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The Joffre banded iron formation, Hamersley Group, Western Australia: Assessing the palaeoenvironment through detailed petrology and chemostratigraphy
Accepted Manuscript Title: The Joffre Banded Iron Formation, Hamersley Group, Western Australia: Assessing the Palaeoenvironment through detailed Petrology and Chemostratigraphy Author: Rasmus Haugaard Ernesto Pecoits Stefan Lalonde Olivier Rouxel Kurt Konhauser PII: DOI: Reference:
S0301-9268(15)00362-9 http://dx.doi.org/doi:10.1016/j.precamres.2015.10.024 PRECAM 4398
To appear in:
Precambrian Research
Received date: Revised date: Accepted date:
8-3-2015 28-9-2015 17-10-2015
Please cite this article as: Haugaard, R., Pecoits, E., Lalonde, S., Rouxel, O., Konhauser, K.,The Joffre Banded Iron Formation, Hamersley Group, Western Australia: Assessing the Palaeoenvironment through detailed Petrology and Chemostratigraphy, Precambrian Research (2015), http://dx.doi.org/10.1016/j.precamres.2015.10.024 This is a PDF file of an unedited manuscript that has been accepted for publication. As a service to our customers we are providing this early version of the manuscript. The manuscript will undergo copyediting, typesetting, and review of the resulting proof before it is published in its final form. Please note that during the production process errors may be discovered which could affect the content, and all legal disclaimers that apply to the journal pertain.
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We present petrology and geochemistry of the largest single known banded iron formation in the world Dominant rock types are oxide BIF, silicate-carbonate-oxide BIF with minor stilpnomelane mudrock and stilpnomelane-rich tuffaceous mudrock. Submarine hydrothermal input and fine-grained pyroclastic detritus were the main sources to the seawater Dominant volcanic sources were likely felsic of composition Dispersed stilpnomelane represents diluted ash material that interacted with Fe-rich seawater prior to deposition
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Highlights
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The Joffre Banded Iron Formation, Hamersley Group, Western
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Australia: Assessing the Palaeoenvironment through detailed Petrology
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and Chemostratigraphy
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Rasmus Haugaard1*, Ernesto Pecoits2, Stefan Lalonde3, Olivier Rouxel3, Kurt Konhauser1
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Department of Earth and Atmospheric Sciences, University of Alberta, Edmonton, Alberta, T6G 2E3, Canada
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Equipe Géobiosphère, Institut de Physique du Globe-Sorbonne Paris Cité, Université Paris Diderot, CNRS, 1 place
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Universite Europe ene de Bretagne, Institut Universitaire Europe en de la Mer, Plouzane 29280, France
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Corresponding author: [email protected]
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Jussieu, 75238 Paris, France
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Abstract
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The Joffre Member of the Brockman Iron Formation is by volume the largest
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single known banded iron formation (BIF) in the world. Here we present detailed
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petrology and chemostratigraphy through the entire 355 m core section of this ~2.45
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billion year old unit. Oxide BIF and silicate-carbonate-oxide BIF dominate the lithology,
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with minor amounts of interbedded stilpnomelane mudrock, stilpnomelane-rich
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tuffaceous mudrock and calcareous mudrock. Beside chert and magnetite, the prominent
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mineralogy is riebeckite, ankerite, hematite, stilpnomelane and crocidolite. The BIF is
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characterised by an average of 50 wt.% SiO2 and 44.5 wt.% Fe2O3 and an overall low
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abundance of Al2O3 (<1 wt.%), TiO2 (<0.04 wt.%), and trace metals such as Cr (<10
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ppm), Ni (<5 ppm) and Mo (<0.5 ppm). It has a high ∑REE (rare earth element) content
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(up to 41 ppm) and a fractionated shale-normalised (SN) seawater REY (rare earth
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element + yttrium) pattern having an enrichment of HREE (heavy rare earth elements)
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relative to LREE (light rare earth elements) with an average (Pr/Yb)SN of 0.24. The REY
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patterns also show a positive LaSN anomaly, no CeSN anomaly and a weakly developed
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positive YSN anomaly. Iron isotopes (56Fe) with positive 56Fe values of +0.04‰ to
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+1.21‰ suggest that a large part of the hydrothermal iron was partly oxidized in the
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upper water column and subsequently precipitated as ferric oxyhydroxides. No epiclastic
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grains have been found; rather submarine hydrothermal fluids and fine-grained
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volcanogenic detritus controlled BIF chemistry. The former source is reflected through a
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constant positive EuSN anomaly throughout the core (average EuSN anomaly of 1.6 with a
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peak of 2.1 between 100-155 m depth), while the latter source is best reflected through
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the stilpnomelane-rich tuffaceous mudrock consisting of volcanic ash-fall tuff with relict
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shards set in a stilpnomelane matrix. The mudrock is overlain by well-preserved wavy
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laminae and laminae sets of stilpnomelane microgranules that likely originated from re-
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worked volcanic ash formed either on the seafloor or in the water column prior to
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deposition. An enriched HREE-to-LREE pattern, a high iron content (~30 wt%), and a
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56Fe value of +0.59‰ collectively imply that the mudrock facies interacted with the Fe-
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rich seawater prior to deposition. The TiO2-Zr ratio of the BIF and the associated
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mudrocks suggest a felsic-only-source related to the same style of volcanics as the
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slightly younger Woongarra rhyolites. Given the observation that the dominant control on
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the seawater chemistry was associated with felsic volcanics, we speculate that the fine-
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grained pelagic ash particles may have sourced bio-available nutrients to the surface
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water. This would have facilitated enhanced biological productivity, including bacterial
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Fe(II)-oxidation which is now recorded as the positively fractionated
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Fe iron oxide
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minerals in the Joffre BIF. Alongside submarine hydrothermal input to the basin, the
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dominant control on the ocean chemistry seems to have been through volcanic and
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pyroclastic pathways, thereby making the Joffre BIF poorly suited as a chemical proxy
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for the study of atmospheric oxygen and its weathering impact on local landmasses.
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Keywords: Palaeoproterozoic; Hamersley Group; Joffre Banded Iron Formation; Seawater chemistry;
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Provenance; Stilpnomelane
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1. Introduction
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Banded iron formations (BIF) are iron-rich (15-40 wt.% Fe2O3) and siliceous (40-60
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wt.% SiO2) chemical sedimentary deposits that precipitated from seawater throughout
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much of the Archaean and Palaeoproterozoic (3.8–1.85 Ga). They are also, more often
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than not, laminated, with banding observed on a wide range of scales, from coarse
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macrobands (meters in thickness) to the characteristic mesobands (centimeter-thick units)
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by which they are typically defined (i.e., banded iron formation), to microbands
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(millimeter to submillimeter). They typically contain low concentrations of Al2O3 (<1
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wt.%) and incompatible elements (Ti, Zr, Th, Hf and Sc <20 ppm), which indicate
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minimal detrital input to the depositional basin, although this does not hold for all type of
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iron formations (see Bekker et al., 2010 for review). For instance, granular iron
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formations (GIF) typically lack banding and are made of granules of chert and iron
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oxides or silicates with early diagenetic chert cement filling pore space. Their texture
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implies that they formed in high-energy environments, with the granules being derived by
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sedimentary re-working of iron-rich clays, mudstone, arenites, and even stromatolites
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(e.g., Ojakangas, 1983; Simonson and Goode, 1989). Towards the end of Archaean, the marine depositional setting for BIF formation
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changed from depositional basins with rapid thermal subsidence and deposition of large
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volume of interbedded volcanic and volcanogenic greywackes (e.g., Lowe and Tice,
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2007; Haugaard et al. 2013), to a more stable style of sedimentation in extensive shallow
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marine basins along stable continental platforms (e.g., Taylor and McLennan, 1981;
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Eriksson et al., 2001; Condie, 2004; Barley et al., 2005). BIF deposited in the former
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setting are considered Algoma-type, whereas BIF formed in the latter setting are
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Superior-type. The latter includes the major BIF of the earliest Paleoproterozoic, such as
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the Hamersley Group BIF (Gross, 1980).
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The mineralogy of BIF from the best-preserved sequences is remarkably uniform,
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comprising mostly chert, magnetite, hematite, Fe-rich silicate minerals (stilpnomelane,
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greenalite, minnesotaite, and riebeckite), carbonate minerals (siderite, ankerite, calcite,
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and dolomite), and minor sulphides (pyrite and pyrrhotite); the presence of both ferric
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and ferrous minerals gives BIF an average oxidation state of Fe2.4+ (Klein and Beukes,
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1992). It is generally agreed that none of the minerals in BIF are primary. Instead, the
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minerals reflect significant post-depositional alteration under diagenetic and metamorphic
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conditions (including, in some cases, post-depositional fluid flow). The effect of
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increasing temperature and pressure is manifested by the progressive change in
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mineralogy through replacement and recrystallisation, increase in crystal size and
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obliteration of primary textures (Klein, 2005; Bekker et al., 2010).
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The abundance of BIF in Precambrian successions was used in early studies to
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argue for a largely anoxic atmosphere and ocean system (e.g., Cloud, 1973; Holland,
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1984) because the accumulation of such large masses of iron found in the form of
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Superior-type BIF required the transport of Fe(II); Fe(III) is essentially insoluble at
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circumneutral pH values. Early studies invoked a continental source of iron for BIF
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because Fe(II) would have been much more mobile in the absence of atmospheric O2
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(e.g., James, 1954; Lepp and Goldich, 1964) and the continents were more mafic in
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composition (Condie, 1993). However, detailed studies that followed in the Hamersley
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basin, Western Australia, suggested that the amount of iron deposited there was on the
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order of 1 x 1013 ton (Trendall and Blockley, 1970; Trendall and Blockley, 2004). This
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estimate would have required rivers the size of the modern Amazon to transport orders of
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magnitude more iron than they do today. This led Holland (1973) to suggest that iron was
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instead sourced from deep marine waters and supplied to the depositional settings via
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upwelling. Recently, however, a new model based on Fe- and Nd-isotopes suggests that a
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large part of the iron in BIF was continental and mobilized by microbial Fe(III) reduction
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and transported through a benthic iron shuttle to the BIF depositional basin (Li et al.,
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2015).
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Based on rare earth element (REE) composition of BIF, it is now generally
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accepted that deep-sea hydrothermal processes are the most likely source of Fe. Shale
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normalised (SN) europium (Eu) anomalies have been central in the use of REE to trace
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the hydrothermal input. Eu enrichment in chemical sedimentary rocks precipitated from
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seawater indicates a strong influence of high-temperature hydrothermal fluids on the
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seawater dissolved REE load (e.g., Klinkhammer et al., 1983; Derry and Jacobsen, 1988;
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1990). It is generally assumed that Fe and REE will not be fractionated during transport
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from spreading ridges or other exhalative centres, and, therefore, a strong positive EuSN
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anomaly indicates that the iron in the BIF precursor sediment was hydrothermally derived
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(e.g., Slack et al., 2007). In addition to REE concentrations, Sm-Nd isotopes have been
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used to constrain REE and Fe sources to seawater (e.g., Miller and O'Nions, 1985; Derry
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and Jacobsen, 1990; Alibert and McCulloch, 1993). The Archaean and Palaeoproterozoic
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oceans were likely strongly heterogeneous in their Nd(t) values, with +1 to +2 values
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typical of the deep-waters dominated by hydrothermal sources and lower values, down to
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-3, typical of shallow-waters dominated by terrestrial sources (Frei et al., 1999, 2007,
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2008; Alexander et al., 2008).
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Within the large Superior-type BIF, the presence of diagenetic to low-grade
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metamorphic phyllosilicate minerals, such as greenalite [(Mg,Fe)3Si2O5(OH)4] and
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stilpnomelane [K0.6(Fe2+,Fe3+,Mg)6Si8Al(O,OH)27nH2O], mostly occur as dense bands
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interbedded with chemical precipitated minerals such as amorphous silica and ferric
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hydroxides (La Berge, 1966; Ayres, 1972; Morris, 1983). The precursor of stilpnomelane
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are believed to have been an iron(III)-rich clay (smectite) derived from volcaniclastic
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sources likely of basaltic provenance (Trendall and Blockley, 1970; Ewers and Morris,
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1981; Pickard 2002; Krapež et al. 2003). Where all primary minerals have been
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overprinted by diagenesis and low-grade metamorphism, preservation of primary textures
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is rare. As for stilpnomelane, unique preservation of silt-sized microgranules, or
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spheroids, opens up the possibility to study the precursor sediment. These stilpnomelane
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microgranules are relatively uncommon but has been observed in Superior-type BIF in
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the Lake Superior region, Canada (Van Hise and Leith, 1911; Moore, 1918), in the
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Kuruman Iron Formation, South Africa (Beukes, 1973) and in the Brockman Iron
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Formation, Western Australia (Ayres, 1972, Krapež et al. 2003). Recently, Rasmussen et
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al., (2013, 2014) presented well-documented lamina sets of stilpnomelane microgranules
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from the Dales Gorge BIF of the Brockman Iron Formation and proposed that they were
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generated by flocculation of iron-rich, Al-poor hydrous silicates either on the seabed or
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within the water column and subsequently shaped and reworked by density currents.
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In this work, we investigate the ~2.45 Ga (Pickard et al., 2002) Joffre Member
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from the Brockman Iron Formation. This is the single largest known BIF worldwide,
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containing approximately 4.3x1013 tonnes of iron at the time of deposition (Trendall and
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Blockley, 2004). This laterally extensive Superior-type BIF, therefore, represents the
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composition of a large volume of ocean water. We conducted detailed petrologic and
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geochemical analyses of a core section drilled through the entire ~355 m of stratigraphic
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depth of the Joffre BIF. Unlike the well-explored Dales Gorge BIF (e.g., Ayres, 1972;
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Ewers and Morris, 1981; Krapež et al., 2003; Pickard et al., 2004; Pecoits et al., 2009;
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Rasmussen et al., 2013, 2014, 2015), this key Hamersley BIF has not previously been
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analysed at high-resolution, nor has a detailed comparison of chemostratigraphy between
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different lithologies through a complete succession ever been attempted. Furthermore, by
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directly preceding the Great Oxidation Event (GOE) and the marked increase in Cr
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dissolution found in the Weeli Wolli Formation (see Konhauser et al., 2011), the Joffre
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BIF potentially covers the transition from an anoxic to a partially oxygenated Earth. As
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such, high-resolution petrography, geochemistry and isotopic studies will provide vital
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information about the evolution of the sediment, the depositional basin and, in particular,
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the source inputs controlling the seawater composition.
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173 2. The Hamersley Group
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The Joffre BIF is a member of the ~620 m thick Brockman Iron Formation which makes
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up part of the 2.63-2.45 Ga Hamersley Group (Trendall et al., 2004). The Hamersley
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Group comprises almost 2.5 km of consecutive sedimentary and volcanic rocks located
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within the ca. 80,000 km2 Hamersley Province of the Pilbara craton in North West
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Australia, approximately 1100 km north of Perth (Fig. 1). In the lower part, it consists of
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dolomite, shale and BIF, while the upper part consists of dolerite, various lava types and
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BIF with minor amounts of tuffs and shales (Trendall and Blockley, 1970). Underlying
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the Hamersley Group, is the 2.78-2.63 Ga (Arndt et al., 1991) Fortescue Group, which
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consists of flood basalts and rhyolites. These volcanics were laid down on the uplifted
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and eroded Pilbara block (Trendall, 1968). This volcanic succession may contain
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remnants of a Large Igneous Province (LIP) as suggested by Ernst et al. (2004).
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The Brockman Iron Formation of the Hamersley Group is divided into four sub-
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lithostratigraphic units, namely the lowermost Dales Gorge Member (BIF), the
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Whaleback Shale Member, the Joffre Member (BIF), and the uppermost Yandicoogina
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Shale Member (Fig. 1B). After deposition, these laterally extensive BIF have all
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experienced minor folding and basinal uplift along with low-grade regional
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metamorphism - from burial prehnite-pumpellyite facies to greenschist facies (Smith et
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al., 1982).
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Geochronological constraints (see Fig. 1B) on the Brockman Iron Formation were
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established by Pickard (2002), and absolute U-Pb zircon ages with interpolated
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stratigraphic age boundaries of the Hamersley Group are presented in Trendall et al. 9
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(2004). The best analytical age estimates for the deposition of the Joffre BIF is 2454±3
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Ma (Pickard, 2002). This age has been established by SHRIMP U-Pb zircon ages
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analysed on 19 zircon grains from interbedded tuffaceous mudrock facies at the top of the
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Joffre BIF (Fig. 1B). The best depositional age estimate for the base of the succession is
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2459±3 Ma established by only 6 zircon grains (Pickard, 2002, Fig. 1B). Without taking
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the uncertainties into consideration, there are 166 m between the above two age peaks,
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which potentially yields a compacted sedimentation rate of 33 m/million year (Pickard,
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2002).
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General consensus exists regarding the depositional model of the Brockman Iron
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Formation. According to this model, the succession was deposited on a large, stable, and
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clastic-starved, continental platform, which was influenced by episodic inputs of fine-
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grained tuffaceous detritus (e.g., Gross, 1983; Morris, 1983; Krapež et al., 2003). Blake
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and Barley (1992) proposed a gradually subsiding open-shelf developed within a backarc
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setting under influence of tuffaceous material sourced from a subduction-related
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magmatic arc. In addition, Barley et al. (1997) found that deposition of the Hamersley
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BIF were possibly linked to major submarine magmatic plume activity in the form of a
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LIPs. Morris (1993) also suggested that the depositional environment for the Hamersley
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BIF included a steady source of silica and iron with minor lateral variation in the
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deposition, and a water depth that was deeper than the formation of GIF but shallow
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enough to form the large carbonate platforms. In the absence of any shoreline facies and
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the lack of siliciclastics within the BIFs, Morris and Horwitz (1983) further argued that
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BIF precipitation and deposition took place on an outer shelf that was isolated by a
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carbonate barrier.
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219 3. Rock core, sampling and analytical methods
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The drill core, DD98SGP001 (diamond drillcore 1998 Silvergrass Peak #001) was drilled
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as part of the Rio Tinto project in the Silvergrass Peak area (see Fig. 1A for location).
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The core samples of DD98SGP001 were obtained at the Rio Tinto core library in Perth,
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Western Australia. A total of 31 core samples (DD98-1 - DD98-30), each ~30 cm long,
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were obtained through a total core length of 354.5 m (94 m to 448.5 m depth).
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The core samples were split and one half was stored as future reference material.
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A total of 40 thin sections were processed and examined using reflected and transmitted
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light microscopy. In addition, six thin section slabs were polished and carbon coated for
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electron microprobe analyses. Backscatter electron images, elemental distribution maps,
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EDS (energy dispersive spectrometry) and WDS (wavelength dispersive spectrometry)
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were obtained with a JEOL Microprobe 8900 at the University of Alberta. The current
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was set to ~20 nA and the probe beam diameter was set to 10 microns. Counting time
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was 20 sec on peak, and 10 sec on background. Dwell time was 10 ms. Standardization to
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minerals with known element concentration were done on hematite, chromite, kaersutite
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and diopside for each 50 measurements.
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A total of 30 samples were selected for trace element analyses. Approximately 4
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cm slabs from each sample were divided into chips and subsequently crushed on an agate
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mill. The crushed rock powders were dissolved with HF+HNO3 and analyzed using a
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PerkinElmer Elan6000 Quad-ICPMS (quadrupole inductively coupled plasma mass
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spectrometer) at the University of Alberta. Accuracy and precision of the analytical
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protocol was verified with the use of the well-established international whole-rock basalt
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standard (CRPG Nancy). Errors on this standard and on duplicates are both below 10%.
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Oxide interferences on Ce show that CeO/Ce < 3% and any oxide interferences are
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therefore considered negligible. For major elements, 17 samples were further analysed by
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Code 4C (11+) Whole Rock Analysis-XRF at Activation Laboratories Ltd., Ontario,
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Canada. For iron isotopes, a total of 42 whole rock powder samples were selected.
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Whole rock Fe isotope compositions were analysed at the French oceanographic
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institution IFREMER, Brest campus, following previously published methods (Rouxel et
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al., 2005; 2008). Briefly, 50-100 mg of sample powder was digested overnight at 80 °C in
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4ml 1:1 HF-HNO3 followed by 4 ml aqua regia, with complete evaporation in between.
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Samples were then taken up in 4ml 6N HCl, from which Fe was purified on Bio-Rad
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AG1X8 anion resin (2 ml wet resin bed) using 6N HCl for matrix elution followed by
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0.24N HCl for Fe elution. Fe isotope compositions were determined using a Thermo
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Scientific Neptune multicollector inductively coupled plasma mass spectrometer
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operating at medium resolution to resolve isobaric interferences such as 40Ar14N on 54Fe,
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correction, introduced to the instrument using an Apex Q desolvating nebuliser
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(Elemental Scientific, Omaha, NE, USA), and ‘sample-standard bracketing’ was used for
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data normalisation to a Fe isotope standard solution of IRMM-14 run before and after
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each
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0.09±0.09‰ and 0.65±0.14‰, respectively, consistent with previous work (e.g.,
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Planavsky et al., 2012), and repeated measurements (n=59) of the reference material
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IRMM-14 (Taylor et al., 1992) constrained average internal precision over the analytical
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sessions to better than ±0.065 (2 SD).
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Fe, and
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Ar16O1H on
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unknown.
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Fe. Solutions were doped with Ni for mass bias
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Concentrations of REE and Y were shale normalised (SN) to Post-Archaean Australia
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Shale (PAAS) after Taylor and McLennan (1985). Potential anomalies of La (La/La*SN)
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and Ce (Ce/Ce*SN) were obtained by the procedure proposed by Bau and Dulski (1996)
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using the combination of Ce/Ce*PAAS = Ce/(0.5*La+0.5*Pr) and Pr/Pr*SN =
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Pr/(0.5*Ce+0.5*Nd). If Ce/Ce* < 1 but Pr/Pr*SN ≈ 1 a positive La anomaly is evident. If
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Ce/Ce*SN < 1 but Pr/Pr*SN > 1.05 a negative Ce anomaly is evident. The Eu anomaly
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(Eu/Eu*SN) was calculated as Eu/Eu*SN = Eu/[0.67Sm+0.33Tb].
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272 4. Petrography and mineralogy
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4.1. Rock types
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Based on the dominant mineralogy, we define five lithological subdivisions within the
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Joffre BIF (Fig. 2): (1) oxide BIF (Fig. 3), (2) silicate-carbonate-oxide BIF (Fig. 4), (3)
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stilpnomelane-rich tuffaceous mudrock (Figs. 5 and 6), (4) stilpnomelane mudrock (Fig.
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5A), and (5) calcareous mudrock. As shown in Fig. 2, a large portion of minerals, such as
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chert, magnetite, hematite, riebeckite, carbonate and occasionally stilpnomelane exist,
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although in different proportions within the first three lithologies. In particular, a
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transitional shift exists between oxide BIF, which is the most dominant rock type, and
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silicate-carbonate-oxide BIF, which is the second most dominant rock type. Therefore,
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any clear petrographic split between those rock types is not feasible. Magnetite is the
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most abundant iron oxide phase and occurs in variable amounts in all of the rock types
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except in calcareous mudrock.
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4.1.1. Oxide BIF
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This unit is dominated by micro- and mesobands of chert, magnetite and lesser amounts
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of hematite (Figs. 3A to 3D). Microcrystalline (~0.05 mm) chert appears both as
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mesobands (~1-5 cm) and microbands (0.25-1 mm). The chert ranges in color from white
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to greyish and to a more red variety (jasperlitic) as a result of interstital hematite grains
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(Fig. 3A). It is often found as microbands with various amount of hematite crocidolite
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carbonate, alternating with microbands of magnetite hematite. A few pure chert
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microbands that alternate with magnetite layers are observed sporadically (Figs. 3A to
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3C).
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Major parts of the magnetite bands are black and opaque while minor parts of the
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bands are dark grey as a result of interlayered chert microbands (Fig. 3A to 3C). The
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bands occur both as mesobands (1-3 cm thick) and microbands (down to 0.1 mm thick).
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The magnetite ranges from fine-grained to coarser-grained with a well-crystallized habit
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and it is commonly coarser-grained than coexisting chert and hematite (Fig. 3D).
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Hematite is found both as microcrystalline cement in relation to chert mirco- and
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mesobands (Fig. 3D) and as <0.1 mm micro-platy crystals (martite), which are locally
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found associated with magnetite meso- and microbands (Fig. 3D). Martite is formed due
304
to secondary oxidation of magnetite and is thus often found related to magnetite. Micro-
305
platy hematite is also found "floating" in a more brownish to yellowish coherent cement
306
probably of a more goethitic composition. The microcrystalline hematite and chert
307
microbands, consisting of red jasperitic mesobands, are more dominant in the middle
308
section of the Joffre BIF. Some hematite grains are associated with microbands of chert,
309
riebeckite and altered carbonate. This association is likely due to the result of micro-platy
Ac ce p
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d
M
296
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Page 14 of 96
hematite replacing part of the chert and carbonate as suggested by Clout and Simonson
311
(2005). Some sections of the oxide BIF furthermore reveal thin (<0.01 mm) microbands
312
containing brown microgranules of stilpnomelane,
313
K(Fe2+,Mg,Fe3+)8(Si,Al)12(O,OH)27·n(H2O).
ip t
310
cr
314 4.1.2. Silicate-carbonate-oxide BIF
316
This rock type is distinctive by having relatively more riebeckite and carbonate and less
317
iron oxides than the oxide BIF (Figs. 4A, 4B and 4C).
an
us
315
Chert mesobands and chert nodules display two different wavy microbands with
319
different grain size (Figs. 4D to 4F). One chert fraction comprises 0.05 mm grains
320
occasionally with a braided network of hematite and minor goethite. These wavy
321
microbands have a thickness of ~0.5 mm. The other chert fraction is finer grained,
322
consisting of <0.02 mm chert grains (Fig. 4F). These microbands are up to 1 mm thick.
323
The finer grained chert microbands are solely restricted to rhombic ankerite and fibrous
324
crocidolite (Fig. 4F). Generally, all chert grains are irregularly shaped and exhibit slight
325
undulatory extinction. Only in few places has the chert been recrystallized to coarser
326
grained (~0.3-0.4 mm) fractions.
d
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Ac ce p
327
M
318
Alongside chert, dense micro- and mesobands of magnetite are found as a major
328
constituent (Fig. 4B). A minor portion of the magnetite displays secondary alteration and
329
oxidation into hematite and other Fe oxides such as goethite (Figs. 4G and 4H).
330 331
After
chert
and
magnetite,
riebeckite,
a
sodium-rich
amphibole
(Na2(Fe2+)3(Fe3+)2(Si8O22)(OH)2), is the most abundant constituent, both within this rock
15
Page 15 of 96
type and within the entire core (Figs. 4B, 4G and 4H). It occurs as dark blue, dense
333
microbands that range from 0.1 mm to 0.5 cm in thickness and as single (0.2 mm long)
334
crystals (Fig. 4I). It is found predominantly within chert-rich microbands and at the
335
interface between chert and iron oxide microbands (Fig. 4G). Locally, riebeckite
336
microbands are associated with magnetite microbands containing randomly dispersed
337
brownish Fe oxide grains (Fig. 4H). The presence of riebeckite-rich fluid escape
338
structures (riebeckite veins in Fig. 4A) indicates remobilization from presumably a
339
hydrous silica-iron-sodium gel during compression of the sediments. Riebeckite is often
340
found associated with acicular and fibrous crocidolite (Figs. 4F and 4I). Crocidolite
341
mostly appears as thin, blue fibrous needles (blue asbestos) in close association with
342
more massive riebeckite crystals or bands. It also occurs within bands of transparent chert
343
microbands where the long fibers of crocidolite tend to grow perpendicular to the
344
bedding planes. Fibrous crocidolite tends to spray out from the rims of the denser
345
riebeckite (Fig. 4I). It coexists with the wavy chert + carbonate microbands indicating,
346
and as proposed by Miyano and Klein (1983), that the crocidolite has grown at the
347
expense of Fe-carbonate. For a more detailed description of both riebeckite and
348
crocidolite as a diagenetic product see Miles (1942), Ryan and Blockley (1965) and
349
Miyano and Klein (1983).
cr
us
an
M
d
te
Ac ce p
350
ip t
332
Various compositional pale-brown carbonates occur throughout as individual
351
crystals with a predominantly rhombic habit occasionally displaying internal zoning. The
352
carbonate crystals mostly occur within microbands composed of very fine-grained chert
353
with riebeckite and crocidolite magnetite hematite. Semi-quantitative EDS
354
observations of the carbonate reveal high elemental peaks corresponding to Ca and Fe,
16
Page 16 of 96
355
with minor peaks for Mg and Mn. Based on this, we interpret the carbonates as belonging
356
to the dolomite-ankerite series, with a predominantly ankeritic composition. Talc- and chlorite-alteration plays a minor role in the Joffre BIF. However, small
358
colorless and non-pleocroic needles of talc (presumably minnesotaite) are observed in
359
relation to green to pale-green chlorite microbands. These needles are oriented in the
360
direction of stretching, indicating that they have been growing during the main
361
compaction phase of the BIF package. In addition, talc alteration is locally visible in
362
between microlaminae of carbonate, magnetite and chert (Fig. 4J).
363
an
us
cr
ip t
357
4.1.3. Stilpnomelane-rich tuffaceous mudrock
365
This rock type is volumetrically minor but it shows important mineralogical features (see
366
Fig. 5A). One of the characteristics of this rock type is the intimate relation between tuff
367
material and stilpnomelane microgranules. The preservation of a 1-2 cm thick bed of
368
pale-green tuff represents direct evidence of volcanogenic ash-fall into the basin (Fig.
369
5B). Texturally, this bed consists of recrystallized shards set in a very fine-grained,
370
greenish-brownish, stilpnomelane matrix (Figs. 5C and 6A). The shards are feldspar-
371
pseudomorphs likely formed from volcanic glass. During compaction and burial
372
metamorphism, the glass devitrified and recrystallized into feldspar. Electron microprobe
373
analysis shows that the feldspar is almost 100% sanidine in composition, with very little
374
iron, calcium and sodium (see Table 1).
Ac ce p
te
d
M
364
375
In addition to the tuffaceous layer, stilpnomelane is cemented by chert forming
376
plane- or wavy-lamina microbands (Figs. 5B, 5D-F and 6B to 6E). These bands are 0.25-
17
Page 17 of 96
1 mm thick and alternate with microbands of almost pure chert with few grains of fibrous
378
crocidolite and platy hematite. The microbands consist of either extremely fine-grained
379
(down to 0.01 mm) stilpnomelane microgranules or spheroids (Fig. 5E) or as medium
380
grained (0.5-1 mm) stilpnomelane flake- and lath-shaped aggregates that are occasionally
381
radiating (Figs. 5F and 5G). The microgranules are pale to dark green to brown
382
(depending on the Fe2+/Fe3+ ratio). Some brown stilpnomelane-rich bands also show
383
evidence of shrinkage texture possibly caused by dehydration of a precursor phase to
384
stilpnomelane (Fig. 5F). EDS images in Figs. 6D and 6E show that elements, such as Al
385
and K, are enriched in the stilpnomelane microlaminae. As for the shards, the
386
stilpnomelane microgranules may often only be preserved where early diagenetic chert
387
formation prevented compaction of the granules (see Rasmussen et al., 2013). However,
388
throughout parts of the core, very fine and diluted laminae of these microgranules can be
389
found interbedded with chert and iron oxide microbands making it difficult to separate
390
true oxide-BIF from weak silicate dominated BIF. The composition of the microgranules
391
and the ash matrix are presented in Table 1 and Fig. 7, wherein a large part of the
392
measurements plot within the stilpnomelane field.
cr
us
an
M
d
te
Ac ce p
393
ip t
377
Ankeritic carbonate (Fe-rich dolomite) occurs as individual crystals having
394
predominantly a rhombic habit with clear internal zoning visible in some of the larger
395
(0.2-0.4 mm) crystals (Figs. 5F and 6B). As within the silicate-carbonate-oxide BIF,
396
considerable recrystallization of the carbonate crystals can be seen by the perfect
397
euhedral outline to neighbouring minerals. The large, randomly distributed rhombic
398
carbonate crystals are normally dispersed in very fine-grained chert cement, along with
399
stilpnomelane (Fig. 6B). EDS images shows that these carbonates are both Fe- and Mg-
18
Page 18 of 96
rich (Figs. 6F and 6G). The likeliness that the carbonates have grown during burial late
401
stage diagenesis and metamorphism can be seen in Fig. 6H. It shows a trail of
402
stilpnomelane grains together with single chert grains engulfed by a rhombohedral
403
carbonate crystal, indicating the secondary growth of the carbonate crystal after both
404
stilpnomelane and chert. Different styles of carbonate growth can be viewed in Figs. 5D
405
and 6C; both showing excellent preserved lamina of single prismatic carbonate crystals.
406
In contrast to the single rhombic ankerite crystals, EDS observations of these carbonates
407
reveal a more dolomitic (Mg-rich) composition. The distinct symmetry of these dolomite
408
crystals were also found by Morris (1993) in the Marra Mamba BIF (see Fig. 1B), where
409
it was suggested they were pseudomorphs after swallowtail gypsym crystals, and as such,
410
they represent shallow water in this vicinity.
M
an
us
cr
ip t
400
Another type of stilpnomelane sedimentation is the existence of ultra-thin
412
microbeds consisting of stilpnomelane with fragments of quarts, chlorite and other sheet
413
minerals (likely muscovite, Fig. 5I). The stilpnomelane is represented as both
414
groundmass and as single individual flakes.
te
Ac ce p
415
d
411
In terms of accessory phases, few randomly dispersed cubic pyrite crystals appear
416
within stilpnomelane laminae, whereas a single-crystal lamina of pyrite is seen in Fig.
417
5G. In between the ash bed and the well-preserved bed of dolomite crystals, electron
418
microprobe EDS analysis reveals a 0.5 mm thin lamina containing zircon, ilmenite,
419
monazite, pyrite and apatite (Figs. 6C and 6I).
