(5) Understanding the Mechanisms for Precambrian Banded Iron Formation
Banded iron formations (BIF) are iron-rich (20-40% Fe) and siliceous (40-50% SiO2) sedimentary deposits that precipitated throughout much of the Precambrian (see Bekker et al., 2010 ; Bekker et al., 2014 ; Konhauser et al., 2017 for reviews). They are characteristically laminated, with alternating Fe-rich and Si-rich layers that can be observed on a wide range of scales. The mineralogy of the least metamorphosed BIF consists of chert, magnetite, hematite, carbonates (siderite, dolomite-ankerite), greenalite, stilpnomelane and riebeckite, but it is generally agreed that these minerals reflect both diagenetic and metamorphic overprinting: the primary iron minerals were most likely ferric hydroxide (e.g., ferrihydrite), ferric-silica gel, amorphous silica, and siderite.
(Left Figure) Detailed image of the mesobanding and microbanding between the chert and hematite layers in BIF. (Right Figure) Overview of the BHP Iron Ore Mine at Newman, Western Australia.
BIF Depositional Environments
Over the past decade, my research group, along with various colleagues, have studied a number BIF in an attempt to shed new light on the depositional environment of BIF. This is important if we are to use the composition of BIF to make inferences about bulk seawater composition, the evolution of marine hydrothermal systems, and/or the magnitude of continental weathering. For instance, based on rare earth element (REE) composition of iron formations, it is generally accepted that deep-sea hydrothermal processes are the most likely source of Fe. Europium (Eu) anomalies and neodymium (Nd) isotope ratios have been central in the use of REE to trace Fe sources because an Eu enrichment and positive 143Nd/144Nd ratio in chemical sedimentary rocks precipitated from seawater indicates a strong influence of hydrothermal fluids on the seawater dissolved REE load. In this regard, our own analyses of REE in the 3.75 Ga Nuvvuagittuq Supracrustal Belt in northern Québec (Mloszewska et al., 2012; Mloszewska et al., 2013), the 2.9 Ga Itilliarsguk Supracrustal Belt in West Greenland (Haugaard et al., 2012), the 2.85 and 2.62 BIFs from the Slave craton in the Norwest Territories (Haugaard et al., 2016; Haugaard et al., 2017), the 2.48 Ga Dales Gorge Member in Western Australia (Pecoits et al., 2009), and the 2.3-2.2 Ga Yuanjiacun BIF in China (Wang et al., 2015) all appear to confirm the hydrothermal sourcing of Fe to those BIF depositional basins. By contrast, REE patterns in the 2.46 Ga Joffre Member in Western Australia reveals that small proportions of pelagic ash particles have a major impact on the overall REE signature and the total REE content of the seawater (Haugaard et al., 2016), while Wang et al. (2016) suggested that periodically surface seawater from which the Yuanjiacun BIF precipitated obtained its REY and Nd signatures from weathering of nearby landmasses. Collectively, these results demonstrate that BIF were differentially influenced by hydrothermal, terrigenous and volcanic inputs through time, and no one universal model for BIF depositional origins holds true.
For those BIF where Fe(II) was hydrothermally-sourced, the proximity of the vents to the site of deposition also remains unclear. On the one hand, Fe could have been delivered from the deep ocean to the outer continental shelf by upwelling currents from a mid-ocean ridge system. Accordingly, BIF would sediment from below the wave base (without influence of wave- and storm-induced currents) onto partially submerged platforms of the continental shelves. On the other hand, the direct supply of Fe(II) into the photic zone by hydrothermal plumes associated with shallow seamount type systems would curtail the difficulties introduced by the high upwelling rates needed to bring sufficient iron from the deep sea onto the continental shelf (Konhauser, 2007a). Similarly, the source of the silica in BIF is enigmatic. While a former PhD student at the University of Hannover (Grit Steinhofel) used Si isotopes in the 2.5 Ga Dales Gorge Member BIF (Hamersley, Australia) formation to show that silica was hydrothermally sourced (Steinhoefel et al., 2010), another PhD student of mine from Leeds (Tristan Hamade) used Ge/Si ratios to show that weathering of landmass was the predominant source of silica in cherty and varved mesobands (Hamade et al., 2003). However, when ratios were measured in Fe-rich mesobands, they reflected a hydrothermal input. This confirms that the sources of silica and iron were decoupled during BIF deposition. Moreover, significant variations between different mesobands (representing tens of thousands of years) indicate that the ratios were not at steady state during the Precambrian.
