nai

Project Progress

Banded iron formations (BIFs) are Precambrian chemical marine sedimentary rocks that represent the major source of Fe used in today’s society. Early studies suggested a continental source of Fe for BIFs (Cloud, 1973), although direct riverine input of Fe has been questioned because of the low-detritus components in the large superior-type BIFs (Holland, 1978). The discovery of mid-ocean ridge (MOR) hydrothermal systems in the 1970s and the similarity of certain rare earth element (REE) signatures (e.g., positive Eu anomaly) between BIFs and MOR hydrothermal fluids led to a commonly accepted model, where BIFs were formed by oxidation of hydrothermally sourced aqueous Fe(II) (Bau and Möller, 1993). More recent work, particularly the combination of Nd isotopes and REEs, suggests a more complex origin for REEs in BIFs, where a significant component is sourced to the continents (Alexander et al., 2009). Interpretations of Fe sources for BIFs using REE patterns and Nd-isotope ratios are, however, based on the underlying assumption that REEs and Fe pathways were coupled during transport and deposition of materials for BIFs, although this assumption has not been independently tested.

We focused this project (Li et al., 2015) on the Dales Gorge member of the 2.5 Ga age Brockman Iron Formation (Hamersley Basin, Western Australia, Australia), the world’s most extensive superior-type BIF that represents the climax of BIF deposition in the geologic record (Klein, 2005). The BIF samples analyzed in this study come from the type section diamond drill core for the Dales Gorge member (Trendall and Blockley, 1968). The depositional age of the Dales Gorge member is between 2.50 and 2.45 Ga (Trendall et al., 2004). The Dales Gorge member is ∼160- to 140-m thick, consisting of 17 iron-rich, meter-scale macrobands and 16 shale macrobands, named BIF0–BIF16 and S1–S16, respectively (Figure 1), and preserved sections reflect deposition in deep water conditions. The meterscale, iron-rich macrobands are each composed of centimeter scale, iron-rich mesobands (Figure 1), which in turn, contain numerous sub-millimeter microbands (Trendall et al., 2004). Iron-isotope compositions were measured at various sample scales from bulk (∼300-mg samples) measurements on the same aliquot used for Nd and REE analysis by isotope dilution mass spectrometry for mesoband samples in different macrobands to in situ analysis of Fe-isotope ratios and REE contents using femtosecond laser ablation for microbands. Measured Fe- and Nd-isotope compositions reveal a large variation in both isotope systems: from −0.83‰to +1.30‰ in δ⁵⁶Fe and −2.2 to +3.0 in ε_Nd (Figure 1). Between macrobands and within a macroband (e.g., BIF16), Fe- and Nd-isotope compositions oscillate over meter scales along the drill core (Figure 1). In contrast, there is limited variation (<±0.2‰) in δ⁵⁶Fe values among microbands over centimeter scales of drill core depth (Figure 1). Overall, there is a broad positive correlation between the δ56Fe and ε_Nd values (Figures 1 and 2).

The large range in εNd values suggests mixing between a low-ε_Nd continental source and a high-ε_Nd mantle source for the REEs (DePaolo, 1988). The relative slopes of ε_Nd–δ⁵⁶Fe variations (Figure 2) are a function of the partition coefficients (K_d) for the REEs in iron oxides (Koeppenkastrop and De Carlo, 1992; Quinn et al., 2006) as well as the contrast in Nd-isotope compositions of the hydrothermal plume relative to the ambient ocean that had a continental Nd isotope signature (Piepgras and Wasserburg, 1980). Although the Kd for REE partitioning into iron hydroxides can exceed 10⁵ (typically >1,000) (Koeppenkastrop and De Carlo, 1992; Quinn et al., 2006), even a modest Kd of 10–50 shows that Nd contents of the hydrothermal plume will be rapidly depleted, such that mixing with ambient seawater will move the hydrothermal precipitates horizontally to the left in Figure 2, failing to reproduce the observed εNd–δ56Fe variations. We conclude (Li et al., 2015) that the εNd–δ56Fe variations cannot be produced through partial oxidation of a single source of hydrothermally derived Fe but instead, reflect mixing of water masses that had distinct Fe- and Nd-isotope compositions (continental vs. mantle sources) in terms of dissolved Fe(II)aq and Nd followed by post-mixing oxidation and precipitation as ferric oxyhydroxides, ultimately forming the BIFs.

Microbial dissimilatory iron reduction (DIR) in coastal sediments is a mechanism that can release significant quantities of isotopically light Fe(II)_aq to the oceans (Severmann et al., 2006; Johnson et al., 2008). Support for a DIR Fe shuttle as the source of the low-ε_Nd and -δ⁵⁶Fe component in the 2.5-Ga BIFs of this study comes from ε_Nd–Sm/Nd relations (Figure 3). Microbial dissolution of iron hydroxides in modern marine sediments is accompanied by significant REE fractionation, where Fe(II)-rich pore waters contain significantly higher Sm/Nd ratios than bulk sediments (Haley et al., 2004), and this relation matches that seen for Sm/Nd ratios of the low-εNd continental component relative to Sm/Nd ratios for Archean shales (Figure 3). For the hydrothermal component, ε_Nd–Sm/Nd relations indicate low Sm/Nd ratios (Figure 3), which likely reflect the effect of Fe(III) hydroxide precipitation given the fact that both laboratory experiments and field studies have shown that adsorption to particulate Fe(III) hydroxide fractionates REEs and that Sm is more strongly adsorbed onto Fe(III) particulates than Nd (Koeppenkastrop and De Carlo, 1992), decreasing the Sm/Nd ratio in the remaining solution (Figure 3). In short, the observed REE data (Eu anomaly and Sm/Nd ratio) are consistent with the proposed mixing model between two end members with fractionated REE signatures.