420 421
4.1.4. Stilpnomelane mudrock
19
Page 19 of 96
Stilpnomelane occurs also as massive, almost opaque, mesobands (Fig. 5A). Detailed
423
petrography shows that the apparent structure-less bands are plane-laminated on minute
424
scale (<0.1 mm). The bands, which have a sharp base and top, vary in thickness from 0.5
425
to 2.0 cm, and contain disseminated quartz and K-feldspar fragments that seemingly
426
indicate volcanogenic provenance. The stilpnomelane mudrock bands are volumetrically
427
minor throughout the BIF but are often found in relation to chert-rich and magnetite-poor
428
sections of the core (Fig. 5A).
us
cr
ip t
422
an
429 4.1.5. Calcareous mudrock
431
At the top part of the core section, a noteworthy ~5 cm thick mesoband of calcareous
432
mudrock occurs. It contains very fine-grained white to pale-grey calcite-dolomite that
433
grades into a fine-grained greenish chloritized material. Angular fragments of carbonate
434
are dispersed within the former, whereas fragments of mostly quartz are dispersed in the
435
latter. The precursor sediment to the calcareous mudrock has erosionally truncated the
436
underlying oxide BIF.
d
te
Ac ce p
437
M
430
438
4.2. Sedimentary structures
439
Primary sedimentary structures are generally absent throughout the Joffre BIF. This is in
440
full agreement with other core sections from the Brockman Iron Formation (e.g., Trendall
441
et al., 2002). However, here we present possible current-generated sedimentary structures
442
developed prior to compaction and lithification of the BIF package. Around the middle-
443
part of the oxide BIF core (Figs. 3B and 3C), a chert rich mesoband has a wavy
20
Page 20 of 96
appearance composed of two coherent and symmetric, concave-down, structures each ca.
445
1 cm thick. The interesting observation is that the planar laminae immediately above and
446
within the trough is not disturbed by the underlying wavy bedding, making it difficult to
447
interpret it as a post-depositional process. Rather, the concave-down structures were
448
generated prior to the deposition of the above laminae. Another, although weaker
449
developed wavy structure can be seen in the upper left corner of Fig. 3C. Note the planar
450
lamination just above the wavy features. Weak sediment slumping is also seen between
451
bedding planes of chert mesobands and iron oxide mesobands (white arrow Fig. 3A). In
452
few of the magnetite bands, soft sediment deformation is evident by the presence of
453
micro-flame structures (white arrow Fig. 3B). The latter two structures are most likely of
454
pre-diagenetic origin.
M
an
us
cr
ip t
444
The chert bedding in the Joffre BIF is normally plane- to wavy laminated, but
456
occasionally the chert forms lenses or nodules (Figs. 4A and 4C). These chert nodules (up
457
to 1 cm thick) still contain internal wavy riebeckite-carbonate microbands. The features
458
most likely developed during burial metamorphism where the lateral termination
459
happened by compaction of more iron-rich bands above and below the original chert
460
layer (white arrow in Fig. 4C). The internal riebeckite-carbonate laminae get compacted
461
from the inner part of the nodule to the outer part. As such, these chert nodules should be
462
interpreted as a result of syn-compaction rather than erosional features.
Ac ce p
te
d
455
463 464
5. Bulk rock geochemistry
465
5.1. Major and trace elements
21
Page 21 of 96
Geochemical data for major and trace elements are presented in Tables 2 and 3. SiO2-
467
Fe2O3-Al2O3 and SiO2-Fe2O3-CaO+MgO ternary diagrams are shown in Figs. 8A and 8B.
468
The evolution of selected major and trace elements with depth are presented in Figs. 9A
469
to 9J. Relative to chemically precipitated elements, such as silica and iron, the Al2O3
470
content in the BIF samples (Fig. 8A) is very low (<1 wt.%). In contrast, CaO and MgO
471
are more elevated, reflecting the appearance of well-developed carbonates throughout the
472
core (Fig. 8B). A constant low input of Al (Fig. 9B), Ti (Fig. 9C) and high field strength
473
elements such, as Zr and Nb (Figs. 9K and 9L), is noteworthy. By contrast, elevated
474
abundances of these insoluble elements are found within the calcareous mudrock, the
475
stilpnomelane-rich tuffaceous mudrock and the massive stilpnomelane mudrock. More
476
variable concentrations with depth are observed for Fe, Na, Mn and REEs (Figs. 9A, 9G,
477
9I and 9J). Phosphorous (Fig. 9H) shows low to moderate concentrations throughout the
478
core except for two BIF outliers, which have up to four times as much P as the other BIF
479
samples likely as a result of very fine grained apatite.
cr
us
an
M
d
te
In terms of elemental correlations (diagrams not shown), significant R-values are
Ac ce p
480
ip t
466
481
found for Al vs. Ti and Ti vs. Zr, indicating adsorption of these components to the fine-
482
grained pelagic clay fraction. The absence of correlation between REE and P or Zr
483
suggests minor contribution from sedimentary apatite, monazite and zircon. Interesting,
484
soluble elements, such as K and Ba, are moderately correlated with more conservative
485
elements, such as Al, Ti and Nb. For example Al vs. K yields an R-value of 0.73. Sodium
486
(Na), which shares a similar degree of mobility as K and Ba, is unrelated to Al (R = -
487
0.18). The Al vs. the REE is relatively uncorrelated. However, the variation in LREE
488
with Al, represented by the Al vs. Pr, are better correlated (R = 0.67) than the variation in
22
Page 22 of 96
489
HREE with Al, as represented by Al vs. Yb (R = 0.27). Also phosphorous, is unrelated to
490
Al2O3 (R = 0.24). The maximum, minimum and average shale normalized REE patterns for the 27
492
BIF samples are displayed in Fig. 10. The absolute concentrations of REEs in the Joffre
493
BIF (average of 17 ppm) are highly elevated relative to the undelaying Dales Gorge BIF,
494
although the REE patterns in both BIF are very similar (Pecoits et al., 2009). A general
495
HREE to LREE enrichment (average (Pr/Yb)SN = 0.24), together with a pronounced EuSN
496
anomaly (average (Eu/Eu*)SN = 1.56) is observed (Fig. 10, Table 3). Part of the BIF
497
package shows a positive YSN anomaly, (Y/Ho)SN, reflecting more Y to Ho in the
498
seawater column (Fig. 10, Table 3). However, the YSN anomaly is not significant enough
499
to be clearly evident in the average REY pattern.
M
an
us
cr
ip t
491
In a primitive mantle-normalized spider diagram (Fig. 11), average Joffre BIF is
501
plotted with relevant associated lithologies. There is the same variation for many
502
elements between the BIF and the three Joffre lithologies and, although less pronounced
503
with respect to the upper continental crust. Exceptions include the soluble elements that
504
tend to be concentrated in seawater, and hence in the BIF (e.g., P, Na and Sr). Notably,
505
the P anomaly seen in the BIF is even higher than the average continental crust. The
506
negative anomalies of the Nb and Ti, two elements that are depleted in newly generated
507
continental crust, show an even larger negative anomaly in the BIF likely due to the high
508
insolubility of those elements in seawater (Fig. 11). Similarly, insoluble elements, such as
509
Th, Zr and Hf, which have positive anomalies for the associated lithologies, show a
510
trough in the spider diagram for the BIF.
Ac ce p
te
d
500
511
23
Page 23 of 96
5.2. Fe-isotopes
513
Fe isotopes, as represented by 56Fe values and their standard deviations, are presented in
514
Table 4 and are plotted with the stratigraphic depth in Fig. 12. Inserted, for comparison,
515
are average 56Fe values for common igneous rocks and average mid ocean ridge (MOR)
516
fluids (Sharma et al., 2001; Johnson et al., 2003). Isotopic fractionation that occurs during
517
Fe(II) oxidation is represented by more positive 56Fe values than the typical values for
518
igneous rocks and hydrothermal fluids. At the top part of the core, from 90 m to 220 m, a
519
large amount of the samples have positive 56Fe values averaging +0.33‰. Between ca.
520
220 m to 360 m the 56Fe values have lower values averaging -0.21‰. In the bottom part
521
of the core, from 360 m to 450 m, a change to more positive values is seen by 56Fe
522
values averaging +0.26‰.
M
an
us
cr
ip t
512
te
d
523 6. Discussion
525
6.1. Post-depositional history
526
6.1.1. Diagenetic mineral paragenesis and burial metamorphism
527
During diagenesis and burial of the precursor BIF sediment, the main factors controlling
528
the flow of metasomatic fluids are dependent on various conditions within the
529
sedimentary basin, such as a high water-to-rock ratio, the permeability, the temperature
530
gradient, and the lithostatic pressure differential, amongst others (e.g., Smith, 1980, 1982;
531
Bau, 1993). Most primary hematite has been replaced by dense magnetite bands and
532
secondarily by larger disseminated magnetite grains. The latter can partly obscure the
533
original fine-scale magnetite bedding as seen in Fig. 4H. The generation of magnetite in
Ac ce p
524
24
Page 24 of 96
BIF could have occurred either through dissimilatory Fe(III) reduction in which primary
535
ferric oxyhydroxides were reduced at the expense of organic carbon oxidation during
536
diagenesis (e.g., Konhauser et al., 2005; Li et al., 2011) or at a later stage during
537
metamorphism, for instance through the reaction between hematite and a ferrous iron
538
phase such as siderite (Miyano, 1987; Li et al., 2013),
539
8Fe(OH)3 + CH3COO- (acetate) 8Fe2+ + HCO3- + 15OH- + 5H2O
540
FeCO3 (siderite) + Fe2O3 Fe3O4 + CO2
541
Evidence for any siderite or primary hematite has not been observed in this work. Since
542
siderite is considered to be of depositional or early diagenetic origin (e.g., Ayers, 1972),
543
we infer that during progressive burial siderite reacted with hematite to form the
544
dominant magnetite phase.
M
an
us
cr
ip t
534
The timing of the ankerite-ferroan dolomite crystals seams to be of a later stage
546
than magnetite formation. By engulfing both early chert and very late stage burial-to-low
547
metamorphic stilpnomelane (Fig. 6H), these euhedral rhombic carbonates were formed
548
very late in the post-depositional story.
te
Ac ce p
549
d
545
Riebeckite in BIF is believed to be of pre- to syn-metamorphic origin that either
550
completely or partially replaced chert and chert-magnetite bands (e.g., Beukes, 1973),
551
whilst the crocodolite formation is ascribed to later regional folding generating
552
overpressured zones during vertical extension (e.g., Krapež et al., 2003). Timing of
553
riebeckite growth in the Joffre BIF was before the formation of the chert nodules (Figs.
554
4C and 4D), indicating an earlier formation stage than the main compaction in the basin.
555
However, mobilization of some riebeckite after the main compaction has resulted in thin
25
Page 25 of 96
veins as seen in Fig. 4A. The abundant riebeckite is likely related to the migration of
557
alkali-bearing solutions with high Na+ activity (see also Miyano and Klein, 1983). Na is
558
uncorrelated with immobile elements (e.g., Al2O3, TiO2, Nb, La, Zr, Hf and Th) found in
559
the fine-grained ash and stilpnomelane. This is a result of Na being hosted within
560
riebeckite and crocidolite only. The average Na2O concentration is 1.1 wt.% and shows
561
that the sodium content in the Joffre BIF lies well above other Hamersley BIF, such as
562
the Dales Gorge and the upper and lower parts of the Marra Mamba (Fig. 13). In fact, K
563
and Ba are moderately correlated to the aforementioned immobile elements, as well as
564
with various trace metals, such as V, Cr, Ni and Co, suggesting a minimum degree of K-
565
and Ba-mobilization during burial metamorphism. This indicates that even with the
566
activity of Na-bearing fluids, the degree of alteration of the primary elements has been
567
minor. Interestingly, the high riebeckite content in Joffre BIF is not seen reflected in a
568
lower silica content (see Fig. 14 and section 6.1.2), which would be expected if chert was
569
being replaced by riebeckite. Alternatively, the formation of laminated microbands of
570
dense to fibrous riebeckite restricted to chert laminae suggest that the precursor of the
571
chert was magadiite (NaSi7O13(OH)3), a sodium rich silica gel (e.g., Eugster and Chou
572
1973; Drever, 1974; Miyano and Klein, 1983; Morris, 1993). Thus, it cannot be excluded
573
that Na could have been part of the originally precipitated components of the BIF.
cr
us
an
M
d
te
Ac ce p
574
ip t
556
575
6.1.2. Supergene enrichment
576
A second means by which the BIF sediment can be altered is via supergene weathering.
577
For instance, the downward flow of meteoric oxidative fluids would cause oxidation of
578
any reduced phases in the BIF sediments, resulting in phase changes of magnetite into
26
Page 26 of 96
high-grade hematite. Depending on pH, those fluids may also have led to the dissolution
580
of carbonate minerals (low pH) or silicate minerals (high pH). Low pH solutions would
581
cause the loss of MgO, CaO, and perhaps even Al2O3 (e.g., Weeb et al., 2003). This,
582
however, is not evident in the Joffre BIF when compared to the other associated unaltered
583
BIF (as shown in Fig. 13). High pH solutions would dissolve silica resulting in chert
584
depletion and the concomitant formation of martite and high-grade hematite ore
585
formation as evident in the Mt. Whaleback and Mt. Tom Price BIF from the Hamersley
586
Group (Ewers and Morris, 1981; Taylor et al., 2001; Webb et al., 2003). These high-
587
grade-iron and low-chert-content iron formations are not features seen in the Joffre BIF
588
chemistry. In fact, the Joffre BIF resembles the more unaltered BIF from the Marra
589
Mamba and Wittenoom formations (Fig. 14). Indeed, as seen in Fig. 14, the Joffre BIF
590
has SiO2 and Fe2O3 values within the range of expected values for relatively non-
591
enriched, Hamersley style BIF. With that stated, late-diagenetic magnetite is observed
592
locally in the Joffre core to have been partly oxidized to post-metamorphic fine-grained
593
hematite (Fig. 4H). Furthermore, very fine-grained hematite-goethite grains (possible a
594
variety of martite) are developed sporadically throughout (Fig. 4I).
cr
us
an
M
d
te
Ac ce p
595
ip t
579
596
6.2. Seawater chemistry
597
6.2.1. The REE budget
598
It is generally accepted that REEs measured in Archaean and Paleoproterozoic BIF can
599
potentially mimic the REE composition in the contemporaneous ocean water at the time
600
of precipitation (e.g., Dymek and Klein, 1988; Bau and Dulski, 1996; Bolhar et al.,
27
Page 27 of 96
2004). If true, it implies that (1) all the REEs were dissolved in seawater prior to
602
precipitation, and (2) post-depositional metamorphism did not remove or add any of the
603
elements. The REE patterns all have the characteristic fractionated HREE to LREE
604
enriched patterns (Fig. 10), and this suggests an overall minor contribution from
605
terrigenous sources on the REE budget. The REE vs. the degree of crustal
606
contamination represented by the (Pr/Yb)SN ratios is plotted in Fig. 15A. The diagram
607
reflects a higher concentrations of REEs in the seawater and hence, an increase in the
608
REE contribution to the Joffre BIF basin relative to the underlying Dales Gorge BIF (see
609
Pecoits et al., 2009). A large portion of the samples from the ca. 2.5 Ga Kuruman BIF
610
(South Africa) also contain higher REE than the Dales Gorge BIF, but still less than a
611
large portion of the Joffre BIF samples (Fig. 15A). Furthermore, the Marra Mamba BIF
612
(not shown here) contains lower overall REE concentrations than the Joffre BIF (see
613
Alibert and McCulloch, 1993). Whether this is due to a higher input of submarine
614
hydrothermal fluids and/or volcanogenic input to the basin is unknown. However, the
615
REE systematics show that while all samples exhibit the characteristic fractionated shale-
616
normalised (Pr/Yb)SN < 1) seawater pattern, the portion of the Joffre BIF samples that
617
have higher amount of total REEs (Fig. 15A) appear to be related to a weak increase in
618
Al and LREEs (Fig. 15B). This, in turn, may be controlled by the same volcanic sources,
619
which are represented in the three Joffre samples with intermixed volcanogenic detritus
620
(Fig. 15B). This reveals that small proportions of pelagic ash particles may have had an
621
impact on the overall REE signature and the total REE content of the seawater.
622
Petrographically, evidence of stilpnomelane microgranules associated with both the shard
623
bearing ash bed, as well as intermixed with some of the silica- and iron-oxide microbands
Ac ce p
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601
28
Page 28 of 96
624
in the BIF (Table 1), may explain the weakly higher LREE and Al content displayed in
625
Fig. 15B. Many Precambrian chemical sediments are also evaluated on the basis of their
627
shale-normalised LaSN- and CeSN-anomalies. The cause for the anomalous behaviour of
628
La reflects enhanced stability of La in solution and, accordingly, may be related to the
629
absence of inner 4f electrons (e.g., De Baar et al., 1985). CeSN anomalies reflect the redox
630
state of the water column from which the particles precipitated. In general, oxygenated
631
marine settings show a strong negative CeSN anomaly, whereas suboxic and anoxic
632
waters lack large negative CeSN anomalies (e.g., German et al., 1991; Byrne and
633
Sholkovitz, 1996). Oxidation of Ce(III) to Ce(IV) greatly reduces Ce solubility, resulting
634
in preferential removal onto Mn-Fe oxyhydroxides, organic matter, and clay particles. In
635
contrast, suboxic and anoxic waters lack significant negative CeSN anomalies due to
636
reductive dissolution of settling Mn-Fe-rich particles. As shown in Fig. 16A, all the
637
samples plot in the two fields of no Ce anomaly, which likely reflects oxygen levels too
638
low to oxidize Ce(III) to Ce(IV), with the concomitant scavenging of Ce(IV) and a
639
resulting negative CeSN anomaly in the BIF. In contrast, a large fraction of the samples
640
show a positive LaSN anomaly, which suggests that La has been stabilized and weakly
641
fractionated relative to the other LREE in the seawater column prior to precipitation.
cr
us
an
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d
te
Ac ce p
642
ip t
626
643
6.2.2. The hydrothermal input
644
By examining Sm-Nd isotopes in the Joffre BIF, Jacobsen and Pimentel-Klose (1988)
645
obtained an average depleted Nd value of +2.1 (n=4), and, therefore, linked a large
29
Page 29 of 96
portion of the REEs in the Joffre BIF to submarine hydrothermal alteration of the
647
seafloor. Similarly, Alibert and McCulloch (1993) reported a gradual change in the Nd
648
values from the lower Marra Mamba BIF with Nd values of -0.6 to more depleted Nd
649
values of +1 for the Dales Gorge and Joffre BIFs (see Fig. 1B for stratigraphic position).
650
Alibert and McCulloch (1993) additionally made Nd mass balances suggesting that mid-
651
ocean hydrothermal fluids mixed with seawater could explain around 50% of the sourced
652
Nd (and hence the REEs) in Joffre BIF. In this regard, the hydrothermal evolution of the
653
Joffre BIF basin can be discerned by the variation of the EuSN anomaly (Eu/Eu*SN)
654
throughout the BIF sedimentary succession.
an
us
cr
ip t
646
The (Eu/Eu*)SN in modern seawater is identical to the Post-Archaean Average
656
Shale, whereas in modern submarine hydrothermal solutions Eu is enriched with
657
(Eu/Eu*)SN values > 1 (e.g., Danielson et al., 1992; Kato et al., 1998). Fig.16B shows the
658
evolution of (Eu/Eu*)SN with depth in the Joffre BIF. The graph shows that the mixture
659
of hydrothermal fluids with seawater resulted in Eu anomaly well above 1 throughout the
660
entire core depth. However, the input of hydrothermal fluids affected the BIF to varying
661
degrees. An increase in dissolved Eu2+ resulted in the increase of the (Eu/Eu*)SN, that
662
peaked between 100 m to 155 m of the BIF sequence (Fig. 15B). This section reflects a
663
larger hydrothermal input, with a maximum EuSN anomaly of ~2.1.
d
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664
M
655
665
6.3. Volcanic activity
666
6.3.1. Ash-fall tuff and microgranules
30
Page 30 of 96
Relict shards within the tuff bed of the Joffre BIF (see Figs. 5C and 6A) indicate a
668
volcanic provenance most likely from explosive magmatic eruptions (e.g., Fisher and
669
Schmincke, 1984). Well-preserved wavy lamina and lamina sets of microgranules
670
deposited on top of the graded tuff bed have been found in the upper part of the Joffre
671
BIF. The granular texture and the well-preserved shards can only exist if early formation
672
of diagenetic silica prevented later compaction of the lamina. Thin section and
673
microprobe analysis of the microgranules and the tuff matrix clearly suggest a
674
stilpnomelane phase. Thus, the stilpnomelane microgranules are linked to the ash bed
675
underneath, and as such, the microgranules most likely represent ash material that has
676
been reworked by unknown processes within the water column or at the seafloor, e.g., by
677
density currents or contourites (e.g., Krapež et al., 2003; Rasmussen et al., 2013). Direct
678
evidence for the latter processes are rare due to burial and metamorphic overprinting but
679
the preserved current generated structure developed in the oxide-BIF (Figs. 3B and 3C) is
680
most likely a result of density currents on the seafloor during sedimentation rather than a
681
post-depositional feature (e.g., Krapež et al., 2003). Those authors linked the origin of
682
locally preserved granules to hydrothermal mud that was transported, deposited and re-
683
sedimented on submarine volcanic flanks by density currents. They suggested that the
684
granular muds may have been the precursor for the BIF, rather than direct precipitates of
685
amorphous silica- and iron oxides from seawater.
Ac ce p
te
d
M
an
us
cr
ip t
667
686
Recently, an interesting model on the formation of the stilpnomelane
687
microgranules within the Dales Gorge BIF has been proposed by Rasmussen et al. (2013)
688
who suggested flocculation of an Fe(III)-rich and Al-poor hydrous silicate either in the
689
water column or on the seabed, which was subsequently reworked by density currents to
31
Page 31 of 96
form lamina sets with a basal granular bed and a more granular diluted-amorphous mud
691
lamina on top. From a geochemical perspective, the high Fe(III)-oxide content and the
692
enriched HREE (Y, Yb, Lu) relative to both primitive mantle and continental crustal
693
values (Fig. 11) strongly support that the precursor to the stilpnomelane mudrock and the
694
stilpnomelane-rich tuffaceous mudrock in the Joffre BIF reacted with the seawater prior
695
to settling, a finding in agreement with the model by Rasmussen et al. (2013). In addition
696
to iron and silica, the stilpnomelane microgranules in the Joffre BIF also contain various
697
amounts of other components, such as Al2O3, TiO2, K2O, Zr, Nb, Th and trace metals.
698
We thus propose that the precursor to the stilpnomelane microgranules in Joffre BIF may
699
have been very fine, reactive ash particles that chemically interacted with seawater,
700
thereby stripping iron from seawater prior to their deposition at the seafloor. In the Joffre
701
BIF, stilpnomelane also exists with other textures than microgranules. For example, the
702
very thin microbed containing stilpnomelane groundmass and flakes alongside quartz
703
fragments and various phyllosilicates (Fig. 5I) may represent settling of volcanic detritus
704
through the water column.
cr
us
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Ac ce p
705
ip t
690
706
6.3.2. The source of the detritus
707
Seawater composition is inevitably linked to the composition of surrounding continental
708
crust and to changes in the degree of submarine hydrothermal alteration of the seafloor.
709
As a result of a change in the overall heat regime and style of plate tectonics, major
710
geochemical changes took place at the Archaean-Proterozoic boundary (e.g., Taylor and
711
McLennan, 1981; 1985; 2009). For example, the upper continental crust and their
712
sedimentary derivatives became enriched in Large Ion Lithophile Elements (LILE) (e.g.,
32
Page 32 of 96
K, Rb), LREE, Zr, Th and Hf, while they became relatively depleted in various transition
714
metals (e.g., Cr, Ni), Fe and Mg (McLennan, 1985; Condie, 1993). Stilpnomelane hosts
715
the dominant part of the detritus in the Joffre BIF. The occurrence of stilpnomelane is
716
likely the replacement product of greenalite (a ferrous-ferric phyllosilicate of the
717
kaolinite-serpentine group), and suggested to be a key indicator of mafic volcanogenic
718
derived material (LaBerge, 1966; Winkler, 1979; Pickard et al., 2004). Unfortunately, the
719
nature of the volcanic precursor for the Joffre BIF, and other BIF in general, is difficult to
720
deduce since the ash has been deposited in a basin influenced by Fe(II)-rich seawater
721
which subsequently interacted with the highly reactive ash particles (see section 6.3.1).
722
However, the primitive mantle normalized spider diagram in Fig. 11 shows that apart
723
from overall lower abundances, a large part of the insoluble elements in the Joffre BIF
724
follow the same pattern as that for the two associated volcanic deposits (stilpnomelane-
725
rich tuffaceous mudrock and stilpnomelane mudrock) and on a broader scale, the upper
726
continental crust. Therefore, it is important to look in detail at those elements to be able
727
to characterize the provenance of the detritus.
cr
us
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M
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te
Ac ce p
728
ip t
713
The low concentrations of insoluble elements (e.g., Nb, Ti, Th, Zr, Hf) can be
729
used as a fingerprint for what lithologies were exposed to weathering on the continents
730
proximal to the BIF depositional basin. In Fig. 17 (TiO2 versus Zr), two insoluble
731
elements that remain adsorbed to the pelagic detritus are plotted for a range of local
732
lithologies that all potentially could have had an influence on the seawater chemistry. A
733
clear dominance from rhyolite type volcanics (here represented by the Woongarra
734
rhyolites) is seen for the Joffre BIF. The regression line is made on data from the Joffre
735
sequence, including the BIF, stilpnomelane-rich tuffaceous mudrock, stilpnomelane
33
Page 33 of 96
mudrock, calcareous mudrock (this study) and 4 tuffaceous mudrock samples from
737
Pickard et al. (2003). All of those samples plot intermediate between the igneous
738
lithologies and the Joffre BIF, indicating a mixture of volcanogenic material and
739
chemical sediment. Similar to the Dales Gorge S-bands (dashed regression line), the
740
massive stilpnomelane mudrock seems to be related to a more intermediate (average
741
upper continental crust) than basaltic TiO2-Zr composition. Interesting, none of the TiO2
742
and Zr in the Joffre BIF is sourced from ultramafic sources or any of the mafic rock
743
suites presented in Fig. 17. This could also explain the low concentrations of trace metals
744
in the Joffre basin (etc., Ni and Cr). The Weeli Wolli tuff, which represents volcanic
745
activity after deposition of Joffre BIF, but before the occurrence of Woongarra rhyolites
746
(see Fig. 1B), also represents a more felsic TiO2-Zr composition and indicates the
747
continuing dominance of felsic volcanic activity on the later stages of Hamersley
748
deposition.
te
d
M
an
us
cr
ip t
736
In this study, no direct evidence of any shelf-derived epiclastic material to the
750
Joffre BIF basin has been found throughout the 355 meter of core section. This is in
751
agreement with other workers who made similar observations for the Hamersley Group
752
BIF as a whole (e.g., Ewers and Morris, 1981; Morris and Horwitz, 1983; Morris, 1993).
753
Instead, it seems likely that most of the continental input to the Joffre basin during BIF
754
precipitation was through volcanic pathways in the form of pyroclastic input and not
755
from an erosive continent. If that is the case, then the lack of terrigenous clastics suggests
756
that the Joffre BIF sequence was either formed in a deep-water setting or in shallower
757
water but within a continental starved and transgressed shelf margin with only minor
758
contribution to the elemental budget from an erosive continental crust.
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749
34
Page 34 of 96
759 6.4. Iron isotopes and Fe(II) oxidation
761
The stratigraphic variations in 56Fe in the Joffre BIF (Table 4 and Fig. 12) suggest that
762
the Fe-cycle in the early Paleaoproterozoc seawater was affected by different
763
fractionation mechanisms (e.g., Steinhofel et al., 2010). For igneous rocks, the
764
isotope ratio (expressed as 56Fe) is generally unfractionated with values around
765
00.15‰, while for mid-ocean hydrothermal fluids the 56Fe is slightly negative, ranging
766
from -0.3 to -0.6‰ (Sharma et al., 2001; Johnson et al., 2003). However, at low
767
temperatures, and influenced by redox processes, it is possible to significantly fractionate
768
the heavy and lighter Fe-isotopes, yielding a range of different 56Fe values as evident in
769
the various BIF facies analyzed here (Fig. 12). The probability histogram for Joffre BIF
770
(Fig. 18A) and Dales Gorge BIF (Fig. 18B) shows the former displaying higher positive
771
56Fe values than the Dales Gorge BIF. This is surprising considering that the measured
772
56Fe (n=40) for the Dales Gorge BIF was on magnetite only, while for the Joffre BIF the
773
measured 56Fe represents bulk analyses (n=42). In contrast to Fe-carbonates, which
774
frequently exhibits negative 56Fe values (Johnson et al., 2003), secondary hematite and
775
magnetite are often the only minerals displaying positive 56Fe values (Johnson et al.,
776
2003; Rouxel et al., 2005). This implies that the main control on the bulk 56Fe signature
777
measured in Joffre BIF is likely from a ferric oxyhydroxide precursor to magnetite (e.g.,
778
ferrihydrite) since magnetite is by far the most dominating mineral. Taking all the
779
ankerite into account, a simple mass balance consideration also suggests a more skewed
780
distribution towards the positive excursion if magnetite was measured only. Bulk BIF
56
Fe/54Fe
Ac ce p
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760
35
Page 35 of 96
samples with overall positive 56Fe values can also be found in many other Archaean and
782
Palaeoproterozoic BIFs (Planavsky et al., 2012) and show that these were a sink for
783
isotopically heavy Fe. The role of secondary riebeckite on the 56Fe budget of BIF is
784
currently unknown.
ip t
781
This Fe-isotope fractionation can be attributed to either inorganic or organic
786
processes where hydrothermal controlled Fe(II) is oxidized to Fe(III) resulting in
787
precipitation of ferric oxyhydroxides with positive 56Fe values (Johnson et al., 2003).
788
Experimental work by Croal et al. (2004) showed that the above fractionation is possible
789
by the existence of anaerobic photoautotrophic Fe(II)-oxidizing bacteria that use Fe2+ as
790
an electron donor and precipitate ferrihydrite enriched in 56Fe by up to +1.5‰.
791
However, since direct oxidation of Fe2+ by free O2 would generate similar positive 56Fe
792
patterns, this "fingerprint" of anaerobic biogenic fractionation can only be predicted if
793
there is independent evidence for an anoxygenic ocean-atmosphere (Croal et al., 2004).
794
We suggest that during deposition of the Joffre BIF submarine hydrothermal injected
795
fluids, together with pyroclastic detritus, played a greater role on seawater chemistry
796
compared to continental derived epiclastic material. This is not unexpected given that the
797
Brockman Iron Formation as a whole has been linked to major plume breakouts (e.g.,
798
Barley et al., 1997) that developed in association with the emergence of new continental
799
crust during supercontinent assembly (e.g., Condie, 2005). A plume breakout would
800
cause shallower mid-ocean ridges (e.g., Ernst et al., 2004), which in turn, would lead to
801
transgressive events that submerged the shallow shelf in waters directly influenced by
802
submarine mantle degassing and hydrothermal alteration of the oceanic crust (seen
803
through the (Eu/Eu*)SN evolution in Fig. 16B). A combination of high Fe2+, abundant
Ac ce p
te
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M
an
us
cr
785
36
Page 36 of 96
reduced gases from the mantle (e.g., CH4, H2, H2S), and alteration of new oceanic crust
805
would have acted as O2 sinks and likely promoted marine anoxia in the Joffre basin. This
806
scenario could explain that the fractionated positive Fe-isotopes were a result of
807
anaerobic photosynthetic Fe(II)-oxidizing bacteria consuming lighter Fe-isotopes faster
808
than heavier. For that to happen, a stratified ocean having a large Fe(II)-rich deep water
809
pool and a shallower upper water pool where Fe(II) oxidation and enrichment of Fe
810
isotopes seems plausible. It is important to note that during diagenesis and pore water
811
interaction, it has been shown by Busginy et al. (2014) that loss of light Fe from the pore
812
waters is unlikely to generate positive Fe isotope values in the sediment.
an
us
cr
ip t
804
One interesting aspect is the 56Fe value of +0.59‰ for the BIF dominated by
814
volcanic ash-fall. This value is significantly higher than igneous rock values (Fig. 12),
815
reflecting contemporaneous Fe(II) oxidation during volcanic activity. A portion of the ash
816
formed microgranules consisting of stilpnomelane, and as such, the iron in the granules
817
must have been fractionated in the water column before deposition and reworking. The
818
depositional link between ash-fall, the stilpnomelane microgranules, and evidence of
819
oxidation fractionating the Fe is interesting in that pulses of fine-grained volcanic ash
820
may have promoted higher bacteria production throughout the upper water column
821
(photic zone), speeding up Fe(II) oxidation and precipitation of the ferric oxyhydroxides
822
that subsequently mixed with the ash-particles during deposition. Concomitantly,
823
increased biomass may have settled to the seafloor, allowing for the needed reductants to
824
dissimilatory Fe(III) reduction and magnetite formation to occur during diagenesis and
825
metamorphism (Konhauser et al., 2005; Kolo et al., 2009; Li et al., 2011, 2013). In
826
contrast, the massive stilpnomelane band in the bottom of the core, having affinities to a
Ac ce p
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M
813
37
Page 37 of 96
more basaltic composition, shows a high negative 56Fe value of -0.74, possibly
828
mirroring diagenetic pyrite with high negative 56Fe values. Today, the fertilizing effect
829
of volcanic ash and the importance of volcanism for the marine biogeochemical cycle has
830
been described (for a review see Duggen et al., 2010).
ip t
827
cr
831 6.5. The GOE: sulphur and trace metals
833
The late Archaean to early Palaeoproterozoic (~2.7-2.4 Ga) marks one of the most
834
important periods in Earth history with a number of major interlinked environmental
835
events. These include, amongst others, (1) major episodes of mantle plume activity and
836
continental assembly (Erikkson et al., 2001; Condie, 2004; Barley et al., 2005), (2) a peak
837
in BIF deposition (Isley and Abott, 1999; Bekker et al., 2010); (3) development of an
838
oxic layer in the oceans (Reinhard et al., 2009; Kendall et al., 2010), (4) evolution of
839
aerobically-respiring organisms (Eigenbrode and Freeman, 2006; Godfrey and
840
Falkowski, 2009; Konhauser et al., 2011), and ultimately (5) the oxygenation of the
841
atmosphere, the so-called Great Oxidation Event (GOE) (Holland, 2002; Bekker et al.,
842
2004).
an
M
d
te
Ac ce p
843
us
832
The GOE represents a transition in time from an atmosphere essentially devoid of
844
free oxygen (O2<10-5 times the present atmospheric level, PAL) to oxygen concentrations
845
higher than 10-5 PAL (Pavlov and Kasting, 2002; Kopp et al., 2005; Buick, 2008). The
846
GOE is best defined by a loss of mass-independent sulphur isotope fractionations (S-
847
MIF) in sedimentary rocks (Faquhar, 2001), with data from various locations worldwide
848
showing that S-MIF continued to persist in the rock record until sometime between 2.45
38
Page 38 of 96
and 2.32 Ga (Guo et al., 2009; Canfield and Farquhar, 2009). This rise of free
850
atmospheric oxygen facilitated the onset of oxidative continental weathering reactions
851
and increased the fluxes of sulphate and redox-sensitive trace elements to the oceans
852
(Canfield, 2005; Anbar et al., 2007; Frei et al., 2009; Reinhard et al., 2009; Konhauser et
853
al., 2011). Collectively, these studies show that the GOE was a protracted process that
854
took hundreds of millions of years (Lyons et al., 2014).
cr
ip t
849
The picture of how the GOE emerged has recently gained more clarity through a
856
compilation of Cr concentrations in BIF through time, which showed a significant
857
enrichment beginning at 2.45 Ga in the Weeli Wolli Formation (Konhauser et al., 2011).
