kuruman to Griquatown IF is gradual, characterised by intercalated microbanded and granular facies, and in itself is intriguing given that transitions from microbanded BIF to granular GIFs are uncommon in the pre-GOE successions (Beukes and Klein, 1990). The clastic texture in the Griquatown Formation is manifested by mm-scale granules of varying composition, typically in the same bed, developed locally into well-sorted grainstones [START_REF] Beukes | Origin and paleoenvironmental significance of major iron formations at the Archean-Paleoproterozoic boundary[END_REF]. Other characteristic textures include soft-sediment deformation features and dmscale chert lenses and pods that contain sprays of riebeckite in lower stratigraphic portions of the Griquatown GIF.

The Koegas Subgroup only develops in the deeper, southwestern part of the Griqualand West Basin. Its absence in the shallower parts -where the Griquatown GIF is directly overlain by the Makganyene diamictite -is explained by palaeobathymetric considerations [START_REF] Moore | Deconstructing the Transvaal Supergroup, South Africa: Implications for Palaeoproterozoic palaeoclimate models[END_REF][START_REF] Polteau | The geology and geochemistry of the Palaeoproterozoic Makganyene diamictite[END_REF]. The Koegas BIF represents sea-level fluctuations with alternating regressive and transgressive cycles. Iron formation deposition is thought to have taken place during transgressions, punctuated by siliciclastic deposition during regressions [START_REF] Schröder | From BIF to red beds: Sedimentology and sequence stratigraphy of the Paleoproterozoic Koegas Subgroup (South Africa)[END_REF].

Magnetite and Fe carbonates are the major iron minerals within the South African IFs (Beukes and Klein, 1990;[START_REF] Kurzweil | Manganese oxide shuttling in pre-GOE oceans-evidence from molybdenum and iron isotopes[END_REF]. Although hematite micro-particles are present, they are volumetrically extremely rare [START_REF] Klein | Geochemistry and sedimentology of a facies transition from limestone to iron-formation deposition in the early Proterozoic Transvaal Supergroup, South Africa[END_REF]. The IFs are also generally sulfurpoor and contain minimal detrital input as recorded in very low bulk Al and Ti abundances [START_REF] Beukes | Sedimentology of the Kuruman and Griquatown iron-formations, Transvaal Supergroup, Griqualand West, South Africa[END_REF]Beukes and Klein, 1990;[START_REF] Horstmann | Chemical composition of banded iron-formations of the Griqualand West Sequence, Northern Cape Province, South Africa, in comparison with other Precambrian iron formations[END_REF]. Whole-rock REE+Y patterns and their interpretations are in good agreement with those of the aforementioned Hamersley IFs of Australia [START_REF] Bau | Distribution of yttrium and rare-earth elements in the Penge and Kuruman iron-formations, Transvaal Supergroup, South Africa[END_REF]Planavsky et al., 2010a). Manganese -deposited as discrete Mnoxide layers in the stratigraphically younger Hotazel Formation -is incorporated entirely in iron carbonate minerals in both the Griquatown and Koegas IFs, where its bulk-rock abundance locally reaches percent levels [START_REF] Tsikos | Petrography and geochemistry of the Paleoproterozoic Hotazel Iron-Formation, Kalahari manganese field, South Africa: Implications for Precambrian manganese metallogenesis[END_REF][START_REF] Johnson | Manganese mineralogy and diagenesis in the sedimentary rock record[END_REF][START_REF] Kurzweil | Manganese oxide shuttling in pre-GOE oceans-evidence from molybdenum and iron isotopes[END_REF].

Major IF deposits in Canada are hosted in several 2.9-2.8 Ga and 2.7-2.6 Ga greenstone belts on the Superior, Slave, Rae, and Hearne cratons. Several of these are Algoma-type IF successions that contain major gold deposits, including the ~2.71 Ga Woodburn Lake greenstone belt, the ~2.97

North Caribou greenstone belt, the ~2.7 Ga Bearmore-Geraldton greenstone belt, the ~2.7-2.6 Ga Rankin Inlet greenstone belt (Goucerol et al., 2016), the ~2.7 Ga Hope Bay greenstone belt [START_REF] Sherlock | Volcanic stratigraphy, geochronology, and gold deposits of the Archean Hope Bay Greenstone Belt, Nunavut, Canada[END_REF], and the ~2.62 Ga BIF from Damoti Lake and Back River (Haugaard et al., 2017a,b). Recently, important petrological and geochemical work was carried out on the Slave Craton IF that consists of Mesoarchaean BIF of the ~2.85 Ga Central Slave Cover Group deposited in a continental rift-setting (Haugaard et al., 2016a), and Neoarchaean BIF of the 2.62 Ga Slemon Group (Haugaard et al., 2017a) deposited in a back-arc basin (Haugaard et al., 2017b).

Iron formations in the Quadrilátero Ferrífero region in Brazil, Middleback Ridge (Gawler Craton) in South Australia, Nimba and Simandou Ranges in Liberia and Guinea, Krivoy Rog area in Ukraine, and Kursk Magnetic Anomaly (KMA) region in Russia are broadly similar in age based on available geochronological and chemostratigraphical constraints, and also display similar patterns of macrobanding [START_REF] Prilutzky | Reconstruction of environmental conditions of early Proterozoic carbonate deposits of the Ukrainian and Baltic Shields based on isotopic studies[END_REF][START_REF] Kulik | Lithological and geochemical evidence of Fe and Mn pathways during deposition of Lower Proterozoic banded iron formation in the Krivoy Rog Basin (Ukraine)[END_REF][START_REF] Egal | Late Eburnean granitization and tectonics along the western and northwestern margin of the Archean Kenema-Man domain (Guinea, West African Craton)[END_REF][START_REF] Bekker | Chemostratigraphy of Paleoproterozoic carbonate successions of the Wyoming Craton: Tectonic forcing of biogeochemical change?[END_REF][START_REF] Phillips | U-Th-Pb detrital zircon geochronology from the southern Prince Charles Mountains, East Antarctica -Defining the Archaean to Neoproterozoic Ruker Province[END_REF][START_REF] Spier | Geochemistry and genesis of the banded iron formations of the Cauê Formation, Quadrilátero Ferrífero, Minas Gerais, Brazil[END_REF][START_REF] Szpunar | Isotopic and geochemical constraints on the Paleoproterozoic Hutchison Group, southern Australia: Implications for Paleoproterozoic continental reconstructions[END_REF]. These IFs were also deposited on reactivated continental margins, and are separated by prominent unconformities from overlying Palaeoproterozoic sequences, which correspond to a long gap in sedimentation following the supercontinent assembly at ca. 2.40 Ga.

Minor iron formation deposition after the GOE and before ca. 1.88 Ga

Giant IFs were not deposited between ca. 2.40 and 1.88 Ga, but oolitic hematitic ironstones of the 2.32 Ga lower Timeball Hill Formation in South Africa were deposited in shallow water, above fairweather wave base [START_REF] Schweigart | Genesis of the iron ores of the Pretoria Series, South Africa[END_REF][START_REF] Dorland | Paleoproterozoic laterites, red beds and ironstones of the Pretoria Group with reference to the history of atmospheric oxygen[END_REF]. Few geochemical data are available for this unit, partly because it contains significant amounts of siliciclastic material. Its deposition at ca. 2.32

Ga [START_REF] Hannah | Primitive Os and 2316 Ma age for marine shale: Implications for Paleoproterozoic glacial events and the rise of atmospheric oxygen[END_REF]) may coincide with a magmatic event at that time (e.g., Eriksson et al., 1994a, b;[START_REF] Fetter | U-Pb and Sm-Nd geochronological constraints on the crustal evolution and basement architecture of Ceará state, NW Borborema Province, NE Brazil: Implications for the existence of the Paleoproterozoic supercontinent "Atlantica[END_REF][START_REF] Berman | Tectonometamorphism at ca. 2.35 and 1.85 Ga in the Rae Domain, western Churchill Province, Nunavut, Canada: Insight from structural, metamorphic, and in situ geochronologic analysis of the southwestern Committee Bay Belt[END_REF][START_REF] Hartlaub | Circa 2.3 Ga magmatism of the Arrowsmith Orogeny, Uranium City Region, Western Churchill Craton[END_REF][START_REF] Rasmussen | Correlation of Paleoproterozoic glaciations based on U-Pb zircon ages for tuff beds in the Transvaal and Huronian supergroups[END_REF][START_REF] Partin | Filling in the juvenile magmatic gap: Evidence for uninterrupted Paleoproterozoic plate tectonics[END_REF]. However, the significance of this magmatic event is poorly known. Correlative

IFs are not documented on other continents.

The 2.43 Ga Hotazel Formation, whose date of formation has recently been revised from earlier estimates of 2.2 Ga [START_REF] Gumsley | Timing and tempo of the Great Oxidation Event[END_REF], contains the largest known manganese deposit in the world [START_REF] Tsikos | Deposition, diagenesis, and secondary enrichment of metals in the Paleoproterozoic Hotazel Iron-Formation, Kalahari manganese field, South Africa[END_REF]. The manganese (and iron) formation lies above basaltic andesites of the Ongeluk Formation and has been interpreted previously as a volcanogenic deposit [START_REF] Cornell | A volcanic-exhalative origin for the world's largest (Kalahari) manganese field[END_REF]. The iron-and manganese-rich succession of the Hotazel Formation consists of three symmetric cycles of Mn-rich carbonate-braunite lutite, intercalated with magnetite-rich microbanded BIF [START_REF] Tsikos | Petrography and geochemistry of the Paleoproterozoic Hotazel Iron-Formation, Kalahari manganese field, South Africa: Implications for Precambrian manganese metallogenesis[END_REF]. The entire sequence is thought to have been deposited in a continental shelf environment that was starved of terrigenous siliciclastic input [START_REF] Schneiderhan | The chemostratigraphy of a Paleoproterozoic MnFe-and BIF succession -the Voëlwater Subgroup of the Transvaal Supergroup in Griqualand West, South Africa[END_REF]. The Hotazel Mn-Fe deposits lack positive Eu anomalies, pointing to an extremely dilute hydrothermal signal at best. The significance of weak Ce anomalies remains contentious considering co-existing positive La anomalies [START_REF] Tsikos | Petrography and geochemistry of the Paleoproterozoic Hotazel Iron-Formation, Kalahari manganese field, South Africa: Implications for Precambrian manganese metallogenesis[END_REF]. Whereas the Ongeluk volcanics cannot be discounted as a possible source of at least some of the Fe and Mn in the basin through low-temperature seafloor alteration, recent applications of combined iron isotope geochemistry and Mn/Fe ratios suggest that enrichment in the Hotazel palaeobasin in Mn and isotopically light Fe could have ensued progressively through prior protracted deposition of Mnpoor, high d 56 Fe BIF [START_REF] Tsikos | Iron isotopes constrain biogeochemical redox cycling of iron and manganese in a Paleoproterozoic stratified basin[END_REF]. Indeed, the new age constraints for the Hotazel Formation [START_REF] Gumsley | Timing and tempo of the Great Oxidation Event[END_REF], and the Fe isotope and Mn-enriched signature of the uppermost Griquatown and Koegas IFs in the Transvaal Supergroup [START_REF] Kurzweil | Manganese oxide shuttling in pre-GOE oceans-evidence from molybdenum and iron isotopes[END_REF][START_REF] Oonk | Fraction-specific geochemistry across the Asbestos Hills BIF of the Transvaal Supergroup, South Africa: Implications for the origin of BIF and the history of atmospheric oxygen[END_REF], lend strong support to this interpretation.

The 2.3-2.2 Ga Yuanjiacun BIF, located in Shanxi Province, is likely the most representative Superior-type IF in the North China Craton (NCC) and arguably the largest and most extensive IF (>10 Gt) deposited during this time period [START_REF] Li | Characteristics of rare earth elements and geological significations of BIFs from Jidong, Wutai and Lüliang area[END_REF][START_REF] Hou | Geochemistry and Si-O-Fe isotope constraints on the origin of banded iron formations of the Yuanjiacun Formation, Lvliang Group, Shanxi, China[END_REF]. The transition from deep-water carbonate facies to shallow-water oxide facies is accompanied by a change in mineralogical composition from siderite through magnetite-ankerite and magnetite-stilpnomelaneassemblages in the transition zone, to magnetite, and then hematite near the palaeoshore. The deposition of BIF immediately following the onset of the GOE confirms that deep marine waters were still episodically ferruginous at that time, but that shallow waters were sufficiently oxygenated such that Fe(II) oxidation no longer needed to be directly tied to proximal and in situ cyanobacterial activity [START_REF] Wang | Depositional environment of the Paleoproterozoic Yuanjiacun banded iron formation in Shanxi Province, China[END_REF][START_REF] Wang | Decoupled sources of the 2.3-2.2 Ga Yuanjiacun banded iron formation: Implications for the Nd cycle in Earth's early oceans[END_REF].

