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he Joffre banded iron formation, Hamersley Group, Westernustralia: Assessing the palaeoenvironment through detailedetrology and chemostratigraphy
asmus Haugaarda,∗, Ernesto Pecoitsb, Stefan Lalondec, Olivier Rouxelc, Kurt Konhausera
Department of Earth and Atmospheric Sciences, University of Alberta, Edmonton, Alberta T6G 2E3, CanadaEquipe Géobiosphère, Institut de Physique du Globe-Sorbonne Paris Cité, Université Paris Diderot, CNRS, 1 Place Jussieu, 75238 Paris, FranceUniversite Europe ene de Bretagne, Institut Universitaire Europe en de la Mer, Plouzane 29280, France
r t i c l e i n f o
rticle history:eceived 8 March 2015eceived in revised form8 September 2015ccepted 17 October 2015vailable online 19 November 2015
eywords:alaeoproterozoicamersley Group
offre banded iron formationeawater chemistryrovenancetilpnomelane
a b s t r a c t
The Joffre Member of the Brockman Iron Formation is by volume the largest single known banded iron for-mation (BIF) in the world. Here we present detailed petrology and chemostratigraphy through the entire355 m core section of this ∼2.45 billion year old unit. Oxide BIF and silicate–carbonate–oxide BIF dominatethe lithology, with minor amounts of interbedded stilpnomelane mudrock, stilpnomelane-rich tuffaceousmudrock and calcareous mudrock. Besides chert and magnetite, the prominent mineralogy is riebeckite,ankerite, hematite, stilpnomelane and crocidolite. The BIF is characterized by an average of 50 wt.% SiO2
and 44.5 wt.% Fe2O3 and an overall low abundance of Al2O3 (<1 wt.%), TiO2 (<0.04 wt.%), and trace metalssuch as Cr (<10 ppm), Ni (<5 ppm) and Mo (<0.5 ppm). It has a high
∑REE (rare earth element) content
(up to 41 ppm) and a fractionated shale-normalized (SN) seawater REY (rare earth element + yttrium)pattern having an enrichment of HREE (heavy rare earth elements) relative to LREE (light rare earthelements) with an average (Pr/Yb)SN of 0.24. The REY patterns also show a positive LaSN anomaly, noCeSN anomaly and a weakly developed positive YSN anomaly. Iron isotopes (�56Fe) with positive �56Fevalues of +0.04‰ to +1.21‰ suggest that a large part of the hydrothermal iron was partly oxidized in theupper water column and subsequently precipitated as ferric oxyhydroxides. No epiclastic grains havebeen found; rather submarine hydrothermal fluids and fine-grained volcanogenic detritus controlled BIFchemistry. The former source is reflected through a constant positive EuSN anomaly throughout the core(average EuSN anomaly of 1.6 with a peak of 2.1 between 100 and 155 m depth), while the latter sourceis best reflected through the stilpnomelane-rich tuffaceous mudrock consisting of volcanic ash-fall tuffwith relict shards set in a stilpnomelane matrix. The mudrock is overlain by well-preserved wavy lam-inae and laminae sets of stilpnomelane microgranules that likely originated from re-worked volcanicash formed either on the seafloor or in the water column prior to deposition. An enriched HREE-to-LREEpattern, a high iron content (∼30 wt.%), and a �56Fe value of +0.59‰ collectively imply that the mudrockfacies interacted with the Fe-rich seawater prior to deposition. The TiO2–Zr ratio of the BIF and the associ-ated mudrocks suggest a felsic-only-source related to the same style of volcanics as the slightly youngerWoongarra rhyolites. Given the observation that the dominant control on the seawater chemistry wasassociated with felsic volcanics, we speculate that the fine-grained pelagic ash particles may have sourcedbio-available nutrients to the surface water. This would have facilitated enhanced biological productivity,
56
including bacterial Fe(II)-oxidation which is now recorded as the positively fractionated Fe iron oxideminerals in the Joffre BIF. Alongside submarine hydrothermal input to the basin, the dominant controlon the ocean chemistry seems to have been through volcanic and pyroclastic pathways, thereby makingthe Joffre BIF poorly suited as a chemical proxy for the study of atmospheric oxygen and its weatheringimpact on local landmasses.