858
Given the insolubility of Cr minerals, its mobilization and incorporation into BIF
859
indicates enhanced chemical weathering at that time, most likely associated with the
860
evolution of aerobic continental pyrite oxidation. Interestingly, evidence for a ‘whiff’ of
861
oxygen was previously noted for the underlying 2.5 Ga Mt. McRae shale (Fig. 1B) where
862
molybdenum (Mo) concentrations increase from <5 ppm (near the crustal level) to 40
863
ppm, and then decrease back to <10 ppm (Anbar et al., 2007). These patterns were
864
interpreted as reflecting a transient oxygenation event in the atmosphere, although this
865
view has been complicated by the suggestion that features indicative of oxidative
866
weathering in the pre-GOE rock record may instead stem from localized O2 production in
867
association with biological soil crusts and freshwater microbial mats covering riverbed,
868
lacustrine, and estuarine sediments (Lalonde and Konhauser, 2015).
Ac ce p
te
d
M
an
us
855
869
The rise in oxygen between 2.4 to 2.3 Ga permitted the increased delivery of
870
sulfate to the oceans via enhanced oxidative sulfide weathering on land. Once in
871
seawater, the sulfate had two major sinks, (1) iron sulfide precipitation as a consequence
39
Page 39 of 96
of bacterial sulfate reduction in the water column, or (2) evaporitic precipitation of
873
gypsum. In the first instance, Canfield (1993) proposed that increased levels of sulfide in
874
the oceans effectively titrated out any remaining Fe2+ in seawater, leading to the end of
875
BIF deposition. Evidence in support of higher sulfide production comes from increasing
876
fractionation between sulfur isotopes; values for
877
values (0‰) prior to around 2.45 Ga but then increase to around 25‰ after 2.45 Ga
878
(Canfield and Farquhar, 2009). In the second instance, primary sulfate evaporites are
879
rarely reported before 2.45 Ga (Schröder et al., 2008), confirming insufficient dissolved
880
sulfate availability before that time. In the Joffre BIF, neither petrographic nor
881
geochemical data support the presence of major sulfide or sulfate mineral phases; pyrite
882
has only been documented as a small mineral constituent in the associated stilpnomelane-
883
rich tuffaceous mudrock.
S/32S (34S) are centered on mantle
d
M
an
us
cr
34
ip t
872
The overall low abundances of trace metals in the Joffre BIF relative to upper
885
continental crust values (Fig. 11) reflects either (1) low solubility related to a low degree
886
of oxidative continental weathering of exposed mineral sulfides; (2) that the adjacent land
887
masses to the Joffre basin were of a composition that did not amply supply those metals -
888
for instance, a lack of ultramafic-mafic sources; or (3) continental weathering had only a
889
subordinate control on the seawater chemistry within the Joffre basin.
Ac ce p
890
te
884
In the case of Mo and Cr, their exceedingly low concentrations in the Joffre BIF
891
(Figs. 9M and 9O) suggests that their parent minerals, such as molybdenite (MoS2) and
892
chromite ([FeCr]2O4), were not significantly dissolved. As demonstrated by Anbar et al.
893
(2007), increased concentration of Mo in the 2.5 Ga Mount McRae shale of Western
894
Australia reflected a 'whiff' of oxygen before the GOE; the Mo was sourced from
40
Page 40 of 96
oxidative weathering of Mo-bearing sulfides in crustal rocks. By contrast, low levels of
896
Ni (Fig. 9N) in the Joffre BIF is in full agreement with the findings of Konhauser et al.
897
(2009) who suggested that the cooling of the upper mantle led to decreased eruption of
898
komatiite lavas (with high Ni content), reduced supply of Ni to seawater, and thus less
899
incorporation into marine chemical precipitates, such as BIF after 2.7 Ga. Despite the
900
predicted low levels of Mo, Cr and Ni in seawater at 2.46 Ga, those metals are mainly
901
controlled by the stilpnomelane suggesting that those elements were, to some extent,
902
controlled by the volcanic input best represented by the tuff material in the
903
stilpnomelane-rich tuffaceous mudrock.
an
us
cr
ip t
895
M
904 7. The Palaeoenvironment
906
In the epiclastic-starved basin, seawater composition during Joffre BIF deposition was
907
controlled by two supply systems: (1) convective upwelling of deep, hydrothermally-
908
enriched seawater to the outer continental platform, and (2) volcanic pyroclastic detritus
909
from distal volcanic centers (see Fig. 19 for a simplified model). During convective
910
upwelling, Fe(II) rich deeper waters flooded the platform. The upwelling was likely
911
discontinous because this process depends on seasonal variations in the oceanic currents
912
in the surface water as previously suggested by Morris (1993) for the Marra Mamba BIF
913
(Fig. 1B). Evidence that upwelling varied in time is reflected by the relatively large
914
variations in both the EuSN anomaly and the thickness and distribution of magnetite bands
915
in the core. Alongside Fe(II), the concentration of REY, including Eu(II), were likely to
916
be enriched in the deeper water. If seafloor plumes and submarine volcanism facilitated
917
BIF deposition (as proposed by Barley et al., 1997; Isley and Abbott, 1999), then the
Ac ce p
te
d
905
41
Page 41 of 96
deeper water was also likely enriched in reduced gases from the mantle (e.g., CH4, H2,
919
H2S) that could act as a sink for oxygen and thus promote marine anoxia. The upwelling,
920
along with pyroclastic material, brought important nutrients to the photic zone speeding
921
up the photoautotrophic Fe(II) oxidation. When upwelling ceased (Fig. 19B), the
922
precursor sediment to chert then precipitated from the surface waters due to the high
923
concentrations of dissolved silica in the oceans at that time (e.g., Konhauser et al., 2007).
924
In addition, dissolved bicarbonate or silica reacted with Fe(II) and formed fine-grained
925
siderite (Ayers, 1972; Klein and Beukes, 1989) or greenalite (Rasmussen et al., 2013),
926
respectively.
an
us
cr
ip t
918
Superimposed on this internal marine dynamic system is the pyroclastic source
928
that provided small amounts of fine-grained diluted ash particles. In the photic zone the
929
ash intermixed with the BIF precipitates leading to the different styles of stilpnomelane
930
mineralogy, that is, the two stilpnomelane-rich rock facies and those disseminated within
931
the more typical oxide facies BIF. From the above two sources a large amount, relative to
932
the underlying Dales Gorge BIF, of REY became incorporated into the iron and silica
933
precipitates on the clastic-starved continental platform (Fig. 19). The small rise in total
934
LREE and Al in about half of the samples (Fig. 15B) could thus reflect the influence of
935
very diluted ash particles on the REY budget of the general BIF precipitate (see also
936
Table 1).
Ac ce p
te
d
M
927
937
In the photic zone, enhanced productivity of anaerobic photoautotrophic Fe(II)-
938
oxidising bacteria led to heavier iron isotope fractionation in the precipitated ferric-
939
hydroxides and subsequently in the various style of magnetite bands (Fig. 19A). In
940
clastic-starved environments, without input of bioessential nutrients through erosion of
42
Page 42 of 96
the nearby landmasses, other sources must have controlled biological productivity.
942
Perhaps the delivery of fine-grained ash particles, coupled to upwelling of
943
hydrothermally-enriched seawater, were those nutrient sources? In the case of the former,
944
experiments in seawater have shown that airborne volcanic ash particles have soluble
945
coatings containing important micronutrients (Cu, Zn etc.) and macronutrients (P, K and
946
NH4+) that upon contact with seawater will be released within minutes. This implies that
947
most of the released nutrients are accessible in the photic zone (see Duggen et al., 2010
948
and references therein).
us
cr
ip t
941
This source is hard to quantify since distal tephra input can be very dispersed in
950
the atmosphere depending on the size of the eruption and numerous meteorological
951
factors. However, seafloor drilling carried out to quantify the tephra input to the Pacific
952
Ocean basin, the largest and oldest (174 Ma) ocean basin on Earth, shows that about 25
953
vol.% of the existing oceanic sediments are tephra material, half of which comes from
954
subaerial arc volcanism (Straub and Schmincke, 1998). A major portion of this material
955
is not necessary deposited as distinct ash layers but instead occurs as dispersed ash
956
particles in the marine sediments (Duggen et al., 2010). Therefore, the role of volcanic
957
ash acting as a fertilizing agent for oceanic biota may have been an underestimated factor
958
on the Precambrian Earth.
M
d
te
Ac ce p
959
an
949
960
8. Summary
961
The ca. 2.45 Ga old Joffre BIF can be subdivided into two major rock types (oxide-facies
962
and silicate-carbonate-oxide facies) and three minor rock types (stilpnomelane mudrock,
963
stilpnomelane-rich tuffaceous mudrock and calcareous mudrock). The oxide-facies is 43
Page 43 of 96
dominated by chert, magnetite and hematite, with lesser amount of riebeckite, carbonate,
965
crocidolite and stilpnomelane. The silicate-carbonate-oxide facies is dominated by chert,
966
magnetite, riebeckite and ankerite, with minor hematite, crocidolite and stilpnomelane. A
967
clear lithological boundary between these two facies types is not feasible; rather it is
968
gradational. Although the dominant parts of the sedimentary structures found are of
969
secondary origin (chert nodules, flame structures, etc.), in rare cases the oxide-facies
970
contains primary syn-depositional features presumably of current-generated origin. This
971
illustrates the existence of weak bottom (etc., density) currents within the BIF basin.
us
cr
ip t
964
There is no evidence for epiclastic material, sourced from an erosive continent,
973
within the Joffre BIF. Instead, petrographical studies show that the three minor rock types
974
all have detritus of volcanogenic origin. Furthermore, the chemostratigraphy shows that
975
soluble elements, such as K2O and Ba, co-vary with insoluble elements, such as Al2O3,
976
Ti, Zr and Nb, suggesting only minor modification of the geochemistry. This relationship
977
supports the notion that adjacent volcanoes, delivering pyroclastic sediment to the Joffre
978
BIF basin, were the main source of both insoluble and soluble elements. Indeed, the
979
stilpnomelane-rich tuffaceous mudrock consists of volcanic tuff with well-preserved
980
shards overlain by wavy laminae and laminae sets of stilpnomelane microgranules. These
981
granules most likely originated from re-worked volcanic ash formed either on the
982
seafloor or in the water column. Since the matrix within the tuff bed is of stilpnomelane
983
composition, it is likely that the felsic ash particles reacted with ferric oxyhydroxides in
984
the water column during settling, thereby gaining the iron to form stilpnomelane. This
985
notion is supported by the high 56Fe value of +0.59‰ for this rock type. In fact, detailed
986
geochemistry and petrography shows that small proportions of pelagic ash particles (now
Ac ce p
te
d
M
an
972
44
Page 44 of 96
in the form of stilpnomelane) have had a minor impact on the overall REE signature.
988
While all BIF samples exhibit the characteristic fractionated shale normalised (SN)
989
seawater pattern with (Pr/Yb)SN < 1, a large portion of the samples have high total REEs
990
(relative to other similar BIF) and weakly elevated, although still below 1, (Pr/Yb)SN
991
ratios. These correspond to a slight increase in Al and LREEs, which is directly linked to
992
the volcanic sources adjacent to the Joffre BIF basin. The TiO2-Zr ratio of Joffre BIF and
993
the mudrocks indicates a felsic source related to the same style of volcanics as the
994
slightly younger Woongarra rhyolites. We demonstrate here that the precursor to
995
stilpnomelane does not have to be an indicator of mafic volcanism but instead it could
996
have been felsic volcanic ash that interacted with Fe-rich seawater
M
an
us
cr
ip t
987
In addition to the volcanic contribution, the input of submarine hydrothermal
998
fluids to the seawater played a significant role as a source of solutes to the BIF. This is
999
most clearly evidenced by the EuSN anomaly, (Eu/Eu*)SN, which is above 1 throughout
1000
the entire succession, with a peak value of ~2.1 between 100-155 m of core depth. Within
1001
the water column, a large fraction of the Fe(II) sourced from the mid-ocean ridge
1002
environment underwent heavy isotopic fractionation where Fe(II)-oxidation and
1003
subsequently precipitation of ferric oxyhydroxides resulted in high positive 56Fe values
1004
ranging between +0.04‰ to +1.21‰ (average +0.46‰). This process seems to have been
1005
more pronounced relative to the underlying Dales Gorge BIF.
Ac ce p
te
d
997
1006
Evidence for elevated abundances of sulfur and redox sensitive trace metals (e.g.,
1007
Mo, Cr) have not been found in the BIF, presumably implying a low degree of oxidative
1008
continental weathering. This is not surprising given the lack of epiclastic components and
1009
the fact that the dominant control on the detritus was from felsic volcanics.
45
Page 45 of 96
Correspondingly, it is clear that the Joffre BIF is poorly suited as a chemical proxy for
1011
the study of atmospheric oxygen and its weathering impact on local landmasses.
1012 1013 1014
9. Acknowledgement
1015
We are grateful to Rio Tinto in Australia for providing access to their core samples. A
1016
number of colleagues at the University of Alberta, including Mark Labbe, Martin Von
1017
Dollen, Ilona Ranger, Tom Chako, Andrew Locock and Igor Jakab, are highly
1018
appreciated for their help, as well as Birger Rasmussen for instructive discussions
1019
concerning BIF petrology. The Natural Sciences and Engineering Research Council of
1020
Canada (NSERC) supported this work.
1021
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1273
Pickard, A., 2003. SHRIMP U–Pb zircon ages for the Palaeoproterozoic Kuruman Iron
1275
Formation, Northern Cape Province, South Africa: evidence for simultaneous BIF deposition on
1276
Kaapvaal and Pilbara Cratons. Precambrian Research 125, 275–315.
d
te
1277
M
1274
Pickard, A. L., Barley, M. E., Krapež, B., 2004. Deep-marine depositional setting of banded iron
1279
formation: sedimentological evidence from interbedded clastic sedimentary rocks in the early
1280
Palaeoproterozoic Dales Gorge Member of Western Australia. Sedimentary Geology 170, 37–62.
1281
Ac ce p
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1282
Pecoits, E., Gingras, M. K., Barley, M. E., Kappler, A., Posth, N. R., Konhauser, K. O., 2009.
1283
Petrography and geochemistry of the Dales Gorge banded iron formation: Paragenetic sequence,
1284
source and implications for palaeo-ocean chemistry. Precambrian Research 172, 163–187.
1285 1286
Planavsky, N., Rouxel, O. J., Bekker, A., Hofmann, A., Little, C. T. S., Lyons, T. W., 2012. Iron
1287
isotope composition of some Archean and Proterozoic iron formations. Geochimica et
1288
Cosmochimica Acta 80, 158–169.
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1289 1290
Rasmussen, B., Meier, D., Krapež, B., Muhling, J., 2013. Iron silicate microgranules as precursor
1291
sediments to 2.5-billion-year-old banded iron formations. Geology 41, 435–438.
ip t
1292 Rasmussen, B., Krapez, B., Meier, D. B., 2014. Replacement origin for hematite in 2.5 Ga
1294
banded iron formation: Evidence for postdepositional oxidation of iron-bearing minerals.
1295
Geological Society of America Bulletin 126 (3-4), 438–446.
us
1296
cr
1293
Rasmussen, B., Krapež, B., Muhling, J. R., 2015. Seafloor silicification and hardground
1298
development during deposition of 2.5 Ga banded iron formations. Geology 43, 235–238.
an
1297
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Reinhard, C. T., Raiswell, R., Scott, C., Anbar, A. D., Lyons, T. W., 2009. A Late Archean
1301
sulfidic sea stimulated by early oxidative weathering of the continents. Science 326, 713–716.
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1300
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Rouxel, O. J., Bekker, A., Edwards, K. J. 2005. Iron isotope constraints on the Archean and
1304
Paleoproterozoic ocean redox state. Science 307, 1088–91.
1305
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1306
Rouxel, O. J., Shanks III, W. C., Bach, W., Edwards, K. J., 2008. Integrated Fe- and S-isotope
1307
study of seafloor hydrothermal vents at East Pacific Rise 9–10°N. Chemical Geology 252, 214–
1308
227.
1309 1310
Ryan, G. R., Blockley, J. G., 1965. Progress report on the Hamersley blue asbestos survey:
1311
Western Australia Geol. Survey Record No. 1965/32, (unpublished open file report).
1312 1313
Schröder, S., Bekker, A., Beukes, N. J., Strauss, H., van Niekerk, H. S., 2008. Rise in seawater
1314
sulphate concentration associated with the Paleoproterozoic positive carbon isotope excursion:
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1315
evidence from sulphate evaporites in the 2.2–2.1 Gyr shallow-marine Lucknow Formation, South
1316
Africa. Terra Nova 20, 108–117.
1317 Sharma M., Polizzotto M., Anbar A. D. 2001. Iron isotopes in hot springs along the Juan de Fuca
1319
Ridge. Earth and Planetary.Scince Letters 194, 39–51
ip t
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cr
1320
Shimizu H., Umemoto N., Masuda A., Appel P. W. U., 1990. Sources of iron-formations in the
1322
Isua and Malene supracrustals, West Greenland: Evidence from La-Ce and Sm-Nd isotopic data
1323
and REE abundances. Geochimical Cosmochimical Acta 54, 1147–1154.
us
1321
an
1324
Straub S. M., Schmincke H. U., 1998. Evaluating the tephra input into Pacific Ocean sediments:
1326
Distribution in space and time. Geologische Rundschau 87, 461–476.
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1325
1327
Simonson, B. M., Goode, A. D. T., 1989. First discovery of ferruginous chert arenites in the early
1329
Precambrian Hamersley Group of Western Australia. Geology 17, 269-272.
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1331
Smith, R., Perdrix, J., Parks, T., 1982. Burial metamorphism in the Hamersley basin, Western
1332
Australia. Journal of Petrology 23, 75–102.
1333 1334
Steinhofel, G., von Blackenburg, F., Horn, I., Konhauser, K. O., Beukes, N., and Gutzmer, J.,
1335
2010. Deciphering formation processes of banded iron formations from the Transvaal and the
1336
Hamersley Sequence by combined Si and Fe isotope analysis using UV femtosecond laser
1337
ablation. Geochimica et Cosmochimica Acta 74, 2677-2696.
1338
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1339
Taylor, S. R., McLennan, S. M., 1981. The Composition and Evolution of the Continental Crust:
1340
Rare Earth Element Evidence from Sedimentary Rocks. Philosophical Transactions of the Royal
1341
Society of London 301, 381-399.
ip t
1342 Taylor, S. R., McLennan, S. M., 1985. The Continental Crust: Its Composition and Evolution.
1344
Blackwell Scientific, Oxford, 312 pages.
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1343
1345
Taylor, S. R., McLennan, S. M., 2009. Planetary Crusts: Their composition, origin and evolution.
1347
Cambridge University Press, New York, 378 pages.
an
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us
1346
Taylor, P., Maeck, R., De Bievre, P., 1992. Determination of the absolute isotopic composition
1350
and atomic weight of a reference sample of natural iron. International Journal of Mass
1351
Spectrometry Ion Processes 121, 111–125
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1349
Taylor, D., Dalstra, H., 2001. Genesis of high-grade hematite orebodies of the Hamersley
1354
Province, Western Australia. Economic Geology 96, 837–873.
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te
1353
1356
Trendall, A. F., Compston, W., Nelson, D. R., De Laeter, J. R., Bennett, V. C., 2004. SHRIMP
1357
zircon ages constraining the depositional chronology of the Hamersley Group, Western Australia.
1358
Australian Journal of Earth Sciences 51, 621-644.
1359 1360
Trendall, A. F., 1968. Three Great Basins of Precambrian Banded Iron Formation Deposition: A
1361
Systematic Comparison. Geological Society of America Bulletin 79, 1527-1544.
1362 1363
Trendall, A. F., Blockley, J. G., 1970. The iron-formations of the Precambrian Hamersley
1364
Group,Western Australia. Geological Survey Western Australia Bulletin 119.
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1365 Trendall, A. F., Blockley, J. G., 2004. Precambrian iron-formation. In: Eriksson, P.G., Altermann,
1367
W., Nelson, D.R., Mueller, W.U., Catuneanu, O. (Eds.), The Precambrian Earth: Tempos and
1368
Events. Developments in Precambrian Geology, vol. 12. Elsevier, Amsterdam, p. 403–421.
ip t
1366
1369
Van Hise, C. R., Leith, C. K., 1911. The geology of the Lake Superior region: U.S. Geological
1371
Survey Monograph 52, 641 pages.
cr
1370
us
1372
Webb, A. D., Dickens, G. R., Oliver, N. H. S., 2003. From banded iron-formation to iron ore:
1374
geochemical and mineralogical constraints from across the Hamersley Province, Western
1375
Australia. Chemical Geology 197, 215–251.
M
1376
an
1373
Wilke, M., Schmidt, C., Dubrail, J., Appel, K., Borchert, M., Kvashnina, K., Manning, C. E.,
1378
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1379
pressure and temperature. Earth and Planetary Science Letters 349-350, 15-25.
1381 1382 1383
te
Ac ce p
1380
d
1377
Captions
1384
Table 1. Major element oxides measured on the microgranules in the BIF (DD98-26A)
1385
and in the stilpnomelane-rich tuffaceous mudrock (DD98-6) and the composition of the
1386
shards and the shard matrix within the ash bed in DD98-6. Note final wt.% with H2O
60
Page 60 of 96
1387
calculated from OH content. The composition of the mictogranules are illustrated in
1388
Figure 7.
ip t
1389 Table 2. The major elements as oxides with core depth in the Joffre BIF. Note the three
1391
different lithologies.
cr
1390
us
1392
Table 3. Selected trace elements and relevant REY anomalies with core depth in the
1394
Joffre BIF. Note the three different lithologies.
an
1393
M
1395
Table 4. Fe isotope values (expressed as 56Fe), and their individual uncertainties, with
1397
depth in the Joffre BIF. Note the three different lithologies.
te
1398
d
1396
Figure 1. (A) The general geology of the Pilbara craton with the location of drill hole
1400
DD98. See text for further explanation. (B) The general stratigraphy of the extensive
1401
banded iron formations and associated rocks of the Hamersley Group. Significant other
1402
BIFs are the Marra Mamba BIF, the Dales Gorge BIF and the BIF of Weeli Wolli
1403
Formation. Note the important U-Pb zircon ages from individual tuff layers. Note the age
1404
of 2454 3 Ma for the upper part the Joffre BIF.
Ac ce p
1399
1405 1406
Figure 2. The main rock types and their mineralogy in the Joffre BIF. The major
1407
minerals are furthermore listed in increasing abundances throughout the core.
1408 61
Page 61 of 96
Figure 3. Photos and photomicrographs illustrating the main petrographic characteristics
1410
of the oxide BIF with representative core samples (A and B) showing pale grey micro-
1411
and mesobands of alternating chert and magnetite, dark grey magnetite mesobands and
1412
reddish and bluish (<<1 mm thin) microbands of chert-hematite-riebeckite crocidolite.
1413
With white arrows, sedimentary slumping can be seen in the upper half of (A) whilst soft
1414
sediment deformation in a magnetite microband is seen by the presence of micro-flame
1415
structures in the top part of (B). Possible current generated sedimentary structures are
1416
presented in (C). The fact that the planar lamina immediately above the wavy band is
1417
undisturbed indicates a primary origin of the two symmetric wavy features. Note also the
1418
two wavy features with internal wavy bedding in the upper left corner of (C). (D) Coarse
1419
and fine grained magnetite microbands with chert and very fine hematite. The latter likely
1420
is a product of secondary magnetite oxidation. Note the single grains of martite upper
1421
left. Ch = chert, Mag = magnetite, Hem = hematite, Rbk = riebeckite.
cr
us
an
M
d
te
1422
ip t
1409
Figure 4. Photos and photomicrographs illustrating the main petrographic characteristics
1424
of the silicate-carbonate-oxide BIF with representative core samples (A and B) and thin
1425
section images (C-J). This rock type is dominated by chert + magnetite + riebeckite +
1426
ankerite + crocidolite stilpnomelane. (A) and (B) Dense mesobands composed of
1427
alternating microbands of chert + riebeckite + crocidolite oxide can be seen in the
1428
middle part of (A) and the upper half of (B). Dense, dark grey magnetite micro- and
1429
macrobands and pale grey chert mesobands with magnetite microbands are evident in
1430
core (B), which furthermore have mesobands of riebeckite with chert and hematite
1431
microbands (lower part). The white arrows in (A) illustrates weak mobilisation of
Ac ce p
1423
62
Page 62 of 96
riebeckite microbands. (C) Thin section PPL image (with inset close up image) of planar
1433
riebeckite microbanding. (D) Chert mesobands and nodules containing wavy-laminated
1434
chert and riebeckite+carbonate+crocidolite microbands. The distinct chert nodules were
1435
likely formed during compaction of the BIF package illustrated with the condensed
1436
internal microbands at the white arrow. These internal laminae are further illustrated in
1437
(E) and (F). Note in (E) the finer grained chert fraction is restricted to the carbonate and
1438
riebeckite-crocidolite laminae only with coarser grained chert in between. (G) PPL thin
1439
section image of blue riebeckite microbands underlain by red hematite and opaque
1440
magnetite microbands. (H) Riebeckite and magnetite microbands in relation to very fine-
1441
grains of other oxides (presumably hematite and goethite). (I) Close up PPL image of the
1442
fibrous bluish crocidolite in a cherty groundmass. (J) Fe-talc alteration occur occasionally
1443
in relation with magnetite. Ch = chert, Mag = magnetite, Hem = hematite, Rbk =
1444
riebeckite, Ank = ankerite, Cr = crocidolite, Ox = oxides (hematite/goethite).
cr
us
an
M
d
te
1445
ip t
1432
Figure 5. Photos and photomicrographs illustrating the main petrographic characteristics
1447
of the stilpnomelane-rich tuffaceous mudrock and stilpnomelane mudrock. (A) Typical
1448
core section influenced by various brown to greenish stilpnomelane mudrock mesobands
1449
alternating with thick white chert bands with internal wavy microbands of stilpnomelane
1450
+ ankeritic dolomite. Note only relative sparse representation of iron oxides and
1451
magnetite. Riebeckite microbands are visible in the middle and lower part of the core
1452
section. (B) Whole thin section PPL image showing stilpnomelane-rich tuffaceous
1453
mudrock having a bottom greenish tuff bed grading up in to plane- and wavy lamination
1454
of stilpnomelane microgranules ending with a magnetite band intermixed with
Ac ce p
1446
63
Page 63 of 96
stilpnomelane. (C) Close up XPL image of the tuff bed in (B) containing well-preserved
1456
shards in a stilpnomelane-rich matrix. (D) A PPL image of the wavy stilpnomelane
1457
laminae set from (B). Each laminae set ends with almost pure laminae of chert. (E) A
1458
close up PPL image of the single crystal dolomite bed in (B) with the stilpnomelane rich
1459
tuff bed underneath and wavy stilpnomelane microgranular laminae above. (F) A close
1460
up PPL image of the stilpnomelane microgranules. (G) PPL image of chert with
1461
stilpnomelane laminae and randomly distributed euhedral ankerite crystals. Note
1462
stilpnomelane occur both as microgranules and as flakes. (H) Stilpnomelane and chert
1463
laminae containing radiating flakes and silt size shrinkage texture of stilpnomelane. A
1464
single laminae of pyrite crystals is seen in the middle part of the image. (I) Ultra-thin
1465
microbed (in PPL) of stilpmomelane rich bearing quartz + chlorite + other sheet silicates
1466
(possible muscovite) illustrating another style of volcanic input to the BIF basin than in
1467
(B). The bed is from the oxide BIF. Ch = chert, Mag = magnetite, Hem = hematite, Rbk =
1468
riebeckite, Ank = ankerite, Cr = crocidolite, Ox = oxides (hematite/goethite), Stp =
1469
stilpnomelane, Chl = chlorite, Qtz = quartz, Py = pyrite, Dol = dolomite.
cr
us
an
M
d
te
Ac ce p
1470
ip t
1455
1471
Figure 6. Backscatter electron images and microprobe elemental maps of various
1472
textures and minerals from the stilpnomelane-rich tuffaceous mudrock. (A) Backscatter
1473
image of Figure 5C, showing dark K-feldspar shards in a stilpnomelane-rich matrix. (B)
1474
Well-crystallised Ankeritic dolomite with internal zoning distributed in chert with
1475
stilpnomelane microgranules. (C) Tuff bed overlain by perfect crystal-shaped dolomite
1476
laminae with homogenous stilpnomelane and dolomite laminae on top. The dominating
1477
wavy-laminae of chert and stilpnomelane microgranules are seen in the upper part. (D
64
Page 64 of 96
and E) Elemental maps showing the distribution of Al and K across wavy micro-laminae
1479
of stilpnomelane and chert. Bright blue (D) and bright green to yellow (E) represent
1480
stilpnomelane-rich laminae having relative higher Al and K compare with darker chert
1481
rich laminae. (F and G) Shows distribution of Fe and Mg across micro-laminae of chert
1482
and stilpnomelane with euhedral ankeritic dolomite crystals. Note the zonation of higher
1483
Fe content (brighter bluish) in the crystals in (F) and the relative high content of Mg
1484
(brighter greenish colors) in (G). Note also brighter zones with respect to Fe and Mg are
1485
evident for the stilpnomelane-rich laminae. (H) A single ankerite crystal that encapsulates
1486
a laminae of single-grain stilpnomelane proving a very late growth of the carbonate. Note
1487
also the engulfment of various amount of chert grains. (I) A close up backscatter image of
1488
the bright laminae in (C). EDS analyses show that this laminae contains various amount
1489
of heavy minerals such as zircon, monazite, ilmenite, pyrite and also apatite (not shown).
1490
Ch = chert, Mag = magnetite, Rbk = riebeckite, Ank = ankerite, Stp = stilpnomelane, Py
1491
= pyrite, Dol = dolomite, Zrn = zircon, Mnz = monazite, Ilm = ilmenite.
cr
us
an
M
d
te
Ac ce p
1492
ip t
1478
1493
Figure 7. Electron microprobe data from the stilpnomelane-rich tuffaceous mudrock
1494
showing a SiO2-FeO+MgO-Al2O3+K2O ternary diagram with 27 measurements of the
1495
microgranules and 10 of the ash matrix (see also Table 4). Almost all of the
1496
measurements fall within the stilpnomelane compositional field. The stilpnomelane,
1497
minnesotaite and greenalite fields are composed from microprobe data from the Marra
1498
Mamba BIF (Klein and Gole, 1981) and crosschecked with mineral data from James
1499
(1954).
1500
65
Page 65 of 96
1501
Figure 8. Major bulk element data from the Joffre BIF plotted in ternary diagrams. (A)
1502
SiO2-Fe2O3(t)-Al2O3 and (B) SiO2-Fe2O3(t)-CaO+MgO. See text for further explanation.
1503 Figure 9. Evolution with depth for selected major and trace elements. See text for details.
ip t
1504 1505
Figure 10. Shale normalised (PAAS) REY pattern from the Joffre BIF (grey area
1507
represents 27 samples) showing pronounced fractionated HREE to LREE pattern along
1508
with a pronounced EuSN anomaly, (Eu/Eu*)SN. Note the high abundances of REY relative
1509
to the underlying Dales Gorge BIF. Average data for Dales Gorge BIF from Pecoits et al.
1510
(2009).
an
us
cr
1506
M
1511
Figure 11. Primitive mantle normalised spider diagram showing the Joffre BIF and the
1513
intermixed massive stilpnomelane mudrock and the stilpnomelane-rich tuffaceous
1514
mudrock. Blue line shows the average upper continental crust from Rudnick and Gao
1515
(2003). For the BIF, note the opposite pattern with very low abundances among major
1516
parts of the insoluble elements (Th, Nb, Zr, Hf, Ti) and the distinctive positive anomalies
1517
of soluble elements (etc., P, Na and Sr) relative to the stilpnomelane mudrock and
1518
stilpnomelane-rich tuffaceous mudrock. Note also the high positive plateau for the
1519
HREEs (Y, Yb, Lu) both for the BIF and for the intermixed tuffaceous detritus. A
1520
significant drop in the trace metals (V, Cr and Ni) is seen in all of the lithologies.
