The current understanding of the evolution of the atmosphere, hydrosphere, and biosphere on early Earth has been strongly influenced by the following six major paradigms for the geochemical cycles of oxygen, iron, and sulfur: (1) a dramatic change from a reducing to an oxidizing atmosphere at ca. 2.4–2.2 Ga, termed the “Great Oxidation Event” (GOE); (2) Fe-rich oceans until ca. 1.85 Ga; (3) a hydrothermal origin for the global oceanic Fe; (4) SO 4 2− -poor oceans before the GOE; (5) an atmospheric origin for the oceanic sulfur species; and (6) the existence of sulfidic Proterozoic oceans. Each of the six paradigms has been built on other paradigms, such as those concerning: (1) the behavior of Fe during soil formation, (2) the environments and processes required for the formation of Fe III oxides in banded iron formations (BIFs), and (3) the origins of siderite and pyrite, as well as (4) the origin of anomalous isotope fractionation of sulfur (AIF-S) in Archean sedimentary rocks. Here, we show that some of the paradigms contradict each other, and that each has serious flaws (contradictions, problems) when they are compared to a variety of observations (geological, mineralogical, or geochemical data from natural samples; laboratory experimental data; results of theoretical studies). In contrast, all of these observations are better explained by the Dimroth-Ohmoto model for Earth's evolution, which postulates that a fully oxygenated atmosphere-ocean system developed by ca. 3.5 Ga. Examination of the available data from natural and experimental systems has also led us to suggest the following: (1) The geochemical cycles of O, Fe, and S (and other redox-sensitive elements) through the atmosphere–ocean–oceanic crust–mantle–continental crust have been basically the same as today since at least ca. 3.5 Ga. (2) The anaerobic and aerobic microbial biospheres, both in the oceans and on land, developed by ca. 3.5 Ga, playing an important role in the geochemical cycles of nutrients and other elements. (3) The geochemistry of sedimentary rocks (shales, carbonates, cherts) has been basically the same since ca. 3.5 Ga. (4) Fe III oxides in BIFs were formed by reactions between locally discharged Fe 2+ - and silica-rich submarine hydrothermal fluids and O 2 -rich deep seawater. (5) Magnetite in BIFs was formed during high-temperature diagenetic stages of BIFs through reactions between primary goethite or hematite and Fe 2+ -rich hydrothermal fluids. (6) BIFs were formed throughout geologic history. (7) Sulfidic oceans (i.e., the “Canfield ocean”) did not exist during the Proterozoic Eon. However, regional sulfidic seas, like the Black Sea, have existed in globally oxygenated oceans throughout geologic history. (8) The primary carbonate in Archean oceans, as in younger oceans, was Fe-poor calcite. Furthermore, (9) the pre–1.8 Ga atmosphere was CO 2 rich with the p CO 2 level greater than ~100 present atmospheric level (PAL). CO 2 alone provided the green-house effect necessary to compensate for the young Sun's lower luminosity. (10) The Archean pH values were 4.0–4.5 for rainwater, between 4.5 and 6.0 for river water, and 7.0 ± 0.5 for ocean water. The oceans were saturated with calcite but under-saturated with siderite. (11) The δ 18 O of Archean oceans was ~0‰, as today. (12) Fe-rich carbonates (siderite, ankerite) have formed during the diagenesis of sediments throughout geologic history by reactions between the primary calcite and Fe 2+ -rich solutions, either hydrothermal solutions or those derived from biological or abiological dissolution of Fe III -(hydr)oxides within the sediments. Other suggestions include: (13) The ranges of δ 34 S values of pyrite and sulfates in Archean sedimentary rocks are much larger than those quoted in the literature and comparable to those in Proterozoic sedimentary rocks. (14) Pyrites in organic C–rich black shales associated with BIFs were formed by a reaction between Fe 2+ - and SO 4 2− -rich hydrothermal solutions and organic C–rich shales during early diagenetic stages of the host sediments. This reaction also created AIF-S in the pyrite and the residual SO 4 2− . (15) The AIF-S signatures in Archean and younger rocks were not created by the ultraviolet photolysis of volcanic SO 2 in a reducing atmosphere. AIF-S signatures are not evidence for a reducing atmosphere. (16) Contrary to a popular belief that AIF-S–forming events ceased at ca. 2.45 Ga, AIF-S was also formed at later geologic times. (17) The presence of AIF-S signatures in some pre–2.4 Ga rocks, but the lower abundance of AIF-S in post–2.4 Ga rocks, may reflect changes in the mantle-crust dynamics, including changes in the thickness and movements of oceanic lithosphere.