Hematitic oolites and hematite-rich sandstones continued to be deposited in shallow-marine environments during the ca. 2.25 to 2.10 Ga Lomagundi carbon isotope excursion in South Africa (Silverton Formation; [START_REF] Schweigart | Genesis of the iron ores of the Pretoria Series, South Africa[END_REF] and on the Kola Peninsula in Russia (Kuetsjärvi Sedimentary Formation; Akhmedov, 1972a). The IFs deposited during the Lomagundi excursion include those within the Ijil Group, Mauritania [START_REF] Bronner | Precambrian banded iron-formations of the Ijil Group (Kediat Ijil, Reguibat Shield, Mauritania)[END_REF]) and the Lomagundi Group, Zimbabwe [START_REF] Master | Stratigraphy, tectonic setting, and mineralization of the Early Proterozoic Magondi Supergroup, Zimbabwe: A review[END_REF]. The former belongs to the 2.20-2.10 Ga Birimian Basin in West Africa, which also contains iron and manganese formations in the Francevillian Basin, Gabon [START_REF] Leclerc | Geology and genesis of the Moanda manganese deposits, Republic of Gabon[END_REF] and in the Nigerian schist belts [START_REF] Mücke | The Nigerian manganese-rich iron-formations and their host rocks-from sedimentation to metamorphism[END_REF]. The latter developed within the ca. 2.2-2.1 Ga Magondi Basin [START_REF] Master | Stratigraphy, tectonic setting, and mineralization of the Early Proterozoic Magondi Supergroup, Zimbabwe: A review[END_REF][START_REF] Master | A review of the stratigraphy and geological setting of the Palaeoproterozoic Magondi Supergroup, Zimbabwe -type locality for the "Lomagundi" carbon isotope excursion[END_REF]. Volcanic-hosted IFs deposited during the Lomagundi carbon isotope excursion are known in Brazil (e.g., Aimbė Formation, Guarinos Group; Resende and Jost, 1995; Itapicuru Complex of the Rio Itapicuru greenstone belt; Dalton de [START_REF] Dalton De Souza | Guia de excursão -geologia do segmento norte do orógeno Itabuna-Salvador-Curaçá[END_REF], and Norway (Iddjajav'ri Group, Karasjok greenstone belt; [START_REF] Often | The early Proterozoic Karasjok Greenstone Belt, Norway: A preliminary description of lithology, stratigraphy and mineralization[END_REF].

Iron formations deposited between 2.40 and 2.10 Ga are distinctly different in scale, and in most cases, the depositional setting differs from those in the Archaean. It seems that either the deepocean redox state was too oxidised, relative to the strength of hydrothermal input, to form giant sedimentary iron deposits during this time interval, or that marine sulfate levels were high enough that the sulfide produced during bacterial sulfate reduction exceeded the hydrothermal Fe(II) supply -in this case iron sulfides would precipitate instead of iron oxides (see [START_REF] Kump | Hydrothermal Fe fluxes during the Precambrian: Effect of low oceanic sulfate concentrations and low hydrostatic pressure on the composition of black smokers[END_REF] for further discussion). Several independent lines of evidence point toward high (mM) sulfate levels during the Lomagundi Event [START_REF] Schröder | Rise in seawater sulphate concentration associated with the Paleoproterozoic positive carbon isotope excursion: Evidence from sulphate evaporites in the ~2.2-2.1 Gyr shallow-marine Lucknow Formation, South Africa[END_REF]Planavsky et al., 2012b;[START_REF] Scott | Pyrite multiple-sulfur isotope evidence for rapid expansion and contraction of the early Paleoproterozoic seawater sulfate reservoir[END_REF], consistent with the idea that growth of the marine sulfate reservoir consumed the hydrothermal Fe(II) flux, which ultimately exerted a first-order control on the distribution and abundance of IF.

Following the end of the Lomagundi carbon isotope excursion at ca. 2.10 Ga, small Algomatype IFs were deposited in several basins in North America (e.g., Homestake Iron Formation, Black Hills, South Dakota; Frei et al., 2008) and Finland [START_REF] Paakola | The volcanic complex and associated manganiferous iron formation of the Porkonen-Pahtavaara area in Finnish Lapland[END_REF][START_REF] Laajoki | On the geology and geochemistry of the Precambrian iron formations in Väyrylänkylä, South Puolanka area, Finland[END_REF].

Oolitic hematitic ironstone is also present in the Kolasjoki Formation, Kola Peninsula, Russia (Akhmedov, 1972b). Finally, economic-grade, shallow-water IFs of the 2.05 Ga Costa Sena Group (underlying the Espinhaço Group) and likely correlative Serra do Sapo Formation of the Serra de Serpentina Group were deposited in an intracratonic basin in central Brazil [START_REF] Rolim | The Orosirian-Statherian banded iron formation-bearing sequences of the southern border of the Espinhaco Range, southeast Brazil[END_REF].

Combined, these data tentatively suggest that dynamic ocean redox conditions were established in the aftermath of the GOE, with periodic upwelling of Fe(II) into shallow-water settings above storm and fair-weather wave base. However, it is also possible that these small, locally-developed IFs were linked to local, continental sources of iron (e.g., Fe-rich groundwater) rather than to upwelling of ferruginous waters. This alternative explanation has not been tested so far and requires further consideration.

An exciting, and perhaps unexpected, finding to come out of some recent studies on secular trends in trace metal contents of IFs (and black shales) is that Earth's redox state during the Precambrian may have been much more complex than previously thought. Partin et al. (2013a, b) recently showed that directly following a dramatic increase in the oceanic uranium reservoir during the GOE, the U reservoir decreased significantly after ca. 2.10 Ga, resulting in a U content in IFs and shales that was only marginally higher than pre-GOE levels. Given that the oxygen content of the oceans directly controls seawater U concentrations once oxidative continental weathering was established, the observed decrease could be due to a subsequent drop in the level of atmospheric and seawater oxygen following the Lomagundi event. It seems plausible that net organic carbon burial associated with oxygen production during the Lomagundi event later became an oxygen sink as organic matter became oxidised, driving oxygen to low levels that may have persisted for some hundreds of million years thereafter [START_REF] Bekker | Oxygen overshoot and recovery during the early Paleoproterozoic[END_REF][START_REF] Canfield | Oxygen dynamics in the aftermath of the Great Oxidation of Earth's atmosphere[END_REF].

The ca. 1.88 Ga resurgence in iron formations

Extensive and large IFs reappeared after an approximately 500 m.y. gap, at about 1.88 Ga. These successions contain an abundance of GIF, relative to Archaean and early Palaeoproterozoic BIF.

The most extensive 1.88 Ga GIF occur in North America along the southern, and eastern margins of the Superior Craton (Animikie and Mistassini basins and Labrador Trough; [START_REF] Simonson | Origin and evolution of large Precambrian iron formations[END_REF] and in Western Australia (Nabberu Basin; [START_REF] Rasmussen | Deposition of 1.88-billion-year-old iron formations as a consequence of rapid crustal growth[END_REF]. These GIFs are coeval with emplacement of a ca. 1.88 Ga ultramafic to mafic LIP [START_REF] Heaman | Precise U-Pb zircon ages for the Molson dyke swarm and the Fox River sill: Constraints for Early Proterozoic crustal evolution in northeastern Manitoba, Canada[END_REF][START_REF] Heaman | Timing and geochemistry of 1.88 Ga Molson Igneous Events, Manitoba: Insights into the formation of a craton-scale magmatic and metallogenic province[END_REF][START_REF] Hulbert | U-Pb zircon and Re-Os isotope geochronology of mineralized ultramafic intrusions and associated nickel ores from the Thompson nickel belt, Manitoba, Canada[END_REF] that is potentially related to a mantle plume breakout event during the early assembly of the supercontinent Nuna/Columbia [START_REF] Hamilton | Wide-spread and short-lived 1870 Ma mafic magmatism along the northern Superior Craton margin: EOS Transactions[END_REF]; for an opposing view, see [START_REF] Heaman | Timing and geochemistry of 1.88 Ga Molson Igneous Events, Manitoba: Insights into the formation of a craton-scale magmatic and metallogenic province[END_REF]. Recognised now to be correlative based on high-precision geochronology (Findlay et al., 1995;[START_REF] Machado | U-Pb ages for magmatism and deformation in the New Québec Orogen[END_REF]Fralick et al., 2002;[START_REF] Schneider | Age of volcanic rocks and syndepositional iron formations, Marquette Range Supergroup: Implications for the tectonic setting of Paleoproterozoic iron formations of the Lake Superior region[END_REF][START_REF] Stott | Archean granitoid geochronology and interpretations, Hudson Bay Lowland, in: Summary of Field Work and Other Activities[END_REF], the North American GIFs extend discontinuously at the surface for more than 3000 km along the southern and eastern margins of the Superior Craton, from Ontario, through Minnesota, Wisconsin, and Michigan, to Québec (Mistassini Basin) and the Labrador Trough. Despite poor exposure, these iron deposits are easily traceable by magnetic anomalies. Correlative and texturally similar GIF are also recognised in the northern part of the Superior Craton in the Hudson Bay region (Richmond Bay and Belchers islands) and in the Sutton Inliers, and are considered to have been deposited in extensional basins coeval with submarine basaltic volcanism (Fralick et al., 2002;[START_REF] Schulz | The Penokean orogeny in the Lake Superior region[END_REF][START_REF] Ricketts | Volcaniclastic rocks and volcaniclastic facies in the middle Precambrian (Aphebian) Belcher Group, Northwest Territories, Canada[END_REF]. Some workers have proposed a back-arc setting (Fralick et al., 2002;[START_REF] Schulz | The Penokean orogeny in the Lake Superior region[END_REF], whereas others have advocated a foreland basin setting [START_REF] Hoffman | Early Proterozoic foredeeps, foredeep magmatism, and Superior-type iron formations of the Canadian Shield[END_REF][START_REF] Ojakangas | Basin evolution of the Paleoproterozoic Karelian Supergroup of the Fennoscandian (Baltic) Shield[END_REF][START_REF] Schneider | Age of volcanic rocks and syndepositional iron formations, Marquette Range Supergroup: Implications for the tectonic setting of Paleoproterozoic iron formations of the Lake Superior region[END_REF].

An intriguing question is whether deposition of these GIF represents local, basin-scale conditions or the global ocean composition and redox state. This is a critical issue because the occurrences and ages of these rocks have been used by some workers to infer deep-water anoxic conditions in the coeval global ocean (e.g., [START_REF] Poulton | The transition to a sulphidic ocean ~1.84 billion years ago[END_REF][START_REF] Slack | Extraterrestrial demise of banded iron formations 1.85 billion years ago[END_REF].

However, if deposition of these GIF reflects restricted, basin-scale conditions [START_REF] Pufahl | Does the Paleoproterozoic Animikie Basin record the sulfidic ocean transition?[END_REF], our understanding of the ocean redox state may be incorrect. Present palaeogeographic reconstructions are insufficient to answer this question; although tidal signatures have been observed in GIF and interbedded sedimentary rocks in Minnesota and the Hudson Bay region [START_REF] Ojakangas | Tidal deposits in the Early Proterozoic basin of the Lake Superior regionthe Palms and the Pokegama Formations: Evidence for subtidal shelf deposition of Superior type banded iron formation[END_REF][START_REF] Chandler | Metallogenesis of an Early Proterozoic foreland sequence, eastern Hudson Bay, Canada[END_REF], which is consistent with at least periodically open-marine conditions, albeit very shallow, during IF deposition.

The ca. 1.88 Ga Frere Iron Formation, part of the Yilgarn Craton of Western Australia [START_REF] Goode | The Nabberu basin of Western Australia[END_REF][START_REF] Rasmussen | Deposition of 1.88-billion-year-old iron formations as a consequence of rapid crustal growth[END_REF], and GIF of the Gibraltar Formation in the Kahochella Group, from the southeastern margin of the Slave Craton that is interbedded with the ~1.9 Ga Akaitcho Formation at its base [START_REF] Bowring | U-Pb zircon ages from Athapuscow aulacogen, east arm of Great Slave Lake, NWT, Canada[END_REF][START_REF] Roscoe | Mineral resource assessment of the area in the East Arm (Great Slave Lake) and Artillery Lake region, N.W.T. proposed as a National Park[END_REF][START_REF] Hoffman | Birthdate for the Coronation paleocean: Age of initial rifting in Wopmay orogen, Canada[END_REF], are age equivalent to the Lake Superior GIF. Notably, a ~20-m-thick magnetite-hematite oolitic GIF is present in the middle member of the Watterson Formation (Hurwitz Group) on the Hearne Craton [START_REF] Miller | Iron-formation, evaporite, and possible metallogenetic implications for the Lower Proterozoic Hurwitz Group, District of Keewatin, Northwest Territories[END_REF]. The age of this oolitic unit is not well constrained, but ages of detrital zircons in the Hurwitz Group suggest that it is younger than 1.9 Ga [START_REF] Davis | Detrital zircon geochronology of the Paleoproterozoic Hurwitz and Kiyuk groups, western Churchill Province, Nunavut[END_REF]. In addition, the presence of IF in the ca. 1.88 Ga Rochford Formation developed on the eastern margin of the Wyoming Craton supports synchronous deposition of IF on several cratons, even though this IF deposit is poorly dated and not granular (Frei et al., 2008).