Ac ce p
te
d
1512
1521 1522
Figure 12. Shows the evolution of the Fe isotopes (56Fe) throughout the Joffre BIF.
1523
Fractionation mechanism during Fe oxidation is represented by more positive 56Fe
66
Page 66 of 96
values than the typical values for igneous rocks and mid ocean ridge (MOR) fluids.
1525
Interesting to note is that the stilpnomelane-rich tuffaceous BIF has a high positive 56Fe
1526
value of 0.74. See text for further details. Field of igneous rocks and MOR fluids from
1527
Sharma et al. (2001); Johnson et al. (2003).
ip t
1524
1528
Figure 13. Average major element plot for some of the least altered BIF from the
1530
Hamersley Group. Wittenoom BIF from Webb et al. (2003); Marra Mamba BIF from
1531
Klein and Gole (1981). Notice the high amount of Na2O in Joffre BIF compared with the
1532
other BIFs.
an
us
cr
1529
1533
Figure 14. Bar charts illustrating the same BIFs as in Figure 6 but compared with one of
1535
the classic altered BIF from Mt. Tom Price. A gain of iron and a loss of silica will be a
1536
natural outcome during supergene enrichment. This is not seen for the other BIFs in
1537
general and for the Joffre BIF in particularly. Note the "sum-to-100" problem due to CO2
1538
and H2O in some of the BIFs. Data from Mt. Tom Price BIF from Taylor et al. (2001).
d
te
Ac ce p
1539
M
1534
1540
Figure 15. (A) The REE as a function of (Pr/Yb)SN for both the Joffre BIF, the Dales
1541
Gorge BIF and the same style 2.5 Ga old Kuruman BIF (South Africa). A large portion of
1542
the Joffre BIF samples have high input of REE and are weakly elevated in (Pr/Yb)SN
1543
values. The field of intermixed volcanogenic detritus is defined by the three Joffre
1544
samples (stilpnomelane-rich tuffaceous mudrock, stilpnomelane mudrock and calcareous
1545
mudrock). (B) Shows the similar trend as in (A) but here with Al vs. LREE. The same
1546
samples have an increase in Al and LREE likely as a consequence of ash particles mixed
67
Page 67 of 96
1547
with the BIF. Dales Gorge BIF from Pecoits et al. (2009); Kuruman BIF from Bau and
1548
Dulski (1996).
1549 Figure 16. (A) Graph illustrating shale-normalised depletion/enrichment of La and Ce.
1551
Most of the BIF samples plot within the field of (Ce/Ce*)SN < 1 and (Pr/Pr*)SN ~ 1,
1552
meaning positive La anomaly and no Ce anomaly (diagram modified from Bau and
1553
Dulski, 1996). (B) Shows the hydrothermal evolution (represented by the EuSN anomaly
1554
(Eu/Eu*)SN) during precipitation of Joffre BIF. Note the steady increase of the EuSN
1555
anomaly from the bottom to the top of the core with a peak of ~2.1 around 250 m depth.
an
us
cr
ip t
1550
1556
M
1557
Figure 17. TiO2-Zr, log-log plot showing the Joffre BIF and the relation of these
1559
elements to other voluminous lithologies that may have had an influence on the seawater
1560
chemistry. In conjunction with Joffre tuffaceous mudrock (from Pickard et al., 2003), the
1561
stilpnomelane-rich tuffaceous mudrock, the calcareous mudrock and the stilpnomelane
1562
mudrock, the Joffre BIF produce a (power) regression that is linked to a rhyolite-only-
1563
source represented by the Woongarra rhyolites. In contrast, a more bimodal TiO2-Zr
1564
contribution, best represented by the Dales Gorge S-bands, which clearly have affinities
1565
to average continental crust. The Joffre tuffaceous mudrock and Woongarra rhyolites
1566
from Barley et al. (1997) and Pickard et al. (2003); original linear regression line of Dales
1567
Gorge S-bands (Zr = 244TiO2(wt.%) - 2.1) from Ewers and Morris (1981); field of
1568
submarine komatiite (3.2 Ga Ruth Well Fm.), submarine basalt (2.72 Ga Kylena basalt),
1569
subaerial basalt on cont. crust (2.69 Ga Medina basalt) from Arndt et al. (2001); field of
Ac ce p
te
d
1558
68
Page 68 of 96
pillow basalt (upper Fortescue Group), flood basalt (2.78 Ga Mt. Roe Fm.) from Nelson
1571
et al. (1992); field of dolerite and tuff (Weeli Wolli Fm.) from Barley et al. (1997) and
1572
Pickard et al. (2003); Dales Gorge tuff (S13 and S15) from Pickard et al. (2003); average
1573
upper continental crust from Taylor and McLennan (2009); average Fortescue shale from
1574
Taylor and McLennan (1981).
ip t
1570
cr
1575
Figure 18. 56Fe histogram of Joffre BIF compared with that of Dales Gorge BIF. The
1577
Joffre BIF is 42 analyses of whole rock whilst the Dales Gorge is 40 analyses on
1578
magnetite alone. Despite that fact, the Joffre BIF shows a more skewed distribution
1579
towards more positive 56Fe values, which indicate higher rate of iron oxidation in the
1580
contemporaneous seawater. See text for further interpretation. Values for Dales Gorge
1581
BIF are from Rouxel et al. (2005).
an
M
d te
1582
us
1576
Figure 19. A simplified palaeoenvironmental model for the formation of the Joffre BIF.
1584
(A) Represent the situation during formation of ferric-hydroxide precipitation from the
1585
photic zone. This scenario has nutrient-rich upwelling onto the platform from deeper
1586
waters influenced by hydrothermal activity. Together with diluted fine-grained ash
1587
particles this increase the nutrient level in the surface water speeding up the Fe(III)
1588
oxidation through photoautotrophic Fe(II) oxidation and leaves a positive 56Fe signature
1589
in the sediment. (B) Show the general situation during silica and carbonate (siderite)
1590
formation. The shut off of the upwelling resulted in relative pure deposition of Si(OH)4
1591
and FeCO3 from the upper seawater. See text for further explanations. The concentration
1592
values of Fe(II) in the upper and deeper seawater is taken from Morris (1993).
Ac ce p
1583
69
Page 69 of 96
DD98-26A* (Microgranules) 1
2
3
4
5
SiO2 TiO2 Al2O3
46.24
47.02
45.57
45.65
0.00
0.06
0.01
2.85
2.85
FeO Fe2O3 Cr2O3
26.07
MnO MgO CaO Na2O K2O H2Ocalc
Ave
Ave (n=22)
Ave (n=40)
45.36
45.97
46.54
0.01
0.02
0.02
0.03
2.77
2.86
2.71
2.81
26.35
25.48
25.32
26.20
25.88
25.35
7.24
7.32
7.08
7.03
7.28
7.19
7.04
0.02
0.02
0.00
0.01
0.00
0.01
0.12
0.15
0.13
0.13
0.19
0.14
6.86
6.83
7.13
6.83
7.35
0.03
0.04
0.05
0.03
0.03
0.21
0.15
0.32
0.54
0.20
2.07
1.85
2.48
3.62
64.34 0.01 18.30 0.56 0.02 0.04 0.00 0.05 16.59
8.69
8.80
8.53
8.10
100.40
101.44
cr
4.67
0.01
us
0.09
7.00
5.70
0.04
0.09
0.28
0.37
1.99
2.40
2.83
8.96
8.62
8.05
99.54 100.13 100.29 100.36
100.79
99.91
Ac ce p
te
d
* BIF, ** Stilpnomelane-rich tuffaceous mudrock
ip t
(Shards)
an
Total
DD98-6**
M
Element
DD98-6** (Microgranules)
Page 70 of 96
DD98-6** (Shard matrix)
2
3
4
5
6
7
8
9
10
Ave
47.77
46.41
48.52
47.57
47.40
47.38
49.14
50.94
48.12
46.99
48.02
0.11
N.D.
0.05
0.02
0.06
0.02
0.09
N.D.
0.04
0.03
0.04
5.19
4.43
5.66
4.45
4.47
4.51
5.93
7.22
4.92
4.20
5.10
21.43
22.50
20.66
22.72
23.00
22.93
21.06
17.49
22.50
23.26
21.75
5.95
6.25
5.74
6.31
6.39
6.37
5.85
4.86
6.25
6.46
6.04
0.01
0.00
0.00
0.00
0.05
0.00
0.04
0.03
0.00
0.03
0.02
0.07
0.07
0.05
0.07
0.04
0.06
0.05
0.07
0.04
0.01
0.05
7.03
6.95
6.51
7.18
7.18
7.58
6.58
0.04
0.04
0.04
0.05
0.05
0.03
0.04
1.28
0.89
0.76
0.69
0.98
1.12
0.95
3.85
2.94
4.21
2.77
2.66
2.85
7.34
7.60
7.26
7.83
7.82
7.87
100.08
98.07
99.45
99.66
us
cr
ip t
1
7.28
7.46
6.93
0.04
0.05
0.04
0.04
0.84
0.55
0.64
0.87
an
5.58
4.10
5.43
3.70
3.26
3.58
7.39
6.56
7.79
7.93
7.53
99.05 101.25 100.31 99.98
Ac ce p
te
d
M
100.08 100.72 101.21
Page 71 of 96
Element oxide
SiO2 (wt.%) Al2O3 Fe2O3 MnO
MgO CaO
Na2O K2O
TiO2 P2O5 LOI
Detection Limit
0.01
0.01
0.01
0.001
0.01
0.01
0.01
0.01
0.01
0.01
94.0 DD98-1
46.31
0.62
59.71
0.012
2.21
1.3
0.34
1.17
0.04
0.01
-11.3
100.4
110.8 DD98-3B
60.44
0.26
46.76
0.001
2.18
0.11
3.52
0.26
0.03
0.04
-13.4
100.2
123.0 DD98-5A*
17.33
3.35
6.07
0.114
5.31
33.01
0.17
1.82
0.12
0.05
32.31 99.65
126.8 DD98-6**
65.76
3.86
31.9
0.07
2.49
1.48
0.19
3.31
0.09
0.04
-9.52
99.67
197.0 DD98-10A
72.95
0.29
32.6
0.026
2.91
0.69
2.82
0.37
0.03
0.04
-12.2
100.5
208.0 DD98-12
45.38
0.08
44.6
0.035
2.43
1.05
3.09
0.23
0.03
0.1
3.49
100.5
216.8 DD98-13B
48.05
0.61
43.5
0.07
1.64
1.33
0.57
0.71
0.04
0.04
3.26
99.82
221.5 DD98-14B
59.52
0.57
34.79
0.033
1.7
0.17
1.43
0.89
0.02
0.02
1.26
100.4
209.6 DD98-15B
35.57
0.66
48.37
0.197
2.77
3.95
0.23
0.99
0.04
0.46
6.74
99.98
232.8 DD98-16
37.46
0.55
50.33
0.185
1.58
3.18
0.86
0.57
0.04
1.86
3.51
100.1
270.2 DD98-20A
32.7
0.89
49.29
0.149
3.62
3.77
0.13
1.12
0.05
0.01
8.17
99.89
270.6 DD98-20B
76.4
0.10
20.39
0.011
0.97
0.06
1.35
0.28
0.02
0.03
0.9
100.5
281.6 DD98-21A
59.94
0.73
33.37
0.089
1.32
1.37
0.3
0.61
0.04
0.46
2.19
100.4
411.2 DD98-26A
64.83
0.04
33.3
0.025
0.54
0.82
0.01
0.08
0.01
0.15
0.26
100.1
435.0 DD98-28***
32.59
3.90
32.43
0.504
5.77
5.98
0.26
2.76
0.18
0.07
15.99 100.4
444.0 DD98-29A
23.45
0.14
68.96
0.016
1.39
2.94
0.23
0.26
0.02
0.01
2.79
100.2
448.5 DD98-30A
38.72
0.15
55.37
0.009
1.48
1.16
0.35
0.44
0.02
0.03
2.08
99.81
Depth (m)
Total
Ac c
ep
te
d
M
an
us
cr
ip
t
* Calcareous mudrock, ** Stilpnomelane-rich tuffaceous mudrock, *** Stilpnomelane mudrock
Page 72 of 96
Element
Al (wt.%) Fe
Mn
P
K
0.030 0.5
5
6
Ti (ppm) V
Cr
Ni
Zn
0.09
0.05
0.05
0.06
0.08
0.23
31.84 0.012 0.17 0.004 0.72
113
5.57
6.07
2.99
7.70
110.8 D098-3B
0.08
27.29 0.005 2.09 0.013 0.13
71.1
5.74
6.84
3.42
12.0
120.3 D098-4
0.12
29.36 0.378 0.20 0.108 0.31
56.1
3.75
7.80
3.57
8.31
2.01
5.92
0.081 0.01 0.022 2.84
842
17.7
37.0
24.9
16.5
126.8 D098-6**
1.73
17.86 0.046 0.09 0.012 2.29
336
6.53
9.12
4.79
15.0
161.9 D098-8
0.24
19.43 0.059 0.39 0.004 0.61
63.9
3.42
3.46
1.86
4.12
191.3 D098-9A
0.31
38.70 0.033 0.55 1.094 0.41
134
7.03
6.11
2.85
4.44
123 D098-5A*
ip t
94 D098-1
cr
Depth (m) Detection Limit 0.2 (ppm) 3.7
Na
0.14
21.18 0.020 1.89 0.012 0.20
68.5
5.70
1.69
2.01
6.87
0.35
24.54 0.028 1.71 0.097 0.48
111
6.49
4.90
2.12
7.36
208 D098-12
0.07
24.06 0.040 2.54 0.032 0.14
35.4
4.13
3.01
3.01
8.88
216.8 D098-13B
0.36
31.60 0.060 0.48 0.018 0.67
138
5.39
2.18
1.74
4.79
221.5 D098-14B
0.29
23.70 0.027 0.94 0.010 0.66
55.6
3.71
1.07
1.24
7.00
209.6 D098-15B
0.33
33.70 0.160 0.11 0.169 0.77
136
5.98
6.01
3.40
5.85
232.8 D098-16
0.29
36.62 0.157 0.60 0.850 0.43
119
5.00
7.71
5.09
5.73
248.3 D098-18
0.36
33.99 0.339 0.59 0.144 0.64
122
4.93
9.25
3.65
5.23
251.7 D098-19A
0.27
38.28 0.371 0.29 0.030 0.56
112
5.08
5.64
2.52
6.24
270.2 D098-20A
0.42
37.72 0.099 0.11 0.004 0.79
158
5.79
5.92
2.77
5.16
270.6 D098-20B
0.06
15.30 0.007 0.95 0.011 0.19
22.5
2.58
2.20
1.21
4.36
281.6 D098-21A
0.41
24.21 0.076 0.21 0.172 0.49
134
6.24
8.63
3.17
3.83
281.9 D098-21B
0.17
31.88 0.029 1.83 0.014 0.24
102
6.25
2.56
1.69
5.52
300.2 D098-22
0.61
34.46 0.100 0.43 0.022 0.64
237
9.50
7.43
3.78
6.71
358.5 D098-24B
0.32
29.90 0.248 0.61 0.039 0.31
99.4
9.03
1.29
1.84
6.49
364.3 D098-25A
0.35
34.40 0.319 0.29 0.008 0.34
130
6.64
4.41
2.12
6.71
Ac ce p
te
d
M
an
us
197 D098-10A 197.3 D098-10B
364.6 D098-25B
0.46
29.53 0.199 1.07 0.018 0.45
160
6.84
6.64
3.50
7.89
411.2 D098-26A
0.03
23.40 0.026 0.01 0.059 0.06
12.0
1.67
0.92
0.41
1.34
435 D098-28***
2.02
19.56 0.355 0.16 0.026 1.94
872
32.3
26.1
20.4
44.5
444 D098-29A
0.07
46.68 0.020 0.16 0.006 0.22
27.8
4.84
2.91
2.29
2.29
444.3 D098-29B
0.02
39.03 0.016 0.35 0.009 0.13
11.8
3.03
0.74
0.50
1.77
448.5 D098-30A
0.08
40.23 0.013 0.26 0.014 0.37
36.8
5.16
8.38
4.15
4.09
448.8 D098-30B
0.06
32.83 0.027 0.19 0.124 0.35
19.9
3.31
1.09
0.73
4.67
* Calcareous mudrock, ** Stilpnomelane‐rich tuffaceous mudrock, *** Stilpnomelane mudrock
Page 73 of 96
Rb
Sr
Y
Zr
Nb
Mo
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
0.02
0.04
0.03
0.02
0.09
0.04
0.02
0.03
0.03
0.03
0.004
0.03
0.04
0.03
0.03
0.03
2.42
87.6
39.3
1.94
8.69
0.63
0.16
39.4
4.13
8.37
0.96
3.69
0.62
0.19
0.53
0.06
3.92
8.99
3.29
3.09
4.53
0.50
0.13
42.8
2.05
4.05
0.48
1.97
0.38
0.15
0.53
0.08
1.54
30.7
32.0
9.76
3.65
0.28
0.40
30.5
1.51
2.59
0.30
1.23
0.27
0.13
0.61
0.11
0.13
31.9
243
6.91
32.0
1.63
0.85
267
7.33
15.4
1.78
6.36
1.14
0.32
1.33
0.18
0.74
68.2
15.8
23.6
57.7
5.73
0.40
145
9.43
20.4
2.49
9.50
1.06
56.3
80.2
1.67
5.20
0.26
0.09
33.8
0.96
1.91
0.20
0.70
3.11
26.6
204
31.8
7.82
0.59
0.14
42.8
5.04
10.3
1.36
6.08
2.87
10.4
11.2
3.52
3.18
0.27
0.05
22.6
1.43
2.69
0.33
2.78
27.5
31.1
11.3
8.07
0.46
0.10
37.3
2.80
5.88
0.79
3.56
5.66
20.0
5.69
1.88
0.19
0.06
39.7
0.98
2.01
0.28
cr
1.35
41.3
21.6
3.66
3.52
0.38
0.05
116
3.25
5.88
2.81
79.7
6.68
4.30
2.72
0.20
0.03
55.8
1.74
3.22
2.40
107
109
16.0
8.67
0.44
0.15
41.0
3.72
7.24
2.12
38.2
128
17.2
7.94
0.47
0.13
94.9
4.32
2.94
73.3
26.4
8.17
8.63
0.43
0.16
91.2
2.15
51.4
27.5
5.97
5.33
0.37
0.12
1.69
108
95.8
3.58
7.88
0.57
0.85
22.8
6.04
2.40
2.21
1.12
39.7
54.1
10.5
3.62
17.7
5.31
3.27
1.99
47.4
23.7
2.83
18.0
3.24
22.9
2.30
24.1
1.97
5.11
0.87
129
4.01
31.2
2.93
19.0
1.64
68.4
1.23
62.6
ip t
Ge
2.01
0.25
2.88
0.57
0.12
0.05
0.17
0.03
1.53
0.75
3.31
0.56
0.31
0.12
0.47
0.07
3.55
0.95
0.38
1.60
0.26
1.37
0.34
0.16
0.63
0.10
0.67
2.54
0.43
0.17
0.58
0.08
0.39
1.62
0.31
0.13
0.45
0.07
0.89
3.75
0.79
0.33
1.31
0.22
9.03
1.22
5.59
1.55
0.87
2.87
0.47
3.45
6.72
0.83
3.50
0.74
0.31
1.16
0.18
62.9
3.10
5.83
0.72
3.04
0.61
0.27
0.93
0.14
0.10
35.8
2.99
6.04
0.73
2.87
0.52
0.16
0.79
0.11
0.14
0.04
49.8
0.86
1.77
0.24
1.02
0.22
0.09
0.36
0.06
8.91
0.47
0.06
107
2.74
6.25
0.85
3.66
0.94
0.40
1.80
0.31
2.96
0.33
0.04
102
1.30
2.52
0.31
1.24
0.26
0.11
0.56
0.09
4.80
12.4
0.86
0.16
106
4.53
9.69
1.17
4.48
0.83
0.28
1.29
0.19
24.7
6.60
4.23
0.67
0.20
64.2
3.18
6.85
0.85
3.61
0.78
0.30
1.66
0.27
16.8
5.67
7.18
0.57
0.15
76.3
3.37
6.86
0.82
3.19
0.60
0.22
1.17
0.19
an
M
d
te
Ac ce p
us
1.39
24.6
8.01
8.10
0.52
0.09
90.6
3.83
7.92
0.98
3.90
0.80
0.29
1.71
0.28
13.2
4.44
1.58
0.11
0.06
7.30
1.13
2.06
0.26
1.09
0.22
0.10
0.50
0.08
57.4
14.8
34.8
2.59
0.73
284
13.3
38.3
3.85
15.2
2.89
0.88
4.64
0.72
56.8
4.41
4.17
0.16
0.10
37.7
3.98
8.02
1.00
4.17
0.66
0.21
1.05
0.14
18.6
3.06
1.51
0.11
0.04
17.9
1.48
2.21
0.25
1.02
0.18
0.08
0.46
0.07
27.1
4.35
3.38
0.25
0.12
38.7
2.21
3.94
0.47
1.93
0.33
0.12
0.80
0.13
49.8
9.10
1.84
0.14
0.09
19.3
1.55
2.48
0.32
1.50
0.33
0.17
1.12
0.19
Page 74 of 96
Dy
Ho
Er
Tm
Yb
Lu
Hf
Pb
Th
U
0.04
0.02
0.04
0.006
0.05
0.04
0.05
0.03
0.01
0.03
0.33
0.07
0.23
0.04
0.36
0.08
0.28
1.05
0.77
0.12
0.84
1.20
1.00
0.52
0.13
0.51
0.11
0.99
0.19
0.12
0.71
0.34
0.06
0.16
2.91
0.88
0.88
0.26
0.96
0.15
1.03
0.18
0.11
0.44
0.32
0.07
0.09
2.37
1.36
1.20
0.28
0.89
0.13
0.87
0.14
1.32
7.23
4.26
1.16
0.65
2.53
0.91
3.92
0.91
2.96
0.45
3.00
0.46
1.47
14.4
3.55
2.57
0.27
0.21
0.06
0.21
0.04
0.32
0.06
0.12
0.40
0.24
0.06
0.20
4.05
1.05
3.37
0.46
2.77
0.43
0.22
1.16
0.75
0.91
0.16
0.49
0.12
0.45
0.08
0.71
0.14
0.06
0.52
0.21
0.13
0.15
2.89
1.09
1.75
0.40
1.22
0.17
1.12
0.18
0.19
2.61
0.54
0.42
0.22
2.38
1.04
0.72
0.20
0.75
0.13
1.01
0.19
0.41
0.08
0.09
2.71
1.05
0.57
0.14
0.49
0.08
0.59
0.11
0.07
0.61
0.22
0.16
0.36
2.46
0.95
0.54
0.14
0.50
0.09
0.65
0.12
0.07
0.35
0.19
0.07
0.19
3.02
1.14
1.63
0.44
1.55
0.25
1.70
0.28
0.23
0.96
0.63
0.26
0.17
2.30
1.33
2.80
0.59
1.57
0.19
1.11
0.17
0.25
1.54
0.72
0.36
0.35
2.61
1.07
1.18
0.28
0.90
0.13
0.83
0.14
0.30
1.67
0.74
0.18
0.32
2.54
1.08
0.91
0.22
0.71
0.11
0.74
0.13
0.17
0.74
0.43
0.12
0.31
2.47
0.98
0.74
0.17
0.58
0.09
0.67
0.12
0.32
1.62
0.83
0.25
0.35
3.38
0.76
0.35
0.08
0.30
0.05
0.44
0.08
0.52
0.07
0.16
0.17
2.83
1.08
1.96
0.43
1.19
0.14
0.81
0.12
0.29
2.16
0.82
0.22
0.33
2.80
0.90
0.61
0.15
0.57
0.11
0.90
0.17
0.06
0.88
0.22
0.17
0.11
2.77
0.81
1.22
0.28
0.90
0.14
1.07
0.18
0.47
2.48
1.58
0.51
0.35
2.60
0.64
1.78
0.41
1.29
0.19
1.30
0.22
0.10
5.52
0.34
0.28
0.21
3.27
0.59
1.35
0.33
1.12
0.18
1.29
0.22
0.33
4.41
0.95
0.43
0.20
2.47
0.62
1.94
0.46
0.60
0.16
4.65
1.01
0.96
0.23
0.53
0.14
0.95
0.25
1.43
0.39
ip t 2.60
0.95
2.26
1.05
2.88
1.11
cr
us
an
M
d
te
Ac ce p
(Pr/Yb)PAAS (Eu/Eu*)PAAS (Y/Ho)PAAS
1.48
0.23
1.47
0.23
0.35
2.59
0.94
0.29
0.21
2.90
0.63
0.52
0.08
0.53
0.09
0.21
0.05
0.15
2.95
1.04
3.16
0.47
3.25
0.52
1.43
7.20
4.83
1.60
0.38
2.54
0.53
0.75
0.11
0.78
0.14
0.10
0.60
0.21
0.03
0.41
2.73
0.70
0.49
0.08
0.55
0.10
0.44
0.06
0.15
2.86
0.79
0.83
0.13
0.94
0.17
0.07
0.78
0.13
0.03
0.16
2.65
0.65
1.33
0.19
1.25
0.20
0.24
0.04
0.08
3.05
0.86
Page 75 of 96
0.52
0.06
0.42
0.13
-0.71
0.05
-0.29
0.08
-0.40
0.02
0.59
0.09
0.24
0.05
0.86
0.04
0.05
0.04
0.31
0.08
0.32
0.06
0.65
0.05
1.21
0.04
0.23
0.11
0.17
0.03
0.04
0.07
0.12
0.10
-0.43
0.10
-0.15
0.11
-0.31
0.07
-0.47
0.05
-0.51
0.08
-0.57
0.08
ip t
0.10
0.12
cr
0.77
-0.23
0.47
0.12
-0.19
0.07
-0.30
0.12
-0.23
0.06
-0.42
0.06
0.00
0.02
-0.22
0.11
-0.23
0.06
0.30
0.10
0.16
0.03
-0.19
0.10
-0.74
0.08
0.48
0.15
0.56
0.03
0.82
0.10
0.65
0.07
us
0.05
δ56Fe ±2SD
an
0.72
Sample DD98-16 DD98-17 DD98-18 DD98-19A DD98-19B DD98-20A DD98-21A DD98-21B DD98-23 DD98-24A DD98-24B DD98-25A DD98-25B DD98-26A DD98-26B DD98-27 DD98-28*** DD98-29A DD98-29B DD98-30A DD98-30B
M
±2SD
d
δ56Fe
te
Sample DD98-1 DD98-2 DD98-3A DD98-3B DD98-4 DD98-5A* DD98-5B DD98-6** DD98-7 DD98-8 DD98-9A DD98-10A DD98-10B DD98-11 DD98-12 DD98-13A DD98-13B DD98-14A DD98-14B DD98-15A DD98-15B
Ac ce p
* Calcareous mudrock, ** Stilpnomelane-rich tuffaceous mudrock, *** Stilpnomelane mudrock
Page 76 of 96
ip t
Figure 1
B Not drawn for scale
M an
us
Pilbara craton
X DD98
Woongarra volcanics
2445±5 Ma (Trendall et al. 2004)
Weeli Wolli Fm.
Hamersley Group
ed ce pt
Pannawonica
Yandicoogina Shale
2454±3 Ma (Pickard et al. 2002)
Brockman Iron Fm.
300 m
Ac
Mt. McRae Shale 2501±8 Ma (Anbar et al., 2007)
Marra Mamba BIF
Mt Tom Price
Basement (granite-greenstone terrain) Fortescue Group (volcanic dominated)
2459±3 Ma (Pickard et al. 2002)
Whaleback Shale
Wittenoom Dolomite
Mt Whaleback
2629±5 Ma (Trendall et al. (2004)
Not drawn for scale
100 km
Joffre BIF
Dales Gorge BIF
Wittenoom
Woongarra
Boolgeeda BIF
cr
A
Fortescue Group
Hamersley Group 2775±10 Ma (Arndt et al. 1991)
Archaean basement
Page 77 of 96
cr
ip t
Figure 2
Mineralogy
Oxide BIF
Chert + magnetite + hematite ± riebeckite ± ankerite ± stilpnomelane
us
Rock type
Chert + magnetite + riebeckite + crocidolite + ankerite + hematite ± stilpnomelane ± Fe-talc ± chlorite
Stilpnomelane-rich tuffaceous mudrock
Chert + stilpnomelane + ankerite + K-feldspar + magnetite + crocidolite + accessory phases: Chlorite + quartz + mica + pyrite + zircon + monazite + apatite + ilmenite
Stilpnomelane mudrock
Stilpnomelane + quartz + mica + K-feldspar
ed
Carbonate + quartz + chlorite + K-feldspar + mica
ce pt
Calcareous mudrock
M an
Silicate-carbonate-oxide BIF
Major minerals - whole core
Magnetite
Riebeckite
Ac
Chert
Ankerite
Hematite
Stilpnomelane
Crocidolite
Increasing abundances
Page 78 of 96
Ac
ce
pt
ed
M
an
us
cr
i
Figure 3
Page 79 of 96
Ac ce p
te
d
M
an
us
cr
ip t
Figure 4
Page 80 of 96
Ac ce p
te
d
M
an
us
cr
ip t
Figure 5
Page 81 of 96
Ac ce p
te
d
M
an
us
cr
ip t
Figure 6
Page 82 of 96
us
cr
ip t
Figure 7
M an
FeO(t)+MgO!
65%$
Greenalite!
50%$
Stilpnomelane!
SiO2!
Ac
20%$
ce pt
ed
Minnesatoite!
Microgranules! ! Matrix of air-fall tuff bed! !
Al2O3+K2O!
Page 83 of 96
M an
B
Al2O3
CaO+MgO
SiO2
Fe2O3(t)
Fe2O3(t)
SiO2
S+lpnomelane-‐rich tuffaceous mudrock
Calcareous mudrock
s+lpnomelane mudrock
Ac
Joffre BIF
ce pt
ed
A
us
cr
ip t
Figure 8
Page 84 of 96
3 90 140
190
190
190
240
240
240
290
290
340
340
390
390
440
440
490
490 Ca (wt.%) 10
15
C
Ti (wt.%)
0.1 0 90
G
90
1
2
K (wt.%) 1 2
3 0 90
190
240
240
290
290
290
340
340
340
390
390
390
440
440
440
490
490
490 3 0.0 90
H
P (wt.%) 0.5
1.0
I
Mn (wt.%)
90
0 0.5 90
0.0
140
140
140
140
140
190
190
190
190
190
240
240
240
240
240
290
290
290
290
340
340
340
340
390
390
390
390
390
440
440
440
440
440
490
490
490
490
490
290 340
E
2
Mg (wt.%) 3 5
140
140
M an 0
D
190
ed
5
Na (wt.%)
ce pt
90
F
0
cr
Al (wt.%) 1 2
140
Ac
Depth (m)
B
140
0
Depth (m)
Fe (wt.%) 10 20 30 40 50 0 90
us
0 90
A
ip t
Figure 9
J
ΣREE (ppm) 50
100
Page 85 of 96
L
Nb (ppm) 2 4
6 0 90
140
140
140
190
190
190
240
240
240
290
290
340
340
390
390
440
440
490
490
Cr (ppm) 20
40 0 90
N
140
Ni (ppm) 10 20
O
30
0.0
140
240
240
290
290
290
340
340
340
390
390
390
440
440
440
490
490
ed
M an
190
SHlpnomelane-‐rich tuffaceous mudrock
Calcareous mudrock
1.0
90
190
490
Mo (ppm) 0.5
sHlpnomelane mudrock
ce pt
Joffre BIF
M
ip t
60 0 90
cr
Zr (ppm) 20 40
Ac
Depth (m)
90
K
us
0
Page 86 of 96
us
cr
ip t
Figure 10
M an
10
Maximum
ed
REY/PAAS
1
ce pt
Minimum
Average Dales Gorge BIF
Ac
0.1
Average Joffre BIF
0.01 La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Y
Ho
Er
Tm
Yb
Lu
Page 87 of 96
cr
ip t
Figure 11
us
Joffre BIF
1000
M an
S+lpnomelane mudrock
0.01
0.001
ce pt
1
0.1
S+lpnomelane-‐rich tuffaceous mudrock
ed
10
Upper con+nental crust
Ac
PM normalised
100
Rb
Ba
Th
U
K
Nb
La
Ce
Sr
Nd
P
Hf
Zr
Na
Sm
Eu
Ti
Y
Yb
Lu
V
Ni
Cr
Page 88 of 96
δ56Fe -‐1.00 50
0.00
1.00
2.00
us
Joffre BIF Calcareous mudstone
cr
ip t
Figure 12
100
M an
Stilpnomelane-rich tuffaceous mudstone Stilpnomelane mudstone
150
ed ce pt
250
300
400
450
500
Ac
350
MOR fluids
Depth (m)
200
Igneous rocks
Page 89 of 96
cr
ip t
Figure 13
10$
us
Wi2enoom&BIF&(Dales&Gorge)&
Marra&Mamba&BIF&(upper&part)&
M an
Marra&Mamba&BIF&(lower&part)& Joffre&BIF&(this&study)&
Wt.%$
1$
ce pt
Al2O3& Al 2O3$
MgO&$ MgO
CaO&$ CaO
Na2O& Na 2O$
K2O& K 2O$
TiO2& TiO 2$
PP2O5& 2O5$
Ac
0.01$
ed
0.1$
Page 90 of 96
us
cr
ip t
Figure 14
M an
Gain'of'iron'
Altered'BIF'(Mt.'Tom'Price)' Marra'Mamba'BIF'(upper'part)'
ed
Fe2O3"
Marra'Mamba'BIF'(lower'part)'
ce pt
Loss'of'silica'
Ac
SiO2"
0"
10"
20"
30" Wt.%"
Wi;enoom'BIF'(Dales'Gorge)' Joffre'BIF'(this'study)'
40"
50"
60"
70"
Page 91 of 96
ip t
Figure 15
80
B
us
ΣLREE
cr
70
50 40
Intermixed volcanogenic detritus
30 20 10 0 0.01
0.1
1
Al (wt.%)
ed
M an
A
10
60
ce pt
(Pr/Yb)SN
1
Dales Gorge BIF Kuruman BIF S?lpnomelane mudrock
Ac
0.1
Joffre BIF (this study)
(DD98-‐28, this study)
Intermixed volcanogenic detritus
S?lpnomelane-‐rich tuffaceous mudstone (DD98-‐6, this study)
Calcareous mudstone (DD98-‐5, this study)
0.01 1
10
100
ΣREE
Page 92 of 96
cr
ip t
Figure 16
0.5
1
1.5
2
2.5
B
1.2
250
Nega9ve La anomaly, no Ce anomaly
Posi9ve Ce anomaly
ed ce pt
1
0.9
Ac
Ce/Ce* shale normalized
1.1
0.8
0.7
150
M an
Depth (m)
us
50
Eu/Eu*shale normalized
0.8
450
No Ce or La anomaly
550
Nega9ve Ce anomaly Posi9ve La anomaly, no Ce anomaly
A 0.7
350
0.9
1
1.1
1.2
Pr/Pr* shale normalized Page 93 of 96
ip t
Figure 17
cr
1000
Flood basalt
M an
us
Tuff (Weeli Wolli Fm.)