The tension between CO 2 dissolved at relatively high atmospheric pressure in the Hadean ocean, and H 2 generated as ocean water oxidized ferrous iron during convection in the oceanic crust, was resolved by the onset of life. We suggest that this chemosynthetic life emerged within hydrothermal mounds produced by alkaline solutions of moderate temperature in the relative safety of the deep ocean floor. Exothermic reaction between hydrothermal H 2 , HCOO − and CH 3 S − with CO 2 was catalyzed in inorganic membranes near the mound's surface by mackinawite (FeS) nanocrysts and “ready-made” clusters corresponding to the greigite (Fe 5 NiS 8 ) structure. Such clusters were precursors to the active centers (e.g., the C-cluster, Fe 4 NiS 5 ) of a metalloenzyme that today catalyzes acetate synthesis, viz., the bifunctional dehydrogenase enzyme (ACS/CODH). The water, and some of the acetate (H 3 C.COO − ), produced in this way were exhaled into the ocean together as fluid waste. Glycine ( + H 3 N.CH 2 .COO − ) and other amino acids, as well as tiny quantities of RNA, generated in the same milieu were trapped within tiny iron sulfide cavities. Energy from the acetate reaction, augmented by a proton gradient operating through the membrane, was spent polymerizing glycine and other amino acids into short peptides upon the phosphorylated mineral surface. In turn these peptides sequestered, and thereby protected, the catalytically and electrochemically active pyrophosphate and iron/nickel sulfide clusters, from dissolution or crystallization. Intervention of RNA as a polymerizing agent for amino acids also led to an adventitious, though crude, process of regulating metabolism—a process that was also to provide genetic information to offspring. The fluxes of energy and nutrient available in the hydrothermal mound—commensurate with the requirements of life—encouraged differentiation of the first microbes into two separate domains. At the bifurcation the Bacteria were to specialize in acetogenesis and the Archaea into methanogenesis. Representatives of both these domains left the mound by way of the ocean floor and crust to colonize the deep biosphere. Once life had emerged and evolved to the extent of being able to reduce nitrogen for use in peptides and nucleic acids, light could have been used directly as an energy source for biosynthesis. Certain bacteria may have been able to do this, where protected from hard UV by a thin coating of chemical sediment produced at a sub-aerial hot spring operating in an obducted and uplifted portion of the deep biosphere. Embedded in fresh manganiferous exhalites, early photosynthetic bacteria could further protect themselves from radiation by adsorbing manganese on the membrane. Organization of the manganese with calcium, within a membrane protein, happened to result in a CaMn 3 O 4 cluster. In Mn(IV) mode this structure could oxidize two molecules of water, evolve waste oxygen, and gain four electrons and four protons in the process to fix CO 2 for biosynthesis. All these biosynthetic pathways had probably evolved before 3.7 Ga, though the reduced nature of the planet prevented a buildup of free atmospheric oxygen until the early Proterozoic.

Carbon and sulfur isotopes have been measured on samples from four Archean greenstone belts dating from 3.8 Ga to 2.7 Ga, in order to trace metabolic changes as life evolved over this one-billion-year period. In the Isua Greenstone Belt (3.8 Ga), Greenland, δ 34 S in sulfide minerals from sedimentary sequences range from −3.8‰ to +3.4‰. δ 13 C red measured in BIFs, turbidites and conglomerates vary from −29.6‰ to −14.7‰; this range permits us to hypothesize the presence of hyperthermophilic and chemotrophic species in transient settings, or possibly pelagic photoautotrophic microbes, or both. In the Barberton Greenstone Belt, South Africa, sulfide minerals show δ 34 S values from +1.5‰ to +5.6‰. Black shales have δ 13 C red values from −32.4‰ to −5.7‰, suggesting that oxygenic photosynthetic and sulfate-reducing bacteria were present by ca. 3.24 Ga. The δ 13 C red measured in the stromatolites of Steep Rock Lake (3.0 Ga), Ontario, Canada, are from −30.6‰ to −21.6‰, giving clear evidence for occupation of a shallow water environment by cyanobacteria. The wide isotopic ranges for δ 34 S in sulfides from −21.1‰ to +16.7‰ and for δ 13 C red in carbon-rich cherts and black shales from −43.4‰ to −7.2‰ in the Belingwe Greenstone Belt, Zimbabwe, indicate that photosynthetic microbial mat communities were well established at 2.7 Ga. In these well-preserved Late Archean formations, modern-style biological sulfur and carbon cycles may have been in operation. The δ 34 S and δ 13 C red ranges, respectively 37‰ and 36‰, indicate a great variety of biological processes interacting with each other.

Archean cratons are fragments of old continents that are believed to be more richly endowed with mineral deposits than younger terrains. The mineral deposits of different cratons are also diversely enriched with useful (to humankind) chemical elements. Cratons are therefore mineral diversity hotspots that represent regional geochemical heterogeneities of early Earth, the evidence for which remains encoded on each craton as unique metallogenic “fingerprints.” Using six selected elements groups from our extensive in-house GIS database of Gondwana mineral deposits, we derive the metallogenic fingerprints of 11 Archean cratons of the Southern Hemisphere, and compare these against metallogenic fingerprints of the same selected elements in younger crust of three of their host continents (Africa, Australia, and South America). After adjusting the mineral inventory of each craton to account for underexploration of regions lacking infrastructure and other political and economic conditions for mineral investment, we show that mineral deposit density and diversity of Earth's continental lithosphere has decreased with time. We conclude that metallogenic elements were transferred more efficiently from the mantle to the continental lithosphere in the Archean and/or that subsequent (younger than 2.5 Ga) recycling of these elements (mineral deposits) back into the mantle has become more effective. How most of these fragments of old continents inherited their rich and diverse metallogenic characteristics is unresolved, because different cratons are likely to represent only small remnants of once much larger and possibly more varied Archean continents, and part of the total metal inventory of Archean continents must have been recycled back into in the mantle. The latter has implications for understanding the secular change in the redox state of the Archean mantle and fluid envelope.