recipitated from seawater throughout much of the Archaean andalaeoproterozoic (3.8–1.85 Ga). They are also, more often thanot, laminated, with banding observed on a wide range of scales,
rom coarse macrobands (meters in thickness) to the characteris-ic mesobands (centimeter-thick units) by which they are typicallyefined (i.e., banded iron formation), to microbands (millimetero submillimeter). They typically contain low concentrations ofl2O3 (<1 wt.%) and incompatible elements (Ti, Zr, Th, Hf and Sc20 ppm), which indicate minimal detrital input to the depositionalasin, although this does not hold for all type of iron formationssee Bekker et al., 2010 for review). For instance, granular ironormations (GIF) typically lack banding and are made of granulesf chert and iron oxides or silicates with early diagenetic chertement filling pore space. Their texture implies that they formedn high-energy environments, with the granules being derived byedimentary re-working of iron-rich clays, mudstone, arenites, andven stromatolites (e.g., Ojakangas, 1983; Simonson and Goode,989).
Toward the end of Archaean, the marine depositional settingor BIF formation changed from depositional basins with rapidhermal subsidence and deposition of large volume of interbed-ed volcanic and volcanogenic greywackes (e.g., Lowe and Tice,007; Haugaard et al., 2013), to a more stable style of sedimen-ation in extensive shallow marine basins along stable continentallatforms (e.g., Taylor and McLennan, 1981; Eriksson et al., 2001;ondie, 2004; Barley et al., 2005). BIF deposited in the former set-ing are considered Algoma-type, whereas BIF formed in the latteretting are Superior-type. The latter includes the major BIF of thearliest Paleoproterozoic, such as the Hamersley Group BIF (Gross,980).
The mineralogy of BIF from the best-preserved sequencess remarkably uniform, comprising mostly chert, magnetite,ematite, Fe-rich silicate minerals (stilpnomelane, greenalite, min-esotaite, and riebeckite), carbonate minerals (siderite, ankerite,alcite, and dolomite), and minor sulfides (pyrite and pyrrhotite);he presence of both ferric and ferrous minerals gives BIF an aver-ge oxidation state of Fe2.4+ (Klein and Beukes, 1992). It is generallygreed that none of the minerals in BIF are primary. Instead, theinerals reflect significant post-depositional alteration under dia-
enetic and metamorphic conditions (including, in some cases,ost-depositional fluid flow). The effect of increasing temperaturend pressure is manifested by the progressive change in mineral-gy through replacement and recrystallisation, increase in crystalize and obliteration of primary textures (Klein, 2005; Bekker et al.,010).
The abundance of BIF in Precambrian successions was used inarly studies to argue for a largely anoxic atmosphere and oceanystem (e.g., Cloud, 1973; Holland, 1984) because the accumula-ion of such large masses of iron found in the form of Superior-typeIF required the transport of Fe(II); Fe(III) is essentially insolublet circumneutral pH values. Early studies invoked a continentalource of iron for BIF because Fe(II) would have been much moreobile in the absence of atmospheric O2 (e.g., James, 1954; Lepp
nd Golditch, 1964) and the continents were more mafic in com-osition (Condie, 1993). However, detailed studies that followed inhe Hamersley basin, Western Australia, suggested that the amountf iron deposited there was on the order of 1 × 1013 t (Trendall andlockley, 1970; Trendall and Blockley, 2004). This estimate wouldave required rivers the size of the modern Amazon to transportrders of magnitude more iron than they do today. This led Holland1973) to suggest that iron was instead sourced from deep marineaters and supplied to the depositional settings via upwelling.
ecently, however, a new model based on Fe- and Nd-isotopesuggests that a large part of the iron in BIF was continental andobilized by microbial Fe(III) reduction and transported through a
enthic iron shuttle to the BIF depositional basin (Li et al., 2015).