Precipitation of BIF Minerals
Irrespective of depositional setting, the mineralogy of BIF dictates that some oxidation of Fe(II) was necessary for their formation, yet which mechanism(s) dominated is uncertain. Prior to the rise of atmospheric oxygen and the development of a protective ozone layer, the Earth’s surface was subjected to high levels of ultraviolet radiation. Bulk ocean waters that were anoxic at this time could have supported high concentrations of dissolved Fe(II). Under such conditions, dissolved ferrous iron species, such as Fe2+ or Fe(OH)+, absorb radiation in the 200-400 nm range, leading to the formation of dissolved ferric iron [reaction 1], which in turn, hydrolyses to form ferric hydroxide at circumneutral pH.
 2Fe2+(aq) + 2H+ + hv --> 2Fe3+(aq) + H2
However, these experiments focused on determining the specific rates of Fe(II) photochemical oxidation, and did not simulate the complex, disequilibrium water chemistry characteristic of an ocean where Fe(II)-rich hydrothermal waters reacted with ambient Si-saturated seawater that also contained high concentrations of HCO3-. Indeed, in fluids with high dissolved Fe(II), Si(OH)4 and HCO3-, the oxidation effects of either UVA or UVC were found to be negligible compared to the precipitation of ferrous-iron-silicates (Konhauser et al., 2007a).
More recently, my former PhD student (Ernesto Pecoits, now a Professor at the Uruguay University of Technology) examined the possibility of photochemical reactions occurring in the atmosphere that might generate hydrogen peroxide (H2O2), a recognized strong oxidant for ferrous iron. We modelled the amount of H2O2 that could be produced in an Eoarchean atmosphere using updated solar fluxes and CO2 mixing ratios, and demonstrated that insufficient levels of H2O2 were likely generated to account for IF deposition (Pecoits et al., 2015) Nonetheless, the H2O2 levels produced at that time could still have induced oxidative continental weathering and the biological development of enzymes (catalases and peroxidases) capable of protecting organisms against such oxidants.
As an alternative to the abiological models, the presence of ferric iron minerals in BIF has also been ascribed to the metabolic activity of planktonic bacteria in the oceans’ photic zone (for reviews see Koehler et al., 2010 and Posth et al., 2013a, both former PhD students at the University of Tubingen for reviews). Chemical oxidation of Fe(II) by photosynthetically produced O2 is one possibility, allowing for the indirect biogenic precipitation of ferric hydroxide. Under an anoxic atmosphere, this O2 could have been confined to localized "oxygen oases” associated with cyanobacterial blooms in coastal settings. Once oxygen was present, aerobic chemolithoautotrophic Fe(II) oxidizers, such as Gallionella sp., could also have precipitated biogenic Fe minerals. It is also possible that light, not O2, may have coupled the carbon and iron cycles, via photosynthesis that used Fe(II) rather than H2O as an electron donor, producing Fe(III) rather than O2 [reaction 2].