Phanerozoic granitoids have been classified into magnetite and ilmenite series based on the abundance of magnetite, which is related to the Fe 2 O 3 /FeO ratio of the rock and the oxygen fugacity ( f O 2 ) of its parent magma. We have examined the temporal and spatial distributions of both series in Archean granitoids from the Barberton region and the Johannesburg Dome of the Kaapvaal Craton, South Africa. The oldest syntectonic TTG (tonalite-trondhjemite-granodiorite) granitoids (ca. 3450 Ma in age) were found to be ilmenite series, whereas some intermediate-series granitoids occurred locally. Younger and larger syntectonic TTGs (e.g., the 3230 Ma Kaap Valley plutons) comprise nearly equal quantities of magnetite and ilmenite series. The major 3105 Ma calc-alkaline batholiths (e.g., Nelspruit batholith), emplaced during the late-tectonic stage, comprise mostly magnetite-series granitoids, suggesting that an oxidized continental crust already existed by this time. The rare earth element ratios and δ 18 O values, as well as the Fe 2 O 3 /FeO ratios, of the Archean magnetite-series granitoids suggest that their magmas were generated from the partial melting of subducted oceanic basalts that had been oxidized by interaction with seawater on mid-oceanic ridges; the processes of magma generation were much like those for Phanerozoic magnetite-series granitoids. This further suggests that the concentrations of oxidants (O 2 and/or SO 4 2− ) in the Archean oceans were similar to those in Phanerozoic oceans. Low concentrations of chlorine in the magmas, as well as deep levels of granite erosion, appear to explain the absence of major mineral deposits associated with the Kaapvaal granitoids.

New δ 15 N analyses combined with a literature compilation reveal that shale kerogen, VMS-micas, and late-metamorphic vein micas show a secular trend from enriched values in the Archean, through intermediate values in Proterozoic terranes, to the Phanerozoic mode of 3‰–4‰. Kerogen in metashales from the 2.7 Ga Sandur Greenstone Belt, eastern Dharwar Craton, India, is characterized by δ 15 N 13.1‰ ± 1.3‰, and C/N 303 ± 93. A second population has δ 15 N 3.5‰ ± 0.9‰, and C/N 8 ± 0.4, close to the Redfield ratio of modern microorganisms, and is interpreted as precipitates of Proterozoic or Phanerozoic oilfield brines that penetrated the Archean basement. Kerogen from 1.7 Ga carbonaceous shales of the Cuddapah Basin average 5.0‰ ± 1.2‰, close to the mode at 3‰–4‰ for kerogen and bulk rock of Phanerozoic sediments. Biotites from late-metamorphic quartz-vein systems of the 2.6 Ga Kolar gold province, E. Dharwar Craton, that proxy for average crust, are also enriched at 14‰–21‰ for three samples, confirming that the N–budget of the hydrothermal fluids is dominated by sedimentary rocks. Muscovites from altered volcanic rocks in 2.7 Ga Abitibi belt VMS deposits have δ 15 N 12‰–20‰, in keeping with published data for shale kerogen from the same terrane, whereas equivalents in the 1.8 Ga Jerome VMS span 11.7‰–14.1‰. 15 N-enriched values in Precambrian rocks cannot be caused by N-isotopic shifts due to metamorphism or Rayleigh fractionation because (1) pre-, and post-metamorphic samples from the same terrane are both enriched in 15 N; (2) there is no covariation of δ 15 N with N, C/N ratios, or metamorphic grade; and (3) the magnitude of fractionations of 1‰ (greenschist) to 3‰ (amphibolite facies) during progressive metamorphism of sedimentary rocks, as constrained from empirical observations and experimental studies, is very small. Nor can 15 N-enriched values stem from long-term preferential diffusional loss of 14 N, as samples were selected from terranes where 40 Ar/ 39 Ar ages are within a few million years of concordant U-Pb ages; nitrogen is structurally bound in micas, whereas Ar is not. It is possible that the 15 N-enriched values stem from a different N-cycle in the Archean, with large biologically mediated fractionations, yet the magnitude of the fractionations between atmospheric N 2 and organic nitrogen observed exceeds any presently known, and chemoautotrophic communities tend to depleted values. Earlier results on Archean cherts show a range of δ 15 N from −6‰ to 30‰. Given the temporal association of chert–banded iron formation (BIF) with mantle plumes, the range is consistent with mixing between mantle N 2 of −5‰ and the 15 N-enriched marine reservoir identified in this study. The 15 N-enriched Archean atmosphere-hydrosphere reservoir does not robustly constrain Archean redox-state. We attribute the 15 N-enriched reservoir to a secondary atmosphere derived from CI-chondrite-like material and comets with δ 15 N of +30‰ to +42‰. Shifts of δ 15 N to its present atmospheric value of 0‰ can be accounted for by a combination of early growth of the continents with sequestration of atmospheric N 2 into crustal rocks, and degassing of mantle N ∼−5‰. If Earth's surface environment became oxygenated ca. 2 Ga, then there were no associated large N-isotope excursions.