Deposition of IF on the Superior Craton is coincident with a peak in tonnage for VMS deposits, some of which were positioned in island arcs. VMS deposits of this age are known in the hinterland to the south of the Animikie Basin [START_REF] Schulz | The Penokean orogeny in the Lake Superior region[END_REF], in the Labrador Trough [START_REF] Barrett | Two-stage hydrothermal formation of a Lower Proterozoic sediment-hosted massive sulfide deposit, northern Labrador Trough, Quebec[END_REF], and in the Trans-Hudson orogen [START_REF] Syme | Stratigraphic and tectonic setting of Early Proterozoic volcanogenic massive sulfide deposits, Flin Flon, Manitoba[END_REF]. Submarine volcanic rocks of the Pembine-Wausau terrane of northern Wisconsin were deposited at ca. 1.88 Ga, contemporaneously with GIF of the Animikie Basin [START_REF] Schulz | Synchronous deposition of Paleoproterozoic Superior-type banded iron-formations and volcanogenic massive sulfides in the Lake Superior region: Implications for the tectonic evolution of the Penokeon orogen [abs[END_REF]. These data also suggest that the hydrothermal systems were likely the source of iron, consistent with earlier models (e.g., [START_REF] Isley | Hydrothermal plumes and the delivery of iron to banded iron formation[END_REF]. Iron-oxide exhalites are conspicuously absent at or near these deep-water VMS deposits that presumably formed under open-marine conditions. This observation is unlikely to reflect preservational bias because slightly younger 1.84, 1.79, and 1.78 Ga Cu-rich VMS deposits that similarly formed in arc settings contain abundant hematite and magnetite exhalites, jasper, and IF [START_REF] Slack | Suboxic deep seawater in the late Paleoproterozoic: Evidence from hematitic chert and iron formation related to seafloorhydrothermal sulfide deposits, central Arizona, USA[END_REF][START_REF] Slack | Extraterrestrial demise of banded iron formations 1.85 billion years ago[END_REF][START_REF] Bekker | Iron formation: The sedimentary product of a complex interplay among mantle, tectonic, oceanic, and biospheric processes[END_REF]. This association points to an anoxic and ferruginous composition of deep waters in open-marine settings at ca. 1.88 Ga.

Furthermore, assuming that the iron was hydrothermally derived, there must have been a crash, potentially to Archaean levels, of the marine sulfate reservoir, which would have fostered iron-rich rather than sulfide-rich hydrothermal fluids [START_REF] Kump | Hydrothermal Fe fluxes during the Precambrian: Effect of low oceanic sulfate concentrations and low hydrostatic pressure on the composition of black smokers[END_REF]. Again, this conclusion is based on the premise that the redox state and chemical composition of seawater-derived hydrothermal fluids are strongly influenced by the concentration of the main oxidant in seawatersulfate. Consistent with this interpretation, the ca. 1.88 Ga GIF are commonly interbedded with, or overlain by, black sulfidic shales, which record the development of euxinic conditions in parts of the Animikie Basin (e.g., [START_REF] Poulton | The transition to a sulphidic ocean ~1.84 billion years ago[END_REF][START_REF] Poulton | Spatial variability in oceanic redox structure 1.8 billion years ago[END_REF].

The Animikie Basin contains another stratigraphic level with regionally extensive IFs. This level is stratigraphically above the 1.85 Ga Sudbury impact ejecta layer and is older than the ca.

1.83 Ga regional metamorphic event related to the Penokean orogeny [START_REF] Cannon | The Sudbury impact layer in the Paleoproterozoic iron ranges of northern Michigan, USA[END_REF]. These

IFs are mineralogically and texturally different from the ca. 1.88 Ga GIF and were likely deposited in deeper waters, below fair-weather wave base, and, probably, even below storm wave base. They are developed in the: (1) Marquette Iron Range, Michigan (~60-m-thick Bijiki Iron-Formation Member of the Michigamme Slate containing siderite, chert, iron oxides, and silicates; [START_REF] Ojakangas | Sedimentology and provenance of the Early Proterozoic Michigamme Formation and Goodrich Quartzite, Northern Michigan-Regional stratigraphic implications and suggested correlations[END_REF][START_REF] Ojakangas | Basin evolution of the Paleoproterozoic Karelian Supergroup of the Fennoscandian (Baltic) Shield[END_REF][START_REF] Cannon | The Sudbury impact layer in the Paleoproterozoic iron ranges of northern Michigan, USA[END_REF]; (2) Iron River-Crystal Falls Iron Ranges in Michigan (~15-m-thick chert-siderite slate of the Stambaugh Formation; [START_REF] James | Geology and ore deposits of the Iron River-Crystal Falls district, Iron County, Michigan[END_REF]; and (3) Gogebic Iron Range, Wisconsin (~47-m-thick IF of the Tyler Formation consisting of chert and siderite; [START_REF] Schmidt | The Marquette Range Supergroup in the Gogebic iron district, Michigan and Wisconsin[END_REF][START_REF] Cannon | The Gogebic Iron Range -A sample of the northern margin of the Penokean fold and thrust belt[END_REF]. Deposition of these IFs might be genetically linked to submarine mafic volcanism in the Animikie Basin based on spatial association with, for example, the Badwater Greenstone, but this relationship has not been documented in detail. The presence of these units indicates that the conditions necessary for IF deposition in marine settings were reestablished in the Animikie Basin after 1.85 Ga, in association with mafic volcanism.

Proterozoic age gap in Superior-type IF deposition

It is generally assumed that after ca. 1.85 Ga, Superior-type IFs were not deposited for approximately 1.1 b.y. until precipitation of Rapitan-type IF, associated with the Sturtian glaciation and extensional magmatism linked to the breakup of Rodinia [START_REF] Isley | Plume-related mafic volcanism and the deposition of banded iron formation[END_REF][START_REF] Huston | Barite, BIFs and bugs: Evidence for the evolution of the Earth's early hydrosphere[END_REF][START_REF] Klein | Some Precambrian banded iron-formations (BIFs) from around the world: Their age, geologic setting, mineralogy, metamorphism, geochemistry, and origin[END_REF][START_REF] Slack | Extraterrestrial demise of banded iron formations 1.85 billion years ago[END_REF]. This gap has been explained by a shift to fully oxic [START_REF] Holland | The Chemical Evolution of the Atmosphere and Oceans[END_REF], sulfidic [START_REF] Canfield | A new model for Proterozoic ocean chemistry[END_REF], or suboxic [START_REF] Slack | Suboxic deep seawater in the late Paleoproterozoic: Evidence from hematitic chert and iron formation related to seafloorhydrothermal sulfide deposits, central Arizona, USA[END_REF][START_REF] Heubeck | An early ecosystem of Archean tidal microbial mats (Moodies Group, South Africa, ca. 3.2 Ga)[END_REF] deep-ocean conditions. The earlier suggestion of oxic deep-ocean conditions after ca. 1.88 Ga [START_REF] Holland | The Chemical Evolution of the Atmosphere and Oceans[END_REF] was challenged by the proposal that deep-ocean conditions were predominantly euxinic (anoxic and sulfidic) until full ocean ventilation during the late Neoproterozoic or earliest Phanerozoic (e.g., [START_REF] Canfield | A new model for Proterozoic ocean chemistry[END_REF][START_REF] Poulton | The transition to a sulphidic ocean ~1.84 billion years ago[END_REF]. This interpretation was grounded in a simple modelling approach suggesting that full-ocean ventilation would be difficult, given most current estimates for atmospheric oxygen levels in the Mesoproterozoic.

Although there are a few instances of locally developed mid-Proterozoic euxinia (e.g., [START_REF] Shen | Evidence for low sulphate and anoxia in a mid-Proterozoic marine basin[END_REF][START_REF] Brocks | Biomarker evidence for green and purple sulphur bacteria in a stratified Palaeoproterozoic sea[END_REF][START_REF] Lyons | Tracking euxinia in the ancient ocean: A multiproxy perspective and Proterozoic case study[END_REF], the idea of global-scale deep euxinia has also recently fallen out of favour. All available data that constrain the marine landscape in the mid-Proterozoic ocean are consistent with only locally developed euxinic conditions along productive continental margins (e.g., in oxygen-minimum zones) or within intracratonic basins [START_REF] Scott | Tracing the stepwise oxygenation of the Proterozoic ocean[END_REF][START_REF] Poulton | Spatial variability in oceanic redox structure 1.8 billion years ago[END_REF][START_REF] Planavsky | Widespread iron-rich conditions in the mid-Proterozoic ocean[END_REF][START_REF] Lyons | Geobiology of the Proterozoic eon[END_REF][START_REF] Reinhard | Proterozoic ocean redox and biogeochemical stasis[END_REF]. The emerging consensus is that the redox state of the deep ocean was spatially and temporally variable during the mid-Proterozoic but was generally at a low oxidation state (e.g., [START_REF] Planavsky | Widespread iron-rich conditions in the mid-Proterozoic ocean[END_REF]. [START_REF] Poulton | Spatial variability in oceanic redox structure 1.8 billion years ago[END_REF] argued that euxinic shales deposited on continental margins in oxygenminimum zones and in intracratonic basins could have been a major sink for the hydrothermal iron flux to the mid-Proterozoic oceans and were thus responsible for the absence or scarcity of mid-Proterozoic IFs. At present, this suggestion has not been quantitatively evaluated, but it predicts very high iron contents in euxinic shales during this time interval. Data in [START_REF] Kump | Iron in Precambrian rocks: Implications for the global oxygen budget of the ancient Earth[END_REF] and [START_REF] Bekker | Oxygen overshoot and recovery during the early Paleoproterozoic[END_REF] do show that the average Proterozoic shale has more iron than the average Phanerozoic shales but less than the average Archaean shale.

Large sedimentary iron deposits that formed during this time gap (1.85-0.75 Ga) are indeed absent, but several small IFs and Fe-rich lithologies in sedimentary rock-dominated successions are known (see [START_REF] Bekker | Iron formation: The sedimentary product of a complex interplay among mantle, tectonic, oceanic, and biospheric processes[END_REF]. Examples within this age group include the ca. 1.7 Ga Freedom Formation of the Lower Baraboo Series, Wisconsin, which contains banded ferruginous chert in the lower part interlayered with siderite-and kaolinite-rich slate, collectively ranging from 60 to 160 m thick [START_REF] Weidman | The Baraboo iron-bearing district of Wisconsin[END_REF][START_REF] Leith | The pre-Cambrian of the Lake Superior region, the Baraboo district, and other isolated areas in the upper Mississippi Valley[END_REF][START_REF] Van Wyck | Detrital zircon ages from early Proterozoic quartzites, Wisconsin, support rapid weathering and deposition of mature quartz arenites[END_REF]. The broadly correlative metasedimentary succession of the Tomiko terrane in Ontario contains a magnetite-chert IF [START_REF] Easton | The Grenvillian Tomiko quartzites of Ontario: Correlatives of the Baraboo quartzites of Wisconsin, the Mazatzal orogen of New Mexico, or unique? Implications for the tectonic architecture of Laurentia in the Great Lakes region[END_REF]. The ca. 1.7 Ga Canjica Formation of the Serra de São José Group of the Serpintina Range, and the correlative and economically important Morro Escuro Ridge Sequence of the Espinhaço Group (both in Brazil), were deposited in an intracratonic (rift) basin and contain alternations of hematite and magnetite with quartz [START_REF] Braga | The Statherian itabirite-bearing sequence from the Morro Escuro Ridge, Santa Maria de Itabira, Minas Gerais, Brazil[END_REF][START_REF] Rolim | The Orosirian-Statherian banded iron formation-bearing sequences of the southern border of the Espinhaco Range, southeast Brazil[END_REF]. Additionally, the Chuanlinggou Formation of the North China Craton, a classic GIF deposit, appears to be latest Palaeoproterozoic in age (ca. 1.7 Ga; [START_REF] Wan | SHRIMP ages of detrital zircons from the Changcheng System in the Ming Tombs area, Beijing: Constraints on the protolith nature and maximum depositional age of the Mesoproterozoic cover of the North China Craton[END_REF][START_REF] Dai | Fossil bacteria in Xuanlong iron ore deposits of Hebei Province[END_REF].

Even though the IFs deposited during this time interval are spatially limited, relative to giant Neoarchaean and Palaeoproterozoic deposits, their existence is significant. For example, these deposits may provide evidence for episodic upwelling of hydrothermal iron from deep-water oceanic settings from ca. 1.85 to 0.8 Ga. That said, these units are still poorly studied, and we cannot rule out alternative depositional models (e.g., Fe-rich groundwater, local shallow-water hydrothermal iron sources, and iron recycled from sediments deposited in inner-shelf mud belts).

Between 1.85 and 0.8 Ga, open-marine, deep-water environments may, in some cases, have been sufficiently oxygenated to oxidise and precipitate Fe(II) from hydrothermal plumes as Fe(III)

oxyhydroxides (e.g., [START_REF] Slack | Suboxic deep seawater in the late Paleoproterozoic: Evidence from hematitic chert and iron formation related to seafloorhydrothermal sulfide deposits, central Arizona, USA[END_REF]Slack et al., , 2009)), whereas evidence for Mn(II) oxidation in the deep ocean is generally lacking. Although this record of the Fe cycle does not provide strong quantitative constraints on the deep-ocean redox state, it questions models invoking either fully oxic or fully sulfidic conditions throughout the deep ocean in the mid-Proterozoic. Again, in terms of time and space, a low but variable redox state in the deep ocean is most consistent with the record of Mesoproterozoic iron deposits.