esearch 273 (2016) 12–37 13
Based on rare earth element (REE) composition of BIF, it isnow generally accepted that deep-sea hydrothermal processes arethe most likely source of Fe. Shale normalized (SN) europium(Eu) anomalies have been central in the use of REE to trace thehydrothermal input. Eu enrichment in chemical sedimentary rocksprecipitated from seawater indicates a strong influence of high-temperature hydrothermal fluids on the seawater dissolved REEload (e.g., Klinkhammer et al., 1983; Derry and Jacobsen, 1988,1990). It is generally assumed that Fe and REE will not be fraction-ated during transport from spreading ridges or other exhalativecenters, and, therefore, a strong positive EuSN anomaly indicatesthat the iron in the BIF precursor sediment was hydrothermallyderived (e.g., Slack et al., 2007). In addition to REE concentrations,Sm–Nd isotopes have been used to constrain REE and Fe sourcesto seawater (e.g., Miller and O’Nions, 1985; Derry and Jacobsen,1990; Alibert and McCulloch, 1993). The Archaean and Palaeopro-terozoic oceans were likely strongly heterogeneous in their εNd(t)values, with +1 to +2 values typical of the deep-waters dominatedby hydrothermal sources and lower values, down to −3, typical ofshallow-waters dominated by terrestrial sources (Frei et al., 1999,2008; Frei and Polat, 2007).
Within the large Superior-type BIF, the presence of diage-netic to low-grade metamorphic phyllosilicate minerals, such asgreenalite [(Mg, Fe)3Si2O5(OH)4] and stilpnomelane [K0.6(Fe2+,Fe3+, Mg)6Si8Al(O, OH)27·nH2O], mostly occur as dense bandsinterbedded with chemical precipitated minerals such as amor-phous silica and ferric hydroxides (Ayres, 1972; Morris, 1980).The precursor of stilpnomelane are believed to have been aniron(III)-rich clay (smectite) derived from volcaniclastic sourceslikely of basaltic provenance (Trendall and Blockley, 1970; Ewersand Morris, 1981; Pickard, 2002; Krapez et al., 2003). Where allprimary minerals have been overprinted by diagenesis and low-grade metamorphism, preservation of primary textures is rare.As for stilpnomelane, unique preservation of silt-sized micro-granules, or spheroids, opens up the possibility to study theprecursor sediment. These stilpnomelane microgranules are rel-atively uncommon but has been observed in Superior-type BIFin the Lake Superior region, Canada (Van Hise and Leith, 1911),in the Kuruman Iron Formation, South Africa (Beukes, 1973) andin the Brockman Iron Formation, Western Australia (Ayres, 1972;Krapez et al., 2003). Recently, Rasmussen et al. (2013, 2014)presented well-documented lamina sets of stilpnomelane micro-granules from the Dales Gorge BIF of the Brockman Iron Formationand proposed that they were generated by flocculation of iron-rich, Al-poor hydrous silicates either on the seabed or within thewater column and subsequently shaped and reworked by densitycurrents.
In this work, we investigate the ∼2.45 Ga (Pickard, 2002)Joffre Member from the Brockman Iron Formation. This is thesingle largest known BIF worldwide, containing approximately4.3 × 1013 t of iron at the time of deposition (Trendall and Blockley,2004). This laterally extensive Superior-type BIF, therefore, rep-resents the composition of a large volume of ocean water. Weconducted detailed petrologic and geochemical analyses of a coresection drilled through the entire ∼355 m of stratigraphic depthof the Joffre BIF. Unlike the well-explored Dales Gorge BIF (e.g.,Ayres, 1972; Ewers and Morris, 1981; Krapez et al., 2003; Pickardet al., 2004; Pecoits et al., 2009; Rasmussen et al., 2013, 2014,2015), this key Hamersley BIF has not previously been analyzedat high-resolution, nor has a detailed comparison of chemostratig-raphy between different lithologies through a complete successionever been attempted. Furthermore, by directly preceding the Great
Oxidation Event (GOE) and the marked increase in Cr dissolu-tion found in the Weeli Wolli Formation (see Konhauser et al.,2011), the Joffre BIF potentially covers the transition from ananoxic to a partially oxygenated Earth. As such, high-resolution
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etrography, geochemistry and isotopic studies will provide vitalnformation about the evolution of the sediment, the depositionalasin and, in particular, the source inputs controlling the seawateromposition.