The 3.05 Ga U-Pb dating of uraninite grains in Dominion and Witwatersrand conglomerates has established that they were older than the onset of Witwatersrand sedimentation at 2.97 Ga. and therefore that they are detrital in origin. A precise Re-Os isochron age of 2.99 Ga obtained for rounded pyrite grains associated with the uraninite indicates that the pyrite is also detrital. Evidence of detrital forms in the gold has confirmed its placer origin prior to modification during metamorphism. Furthermore, rhenium-depletion ages ranging from 3.5 to 2.9 Ga for Witwatersrand gold support numerous other lines of evidence that have been used in the past to interpret the gold, uranium and pyrite concentrates as Archean placers. The sedimentary and stratigraphic history of the Witwatersrand succession indicates that uraninite, pyrite, and gold were part of a sub-aerial sediment load over a period of 180 m.y. Net sedimentation rate during the accumulation of the Central Rand Group is estimated at less than 14 m per million years, reflecting stratigraphic losses due to repeated reworking. Individual paleoplacers would have been exposed over areas of up to 400 km 2 but collectively they covered a region exceeding 2000 km 2 and contained more than 243 million tons of pyrite, 1.5 million tons of uraninite, and 80,000 tons of gold. Because there is no record of either detrital uraninite or pyrite in Proterozoic red-bed sediments on the Kaapvaal Craton, it is concluded that a change in the composition of the atmosphere took place there after 2.64 Ga.

The first serious suggestion that the Archean atmosphere was reducing was based on the interpretation of round uraninite and pyrite grains in the Witwatersrand Basin in the early 1950s. It was then inferred that these minerals were detrital and that they reflected equilibrium with a reducing Archean atmosphere. Over the past 20 years the understanding of the Witwatersrand Basin has changed dramatically with more integrated studies of the basin and the recognition of widespread alteration in close spatial association with the mineralization in every goldfield. Post-depositional mobility of gold, sulfur, and uranium during alteration is widespread and supports hydrothermal ore genesis, or at least substantial modification of the original mineral assemblage. Pseudomorphic replacements of pre-existing detrital minerals (e.g., pyrite after titano-magnetite), and precipitation and/or chemical rounding to generate round mineral shapes (e.g., uraninite in carbon seams) have all been documented. The recognition that the carbon seams formed by the post-depositional introduction and maturation of migrated hydrocarbons is a dramatic departure from earlier models of coalified algal material deposited with the sediments. The enrichment of both gold and uraninite in carbon seams implies that these minerals are hydrothermal and that their shapes do not reflect detrital processes. Uranium mobility in basinal waters may in fact require a relatively oxidizing atmosphere. None of the existing arguments for the Witwatersrand mineralization unambiguously support a placer or modified placer model for the mineralization. Consequently, round uraninite and pyrite of the Witwatersrand Basin do not provide support for a reducing Archean atmosphere.

The ca. 2.45 Ga pyritic uraniferous quartz-pebble conglomerate (UQC) of the Matinenda Formation of the Elliot Lake Group, Huronian Supergroup, was used in this study to investigate the origin of pyrite. A laser-microprobe was used for analysis of the sulfur isotopic compositions of individual pyrite grains, and an electron-probe microanalyzer was used for analysis of the trace element compositions of pyrite grains with overgrowth texture. We found a variation in δ 34 S values among pyrite crystals (73 analyses) of various size and morphologies that occur in a small (∼1 cm 3 ) rock chip: the total range in δ 34 S is −9.0‰ to +5.5‰ with respect to CDT (Cañon Diablo Troilite) with a mean value of +0.6‰ ±2.1‰ (1σ). The widest range of ∼15‰ is found among euhedral pyrite grains whereas variations of ∼4‰ to ∼6‰ are common in anhedral, subhedral, and rounded grains of pyrite. These values are in marked contrast to the δ 34 S values of pyrite from the Matinenda Formation that were obtained by previous investigators using bulk-rock sulfur isotope analyses. We found variable concentrations of Co (below detection to 4700 ppm), Ni (to 1900 ppm), and As (to 3400 ppm) among individual pyrite crystals and within single grains with overgrowth textures. These elemental concentrations are markedly different between core and overgrowth parts of pyrite. We demonstrate that the pyrite grains in the Paleoproterozoic UQC have been isotopically, chemically, and morphologically modified by post-depositional processes, suggesting that the pyrite grains have undergone multiple generations. The results of the present study cannot be explained solely by a detrital process. Therefore, one cannot use the preserved morphology and chemistry of pyrite (and possibly uraninite) to represent the original features at the time of deposition to support the hypothesis of an anoxic atmosphere prior to 2.2 Ga.