Demonstration of the dual sources of Fe for BIFs, one hydrothermal and one DIR-driven Fe shuttle from continental shelves, raises the question of the timescales over which such sources could have operated. It has been proposed, for example, that Fe(III) oxides in BIFs may be produced by photosynthetic oxidation of Fe(II)aq (Kappler et al., 2005), which in turn, might indicate seasonal variations that correlate with light intensity. At the smallest scale of banding, the sub-millimeter microbands that have been interpreted to represent annual varve-like layers (Morris, 1993), in situ Fe-isotope analyses in this study show relatively small variation over timescales of ∼10³ y (Figure 1). These results suggest that the DIR-driven Fe shuttle did not respond to seasonal changes in photosynthetic primary productivity, such as the amount of organic carbon production, at least from the perspective of the site of BIF deposition for the samples studied here. In contrast, at the scale of individual macrobands, which represent periods of ∼10⁵ y, Fe sources were clearly variable, as shown, for example, by macrosampling of BIF16 (Figure 1). Over the ∼50-My period represented by the entire Dales Gorge member (Trendall et al., 2004), the balance between hydrothermal and continental DIR Fe sources also varied (Figure 1). We conclude that the relative proportions of DIR and hydrothermal Fe sources recorded in BIF deposition were controlled by long-timescale changes that reflect variability in basin-wide circulation changes on the order of 10⁵–10⁶ y. It is possible that basin-wide sampling transects might record different scales of isotopic variability depending on conditions that affected the proportion of DIR- and hydrothermally sourced Fe. Nevertheless, the combined Fe- and Nd-isotope analysis indicates that BIFs formed from two sources of Fe and that an active DIR-driven Fe shuttle was operating at 2.5 Ga.

References Cited
Alexander BW, Bau M, Andersson P (2009) Neodymium isotopes in Archean seawater and implications for the marine Nd cycle in Earth’s early oceans. Earth Planet Sci Lett 283(1–4):144–155.

Bau M, Möller P (1993) Rare earth element systematics of the chemically precipitated component in early precambrian iron formations and the evolution of the terrestrial atmosphere-hydrosphere-lithosphere system. Geochim Cosmochim Acta 57(10):2239–2249.

Cloud P (1973) Paleoecological significance of the banded iron-formation. Econ Geol 68(7):1135–1143.

DePaolo DJ (1988) Neodymium Isotope Geochemistry: An Introduction (Springer, Heidelberg).

Haley BA, Klinkhammer GP, McManus J (2004) Rare earth elements in pore waters of marine sediments. Geochim Cosmochim Acta 68(6):1265–1279.

Holland HD (1978) The Chemistry of the Atmosphere and Oceans (Wiley, New York).

Johnson CM, Beard BL, Roden EE (2008) The iron isotope fingerprints of redox and biogeochemical cycling in modern and ancient Earth. Annu Rev Earth Planet Sci 36(1): 457–493.

Kappler A, Pasquero C, Konhauser KO, Newman DK (2005) Deposition of banded iron formations by anoxygenic phototrophic Fe(II)-oxidizing bacteria. Geology 33(11):865–868.

Klein C (2005) Some Precambrian banded iron-formations (BIFs) from around the world: Their age, geologic setting, mineralogy, metamorphism, geochemistry, and origins. Am Mineral 90(10):1473–1499.

Koeppenkastrop D, De Carlo EH (1992) Sorption of rare-earth elements from seawater onto synthetic mineral particles: An experimental approach. Chem Geol 95(3–4):251–263.

Li W, Beard BL, Johnson CM (2015) Biologically recycled continental iron is a major component of banded iron formations. Proc Nat Academ Sci 112(27):8193-8198.

Morris RC (1993) Genetic modelling for banded iron-formation of the Hamersley Group, Pilbara Craton, Western Australia. Precambrian Res 60(1–4):243–286.

Piepgras DJ, Wasserburg GJ (1980) Neodymium isotopic variations in seawater. Earth Planet Sci Lett 50(1):128–138.

Quinn KA, Byrne RH, Schijf J (2006) Sorption of yttrium and rare earth elements byamorphous ferric hydroxide: Influence of pH and ionic strength. Mar Chem 99(1–4):128–150.

Severmann S, Johnson CM, Beard BL, McManus J (2006) The effect of early diagenesis on the Fe isotope compositions of porewaters and authigenic minerals in continental margin sediments. Geochim Cosmochim Acta 70(8):2006–2022.

Trendall AF, Blockley JG (1968) Stratigraphy of the Dales Gorge Member of the Brockman Iron formation in the Precambrian Hamersley Group of Western Australia. Geological Survey of Western Australia Annual Report for 1967, ed Davies AB (The Department of Mines Western Australia, Perth, Australia), pp 86–91.

Trendall AF, Compston W, Nelson DR, De Laeter JR, Bennett VC (2004) SHRIMP zircon ages constraining the depositional chronology of the Hamersley Group, Western Australia. Aust J Earth Sci 51(5):621–644.