. The Hamersley Group
The Joffre BIF is a member of the ∼620 m thick Brockman Ironormation which makes up part of the 2.63–2.45 Ga Hamersleyroup (Trendall et al., 2004). The Hamersley Group compriseslmost 2.5 km of consecutive sedimentary and volcanic rocksocated within the ca. 80,000 km2 Hamersley Province of the Pil-ara craton in North West Australia, approximately 1100 km northf Perth (Fig. 1). In the lower part, it consists of dolomite, shale andIF, while the upper part consists of dolerite, various lava types andIF with minor amounts of tuffs and shales (Trendall and Blockley,970). Underlying the Hamersley Group, is the 2.78–2.63 Ga (Arndtt al., 1991) Fortescue Group, which consists of flood basalts andhyolites. These volcanics were laid down on the uplifted androded Pilbara block (Trendall, 1968). This volcanic succession mayontain remnants of a Large Igneous Province (LIP) as suggested byrnst et al. (2004).
The Brockman Iron Formation of the Hamersley Group is dividednto four sub-lithostratigraphic units, namely the lowermost Dalesorge Member (BIF), the Whaleback Shale Member, the Joffreember (BIF), and the uppermost Yandicoogina Shale Member
Fig. 1B). After deposition, these laterally extensive BIF have allxperienced minor folding and basinal uplift along with low-gradeegional metamorphism – from burial prehnite–pumpellyite facieso greenschist facies (Smith et al., 1982).
ig. 1. (A) The general geology of the Pilbara craton with the location of drill hole DD98. Seron formations and associated rocks of the Hamersley Group. Significant other BIFs are thhe important U–Pb zircon ages from individual tuff layers. Note the age of 2454 ± 3 Ma f
esearch 273 (2016) 12–37
Geochronological constraints (see Fig. 1B) on the BrockmanIron Formation were established by Pickard (2002), and absoluteU–Pb zircon ages with interpolated stratigraphic age boundariesof the Hamersley Group are presented in Trendall et al. (2004).The best analytical age estimates for the deposition of the JoffreBIF is 2454 ± 3 Ma (Pickard, 2002). This age has been establishedby SHRIMP U–Pb zircon ages analyzed on 19 zircon grains frominterbedded tuffaceous mudrock facies at the top of the Joffre BIF(Fig. 1B). The best depositional age estimate for the base of the suc-cession is 2459 ± 3 Ma established by only 6 zircon grains (Pickard,2002, Fig. 1B). Without taking the uncertainties into consideration,there are 166 m between the above two age peaks, which poten-tially yields a compacted sedimentation rate of 33 m/million year(Pickard, 2002).
General consensus exists regarding the depositional modelof the Brockman Iron Formation. According to this model, thesuccession was deposited on a large, stable, and clastic-starved,continental platform, which was influenced by episodic inputs offine-grained tuffaceous detritus (e.g., Gross, 1983; Morris, 1980;Krapez et al., 2003). Blake and Barley (1992) proposed a gradu-ally subsiding open-shelf developed within a backarc setting underinfluence of tuffaceous material sourced from a subduction-relatedmagmatic arc. In addition, Barley et al. (1997) found that deposi-tion of the Hamersley BIF were possibly linked to major submarinemagmatic plume activity in the form of a LIPs. Morris (1993) alsosuggested that the depositional environment for the Hamersley BIFincluded a steady source of silica and iron with minor lateral vari-
ation in the deposition, and a water depth that was deeper thanthe formation of GIF but shallow enough to form the large carbon-ate platforms. In the absence of any shoreline facies and the lackof siliciclastics within the BIFs, Morris and Horwitz (1983) further
e text for further explanation. (B) The general stratigraphy of the extensive bandede Marra Mamba BIF, the Dales Gorge BIF and the BIF of Weeli Wolli Formation. Noteor the upper part the Joffre BIF.
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rgued that BIF precipitation and deposition took place on an outerhelf that was isolated by a carbonate barrier.
. Rock core, sampling and analytical methods
The drill core, DD98SGP001 (diamond drillcore 1998 Silvergrasseak #001) was drilled as part of the Rio Tinto project in the Sil-ergrass Peak area (see Fig. 1A for location). The core samples ofD98SGP001 were obtained at the Rio Tinto core library in Perth,estern Australia. A total of 31 core samples (DD98-1 – DD98-30),
ach ∼30 cm long, were obtained through a total core length of54.5 m (94–448.5 m depth).