Natural fission reactors at the Oklo uranium deposits in Gabon appear to have formed in a short interval of geologic time during which uranium could migrate to form deposits and the 235 U/ 238 U ratio was still high enough to trigger fission reactions. At the time of sediment deposition in the ore-hosting Franceville Basin ∼2100 m.y. ago, the oxygen deficient atmosphere would have inhibited uranium dissolution and therefore its migration to form deposits. Dissolution and migration of uranium probably began only during later diagenesis after ca. 2050 Ma, and local reduction reactions in the presence of hydrocarbons allowed formation of high-grade uranium deposits. At this time the 235 U/ 238 U ratio was still significantly higher than it is today, thus triggering nuclear fission reactions. Before 2.0 Ga, the 235 U/ 238 U ratio was also high enough to allow fission reactions but no mechanisms were able to produce high-grade uranium ores. Thus, oxygen in the atmosphere was probably the main factor controlling the occurrence of natural nuclear fission reactions. This conclusion is in agreement with earlier suggestions that oxygen contents in atmosphere increased during a “transition phase” some 2450–2100 m.y. ago.

Sedimentary exhalative (SEDEX) Zn-Pb-sulfide mineralization first occurred on a large scale during the late Paleoproterozoic. Metal sulfides in most Proterozoic deposits have yielded broad ranges of predominantly positive δ 34 S values traditionally attributed to bacterial sulfate reduction. Heavy isotopic signatures are often ascribed to fractionation within closed or partly closed local reservoirs isolated from the global ocean by rifting before, during, and after the formation of Rodinia. Although such conditions likely played a central role, we argue here that the first appearance of significant SEDEX mineralization during the Proterozoic and the isotopic properties of those deposits are also strongly coupled to temporal evolution of the amount of sulfate in seawater. The ubiquity of 34 S-enriched sulfide in ore bodies and shales and the widespread stratigraphic patterns of rapid δ 34 S variability expressed in both sulfate and sulfide data are among the principal evidence for global seawater sulfate that was increasing during the Proterozoic but remained substantially lower than today. Because sulfate is produced mostly through weathering of the continents in the presence of oxygen, low Proterozoic concentrations imply that levels of atmospheric oxygen fell between the abundances of the Phanerozoic and the deficiencies of the Archean, which are also indicated by the Precambrian sulfur isotope record. Given the limited availability of atmospheric oxygen, deep-water anoxia may have persisted well into the Proterozoic in the presence of a growing sulfate reservoir, which promoted prevalent euxinia. Collectively, these observations suggest that the mid-Proterozoic maximum in SEDEX mineralization and the absence of Archean deposits reflect a critical threshold in the accumulation of oceanic sulfate and thus sulfide within anoxic bottom waters and pore fluids—conditions that favored both the production and preservation of sulfide mineralization at or just below the seafloor. Consistent with these evolving global conditions, the appearance of voluminous SEDEX mineralization ca. 1800 Ma coincides generally with the disappearance of banded iron formations—marking the transition from an early iron-dominated ocean to one more strongly influenced by sulfide availability. In further agreement with this conceptual model, Proterozoic SEDEX deposits in northern Australian formed from relatively oxidized fluids that required reduced conditions at the site of mineralization. By contrast, the generally more oxygenated Phanerozoic ocean may have only locally and intermittently favored the formation and preservation of exhalative mineralization, and most Phanerozoic deposits formed from reduced fluids that carried some sulfide to the site of ore precipitation.