 4Fe2+ + HCO3- + 10H2O --> 4Fe(OH)3 + CH2O + 7H+
With the evolution of this photoferrotrophy, biological Fe(II) oxidation would have superseded photochemical oxidation as the bacteria could grow deeper in the water column where UV radiation would be effectively absorbed (Konhauser et al., 2002). Modelling of experimentally determined photosynthetic Fe(II) oxidation rates even suggests that such microorganisms could have accounted for all of the Fe(III) initially deposited in primary BIF sediment (Kappler et al., 2005). It has also been shown that photoferrotrophy is not hindered by the high concentrations of silica (2 mM) estimated for the Archean oceans (Konhauser et al., 2007b). Indeed, the temperature dependence of anoxygenic Fe(II) oxidation in growth medium containing dissolved Fe(II) and dissolved silica suggests the potential for a coupled abiotic-biotic mechanism for the deposition of the alternating Fe-rich and Si-rich BIF layering (Posth et al., 2008). Although, this was demonstrated in a simplified laboratory system, these experiments suggest that at temperatures suitable for Fe(II)-oxidizing phototrophs, the ferric hydroxides they produce will precipitate, while silica largely stays in solution. At cooler temperatures, silica will precipitate out abiotically while the bacteria cease measureable Fe(II) oxidation until temperatures once again increase.
(Left Figures) Summary of Fe(II) oxidation processes potentially involved in BIF deposition. (A) Abiotic photooxidation of hydrothermal Fe(II) by UV light with sedimentation of ferric iron minerals. (B) Chemical oxidation of hydrothermal Fe(II) by cyanobacterial produced O2 and Fe(II) oxidation by aerobic Fe(II)-oxidising, chemolithoautotrophic bacteria with sedimentation of either iron minerals only (chemical oxidation) or sedimentation of joint biomass and iron minerals. (C) Direct biological oxidation of Fe(II) by anoxygenic phototrophic Fe(II)-oxidising bacteria (photoferrotrophy) with joint sedimentation of biomass and iron minerals. From Koehler et al. (2010). (Right Figure) Possible deposition of alternating iron and silicate mineral layers triggered by temperature fluctuations in the ocean water - based on Posth et al. (2008). (1) and (3) Moderate to higher temperatures yield relatively high photoautotrophic bacterial Fe(II) oxidation (photoferrotrophic) rates and thus Fe(III) mineral formation. Therefore, biomass and Fe(III) settle together to the seafloor. (2) With decreasing temperatures photoautotrophic oxidation rates slow down and at the same time lower temperatures initiate abiotic Si precipitation from Si oversaturated ocean waters. SiO2 then settles to the seafloor. From Koehler et al. (2010).
We have also conducted systematic analyses of the physico-chemical characteristics of the cell-Fe(III) mineral aggregates, such as shape, size, density and chemical composition (Posth et al., 2010; Klueglein et al., 2014; Gauger et al., 2016a; Gauger et al., 2016b). For instance, we have shown experimentally that most aggregates are bulbous or ragged in shape, with an average particle size of 10-40 µm, and densities that typically range between 2.0 and 2.4 g/cm3; the cell fraction of the aggregates increased and their density decreased with initial Fe(II) concentration. The mineralogy of the ferric iron phase depended on the composition of the medium: goethite formed in cultures grown by oxidation of dissolved Fe(II) medium in the presence of low phosphate concentrations, while poorly ordered ferrihydrite and/or Fe(III) phosphates formed when high concentrations of phosphate were initially present. These results not only have an important bearing on nutrient and trace element cycling in the modern water column, but the size, shape and composition of the aggregates can be used to estimate aggregate reactivity during sediment diagenesis over short and geologic time scales (Posth et al., 2013a).
In a number of the biological Fe(II) oxidation experiments, a fraction of the cells remained planktonic, demonstrating a constant stoichiometric excess of Fe(III) in the precipitate (Konhauser et al., 2005; Posth et al., 2010; Gauger et al., 2016b). This observation has three important implications. First, the rather loose cell-mineral associations observed for photoferrotrophic species, such as the green-sulfur Fe(II)-oxidizing bacterium Chlorobium ferrooxidans KoFox, will have consequences for their preservation into the rock record as compared to other Fe(II)-oxidizing organisms that form mineral crusts around the cells. Indeed, it has always been interesting that so few microfossils have been recovered from BIF. Second, the presence or absence of a strong association of cells with minerals or encrustation in minerals also influences the rate and extent of sedimentation of the microbial cells, i.e., lesser amounts of organics are sedimented to the seafloor than cells with strong associations to Fe(III) minerals. Third, the segregation of ferric iron (that makes BIF) from planktonic biomass means that the latter may potentially contribute organic carbon to other sediments, such as black shales (Konhauser et al., 2018).
