Description: © The Author(s), 2017. This is the author's version of the work. It is posted here under a nonexclusive, irrevocable, paid-up, worldwide license granted to WHOI. It is made available for personal use, not for redistribution. The definitive version was published in Earth-Science Reviews 172 (2017): 140-177, doi:10.1016/j.earscirev.2017.06.012.

Description: Iron formations (IF) represent an iron-rich rock type that typifies many Archaean and Proterozoic supracrustal successions and are chemical archives of Precambrian seawater chemistry and postdepositional iron cycling. Given that IF accumulated on the seafloor for over two billion years of Earth’s early history, changes in their chemical, mineralogical, and isotopic compositions offer a unique glimpse into environmental changes that took place on the evolving Earth. Perhaps one of the most significant events was the transition from an anoxic planet to one where oxygen was persistently present within the marine water column and atmosphere. Linked to this progressive global oxygenation was the evolution of aerobic microbial metabolisms that fundamentally influenced continental weathering processes, the supply of nutrients to the oceans, and, ultimately, diversification of the biosphere and complex life forms. Many of the key recent innovations in understanding IF genesis are linked to geobiology, since biologically assisted Fe(II) oxidation, either directly through photoferrotrophy, or indirectly through oxygenic photosynthesis, provides a process for IF deposition from mineral precursors. The abundance and isotope composition of Fe(II)-bearing minerals in IF additionally suggests microbial Fe(III) reduction, a metabolism that is deeply rooted in the Archaea and Bacteria. Linkages among geobiology, hydrothermal systems, and deposition of IF have been traditionally overlooked, but now form a coherent model for this unique rock type. This paper reviews the defining features of IF and their distribution through the Neoarchaean and Palaeoproterozoic. This paper is an update of previous reviews by Bekker et al. (2010, 2014) 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.

Description: 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 for providing research funding.

Description: Precambrian Si-rich sedimentary rocks, including cherts and banded iron formations (BIFs), record a 〉7‰ spread in 30Si/28Si ratios (δ30Si values), yet interpretation of this large variability has been hindered by the paucity of data on Si isotope exchange kinetics and equilibrium fractionation factors in systems that are pertinent to Precambrian marine conditions. Using the three-isotope method and an enriched 29Si tracer, a series of experiments were conducted to constrain Si isotope exchange kinetics and fractionation factors between amorphous Fe(III)–Si gel, a likely precursor to Precambrian jaspers and BIFs, and aqueous Si in artificial Archean seawater under anoxic conditions. Experiments were conducted at room temperature, and in the presence and absence of aqueous Fe(II) (Fe(II)aq). Results of this study demonstrate that Si solubility is significantly lower for Fe–Si gel than that of amorphous Si, indicating that seawater Si concentrations in the Precambrian may have been lower than previous estimates. The experiments reached ∼70–90% Si isotope exchange after a period of 53–126 days, and the highest extents of exchange were obtained where Fe(II)aq was present, suggesting that Fe(II)–Fe(III) electron-transfer and atom-exchange reactions catalyze Si isotope exchange through breakage of Fe–Si bonds. All experiments except one showed little change in the instantaneous solid–aqueous Si isotope fractionation factor with time, allowing extraction of equilibrium Si isotope fractionation factors through extrapolation to 100% isotope exchange. The equilibrium 30Si/28Si fractionation between Fe(III)–Si gel and aqueous Si (Δ30Sigel–aqueous) is −2.30 ± 0.25‰ (2σ) in the absence of Fe(II)aq. In the case where Fe(II)aq was present, which resulted in addition of ∼10% Fe(II) in the final solid, creating a mixed Fe(II)–Fe(III) Si gel, the equilibrium fractionation between Fe(II)–Fe(III)–Si gel and aqueous Si (Δ30Sigel–aqueous) is −3.23 ± 0.37‰ (2σ). Equilibrium Si isotope fractionation for Fe–Si gel systems is significantly larger in magnitude than estimates of a near-zero solid–aqueous fractionation factor between pure Si gel and aqueous Si, indicating a major influence of Fe atoms on Si–O bonds, and hence the isotopic properties, of Fe–Si gel. Larger Si isotope fractionation in the Fe(II)-bearing systems may be caused by incorporation of Fe(II) into the solid structure, which may further weaken Fe–Si bonds and thus change the Si isotope fractionation factor. The relatively large Si isotope fractionation for Fe–Si gel, relative to pure Si gel, provides a new explanation for the observed contrast in δ30Si values in the Precambrian BIFs and cherts, as well as an explanation for the relatively negative δ30Si values in BIFs, in contrast to previous proposals that the more negative δ30Si values in BIFs reflect hydrothermal sources of Si or sorption to Fe oxides/hydroxides.

Description: Microbial dissimilatory iron reduction (DIR) is a deeply rooted metabolism in the Bacteria and Archaea. In the Archean and Proterozoic, the most likely electron acceptor for DIR in marine environments was Fe(III)–Si gels. It has been recently suggested that the Fe and Si cycles were coupled through sorption of aqueous Si to iron oxides/hydroxides, and through release of Si during DIR. Evidence for the close association of the Fe and Si cycles comes from banded iron formations (BIFs), which consist of alternating bands of Fe-bearing minerals and quartz (chert). Although there has been extensive study of the stable Fe isotope fractionations produced by DIR of Fe(III)–Si gels, as well as studies of stable Fe isotope fractionations in analogous abiologic systems, no studies to date have investigated stable Si isotope fractionations produced by DIR. In this study, the stable Si isotope fractionations produced by microbial reduction of Fe(III)–Si gels were investigated in simulated artificial Archean seawater (AAS), using the marine iron-reducing bacterium Desulfuromonas acetoxidans. Microbial reduction produced very large 30Si/28Si isotope fractionations between the solid and aqueous phase at ∼23 °C, where Δ30Sisolid–aqueous isotope fractionations of −3.35 ± 0.16‰ and −3.46 ± 0.09‰ were produced in two replicate experiments at 32% Fe(III) reduction (solid-phase Fe(II)/FeTotal = 0.32). This isotopic fractionation was substantially greater than that observed in two abiologic controls that had solid-phase Fe(II)/FeTotal = 0.02–0.03, which produced Δ30Sisolid–aqueous isotope fractionations of −2.83 ± 0.24‰ and −2.65 ± 0.28‰. In a companion study, the equilibrium Δ30Sisolid–aqueous isotope fractionation was determined to be −2.3‰ for solid-phase Fe(II)/FeTotal = 0. Collectively, these results highlight the importance of Fe(II) in Fe–Si gels in producing large changes in Si isotope fractionations. These results suggest that DIR should produce highly negative δ30Si values in quartz that is the product of diagenetic reactions associated with Fe–Si gels. Such Si isotope compositions would be expected to be associated with Fe-bearing minerals that contain Fe(II), indicative of reduction, such as magnetite. Support for this model comes from recent in situ Si isotope studies of oxide-facies BIFs, where quartz in magnetite-rich samples have significantly more negative δ30Si values than quartz in hematite-rich samples.