We have evaluated the temporal distribution of Mississippi Valley-type (MVT) Zn-Pb deposits with special attention to the nature and number of deposits of Precambrian age. Our evaluation is based on the widely used model for MVT mineralization involving metal-bearing brines that lack reduced S and that deposit sulfides only where they encounter a reservoir of sulfide or where sulfate in the metal-bearing brine is reduced to sulfide. For MVT systems of this type, basins with abundant sulfate would be most favorable for development of MVT mineralization because these would allow transport of metals in sulfate-rich brines and deposition of metals in areas where the sulfate was reduced. Because abundant sulfate requires abundant atmospheric oxygen, the distribution of MVT deposits through time might reflect compositional changes in Earth's atmosphere, especially the suggested Great Oxidation Event (GOE). A compilation of new data for the Bushy Park-Pering district in the Transvaal Supergroup of South Africa, the world's oldest known MVT province, and published information on other Precambrian MVT deposits in the Ediacara, Berg Aukas/Abenab, Gayna River, Warrabarty, Nanisivik, Kamarga (Century), McArthur River (Coxco), Ramah, and Esker districts shows that they are generally similar in geologic setting and mineralogy to those in Phanerozoic rocks. Fluid inclusions in some Neo-proterozoic deposits, including Berg Aukas/Abenab, Gayna River, Warrabarty, and Nanisivik, record higher temperatures and salinities than found in most Phanerozoic deposits, possibly reflecting igneous activity or a more proximal basinal setting during Precambrian time. Fluid inclusion leachate data for several Precambrian MVT deposits suggest that their parent brines formed by evaporation of seawater, and S isotope compositions indicate that the S was derived largely from coeval seawater sulfate. Comparisons of data from all deposits show no evidence for a gradual increase in temperature or salinity backward through time, such as might be caused by higher heat flow during early stages of Earth history, although the magnitude of this effect might be lost in the uncertainty of most fluid inclusion measurements. These observations confirm that MVT deposits reflect the chemistry of their source basins, which are as old as 2.6 Ga. No MVT deposits or suitable host rocks of an older age are known. Precambrian MVT deposits do differ from their Phanerozoic analogues in the magnitude of mineralization. Precambrian deposits and districts formed at an estimated rate of 5.5 per billion years versus a significantly larger rate of ∼60 per billion years for Phanerozoic deposits, and the Phanerozoic deposits are considerably larger. Furthermore, the transition from low-magnitude, Precambrian-type to high-magnitude, Phanerozoic-type MVT mineralization took place at the beginning of Cambrian time rather than at the 2.3 Ga GOE. This appearance of widespread MVT mineralization is closer to the time at which sulfate concentrations in the world ocean are estimated to have reached present-day levels. Although these conclusions are subject to considerable uncertainty because of the limited number of Precambrian deposits, the lack of an increase in the frequency of MVT mineralization at the GOE suggests that widespread MVT mineralization requires higher levels of sulfate than could have been provided by this event, or that the appearance of sulfate in the ocean was considerably delayed. Finally, the presence of MVT deposits in basins that formed considerably before the GOE suggests that local sulfate concentrations were available at even early points in Earth's history.

Black shales and Mn carbonates interbedded with glacial deposits from the Neoproterozoic of southern China exhibit extremely heavy values of pyrite S isotopes that may reflect the peculiar environment of Earth at this time. δ 34 S averages +30‰ at Tanganshan and +44‰ at Xiangtan, compared with typical values of 0‰ to +5‰ found in younger deposits. Furthermore there is no distinction between the shales and the Mn carbonate ores in the Neoproterozoic, unlike the younger deposits, which show much lighter δ 34 S in the shales than in the Mn ores (the spread is 25‰). Most other chemical parameters are very similar to both the younger Mn deposits and those from the Paleoproterozoic. The exception is the rare earth elements (REE). All Neoproterozoic Fe ores and most Neoproterozoic Mn ores lack the positive Eu anomaly that characterizes Archean and Paleoproterozoic Fe-Mn accumulations. On the other hand, Neoproterozoic Mn deposits have positive Ce anomalies on North American Shale Composite (NASC) normalized plots, in contrast to other MnCO 3 ores. The ΣREE is also higher than in other Mn deposits, but lower than in modern deep-sea crusts. Sulfide S values in all Neoproterozoic shales tend to be exceptionally variable and to often show much heavier values than can be found in marine strata from the Phanerozoic. Therefore the anomalous δ 34 S values we observed reflect peculiar conditions in the world oceans at this time rather than purely local effects. Times of enrichment of seawater sulfate in 34 S do not correspond to periods of glaciation, so the likely cause of the S isotope patterns is not worldwide glaciation, but a generally low level of dissolved sulfate S in the Neoproterozoic oceans that allowed modest increases in the amounts of S removed as pyrite to drive down the oceanic S reservoir enough to produce strong Rayleigh reservoir effects. The abundance worldwide of Sturtian-age Mn and Fe deposits indicates an increase in Fe flux to the oceans that would have been sufficient to depress SO 4 2- levels severely and to result in residual dissolved S extremely enriched in 34 S. REE evidence indicates that most of this enhanced Fe and Mn flux came from diagenetic remobilization of detrital oxides rather than from ridge-crest hydrothermal systems, in contrast to the Paleoproterozoic banded iron formations. Rapid introduction of lateritic soil residues to restricted basins by low-latitude glaciation could have provided the needed excess Fe and Mn to drive this system.

REE and Nd isotope data indicate that most of the iron in banded iron formations is derived from hydrothermal sources but do not exclude a significant contribution from terrestrial sources, such as diagenetic recycling. A diagenetic model has been used to estimate the recycling of iron into overlying seawater, due to microbial reduction and dissolution at depth in anoxic sediment pore waters, followed by diffusion upward through a surface layer of sediment that contains oxygenated pore waters. Rates of iron recycling increase with higher pore-water dissolved iron concentrations, decreasing pH and temperature, and smaller thicknesses of the surface oxygenated layer. Iron can be recycled at rates of 1000–5000 µg cm −2 yr −1 from Proterozoic (pO 2 = PAL) pore waters with dissolved Fe 2+ = 1–5 µg cm −3 , pH 6.5 (and T < 65 °C), or pH 7.0 (and T < 40 °C), or pH 7.5 (and T < 20 °C), provided the thickness of the surface oxygenated layer is less than 0.1 cm. Lower pO 2 levels and more weakly oxygenated surface layers do not significantly increase the maximum recycling rates but enable these to be achieved at larger thicknesses of the surface layer, for all pH 6.5–7.5 and temperatures from 10 to 65 °C. Rates of iron supplied by diagenetic recycling can be substantially modified by the export efficiency (ϵ) from the source area and by the ratio (Source Area)/(Sink Area), which can either disperse or concentrate the recycling flux that is delivered to a sink area of banded iron formation. Banded iron formations that require maximum iron delivery rates of 22500 µg cm −2 yr −1 can be produced only by recycling rates of 5000 µg cm −2 yr −1 (and ϵ = 1) from a source area that is at least four times larger than the area of banded iron formation. Modern basins have ratios of shelf area (<200 m water depth) to deep basin area that commonly range from 0.2 to 4. Basins at either extreme have ratios of (Deep Basin Area)/(Shelf Area) or (Shelf Area)/(Deep Basin Area) that exceed 4 and are potentially able to concentrate iron either from a deep basin source area to banded iron formation on the shelf, or from a shelf source area to a banded iron formation depositing in the deep basin. However, these mass balance constraints require the existence of substantial areas of contemporaneous source sediments (or smaller areas of iron-enriched sediments) located either on the shelf or in the deep basin.