The core samples were split and one half was stored as futureeference material. A total of 40 thin sections were processed andxamined using reflected and transmitted light microscopy. Inddition, six thin section slabs were polished and carbon coatedor electron microprobe analyses. Backscatter electron images, ele-
ental distribution maps, EDS (energy dispersive spectrometry)nd WDS (wavelength dispersive spectrometry) were obtainedith a JEOL Microprobe 8900 at the University of Alberta. The cur-
ent was set to ∼20 nA and the probe beam diameter was set to0 �m. Counting time was 20 s on peak, and 10 s on background.well time was 10 ms. Standardization to minerals with known ele-ent concentration were done on hematite, chromite, kaersutite
nd diopside for each 50 measurements.A total of 30 samples were selected for trace element anal-
ses. Approximately 4 cm slabs from each sample were dividednto chips and subsequently crushed on an agate mill. The crushedock powders were dissolved with HF + HNO3 and analyzed using aerkinElmer Elan6000 Quad-ICPMS (quadrupole inductively cou-led plasma mass spectrometer) at the University of Alberta.ccuracy and precision of the analytical protocol was verified with
he use of the well-established international whole-rock basalttandard (CRPG Nancy). Errors on this standard and on dupli-ates are both below 10%. Oxide interferences on Ce show thateO/Ce < 3% and any oxide interferences are therefore consideredegligible. For major elements, 17 samples were further analyzedy Code 4C (11+) Whole Rock Analysis-XRF at Activation Laborato-ies Ltd., Ontario, Canada. For iron isotopes, a total of 42 whole rockowder samples were selected.
Whole rock Fe isotope compositions were analyzed at the Frenchceanographic institution IFREMER, Brest campus, following previ-usly published methods (Rouxel et al., 2005; Rouxel et al., 2008).riefly, 50–100 mg of sample powder was digested overnight at0 ◦C in 4 ml 1:1 HF–HNO3 followed by 4 ml aqua regia, with com-lete evaporation in between. Samples were then taken up in 4 ml
N HCl, from which Fe was purified on Bio-Rad AG1X8 anion resin2 ml wet resin bed) using 6 N HCl for matrix elution followed by.24 N HCl for Fe elution. Fe isotope compositions were determinedsing a Thermo Scientific Neptune multicollector inductively cou-led plasma mass spectrometer operating at medium resolutiono resolve isobaric interferences such as 40Ar14N on 54Fe, 40Ar16On 56Fe, and 40Ar16O1H on 57Fe. Solutions were doped with Ni forass bias correction, introduced to the instrument using an Apex Q
esolvating nebulizer (Elemental Scientific, Omaha, NE, USA), andsample-standard bracketing’ was used for data normalization to ae isotope standard solution of IRMM-14 run before and after eachnknown. Geostandards BHVO-2 and IF-G yielded �56Fe values of.09 ± 0.09‰ and 0.65 ± 0.14‰, respectively, consistent with previ-us work (e.g., Planavsky et al., 2012), and repeated measurementsn = 59) of the reference material IRMM-14 (Taylor et al., 1992) con-
trained average internal precision over the analytical sessions toetter than ±0.065 (2SD).
Concentrations of REE and Y were shale normalized (SN) toost-Archaean Australia Shale (PAAS) after Taylor and McLennan
esearch 273 (2016) 12–37 15
(1985). Potential anomalies of La (La/La*SN) and Ce (Ce/Ce*SN)were obtained by the procedure proposed by Bau and Dulski(1996) using the combination of Ce/Ce*PAAS = Ce/(0.5*La + 0.5*Pr)and Pr/Pr*SN = Pr/(0.5*Ce + 0.5*Nd). If Ce/Ce* < 1 but Pr/Pr*SN ≈ 1 apositive La anomaly is evident. If Ce/Ce*SN < 1 but Pr/Pr*SN > 1.05 anegative Ce anomaly is evident. The Eu anomaly (Eu/Eu*SN) wascalculated as Eu/Eu*SN = Eu/[0.67Sm + 0.33Tb].
4. Petrography and mineralogy
4.1. Rock types
Based on the dominant mineralogy, we define five lithologicalsubdivisions within the Joffre BIF (Fig. 2): (1) oxide BIF (Fig. 3), (2)silicate–carbonate–oxide BIF (Fig. 4), (3) stilpnomelane-rich tuffa-ceous mudrock (Figs. 5 and 6), (4) stilpnomelane mudrock (Fig. 5A),and (5) calcareous mudrock. As shown in Fig. 2, a large portionof minerals, such as chert, magnetite, hematite, riebeckite, car-bonate and occasionally stilpnomelane exist, although in differentproportions within the first three lithologies. In particular, a tran-sitional shift exists between oxide BIF, which is the most dominantrock type, and silicate–carbonate–oxide BIF, which is the secondmost dominant rock type. Therefore, any clear petrographic splitbetween those rock types is not feasible. Magnetite is the mostabundant iron oxide phase and occurs in variable amounts in all ofthe rock types except in calcareous mudrock.