Iron Formations, Primary Productivity, and Atmospheric Oxygenation

Although there is still debate about the dynamics of initial atmospheric oxygenation, herein we present a coherent model for the rise of atmospheric O 2 . Underpinning this transition was the evolution of cyanobacteria and the oxygen they generated. Importantly, it needs to be reiterated that oxygenation of the shallow oceans likely preceded that of the atmosphere. As discussed above, several lines of evidence tentatively point to their evolution by at least 3.0 Ga, and perhaps even earlier.

If cyanobacteria evolved by 3.0 Ga or before, the obvious question is why did it take several hundred million years for oxygen to accumulate in the atmosphere? To maintain low atmospheric oxygen levels despite O 2 production in the oceans requires that oxygen sinks must have been much larger than they are now (including reduced solutes in seawater, volcanic gases, or crust), or the oxygen flux was greatly reduced (due to limited cyanobacterial productivity), or both. It is vital to remember that most organic matter produced in the upper water column will be remineralised before reaching the seafloor, even if there is a surge in productivity associated with the evolution of cyanobacteria. And, this relationship will likely hold even if marine sulfate levels were very low (e.g., 100 µM). Therefore, oxygen flux calculations on local or global scales need to consider the very large offset between export and primary production. That being said, there should be an oxygen flux to the atmosphere in the Archaean in the presence of cyanobacterial production; exactly how oxygen levels were buffered at very low levels remains an unresolved question. In this regard, there are several possibilities: (1) cyanobacteria were nutrient starved and hence failed to generate appreciable O 2 ;

(2) cyanobacteria were marginalised by factors other than nutrients;

(3) an overabundance of reductants consumed any O 2 produced; and (4) there was limited emergence of continental landmass, which may have limited land-based cyanobacterial activity, as well as affected consumption of oxygen during weathering.

(1) The extent to which phosphorous may have limited cyanobacterial productivity in the Archaean and Proterozoic oceans remains an area of active debate. Based on the low P/Fe ratios in Archaean and Palaeoproterozoic IFs, it was proposed that seawater at that time contained limited marine phosphorous. This, in turn, would have reduced levels of photosynthesis and carbon burial, thereby inhibiting long-term oxygen production on the early Earth [START_REF] Bjerrum | Ocean productivity before about 1.9 Ga ago limited by phosphorus adsorption onto iron oxides[END_REF]. This model was based on partitioning coefficients (K D ) derived for P to Fe(III) oxyhydroxides associated with modern hydrothermal plumes. However, because the K D value for P sorption to Fe(III) oxyhydroxides diminishes in the presence of dissolved silica, later workers suggested that IF P/Fe ratios suggest that Archaean phosphate concentrations may not have been limiting (Konhauser et al., 2007a;Planavsky et al., 2010b).

From a modelling perspective, [START_REF] Laakso | Regulation of atmospheric oxygen during the Proterozoic[END_REF] and [START_REF] Derry | Causes and consequences of mid-Proterozoic anoxia[END_REF] argued for low P levels in the Proterozoic through the establishment of an Fe trap -either through scavenging by Fe(III) particles or precipitation of ferrous phosphates, such as vivianite, Fe 3 (PO 4 ) 2 . Similarly, [START_REF] Reinhard | Evolution of the global phosphorus cycle[END_REF] presented a large compilation of P data from siliciclastic sediments and identified a four-fold increase in P burial at the start of the Cryogenian (~720 Ma), which they linked to progressive oxygenation of the Earth and loss of the deep-sea Fe trap. [START_REF] Poulton | Biogeochemistry: Early phosphorus redigested[END_REF] countered that the record presented by [START_REF] Reinhard | Evolution of the global phosphorus cycle[END_REF], and the large observed increase in P burial in the Neoproterozoic, may alternatively be explained by high P liberation and recycling from the sediment back into the water column. Accordingly, the extent to which P may have limited cyanobacterial productivity in the Proterozoic ocean remains unresolved.

Although P is likely to have been the ultimate limiting nutrient (e.g., [START_REF] Reinhard | Evolution of the global phosphorus cycle[END_REF], limitation need not have been restricted to this element. With denitrification in the ocean interiors, N fixation in surface waters is needed to balance any P input to the oceans. Despite N isotopes having been used to argue for the antiquity of Mo-nitrogenase to a least 3.2 Ga (Stüeken et al., 2015a), prior to the Neoproterozoic rise in oxygen, the bulk marine Mo reservoir is likely to have been depressed owing to limited oxidative weathering, which in turn could have limited nitrogen fixation and overall primary productivity by oxygenic phototrophs in the Proterozoic [START_REF] Anbar | Proterozoic ocean chemistry and evolution: A bioinorganic bridge?[END_REF][START_REF] Scott | Tracing the stepwise oxygenation of the Proterozoic ocean[END_REF][START_REF] Reinhard | Proterozoic ocean redox and biogeochemical stasis[END_REF]. What role other biologically-essential nutrients played in ancient primary productivity is a current avenue of research (e.g., [START_REF] Robbins | Authigenic iron oxide proxies for marine zinc over geological time and implications for eukaryotic metallome evolution[END_REF][START_REF] Scott | Bioavailability of zinc in marine systems through time[END_REF][START_REF] Swanner | Cobalt and marine redox evolution[END_REF]. Overall, the possible suppression of cyanobacterial productivity by nutrient limitation in the Proterozoic is consistent with the current view of Earth's oxygenation history (e.g., [START_REF] Lyons | The rise of oxygen in Earth's early ocean and atmosphere[END_REF] and with either limited or transient oxygen presence in the surface ocean (e.g., [START_REF] Hardisty | An iodine record of Paleoproterozoic surface ocean oxygenation[END_REF][START_REF] Hardisty | Perspectives on Proterozoic surface ocean redox from iodine contents in ancient and recent carbonate[END_REF].

(2) Instead of simply being starved, is it possible that other factors caused the marginalisation of cyanobacteria? One avenue worth considering is the analysis of nickel concentrations (expressed as molar Ni/Fe ratios) in IFs through time. The pattern shows a dramatic drop in Ni availability in the oceans around 2.7-2.6 Ga [START_REF] Pecoits | Petrography and geochemistry of the Dales Gorge banded iron formation: Paragenetic sequence, source and implications for palaeo-ocean chemistry[END_REF][START_REF] Konhauser | Oceanic nickel depletion and a methanogen famine before the Great Oxidation Event[END_REF][START_REF] Konhauser | The Archean nickel famine revisited[END_REF], although the timing of this dramatic decrease has been suggested to have occurred closer to the GOE, once the effects of cellular organic matter on Ni adsorption to ferrihydrite particles are considered [START_REF] Eickhoff | Nickel partitioning in biogenic and abiogenic ferrihydrite: The influence of silica and implications for ancient environments[END_REF]. The decrease in seawater Ni content after 2.7 Ga is consistent with a progressively cooling Archaean mantle, whereby the volume of ultramafic melts produced by partial mantle melting and the eruption of Ni-bearing ultramafic volcanic rocks (known as komatiites) decreased, and, subsequently, less Ni was dissolved into seawater. Since first identified, the decrease in Ni observed by [START_REF] Konhauser | Oceanic nickel depletion and a methanogen famine before the Great Oxidation Event[END_REF] has been suggested to be an artefact of limited IF sampling; however, after nearly doubling the number of IF analyses, the unidirectional decrease in Ni has remained stark and robust [START_REF] Konhauser | The Archean nickel famine revisited[END_REF]. This decrease in Ni availability could have had profound consequences for microorganisms that depended on it, specifically methane-producing bacteria or methanogens [START_REF] Konhauser | Oceanic nickel depletion and a methanogen famine before the Great Oxidation Event[END_REF]. These bacteria have a unique Ni requirement for their methane-producing enzymes, therefore, a limitation in this metal could have decreased their population in seawater. Compounding their plight, an increased supply of sulfate to the oceans after ca. 2.5 Ga (e.g., [START_REF] Cameron | Evidence from early Proterozoic anhydrite for sulfur isotopic partitioning in Precambrian oceans[END_REF] would have fostered competition between the starving methanogens and the increasingly abundant sulfate-reducing bacteria, the latter having a competitive edge in terms of substrate utilisation [START_REF] Lovley | Kinetic analysis of competition between sulfate reducers and methanogens for hydrogen in sediments[END_REF]. It is thus possible that a Ni famine eventually led to a cascade of events that began with reduced methane production, the expansion of cyanobacteria into deep-water settings previously occupied by methanogens, and ultimately increased oxygenic photosynthesis that tipped the atmospheric redox balance in favour of oxygen, leading to the GOE. At the heart of this model is the simple idea that a decreased flux of methane led to greater oxygen concentrations in the atmosphere even if the marine O 2 flux to the atmosphere remained constant [START_REF] Zahnle | The loss of mass-independent fractionation in sulfur due to a Paleoproterozoic collapse of atmospheric methane[END_REF].

Is there any other evidence in the rock record to support a change in biological dynamics between methanogens and cyanobacteria? The answer may come from Neoarchaean shallow and deep-water sediments from the Hamersley Province that record 150 m.y. of Archaean history. As discussed above, [START_REF] Eigenbrode | Late Archean rise of aerobic microbial ecosystems[END_REF] observed highly negative δ 13 C values in the organic carbon fraction, a feature of the Neoarchaean that has been known for some time, most likely representing microbial habitats influenced by the assimilation of methane, i.e., methanotrophy (e.g., [START_REF] Hayes | Geochemical evidence bearing on the origin of aerobiosis, a speculative hypothesis[END_REF]. Intriguingly, there is a 10‰ increase in δ 13 C values in kerogen in the post-2.7

Ga shallow-water facies relative to the deep-water settings, suggesting that the shallow waters became oxygenated (or oxidant-rich) as cyanobacteria became more abundant. At the same time, the deep waters remained a viable environment for the methanogens and methanotrophs. As populations of sulfate-reducing bacteria in the oceans increased, owing to the rise of atmospheric oxygen and concomitantly higher seawater sulfate concentrations, such bacteria would have continued to marginalise the methanogens and, indirectly, restrict the latter to even greater depths in the oceans, limiting their distribution to bottom sediments. Restricting methane production largely to the sediment pile would greatly increase the chances of anaerobic methane oxidation by either sulfate or solid-phase ferric iron, further reducing the methane flux to the atmosphere (Konhauser et al., 2005;[START_REF] Catling | Anaerobic methanotrophy and the rise of atmospheric oxygen[END_REF].

It is also possible that prior to the GOE cyanobacterial activity was depressed by high Fe(II) concentrations in surface waters. [START_REF] Swanner | Modulation of oxygen production in Archaean oceans by episodes of Fe (II) toxicity[END_REF] recently discussed the implications of Fe toxicity to cyanobacteria in the range of 10s to 100s of µM range of Fe(II). They suggested that elevated Fe(II) concentrations could have provided an external environmental stress on cyanobacterial populations, leading to increased production of intracellular reactive oxygen species, and attenuated growth rates and oxygen production. The potential for Fe(II) toxicity represents a previously unrecognised environmental limitation of planktonic phototrophs, and could conceivably have been important given estimates for marine Fe(II) concentrations from the low µM to mM range (e.g., [START_REF] Holland | The Chemical Evolution of the Atmosphere and Oceans[END_REF][START_REF] Czaja | Evidence for free oxygen in the Neoarchean ocean based on coupled ironmolybdenum isotope fractionation[END_REF][START_REF] Tosca | An authigenic origin for Precambrian greenalite: Implications for iron formation and the chemistry of ancient seawater[END_REF]. As explored above, photoferrotrophs are likely to drive quantitative Fe(II) oxidation in an upwelling water mass in some, but not all, marine environments.

(3) It was hypothesised decades ago that the abundance of reduced solutes in the oceans could have served as a buffer to the increasing amounts of O 2 produced (e.g., [START_REF] Cloud | Paleoecological significance of banded iron-formation[END_REF]. For instance, the oceans were rich in dissolved Fe(II), as testified by the global deposition of IFs between 2.60-2.45

Ga. Yet, as pointed out by [START_REF] Towe | Earth's earliest atmosphere: Constraints and opportunities for early evolution[END_REF][START_REF] Towe | Environmental oxygen conditions during the origin and early evolution of life[END_REF], it is unlikely that ferrous iron buffering alone could balance a moderate oxygen flux because for every 1 mole of O 2 that reacts with ferrous iron in solution, 4 moles of Fe 2+ are consumed [see reactions 7-8].