It is suggested that sedimentary deposition of iron in sediments from the Archean to the present day can be attributed largely to microbial mediation. Results from laboratory experiments, using a microbial consortium enriched from the source of a biofilm growing on a rock face in an underground research laboratory, are used to advance a plausible explanation for the mobilization, precipitation, transport, and deposition of iron. The consortium produces its own local environments independent of the prevailing atmosphere. It is able to extract iron from minerals such as biotite and magnetite, as well as from a chelated solution; this iron is then metabolized, mainly through dissimulatory iron reduction, to provide cell energy, after which it is immediately precipitated. All organisms require energy for growth and reproduction, but because iron redox reactions are inefficient a large amount of iron must be processed, either directly through metabolism or indirectly due to the local microbial redox microenvironment, with the result that vast quantities accumulate as waste. This precipitate could be the main source of iron in sedimentary iron formations. Siderite and ferrihydrite, the main precipitates, may occur in close juxtaposition within a biofilm. The oxidation state of the iron precipitate is controlled by the nutrient supply, which in turn influences the metabolism of the biofilm organisms and hence their redox. Subsequently, this iron, enmeshed within the biofilm, is either deposited in sediments as fine layers or rolled by wave or current action into particles and granules, which can form structures similar to those found in banded iron formations (BIFs).

High-grade BIF-hosted iron ore deposits are widely believed to have formed by epigenetic residual enrichment of hematite at the expense of other constituents, most notably chert. Processes responsible for the enrichment to high-grade iron ores are, however, only poorly understood and a range of metallogenetic models have been proposed. Field relationships have been used to distinguish three major groups of BIF-hosted high-grade iron ore deposits, namely deposits of ancient supergene, hydrothermal, and supergene-modified hydrothermal origin. Iron ores from all three deposit types are essentially composed of hematite; among different morphological types of hematite, microcystalline platy hematite and martite predominate. In this contribution, the oxygen isotope geochemistry of ore-forming hematite and martite from several high-grade iron ore deposits is examined, in an attempt to differentiate deposits of hydrothermal origin from those formed in ancient supergene environments. The δ 18 O composition of martite and microplaty hematite from deposits presumed to be of hydrothermal origin range from +0.9‰ to −7.3‰. Microcrystalline platy hematite from high-grade ores of ancient supergene origin, in contrast, has δ 18 O values ranging between +2.0‰ and −3.9‰. The latter range overlaps with the range that is defined by hematite and magnetite from weakly metamorphosed Archean–Paleoproterozoic BIF (+5‰ to −4‰). The results obtained for ancient supergene deposits developed along the 2.2 Ga Gamagara-Mapedi unconformity strengthen the argument that the Paleoproterozoic atmosphere-hydrosphere-lithosphere system was very similar to that of modern Earth. The marked shift to negative δ 18 O values displayed by hematite and martite from hydrothermal iron ore deposits, on the other hand, provides support for the suggestion that aqueous fluids of shallow crustal origin were responsible for the hydrothermal enrichment of banded iron formations to high-grade iron ores.

