Banded iron formations (BIF) are the protolith to most of the world’s largest iron ore deposits. Previous hypogene genetic models for Paleoproterozoic “Lake Superior” BIF-hosted deposits invoke upwards, down-temperature flow of basinal brines via complex silica and carbonate precipitation/dissolution processes. Such models are challenged by the necessary SiO2 removal. Thermodynamic and mass balance constraints are used to refine conceptual models of the formation of BIF-hosted iron ore. These constraints, plus existing isotope and halogen ratio evidence, are consistent with removal of silica by down- or up-directed infiltration of high-pH hypersaline brines, with or without a contribution from basinal brines. The proposed link to surface environments suggest that Paleoproterozoic BIF-ore upgrade may provide a record of a critical time in the evolution of the Earth’s biosphere and hydrosphere.


Banded iron formations (BIFs) record changing environmental conditions and are a precursor to the world’s largest iron ore deposits. Vast volumes of BIFs were deposited on passive margins from ca. 2.6 Ga to the Great Oxygenation Event (GOE) at ca. 2.4 Ga (e.g., Beukes and Gutzmer, 2008). After 1.85 Ga, formation of iron formations essentially ceased until a restricted resurgence in the late Neoproterozoic (Young, 1976; Klein and Beukes, 1993) and more recent minor occurrences.

The temporal distribution of BIFs records a complex interplay between a cooling Earth and changes in mantle plume events, continental growth and tectonics, evolution of the biosphere and an increased flux of iron to the hydrosphere, which in turn had a fundamental control on the oxygen contents of the hydrosphere and redox state of the oceans (Isley and Abbott, 1999; Holland, 2005; Bekker et al., 2010). In most giant Paleoproterozoic BIF-hosted iron ore deposits (∼35 wt% Fe) the formation of high-grade (>58 wt% Fe) iron ore from Lake Superior–type BIF is thought to have occurred post 2.2 Ga, after the GOE (Taylor et al., 2001; Rasmussen et al., 2007; Thorne et al., 2009), so the BIF-ore upgrade provides a record of the changing chemistry of the hydrosphere and atmosphere in the Paleoproterozoic.

In the Hamersley province, Western Australia, the BIF-ore upgrade is proposed to involve basinal brines, meteoric fluids, and supergene enrichment (e.g., Morris et al., 1980; Barley et al., 1999; Taylor et al., 2001; Thorne et al., 2004, 2008). Silica removal is proposed to occur by upward, down-temperature flow of basinal brines (e.g., Thorne et al., 2004; Gutzmer et al., 2006; Thorne et al., 2008), and it is this stage of the BIF-ore upgrade that is explored in this paper.

The challenge for models that involve removal of SiO2 from BIF by down-temperature flow is that quartz solubility decreases with decreasing temperature (e.g., Manning, 1994), so fluid that moves down-temperature is a poor agent for quartz removal, unless large volumes of fluids enter the BIF in a silica-undersaturated state. Here, we combine numerical constraints with petrological, geochemical and isotopic characteristics of BIF-hosted iron ore from the Hamersley province, Western Australia, to constrain the conditions associated with silica loss and carbonate precipitation, and discuss the findings in the context of the temporal evolution of the world’s atmosphere and oceans.


The mineralogy of BIF, hydrothermally altered BIF, and iron ore varies (e.g., Thorne et al., 2004; Rosiere et al., 2008; Mukhopadhyay et al., 2008; Angerer and Hagemann, 2010) but some features are sufficiently common to allow generalization. BIF protolith (Figs. 1A, 1B1, and 1B2) consists of magnetite- and chert-rich bands, iron silicates ± diagenetic carbonates. Initial hydrothermal alteration (Figs. 1A, 1B3, and 1B4), which may be gradational, involves conversion of some of the silicates to carbonates. Subsequently, chert and quartz are replaced by iron-bearing carbonate with, in some cases, conversion of magnetite to hematite, often with no apparent change in volume. Reaction fronts may be gradual (millimeter to decimeter scale) or knife-sharp on the scale of individual bands. Quartz may be precipitated locally in fault zones within shear veins (e.g., Hagemann et al., 1999; Thorne et al., 2010).

The carbonate-bearing alteration assemblage (Figs. 1A, 1B5, and 1B6) is separated from almost pure (hypogene) hematite iron ore (Figs. 1A, 1B7, and 1B8) by a sharp or gradational reaction front. Carbonate loss is thought to occur without iron mobility on a scale greater than a few cm (Taylor et al., 2001). Carbonate dissolution continues in the weathering environment.

Upgrade of BIF in the Hamersley province, Western Australia, is thought to have occurred from 2.15 Ga onwards during the waning stages of the Opthalmian orogeny (Rasmussen et al., 2007). There is evidence for extensional faulting and orogenic collapse at the proposed time of ore upgrade (Müller et al., 2005) and it has been suggested that topographic relief drove circulation of surface-derived waters through the fault system at this time (e.g., Hagemann et al., 1999; Oliver and Dickens, 1999).