[

7] H 2 O + CO 2 à CH 2 O + O 2 [8] 4Fe 2+ + O 2 + 10H 2 O à 4Fe(OH) 3 + 8H +
Other workers have proposed that the additional buffering capacity was provided by reduced volcanic and metamorphic gases (e.g., [START_REF] Kump | Rise of atmospheric oxygen and the ''upside-down[END_REF]. However, for these sinks to be larger in the Archaean than the Proterozoic, volcanic and metamorphic gases must have been more reducing unless the O 2 production flux was smaller. [START_REF] Kump | Increased subaerial volcanism and the rise of atmospheric oxygen 2.5 billion years ago[END_REF] proposed that in the Archaean volcanism was predominantly restricted to submarine settings and that volatiles released during submarine volcanism were more reducing than in subaerial settings (as they are today) because the former contain more CO, CH 4 , and H 2 S. Following the Archaean-Proterozoic transition, subaerial volcanism appears to have become more pervasive due to stabilisation of the continents. A shift from predominantly submarine to subaerial volcanism, releasing more oxidising volatiles such as H 2 O, CO 2 , and SO 2 , would have reduced the overall sink for oxygen and led to the rise of atmospheric oxygen. Along similar lines, [START_REF] Gaillard | Atmospheric oxygenation caused by a change in volcanic degassing pressure[END_REF] proposed that a decrease in the average pressure of volcanic degassing changed the oxidation state of sulfur in volcanic gases from predominantly H 2 S to SO 2 . Interestingly, the increased degassing of SO 2 over H 2 S into the atmosphere was used by those same authors to not only explain the proliferation of S-MIF recorded in marine sediments between 2.7-2.5 Ga, but also that the production of SO 2 in the gas is accompanied by the formation and release of H 2 [reaction 9], which then escaped the atmosphere and aided in oxygenation of the planet [START_REF] Catling | Biogenic methane, hydrogen escape, and the irreversible oxidation of early Earth[END_REF].

[9] H 2 S + 2H 2 O à SO 2 + 3H 2 (4) Another factor might have been the limited emergence of continental land masses. Signs of oxidative weathering likely reflect the ancient presence of cyanobacteria, which are rarely considered with respect to having formed terrestrial mats versus existence as marine phytoplankton.

In this regard, [START_REF] Lalonde | Benthic perspective on Earth's oldest evidence for oxygenic photosynthesis[END_REF] hypothesised that localised O 2 production within the confines of ancient biological communities on land could have generated sufficient oxidising power to mobilise trace metals and SO 4 2-from sulfide-bearing minerals in underlying substrata, without affecting the redox balance of the ocean-atmosphere system prior to the GOE. Under such a scenario, oxygen produced by the benthic microbial community would be consumed proximal to the mat, much as is observed in some modern mat systems (e.g., [START_REF] Gingras | Possible evolution of mobile animals in association with microbial mats[END_REF][START_REF] Sumner | Antarctic microbial mats: A modern analog for Archean lacustrine oxygen oases[END_REF].

New evidence that the early Archaean continents were greater in size and that a larger proportion was emergent than previously thought, highlights the potential importance of consumption of O 2 through oxidative weathering of the continents and suppression of atmospheric O 2 until the GOE [START_REF] Satkoski | A high continental weathering flux into Paleoarchean seawater revealed by strontium isotope analysis of 3.26 Ga barite[END_REF][START_REF] Satkoski | Initiation of modern-style plate tectonics recorded in Mesoarchean marine chemical sediments[END_REF] [START_REF] Satkoski | A high continental weathering flux into Paleoarchean seawater revealed by strontium isotope analysis of 3.26 Ga barite[END_REF][START_REF] Satkoski | Initiation of modern-style plate tectonics recorded in Mesoarchean marine chemical sediments[END_REF].

Irrespective of the causes for the delay in the rise of atmospheric O 2 , by 2.5 Ga we see evidence for oxygen accumulation in the atmosphere from the Mt. McRae Shale in Western Australia, recording transient oxidised conditions now widely known as the 'whiff' of oxygen [START_REF] Anbar | A whiff of oxygen before the Great Oxidation Event?[END_REF]. The whiff is recorded by spikes in Mo and Re concentrations, elevated nitrogen isotope values, and iron speciation data [START_REF] Anbar | A whiff of oxygen before the Great Oxidation Event?[END_REF][START_REF] Reinhard | A Late Archean sulfidic sea stimulated by early oxidative weathering of the continents[END_REF][START_REF] Garvin | Isotopic evidence for an aerobic nitrogen cycle in the latest Archean[END_REF]; based on Re-Os isotope analyses, these data can be confidently considered primary signals [START_REF] Kendall | Transient episodes of mild environmental oxygenation and oxidative continental weathering during the late Archean[END_REF]. The shales lie near the middle of the 2.5-km-thick Hamersley Group (dated between 2.63 and 2.45 Ga), and directly underlie (in ascending order) the Brockman Iron Formation, Weeli Wolli Formation, Woongarra Volcanics, and Boolgeeda Iron Formation. Scott et al. (2011) have subsequently shown that euxinic conditions could develop under an Archaean anoxic atmosphere, and alternatives to the 'whiff' model have been proposed.

A recent compilation of the Cr contents in IFs has also contributed to our understanding of the GOE. For instance, a moderate enrichment in Cr (as expressed by the molar Cr/Ti ratio) beginning at 2.45 Ga in the Weeli Wolli Formation, as well as the similarly aged Cauê Iron Formation in Brazil, was followed by a spike in Cr enrichment in oolitic and pisolitic ironstone associated with the ca. 2.32 Ga Timeball Hill Formation [START_REF] Konhauser | Aerobic bacterial pyrite oxidation and acid rock drainage during the Great Oxidation Event[END_REF]. Cr enrichment in the face of muted Cr isotope fractionation at this time (Planavsky et al., 2014b) points to a supply mechanism that involved predominately the reduced Cr(III) form [START_REF] Konhauser | Aerobic bacterial pyrite oxidation and acid rock drainage during the Great Oxidation Event[END_REF] (Garrels and Mackenzie, 1971) -until the volume of easily oxidisable pyrite in the weathered crust was diminished [START_REF] Konhauser | Aerobic bacterial pyrite oxidation and acid rock drainage during the Great Oxidation Event[END_REF][START_REF] Bekker | Oxygen overshoot and recovery during the early Paleoproterozoic[END_REF].

Interestingly, not all IFs in this time window display elevated Cr enrichments (e.g., Dales Gorge and Joffre members of the Brockman Iron Formation). Considering the low solubility of Cr(III) at marine pH levels, the rapid reduction of Cr(VI) by aqueous Fe(II), and near-instantaneous coprecipitation of Cr(III) with Fe(III)-oxyhydroxide (e.g., [START_REF] Fendorf | Surface reactions of chromium in soils and waters[END_REF], dissolved Cr dispersal would be limited upon delivery to the oceans. In this regard, proximity of the depositional site to shore played a strong role in determining which IF would record a continental Cr input. Indeed, IFs having some of the highest Cr values, such as the Cauê and Timeball Hill formations, show evidence of sediment re-working and grade into GIF, indicative of shallow-water deposition [START_REF] Eriksson | The Timeball Hill Formation: A fossil delta[END_REF][START_REF] Spier | Geochemistry and genesis of the banded iron formations of the Cauê Formation, Quadrilátero Ferrífero, Minas Gerais, Brazil[END_REF]. The oolitic and pisolitic Timeball Hill Formation records the highest degree of Cr enrichment among Palaeoproterozoic IF [START_REF] Konhauser | Aerobic bacterial pyrite oxidation and acid rock drainage during the Great Oxidation Event[END_REF], which is not surprising considering its shallow-water depositional environment (pro-deltaic to offshore).

Evidence for increased oxidative weathering of the continents is also supported by a recent The ability of cyanobacteria in the ocean's photic zones to remotely facilitate their own nutrient supply via O 2 -enhanced chemical weathering may represent an unrecognised positive biological feedback [START_REF] Konhauser | Aerobic bacterial pyrite oxidation and acid rock drainage during the Great Oxidation Event[END_REF]. Thus the ca. 2.22-2.10 Ga Lomagundi carbon isotope excursion may, together with the Cr enrichment and Cu isotope changes in early Palaeoproterozoic sediments, be a manifestation of acid-driven dissolution and the first sedimentary cycle of oxidative weathering [START_REF] Holland | Volcanic gases, black smokers, and the Great Oxidation Event[END_REF][START_REF] Bekker | Oxygen overshoot and recovery during the early Paleoproterozoic[END_REF].

Although multiple palaeoredox indicators point to a dramatic increase in surface oxygenation during the GOE, and in particular during its later part in association with the Lomagundi carbon isotope excursion (e.g., [START_REF] Bekker | Oxygen overshoot and recovery during the early Paleoproterozoic[END_REF], we know surprisingly little regarding how high atmospheric and ocean oxygen contents rose. For instance, paleosol data from that time suggests that atmospheric oxygen might have risen to levels of 0.03 atm, and perhaps even as high as 0.1 atm [START_REF] Bekker | Oxygen overshoot and recovery during the early Paleoproterozoic[END_REF] (Fig. 10). The emerging picture from black shales indicates that Precambrian seawater reservoirs of redox-sensitive elements were nowhere close to those of the Phanerozoic (e.g., [START_REF] Scott | Tracing the stepwise oxygenation of the Proterozoic ocean[END_REF]Partin et al., 2013b;[START_REF] Reinhard | Proterozoic ocean redox and biogeochemical stasis[END_REF]. Furthermore, Se isotope records of black shales suggest that while shallow-marine settings were suboxic and maintained enough oxygen for complex life to evolve, deep oceans remained at a low oxygenation state during the Lomagundi carbon isotope excursion, with euxinic conditions developed in intracratonic basins and on continental margins where nutrients were upwelled and high biological productivity was established [START_REF] Kipp | Selenium isotopes record extensive marine suboxia during the Great Oxidation Event[END_REF]. It is unclear why IFs, especially GIFs that are consistent with shallow-water depositional environments, are not more abundant during >100 m.y.

of the Lomagundi carbon isotope excursion.

Following the end of the Lomagundi carbon isotope excursion, multiple redox indicators point to the collapse of seawater sulfate and atmospheric/oceanic oxygenation (e.g., Planavsky et al., 2012b;[START_REF] Scott | Tracing the stepwise oxygenation of the Proterozoic ocean[END_REF][START_REF] Scott | Pyrite multiple-sulfur isotope evidence for rapid expansion and contraction of the early Paleoproterozoic seawater sulfate reservoir[END_REF][START_REF] Bekker | Oxygen overshoot and recovery during the early Paleoproterozoic[END_REF]Partin et al., 2013a, b). In the case of oxygen, estimates range from values as high as 40% PAL to account for persistent anoxia in the oceans [START_REF] Canfield | The early history of atmospheric oxygen: Homage to Robert M. Garrels[END_REF] to values as low as 0.1% PAL based on Cr isotopes in various Proterozoic strata (Planavsky et al., 2014b) (Fig. 10). Consistent with these observations, IFs resurface, reaching their peak at ca. 1.88 Ga. At this time, shallow-marine GIFs were deposited above wave-base on several cratons (e.g., [START_REF] Rasmussen | Deposition of 1.88-billion-year-old iron formations as a consequence of rapid crustal growth[END_REF][START_REF] Bekker | Iron formations: Their origins and implications for ancient seawater chemistry[END_REF] indicating a large flux of Fe(II)

and Mn(II) from submarine-hydrothermal vents in association with an extensive LIP event. Deep oceans must have returned for a protracted period to the Archaean ferruginous (anoxic) state for the last time in Earth history, yet sedimentologically these ca. 1.88 Ga GIFs do not have equivalents in the Archaean. Although details remain unclear, it is certain that interplay of the strength of the hydrothermal flux linked to magmatic activity, ocean redox state, and palaeogeography (e.g., configuration of basins) led to this unprecedented rate of IF deposition for post-GOE time [START_REF] Bekker | Iron formation: The sedimentary product of a complex interplay among mantle, tectonic, oceanic, and biospheric processes[END_REF](Bekker et al., , 2013)). The extent and volume of these GIFs point to a rather peculiar state of the Earth System characterised by fully Fe-charged deep oceans like in the Archaean, and oxygenated nearshore conditions like in the Phanerozoic.

It is generally assumed that atmospheric and ocean oxygen levels remained low until ca. 1.2

Ga [START_REF] Kah | Low marine sulphate and protracted oxygenation of the Proterozoic biosphere[END_REF] or even ca. 800 Ma [START_REF] Thomson | Chemostratigraphy of the Shaler Supergroup, Victoria Island, NW Canada: A record of ocean composition prior to the Cryogenian glaciations[END_REF][START_REF] Turner | Thick sulfate evaporate accumulations marking a mid-Neoproterozoic oxygenation event (Ten Stone Formation, Northwest Territories, Canada)[END_REF]Planavsky et al., 2014b). Similarly, the deep oceans are universally thought to have remained in a low-oxygen state, although debate continues on the exact level of deep-ocean oxygen concentrations with correspondingly anoxic (=ferruginous) and suboxic designations debated [START_REF] Planavsky | Widespread iron-rich conditions in the mid-Proterozoic ocean[END_REF][START_REF] Slack | Suboxic deep seawater in the late Paleoproterozoic: Evidence from hematitic chert and iron formation related to seafloorhydrothermal sulfide deposits, central Arizona, USA[END_REF]Slack et al., , 2009)). Puzzlingly, economic-grade IFs are rather rare for this billion-year-long interval. However, a growing worldwide database for IFs and much more precise ages point to a dynamic redox state of the deep-oceans with IFs speaking to the water column's episodic anoxic (=ferruginous) state; for the rest of the time low, but substantial deepocean oxygen levels restricted Fe(II) transport in solution from hydrothermal vents. Similarly, Mn deposits are uncommon during this time period. This low-oxygen redox state in the deep oceans formed a prelude to Snowball glaciations, the rise of atmospheric and oceanic oxygen, and the emergence of complex life; the only return of massive IFs was in association with the Sturtian glaciation (e.g., [START_REF] Klein | Sedimentology and geochemistry of the glaciogenic Late Proterozoic Rapitan Iron Formation in Canada[END_REF]Hofmann et al., 1998). The Phanerozoic is a logical continuation of this trend during which higher atmospheric oxygen and a more intense biological pump restricted ironstones to relatively short time periods when the emplacement of LIPs resulted in greenhouse conditions, sea level rise, and extensive anoxic conditions in ocean basins.

