Isotopic Constraints on Biogeochemical Cycling of Fe

INTRODUCTION Cycling of redox-sensitive elements such as Fe is affected by not only ambient Eh-pH conditions, but also by a signifi cant biomass that may derive energy through changes in redox state (e. The evidence now seems overwhelming that biological processing of redox-sensitive metals is likely to be the rule in surface-and near-surface environments, rather than the exception. The Fe redox cycle of the Earth fundamentally begins with tectonic processes, where " juvenile " crust (high-temperature metamorphic and igneous rocks) that contains Fe which is largely in the divalent state is continuously exposed on the surface. If the surface is oxidizing, which is likely for the Earth over at least the last two billion years (e.g., Holland 1984), exposure of large quantities of Fe(II) at the surface represents a tremendous redox disequilibrium. Oxidation of Fe(II) early in Earth's history may have occurred through increases in ambient O 2 contents through photosynthesis (e. Iron oxides produced by oxidation of Fe(II) represent an important sink for Fe released by terrestrial weathering processes, which will generally be quite reactive. In turn, dissimilatory microbial reduction of ferric oxides, coupled to oxidation of organic carbon and/or H 2 , is an important process by which Fe(III) is reduced in both modern and ancient sedimentary environments 360 1998), together with a wealth of geochemical information, suggests that microbial Fe(III) reduction may have been one of the earliest forms of respiration on Earth. It therefore seems inescapable that biological redox cycling of Fe has occurred for at least several billion years of Earth's history. Signifi cant Fe isotope variations in nature are generally restricted to relatively low-temperature systems, including hydrothermal fl uids and chemically precipitated minerals metabolic processing of Fe have shown that measurable Fe isotope fractionations are produced during dissimilatory Fe(III) reduction by bacteria (Beard et al. 1999, 2003a; Icopini et al. 2004; Johnson et al. 2004a), as well as anaerobic photosynthetic Fe(II) oxidation (Croal et al. 2004). In addition, the role of organic ligands in promoting mineral dissolution has been investigated in experiments (Brantley et al. 2001, 2004). Iron isotopes may also be fractionated in abiologic systems, including ion-exchange chromatography (Anbar et al. 2000; Roe et al. 2003), abiotic precipitation of ferric oxides or oxyhydroxides (Bullen et al. 2001; Skulan et al. 2002), and sorption of aqueous Fe(II) to ferric hydroxides (Icopini et al. 2004). The largest abiotic fractionations in experiment have been measured between …