Constraints on ore upgrade in the Hamersley are provided by fluid inclusion, isotopic, and halide concentration measurements. Fluid inclusions in carbonates are hypersaline, up to 24 wt% CaCl2 equivalent, which were trapped at temperatures between 150 and 400 °C (Thorne et al., 2004; Brown et al., 2004), mainly at the lower end of this range. Quartz-hosted fluid inclusions (e.g., Brown et al., 2004) records periods of quartz growth, not quartz removal, so such inclusions may record either a different event, or a later stage of the silica dissolution event that is of interest here.

Iron oxide oxygen isotopes in unaltered BIF have δ18OVSMOW (Vienna Standard Mean Ocean Water) between 4‰ and 13‰, whereas hematite and magnetite in altered and mineralized rocks have δ18OVSMOW between −9‰ and −2.9‰ (Thorne et al., 2009). Carbon isotopes of carbonate minerals lie between −10‰ and 0‰, relative to Vienna Peedee belemnite (VPDB). The lower values are typical of unaltered BIF. Values in altered rocks could record equilbrium with either Paleoproterozoic ocean water or dolomite in the underlying Wittenoom Formation. Fluid inclusion Na/Br and Cl/Br ratios record overlapping populations of fluids: (1) seawater that has evaporated to halite saturation; and (2) meteoric waters that interacted with evaporites (Thorne et al., 2010).


Fluids in Equilibrium with BIF

Quartz solubility is sensitive to pressure, temperature (e.g., Manning, 1994), pH (Busey and Mesmer, 1977), and salinity (e.g., Shmulovich et al., 2006). Quartz and carbonate solubilities were calculated as a function of pressure, temperature and salt content (Fig. 2) using the methods outlined in the GSA Data Repository1. Calcite, rather than iron carbonate was used for the models because data for iron carbonates is sparse, but similar trends for iron carbonates are expected. Temperature is the first order control on quartz solubility, whereas calcite solubility depends strongly on pressure (Fig. 2A). Salinity-driven mineral precipitation/dissolution is minor for geothermal gradients <50 °C km−1 (Fig. 2B). There is no geothermal gradient for which silica dissolution is accompanied by carbonate precipitation in the observed quantities, so it is necessary to consider infiltration of out-of-equilibrium fluids.

Fluid out of Equilibrium with BIF

Infiltration of high-pH, out-of-equilibrium fluid can drive quartz removal via up- or down-temperature flow, because quartz is 4 orders of magnitude more soluble at pH 9 than it is at pH 6 (Busey and Mesmer, 1977). The desilicification/carbonation and oxidation fronts observed in BIFs are sufficiently sharp that they can be treated as advective chromatographic fronts; broadening by diffusion, dispersion and kinetic broadening can be neglected. The ratio of the distance traveled by a fluid to that of associated reaction front is given by 
(adapted from Evans et al., 2003). d is the position of the front relative to the infiltration horizon in meters, ci is the concentration of the reactant of interest in moles m−3, and θ is porosity.

Relative positions of carbonation, desilicification, and oxidation fronts were calculated using the values given in Table DR1 (in the Data Repository), which are based on the assumption of infiltration an out-of-equilibrium, pH 9, SiO2-free, CO2-rich, O2-bearing fluid (O2 at 10% of present-day values). Front positions were normalized to that of the most advanced carbonation front (Fig. 1B). The oxidation front is predicted to proceed the smallest distance, as observed, though the position of this front is modified by post-hypogene meteoric fluid and supergene processes. The carbonation front is predicted to advance further than the desilicification front for any reasonable combination of input parameters. In reality, the desilicification front coincides with a sharp increase in carbonate content, which suggests a reaction-induced porosity control on fluid infiltration.

Volumes of Fluid Flow

The relationship between time-integrated fluid flux and silica removal for hydrothermal alteration where pressure, temperature and fluid composition gradients are fixed, and for constant fluid flow rates is given by 
(c.f. Evans and Bickle, 1999, their equation 6). Q is the time-integrated fluid flux, in cubic meters of fluid per square meter of rock, ΔnSiO2 is the SiO2 removed from the volume (moles m−3), cSiO2 is the concentration of SiO2 in the fluid in (moles m−3), and l is distance along the flow path (m). This expression is based on mass balance and can be applied to in- or out-of-equilibrium situations. Calculations were made for flow up- or down-temperature of a pH 9, silica-undersaturated fluid, such as might be derived from an ultramafic lithology or hypersaline evaporite-derived fluid. pH is assumed to drop from 9 to 5 during equilibration. dcSiO2/dpH was calculated from Busey and Mesmer (1977) and checked against HCh output (Shvarov and Bastrakov, 1999).

Flow volumes were calculated for a single cubic meter of rock undergoing pH-driven quartz dissolution and scaled to estimate the volume needed to produce the observed SiO2 depletion on a scale comparable to observed small iron ore deposits (3 × 108 m3). The volume of fluid required is approximately ∼109 m3.