100
Zr (ppm)
Tuff (Dales Gorge)
Joffre samples Zr = 1662.4TiO21.6298
Average upper con;nental crust
Subaerial basalt on cont. crust
Average Fortescue shales Joffre BIF
Pillow basalt
Dolerite
Calcareous mudstone
Submarine basalt
s;lpnomelane-‐rich tuffaceous mudstone s;lpnomelane mudrock Joffre tuffaceous mudrock
ed
10
Woongarra rhyolite
0.1
Ac
1 0.01
ce pt
Dales Gorge S-‐bands Zr = 289.43TiO21.2839
1
10
TiO2 (wt.%)
Page 94 of 96
14
cr
16
ip t
Figure 18
n = 42
A
Joffre BIF
us
12 10
M an
8 6 4 2 0
0
0.5
ed
-‐0.5
1
1.5
δ56Fe
ce pt
20 18
B
n = 40
Dales Gorge BIF
16
14
12
Ac
10
8 6
4
2 0 -‐1
-‐0.5
0
0.5
1
1.5
δ56Fe Page 95 of 96
Pyroclas$c material (distal felsic volcanoes)
us
A
cr
Surface seawater: *High nutrient input from upwelling and diluted ash par$cles. *Leads to photoautotrophic Fe(II)-‐oxidia$on and Fe(III) deposi$on
ip t
Figure 19
Sea level
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
M an
Pho$c zone
Fe(III)-‐ hydroxide
Upwelling Hydrothermal fluids: Fe(II) ~ 4-‐20 ppm REY(solute) Eu(II) CH4, H2, H2S
δ56Fe > 0 Si(OH)4
Land
Joffre BIF
B
Sea level
Fe(II) ~ 1-‐3 ppm
Ac
Pho$c zone
ce pt
ed
Con$nental pla9orm
Surface seawater: *Low nutrient input only from diluted ash par$cles.Diminishes bacterial ac$vity and Fe(II) oxida$on. *Silica and siderite background deposi$on
Pyroclas$c material (distal felsic volcanoes)
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
Si(OH)4 FeCO3
Fe(II)-‐silicates?
Land
Joffre BIF Con$nental pla9orm
Page 96 of 96
S0301-9268(15)00362-9 http://dx.doi.org/doi:10.1016/j.precamres.2015.10.024 PRECAM 4398
To appear in:
Precambrian Research
Received date: Revised date: Accepted date:
8-3-2015 28-9-2015 17-10-2015
Please cite this article as: Haugaard, R., Pecoits, E., Lalonde, S., Rouxel, O., Konhauser, K.,The Joffre Banded Iron Formation, Hamersley Group, Western Australia: Assessing the Palaeoenvironment through detailed Petrology and Chemostratigraphy, Precambrian Research (2015), http://dx.doi.org/10.1016/j.precamres.2015.10.024 This is a PDF file of an unedited manuscript that has been accepted for publication. As a service to our customers we are providing this early version of the manuscript. The manuscript will undergo copyediting, typesetting, and review of the resulting proof before it is published in its final form. Please note that during the production process errors may be discovered which could affect the content, and all legal disclaimers that apply to the journal pertain.
ip t
We present petrology and geochemistry of the largest single known banded iron formation in the world Dominant rock types are oxide BIF, silicate-carbonate-oxide BIF with minor stilpnomelane mudrock and stilpnomelane-rich tuffaceous mudrock. Submarine hydrothermal input and fine-grained pyroclastic detritus were the main sources to the seawater Dominant volcanic sources were likely felsic of composition Dispersed stilpnomelane represents diluted ash material that interacted with Fe-rich seawater prior to deposition
cr
2 3 4 5 6 7 8 9 10
Highlights
us
1
Ac ce p
te
d
M
an
11
1
Page 1 of 96
The Joffre Banded Iron Formation, Hamersley Group, Western
12
Australia: Assessing the Palaeoenvironment through detailed Petrology
13
and Chemostratigraphy
ip t
11
14
Rasmus Haugaard1*, Ernesto Pecoits2, Stefan Lalonde3, Olivier Rouxel3, Kurt Konhauser1
cr
15 16 17
Department of Earth and Atmospheric Sciences, University of Alberta, Edmonton, Alberta, T6G 2E3, Canada
2
us
18
1
Equipe Géobiosphère, Institut de Physique du Globe-Sorbonne Paris Cité, Université Paris Diderot, CNRS, 1 place
19
an
3
Universite Europe ene de Bretagne, Institut Universitaire Europe en de la Mer, Plouzane 29280, France
21
*
Corresponding author: [email protected]
M
20
Jussieu, 75238 Paris, France
22
Abstract
d
23
The Joffre Member of the Brockman Iron Formation is by volume the largest
25
single known banded iron formation (BIF) in the world. Here we present detailed
26
petrology and chemostratigraphy through the entire 355 m core section of this ~2.45
27
billion year old unit. Oxide BIF and silicate-carbonate-oxide BIF dominate the lithology,
28
with minor amounts of interbedded stilpnomelane mudrock, stilpnomelane-rich
29
tuffaceous mudrock and calcareous mudrock. Beside chert and magnetite, the prominent
30
mineralogy is riebeckite, ankerite, hematite, stilpnomelane and crocidolite. The BIF is
31
characterised by an average of 50 wt.% SiO2 and 44.5 wt.% Fe2O3 and an overall low
32
abundance of Al2O3 (<1 wt.%), TiO2 (<0.04 wt.%), and trace metals such as Cr (<10
33
ppm), Ni (<5 ppm) and Mo (<0.5 ppm). It has a high ∑REE (rare earth element) content
34
(up to 41 ppm) and a fractionated shale-normalised (SN) seawater REY (rare earth
Ac ce p
te
24
2
Page 2 of 96
element + yttrium) pattern having an enrichment of HREE (heavy rare earth elements)
36
relative to LREE (light rare earth elements) with an average (Pr/Yb)SN of 0.24. The REY
37
patterns also show a positive LaSN anomaly, no CeSN anomaly and a weakly developed
38
positive YSN anomaly. Iron isotopes (56Fe) with positive 56Fe values of +0.04‰ to
39
+1.21‰ suggest that a large part of the hydrothermal iron was partly oxidized in the
40
upper water column and subsequently precipitated as ferric oxyhydroxides. No epiclastic
41
grains have been found; rather submarine hydrothermal fluids and fine-grained
42
volcanogenic detritus controlled BIF chemistry. The former source is reflected through a
43
constant positive EuSN anomaly throughout the core (average EuSN anomaly of 1.6 with a
44
peak of 2.1 between 100-155 m depth), while the latter source is best reflected through
45
the stilpnomelane-rich tuffaceous mudrock consisting of volcanic ash-fall tuff with relict
46
shards set in a stilpnomelane matrix. The mudrock is overlain by well-preserved wavy
47
laminae and laminae sets of stilpnomelane microgranules that likely originated from re-
48
worked volcanic ash formed either on the seafloor or in the water column prior to
49
deposition. An enriched HREE-to-LREE pattern, a high iron content (~30 wt%), and a
50
56Fe value of +0.59‰ collectively imply that the mudrock facies interacted with the Fe-
51
rich seawater prior to deposition. The TiO2-Zr ratio of the BIF and the associated
52
mudrocks suggest a felsic-only-source related to the same style of volcanics as the
53
slightly younger Woongarra rhyolites. Given the observation that the dominant control on
54
the seawater chemistry was associated with felsic volcanics, we speculate that the fine-
55
grained pelagic ash particles may have sourced bio-available nutrients to the surface
56
water. This would have facilitated enhanced biological productivity, including bacterial
57
Fe(II)-oxidation which is now recorded as the positively fractionated
Ac ce p
te
d
M
an
us
cr
ip t
35
3
56
Fe iron oxide
Page 3 of 96
minerals in the Joffre BIF. Alongside submarine hydrothermal input to the basin, the
59
dominant control on the ocean chemistry seems to have been through volcanic and
60
pyroclastic pathways, thereby making the Joffre BIF poorly suited as a chemical proxy
61
for the study of atmospheric oxygen and its weathering impact on local landmasses.
ip t
58
cr
62
Keywords: Palaeoproterozoic; Hamersley Group; Joffre Banded Iron Formation; Seawater chemistry;
64
Provenance; Stilpnomelane
us
63
an
65 66 67
1. Introduction
M
68
Banded iron formations (BIF) are iron-rich (15-40 wt.% Fe2O3) and siliceous (40-60
70
wt.% SiO2) chemical sedimentary deposits that precipitated from seawater throughout
71
much of the Archaean and Palaeoproterozoic (3.8–1.85 Ga). They are also, more often
72
than not, laminated, with banding observed on a wide range of scales, from coarse
73
macrobands (meters in thickness) to the characteristic mesobands (centimeter-thick units)
74
by which they are typically defined (i.e., banded iron formation), to microbands
75
(millimeter to submillimeter). They typically contain low concentrations of Al2O3 (<1
76
wt.%) and incompatible elements (Ti, Zr, Th, Hf and Sc <20 ppm), which indicate
77
minimal detrital input to the depositional basin, although this does not hold for all type of
78
iron formations (see Bekker et al., 2010 for review). For instance, granular iron
79
formations (GIF) typically lack banding and are made of granules of chert and iron
80
oxides or silicates with early diagenetic chert cement filling pore space. Their texture
81
implies that they formed in high-energy environments, with the granules being derived by
Ac ce p
te
d
69
4
Page 4 of 96
82
sedimentary re-working of iron-rich clays, mudstone, arenites, and even stromatolites
83
(e.g., Ojakangas, 1983; Simonson and Goode, 1989). Towards the end of Archaean, the marine depositional setting for BIF formation
85
changed from depositional basins with rapid thermal subsidence and deposition of large
86
volume of interbedded volcanic and volcanogenic greywackes (e.g., Lowe and Tice,
87
2007; Haugaard et al. 2013), to a more stable style of sedimentation in extensive shallow
88
marine basins along stable continental platforms (e.g., Taylor and McLennan, 1981;
89
Eriksson et al., 2001; Condie, 2004; Barley et al., 2005). BIF deposited in the former
90
setting are considered Algoma-type, whereas BIF formed in the latter setting are
91
Superior-type. The latter includes the major BIF of the earliest Paleoproterozoic, such as
92
the Hamersley Group BIF (Gross, 1980).
M
an
us
cr
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84
The mineralogy of BIF from the best-preserved sequences is remarkably uniform,
94
comprising mostly chert, magnetite, hematite, Fe-rich silicate minerals (stilpnomelane,
95
greenalite, minnesotaite, and riebeckite), carbonate minerals (siderite, ankerite, calcite,
96
and dolomite), and minor sulphides (pyrite and pyrrhotite); the presence of both ferric
97
and ferrous minerals gives BIF an average oxidation state of Fe2.4+ (Klein and Beukes,
98
1992). It is generally agreed that none of the minerals in BIF are primary. Instead, the
99
minerals reflect significant post-depositional alteration under diagenetic and metamorphic
100
conditions (including, in some cases, post-depositional fluid flow). The effect of
101
increasing temperature and pressure is manifested by the progressive change in
102
mineralogy through replacement and recrystallisation, increase in crystal size and
103
obliteration of primary textures (Klein, 2005; Bekker et al., 2010).
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The abundance of BIF in Precambrian successions was used in early studies to
105
argue for a largely anoxic atmosphere and ocean system (e.g., Cloud, 1973; Holland,
106
1984) because the accumulation of such large masses of iron found in the form of
107
Superior-type BIF required the transport of Fe(II); Fe(III) is essentially insoluble at
108
circumneutral pH values. Early studies invoked a continental source of iron for BIF
109
because Fe(II) would have been much more mobile in the absence of atmospheric O2
110
(e.g., James, 1954; Lepp and Goldich, 1964) and the continents were more mafic in
111
composition (Condie, 1993). However, detailed studies that followed in the Hamersley
112
basin, Western Australia, suggested that the amount of iron deposited there was on the
113
order of 1 x 1013 ton (Trendall and Blockley, 1970; Trendall and Blockley, 2004). This
114
estimate would have required rivers the size of the modern Amazon to transport orders of
115
magnitude more iron than they do today. This led Holland (1973) to suggest that iron was
116
instead sourced from deep marine waters and supplied to the depositional settings via
117
upwelling. Recently, however, a new model based on Fe- and Nd-isotopes suggests that a
118
large part of the iron in BIF was continental and mobilized by microbial Fe(III) reduction
119
and transported through a benthic iron shuttle to the BIF depositional basin (Li et al.,
120
2015).
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Based on rare earth element (REE) composition of BIF, it is now generally
122
accepted that deep-sea hydrothermal processes are the most likely source of Fe. Shale
123
normalised (SN) europium (Eu) anomalies have been central in the use of REE to trace
124
the hydrothermal input. Eu enrichment in chemical sedimentary rocks precipitated from
125
seawater indicates a strong influence of high-temperature hydrothermal fluids on the
126
seawater dissolved REE load (e.g., Klinkhammer et al., 1983; Derry and Jacobsen, 1988;
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Page 6 of 96
1990). It is generally assumed that Fe and REE will not be fractionated during transport
128
from spreading ridges or other exhalative centres, and, therefore, a strong positive EuSN
129
anomaly indicates that the iron in the BIF precursor sediment was hydrothermally derived
130
(e.g., Slack et al., 2007). In addition to REE concentrations, Sm-Nd isotopes have been
131
used to constrain REE and Fe sources to seawater (e.g., Miller and O'Nions, 1985; Derry
132
and Jacobsen, 1990; Alibert and McCulloch, 1993). The Archaean and Palaeoproterozoic
133
oceans were likely strongly heterogeneous in their Nd(t) values, with +1 to +2 values
134
typical of the deep-waters dominated by hydrothermal sources and lower values, down to
135
-3, typical of shallow-waters dominated by terrestrial sources (Frei et al., 1999, 2007,
136
2008; Alexander et al., 2008).
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Within the large Superior-type BIF, the presence of diagenetic to low-grade
138
metamorphic phyllosilicate minerals, such as greenalite [(Mg,Fe)3Si2O5(OH)4] and
139
stilpnomelane [K0.6(Fe2+,Fe3+,Mg)6Si8Al(O,OH)27nH2O], mostly occur as dense bands
140
interbedded with chemical precipitated minerals such as amorphous silica and ferric
141
hydroxides (La Berge, 1966; Ayres, 1972; Morris, 1983). The precursor of stilpnomelane
142
are believed to have been an iron(III)-rich clay (smectite) derived from volcaniclastic
143
sources likely of basaltic provenance (Trendall and Blockley, 1970; Ewers and Morris,
144
1981; Pickard 2002; Krapež et al. 2003). Where all primary minerals have been
145
overprinted by diagenesis and low-grade metamorphism, preservation of primary textures
146
is rare. As for stilpnomelane, unique preservation of silt-sized microgranules, or
147
spheroids, opens up the possibility to study the precursor sediment. These stilpnomelane
148
microgranules are relatively uncommon but has been observed in Superior-type BIF in
149
the Lake Superior region, Canada (Van Hise and Leith, 1911; Moore, 1918), in the
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Kuruman Iron Formation, South Africa (Beukes, 1973) and in the Brockman Iron
151
Formation, Western Australia (Ayres, 1972, Krapež et al. 2003). Recently, Rasmussen et
152
al., (2013, 2014) presented well-documented lamina sets of stilpnomelane microgranules
153
from the Dales Gorge BIF of the Brockman Iron Formation and proposed that they were
154
generated by flocculation of iron-rich, Al-poor hydrous silicates either on the seabed or
155
within the water column and subsequently shaped and reworked by density currents.
cr
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In this work, we investigate the ~2.45 Ga (Pickard et al., 2002) Joffre Member
157
from the Brockman Iron Formation. This is the single largest known BIF worldwide,
158
containing approximately 4.3x1013 tonnes of iron at the time of deposition (Trendall and
159
Blockley, 2004). This laterally extensive Superior-type BIF, therefore, represents the
160
composition of a large volume of ocean water. We conducted detailed petrologic and
161
geochemical analyses of a core section drilled through the entire ~355 m of stratigraphic
162
depth of the Joffre BIF. Unlike the well-explored Dales Gorge BIF (e.g., Ayres, 1972;
163
Ewers and Morris, 1981; Krapež et al., 2003; Pickard et al., 2004; Pecoits et al., 2009;
164
Rasmussen et al., 2013, 2014, 2015), this key Hamersley BIF has not previously been
165
analysed at high-resolution, nor has a detailed comparison of chemostratigraphy between
166
different lithologies through a complete succession ever been attempted. Furthermore, by
167
directly preceding the Great Oxidation Event (GOE) and the marked increase in Cr
168
dissolution found in the Weeli Wolli Formation (see Konhauser et al., 2011), the Joffre
169
BIF potentially covers the transition from an anoxic to a partially oxygenated Earth. As
170
such, high-resolution petrography, geochemistry and isotopic studies will provide vital
171
information about the evolution of the sediment, the depositional basin and, in particular,
172
the source inputs controlling the seawater composition.
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173 2. The Hamersley Group
175
The Joffre BIF is a member of the ~620 m thick Brockman Iron Formation which makes
176
up part of the 2.63-2.45 Ga Hamersley Group (Trendall et al., 2004). The Hamersley
177
Group comprises almost 2.5 km of consecutive sedimentary and volcanic rocks located
178
within the ca. 80,000 km2 Hamersley Province of the Pilbara craton in North West
179
Australia, approximately 1100 km north of Perth (Fig. 1). In the lower part, it consists of
180
dolomite, shale and BIF, while the upper part consists of dolerite, various lava types and
181
BIF with minor amounts of tuffs and shales (Trendall and Blockley, 1970). Underlying
182
the Hamersley Group, is the 2.78-2.63 Ga (Arndt et al., 1991) Fortescue Group, which
183
consists of flood basalts and rhyolites. These volcanics were laid down on the uplifted
184
and eroded Pilbara block (Trendall, 1968). This volcanic succession may contain
185
remnants of a Large Igneous Province (LIP) as suggested by Ernst et al. (2004).
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The Brockman Iron Formation of the Hamersley Group is divided into four sub-
187
lithostratigraphic units, namely the lowermost Dales Gorge Member (BIF), the
188
Whaleback Shale Member, the Joffre Member (BIF), and the uppermost Yandicoogina
189
Shale Member (Fig. 1B). After deposition, these laterally extensive BIF have all
190
experienced minor folding and basinal uplift along with low-grade regional
191
metamorphism - from burial prehnite-pumpellyite facies to greenschist facies (Smith et
192
al., 1982).
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193
Geochronological constraints (see Fig. 1B) on the Brockman Iron Formation were
194
established by Pickard (2002), and absolute U-Pb zircon ages with interpolated
195
stratigraphic age boundaries of the Hamersley Group are presented in Trendall et al. 9
Page 9 of 96
(2004). The best analytical age estimates for the deposition of the Joffre BIF is 2454±3
197
Ma (Pickard, 2002). This age has been established by SHRIMP U-Pb zircon ages
198
analysed on 19 zircon grains from interbedded tuffaceous mudrock facies at the top of the
199
Joffre BIF (Fig. 1B). The best depositional age estimate for the base of the succession is
200
2459±3 Ma established by only 6 zircon grains (Pickard, 2002, Fig. 1B). Without taking
201
the uncertainties into consideration, there are 166 m between the above two age peaks,
202
which potentially yields a compacted sedimentation rate of 33 m/million year (Pickard,
203
2002).
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General consensus exists regarding the depositional model of the Brockman Iron
205
Formation. According to this model, the succession was deposited on a large, stable, and
206
clastic-starved, continental platform, which was influenced by episodic inputs of fine-
207
grained tuffaceous detritus (e.g., Gross, 1983; Morris, 1983; Krapež et al., 2003). Blake
208
and Barley (1992) proposed a gradually subsiding open-shelf developed within a backarc
209
setting under influence of tuffaceous material sourced from a subduction-related
210
magmatic arc. In addition, Barley et al. (1997) found that deposition of the Hamersley
211
BIF were possibly linked to major submarine magmatic plume activity in the form of a
212
LIPs. Morris (1993) also suggested that the depositional environment for the Hamersley
213
BIF included a steady source of silica and iron with minor lateral variation in the
214
deposition, and a water depth that was deeper than the formation of GIF but shallow
215
enough to form the large carbonate platforms. In the absence of any shoreline facies and
216
the lack of siliciclastics within the BIFs, Morris and Horwitz (1983) further argued that
217
BIF precipitation and deposition took place on an outer shelf that was isolated by a
218
carbonate barrier.
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219 3. Rock core, sampling and analytical methods
221
The drill core, DD98SGP001 (diamond drillcore 1998 Silvergrass Peak #001) was drilled
222
as part of the Rio Tinto project in the Silvergrass Peak area (see Fig. 1A for location).
223
The core samples of DD98SGP001 were obtained at the Rio Tinto core library in Perth,
224
Western Australia. A total of 31 core samples (DD98-1 - DD98-30), each ~30 cm long,
225
were obtained through a total core length of 354.5 m (94 m to 448.5 m depth).
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The core samples were split and one half was stored as future reference material.
227
A total of 40 thin sections were processed and examined using reflected and transmitted
228
light microscopy. In addition, six thin section slabs were polished and carbon coated for
229
electron microprobe analyses. Backscatter electron images, elemental distribution maps,
230
EDS (energy dispersive spectrometry) and WDS (wavelength dispersive spectrometry)
231
were obtained with a JEOL Microprobe 8900 at the University of Alberta. The current
232
was set to ~20 nA and the probe beam diameter was set to 10 microns. Counting time
233
was 20 sec on peak, and 10 sec on background. Dwell time was 10 ms. Standardization to
234
minerals with known element concentration were done on hematite, chromite, kaersutite
235
and diopside for each 50 measurements.
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A total of 30 samples were selected for trace element analyses. Approximately 4
237
cm slabs from each sample were divided into chips and subsequently crushed on an agate
238
mill. The crushed rock powders were dissolved with HF+HNO3 and analyzed using a
239
PerkinElmer Elan6000 Quad-ICPMS (quadrupole inductively coupled plasma mass
240
spectrometer) at the University of Alberta. Accuracy and precision of the analytical
241
protocol was verified with the use of the well-established international whole-rock basalt
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Page 11 of 96
standard (CRPG Nancy). Errors on this standard and on duplicates are both below 10%.
243
Oxide interferences on Ce show that CeO/Ce < 3% and any oxide interferences are
244
therefore considered negligible. For major elements, 17 samples were further analysed by
245
Code 4C (11+) Whole Rock Analysis-XRF at Activation Laboratories Ltd., Ontario,
246
Canada. For iron isotopes, a total of 42 whole rock powder samples were selected.
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Whole rock Fe isotope compositions were analysed at the French oceanographic
248
institution IFREMER, Brest campus, following previously published methods (Rouxel et
249
al., 2005; 2008). Briefly, 50-100 mg of sample powder was digested overnight at 80 °C in
250
4ml 1:1 HF-HNO3 followed by 4 ml aqua regia, with complete evaporation in between.
251
Samples were then taken up in 4ml 6N HCl, from which Fe was purified on Bio-Rad
252
AG1X8 anion resin (2 ml wet resin bed) using 6N HCl for matrix elution followed by
253
0.24N HCl for Fe elution. Fe isotope compositions were determined using a Thermo
254
Scientific Neptune multicollector inductively coupled plasma mass spectrometer
255
operating at medium resolution to resolve isobaric interferences such as 40Ar14N on 54Fe,
256
40
257
correction, introduced to the instrument using an Apex Q desolvating nebuliser
258
(Elemental Scientific, Omaha, NE, USA), and ‘sample-standard bracketing’ was used for
259
data normalisation to a Fe isotope standard solution of IRMM-14 run before and after
260
each
261
0.09±0.09‰ and 0.65±0.14‰, respectively, consistent with previous work (e.g.,
262
Planavsky et al., 2012), and repeated measurements (n=59) of the reference material
263
IRMM-14 (Taylor et al., 1992) constrained average internal precision over the analytical
264
sessions to better than ±0.065 (2 SD).
56
Fe, and
40
Ar16O1H on
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unknown.
Geostandards
57
Fe. Solutions were doped with Ni for mass bias
BHVO-2
12
and
IF-G
yielded
δ56Fe
values
of
Page 12 of 96
Concentrations of REE and Y were shale normalised (SN) to Post-Archaean Australia
266
Shale (PAAS) after Taylor and McLennan (1985). Potential anomalies of La (La/La*SN)
267
and Ce (Ce/Ce*SN) were obtained by the procedure proposed by Bau and Dulski (1996)
268
using the combination of Ce/Ce*PAAS = Ce/(0.5*La+0.5*Pr) and Pr/Pr*SN =
269
Pr/(0.5*Ce+0.5*Nd). If Ce/Ce* < 1 but Pr/Pr*SN ≈ 1 a positive La anomaly is evident. If
270
Ce/Ce*SN < 1 but Pr/Pr*SN > 1.05 a negative Ce anomaly is evident. The Eu anomaly
271
(Eu/Eu*SN) was calculated as Eu/Eu*SN = Eu/[0.67Sm+0.33Tb].
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272 4. Petrography and mineralogy
274
4.1. Rock types
275
Based on the dominant mineralogy, we define five lithological subdivisions within the
276
Joffre BIF (Fig. 2): (1) oxide BIF (Fig. 3), (2) silicate-carbonate-oxide BIF (Fig. 4), (3)
277
stilpnomelane-rich tuffaceous mudrock (Figs. 5 and 6), (4) stilpnomelane mudrock (Fig.
278
5A), and (5) calcareous mudrock. As shown in Fig. 2, a large portion of minerals, such as
279
chert, magnetite, hematite, riebeckite, carbonate and occasionally stilpnomelane exist,
280
although in different proportions within the first three lithologies. In particular, a
281
transitional shift exists between oxide BIF, which is the most dominant rock type, and
282
silicate-carbonate-oxide BIF, which is the second most dominant rock type. Therefore,
283
any clear petrographic split between those rock types is not feasible. Magnetite is the
284
most abundant iron oxide phase and occurs in variable amounts in all of the rock types
285
except in calcareous mudrock.
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4.1.1. Oxide BIF
288
This unit is dominated by micro- and mesobands of chert, magnetite and lesser amounts
289
of hematite (Figs. 3A to 3D). Microcrystalline (~0.05 mm) chert appears both as
290
mesobands (~1-5 cm) and microbands (0.25-1 mm). The chert ranges in color from white
291
to greyish and to a more red variety (jasperlitic) as a result of interstital hematite grains
292
(Fig. 3A). It is often found as microbands with various amount of hematite crocidolite
293
carbonate, alternating with microbands of magnetite hematite. A few pure chert
294
microbands that alternate with magnetite layers are observed sporadically (Figs. 3A to
295
3C).
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Major parts of the magnetite bands are black and opaque while minor parts of the
297
bands are dark grey as a result of interlayered chert microbands (Fig. 3A to 3C). The
298
bands occur both as mesobands (1-3 cm thick) and microbands (down to 0.1 mm thick).
299
The magnetite ranges from fine-grained to coarser-grained with a well-crystallized habit
300
and it is commonly coarser-grained than coexisting chert and hematite (Fig. 3D).
301
Hematite is found both as microcrystalline cement in relation to chert mirco- and
302
mesobands (Fig. 3D) and as <0.1 mm micro-platy crystals (martite), which are locally
303
found associated with magnetite meso- and microbands (Fig. 3D). Martite is formed due
304
to secondary oxidation of magnetite and is thus often found related to magnetite. Micro-
305
platy hematite is also found "floating" in a more brownish to yellowish coherent cement
306
probably of a more goethitic composition. The microcrystalline hematite and chert
307
microbands, consisting of red jasperitic mesobands, are more dominant in the middle
308
section of the Joffre BIF. Some hematite grains are associated with microbands of chert,
309
riebeckite and altered carbonate. This association is likely due to the result of micro-platy
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Page 14 of 96
hematite replacing part of the chert and carbonate as suggested by Clout and Simonson
311
(2005). Some sections of the oxide BIF furthermore reveal thin (<0.01 mm) microbands
312
containing brown microgranules of stilpnomelane,
313
K(Fe2+,Mg,Fe3+)8(Si,Al)12(O,OH)27·n(H2O).
ip t
310
cr
314 4.1.2. Silicate-carbonate-oxide BIF
316
This rock type is distinctive by having relatively more riebeckite and carbonate and less
317
iron oxides than the oxide BIF (Figs. 4A, 4B and 4C).
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Chert mesobands and chert nodules display two different wavy microbands with
319
different grain size (Figs. 4D to 4F). One chert fraction comprises 0.05 mm grains
320
occasionally with a braided network of hematite and minor goethite. These wavy
321
microbands have a thickness of ~0.5 mm. The other chert fraction is finer grained,
322
consisting of <0.02 mm chert grains (Fig. 4F). These microbands are up to 1 mm thick.
323
The finer grained chert microbands are solely restricted to rhombic ankerite and fibrous
324
crocidolite (Fig. 4F). Generally, all chert grains are irregularly shaped and exhibit slight
325
undulatory extinction. Only in few places has the chert been recrystallized to coarser
326
grained (~0.3-0.4 mm) fractions.
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Alongside chert, dense micro- and mesobands of magnetite are found as a major
328
constituent (Fig. 4B). A minor portion of the magnetite displays secondary alteration and
329
oxidation into hematite and other Fe oxides such as goethite (Figs. 4G and 4H).
330 331
After
chert
and
magnetite,
riebeckite,
a
sodium-rich
amphibole
(Na2(Fe2+)3(Fe3+)2(Si8O22)(OH)2), is the most abundant constituent, both within this rock
15
Page 15 of 96
type and within the entire core (Figs. 4B, 4G and 4H). It occurs as dark blue, dense
333
microbands that range from 0.1 mm to 0.5 cm in thickness and as single (0.2 mm long)
334
crystals (Fig. 4I). It is found predominantly within chert-rich microbands and at the
335
interface between chert and iron oxide microbands (Fig. 4G). Locally, riebeckite
336
microbands are associated with magnetite microbands containing randomly dispersed
337
brownish Fe oxide grains (Fig. 4H). The presence of riebeckite-rich fluid escape
338
structures (riebeckite veins in Fig. 4A) indicates remobilization from presumably a
339
hydrous silica-iron-sodium gel during compression of the sediments. Riebeckite is often
340
found associated with acicular and fibrous crocidolite (Figs. 4F and 4I). Crocidolite
341
mostly appears as thin, blue fibrous needles (blue asbestos) in close association with
342
more massive riebeckite crystals or bands. It also occurs within bands of transparent chert
343
microbands where the long fibers of crocidolite tend to grow perpendicular to the
344
bedding planes. Fibrous crocidolite tends to spray out from the rims of the denser
345
riebeckite (Fig. 4I). It coexists with the wavy chert + carbonate microbands indicating,
346
and as proposed by Miyano and Klein (1983), that the crocidolite has grown at the
347
expense of Fe-carbonate. For a more detailed description of both riebeckite and
348
crocidolite as a diagenetic product see Miles (1942), Ryan and Blockley (1965) and
349
Miyano and Klein (1983).
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332
Various compositional pale-brown carbonates occur throughout as individual
351
crystals with a predominantly rhombic habit occasionally displaying internal zoning. The
352
carbonate crystals mostly occur within microbands composed of very fine-grained chert
353
with riebeckite and crocidolite magnetite hematite. Semi-quantitative EDS
354
observations of the carbonate reveal high elemental peaks corresponding to Ca and Fe,
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Page 16 of 96
355
with minor peaks for Mg and Mn. Based on this, we interpret the carbonates as belonging
356
to the dolomite-ankerite series, with a predominantly ankeritic composition. Talc- and chlorite-alteration plays a minor role in the Joffre BIF. However, small
358
colorless and non-pleocroic needles of talc (presumably minnesotaite) are observed in
359
relation to green to pale-green chlorite microbands. These needles are oriented in the
360
direction of stretching, indicating that they have been growing during the main
361
compaction phase of the BIF package. In addition, talc alteration is locally visible in
362
between microlaminae of carbonate, magnetite and chert (Fig. 4J).
363
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357
4.1.3. Stilpnomelane-rich tuffaceous mudrock
365
This rock type is volumetrically minor but it shows important mineralogical features (see
366
Fig. 5A). One of the characteristics of this rock type is the intimate relation between tuff
367
material and stilpnomelane microgranules. The preservation of a 1-2 cm thick bed of
368
pale-green tuff represents direct evidence of volcanogenic ash-fall into the basin (Fig.
369
5B). Texturally, this bed consists of recrystallized shards set in a very fine-grained,
370
greenish-brownish, stilpnomelane matrix (Figs. 5C and 6A). The shards are feldspar-
371
pseudomorphs likely formed from volcanic glass. During compaction and burial
372
metamorphism, the glass devitrified and recrystallized into feldspar. Electron microprobe
373
analysis shows that the feldspar is almost 100% sanidine in composition, with very little
374
iron, calcium and sodium (see Table 1).