Concluding remarks

Iron formations are a defining part of the Neoarchaean to Palaeoproterozoic marine sedimentary record. These deposits have been studied extensively, given their importance as ore bodies, for over a century. However, in the past decade alone there have been major advances in our understanding of their genesis. Some of these advances have come from revisiting the role that microbial metabolic reactions likely played in their deposition and diagenesis, in phase with equally rapid progress in our general understanding of microbial Fe cycling. Also, as expected, other advances have come from the application of novel isotope systems (e.g., Fe, Cr, U, Mo, Cu isotopes), as well as expansion of in situ isotopic analysis, which allows correlations to mineral paragenesis. With these new findings, there have been shifts from the 'textbook view' of IF genesis established during the pioneering work done in the 1970s and 1980s. Foremost, in contrast to the entrenched view, it is now generally accepted that microbial (i.e., enzymatic) Fe(II) oxidation was likely a key process in the deposition of many IFs. In addition, evidence for extensive microbial Fe(III) reduction (e.g., DIR), on a basin-wide scale, as well as locally in the sediment, seems compelling. It is also clear that the deposition of IF was unlikely to have represented background marine sedimentation.

Instead, a strong hydrothermal flux of iron, coupled with a low epiclastic input, was often the key factor that yielded IF rather than more typical marine sediments. Similarly, a complex set of factors -including the evolution of submarine-hydrothermal systems -is behind the secular distribution, rather than simply the deep marine redox evolution (i.e., ocean oxygenation) as was often imagined by early workers.

Our understanding of IF genesis has certainly moved forward at a rapid pace, as our use of these distinctive chemical sediments as palaeo-oceanographic archives has surged in recent years.

One of the most exciting advances is an increased understanding of the co-evolution of Earth surface processes with microbial metabolism during the Archaean-Proterozoic transition. For example, the IF record has helped shape our view on the evolution of aerobic metabolisms tied to earlier expansion of cyanobacteria throughout the surface oceans during the Archaean. The emergence of an aerobic biosphere was likely brought on by an oxidative component to nutrient cycling and chemical weathering reactions, potentially altering nutrient fluxes to the oceans.

Identification of widespread DIR, a deeply rooted metabolism in both bacteria and archaea, has also allowed us to understand the diversity of microbial ecology in the Archaean. Evidence for these processes continues to be found in trace element and isotopic signatures contained within these Ferich chemical deposits as they continue to be the subject of intense study. Jacobsen and Pimentel-Klose (1998a), [START_REF] Alibert | Rare earth elements in Hamersley BIF minerals[END_REF][START_REF] Alibert | Rare element and neodymium isotopic compositions of the banded iron-formations and associated shales from Hamersley, Western Australia[END_REF]Li et al. (2015).

Acknowledgments

Figure 7 -Histograms of Si isotope compositions for Precambrian cherts in IFs, as well as Fe-free cherts. Also shown are Si isotope fractionation factors for Si sorption onto Fe(III)-oxyhydroxides [START_REF] Delstanche | Silicon isotopic fractionation during adsorption of aqueous monosilicic acid onto iron oxide[END_REF], Fe(III)-Si gels [START_REF] Zheng | Abiologic silicon isotope fractionation between aqueous Si and Fe(III)-Si gel in simulated Archean seawater: Implications for Si isotope records in Precambrian sedimentary rocks[END_REF], and mixed Fe(II)-Fe(III)-Si gels produced by DIR [START_REF] Reddy | Silicon isotope fractionation during microbial reduction of Fe(III)-Si gels under Archean seawater conditions and implications for iron formation genesis[END_REF], referenced to the average δ 30 Si value of Fe-free chert. Lower Member. Earliest magnetite is recorded in magnetite that is fine-grained and has low Si contents, whereas high-temperature magnetite is coarse-grained and/or contains high levels of Si [START_REF] Huberty | Silician magnetite from the Dales Gorge Member of the Brockman Iron Formation, Hamersley Group, Western Australia[END_REF]. Oxygen isotope temperatures were obtained using hematite-quartz and magnetite-quartz fractionations; those for high-temperature magnetite use the average measured δ 18 O value for quartz (δ 18 O ~ +21‰), reflecting the last equilibration of quartz during metamorphism, whereas lowtemperature magnetite and hematite assumes a δ 18 O value for quartz of +23‰ (see discussion in W Li et al., 2013a). Oxygen isotope fractionation factors are from [START_REF] Kita | Oxygen isotope fractionation between amorphous silica and water at 34-93 o C[END_REF], [START_REF] Mandernack | Oxygen and iron isotope studies of magnetite produced by magnetotactic bacteria[END_REF], and [START_REF] Bao | Oxygen isotope fractionation in ferric oxide-water systems: Low temperature synthesis[END_REF]. Atmospheric O 2 curve is relative to Present Atmosphere Level (PAL), adapted from [START_REF] Lyons | The rise of oxygen in Earth's early ocean and atmosphere[END_REF]. GOE marks the "Great Oxidation Event" that may record an "over-shoot" in the first major rise in atmospheric O 2 , followed by low periods in most of the Proterozoic. The green curves represent the range of commonly proposed pO 2 estimates for early Paleoproterozoic. There may also have been transient increases in atmospheric O 2 before the GOE (so-called whiffs), although such increases could also represent oxidative weathering associated with benthic mats [START_REF] Lalonde | Benthic perspective on Earth's oldest evidence for oxygenic photosynthesis[END_REF]. Estimates for maximum O 2 contents in shallow oceans (photic zones) in blue boxes relative to Present Ocean Level (POL) are based on modelling of stable Fe isotopes [START_REF] Czaja | Evidence for free oxygen in the Neoarchean ocean based on coupled ironmolybdenum isotope fractionation[END_REF]W. Li et al., 2013b;[START_REF] Satkoski | A redox-stratified ocean 3.2 billion years ago[END_REF], and potentially provide evidence that the shallow oceans had significant free O 2 levels by 3.2 Ga, hundreds of millions of years prior to the rise in atmospheric O 2 . supply rates are summed, with eddy-diffusion held constant at 0.1 cm 2 sec -1 in all cases (cf., [START_REF] Kappler | Deposition of banded iron formations by anoxygenic phototrophic Fe(II)-oxidizing bacteria[END_REF]. Modelled advection rates are consistent with low and high annual averages for upwelling regimes on continental margins. Eddy-diffusivity value of 0.1 cm 2 sec -1 is also shown on each plot. Dashed lines indicate the photoferrotrophic layer thickness required for Fe(II) oxidation at rates of 1.4 x 10 -4 , 1.4 x 10 -5 , and 1.4 x 10 -6 M day -1 , with 1.4 x 10 -5 M day -1 representing a conservative estimate as determined experimentally by [START_REF] Kappler | Deposition of banded iron formations by anoxygenic phototrophic Fe(II)-oxidizing bacteria[END_REF]. Thickness of photoferrotrophic layer increases dramatically with low Fe(II) oxidation rates when both advection and eddy-diffusional Fe transport are considered, as opposed to solely eddy-diffusional transport. At conservatively estimated Fe(II) oxidation rate for upwelling waters with Fe(II) concentration of 500 µM, photoferrotrophic layer thicknesses of 17.6 and 5.6 m are required when solely considering diffusion at rates of 1 cm 2 sec -1 (not shown) and 0.1 cm 2 s -1 , respectively. These thicknesses are distinctly different than >100 m required when both advection at 5 m day -1 and diffusion at 0.1 cm 2 sec -1 are considered for the same Fe concentrations in upwelling waters. More realistic layer thickness of ~35 m is required for Fe(II) oxidation rate of 1.4 x 10 -5 M day -1 and advection rate and eddy-diffusivity of 5 m day -1 and 0.1 cm 2 s -1 , respectively, at 100 µM Fe in upwelling waters.

Figure 12 -Comparison of (A) typical sediment pore-water profile during early diagenesis in a modern continental marine setting (e.g., [START_REF] Froelich | Early oxidation of organic matter in pelagic sediments of the eastern equatorial Atlantic: Suboxic diagenesis[END_REF]modified from Konhauser, 2007), and (B) plausible Archaean pore-water profile during early diagenesis wherein minerals on right-hand side are dominant controls on pore-water Fe speciation. As shown in Figure 9, delivery of Fe(OH) 3 to sediment is controlled by O 2 concentrations and/or metabolism of photoferrotrophic bacteria.

Figure 13 -Comparison of wt.% organic carbon measured in ca. 2.5 Ga Transvaal IFs (Beukes et al., 1990;Beukes and Klein, 1990;[START_REF] Kaufman | Geochemical and mineralogic effects of contact metamorphism on banded iron-formation: An example from the Transvaal Basin, South Africa[END_REF] with those estimated for the same samples based on organic carbon contents required to account for δ 13 C values of siderite and Fe(II) concentrations in magnetite. Using this approach, average original organic carbon content was 4.70 wt.% for oxide-facies IF, which is comparable to those measured in Neoarchaean black shales, indicating potentially large amounts of initial organic carbon deposition that was later respired to produce siderite and magnetite via DIR. See additional discussion in W Li et al. (2013a). Data from Beukes et al. (1990), Beukes and Klein (1990), [START_REF] Kaufman | Geochemical and mineralogic effects of contact metamorphism on banded iron-formation: An example from the Transvaal Basin, South Africa[END_REF], and Heimann et al. Figure 16 -Iron isotope variations among coexisting hematite, magnetite, and siderite (same sample) from the 2.48-2.47 Ga Dales Gorge Member and Kuruman Formation, respectively, as measured on various scales, from µm scale via in situ analysis by laser ablation, mg-size by bulk analysis, and g-size by bulk analysis. Equilibrium hematite-magnetite and siderite-magnetite Fe isotope fractionations at low temperature (~25 o C) are shown in red [START_REF] Wiesli | Experimental determination of Fe isotope fractionation between aqueous Fe(II), siderite and "green rust" in abiotic systems[END_REF][START_REF] Wu | Stable Fe isotope fractionations produced by aqueous Fe(II)-hematite surface interactions[END_REF]Frierdich et al., 2014), indicating that none of the minerals are in Fe isotope equilibrium.

Grey lines (marked "1:1") reflect origin through reduction of common precursor Fe(III) (modified from Haugaard et al., 2016b). Thickness and age data from: (a) [START_REF] Arndt | The age of the Fortescue Group, Hamersley Basin, Western Australia, from iron microprobe zircon U-Pb results[END_REF](b) Trendall et al., 2004;[START_REF] Anbar | A whiff of oxygen before the Great Oxidation Event?[END_REF]and (d) Pickard et al., 2002.

  Stüeken et al. (2015a) identified d 15 N values from a series of 3.2 to 2.75 Ga fluvial and marine successions that are centred on 0‰, which in their opinion is indicative of biological nitrogen fixation by the Mo-based nitrogenase. This may be difficult to reconcile owing to the low levels of Mo in the ancient ocean as a function of limited oxidative weathering (e.g.,[START_REF] Scott | Tracing the stepwise oxygenation of the Proterozoic ocean[END_REF]; however,Stüeken et al. (2015a) suggested this constraint may have been overcome by either small contributions from oxidative weathering in localised areas or with hydrothermally-sourced Mo. If we assume that the Mo necessary to allow for Mo nitrogenase to be the predominant pathway for biological nitrogen fixation was sourced by weak oxidative weathering, then the implicit implication is that cyanobacteria would have been active since ~3.2 Ga.

  terms of seawater,[START_REF] Satkoski | A redox-stratified ocean 3.2 billion years ago[END_REF] used a combination of stable Fe and radiogenic U-Th-Pb isotope data from the 3.2 Ga Manzimnyama IF of the Fig TreeGroup, South Africa to demonstrate that proximal, shallow-water samples record precipitation under more oxic conditions compared to the distal, deeper-water facies, and that the presence of U enrichment in the shallowwater samples indicates the presence of free oxygen in the photic zone (Fig.10). Results of the same approach applied to the 3.4 Ga Marble Bar Chert led W.Li et al. (2013b) to conclude that Fe(II) oxidation occurred under anoxic marine conditions by photoferrotrophy.[START_REF] Czaja | Biological Fe oxidation controlled deposition of banded iron formation in the ca. 3370 Ma Isua Supracrustal Belt (West Greenland)[END_REF], modelling Fe isotope fractionations, estimated free oxygen contents of <0.001% that of modern seawater during deposition of the Isua IF. Although this approach clearly delineates a shift in the oxidative capacity of surface seawater, the limited extent of predicted Fe(II) oxidation could have been linked to oxidation pathways independent of O 2 . Indeed, as explored below, photoferrotrophs can drive a wide range of extent of water column Fe(II) oxidation-including quantitative oxidation.