The observed silica removal and carbonate precipitation needed for formation of a kilometer-scale BIF-hosted Fe deposit is inconsistent with in-equilibrium fluid flow (Fig. 2), but could be produced by pervasive infiltration of ∼109 m3 of a high pH (pH > 9) fluid flowing up- or down- temperature.

Density- or topography-driven infiltration (McLellan et al., 2004) of dense hypersaline brines evolved on continental margins (Fig. 3) provides a plausible source of high-pH fluid. The pH of hypersaline lakes today is up to 10, and the CO2 content is high; reported alkalinities exceed 10,000 mg L−1 (Gosselin et al., 1994). These brines form on marginal platforms (Leach et al., 2010) and were present on the north Yilgarn margin in the Palaeoproterozoic (El Tabakh et al., 1999) and in Mesoproterozoic dolomitic argillites in the western United States (Gonzalez-Alvarez and Kerrich, 2011). The Yilgarn and Pilbara cratons may not have been adjacent at the time of ore formation, and the latitude of the Pilbara margin at that time is poorly known but existing constraints place the latitude within the range of those required for evaporite formation (Li, 2000). Fault zones in similar environments today are sufficiently permeable that proposed volumes could have flowed on a geologically reasonable timescale (Jones et al., 2002).

Carbon and oxygen isotopes of early carbonate and iron oxides in mineralized BIF are consistent with seawater derivation from Palaeoproterozoic seawater. High salinities and Na/Br and Cl/Br ratios record fluids typical of seawater that has evaporated to halite saturation (Thorne et al., 2010), which is also consistent with an involvement of evaporite-derived fluids. Temperatures of up to 250 °C are consistent with fluid inclusion evidence, and with other environments where voluminous quantities of surface-derived fluids infiltrate basement rocks (e.g., Gleeson and Yardley., 2003). However, temperatures >300 °C recorded by early assemblages at Mount Tom Price require a separate high-temperature alteration stage. SiO2 deposition in outflow zones from the flow system described have not been recorded, to date, but may prove difficult to recognize given that chert was a common sedimentary rock at this time.

Some geometric constraints in mineralized Pilbara BIFs favour an upwards flowing fluid (e.g., Dalstra and Rosiere, 2004). The Wittenoom Formation underlies many of the ore-bearing units, and has been proposed as a way for basinal brines to access the BIF (e.g., Taylor et al., 2001). Dolomite in the Wittenoom Formation contains numerous shale bands (Davy, 1975), and is separated from the BIF by the Si-bearing Mount McRae shale and Sylvia Formation, so fluids from this unit would be at, or close to, SiO2 saturation, unless fluids were focused solely within shale-free regions of the dolomite, such as the Paraburdoo Member (Thorne and Tyler, 1997). If this were the case then the dolomite in the Wittenoom Formation could have acted as an aquifer that supplied overlying iron formations with fluids derived either from faults that focused either down-flowing evaporitic brines with or without a contribution from upward-flowing basinal brines, as suggested by previous workers (e.g., Taylor et al., 2001).

Mafic and ultramafic rocks are present in the Pilbara (e.g., Barnes and Hoatson, 1994) and can equilibrate with water to provide high pH fluids, but such fluids are unlikely to have reached the BIF without interacting with Si-bearing rocks such as the Mount McRae shale.

Oxidized, sulfur-bearing, hypersaline brines produced on basin margins since 2.0 Ga have been linked to base metal sulfide deposition, including MVT (Mississippi Valley type) and SEDEX-Pb-Zn deposits (Leach et al., 2010). Such deposits older than 2.02 Ga have not been found; Archaean and Paleoproterozoic oceans were reduced and largely sulfur-depleted (e.g., Farquhar et al., 2010), so marginal hypersaline brines at 2.2 Ga, at the time of ore upgrade, could have had a significantly different composition to those today. It is interesting to speculate whether redox controls related to mobilization, transport, and deposition of Pb and Zn by Palaeoproterozoic evaporite-derived brines, such as a lack of sulfur (Leach et al., 2010) lack of oxidized source rocks and aquifers, or reduction of sulfate via rock-buffering in the immediate sub-surface, may have prevented formation of these deposit types prior to 2.02 Ga.

We thank Rob Kerrich, Hilke Dalstra, and an anonymous reviewer for perceptive, knowledgeable reviews. Thanks for funding and support are due to a Curtin Research and Teaching Fellowship, Australian Research Council Discovery DP1090475 to Evans, and Core to Crust Fluid Systems funding to McCuaig and Hagemann. This is TiGeR publication number 423.

1GSA Data Repository item 2013024, information on the quartz solubility model and details of values used for calculations, is available online at www.geosociety.org/pubs/ft2013.htm, or on request from editing@geosociety.org or Documents Secretary, GSA, P.O. Box 9140, Boulder, CO 80301, USA.