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375
In addition to the tuffaceous layer, stilpnomelane is cemented by chert forming
376
plane- or wavy-lamina microbands (Figs. 5B, 5D-F and 6B to 6E). These bands are 0.25-
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Page 17 of 96
1 mm thick and alternate with microbands of almost pure chert with few grains of fibrous
378
crocidolite and platy hematite. The microbands consist of either extremely fine-grained
379
(down to 0.01 mm) stilpnomelane microgranules or spheroids (Fig. 5E) or as medium
380
grained (0.5-1 mm) stilpnomelane flake- and lath-shaped aggregates that are occasionally
381
radiating (Figs. 5F and 5G). The microgranules are pale to dark green to brown
382
(depending on the Fe2+/Fe3+ ratio). Some brown stilpnomelane-rich bands also show
383
evidence of shrinkage texture possibly caused by dehydration of a precursor phase to
384
stilpnomelane (Fig. 5F). EDS images in Figs. 6D and 6E show that elements, such as Al
385
and K, are enriched in the stilpnomelane microlaminae. As for the shards, the
386
stilpnomelane microgranules may often only be preserved where early diagenetic chert
387
formation prevented compaction of the granules (see Rasmussen et al., 2013). However,
388
throughout parts of the core, very fine and diluted laminae of these microgranules can be
389
found interbedded with chert and iron oxide microbands making it difficult to separate
390
true oxide-BIF from weak silicate dominated BIF. The composition of the microgranules
391
and the ash matrix are presented in Table 1 and Fig. 7, wherein a large part of the
392
measurements plot within the stilpnomelane field.
cr
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393
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377
Ankeritic carbonate (Fe-rich dolomite) occurs as individual crystals having
394
predominantly a rhombic habit with clear internal zoning visible in some of the larger
395
(0.2-0.4 mm) crystals (Figs. 5F and 6B). As within the silicate-carbonate-oxide BIF,
396
considerable recrystallization of the carbonate crystals can be seen by the perfect
397
euhedral outline to neighbouring minerals. The large, randomly distributed rhombic
398
carbonate crystals are normally dispersed in very fine-grained chert cement, along with
399
stilpnomelane (Fig. 6B). EDS images shows that these carbonates are both Fe- and Mg-
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Page 18 of 96
rich (Figs. 6F and 6G). The likeliness that the carbonates have grown during burial late
401
stage diagenesis and metamorphism can be seen in Fig. 6H. It shows a trail of
402
stilpnomelane grains together with single chert grains engulfed by a rhombohedral
403
carbonate crystal, indicating the secondary growth of the carbonate crystal after both
404
stilpnomelane and chert. Different styles of carbonate growth can be viewed in Figs. 5D
405
and 6C; both showing excellent preserved lamina of single prismatic carbonate crystals.
406
In contrast to the single rhombic ankerite crystals, EDS observations of these carbonates
407
reveal a more dolomitic (Mg-rich) composition. The distinct symmetry of these dolomite
408
crystals were also found by Morris (1993) in the Marra Mamba BIF (see Fig. 1B), where
409
it was suggested they were pseudomorphs after swallowtail gypsym crystals, and as such,
410
they represent shallow water in this vicinity.
M
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400
Another type of stilpnomelane sedimentation is the existence of ultra-thin
412
microbeds consisting of stilpnomelane with fragments of quarts, chlorite and other sheet
413
minerals (likely muscovite, Fig. 5I). The stilpnomelane is represented as both
414
groundmass and as single individual flakes.
te
Ac ce p
415
d
411
In terms of accessory phases, few randomly dispersed cubic pyrite crystals appear
416
within stilpnomelane laminae, whereas a single-crystal lamina of pyrite is seen in Fig.
417
5G. In between the ash bed and the well-preserved bed of dolomite crystals, electron
418
microprobe EDS analysis reveals a 0.5 mm thin lamina containing zircon, ilmenite,
419
monazite, pyrite and apatite (Figs. 6C and 6I).
420 421
4.1.4. Stilpnomelane mudrock
19
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Stilpnomelane occurs also as massive, almost opaque, mesobands (Fig. 5A). Detailed
423
petrography shows that the apparent structure-less bands are plane-laminated on minute
424
scale (<0.1 mm). The bands, which have a sharp base and top, vary in thickness from 0.5
425
to 2.0 cm, and contain disseminated quartz and K-feldspar fragments that seemingly
426
indicate volcanogenic provenance. The stilpnomelane mudrock bands are volumetrically
427
minor throughout the BIF but are often found in relation to chert-rich and magnetite-poor
428
sections of the core (Fig. 5A).
us
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ip t
422
an
429 4.1.5. Calcareous mudrock
431
At the top part of the core section, a noteworthy ~5 cm thick mesoband of calcareous
432
mudrock occurs. It contains very fine-grained white to pale-grey calcite-dolomite that
433
grades into a fine-grained greenish chloritized material. Angular fragments of carbonate
434
are dispersed within the former, whereas fragments of mostly quartz are dispersed in the
435
latter. The precursor sediment to the calcareous mudrock has erosionally truncated the
436
underlying oxide BIF.
d
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437
M
430
438
4.2. Sedimentary structures
439
Primary sedimentary structures are generally absent throughout the Joffre BIF. This is in
440
full agreement with other core sections from the Brockman Iron Formation (e.g., Trendall
441
et al., 2002). However, here we present possible current-generated sedimentary structures
442
developed prior to compaction and lithification of the BIF package. Around the middle-
443
part of the oxide BIF core (Figs. 3B and 3C), a chert rich mesoband has a wavy
20
Page 20 of 96
appearance composed of two coherent and symmetric, concave-down, structures each ca.
445
1 cm thick. The interesting observation is that the planar laminae immediately above and
446
within the trough is not disturbed by the underlying wavy bedding, making it difficult to
447
interpret it as a post-depositional process. Rather, the concave-down structures were
448
generated prior to the deposition of the above laminae. Another, although weaker
449
developed wavy structure can be seen in the upper left corner of Fig. 3C. Note the planar
450
lamination just above the wavy features. Weak sediment slumping is also seen between
451
bedding planes of chert mesobands and iron oxide mesobands (white arrow Fig. 3A). In
452
few of the magnetite bands, soft sediment deformation is evident by the presence of
453
micro-flame structures (white arrow Fig. 3B). The latter two structures are most likely of
454
pre-diagenetic origin.
M
an
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cr
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444
The chert bedding in the Joffre BIF is normally plane- to wavy laminated, but
456
occasionally the chert forms lenses or nodules (Figs. 4A and 4C). These chert nodules (up
457
to 1 cm thick) still contain internal wavy riebeckite-carbonate microbands. The features
458
most likely developed during burial metamorphism where the lateral termination
459
happened by compaction of more iron-rich bands above and below the original chert
460
layer (white arrow in Fig. 4C). The internal riebeckite-carbonate laminae get compacted
461
from the inner part of the nodule to the outer part. As such, these chert nodules should be
462
interpreted as a result of syn-compaction rather than erosional features.
Ac ce p
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d
455
463 464
5. Bulk rock geochemistry
465
5.1. Major and trace elements
21
Page 21 of 96
Geochemical data for major and trace elements are presented in Tables 2 and 3. SiO2-
467
Fe2O3-Al2O3 and SiO2-Fe2O3-CaO+MgO ternary diagrams are shown in Figs. 8A and 8B.
468
The evolution of selected major and trace elements with depth are presented in Figs. 9A
469
to 9J. Relative to chemically precipitated elements, such as silica and iron, the Al2O3
470
content in the BIF samples (Fig. 8A) is very low (<1 wt.%). In contrast, CaO and MgO
471
are more elevated, reflecting the appearance of well-developed carbonates throughout the
472
core (Fig. 8B). A constant low input of Al (Fig. 9B), Ti (Fig. 9C) and high field strength
473
elements such, as Zr and Nb (Figs. 9K and 9L), is noteworthy. By contrast, elevated
474
abundances of these insoluble elements are found within the calcareous mudrock, the
475
stilpnomelane-rich tuffaceous mudrock and the massive stilpnomelane mudrock. More
476
variable concentrations with depth are observed for Fe, Na, Mn and REEs (Figs. 9A, 9G,
477
9I and 9J). Phosphorous (Fig. 9H) shows low to moderate concentrations throughout the
478
core except for two BIF outliers, which have up to four times as much P as the other BIF
479
samples likely as a result of very fine grained apatite.
cr
us
an
M
d
te
In terms of elemental correlations (diagrams not shown), significant R-values are
Ac ce p
480
ip t
466
481
found for Al vs. Ti and Ti vs. Zr, indicating adsorption of these components to the fine-
482
grained pelagic clay fraction. The absence of correlation between REE and P or Zr
483
suggests minor contribution from sedimentary apatite, monazite and zircon. Interesting,
484
soluble elements, such as K and Ba, are moderately correlated with more conservative
485
elements, such as Al, Ti and Nb. For example Al vs. K yields an R-value of 0.73. Sodium
486
(Na), which shares a similar degree of mobility as K and Ba, is unrelated to Al (R = -
487
0.18). The Al vs. the REE is relatively uncorrelated. However, the variation in LREE
488
with Al, represented by the Al vs. Pr, are better correlated (R = 0.67) than the variation in
22
Page 22 of 96
489
HREE with Al, as represented by Al vs. Yb (R = 0.27). Also phosphorous, is unrelated to
490
Al2O3 (R = 0.24). The maximum, minimum and average shale normalized REE patterns for the 27
492
BIF samples are displayed in Fig. 10. The absolute concentrations of REEs in the Joffre
493
BIF (average of 17 ppm) are highly elevated relative to the undelaying Dales Gorge BIF,
494
although the REE patterns in both BIF are very similar (Pecoits et al., 2009). A general
495
HREE to LREE enrichment (average (Pr/Yb)SN = 0.24), together with a pronounced EuSN
496
anomaly (average (Eu/Eu*)SN = 1.56) is observed (Fig. 10, Table 3). Part of the BIF
497
package shows a positive YSN anomaly, (Y/Ho)SN, reflecting more Y to Ho in the
498
seawater column (Fig. 10, Table 3). However, the YSN anomaly is not significant enough
499
to be clearly evident in the average REY pattern.
M
an
us
cr
ip t
491
In a primitive mantle-normalized spider diagram (Fig. 11), average Joffre BIF is
501
plotted with relevant associated lithologies. There is the same variation for many
502
elements between the BIF and the three Joffre lithologies and, although less pronounced
503
with respect to the upper continental crust. Exceptions include the soluble elements that
504
tend to be concentrated in seawater, and hence in the BIF (e.g., P, Na and Sr). Notably,
505
the P anomaly seen in the BIF is even higher than the average continental crust. The
506
negative anomalies of the Nb and Ti, two elements that are depleted in newly generated
507
continental crust, show an even larger negative anomaly in the BIF likely due to the high
508
insolubility of those elements in seawater (Fig. 11). Similarly, insoluble elements, such as
509
Th, Zr and Hf, which have positive anomalies for the associated lithologies, show a
510
trough in the spider diagram for the BIF.
Ac ce p
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500
511
23
Page 23 of 96
5.2. Fe-isotopes
513
Fe isotopes, as represented by 56Fe values and their standard deviations, are presented in
514
Table 4 and are plotted with the stratigraphic depth in Fig. 12. Inserted, for comparison,
515
are average 56Fe values for common igneous rocks and average mid ocean ridge (MOR)
516
fluids (Sharma et al., 2001; Johnson et al., 2003). Isotopic fractionation that occurs during
517
Fe(II) oxidation is represented by more positive 56Fe values than the typical values for
518
igneous rocks and hydrothermal fluids. At the top part of the core, from 90 m to 220 m, a
519
large amount of the samples have positive 56Fe values averaging +0.33‰. Between ca.
520
220 m to 360 m the 56Fe values have lower values averaging -0.21‰. In the bottom part
521
of the core, from 360 m to 450 m, a change to more positive values is seen by 56Fe
522
values averaging +0.26‰.
M
an
us
cr
ip t
512
te
d
523 6. Discussion
525
6.1. Post-depositional history
526
6.1.1. Diagenetic mineral paragenesis and burial metamorphism
527
During diagenesis and burial of the precursor BIF sediment, the main factors controlling
528
the flow of metasomatic fluids are dependent on various conditions within the
529
sedimentary basin, such as a high water-to-rock ratio, the permeability, the temperature
530
gradient, and the lithostatic pressure differential, amongst others (e.g., Smith, 1980, 1982;
531
Bau, 1993). Most primary hematite has been replaced by dense magnetite bands and
532
secondarily by larger disseminated magnetite grains. The latter can partly obscure the
533
original fine-scale magnetite bedding as seen in Fig. 4H. The generation of magnetite in
Ac ce p
524
24
Page 24 of 96
BIF could have occurred either through dissimilatory Fe(III) reduction in which primary
535
ferric oxyhydroxides were reduced at the expense of organic carbon oxidation during
536
diagenesis (e.g., Konhauser et al., 2005; Li et al., 2011) or at a later stage during
537
metamorphism, for instance through the reaction between hematite and a ferrous iron
538
phase such as siderite (Miyano, 1987; Li et al., 2013),
539
8Fe(OH)3 + CH3COO- (acetate) 8Fe2+ + HCO3- + 15OH- + 5H2O
540
FeCO3 (siderite) + Fe2O3 Fe3O4 + CO2
541
Evidence for any siderite or primary hematite has not been observed in this work. Since
542
siderite is considered to be of depositional or early diagenetic origin (e.g., Ayers, 1972),
543
we infer that during progressive burial siderite reacted with hematite to form the
544
dominant magnetite phase.
M
an
us
cr
ip t
534
The timing of the ankerite-ferroan dolomite crystals seams to be of a later stage
546
than magnetite formation. By engulfing both early chert and very late stage burial-to-low
547
metamorphic stilpnomelane (Fig. 6H), these euhedral rhombic carbonates were formed
548
very late in the post-depositional story.
te
Ac ce p
549
d
545
Riebeckite in BIF is believed to be of pre- to syn-metamorphic origin that either
550
completely or partially replaced chert and chert-magnetite bands (e.g., Beukes, 1973),
551
whilst the crocodolite formation is ascribed to later regional folding generating
552
overpressured zones during vertical extension (e.g., Krapež et al., 2003). Timing of
553
riebeckite growth in the Joffre BIF was before the formation of the chert nodules (Figs.
554
4C and 4D), indicating an earlier formation stage than the main compaction in the basin.
555
However, mobilization of some riebeckite after the main compaction has resulted in thin
25
Page 25 of 96
veins as seen in Fig. 4A. The abundant riebeckite is likely related to the migration of
557
alkali-bearing solutions with high Na+ activity (see also Miyano and Klein, 1983). Na is
558
uncorrelated with immobile elements (e.g., Al2O3, TiO2, Nb, La, Zr, Hf and Th) found in
559
the fine-grained ash and stilpnomelane. This is a result of Na being hosted within
560
riebeckite and crocidolite only. The average Na2O concentration is 1.1 wt.% and shows
561
that the sodium content in the Joffre BIF lies well above other Hamersley BIF, such as
562
the Dales Gorge and the upper and lower parts of the Marra Mamba (Fig. 13). In fact, K
563
and Ba are moderately correlated to the aforementioned immobile elements, as well as
564
with various trace metals, such as V, Cr, Ni and Co, suggesting a minimum degree of K-
565
and Ba-mobilization during burial metamorphism. This indicates that even with the
566
activity of Na-bearing fluids, the degree of alteration of the primary elements has been
567
minor. Interestingly, the high riebeckite content in Joffre BIF is not seen reflected in a
568
lower silica content (see Fig. 14 and section 6.1.2), which would be expected if chert was
569
being replaced by riebeckite. Alternatively, the formation of laminated microbands of
570
dense to fibrous riebeckite restricted to chert laminae suggest that the precursor of the
571
chert was magadiite (NaSi7O13(OH)3), a sodium rich silica gel (e.g., Eugster and Chou
572
1973; Drever, 1974; Miyano and Klein, 1983; Morris, 1993). Thus, it cannot be excluded
573
that Na could have been part of the originally precipitated components of the BIF.
cr
us
an
M
d
te
Ac ce p
574
ip t
556
575
6.1.2. Supergene enrichment
576
A second means by which the BIF sediment can be altered is via supergene weathering.
577
For instance, the downward flow of meteoric oxidative fluids would cause oxidation of
578
any reduced phases in the BIF sediments, resulting in phase changes of magnetite into
26
Page 26 of 96
high-grade hematite. Depending on pH, those fluids may also have led to the dissolution
580
of carbonate minerals (low pH) or silicate minerals (high pH). Low pH solutions would
581
cause the loss of MgO, CaO, and perhaps even Al2O3 (e.g., Weeb et al., 2003). This,
582
however, is not evident in the Joffre BIF when compared to the other associated unaltered
583
BIF (as shown in Fig. 13). High pH solutions would dissolve silica resulting in chert
584
depletion and the concomitant formation of martite and high-grade hematite ore
585
formation as evident in the Mt. Whaleback and Mt. Tom Price BIF from the Hamersley
586
Group (Ewers and Morris, 1981; Taylor et al., 2001; Webb et al., 2003). These high-
587
grade-iron and low-chert-content iron formations are not features seen in the Joffre BIF
588
chemistry. In fact, the Joffre BIF resembles the more unaltered BIF from the Marra
589
Mamba and Wittenoom formations (Fig. 14). Indeed, as seen in Fig. 14, the Joffre BIF
590
has SiO2 and Fe2O3 values within the range of expected values for relatively non-
591
enriched, Hamersley style BIF. With that stated, late-diagenetic magnetite is observed
592
locally in the Joffre core to have been partly oxidized to post-metamorphic fine-grained
593
hematite (Fig. 4H). Furthermore, very fine-grained hematite-goethite grains (possible a
594
variety of martite) are developed sporadically throughout (Fig. 4I).
cr
us
an
M
d
te
Ac ce p
595
ip t
579
596
6.2. Seawater chemistry
597
6.2.1. The REE budget
598
It is generally accepted that REEs measured in Archaean and Paleoproterozoic BIF can
599
potentially mimic the REE composition in the contemporaneous ocean water at the time
600
of precipitation (e.g., Dymek and Klein, 1988; Bau and Dulski, 1996; Bolhar et al.,
27
Page 27 of 96
2004). If true, it implies that (1) all the REEs were dissolved in seawater prior to
602
precipitation, and (2) post-depositional metamorphism did not remove or add any of the
603
elements. The REE patterns all have the characteristic fractionated HREE to LREE
604
enriched patterns (Fig. 10), and this suggests an overall minor contribution from
605
terrigenous sources on the REE budget. The REE vs. the degree of crustal
606
contamination represented by the (Pr/Yb)SN ratios is plotted in Fig. 15A. The diagram
607
reflects a higher concentrations of REEs in the seawater and hence, an increase in the
608
REE contribution to the Joffre BIF basin relative to the underlying Dales Gorge BIF (see
609
Pecoits et al., 2009). A large portion of the samples from the ca. 2.5 Ga Kuruman BIF
610
(South Africa) also contain higher REE than the Dales Gorge BIF, but still less than a
611
large portion of the Joffre BIF samples (Fig. 15A). Furthermore, the Marra Mamba BIF
612
(not shown here) contains lower overall REE concentrations than the Joffre BIF (see
613
Alibert and McCulloch, 1993). Whether this is due to a higher input of submarine
614
hydrothermal fluids and/or volcanogenic input to the basin is unknown. However, the
615
REE systematics show that while all samples exhibit the characteristic fractionated shale-
616
normalised (Pr/Yb)SN < 1) seawater pattern, the portion of the Joffre BIF samples that
617
have higher amount of total REEs (Fig. 15A) appear to be related to a weak increase in
618
Al and LREEs (Fig. 15B). This, in turn, may be controlled by the same volcanic sources,
619
which are represented in the three Joffre samples with intermixed volcanogenic detritus
620
(Fig. 15B). This reveals that small proportions of pelagic ash particles may have had an
621
impact on the overall REE signature and the total REE content of the seawater.
622
Petrographically, evidence of stilpnomelane microgranules associated with both the shard
623
bearing ash bed, as well as intermixed with some of the silica- and iron-oxide microbands
Ac ce p
te
d
M
an
us
cr
ip t
601
28
Page 28 of 96
624
in the BIF (Table 1), may explain the weakly higher LREE and Al content displayed in
625
Fig. 15B. Many Precambrian chemical sediments are also evaluated on the basis of their
627
shale-normalised LaSN- and CeSN-anomalies. The cause for the anomalous behaviour of
628
La reflects enhanced stability of La in solution and, accordingly, may be related to the
629
absence of inner 4f electrons (e.g., De Baar et al., 1985). CeSN anomalies reflect the redox
630
state of the water column from which the particles precipitated. In general, oxygenated
631
marine settings show a strong negative CeSN anomaly, whereas suboxic and anoxic
632
waters lack large negative CeSN anomalies (e.g., German et al., 1991; Byrne and
633
Sholkovitz, 1996). Oxidation of Ce(III) to Ce(IV) greatly reduces Ce solubility, resulting
634
in preferential removal onto Mn-Fe oxyhydroxides, organic matter, and clay particles. In
635
contrast, suboxic and anoxic waters lack significant negative CeSN anomalies due to
636
reductive dissolution of settling Mn-Fe-rich particles. As shown in Fig. 16A, all the
637
samples plot in the two fields of no Ce anomaly, which likely reflects oxygen levels too
638
low to oxidize Ce(III) to Ce(IV), with the concomitant scavenging of Ce(IV) and a
639
resulting negative CeSN anomaly in the BIF. In contrast, a large fraction of the samples
640
show a positive LaSN anomaly, which suggests that La has been stabilized and weakly
641
fractionated relative to the other LREE in the seawater column prior to precipitation.
cr
us
an
M
d
te
Ac ce p
642
ip t
626
643
6.2.2. The hydrothermal input
644
By examining Sm-Nd isotopes in the Joffre BIF, Jacobsen and Pimentel-Klose (1988)
645
obtained an average depleted Nd value of +2.1 (n=4), and, therefore, linked a large
29
Page 29 of 96
portion of the REEs in the Joffre BIF to submarine hydrothermal alteration of the
647
seafloor. Similarly, Alibert and McCulloch (1993) reported a gradual change in the Nd
648
values from the lower Marra Mamba BIF with Nd values of -0.6 to more depleted Nd
649
values of +1 for the Dales Gorge and Joffre BIFs (see Fig. 1B for stratigraphic position).
650
Alibert and McCulloch (1993) additionally made Nd mass balances suggesting that mid-
651
ocean hydrothermal fluids mixed with seawater could explain around 50% of the sourced
652
Nd (and hence the REEs) in Joffre BIF. In this regard, the hydrothermal evolution of the
653
Joffre BIF basin can be discerned by the variation of the EuSN anomaly (Eu/Eu*SN)
654
throughout the BIF sedimentary succession.
an
us
cr
ip t
646
The (Eu/Eu*)SN in modern seawater is identical to the Post-Archaean Average
656
Shale, whereas in modern submarine hydrothermal solutions Eu is enriched with
657
(Eu/Eu*)SN values > 1 (e.g., Danielson et al., 1992; Kato et al., 1998). Fig.16B shows the
658
evolution of (Eu/Eu*)SN with depth in the Joffre BIF. The graph shows that the mixture
659
of hydrothermal fluids with seawater resulted in Eu anomaly well above 1 throughout the
660
entire core depth. However, the input of hydrothermal fluids affected the BIF to varying
661
degrees. An increase in dissolved Eu2+ resulted in the increase of the (Eu/Eu*)SN, that
662
peaked between 100 m to 155 m of the BIF sequence (Fig. 15B). This section reflects a
663
larger hydrothermal input, with a maximum EuSN anomaly of ~2.1.
d
te
Ac ce p
664
M
655
665
6.3. Volcanic activity
666
6.3.1. Ash-fall tuff and microgranules
30
Page 30 of 96
Relict shards within the tuff bed of the Joffre BIF (see Figs. 5C and 6A) indicate a
668
volcanic provenance most likely from explosive magmatic eruptions (e.g., Fisher and
669
Schmincke, 1984). Well-preserved wavy lamina and lamina sets of microgranules
670
deposited on top of the graded tuff bed have been found in the upper part of the Joffre
671
BIF. The granular texture and the well-preserved shards can only exist if early formation
672
of diagenetic silica prevented later compaction of the lamina. Thin section and
673
microprobe analysis of the microgranules and the tuff matrix clearly suggest a
674
stilpnomelane phase. Thus, the stilpnomelane microgranules are linked to the ash bed
675
underneath, and as such, the microgranules most likely represent ash material that has
676
been reworked by unknown processes within the water column or at the seafloor, e.g., by
677
density currents or contourites (e.g., Krapež et al., 2003; Rasmussen et al., 2013). Direct
678
evidence for the latter processes are rare due to burial and metamorphic overprinting but
679
the preserved current generated structure developed in the oxide-BIF (Figs. 3B and 3C) is
680
most likely a result of density currents on the seafloor during sedimentation rather than a
681
post-depositional feature (e.g., Krapež et al., 2003). Those authors linked the origin of
682
locally preserved granules to hydrothermal mud that was transported, deposited and re-
683
sedimented on submarine volcanic flanks by density currents. They suggested that the
684
granular muds may have been the precursor for the BIF, rather than direct precipitates of
685
amorphous silica- and iron oxides from seawater.
Ac ce p
te
d
M
an
us
cr
ip t
667
686
Recently, an interesting model on the formation of the stilpnomelane
687
microgranules within the Dales Gorge BIF has been proposed by Rasmussen et al. (2013)
688
who suggested flocculation of an Fe(III)-rich and Al-poor hydrous silicate either in the
689
water column or on the seabed, which was subsequently reworked by density currents to
31
Page 31 of 96
form lamina sets with a basal granular bed and a more granular diluted-amorphous mud
691
lamina on top. From a geochemical perspective, the high Fe(III)-oxide content and the
692
enriched HREE (Y, Yb, Lu) relative to both primitive mantle and continental crustal
693
values (Fig. 11) strongly support that the precursor to the stilpnomelane mudrock and the
694
stilpnomelane-rich tuffaceous mudrock in the Joffre BIF reacted with the seawater prior
695
to settling, a finding in agreement with the model by Rasmussen et al. (2013). In addition
696
to iron and silica, the stilpnomelane microgranules in the Joffre BIF also contain various
697
amounts of other components, such as Al2O3, TiO2, K2O, Zr, Nb, Th and trace metals.
698
We thus propose that the precursor to the stilpnomelane microgranules in Joffre BIF may
699
have been very fine, reactive ash particles that chemically interacted with seawater,
700
thereby stripping iron from seawater prior to their deposition at the seafloor. In the Joffre
701
BIF, stilpnomelane also exists with other textures than microgranules. For example, the
702
very thin microbed containing stilpnomelane groundmass and flakes alongside quartz
703
fragments and various phyllosilicates (Fig. 5I) may represent settling of volcanic detritus
704
through the water column.
cr
us
an
M
d
te
Ac ce p
705
ip t
690
706
6.3.2. The source of the detritus
707
Seawater composition is inevitably linked to the composition of surrounding continental
708
crust and to changes in the degree of submarine hydrothermal alteration of the seafloor.
709
As a result of a change in the overall heat regime and style of plate tectonics, major
710
geochemical changes took place at the Archaean-Proterozoic boundary (e.g., Taylor and
711
McLennan, 1981; 1985; 2009). For example, the upper continental crust and their
712
sedimentary derivatives became enriched in Large Ion Lithophile Elements (LILE) (e.g.,
32
Page 32 of 96
K, Rb), LREE, Zr, Th and Hf, while they became relatively depleted in various transition
714
metals (e.g., Cr, Ni), Fe and Mg (McLennan, 1985; Condie, 1993). Stilpnomelane hosts
715
the dominant part of the detritus in the Joffre BIF. The occurrence of stilpnomelane is
716
likely the replacement product of greenalite (a ferrous-ferric phyllosilicate of the
717
kaolinite-serpentine group), and suggested to be a key indicator of mafic volcanogenic
718
derived material (LaBerge, 1966; Winkler, 1979; Pickard et al., 2004). Unfortunately, the
719
nature of the volcanic precursor for the Joffre BIF, and other BIF in general, is difficult to
720
deduce since the ash has been deposited in a basin influenced by Fe(II)-rich seawater
721
which subsequently interacted with the highly reactive ash particles (see section 6.3.1).
722
However, the primitive mantle normalized spider diagram in Fig. 11 shows that apart
723
from overall lower abundances, a large part of the insoluble elements in the Joffre BIF
724
follow the same pattern as that for the two associated volcanic deposits (stilpnomelane-
725
rich tuffaceous mudrock and stilpnomelane mudrock) and on a broader scale, the upper
726
continental crust. Therefore, it is important to look in detail at those elements to be able
727
to characterize the provenance of the detritus.
cr
us
an
M
d
te
Ac ce p
728
ip t
713
The low concentrations of insoluble elements (e.g., Nb, Ti, Th, Zr, Hf) can be
729
used as a fingerprint for what lithologies were exposed to weathering on the continents
730
proximal to the BIF depositional basin. In Fig. 17 (TiO2 versus Zr), two insoluble
731
elements that remain adsorbed to the pelagic detritus are plotted for a range of local
732
lithologies that all potentially could have had an influence on the seawater chemistry. A
733
clear dominance from rhyolite type volcanics (here represented by the Woongarra
734
rhyolites) is seen for the Joffre BIF. The regression line is made on data from the Joffre
735
sequence, including the BIF, stilpnomelane-rich tuffaceous mudrock, stilpnomelane
33
Page 33 of 96
mudrock, calcareous mudrock (this study) and 4 tuffaceous mudrock samples from
737
Pickard et al. (2003). All of those samples plot intermediate between the igneous
738
lithologies and the Joffre BIF, indicating a mixture of volcanogenic material and
739
chemical sediment. Similar to the Dales Gorge S-bands (dashed regression line), the
740
massive stilpnomelane mudrock seems to be related to a more intermediate (average
741
upper continental crust) than basaltic TiO2-Zr composition. Interesting, none of the TiO2
742
and Zr in the Joffre BIF is sourced from ultramafic sources or any of the mafic rock
743
suites presented in Fig. 17. This could also explain the low concentrations of trace metals
744
in the Joffre basin (etc., Ni and Cr). The Weeli Wolli tuff, which represents volcanic
745
activity after deposition of Joffre BIF, but before the occurrence of Woongarra rhyolites
746
(see Fig. 1B), also represents a more felsic TiO2-Zr composition and indicates the
747
continuing dominance of felsic volcanic activity on the later stages of Hamersley
748
deposition.
te
d
M
an
us
cr
ip t
736
In this study, no direct evidence of any shelf-derived epiclastic material to the
750
Joffre BIF basin has been found throughout the 355 meter of core section. This is in
751
agreement with other workers who made similar observations for the Hamersley Group
752
BIF as a whole (e.g., Ewers and Morris, 1981; Morris and Horwitz, 1983; Morris, 1993).
753
Instead, it seems likely that most of the continental input to the Joffre basin during BIF
754
precipitation was through volcanic pathways in the form of pyroclastic input and not
755
from an erosive continent. If that is the case, then the lack of terrigenous clastics suggests
756
that the Joffre BIF sequence was either formed in a deep-water setting or in shallower
757
water but within a continental starved and transgressed shelf margin with only minor
758
contribution to the elemental budget from an erosive continental crust.
Ac ce p
749
34
Page 34 of 96
759 6.4. Iron isotopes and Fe(II) oxidation
761
The stratigraphic variations in 56Fe in the Joffre BIF (Table 4 and Fig. 12) suggest that
762
the Fe-cycle in the early Paleaoproterozoc seawater was affected by different
763
fractionation mechanisms (e.g., Steinhofel et al., 2010). For igneous rocks, the
764
isotope ratio (expressed as 56Fe) is generally unfractionated with values around
765
00.15‰, while for mid-ocean hydrothermal fluids the 56Fe is slightly negative, ranging
766
from -0.3 to -0.6‰ (Sharma et al., 2001; Johnson et al., 2003). However, at low
767
temperatures, and influenced by redox processes, it is possible to significantly fractionate
768
the heavy and lighter Fe-isotopes, yielding a range of different 56Fe values as evident in
769
the various BIF facies analyzed here (Fig. 12). The probability histogram for Joffre BIF
770
(Fig. 18A) and Dales Gorge BIF (Fig. 18B) shows the former displaying higher positive
771
56Fe values than the Dales Gorge BIF. This is surprising considering that the measured
772
56Fe (n=40) for the Dales Gorge BIF was on magnetite only, while for the Joffre BIF the
773
measured 56Fe represents bulk analyses (n=42). In contrast to Fe-carbonates, which
774
frequently exhibits negative 56Fe values (Johnson et al., 2003), secondary hematite and
775
magnetite are often the only minerals displaying positive 56Fe values (Johnson et al.,
776
2003; Rouxel et al., 2005). This implies that the main control on the bulk 56Fe signature
777
measured in Joffre BIF is likely from a ferric oxyhydroxide precursor to magnetite (e.g.,
778
ferrihydrite) since magnetite is by far the most dominating mineral. Taking all the
779
ankerite into account, a simple mass balance consideration also suggests a more skewed
780
distribution towards the positive excursion if magnetite was measured only. Bulk BIF
56
Fe/54Fe
Ac ce p
te
d
M
an
us
cr
ip t
760
35
Page 35 of 96
samples with overall positive 56Fe values can also be found in many other Archaean and
782
Palaeoproterozoic BIFs (Planavsky et al., 2012) and show that these were a sink for
783
isotopically heavy Fe. The role of secondary riebeckite on the 56Fe budget of BIF is
784
currently unknown.
ip t
781
This Fe-isotope fractionation can be attributed to either inorganic or organic
786
processes where hydrothermal controlled Fe(II) is oxidized to Fe(III) resulting in
787
precipitation of ferric oxyhydroxides with positive 56Fe values (Johnson et al., 2003).
788
Experimental work by Croal et al. (2004) showed that the above fractionation is possible
789
by the existence of anaerobic photoautotrophic Fe(II)-oxidizing bacteria that use Fe2+ as
790
an electron donor and precipitate ferrihydrite enriched in 56Fe by up to +1.5‰.