4.1.1. Oxide BIFThis unit is dominated by micro- and mesobands of chert,
magnetite and lesser amounts of hematite (Fig. 3A–D). Microcrys-talline (∼0.05 mm) chert appears both as mesobands (∼1–5 cm)and microbands (0.25–1 mm). The chert ranges in color from whiteto grayish and to a more red variety (jasperlitic) as a result of inter-stital hematite grains (Fig. 3A). It is often found as microbandswith various amount of hematite ± crocidolite ± carbonate, alter-nating with microbands of magnetite ± hematite. A few pure chertmicrobands that alternate with magnetite layers are observed spo-radically (Fig. 3A–C).
Major parts of the magnetite bands are black and opaquewhile minor parts of the bands are dark gray as a result of inter-layered chert microbands (Fig. 3A–C). The bands occur both asmesobands (1–3 cm thick) and microbands (down to 0.1 mm thick).The magnetite ranges from fine-grained to coarser-grained witha well-crystallized habit and it is commonly coarser-grained thancoexisting chert and hematite (Fig. 3D).
Hematite is found both as microcrystalline cement in relationto chert mirco- and mesobands (Fig. 3D) and as <0.1 mm micro-platy crystals (martite), which are locally found associated withmagnetite meso- and microbands (Fig. 3D). Martite is formeddue to secondary oxidation of magnetite and is thus often foundrelated to magnetite. Micro-platy hematite is also found “float-ing” in a more brownish to yellowish coherent cement probably ofa more goethitic composition. The microcrystalline hematite andchert microbands, consisting of red jasperitic mesobands, are moredominant in the middle section of the Joffre BIF. Some hematitegrains are associated with microbands of chert, riebeckite andaltered carbonate. This association is likely due to the result ofmicro-platy hematite replacing part of the chert and carbonateas suggested by Clout and Simonson (2005). Some sections of theoxide BIF furthermore reveal thin (<0.01 mm) microbands contain-ing brown microgranules of stilpnomelane, K(Fe2+, Mg, Fe3+)8(Si,Al)12(O, OH)27·n(H2O).
4.1.2. Silicate–carbonate–oxide BIFThis rock type is distinctive by having relatively more riebeckite
and carbonate and less iron oxides than the oxide BIF (Fig. 4A–C).
16 R. Haugaard et al. / Precambrian Research 273 (2016) 12–37
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Fig. 2. The main rock types and their mineralogy in the Joffre BIF. The majo
Chert mesobands and chert nodules display two different wavyicrobands with different grain size (Fig. 4D–F). One chert fraction
omprises 0.05 mm grains occasionally with a braided network ofematite and minor goethite. These wavy microbands have a thick-ess of ∼0.5 mm. The other chert fraction is finer grained, consistingf <0.02 mm chert grains (Fig. 4F). These microbands are up to 1 mmhick. The finer grained chert microbands are solely restricted tohombic ankerite and fibrous crocidolite (Fig. 4F). Generally, allhert grains are irregularly shaped and exhibit slight undulatoryxtinction. Only in few places has the chert been recrystallized tooarser grained (∼0.3–0.4 mm) fractions.
Alongside chert, dense micro- and mesobands of magnetite areound as a major constituent (Fig. 4B). A minor portion of the mag-etite displays secondary alteration and oxidation into hematitend other Fe oxides such as goethite (Fig. 4G and H).