  Fe(II) inputs, we add upwelling as a mode of Fe(II) transport to the[START_REF] Kappler | Deposition of banded iron formations by anoxygenic phototrophic Fe(II)-oxidizing bacteria[END_REF] model.The thickness required to quantitatively oxidise upwelling iron is calculated similarly toKappler et al. (2005) using the expression: z = ω×ΔC/(dFe(II)/dt), where ω represents advection or upwelling rate, ΔC is the iron concentration gradient from the bottom to the top of the Fe(II)-oxidising bacterial community, and dFe(II)/dt is the rate of photoferrotrophic Fe(II) oxidation. The calculated thickness is then added to that for eddy-diffusion, using the modern global mean eddy diffusivity of 0.1 cm 2 /sec. Results with a ΔC values of 0.5 mM and 0.1 mM and upwelling rates of 0.5 m/day and 5 m/day are shown in Fig.11. Modelled advection rates (0.5 and 5 m/day) fall within the average and high ranges of annual averages observed in modern continental upwelling zones (e.g.,[START_REF] Canfield | Models of oxic respiration, denitrification and sulfate reduction in zones of coastal upwelling[END_REF].

  [START_REF] Heimann | Fe, C, and O isotope compositions of banded iron formation carbonates demonstrate a major role for dissimilatory iron reduction in ~ 2.5 Ga marine environments[END_REF] recognised that reactions 3, 5, and 6 predict changes in both the δ 13 C and δ 18 O values of IF carbonates, where a decreasing component of seawater carbonate (which has high δ 13 C and δ 18 O values) would produce decreases in both δ 13 C and δ 18 O values of carbonate that would also correlate with a shift from siderite-facies IF to oxide-facies IF, relations that are observed in the large ca. 2.5 Ga Transvaal IFs (Fig.14). Such a model explains the long-standing puzzle of why IF carbonates have lower δ 18 O values than coeval Ca-Mg carbonates, as well as the observation that the average δ 13 C value of carbonates in oxide-facies IF are lower than those in siderite-facies IF (Fig.14). This model also explains the observation that organic carbon contents are lowest in oxide-facies IF as compared to siderite-facies IF and the relation between δ 13 C and wt.% C org (Fig.14), where large proportions of electron equivalents from organic carbon lie in the magnetite inventory of oxide-facies IF. Finally, low δ 18 O values of carbonates from oxide-facies IF likely reflect their formation at a later stage in diagenesis in deeply buried sediments that were disconnected from the water column.To further explore the idea that the observed carbon isotope values result solely from the mixing of DIC from seawater and remineralised organic carbon, we use a basic mass-balance model to explore a range of environmental conditions. For simplicity, carbonate precipitation/dissolution, and accompanying buffering effects, linked with a dynamic DIC reservoir, are ignored. The [DIC] total was calculated using a modified version of the general diagenetic equation presented in[START_REF] Berner | An idealized model of dissolved sulfate distribution in recent sediments[END_REF] that was used in recent work to discuss the potential for sulfate reduction as a driver of the δ 13 C values of authigenic carbonates from the Proterozoic Chuanlingguao Formation in North China(Li et al., 2015); see equation A below. The three terms of this equation describe the processes of ionic diffusion, deposition and compaction, and input of DIC into the pore-waters from organic matter remineralisation, respectively. In this equation, G 0 is the concentration of organic carbon at the sediment/water interface; D and are the diffusion coefficient in the whole sediment and the sum of deposition and compaction, respectively; k is the rate of organic carbon remineralisation (calculated from ;[START_REF] Boudreau | Diagenetic Models and Their Implementation: Modelling Transport and Reactions in Aquatic Sediments[END_REF]; z is depth in the sediment; C is [DIC]; t is time; and L is a stoichiometric coefficient that describes the ratio between the amount of Fe(III)

  boundary condition C = , giving the following analytical solution [equation B]: [B] A basic assumption of the model is that a constant DIC concentration, C 0 , is contributed from seawater to the sediment profile and is equal to the DIC concentration of the overlying water column. Such an assumption allows for the use of the boundary condition, C 0 = [DIC] SW , with equation C providing solution for . Substituting equation C into equation B, the final solution for [DIC] total is obtained [equation D]. difference between [DIC] total and [DIC] SW to determine [DIC] OM , the mass balance relationship can be implemented to determine the carbon isotope composition of the DIC pore-water reservoir [equation E]. The model accounts for isotopic inputs of DIC from seawater (δ 13 C sw = 0‰) and the remineralisation of organic matter during DIR (δ 13 C OM = -25‰), with the isotopic values held constant and only the fluxes of DIC from each source varied with depth. This basic approach is widely used to model modern marine carbon isotope profiles, despite the intrinsic simplifications of the model: [E] [DIC] Total δ 13 C Total = [DIC] OM δ 13 C OM + [DIC] SW δ 13 C SW

  size is also required to explain the δ 13 C values of Ca-Mg (non-IF) carbonates throughout the Archaean (e.g.,[START_REF] Krissansen-Totton | A statistical analysis of the carbon isotope record from the Archean to Phanerozoic and implications for the rise of oxygen[END_REF]. Consequently, the carbonate carbon isotope record in IF is consistent with significant levels of methane cycling in the precursor sediments. In fact, the large spread in IF carbonate δ 13 C values and markedly light values are exactly what would be expected in the presence of Fe(III)-mediated organic carbon and methane oxidation in the upper part of the sediment pile. The most distinct isotopic fingerprint of this process is the observation that neither the C or O isotope compositions of IF carbonates match those expected for equilibrium with seawater.

the

  Nd-Fe isotope interpretations outlined byLi et al. (2015) are correct (see above). If the relations shown in Fig.6, which document basin-scale mixing of Fe sources, are considered alongside those in Fig.16, which document in situ DIR at the site of IF deposition, the comparison seemingly indicates that DIR may operate on both large (basin-wide Fe-shuttle) and local (in situ microbial reduction) scales. Iron may be oxidised, deposited on continental shelves, reduced, shuttled to distal locations, oxidised, settled to the deep basin along with organic carbon, and then reduced in the final location of the sediment before being lithified into IF.A seawater control on bulk-rock δ 56 Fe trends is also seen across multiple drill cores of the Transvaal BIFs[START_REF] Oonk | Fraction-specific geochemistry across the Asbestos Hills BIF of the Transvaal Supergroup, South Africa: Implications for the origin of BIF and the history of atmospheric oxygen[END_REF]. The mineral-fluid fractionation-specific δ 56 Fe values of these samples indicate a similar, consistently lighter, source for carbonate and silicate bound-Fe, whereas the oxide-bound Fe tends to have more positive δ 56 Fe values. The difference in the carbonateversus oxide-bound δ 56 Fe signals could be causally linked to DIR or to varying extents of oxidation on a basin-wide scale (e.g.,[START_REF] Oonk | Fraction-specific geochemistry across the Asbestos Hills BIF of the Transvaal Supergroup, South Africa: Implications for the origin of BIF and the history of atmospheric oxygen[END_REF].

Figure Captions Figure 1 -

 Captions1 Figure Captions

Figure 2 -

 2 Figure 2 -Iron formation deposition in billions of metric tonnes with 50 m.y. bar widths (modified from Bekker et al., 2010). Most deposits before 2.4 Ga comprise BIF; those ca. 2.3-0.8 Ga are predominantly GIF; Neoproterozoic Rapitan-type iron formation (~0.715 Ga) are associated with glacial deposits.

Figure 3 -

 3 Figure 3 -Representative images of Paleoproterozoic BIF. (A) Aerial overview of the 2.48 Ga Dales Gorge Member along the southern ridge at Mt. Tom Price, Western Australia. (B) Field photograph of basal part of rhythmically banded 2.47 Ga Kuruman Formation BIF outcropping at Kuruman Kop locality. (C) Section of banding in 2.46 Ga Joffre Member, Western Australia. (D) Weathered Kuruman BIF showing distinct iron-rich and iron-poor (sub-)millimeter thick microbands. Field of view 20 cm in height. (E-F) Photos and photomicrographs illustrating main petrographic characteristics of oxide-facies BIF in Joffre BIF. (E) Core displays pale grey microand mesobands of alternating chert and magnetite, dark grey magnetite mesobands, and reddish and bluish (1 mm thin) microbands of chert-hematite-riebeckite. Field of view 5 cm in height. (F)Coarse-and fine-magnetite microbands with chert and very fine hematite. The latter likely is a product of secondary oxidation of magnetite. Note single grains of martite in upper left. Field of view 1 cm in height. Abbreviations: ch = chert, mag = magnetite, hem = hematite, rbk = riebeckite.

Figure 4 -

 4 Figure4 -(A-B) Middle Proterozoic GIF from 1.88 Ga Frere Formation, Earaheedy Group, Western Australia. (A) Graded bed comprising granules and ooids. Field of view is 2 cm in width.(B) Rounded granules comprising hematite, magnetite, chlorite, chert, and carbonate 'floating' in a matrix of carbonate cement. Field of view is 6 mm. (C) Thin section of 2.32 Ga GIF of Timeball Hill Formation, South Africa.

Figure 5 -

 5 Figure 5 -PAAS-normalized (McLennan, 1989) Eu anomaly (Eu/Eu*) through time of some of the important IFs. Data for Isua from Frei and Polat (2007); Nuvvuagittuq from Mloszewska et al. (2012, 2013); Fig Tree from Derry and Jacobsen (1990) and Hofmann (2005); Jharkhand-Orissa from Bhattacharya et al. (2007); Pongola from Alexander et al. (2008); Itilliarsuk from Haugaard et al. (2013); Slave Craton from Haugaard et al. (2016a, 2017b); Krivoy Rog from Viehmann et al.(2015); Joffre fromHaugaard et al. (2016b). Kuruman and Penge from[START_REF] Bau | Distribution of yttrium and rare-earth elements in the Penge and Kuruman iron-formations, Transvaal Supergroup, South Africa[END_REF]; Nemo fromFrei et al. (2008); Gunflint from[START_REF] Danielson | The europium anomalies in banded iron formations and the thermal history of the oceanic crust[END_REF]; Braemar from[START_REF] Lottermoser | Geochemistry, petrology and origin of Neoproterozoic ironstones in the eastern part of the Adelaide Geosyncline, South Australia[END_REF]; Rapitan from[START_REF] Halverson | Fe isotope and trace element geochemistry of the Neoproterozoic syn-glacial Rapitan Iron Formation[END_REF].

Figure 6 -

 6 Figure 6 -Neodymium and Fe isotope data from Marra Mamba and Brockman IFs and associated shales of the Hamersley Basin, Western Australia. ε Nd (t) values for IFs scatter between mantle/primitive crust composition and those of old crust, indicating a mix of REE+Y sources. Intermediate ε Nd (t) values for associated shales indicate local continental crust may have near-zero ε Nd (t), which would increase the continental component of REE+Y relative to an assumed continental ε Nd (t) value of -3. Positive correlation between ε Nd (t) and δ 56 Fe seen in 2.48 Ga Dales Gorge Member is interpreted to reflect mixing between a hydrothermal (high-ε Nd (t), high-δ 56 Fe) iron source and iron mobilized by a basinwide Fe shuttle via DIR on continental shelf. Data sources:

δ

  30 Si values of IF cherts cannot be explained solely by Si sorption to Fe(III)-oxyhydroxides, but likely involves Fe(III)-Si gels as the primary phase. As discussed in[START_REF] Reddy | Silicon isotope fractionation during microbial reduction of Fe(III)-Si gels under Archean seawater conditions and implications for iron formation genesis[END_REF], magnetiterich IFs have lower δ 30 Si values than hematite-rich IFs; measured Si isotope fractionations suggest production of magnetite-quartz assemblages that have very low δ 30 Si values through DIR. Data sources:[START_REF] André | Silicon isotopes in ∼3.8 Ga West Greenland rocks as clues to the Eoarchaean supracrustal Si cycle[END_REF],Robert and Chaussidon (2006), van den Boom et al. (2007; 2010),[START_REF] Delvigne | Stratigraphic changes of Ge/Si, REE+Y and silicon isotopes as insights into the deposition of a Mesoarchaean banded iron formation[END_REF]), Steinhoefel et al. (2009, 2010),[START_REF] Abraham | Coupled silicon-oxygen isotope fractionation traces Archaean silicification[END_REF][START_REF] Heck | SIMS analyses of silicon and oxygen isotope ratios for quartz from Archean and Paleoproterozoic banded iron formations[END_REF],[START_REF] Chakrabarti | Si isotope variability in Proterozoic cherts[END_REF],[START_REF] Li | Precambrian banded iron formations in the North China Craton: Silicon and oxygen isotopes and genetic implications[END_REF][START_REF] Geilert | Silicon isotopes and trace elements in chert record early Archean basin evolution[END_REF],[START_REF] Stefurak | Texture-specific Si isotope variations in Barberton Greenstone Belt cherts record low temperature fractionations in early Archean seawater[END_REF],[START_REF] Ding | The δ 30 Si peak value discovered in middle Proterozoic chert and its implication for environmental variations in the ancient ocean[END_REF].