791
However, since direct oxidation of Fe2+ by free O2 would generate similar positive 56Fe
792
patterns, this "fingerprint" of anaerobic biogenic fractionation can only be predicted if
793
there is independent evidence for an anoxygenic ocean-atmosphere (Croal et al., 2004).
794
We suggest that during deposition of the Joffre BIF submarine hydrothermal injected
795
fluids, together with pyroclastic detritus, played a greater role on seawater chemistry
796
compared to continental derived epiclastic material. This is not unexpected given that the
797
Brockman Iron Formation as a whole has been linked to major plume breakouts (e.g.,
798
Barley et al., 1997) that developed in association with the emergence of new continental
799
crust during supercontinent assembly (e.g., Condie, 2005). A plume breakout would
800
cause shallower mid-ocean ridges (e.g., Ernst et al., 2004), which in turn, would lead to
801
transgressive events that submerged the shallow shelf in waters directly influenced by
802
submarine mantle degassing and hydrothermal alteration of the oceanic crust (seen
803
through the (Eu/Eu*)SN evolution in Fig. 16B). A combination of high Fe2+, abundant
Ac ce p
te
d
M
an
us
cr
785
36
Page 36 of 96
reduced gases from the mantle (e.g., CH4, H2, H2S), and alteration of new oceanic crust
805
would have acted as O2 sinks and likely promoted marine anoxia in the Joffre basin. This
806
scenario could explain that the fractionated positive Fe-isotopes were a result of
807
anaerobic photosynthetic Fe(II)-oxidizing bacteria consuming lighter Fe-isotopes faster
808
than heavier. For that to happen, a stratified ocean having a large Fe(II)-rich deep water
809
pool and a shallower upper water pool where Fe(II) oxidation and enrichment of Fe
810
isotopes seems plausible. It is important to note that during diagenesis and pore water
811
interaction, it has been shown by Busginy et al. (2014) that loss of light Fe from the pore
812
waters is unlikely to generate positive Fe isotope values in the sediment.
an
us
cr
ip t
804
One interesting aspect is the 56Fe value of +0.59‰ for the BIF dominated by
814
volcanic ash-fall. This value is significantly higher than igneous rock values (Fig. 12),
815
reflecting contemporaneous Fe(II) oxidation during volcanic activity. A portion of the ash
816
formed microgranules consisting of stilpnomelane, and as such, the iron in the granules
817
must have been fractionated in the water column before deposition and reworking. The
818
depositional link between ash-fall, the stilpnomelane microgranules, and evidence of
819
oxidation fractionating the Fe is interesting in that pulses of fine-grained volcanic ash
820
may have promoted higher bacteria production throughout the upper water column
821
(photic zone), speeding up Fe(II) oxidation and precipitation of the ferric oxyhydroxides
822
that subsequently mixed with the ash-particles during deposition. Concomitantly,
823
increased biomass may have settled to the seafloor, allowing for the needed reductants to
824
dissimilatory Fe(III) reduction and magnetite formation to occur during diagenesis and
825
metamorphism (Konhauser et al., 2005; Kolo et al., 2009; Li et al., 2011, 2013). In
826
contrast, the massive stilpnomelane band in the bottom of the core, having affinities to a
Ac ce p
te
d
M
813
37
Page 37 of 96
more basaltic composition, shows a high negative 56Fe value of -0.74, possibly
828
mirroring diagenetic pyrite with high negative 56Fe values. Today, the fertilizing effect
829
of volcanic ash and the importance of volcanism for the marine biogeochemical cycle has
830
been described (for a review see Duggen et al., 2010).
ip t
827
cr
831 6.5. The GOE: sulphur and trace metals
833
The late Archaean to early Palaeoproterozoic (~2.7-2.4 Ga) marks one of the most
834
important periods in Earth history with a number of major interlinked environmental
835
events. These include, amongst others, (1) major episodes of mantle plume activity and
836
continental assembly (Erikkson et al., 2001; Condie, 2004; Barley et al., 2005), (2) a peak
837
in BIF deposition (Isley and Abott, 1999; Bekker et al., 2010); (3) development of an
838
oxic layer in the oceans (Reinhard et al., 2009; Kendall et al., 2010), (4) evolution of
839
aerobically-respiring organisms (Eigenbrode and Freeman, 2006; Godfrey and
840
Falkowski, 2009; Konhauser et al., 2011), and ultimately (5) the oxygenation of the
841
atmosphere, the so-called Great Oxidation Event (GOE) (Holland, 2002; Bekker et al.,
842
2004).
an
M
d
te
Ac ce p
843
us
832
The GOE represents a transition in time from an atmosphere essentially devoid of
844
free oxygen (O2<10-5 times the present atmospheric level, PAL) to oxygen concentrations
845
higher than 10-5 PAL (Pavlov and Kasting, 2002; Kopp et al., 2005; Buick, 2008). The
846
GOE is best defined by a loss of mass-independent sulphur isotope fractionations (S-
847
MIF) in sedimentary rocks (Faquhar, 2001), with data from various locations worldwide
848
showing that S-MIF continued to persist in the rock record until sometime between 2.45
38
Page 38 of 96
and 2.32 Ga (Guo et al., 2009; Canfield and Farquhar, 2009). This rise of free
850
atmospheric oxygen facilitated the onset of oxidative continental weathering reactions
851
and increased the fluxes of sulphate and redox-sensitive trace elements to the oceans
852
(Canfield, 2005; Anbar et al., 2007; Frei et al., 2009; Reinhard et al., 2009; Konhauser et
853
al., 2011). Collectively, these studies show that the GOE was a protracted process that
854
took hundreds of millions of years (Lyons et al., 2014).
cr
ip t
849
The picture of how the GOE emerged has recently gained more clarity through a
856
compilation of Cr concentrations in BIF through time, which showed a significant
857
enrichment beginning at 2.45 Ga in the Weeli Wolli Formation (Konhauser et al., 2011).
858
Given the insolubility of Cr minerals, its mobilization and incorporation into BIF
859
indicates enhanced chemical weathering at that time, most likely associated with the
860
evolution of aerobic continental pyrite oxidation. Interestingly, evidence for a ‘whiff’ of
861
oxygen was previously noted for the underlying 2.5 Ga Mt. McRae shale (Fig. 1B) where
862
molybdenum (Mo) concentrations increase from <5 ppm (near the crustal level) to 40
863
ppm, and then decrease back to <10 ppm (Anbar et al., 2007). These patterns were
864
interpreted as reflecting a transient oxygenation event in the atmosphere, although this
865
view has been complicated by the suggestion that features indicative of oxidative
866
weathering in the pre-GOE rock record may instead stem from localized O2 production in
867
association with biological soil crusts and freshwater microbial mats covering riverbed,
868
lacustrine, and estuarine sediments (Lalonde and Konhauser, 2015).
Ac ce p
te
d
M
an
us
855
869
The rise in oxygen between 2.4 to 2.3 Ga permitted the increased delivery of
870
sulfate to the oceans via enhanced oxidative sulfide weathering on land. Once in
871
seawater, the sulfate had two major sinks, (1) iron sulfide precipitation as a consequence
39
Page 39 of 96
of bacterial sulfate reduction in the water column, or (2) evaporitic precipitation of
873
gypsum. In the first instance, Canfield (1993) proposed that increased levels of sulfide in
874
the oceans effectively titrated out any remaining Fe2+ in seawater, leading to the end of
875
BIF deposition. Evidence in support of higher sulfide production comes from increasing
876
fractionation between sulfur isotopes; values for
877
values (0‰) prior to around 2.45 Ga but then increase to around 25‰ after 2.45 Ga
878
(Canfield and Farquhar, 2009). In the second instance, primary sulfate evaporites are
879
rarely reported before 2.45 Ga (Schröder et al., 2008), confirming insufficient dissolved
880
sulfate availability before that time. In the Joffre BIF, neither petrographic nor
881
geochemical data support the presence of major sulfide or sulfate mineral phases; pyrite
882
has only been documented as a small mineral constituent in the associated stilpnomelane-
883
rich tuffaceous mudrock.
S/32S (34S) are centered on mantle
d
M
an
us
cr
34
ip t
872
The overall low abundances of trace metals in the Joffre BIF relative to upper
885
continental crust values (Fig. 11) reflects either (1) low solubility related to a low degree
886
of oxidative continental weathering of exposed mineral sulfides; (2) that the adjacent land
887
masses to the Joffre basin were of a composition that did not amply supply those metals -
888
for instance, a lack of ultramafic-mafic sources; or (3) continental weathering had only a
889
subordinate control on the seawater chemistry within the Joffre basin.
Ac ce p
890
te
884
In the case of Mo and Cr, their exceedingly low concentrations in the Joffre BIF
891
(Figs. 9M and 9O) suggests that their parent minerals, such as molybdenite (MoS2) and
892
chromite ([FeCr]2O4), were not significantly dissolved. As demonstrated by Anbar et al.
893
(2007), increased concentration of Mo in the 2.5 Ga Mount McRae shale of Western
894
Australia reflected a 'whiff' of oxygen before the GOE; the Mo was sourced from
40
Page 40 of 96
oxidative weathering of Mo-bearing sulfides in crustal rocks. By contrast, low levels of
896
Ni (Fig. 9N) in the Joffre BIF is in full agreement with the findings of Konhauser et al.
897
(2009) who suggested that the cooling of the upper mantle led to decreased eruption of
898
komatiite lavas (with high Ni content), reduced supply of Ni to seawater, and thus less
899
incorporation into marine chemical precipitates, such as BIF after 2.7 Ga. Despite the
900
predicted low levels of Mo, Cr and Ni in seawater at 2.46 Ga, those metals are mainly
901
controlled by the stilpnomelane suggesting that those elements were, to some extent,
902
controlled by the volcanic input best represented by the tuff material in the
903
stilpnomelane-rich tuffaceous mudrock.
an
us
cr
ip t
895
M
904 7. The Palaeoenvironment
906
In the epiclastic-starved basin, seawater composition during Joffre BIF deposition was
907
controlled by two supply systems: (1) convective upwelling of deep, hydrothermally-
908
enriched seawater to the outer continental platform, and (2) volcanic pyroclastic detritus
909
from distal volcanic centers (see Fig. 19 for a simplified model). During convective
910
upwelling, Fe(II) rich deeper waters flooded the platform. The upwelling was likely
911
discontinous because this process depends on seasonal variations in the oceanic currents
912
in the surface water as previously suggested by Morris (1993) for the Marra Mamba BIF
913
(Fig. 1B). Evidence that upwelling varied in time is reflected by the relatively large
914
variations in both the EuSN anomaly and the thickness and distribution of magnetite bands
915
in the core. Alongside Fe(II), the concentration of REY, including Eu(II), were likely to
916
be enriched in the deeper water. If seafloor plumes and submarine volcanism facilitated
917
BIF deposition (as proposed by Barley et al., 1997; Isley and Abbott, 1999), then the
Ac ce p
te
d
905
41
Page 41 of 96
deeper water was also likely enriched in reduced gases from the mantle (e.g., CH4, H2,
919
H2S) that could act as a sink for oxygen and thus promote marine anoxia. The upwelling,
920
along with pyroclastic material, brought important nutrients to the photic zone speeding
921
up the photoautotrophic Fe(II) oxidation. When upwelling ceased (Fig. 19B), the
922
precursor sediment to chert then precipitated from the surface waters due to the high
923
concentrations of dissolved silica in the oceans at that time (e.g., Konhauser et al., 2007).
924
In addition, dissolved bicarbonate or silica reacted with Fe(II) and formed fine-grained
925
siderite (Ayers, 1972; Klein and Beukes, 1989) or greenalite (Rasmussen et al., 2013),
926
respectively.
an
us
cr
ip t
918
Superimposed on this internal marine dynamic system is the pyroclastic source
928
that provided small amounts of fine-grained diluted ash particles. In the photic zone the
929
ash intermixed with the BIF precipitates leading to the different styles of stilpnomelane
930
mineralogy, that is, the two stilpnomelane-rich rock facies and those disseminated within
931
the more typical oxide facies BIF. From the above two sources a large amount, relative to
932
the underlying Dales Gorge BIF, of REY became incorporated into the iron and silica
933
precipitates on the clastic-starved continental platform (Fig. 19). The small rise in total
934
LREE and Al in about half of the samples (Fig. 15B) could thus reflect the influence of
935
very diluted ash particles on the REY budget of the general BIF precipitate (see also
936
Table 1).
Ac ce p
te
d
M
927
937
In the photic zone, enhanced productivity of anaerobic photoautotrophic Fe(II)-
938
oxidising bacteria led to heavier iron isotope fractionation in the precipitated ferric-
939
hydroxides and subsequently in the various style of magnetite bands (Fig. 19A). In
940
clastic-starved environments, without input of bioessential nutrients through erosion of
42
Page 42 of 96
the nearby landmasses, other sources must have controlled biological productivity.
942
Perhaps the delivery of fine-grained ash particles, coupled to upwelling of
943
hydrothermally-enriched seawater, were those nutrient sources? In the case of the former,
944
experiments in seawater have shown that airborne volcanic ash particles have soluble
945
coatings containing important micronutrients (Cu, Zn etc.) and macronutrients (P, K and
946
NH4+) that upon contact with seawater will be released within minutes. This implies that
947
most of the released nutrients are accessible in the photic zone (see Duggen et al., 2010
948
and references therein).
us
cr
ip t
941
This source is hard to quantify since distal tephra input can be very dispersed in
950
the atmosphere depending on the size of the eruption and numerous meteorological
951
factors. However, seafloor drilling carried out to quantify the tephra input to the Pacific
952
Ocean basin, the largest and oldest (174 Ma) ocean basin on Earth, shows that about 25
953
vol.% of the existing oceanic sediments are tephra material, half of which comes from
954
subaerial arc volcanism (Straub and Schmincke, 1998). A major portion of this material
955
is not necessary deposited as distinct ash layers but instead occurs as dispersed ash
956
particles in the marine sediments (Duggen et al., 2010). Therefore, the role of volcanic
957
ash acting as a fertilizing agent for oceanic biota may have been an underestimated factor
958
on the Precambrian Earth.
M
d
te
Ac ce p
959
an
949
960
8. Summary
961
The ca. 2.45 Ga old Joffre BIF can be subdivided into two major rock types (oxide-facies
962
and silicate-carbonate-oxide facies) and three minor rock types (stilpnomelane mudrock,
963
stilpnomelane-rich tuffaceous mudrock and calcareous mudrock). The oxide-facies is 43
Page 43 of 96
dominated by chert, magnetite and hematite, with lesser amount of riebeckite, carbonate,
965
crocidolite and stilpnomelane. The silicate-carbonate-oxide facies is dominated by chert,
966
magnetite, riebeckite and ankerite, with minor hematite, crocidolite and stilpnomelane. A
967
clear lithological boundary between these two facies types is not feasible; rather it is
968
gradational. Although the dominant parts of the sedimentary structures found are of
969
secondary origin (chert nodules, flame structures, etc.), in rare cases the oxide-facies
970
contains primary syn-depositional features presumably of current-generated origin. This
971
illustrates the existence of weak bottom (etc., density) currents within the BIF basin.
us
cr
ip t
964
There is no evidence for epiclastic material, sourced from an erosive continent,
973
within the Joffre BIF. Instead, petrographical studies show that the three minor rock types
974
all have detritus of volcanogenic origin. Furthermore, the chemostratigraphy shows that
975
soluble elements, such as K2O and Ba, co-vary with insoluble elements, such as Al2O3,
976
Ti, Zr and Nb, suggesting only minor modification of the geochemistry. This relationship
977
supports the notion that adjacent volcanoes, delivering pyroclastic sediment to the Joffre
978
BIF basin, were the main source of both insoluble and soluble elements. Indeed, the
979
stilpnomelane-rich tuffaceous mudrock consists of volcanic tuff with well-preserved
980
shards overlain by wavy laminae and laminae sets of stilpnomelane microgranules. These
981
granules most likely originated from re-worked volcanic ash formed either on the
982
seafloor or in the water column. Since the matrix within the tuff bed is of stilpnomelane
983
composition, it is likely that the felsic ash particles reacted with ferric oxyhydroxides in
984
the water column during settling, thereby gaining the iron to form stilpnomelane. This
985
notion is supported by the high 56Fe value of +0.59‰ for this rock type. In fact, detailed
986
geochemistry and petrography shows that small proportions of pelagic ash particles (now
Ac ce p
te
d
M
an
972
44
Page 44 of 96
in the form of stilpnomelane) have had a minor impact on the overall REE signature.
988
While all BIF samples exhibit the characteristic fractionated shale normalised (SN)
989
seawater pattern with (Pr/Yb)SN < 1, a large portion of the samples have high total REEs
990
(relative to other similar BIF) and weakly elevated, although still below 1, (Pr/Yb)SN
991
ratios. These correspond to a slight increase in Al and LREEs, which is directly linked to
992
the volcanic sources adjacent to the Joffre BIF basin. The TiO2-Zr ratio of Joffre BIF and
993
the mudrocks indicates a felsic source related to the same style of volcanics as the
994
slightly younger Woongarra rhyolites. We demonstrate here that the precursor to
995
stilpnomelane does not have to be an indicator of mafic volcanism but instead it could
996
have been felsic volcanic ash that interacted with Fe-rich seawater
M
an
us
cr
ip t
987
In addition to the volcanic contribution, the input of submarine hydrothermal
998
fluids to the seawater played a significant role as a source of solutes to the BIF. This is
999
most clearly evidenced by the EuSN anomaly, (Eu/Eu*)SN, which is above 1 throughout
1000
the entire succession, with a peak value of ~2.1 between 100-155 m of core depth. Within
1001
the water column, a large fraction of the Fe(II) sourced from the mid-ocean ridge
1002
environment underwent heavy isotopic fractionation where Fe(II)-oxidation and
1003
subsequently precipitation of ferric oxyhydroxides resulted in high positive 56Fe values
1004
ranging between +0.04‰ to +1.21‰ (average +0.46‰). This process seems to have been
1005
more pronounced relative to the underlying Dales Gorge BIF.
Ac ce p
te
d
997
1006
Evidence for elevated abundances of sulfur and redox sensitive trace metals (e.g.,
1007
Mo, Cr) have not been found in the BIF, presumably implying a low degree of oxidative
1008
continental weathering. This is not surprising given the lack of epiclastic components and
1009
the fact that the dominant control on the detritus was from felsic volcanics.
45
Page 45 of 96
Correspondingly, it is clear that the Joffre BIF is poorly suited as a chemical proxy for
1011
the study of atmospheric oxygen and its weathering impact on local landmasses.
1012 1013 1014
9. Acknowledgement
1015
We are grateful to Rio Tinto in Australia for providing access to their core samples. A
1016
number of colleagues at the University of Alberta, including Mark Labbe, Martin Von
1017
Dollen, Ilona Ranger, Tom Chako, Andrew Locock and Igor Jakab, are highly
1018
appreciated for their help, as well as Birger Rasmussen for instructive discussions
1019
concerning BIF petrology. The Natural Sciences and Engineering Research Council of
1020
Canada (NSERC) supported this work.
1021
10. References
1022
Alibert, C., & McCulloch, M. T. M., 1993. Rare earth element and neodymium isotopic
1023
compositions of the banded iron-formations and associated shales from Hamersley, western
1024
Australia. Geochimica et Cosmochimica Acta 57, 1–18.
cr
us
an
M
d
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Ac ce p
1025
ip t
1010
1026
Anbar, A. D., Duan, Y., Lyons, T. W., Arnold, G. L., Kendall, B., Creaser, R. A., Kaufman, A.J.,
1027
Gordon, G.W., Scott, C., Garvin, J., Buick, R., 2007. A whiff of oxygen before the great
1028
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Table 1. Major element oxides measured on the microgranules in the BIF (DD98-26A)
1385
and in the stilpnomelane-rich tuffaceous mudrock (DD98-6) and the composition of the
1386
shards and the shard matrix within the ash bed in DD98-6. Note final wt.% with H2O
60
Page 60 of 96
1387
calculated from OH content. The composition of the mictogranules are illustrated in
1388
Figure 7.
ip t
1389 Table 2. The major elements as oxides with core depth in the Joffre BIF. Note the three
1391
different lithologies.
cr
1390
us
1392
Table 3. Selected trace elements and relevant REY anomalies with core depth in the
1394
Joffre BIF. Note the three different lithologies.
an
1393
M
1395
Table 4. Fe isotope values (expressed as 56Fe), and their individual uncertainties, with
1397
depth in the Joffre BIF. Note the three different lithologies.
te
1398
d
1396
Figure 1. (A) The general geology of the Pilbara craton with the location of drill hole
1400
DD98. See text for further explanation. (B) The general stratigraphy of the extensive
1401
banded iron formations and associated rocks of the Hamersley Group. Significant other
1402
BIFs are the Marra Mamba BIF, the Dales Gorge BIF and the BIF of Weeli Wolli
1403
Formation. Note the important U-Pb zircon ages from individual tuff layers. Note the age
1404
of 2454 3 Ma for the upper part the Joffre BIF.
Ac ce p
1399
1405 1406
Figure 2. The main rock types and their mineralogy in the Joffre BIF. The major
1407
minerals are furthermore listed in increasing abundances throughout the core.
1408 61
Page 61 of 96
Figure 3. Photos and photomicrographs illustrating the main petrographic characteristics
1410
of the oxide BIF with representative core samples (A and B) showing pale grey micro-
1411
and mesobands of alternating chert and magnetite, dark grey magnetite mesobands and
1412
reddish and bluish (<<1 mm thin) microbands of chert-hematite-riebeckite crocidolite.
1413
With white arrows, sedimentary slumping can be seen in the upper half of (A) whilst soft
1414
sediment deformation in a magnetite microband is seen by the presence of micro-flame
1415
structures in the top part of (B). Possible current generated sedimentary structures are
1416
presented in (C). The fact that the planar lamina immediately above the wavy band is
1417
undisturbed indicates a primary origin of the two symmetric wavy features. Note also the
1418
two wavy features with internal wavy bedding in the upper left corner of (C). (D) Coarse
1419
and fine grained magnetite microbands with chert and very fine hematite. The latter likely
1420
is a product of secondary magnetite oxidation. Note the single grains of martite upper
1421
left. Ch = chert, Mag = magnetite, Hem = hematite, Rbk = riebeckite.
cr
us
an
M
d
te
1422
ip t
1409
Figure 4. Photos and photomicrographs illustrating the main petrographic characteristics
1424
of the silicate-carbonate-oxide BIF with representative core samples (A and B) and thin
1425
section images (C-J). This rock type is dominated by chert + magnetite + riebeckite +
1426
ankerite + crocidolite stilpnomelane. (A) and (B) Dense mesobands composed of
1427
alternating microbands of chert + riebeckite + crocidolite oxide can be seen in the
1428
middle part of (A) and the upper half of (B). Dense, dark grey magnetite micro- and
1429
macrobands and pale grey chert mesobands with magnetite microbands are evident in
1430
core (B), which furthermore have mesobands of riebeckite with chert and hematite
1431
microbands (lower part). The white arrows in (A) illustrates weak mobilisation of
Ac ce p
1423
62
Page 62 of 96
riebeckite microbands. (C) Thin section PPL image (with inset close up image) of planar
1433
riebeckite microbanding. (D) Chert mesobands and nodules containing wavy-laminated
1434
chert and riebeckite+carbonate+crocidolite microbands. The distinct chert nodules were
1435
likely formed during compaction of the BIF package illustrated with the condensed
1436
internal microbands at the white arrow. These internal laminae are further illustrated in
1437
(E) and (F). Note in (E) the finer grained chert fraction is restricted to the carbonate and
1438
riebeckite-crocidolite laminae only with coarser grained chert in between. (G) PPL thin
1439
section image of blue riebeckite microbands underlain by red hematite and opaque
1440
magnetite microbands. (H) Riebeckite and magnetite microbands in relation to very fine-
1441
grains of other oxides (presumably hematite and goethite). (I) Close up PPL image of the
1442
fibrous bluish crocidolite in a cherty groundmass. (J) Fe-talc alteration occur occasionally
1443
in relation with magnetite. Ch = chert, Mag = magnetite, Hem = hematite, Rbk =
1444
riebeckite, Ank = ankerite, Cr = crocidolite, Ox = oxides (hematite/goethite).
cr
us
an
M
d
te
1445
ip t
1432
Figure 5. Photos and photomicrographs illustrating the main petrographic characteristics
1447
of the stilpnomelane-rich tuffaceous mudrock and stilpnomelane mudrock. (A) Typical
1448
core section influenced by various brown to greenish stilpnomelane mudrock mesobands
1449
alternating with thick white chert bands with internal wavy microbands of stilpnomelane
1450
+ ankeritic dolomite. Note only relative sparse representation of iron oxides and
1451
magnetite. Riebeckite microbands are visible in the middle and lower part of the core
1452
section. (B) Whole thin section PPL image showing stilpnomelane-rich tuffaceous
1453
mudrock having a bottom greenish tuff bed grading up in to plane- and wavy lamination
1454
of stilpnomelane microgranules ending with a magnetite band intermixed with
Ac ce p
1446
63
Page 63 of 96
stilpnomelane. (C) Close up XPL image of the tuff bed in (B) containing well-preserved
1456
shards in a stilpnomelane-rich matrix. (D) A PPL image of the wavy stilpnomelane
1457
laminae set from (B). Each laminae set ends with almost pure laminae of chert. (E) A
1458
close up PPL image of the single crystal dolomite bed in (B) with the stilpnomelane rich
1459
tuff bed underneath and wavy stilpnomelane microgranular laminae above. (F) A close
1460
up PPL image of the stilpnomelane microgranules. (G) PPL image of chert with
1461
stilpnomelane laminae and randomly distributed euhedral ankerite crystals. Note
1462
stilpnomelane occur both as microgranules and as flakes. (H) Stilpnomelane and chert
1463
laminae containing radiating flakes and silt size shrinkage texture of stilpnomelane. A
1464
single laminae of pyrite crystals is seen in the middle part of the image. (I) Ultra-thin
1465
microbed (in PPL) of stilpmomelane rich bearing quartz + chlorite + other sheet silicates
1466
(possible muscovite) illustrating another style of volcanic input to the BIF basin than in
1467
(B). The bed is from the oxide BIF. Ch = chert, Mag = magnetite, Hem = hematite, Rbk =
1468
riebeckite, Ank = ankerite, Cr = crocidolite, Ox = oxides (hematite/goethite), Stp =
1469
stilpnomelane, Chl = chlorite, Qtz = quartz, Py = pyrite, Dol = dolomite.
cr
us
an
M
d
te
Ac ce p
1470
ip t
1455
1471
Figure 6. Backscatter electron images and microprobe elemental maps of various
1472
textures and minerals from the stilpnomelane-rich tuffaceous mudrock. (A) Backscatter
1473
image of Figure 5C, showing dark K-feldspar shards in a stilpnomelane-rich matrix. (B)
1474
Well-crystallised Ankeritic dolomite with internal zoning distributed in chert with
1475
stilpnomelane microgranules. (C) Tuff bed overlain by perfect crystal-shaped dolomite
1476
laminae with homogenous stilpnomelane and dolomite laminae on top. The dominating
1477
wavy-laminae of chert and stilpnomelane microgranules are seen in the upper part. (D
64
Page 64 of 96
and E) Elemental maps showing the distribution of Al and K across wavy micro-laminae
1479
of stilpnomelane and chert. Bright blue (D) and bright green to yellow (E) represent
1480
stilpnomelane-rich laminae having relative higher Al and K compare with darker chert
1481
rich laminae. (F and G) Shows distribution of Fe and Mg across micro-laminae of chert
1482
and stilpnomelane with euhedral ankeritic dolomite crystals. Note the zonation of higher
1483
Fe content (brighter bluish) in the crystals in (F) and the relative high content of Mg
1484
(brighter greenish colors) in (G). Note also brighter zones with respect to Fe and Mg are
1485
evident for the stilpnomelane-rich laminae. (H) A single ankerite crystal that encapsulates
1486
a laminae of single-grain stilpnomelane proving a very late growth of the carbonate. Note
1487
also the engulfment of various amount of chert grains. (I) A close up backscatter image of
1488
the bright laminae in (C). EDS analyses show that this laminae contains various amount
1489
of heavy minerals such as zircon, monazite, ilmenite, pyrite and also apatite (not shown).
1490
Ch = chert, Mag = magnetite, Rbk = riebeckite, Ank = ankerite, Stp = stilpnomelane, Py
1491
= pyrite, Dol = dolomite, Zrn = zircon, Mnz = monazite, Ilm = ilmenite.
cr
us
an
M
d
te
Ac ce p
1492
ip t
1478
1493
Figure 7. Electron microprobe data from the stilpnomelane-rich tuffaceous mudrock
1494
showing a SiO2-FeO+MgO-Al2O3+K2O ternary diagram with 27 measurements of the
1495
microgranules and 10 of the ash matrix (see also Table 4). Almost all of the
1496
measurements fall within the stilpnomelane compositional field. The stilpnomelane,
1497
minnesotaite and greenalite fields are composed from microprobe data from the Marra
1498
Mamba BIF (Klein and Gole, 1981) and crosschecked with mineral data from James
1499
(1954).
1500
65
Page 65 of 96
1501
Figure 8. Major bulk element data from the Joffre BIF plotted in ternary diagrams. (A)
1502
SiO2-Fe2O3(t)-Al2O3 and (B) SiO2-Fe2O3(t)-CaO+MgO. See text for further explanation.
1503 Figure 9. Evolution with depth for selected major and trace elements. See text for details.
ip t
1504 1505
Figure 10. Shale normalised (PAAS) REY pattern from the Joffre BIF (grey area
1507
represents 27 samples) showing pronounced fractionated HREE to LREE pattern along
1508
with a pronounced EuSN anomaly, (Eu/Eu*)SN. Note the high abundances of REY relative
1509
to the underlying Dales Gorge BIF. Average data for Dales Gorge BIF from Pecoits et al.
1510
(2009).
an
us
cr
1506
M
1511
Figure 11. Primitive mantle normalised spider diagram showing the Joffre BIF and the
1513
intermixed massive stilpnomelane mudrock and the stilpnomelane-rich tuffaceous
1514
mudrock. Blue line shows the average upper continental crust from Rudnick and Gao
1515
(2003). For the BIF, note the opposite pattern with very low abundances among major
1516
parts of the insoluble elements (Th, Nb, Zr, Hf, Ti) and the distinctive positive anomalies
1517
of soluble elements (etc., P, Na and Sr) relative to the stilpnomelane mudrock and
1518
stilpnomelane-rich tuffaceous mudrock. Note also the high positive plateau for the
1519
HREEs (Y, Yb, Lu) both for the BIF and for the intermixed tuffaceous detritus. A
1520
significant drop in the trace metals (V, Cr and Ni) is seen in all of the lithologies.
Ac ce p
te
d
1512
1521 1522
Figure 12. Shows the evolution of the Fe isotopes (56Fe) throughout the Joffre BIF.
1523
Fractionation mechanism during Fe oxidation is represented by more positive 56Fe
66
Page 66 of 96
values than the typical values for igneous rocks and mid ocean ridge (MOR) fluids.
1525
Interesting to note is that the stilpnomelane-rich tuffaceous BIF has a high positive 56Fe
1526
value of 0.74. See text for further details. Field of igneous rocks and MOR fluids from
1527
Sharma et al. (2001); Johnson et al. (2003).
ip t
1524
1528
Figure 13. Average major element plot for some of the least altered BIF from the
1530
Hamersley Group. Wittenoom BIF from Webb et al. (2003); Marra Mamba BIF from
1531
Klein and Gole (1981). Notice the high amount of Na2O in Joffre BIF compared with the
1532
other BIFs.
an
us
cr
1529
1533
Figure 14. Bar charts illustrating the same BIFs as in Figure 6 but compared with one of
1535
the classic altered BIF from Mt. Tom Price. A gain of iron and a loss of silica will be a
1536
natural outcome during supergene enrichment. This is not seen for the other BIFs in
1537
general and for the Joffre BIF in particularly. Note the "sum-to-100" problem due to CO2
1538
and H2O in some of the BIFs. Data from Mt. Tom Price BIF from Taylor et al. (2001).
d
te
Ac ce p
1539
M
1534
1540
Figure 15. (A) The REE as a function of (Pr/Yb)SN for both the Joffre BIF, the Dales
1541
Gorge BIF and the same style 2.5 Ga old Kuruman BIF (South Africa). A large portion of
1542
the Joffre BIF samples have high input of REE and are weakly elevated in (Pr/Yb)SN
1543
values. The field of intermixed volcanogenic detritus is defined by the three Joffre
1544
samples (stilpnomelane-rich tuffaceous mudrock, stilpnomelane mudrock and calcareous
1545
mudrock). (B) Shows the similar trend as in (A) but here with Al vs. LREE. The same
1546
samples have an increase in Al and LREE likely as a consequence of ash particles mixed
67
Page 67 of 96
1547
with the BIF. Dales Gorge BIF from Pecoits et al. (2009); Kuruman BIF from Bau and
1548
Dulski (1996).