Post-Depositional Alteration of BIF
If a biological mechanism was important in the initial process of Fe(II) oxidation in the ancient ocean water column, it is expected that biomass would have settled to the seafloor along with the Fe(III) minerals (see Koehler et al., 2010; Posth et al., 2013a; Posth et al., 2013b; Konhauser et al., 2018). This organic carbon would subsequently have served as an oxidisable substrate during diagenesis and metamorphism, but the relevant question is what terminal electron acceptors were present at the seafloor? The paucity of O2 would have meant minimal nitrate and sulfate availability. By contrast, there was abundant ferric hydroxide deposited as BIF, and given the presence of partially reduced iron phases such as magnetite and siderite, a microbial process coupling the oxidation of organic carbon to the reduction of ferric iron mineral phases seems very likely. In fact, our models suggest that as much as 70% of biologically formed Fe(III) could have been recycled back into the water column via fermentation and organic carbon oxidation coupled to microbial Fe(III) reduction, with some fraction of the original biomass being consumed via methanogenesis, i.e., coupling the oxidation of acetate or H2 to methane formation (Konhauser et al., 2005). Coupling the reduction of Fe(III) minerals to the oxidation of organic matter – and possibly even methane - not only explains the low content of organic carbon in the BIF via the consumption of carbon (BIF have on average <0.5 wt.% organic carbon), but it also explains the textural features in the reduced iron minerals in BIF, such as magnetite overgrowths.
Model for the biological role in Fe cycling in the late Archean-Paleoproterozoic oceans. Based on Konhauser et al. (2005).
Despite our earlier studies that attempted to quantify the significance of this pathway based on models of the ancient Fe cycle, the only direct evidence of a biological role in magnetite formation in BIF are their iron isotope compositions and unique crystallography which are reminiscent of biologically-generated magnetite. In the case of the former, Fe and Si isotope composition of coexisting mineral phases in samples from the ~2.5 billion year old Kuruman Iron Formation (Transvaal Supergroup, South Africa) and Dales Gorges Member were measured. We showed that magnetite exhibits negative 56Fe values, which can be attributed to a variety of diagenetic pathways: the light Fe isotope composition was inherited from the Fe(III) precursor, while heavy Fe(II) was lost by abiotic reduction of the Fe(III) precursor or light Fe(II) was gained from external fluids (Steinhoefel et al., 2010). Diagenetic Fe(III) reduction caused by oxidation of organic matter and Fe redistribution is supported by the C isotope composition of a carbonate-rich sample containing primary siderite. In the case of the latter, a colleague of mine from the University of Hong Kong (Yi-Liang Li) conducted high resolution mineral analyses from the 2.48 Gyr old the Dales Gorge Member BIF and reported magnetite with a lattice constant and Fe(II)/Fe(III) stoichiometry very similar to those produced by DIR bacteria, such as Geobacter, Shewanella and Thermoanaerobacter (Li et al., 2011). He also reported the detection of an Fe(III)-acetate salt, as well as nanocrystals of apatite in association with magnetite; this combination of features points to the original presence of biomass in the BIF sediments, but also that the organic carbon served as an electron donor during bacterial Fe(III) reduction. However, the biogenesis hypothesis lacks an explanation as to why modern biogenic magnetite crystals are generally a few hundred nm or smaller in size, yet the magnetite crystals in BIF are mostly tens micrometers or larger in size. In a follow-up study, (Li et al., 2013) demonstrated that biogenic magnetite crystals can grow to nearly 1 micron in size through a three-stage sequence, (i) beginning with dissimilatory Fe(III) reduction of an initial ferric iron-rich sediment coupled to the oxidation of dead phytoplankton biomass, (ii) followed by magnetite crystal aging, and (iii) ultimately pressure-temperature induced abiotic alteration (compression and heating to 1 kbar and 150°C, respectively) of the biogenic magnetite during metamorphism.
