Abstract:
Cycling of redox-sensitive elements such as Fe is affected by not only ambient Eh-pH conditions, but also by a significant biomass that may derive energy through changes in redox state (e.g., Nealson 1983; Lovely et al. 1987; Myers and Nealson 1988; Ghiorse 1989). 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 O2 contents through photosynthesis (e.g., Cloud 1965, 1968), UV-photo oxidation (e.g., Braterman and Cairns-Smith 1987), or anaerobic photosynthetic Fe(II) oxidation (e.g., Hartman 1984; Widdel et al. 1993; Ehrenreich and Widdel 1994). 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 H2, is an important process by which Fe(III) is reduced in both modern and ancient sedimentary environments (Lovley 1991; Nealson and Saffarini 1994). Recent microbiological evidence (Vargas et al. 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.