Figure 8 -

 8 Figure 8 -Iron-O isotope variations in coexisting hematite and magnetite from the Dales Gorge

Figure 9 -

 9 Figure 9 -Mechanisms of Fe(II) oxidation in Precambrian oceans. Two biologically controlled mechanisms are envisioned: (1) reaction of cyanobacterially-generated O 2 with dissolved Fe(II),and/or (2) direct oxidation via Fe(II)-based anoxygenic photosynthesis (photoferrotrophy). The Fe(II) was sourced from both deep-sea hydrothermal systems and continental weathering, whereas the Fe(III) formed in the photic zone was precipitated as ferrihydrite, Fe(OH) 3 , and deposited onto the seafloor as a precursor sediment for IF. Some ferrihydrite was later reduced either through direct bacterial Fe(III) reduction (DIR) utilizing organic carbon or through a potential metabolic coupling of Fe(III) reduction and methane oxidation (e.g.,[START_REF] Konhauser | The potential significance of microbial Fe(III) reduction during deposition of Precambrian banded iron formations[END_REF].

Figure 10 -

 10 Figure 10 -Secular variations in atmospheric O 2 (green curve) and shallow ocean O 2 (blue boxes).

Figure 11 -

 11 Figure 11 -Calculated thickness of a community of photoferrotrophs required to completely oxidize upwelling, Fe(II)-rich, hydrothermal waters with Fe(II) concentrations of (A) 500 µM and (B) 100 µM, at different rates of Fe(II) oxidation. Unless noted, advection and eddy-diffusion Fe

Figure 14 -

 14 Figure 14 -Co-variation in C and O isotope compositions of carbonates from 2.6-2.5 Ga Kuruman IF and Campbellrand-Malmami carbonate platform in South Africa, and bulk-rock organic carbon contents. Ca-Mg carbonates (limestone and dolostone) shown by green circles, ankerite/siderite from siderite-facies IF by blue diamonds, and ankerite/siderite from oxide-facies IF by red squares.

(

  2010). Transparent grey box in left diagram shows C and O isotope compositions for siderite that would be in equilibrium with seawater, showing that none of the siderite in either siderite-facies IF or oxide-facies IF is in C or O isotope equilibrium with seawater. Parallel grey curves in left diagram show mixing lines between seawater carbonate and carbonate produced by DIR at various ratios of organic C to seawater carbon (see reactions 3, 5, and 6 in main text). Estimate for primary organic carbon content of IF, prior to DIR, shown in right diagram from Figure 13. See Heimann et al. (2010) for details.

Figure 15 -

 15 Figure 15 -Calculated δ 13 C DIC for sediments dominated by anaerobic organic matter remineralisation. Unless noted in legend, sediment and water column parameters are for Santa Barbara Basin (from Berner, 1964). Water column DIC concentration is taken from modern continental margin values; sedimentation rate and organic matter concentrations are varied at ranges typical of continental margin settings. A DIC concentration of 10x modern one is used as estimate for Archaean oceans. (A) δ 13 C DIC vertical pore-water profiles for modern Santa Barbara Basin conditions compared with modelled Archaean settings with 10x DIC and various organic matter fluxes. (B) δ 13 C DIC vertical pore-water profiles for modelled Archaean settings with DIC 10x of modern seawater showing effect of variations in the sum of sedimentation and compaction (ω) from that in the modern Santa Barbara Basin. (C) Effect of variations in ω, at DIC 10x that of modern seawater and labile organic matter load at 0.5x that of the modern Santa Barbara Basin, on δ 13 C DIC vertical pore-water profiles for modelled Archaean settings. (D) Effect of variations in ω, at DIC 10x that of modern seawater and labile organic matter load at 0.1x that of the modern Santa Barbara Basin, on δ 13 C DIC vertical pore-water profiles for modelled Archaean settings.

  oxyhydroxide via "in situ" DIR in sediment prior to lithification. Broad variations in δ 56 Fe values, from positive to negative values, is interpreted to reflect various Fe sources, ranging from hydrothermal (positive δ 56 Fe) to DIR shuttle from continental shelves (negative δ 56 Fe); see Figure 6. Sources: Johnson et al. (2003; 2008b); Steinhoefel et al. (2010); Heimann et al. (2010); Craddock and Dauphas (2011).

Figure 17 -

 17 Figure 17 -Model for the evolution of ocean redox structure based on REE patterns in wellpreserved IF. Mn(IV)-poor IF are likely to record seawater REE+Y patterns. Similar to modern redox-stratified basins (A), REE+Y pattern of late Paleoproterozoic IF (B) records evidence for a shuttle of metal and Ce oxides from oxic shallow seawater across the redoxcline. Mn(IV)-hydroxide dissolution in anoxic water column lowers Y/Ho ratios, raises LREE to HREE ratios, and increases Ce concentration relative to neighboring REE (La and Pr). In contrast, Archaean IF do not display REE+Y pattern indicative of a strong oxide shuttle (C), which implies a lack of significant Mncycling across a redoxcline in many Archaean IF-bearing basins. Absence of discrete redoxcline points toward microbial Fe(II) oxidation, rather than a direct reaction of Fe(II) with O 2 . Modified from Planavsky et al. (2010a).

Figure 18 -

 18 Figure 18 -(A) Regional geology and (B) stratigraphy of the Hamersley Group, Western Australia

  New research on Si isotopes provides insight into the primary solid phase of Fe(III) in IF genesis due to the strong control that Fe-Si bonding exerts on Si isotope fractionations. δ 30 Si values

	for Precambrian cherts, including those associated with IFs, vary by ~7‰, and broad secular trends
	of increasing δ 30 Si with decreasing age have been attributed to decreasing ocean temperature
	(Robert and Chaussidon, 2006), although this proposal has been difficult to test owing to the lack of
	experimental determination of Si isotope fractionation factors. A recent review of the Si isotope
	record has highlighted the fact that δ 30 Si values for IFs are systematically lower than those of
	similar-aged cherts that are not associated with IFs (Marin-Carbonne et al., 2014), which is ascribed
	to the effects of Si sorption to Fe(III)-oxyhydroxides

  [START_REF] Olson | Quantifying the areal extent and dissolved oxygen concentrations of Archean oxygen oases[END_REF] estimated that photic zone O 2 contents could be as high as ~10 µM under an essentially anoxic atmosphere. Based on modelling of Fe isotopes in the Manzimnyama jasper,[START_REF] Satkoski | A redox-stratified ocean 3.2 billion years ago[END_REF] calculated shallow photic zone O 2 contents at ~10 µM (between 10 -3 and 10 -2 POL, or Present Ocean Level; Fig.10), and deep-water O 2 contents at ~0.4 µM. The calculations of Czaja et al. (2012), based on covariation in Mo and Fe isotope compositions in carbonates of the 2.65-2.5 Ga Campbellrand-Malmami carbonate platform, estimated 35 µM O 2 content in the photic zone during this time period, being ~12% of modern level. The potential decoupling of photic zone and atmospheric O 2 content prior to the GOE is important in our understanding of the roles of O 2 sources and sinks, and of atmospheric evolution prior to the GOE.

  is absent in the REE+Y plots of the Joffre Member as well, suggesting O 2 levels below the Ce(III)/Ce(IV) redox boundary. Iron isotope compositions in both the Joffre and the Dales Gorge members show a wide range in d 56 Fe values, indicating multiple Fe sources and pathways (Fig. 6), where average d 56 Fe values in the Joffre Member are heavier than those of the Dales Gorge Member

). The two IFs display similar REE+Y patterns, including large Eu anomalies; however, the Joffre Member lacks a strong positive Y anomaly, which is typically associated with a seawater geochemical signature. A negative Ce anomaly

  . It has been recognised for some time that the volume of continental crust was substantial in the early Archaean (see review by[START_REF] Hawkesworth | The generation and evolution of the continental crust[END_REF], but to date it has been unclear how continental freeboard controlled the amount of emergent continental landmasses. Continental freeboard depends more on mantle temperature, crustal composition, and continental hypsometry (also temperature-dependent) than on the growth of continental crust. Recent models suggest that continental landmasses did not reach modern level of emergence in the Archaean(Korenega et al., 2017). Limited emergent continental crust would not only have severely limited the availability of continental crustal sources for weathering, but it would have compounded nutrient limitation for elements such as Mo supplied dominantly by oxidative weathering of crustal materials (see above discussion). Moreover, the amount of continental freeboard directly affects the relative proportions of submarine versus subaerial volcanism and thereby the reducing equivalents of volcanic gases, in the[START_REF] Kump | Increased subaerial volcanism and the rise of atmospheric oxygen 2.5 billion years ago[END_REF] model discussed above. Such conclusions, however, conflict with the growing evidence for elevated 87 Sr/ 86 Sr ratios in the early Archaean oceans that point to significant exposure of evolved continental crust at that time

  compilation of Cu isotopes in Precambrian marine sediments. Although the Cu concentration record presented by[START_REF] Fru | Cu isotopes in marine black shales record the Great Oxidation Event[END_REF] for IFs and shales is relatively constant in terms of abundancereminiscent of the Zn record raising the possibility of stabilisation of these trace metals by organic ligand complexation (see[START_REF] Robbins | Authigenic iron oxide proxies for marine zinc over geological time and implications for eukaryotic metallome evolution[END_REF][START_REF] Scott | Bioavailability of zinc in marine systems through time[END_REF] -there is a clear trend in the stable isotopes values of Cu. Following the GOE, Cu isotopes in black shales become progressively heavier. This trend has been attributed to secular changes in seawater composition, due to the combined effect of waning IF deposition, which prior to the GOE would have preferentially removed65 Cu with Fe(III)-oxyhydroxides, and increased oxidative supply of 65 Cu derived from continental weathering due to preferential leaching of the heavier Cu isotope. This 65 Cu, in turn, would have become incorporated into planktonic biomass that scavenged Cu from seawater.

  KOK, TJW, RH, CAP and AB would like to thank the Natural Sciences and Engineering Research Council of Canada (NSERC) for its financial support. LJR gratefully acknowledges the support of a Vanier Canada Graduate Scholarship. CMJ, DSH, NJP and TWL acknowledge support from the NASA Astrobiology Institute. SVL acknowledges support from the European Institute for Marine Studies (LabexMER, ANR-10-LABX-19). HT and PBHO thank ASSMANG Ltd and SOUTH32 for providing research funding and access to drillcores, respectively. The authors thank Drs. Philip Fralick and John Slack for comments that greatly improved the manuscript, and Dr. André Strasser for his editorial efforts.

Frei, R., Crowe, S.A., Bau, M., Polat, A., Fowle, D.A., Døssing, L.N., 2016. Oxidative elemental cycling under the low O 2 Eoarchean atmosphere. Scientific Reports. 6, 21058. [START_REF] Oonk | Fraction-specific geochemistry across the Asbestos Hills BIF of the Transvaal Supergroup, South Africa: Implications for the origin of BIF and the history of atmospheric oxygen[END_REF]. Thickness and age data from: (a) [START_REF] Kunzmann | Depositional environment and lithostratigraphy of the Paleoproterozoic Mooidraai Formation, Kalahari manganese field, South Africa[END_REF](b) Tsikos et al., 2003;(c) [START_REF] Polteau | The geology and geochemistry of the Palaeoproterozoic Makganyene diamictite[END_REF][START_REF] Ayres | Genesis of iron-bearing minerals in banded iron formation mesobands in the Dales Gorge Member, Hamersley Group, Western Australia[END_REF] [START_REF] Kendall | Depositional age of the early Paleoproterozoic Klipputs Member, Nelani Formation (Ghaap Group, Transvaal Supergroup, South Africa) and implications for low-level Re-Os geochronology and Paleoproterozoic global correlations[END_REF](e) Sumner, 1997;(f) [START_REF] Bau | Pb, O, and C isotopes in silicified Mooidraai dolomite (Transvaal Supergroup, South Africa): Implications for the composition of Paleoproterozoic seawater and 'dating' the increase in oxygen in the Precambrian atmosphere[END_REF](g) Cornell et al., 1996;(h) [START_REF] Gumsley | Timing and tempo of the Great Oxidation Event[END_REF](i) [START_REF] Moore | The age and correlation of the Postmasburg Group, southern Africa: Constraints from detrital zircon grains[END_REF](j) Trendall et al., 1990;(k) Pickard, 2003;(l) [START_REF] Sumner | U-Pb geochronological constraints on deposition of the Campbellrand Subgroup, Transvaal Supergroup, South Africa[END_REF](m) Altermann and Nelson, 1998;and (n) [START_REF] Walraven | Zircon Pb-evaporation age determinations for the Oak Tree Formation, Chuniespoort Group, Transvaal Sequence: Implications for Transvaal-Griqualand West basin correlations[END_REF]