Coupled with actual BIF analyses, my colleagues at the University of Tubingen have been conducting a number of diagenesis experiments aimed at replicating the conditions under which BIF minerals formed. One of those studies involved incubating mixtures of ferrihydrite (as proxy for BIF precursor sediment) and glucose (as a proxy for degradation phases of microbial biomass) in gold capsules at 1.2 kbar and 170°C to mimic diagenesis (Posth et al., 2013b). Under these conditions, ferrihydrite - Fe(OH)3 - transforms to hematite, magnetite, and siderite; silica-coated ferrihydrite yielded hematite and siderite, but not magnetite. The results showed that electron transfer from organic carbon to Fe(III) minerals and between transitional Fe phase minerals during temperature/pressure diagenesis can drive the production of key BIF minerals. Interestingly, however, further capsule experiments conducted by a former PhD student at Tubingen (Maximilian Halama), with either glucose or actual cell biomass showed different results despite being subjected to similar conditions as above (Halama et al., 2016). No magnetite was formed from Fe(III) minerals when microbial biomass was present as the carbon and electron sources for thermochemical Fe(III) reduction. This could be due to biomass-derived organic molecules binding to the mineral surfaces and preventing solid-state conversion to magnetite. By contrast, diagenetic magnetite was either formed by microbial Fe(III) reduction during early diagenesis, i.e., below 120°C, or by thermochemical Fe(III) reduction with simple organic compounds at higher temperatures. In another study, led by a former PhD student at the University of Tubingen (Inga Köhler), we showed experimentally that spheroidal siderite, which is preserved in many BIF, and could have been a precursor to rhombohedral or massive siderites, formed by reacting ferrihydrite with glucose at similar pressure and temperature conditions to above (Köhler et al., 2013). Depending on abundance of the siderite, she found that it is also possible to draw conclusions about the Fe(III):C ratio of the initial ferrihydrite-biomass sediment. Our results suggested that spherical to rhombohedral siderite structures in deep-water, Fe-oxide BIF can be used as a biosignatures for photoferrotrophy, whereas massive siderites reflect high cyanobacterial biomass loading in highly productive shallow-waters.
Set-up for diagenesis experiments, as run in Tubingen at the Kappler laboratory.
Because a number of recent studies have purported how the trace metal abundances in BIF can be used as proxies for the bioavailability of trace metals in ancient seawater (see next section), it is crucial to demonstrate there was minimal secondary mobilization during burial. In this regard, my current PhD student (Jamie Robbins) measured the mobility of Zn and Ni from ferrihydrite in the absence and presence of organic matter (glucose) during simulated diagenesis (170°C, 1.2 kbar). Quantitative concentration data, coupled with x-ray diffraction analysis and electron microprobe element mapping, showed that both metals are relatively immobile during simulated diagenesis: more than 99% of Zn and 91.9% of Ni were retained under the varied conditions considered here (Robbins et al., 2015). Overall, our results indicated that paleomarine Zn and Ni concentrations are likely to be faithfully recorded in well-preserved BIF deposits.
In addition to the reduced iron minerals in BIF being used as proxies for ancient biomass deposition, there are other BIF features that point to a microbial role during BF diagenesis. Papineau et al. (2010) conducted a detailed petrographic analyses of apatite grains in two BIF from the Paleoproterozoic of Uruguay and Michigan. These BIF typically contain apatite co-existing with carbonaceous matter, which suggests the original presence of diagenetically altered biomass.
The goals for the future are to continue trying to determine the mechanisms by which BIF precipitated and were post-depositionally altered, and ascertain how those chemical sediments can be used as archives of Earth’s surface conditions at the time of their deposition? Major questions being sought are:
How did the different Fe(II) oxidation mechanisms change through time in concert with the changing redox state of the ocean-atmosphere system?
Can the less developed stable metal isotopes (e.g., Si, Ge, Cu) shed new insights into the sources for the iron and silica in BIF?
Can diagenesis experiments at high P/T inform about mineral and compositional changes that BIF precursor sediment underwent during later stage metamorphism?