Rare earth element (REE) analyses of Precambrian banded iron formations (BIFs) show that distinct negative Ce anomalies, although rather weak or moderate (Ce/Ce* = 0.5–0.9), are commonly present in Algoma-type BIFs of the Early and Middle Archean, and even in the 3.8–3.7 Ga Isua iron formation (IF). This indicates that the seawater columns from which the BIFs precipitated were not entirely anoxic and that Ce oxidation mechanisms already existed in the 3.8–3.7 Ga oceans. The presence of pronounced negative Ce anomalies (Ce/Ce* = 0.1–0.5) in Late Archean (2.9–2.7 Ga) Algoma-type BIFs suggests that strongly oxygenated oceanic conditions like today emerged by 2.9–2.7 Ga. This suggestion is consistent with geologic evidence that small but widespread Mn deposits formed during Late Archean time. The Hamersley and Transvaal IFs (2.7–2.4 Ga in age) have less noticeable Ce anomalies (Ce/Ce* = 0.7–1.0). These BIFs were deposited on an evolving rift in a land-locked ocean that became anoxic due to intense hydrothermal activity. The 2.2–2.1 Ga Superior-type IFs exhibit distinct negative Ce anomalies (Ce/Ce* = 0.2–0.7), but the ca. 1.9 Ga IFs and the ca. 0.7 Ga IFs have less distinct Ce anomalies. These variations in Ce/Ce* values of the post-2.7 Ga BIFs may reflect the episodicity in global mantle plume activity that created locally anoxic basins. The Archean Algoma-type BIFs have distinctly positive but variable Eu anomalies (Eu/Eu* = 0.8–7). Their strong positive Eu anomalies suggest large contributions of hydrothermal fluids to the seawater involved in BIF precipitation. The large variation in Eu anomalies in Algoma-type BIFs reflects the large variation in mixing ratios of hydrothermal fluids and ambient seawaters at various depositional sites. However, the post-2.7 Ga Superior-type BIFs exhibit much lower and constant Eu/Eu* values (0.7–1.9). This implies that the REE chemistry of the basin water that hosted voluminous Superior-type BIFs was influenced by a riverine influx from the surrounding continent that grew rapidly due to global mantle plume activity, besides the intense hydrothermal influx.

Geological and geochemical characteristics of banded iron formations (BIFs) suggest that they formed by mixing locally (or regionally) discharged submarine hydrothermal fluids with local seawater, rather than by upwelling deep ocean water. Submarine hydrothermal fluids typically evolved from local seawater by acquiring heat, metals, and sulfur during deep circulation through a variety of rocks (e.g., volcanics, evaporites) in greenstone terranes that developed under a variety of tectonic settings. In general, when the fluids were heated above ∼350 °C, they may have produced Cu- and Zn-rich volcanogenic massive sulfide deposits (VMSDs), whereas those heated less than ∼200 °C were generally poor in H 2 S and heavy metals, except Fe, and may have subsequently produced BIFs. Depending on the salinity contrast between discharging hydrothermal fluids (evolved seawater) and local seawater, hydrothermal fluids may (1) mix rapidly with local seawater to form smoker-type BIFs or (2) create a metal- and silica-rich brine pool, mix slowly with the overlying water body, and form brine pool-type BIFs. BIFs associated with VMSDs and volcanic rocks generally belong to smoker-type BIFs; many formed at seawater depths >2.5 km. Large BIFs, including the 2.6–2.4 Ga BIFs in the Hamersley Basin, Australia, the 2.5 Ga Kuruman IF in South Africa, and the 1.87 Ga BIFs in the Lake Superior region, United States-Canada, belong to brine pool-type BIFs. The Hamersley Basin and possibly other large BIF-hosting basins were probably land-locked seas (like the Black Sea) where river waters diluted the surface water zone and the underlying water bodies were anoxic. During the accumulation of a BIF sequence, the dominant Fe mineralogy frequently changed from ferric (hydr)oxides (oxide BIFs) to siderite (carbonate BIFs) and to pyrite (sulfide BIFs). Such changes were probably caused by changes in the relative amounts of dissolved O 2 (DO), ΣCO 3 2− , and ΣS 2− in local seawater. From the Fe 2+ -O 2 mass balance calculations for the formation of iron oxides in smoker-type BIFs, and the relationship between the atmospheric pO 2 and oceanic O 2 depth profile, we conclude that the atmosphere and oceans have been fully oxygenated since ca. 3.8 Ga, except in local anoxic basins. Thermodynamic analyses of the formational conditions of siderite and analyses of the carbon isotopic composition of siderite associated with major BIFs suggest that the pre–1.8 Ga atmosphere was CO 2 -rich (pCO 2 >100 PAL) and CH 4 -poor (pCH 4 ≈ 10 ppm); therefore, CO 2 , rather than CH 4 , was the major greenhouse gas throughout geologic history. After a decline of hydrothermal fluid flux, BIF-hosting basins generally became euxinic (H 2 S-rich) because of the increased activity of sulfate-reducing bacteria (SRB) and SO 4 2− -rich seawater, and thereby accumulated organic carbon-rich and pyrite-rich black shales (sulfide-type BIFs). The SO 4 2− contents and SRB activity in the oceans have been essentially the same since ca. 3.8 Ga. The Archean oceans were most likely poor in both Fe 2+ and silica, much like modern oceans. Our study also suggests that diverse communities of organisms, including cyanobacteria, SRB, methanogens, methanotrophs, and eukaryotes, evolved very early in Earth's history, probably by the time the oldest BIFs (ca. 3.8 Ga) formed. BIFs have been found in rocks of all geologic age. Therefore, they cannot be indicators of an anoxic atmosphere and/or anoxic oceans as suggested by many previous researchers. Instead, BIFs indicate that the atmosphere and ocean chemistry have been regulated at present compositions (except pCO 2 ) through geologic history by interactions with the biosphere. The general trend of declining size and abundance of BIFs with geologic time reflects the cooling history of Earth's interior.