1549 Figure 16. (A) Graph illustrating shale-normalised depletion/enrichment of La and Ce.
1551
Most of the BIF samples plot within the field of (Ce/Ce*)SN < 1 and (Pr/Pr*)SN ~ 1,
1552
meaning positive La anomaly and no Ce anomaly (diagram modified from Bau and
1553
Dulski, 1996). (B) Shows the hydrothermal evolution (represented by the EuSN anomaly
1554
(Eu/Eu*)SN) during precipitation of Joffre BIF. Note the steady increase of the EuSN
1555
anomaly from the bottom to the top of the core with a peak of ~2.1 around 250 m depth.
an
us
cr
ip t
1550
1556
M
1557
Figure 17. TiO2-Zr, log-log plot showing the Joffre BIF and the relation of these
1559
elements to other voluminous lithologies that may have had an influence on the seawater
1560
chemistry. In conjunction with Joffre tuffaceous mudrock (from Pickard et al., 2003), the
1561
stilpnomelane-rich tuffaceous mudrock, the calcareous mudrock and the stilpnomelane
1562
mudrock, the Joffre BIF produce a (power) regression that is linked to a rhyolite-only-
1563
source represented by the Woongarra rhyolites. In contrast, a more bimodal TiO2-Zr
1564
contribution, best represented by the Dales Gorge S-bands, which clearly have affinities
1565
to average continental crust. The Joffre tuffaceous mudrock and Woongarra rhyolites
1566
from Barley et al. (1997) and Pickard et al. (2003); original linear regression line of Dales
1567
Gorge S-bands (Zr = 244TiO2(wt.%) - 2.1) from Ewers and Morris (1981); field of
1568
submarine komatiite (3.2 Ga Ruth Well Fm.), submarine basalt (2.72 Ga Kylena basalt),
1569
subaerial basalt on cont. crust (2.69 Ga Medina basalt) from Arndt et al. (2001); field of
Ac ce p
te
d
1558
68
Page 68 of 96
pillow basalt (upper Fortescue Group), flood basalt (2.78 Ga Mt. Roe Fm.) from Nelson
1571
et al. (1992); field of dolerite and tuff (Weeli Wolli Fm.) from Barley et al. (1997) and
1572
Pickard et al. (2003); Dales Gorge tuff (S13 and S15) from Pickard et al. (2003); average
1573
upper continental crust from Taylor and McLennan (2009); average Fortescue shale from
1574
Taylor and McLennan (1981).
ip t
1570
cr
1575
Figure 18. 56Fe histogram of Joffre BIF compared with that of Dales Gorge BIF. The
1577
Joffre BIF is 42 analyses of whole rock whilst the Dales Gorge is 40 analyses on
1578
magnetite alone. Despite that fact, the Joffre BIF shows a more skewed distribution
1579
towards more positive 56Fe values, which indicate higher rate of iron oxidation in the
1580
contemporaneous seawater. See text for further interpretation. Values for Dales Gorge
1581
BIF are from Rouxel et al. (2005).
an
M
d te
1582
us
1576
Figure 19. A simplified palaeoenvironmental model for the formation of the Joffre BIF.
1584
(A) Represent the situation during formation of ferric-hydroxide precipitation from the
1585
photic zone. This scenario has nutrient-rich upwelling onto the platform from deeper
1586
waters influenced by hydrothermal activity. Together with diluted fine-grained ash
1587
particles this increase the nutrient level in the surface water speeding up the Fe(III)
1588
oxidation through photoautotrophic Fe(II) oxidation and leaves a positive 56Fe signature
1589
in the sediment. (B) Show the general situation during silica and carbonate (siderite)
1590
formation. The shut off of the upwelling resulted in relative pure deposition of Si(OH)4
1591
and FeCO3 from the upper seawater. See text for further explanations. The concentration
1592
values of Fe(II) in the upper and deeper seawater is taken from Morris (1993).
Ac ce p
1583
69
Page 69 of 96
DD98-26A* (Microgranules) 1
2
3
4
5
SiO2 TiO2 Al2O3
46.24
47.02
45.57
45.65
0.00
0.06
0.01
2.85
2.85
FeO Fe2O3 Cr2O3
26.07
MnO MgO CaO Na2O K2O H2Ocalc
Ave
Ave (n=22)
Ave (n=40)
45.36
45.97
46.54
0.01
0.02
0.02
0.03
2.77
2.86
2.71
2.81
26.35
25.48
25.32
26.20
25.88
25.35
7.24
7.32
7.08
7.03
7.28
7.19
7.04
0.02
0.02
0.00
0.01
0.00
0.01
0.12
0.15
0.13
0.13
0.19
0.14
6.86
6.83
7.13
6.83
7.35
0.03
0.04
0.05
0.03
0.03
0.21
0.15
0.32
0.54
0.20
2.07
1.85
2.48
3.62
64.34 0.01 18.30 0.56 0.02 0.04 0.00 0.05 16.59
8.69
8.80
8.53
8.10
100.40
101.44
cr
4.67
0.01
us
0.09
7.00
5.70
0.04
0.09
0.28
0.37
1.99
2.40
2.83
8.96
8.62
8.05
99.54 100.13 100.29 100.36
100.79
99.91
Ac ce p
te
d
* BIF, ** Stilpnomelane-rich tuffaceous mudrock
ip t
(Shards)
an
Total
DD98-6**
M
Element
DD98-6** (Microgranules)
Page 70 of 96
DD98-6** (Shard matrix)
2
3
4
5
6
7
8
9
10
Ave
47.77
46.41
48.52
47.57
47.40
47.38
49.14
50.94
48.12
46.99
48.02
0.11
N.D.
0.05
0.02
0.06
0.02
0.09
N.D.
0.04
0.03
0.04
5.19
4.43
5.66
4.45
4.47
4.51
5.93
7.22
4.92
4.20
5.10
21.43
22.50
20.66
22.72
23.00
22.93
21.06
17.49
22.50
23.26
21.75
5.95
6.25
5.74
6.31
6.39
6.37
5.85
4.86
6.25
6.46
6.04
0.01
0.00
0.00
0.00
0.05
0.00
0.04
0.03
0.00
0.03
0.02
0.07
0.07
0.05
0.07
0.04
0.06
0.05
0.07
0.04
0.01
0.05
7.03
6.95
6.51
7.18
7.18
7.58
6.58
0.04
0.04
0.04
0.05
0.05
0.03
0.04
1.28
0.89
0.76
0.69
0.98
1.12
0.95
3.85
2.94
4.21
2.77
2.66
2.85
7.34
7.60
7.26
7.83
7.82
7.87
100.08
98.07
99.45
99.66
us
cr
ip t
1
7.28
7.46
6.93
0.04
0.05
0.04
0.04
0.84
0.55
0.64
0.87
an
5.58
4.10
5.43
3.70
3.26
3.58
7.39
6.56
7.79
7.93
7.53
99.05 101.25 100.31 99.98
Ac ce p
te
d
M
100.08 100.72 101.21
Page 71 of 96
Element oxide
SiO2 (wt.%) Al2O3 Fe2O3 MnO
MgO CaO
Na2O K2O
TiO2 P2O5 LOI
Detection Limit
0.01
0.01
0.01
0.001
0.01
0.01
0.01
0.01
0.01
0.01
94.0 DD98-1
46.31
0.62
59.71
0.012
2.21
1.3
0.34
1.17
0.04
0.01
-11.3
100.4
110.8 DD98-3B
60.44
0.26
46.76
0.001
2.18
0.11
3.52
0.26
0.03
0.04
-13.4
100.2
123.0 DD98-5A*
17.33
3.35
6.07
0.114
5.31
33.01
0.17
1.82
0.12
0.05
32.31 99.65
126.8 DD98-6**
65.76
3.86
31.9
0.07
2.49
1.48
0.19
3.31
0.09
0.04
-9.52
99.67
197.0 DD98-10A
72.95
0.29
32.6
0.026
2.91
0.69
2.82
0.37
0.03
0.04
-12.2
100.5
208.0 DD98-12
45.38
0.08
44.6
0.035
2.43
1.05
3.09
0.23
0.03
0.1
3.49
100.5
216.8 DD98-13B
48.05
0.61
43.5
0.07
1.64
1.33
0.57
0.71
0.04
0.04
3.26
99.82
221.5 DD98-14B
59.52
0.57
34.79
0.033
1.7
0.17
1.43
0.89
0.02
0.02
1.26
100.4
209.6 DD98-15B
35.57
0.66
48.37
0.197
2.77
3.95
0.23
0.99
0.04
0.46
6.74
99.98
232.8 DD98-16
37.46
0.55
50.33
0.185
1.58
3.18
0.86
0.57
0.04
1.86
3.51
100.1
270.2 DD98-20A
32.7
0.89
49.29
0.149
3.62
3.77
0.13
1.12
0.05
0.01
8.17
99.89
270.6 DD98-20B
76.4
0.10
20.39
0.011
0.97
0.06
1.35
0.28
0.02
0.03
0.9
100.5
281.6 DD98-21A
59.94
0.73
33.37
0.089
1.32
1.37
0.3
0.61
0.04
0.46
2.19
100.4
411.2 DD98-26A
64.83
0.04
33.3
0.025
0.54
0.82
0.01
0.08
0.01
0.15
0.26
100.1
435.0 DD98-28***
32.59
3.90
32.43
0.504
5.77
5.98
0.26
2.76
0.18
0.07
15.99 100.4
444.0 DD98-29A
23.45
0.14
68.96
0.016
1.39
2.94
0.23
0.26
0.02
0.01
2.79
100.2
448.5 DD98-30A
38.72
0.15
55.37
0.009
1.48
1.16
0.35
0.44
0.02
0.03
2.08
99.81
Depth (m)
Total
Ac c
ep
te
d
M
an
us
cr
ip
t
* Calcareous mudrock, ** Stilpnomelane-rich tuffaceous mudrock, *** Stilpnomelane mudrock
Page 72 of 96
Element
Al (wt.%) Fe
Mn
P
K
0.030 0.5
5
6
Ti (ppm) V
Cr
Ni
Zn
0.09
0.05
0.05
0.06
0.08
0.23
31.84 0.012 0.17 0.004 0.72
113
5.57
6.07
2.99
7.70
110.8 D098-3B
0.08
27.29 0.005 2.09 0.013 0.13
71.1
5.74
6.84
3.42
12.0
120.3 D098-4
0.12
29.36 0.378 0.20 0.108 0.31
56.1
3.75
7.80
3.57
8.31
2.01
5.92
0.081 0.01 0.022 2.84
842
17.7
37.0
24.9
16.5
126.8 D098-6**
1.73
17.86 0.046 0.09 0.012 2.29
336
6.53
9.12
4.79
15.0
161.9 D098-8
0.24
19.43 0.059 0.39 0.004 0.61
63.9
3.42
3.46
1.86
4.12
191.3 D098-9A
0.31
38.70 0.033 0.55 1.094 0.41
134
7.03
6.11
2.85
4.44
123 D098-5A*
ip t
94 D098-1
cr
Depth (m) Detection Limit 0.2 (ppm) 3.7
Na
0.14
21.18 0.020 1.89 0.012 0.20
68.5
5.70
1.69
2.01
6.87
0.35
24.54 0.028 1.71 0.097 0.48
111
6.49
4.90
2.12
7.36
208 D098-12
0.07
24.06 0.040 2.54 0.032 0.14
35.4
4.13
3.01
3.01
8.88
216.8 D098-13B
0.36
31.60 0.060 0.48 0.018 0.67
138
5.39
2.18
1.74
4.79
221.5 D098-14B
0.29
23.70 0.027 0.94 0.010 0.66
55.6
3.71
1.07
1.24
7.00
209.6 D098-15B
0.33
33.70 0.160 0.11 0.169 0.77
136
5.98
6.01
3.40
5.85
232.8 D098-16
0.29
36.62 0.157 0.60 0.850 0.43
119
5.00
7.71
5.09
5.73
248.3 D098-18
0.36
33.99 0.339 0.59 0.144 0.64
122
4.93
9.25
3.65
5.23
251.7 D098-19A
0.27
38.28 0.371 0.29 0.030 0.56
112
5.08
5.64
2.52
6.24
270.2 D098-20A
0.42
37.72 0.099 0.11 0.004 0.79
158
5.79
5.92
2.77
5.16
270.6 D098-20B
0.06
15.30 0.007 0.95 0.011 0.19
22.5
2.58
2.20
1.21
4.36
281.6 D098-21A
0.41
24.21 0.076 0.21 0.172 0.49
134
6.24
8.63
3.17
3.83
281.9 D098-21B
0.17
31.88 0.029 1.83 0.014 0.24
102
6.25
2.56
1.69
5.52
300.2 D098-22
0.61
34.46 0.100 0.43 0.022 0.64
237
9.50
7.43
3.78
6.71
358.5 D098-24B
0.32
29.90 0.248 0.61 0.039 0.31
99.4
9.03
1.29
1.84
6.49
364.3 D098-25A
0.35
34.40 0.319 0.29 0.008 0.34
130
6.64
4.41
2.12
6.71
Ac ce p
te
d
M
an
us
197 D098-10A 197.3 D098-10B
364.6 D098-25B
0.46
29.53 0.199 1.07 0.018 0.45
160
6.84
6.64
3.50
7.89
411.2 D098-26A
0.03
23.40 0.026 0.01 0.059 0.06
12.0
1.67
0.92
0.41
1.34
435 D098-28***
2.02
19.56 0.355 0.16 0.026 1.94
872
32.3
26.1
20.4
44.5
444 D098-29A
0.07
46.68 0.020 0.16 0.006 0.22
27.8
4.84
2.91
2.29
2.29
444.3 D098-29B
0.02
39.03 0.016 0.35 0.009 0.13
11.8
3.03
0.74
0.50
1.77
448.5 D098-30A
0.08
40.23 0.013 0.26 0.014 0.37
36.8
5.16
8.38
4.15
4.09
448.8 D098-30B
0.06
32.83 0.027 0.19 0.124 0.35
19.9
3.31
1.09
0.73
4.67
* Calcareous mudrock, ** Stilpnomelane‐rich tuffaceous mudrock, *** Stilpnomelane mudrock
Page 73 of 96
Rb
Sr
Y
Zr
Nb
Mo
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
0.02
0.04
0.03
0.02
0.09
0.04
0.02
0.03
0.03
0.03
0.004
0.03
0.04
0.03
0.03
0.03
2.42
87.6
39.3
1.94
8.69
0.63
0.16
39.4
4.13
8.37
0.96
3.69
0.62
0.19
0.53
0.06
3.92
8.99
3.29
3.09
4.53
0.50
0.13
42.8
2.05
4.05
0.48
1.97
0.38
0.15
0.53
0.08
1.54
30.7
32.0
9.76
3.65
0.28
0.40
30.5
1.51
2.59
0.30
1.23
0.27
0.13
0.61
0.11
0.13
31.9
243
6.91
32.0
1.63
0.85
267
7.33
15.4
1.78
6.36
1.14
0.32
1.33
0.18
0.74
68.2
15.8
23.6
57.7
5.73
0.40
145
9.43
20.4
2.49
9.50
1.06
56.3
80.2
1.67
5.20
0.26
0.09
33.8
0.96
1.91
0.20
0.70
3.11
26.6
204
31.8
7.82
0.59
0.14
42.8
5.04
10.3
1.36
6.08
2.87
10.4
11.2
3.52
3.18
0.27
0.05
22.6
1.43
2.69
0.33
2.78
27.5
31.1
11.3
8.07
0.46
0.10
37.3
2.80
5.88
0.79
3.56
5.66
20.0
5.69
1.88
0.19
0.06
39.7
0.98
2.01
0.28
cr
1.35
41.3
21.6
3.66
3.52
0.38
0.05
116
3.25
5.88
2.81
79.7
6.68
4.30
2.72
0.20
0.03
55.8
1.74
3.22
2.40
107
109
16.0
8.67
0.44
0.15
41.0
3.72
7.24
2.12
38.2
128
17.2
7.94
0.47
0.13
94.9
4.32
2.94
73.3
26.4
8.17
8.63
0.43
0.16
91.2
2.15
51.4
27.5
5.97
5.33
0.37
0.12
1.69
108
95.8
3.58
7.88
0.57
0.85
22.8
6.04
2.40
2.21
1.12
39.7
54.1
10.5
3.62
17.7
5.31
3.27
1.99
47.4
23.7
2.83
18.0
3.24
22.9
2.30
24.1
1.97
5.11
0.87
129
4.01
31.2
2.93
19.0
1.64
68.4
1.23
62.6
ip t
Ge
2.01
0.25
2.88
0.57
0.12
0.05
0.17
0.03
1.53
0.75
3.31
0.56
0.31
0.12
0.47
0.07
3.55
0.95
0.38
1.60
0.26
1.37
0.34
0.16
0.63
0.10
0.67
2.54
0.43
0.17
0.58
0.08
0.39
1.62
0.31
0.13
0.45
0.07
0.89
3.75
0.79
0.33
1.31
0.22
9.03
1.22
5.59
1.55
0.87
2.87
0.47
3.45
6.72
0.83
3.50
0.74
0.31
1.16
0.18
62.9
3.10
5.83
0.72
3.04
0.61
0.27
0.93
0.14
0.10
35.8
2.99
6.04
0.73
2.87
0.52
0.16
0.79
0.11
0.14
0.04
49.8
0.86
1.77
0.24
1.02
0.22
0.09
0.36
0.06
8.91
0.47
0.06
107
2.74
6.25
0.85
3.66
0.94
0.40
1.80
0.31
2.96
0.33
0.04
102
1.30
2.52
0.31
1.24
0.26
0.11
0.56
0.09
4.80
12.4
0.86
0.16
106
4.53
9.69
1.17
4.48
0.83
0.28
1.29
0.19
24.7
6.60
4.23
0.67
0.20
64.2
3.18
6.85
0.85
3.61
0.78
0.30
1.66
0.27
16.8
5.67
7.18
0.57
0.15
76.3
3.37
6.86
0.82
3.19
0.60
0.22
1.17
0.19
an
M
d
te
Ac ce p
us
1.39
24.6
8.01
8.10
0.52
0.09
90.6
3.83
7.92
0.98
3.90
0.80
0.29
1.71
0.28
13.2
4.44
1.58
0.11
0.06
7.30
1.13
2.06
0.26
1.09
0.22
0.10
0.50
0.08
57.4
14.8
34.8
2.59
0.73
284
13.3
38.3
3.85
15.2
2.89
0.88
4.64
0.72
56.8
4.41
4.17
0.16
0.10
37.7
3.98
8.02
1.00
4.17
0.66
0.21
1.05
0.14
18.6
3.06
1.51
0.11
0.04
17.9
1.48
2.21
0.25
1.02
0.18
0.08
0.46
0.07
27.1
4.35
3.38
0.25
0.12
38.7
2.21
3.94
0.47
1.93
0.33
0.12
0.80
0.13
49.8
9.10
1.84
0.14
0.09
19.3
1.55
2.48
0.32
1.50
0.33
0.17
1.12
0.19
Page 74 of 96
Dy
Ho
Er
Tm
Yb
Lu
Hf
Pb
Th
U
0.04
0.02
0.04
0.006
0.05
0.04
0.05
0.03
0.01
0.03
0.33
0.07
0.23
0.04
0.36
0.08
0.28
1.05
0.77
0.12
0.84
1.20
1.00
0.52
0.13
0.51
0.11
0.99
0.19
0.12
0.71
0.34
0.06
0.16
2.91
0.88
0.88
0.26
0.96
0.15
1.03
0.18
0.11
0.44
0.32
0.07
0.09
2.37
1.36
1.20
0.28
0.89
0.13
0.87
0.14
1.32
7.23
4.26
1.16
0.65
2.53
0.91
3.92
0.91
2.96
0.45
3.00
0.46
1.47
14.4
3.55
2.57
0.27
0.21
0.06
0.21
0.04
0.32
0.06
0.12
0.40
0.24
0.06
0.20
4.05
1.05
3.37
0.46
2.77
0.43
0.22
1.16
0.75
0.91
0.16
0.49
0.12
0.45
0.08
0.71
0.14
0.06
0.52
0.21
0.13
0.15
2.89
1.09
1.75
0.40
1.22
0.17
1.12
0.18
0.19
2.61
0.54
0.42
0.22
2.38
1.04
0.72
0.20
0.75
0.13
1.01
0.19
0.41
0.08
0.09
2.71
1.05
0.57
0.14
0.49
0.08
0.59
0.11
0.07
0.61
0.22
0.16
0.36
2.46
0.95
0.54
0.14
0.50
0.09
0.65
0.12
0.07
0.35
0.19
0.07
0.19
3.02
1.14
1.63
0.44
1.55
0.25
1.70
0.28
0.23
0.96
0.63
0.26
0.17
2.30
1.33
2.80
0.59
1.57
0.19
1.11
0.17
0.25
1.54
0.72
0.36
0.35
2.61
1.07
1.18
0.28
0.90
0.13
0.83
0.14
0.30
1.67
0.74
0.18
0.32
2.54
1.08
0.91
0.22
0.71
0.11
0.74
0.13
0.17
0.74
0.43
0.12
0.31
2.47
0.98
0.74
0.17
0.58
0.09
0.67
0.12
0.32
1.62
0.83
0.25
0.35
3.38
0.76
0.35
0.08
0.30
0.05
0.44
0.08
0.52
0.07
0.16
0.17
2.83
1.08
1.96
0.43
1.19
0.14
0.81
0.12
0.29
2.16
0.82
0.22
0.33
2.80
0.90
0.61
0.15
0.57
0.11
0.90
0.17
0.06
0.88
0.22
0.17
0.11
2.77
0.81
1.22
0.28
0.90
0.14
1.07
0.18
0.47
2.48
1.58
0.51
0.35
2.60
0.64
1.78
0.41
1.29
0.19
1.30
0.22
0.10
5.52
0.34
0.28
0.21
3.27
0.59
1.35
0.33
1.12
0.18
1.29
0.22
0.33
4.41
0.95
0.43
0.20
2.47
0.62
1.94
0.46
0.60
0.16
4.65
1.01
0.96
0.23
0.53
0.14
0.95
0.25
1.43
0.39
ip t 2.60
0.95
2.26
1.05
2.88
1.11
cr
us
an
M
d
te
Ac ce p
(Pr/Yb)PAAS (Eu/Eu*)PAAS (Y/Ho)PAAS
1.48
0.23
1.47
0.23
0.35
2.59
0.94
0.29
0.21
2.90
0.63
0.52
0.08
0.53
0.09
0.21
0.05
0.15
2.95
1.04
3.16
0.47
3.25
0.52
1.43
7.20
4.83
1.60
0.38
2.54
0.53
0.75
0.11
0.78
0.14
0.10
0.60
0.21
0.03
0.41
2.73
0.70
0.49
0.08
0.55
0.10
0.44
0.06
0.15
2.86
0.79
0.83
0.13
0.94
0.17
0.07
0.78
0.13
0.03
0.16
2.65
0.65
1.33
0.19
1.25
0.20
0.24
0.04
0.08
3.05
0.86
Page 75 of 96
0.52
0.06
0.42
0.13
-0.71
0.05
-0.29
0.08
-0.40
0.02
0.59
0.09
0.24
0.05
0.86
0.04
0.05
0.04
0.31
0.08
0.32
0.06
0.65
0.05
1.21
0.04
0.23
0.11
0.17
0.03
0.04
0.07
0.12
0.10
-0.43
0.10
-0.15
0.11
-0.31
0.07
-0.47
0.05
-0.51
0.08
-0.57
0.08
ip t
0.10
0.12
cr
0.77
-0.23
0.47
0.12
-0.19
0.07
-0.30
0.12
-0.23
0.06
-0.42
0.06
0.00
0.02
-0.22
0.11
-0.23
0.06
0.30
0.10
0.16
0.03
-0.19
0.10
-0.74
0.08
0.48
0.15
0.56
0.03
0.82
0.10
0.65
0.07
us
0.05
δ56Fe ±2SD
an
0.72
Sample DD98-16 DD98-17 DD98-18 DD98-19A DD98-19B DD98-20A DD98-21A DD98-21B DD98-23 DD98-24A DD98-24B DD98-25A DD98-25B DD98-26A DD98-26B DD98-27 DD98-28*** DD98-29A DD98-29B DD98-30A DD98-30B
M
±2SD
d
δ56Fe
te
Sample DD98-1 DD98-2 DD98-3A DD98-3B DD98-4 DD98-5A* DD98-5B DD98-6** DD98-7 DD98-8 DD98-9A DD98-10A DD98-10B DD98-11 DD98-12 DD98-13A DD98-13B DD98-14A DD98-14B DD98-15A DD98-15B
Ac ce p
* Calcareous mudrock, ** Stilpnomelane-rich tuffaceous mudrock, *** Stilpnomelane mudrock
Page 76 of 96
ip t
Figure 1
B Not drawn for scale
M an
us
Pilbara craton
X DD98
Woongarra volcanics
2445±5 Ma (Trendall et al. 2004)
Weeli Wolli Fm.
Hamersley Group
ed ce pt
Pannawonica
Yandicoogina Shale
2454±3 Ma (Pickard et al. 2002)
Brockman Iron Fm.
300 m
Ac
Mt. McRae Shale 2501±8 Ma (Anbar et al., 2007)
Marra Mamba BIF
Mt Tom Price
Basement (granite-greenstone terrain) Fortescue Group (volcanic dominated)
2459±3 Ma (Pickard et al. 2002)
Whaleback Shale
Wittenoom Dolomite
Mt Whaleback
2629±5 Ma (Trendall et al. (2004)
Not drawn for scale
100 km
Joffre BIF
Dales Gorge BIF
Wittenoom
Woongarra
Boolgeeda BIF
cr
A
Fortescue Group
Hamersley Group 2775±10 Ma (Arndt et al. 1991)
Archaean basement
Page 77 of 96
cr
ip t
Figure 2
Mineralogy
Oxide BIF
Chert + magnetite + hematite ± riebeckite ± ankerite ± stilpnomelane
us
Rock type
Chert + magnetite + riebeckite + crocidolite + ankerite + hematite ± stilpnomelane ± Fe-talc ± chlorite
Stilpnomelane-rich tuffaceous mudrock
Chert + stilpnomelane + ankerite + K-feldspar + magnetite + crocidolite + accessory phases: Chlorite + quartz + mica + pyrite + zircon + monazite + apatite + ilmenite
Stilpnomelane mudrock
Stilpnomelane + quartz + mica + K-feldspar
ed
Carbonate + quartz + chlorite + K-feldspar + mica
ce pt
Calcareous mudrock
M an
Silicate-carbonate-oxide BIF
Major minerals - whole core
Magnetite
Riebeckite
Ac
Chert
Ankerite
Hematite
Stilpnomelane
Crocidolite
Increasing abundances
Page 78 of 96
Ac
ce
pt
ed
M
an
us
cr
i
Figure 3
Page 79 of 96
Ac ce p
te
d
M
an
us
cr
ip t
Figure 4
Page 80 of 96
Ac ce p
te
d
M
an
us
cr
ip t
Figure 5
Page 81 of 96
Ac ce p
te
d
M
an
us
cr
ip t
Figure 6
Page 82 of 96
us
cr
ip t
Figure 7
M an
FeO(t)+MgO!
65%$
Greenalite!
50%$
Stilpnomelane!
SiO2!
Ac
20%$
ce pt
ed
Minnesatoite!
Microgranules! ! Matrix of air-fall tuff bed! !
Al2O3+K2O!
Page 83 of 96
M an
B
Al2O3
CaO+MgO
SiO2
Fe2O3(t)
Fe2O3(t)
SiO2
S+lpnomelane-‐rich tuffaceous mudrock
Calcareous mudrock
s+lpnomelane mudrock
Ac
Joffre BIF
ce pt
ed
A
us
cr
ip t
Figure 8
Page 84 of 96
3 90 140
190
190
190
240
240
240
290
290
340
340
390
390
440
440
490
490 Ca (wt.%) 10
15
C
Ti (wt.%)
0.1 0 90
G
90
1
2
K (wt.%) 1 2
3 0 90
190
240
240
290
290
290
340
340
340
390
390
390
440
440
440
490
490
490 3 0.0 90
H
P (wt.%) 0.5
1.0
I
Mn (wt.%)
90
0 0.5 90
0.0
140
140
140
140
140
190
190
190
190
190
240
240
240
240
240
290
290
290
290
340
340
340
340
390
390
390
390
390
440
440
440
440
440
490
490
490
490
490
290 340
E
2
Mg (wt.%) 3 5
140
140
M an 0
D
190
ed
5
Na (wt.%)
ce pt
90
F
0
cr
Al (wt.%) 1 2
140
Ac
Depth (m)
B
140
0
Depth (m)
Fe (wt.%) 10 20 30 40 50 0 90
us
0 90
A
ip t
Figure 9
J
ΣREE (ppm) 50
100
Page 85 of 96
L
Nb (ppm) 2 4
6 0 90
140
140
140
190
190
190
240
240
240
290
290
340
340
390
390
440
440
490
490
Cr (ppm) 20
40 0 90
N
140
Ni (ppm) 10 20
O
30
0.0
140
240
240
290
290
290
340
340
340
390
390
390
440
440
440
490
490
ed
M an
190
SHlpnomelane-‐rich tuffaceous mudrock
Calcareous mudrock
1.0
90
190
490
Mo (ppm) 0.5
sHlpnomelane mudrock
ce pt
Joffre BIF
M
ip t
60 0 90
cr
Zr (ppm) 20 40
Ac
Depth (m)
90
K
us
0
Page 86 of 96
us
cr
ip t
Figure 10
M an
10
Maximum
ed
REY/PAAS
1
ce pt
Minimum
Average Dales Gorge BIF
Ac
0.1
Average Joffre BIF
0.01 La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Y
Ho
Er
Tm
Yb
Lu
Page 87 of 96
cr
ip t
Figure 11
us
Joffre BIF
1000
M an
S+lpnomelane mudrock
0.01
0.001
ce pt
1
0.1
S+lpnomelane-‐rich tuffaceous mudrock
ed
10
Upper con+nental crust
Ac
PM normalised
100
Rb
Ba
Th
U
K
Nb
La
Ce
Sr
Nd
P
Hf
Zr
Na
Sm
Eu
Ti
Y
Yb
Lu
V
Ni
Cr
Page 88 of 96
δ56Fe -‐1.00 50
0.00
1.00
2.00
us
Joffre BIF Calcareous mudstone
cr
ip t
Figure 12
100
M an
Stilpnomelane-rich tuffaceous mudstone Stilpnomelane mudstone
150
ed ce pt
250
300
400
450
500
Ac
350
MOR fluids
Depth (m)
200
Igneous rocks
Page 89 of 96
cr
ip t
Figure 13
10$
us
Wi2enoom&BIF&(Dales&Gorge)&
Marra&Mamba&BIF&(upper&part)&
M an
Marra&Mamba&BIF&(lower&part)& Joffre&BIF&(this&study)&
Wt.%$
1$
ce pt
Al2O3& Al 2O3$
MgO&$ MgO
CaO&$ CaO
Na2O& Na 2O$
K2O& K 2O$
TiO2& TiO 2$
PP2O5& 2O5$
Ac
0.01$
ed
0.1$
Page 90 of 96
us
cr
ip t
Figure 14
M an
Gain'of'iron'
Altered'BIF'(Mt.'Tom'Price)' Marra'Mamba'BIF'(upper'part)'
ed
Fe2O3"
Marra'Mamba'BIF'(lower'part)'
ce pt
Loss'of'silica'
Ac
SiO2"
0"
10"
20"
30" Wt.%"
Wi;enoom'BIF'(Dales'Gorge)' Joffre'BIF'(this'study)'
40"
50"
60"
70"
Page 91 of 96
ip t
Figure 15
80
B
us
ΣLREE
cr
70
50 40
Intermixed volcanogenic detritus
30 20 10 0 0.01
0.1
1
Al (wt.%)
ed
M an
A
10
60
ce pt
(Pr/Yb)SN
1
Dales Gorge BIF Kuruman BIF S?lpnomelane mudrock
Ac
0.1
Joffre BIF (this study)
(DD98-‐28, this study)
Intermixed volcanogenic detritus
S?lpnomelane-‐rich tuffaceous mudstone (DD98-‐6, this study)
Calcareous mudstone (DD98-‐5, this study)
0.01 1
10
100
ΣREE
Page 92 of 96
cr
ip t
Figure 16
0.5
1
1.5
2
2.5
B
1.2
250
Nega9ve La anomaly, no Ce anomaly
Posi9ve Ce anomaly
ed ce pt
1
0.9
Ac
Ce/Ce* shale normalized
1.1
0.8
0.7
150
M an
Depth (m)
us
50
Eu/Eu*shale normalized
0.8
450
No Ce or La anomaly
550
Nega9ve Ce anomaly Posi9ve La anomaly, no Ce anomaly
A 0.7
350
0.9
1
1.1
1.2
Pr/Pr* shale normalized Page 93 of 96
ip t
Figure 17
cr
1000
Flood basalt
M an
us
Tuff (Weeli Wolli Fm.)
100
Zr (ppm)
Tuff (Dales Gorge)
Joffre samples Zr = 1662.4TiO21.6298
Average upper con;nental crust
Subaerial basalt on cont. crust
Average Fortescue shales Joffre BIF
Pillow basalt
Dolerite
Calcareous mudstone
Submarine basalt
s;lpnomelane-‐rich tuffaceous mudstone s;lpnomelane mudrock Joffre tuffaceous mudrock
ed
10
Woongarra rhyolite
0.1
Ac
1 0.01
ce pt
Dales Gorge S-‐bands Zr = 289.43TiO21.2839
1
10
TiO2 (wt.%)
Page 94 of 96
14
cr
16
ip t
Figure 18
n = 42
A
Joffre BIF
us
12 10
M an
8 6 4 2 0
0
0.5
ed
-‐0.5
1
1.5
δ56Fe
ce pt
20 18
B
n = 40
Dales Gorge BIF
16
14
12
Ac
10
8 6
4
2 0 -‐1
-‐0.5
0
0.5
1
1.5
δ56Fe Page 95 of 96
Pyroclas$c material (distal felsic volcanoes)
us
A
cr
Surface seawater: *High nutrient input from upwelling and diluted ash par$cles. *Leads to photoautotrophic Fe(II)-‐oxidia$on and Fe(III) deposi$on
ip t
Figure 19
Sea level
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
M an
Pho$c zone
Fe(III)-‐ hydroxide
Upwelling Hydrothermal fluids: Fe(II) ~ 4-‐20 ppm REY(solute) Eu(II) CH4, H2, H2S
δ56Fe > 0 Si(OH)4
Land
Joffre BIF
B
Sea level
Fe(II) ~ 1-‐3 ppm
Ac
Pho$c zone
ce pt
ed
Con$nental pla9orm
Surface seawater: *Low nutrient input only from diluted ash par$cles.Diminishes bacterial ac$vity and Fe(II) oxida$on. *Silica and siderite background deposi$on
Pyroclas$c material (distal felsic volcanoes)
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
Si(OH)4 FeCO3
Fe(II)-‐silicates?
Land
Joffre BIF Con$nental pla9orm
Page 96 of 96