After chert and magnetite, riebeckite, a sodium-rich amphiboleNa2(Fe2+)3(Fe3+)2(Si8O22)(OH)2), is the most abundant con-tituent, both within this rock type and within the entire coreFig. 4B, G and H). It occurs as dark blue, dense microbands thatange from 0.1 mm to 0.5 cm in thickness and as single (0.2 mmong) crystals (Fig. 4I). It is found predominantly within chert-ich microbands and at the interface between chert and iron oxideicrobands (Fig. 4G). Locally, riebeckite microbands are associ-
ted with magnetite microbands containing randomly dispersedrownish Fe oxide grains (Fig. 4H). The presence of riebeckite-ich fluid escape structures (riebeckite veins in Fig. 4A) indicatesemobilization from presumably a hydrous silica–iron–sodium geluring compression of the sediments. Riebeckite is often foundssociated with acicular and fibrous crocidolite (Fig. 4F and I).rocidolite mostly appears as thin, blue fibrous needles (bluesbestos) in close association with more massive riebeckite crys-als or bands. It also occurs within bands of transparent chert
icrobands where the long fibers of crocidolite tend to grow per-endicular to the bedding planes. Fibrous crocidolite tends to sprayut from the rims of the denser riebeckite (Fig. 4I). It coexists withhe wavy chert + carbonate microbands indicating, and as proposed
erals are furthermore listed in increasing abundances throughout the core.
by Miyano and Klein (1983), that the crocidolite has grown at theexpense of Fe–carbonate. For a more detailed description of bothriebeckite and crocidolite as a diagenetic product see Miles (1942),Ryan and Blockley (1965) and Miyano and Klein (1983).
Various compositional pale-brown carbonates occur through-out as individual crystals with a predominantly rhombic habitoccasionally displaying internal zoning. The carbonate crystalsmostly occur within microbands composed of very fine-grainedchert with riebeckite and crocidolite ± magnetite ± hematite. Semi-quantitative EDS observations of the carbonate reveal highelemental peaks corresponding to Ca and Fe, with minor peaks forMg and Mn. Based on this, we interpret the carbonates as belong-ing to the dolomite–ankerite series, with a predominantly ankeriticcomposition.
Talc- and chlorite-alteration plays a minor role in the JoffreBIF. However, small colorless and non-pleocroic needles of talc(presumably minnesotaite) are observed in relation to green topale-green chlorite microbands. These needles are oriented in thedirection of stretching, indicating that they have been growing dur-ing the main compaction phase of the BIF package. In addition, talcalteration is locally visible in between microlaminae of carbonate,magnetite and chert (Fig. 4J).
4.1.3. Stilpnomelane-rich tuffaceous mudrockThis rock type is volumetrically minor but it shows impor-
tant mineralogical features (see Fig. 5A). One of the characteristicsof this rock type is the intimate relation between tuff materialand stilpnomelane microgranules. The preservation of a 1–2 cmthick bed of pale-green tuff represents direct evidence of vol-canogenic ash-fall into the basin (Fig. 5B). Texturally, this bedconsists of recrystallized shards set in a very fine-grained, greenish-
brownish, stilpnomelane matrix (Figs. 5C and 6A). The shards arefeldspar-pseudomorphs likely formed from volcanic glass. Duringcompaction and burial metamorphism, the glass devitrified andrecrystallized into feldspar. Electron microprobe analysis shows
R. Haugaard et al. / Precambrian Research 273 (2016) 12–37 17
Fig. 3. Photos and photomicrographs illustrating the main petrographic characteristics of the oxide BIF with representative core samples (A and B) show-ing pale gray micro- and mesobands of alternating chert and magnetite, dark gray magnetite mesobands and reddish and bluish (�1 mm thin) microbands ofchert–hematite–riebeckite ± crocidolite. With white arrows, sedimentary slumping can be seen in the upper half of (A) while soft sediment deformation in a magnetitemicroband is seen by the presence of micro-flame structures in the top part of (B). Possible current generated sedimentary structures are presented in (C). The fact that thep originw microm agne
tl
bamlfisfl((bdieseucfiBps
lanar lamina immediately above the wavy band is undisturbed indicates a primaryavy bedding in the upper left corner of (C). (D) Coarse and fine grained magnetiteagnetite oxidation. Note the single grains of martite upper left. Ch = chert, Mag = m
hat the feldspar is almost 100% sanidine in composition, with veryittle iron, calcium and sodium (see Table 1).
In addition to the tuffaceous layer, stilpnomelane is cementedy chert forming plane- or wavy-lamina microbands (Figs. 5B, D–Fnd 6B–E). These bands are 0.25–1 mm thick and alternate withicrobands of almost pure chert with few grains of fibrous crocido-
ite and platy hematite. The microbands consist of either extremelyne-grained (down to 0.01 mm) stilpnomelane microgranules orpheroids (Fig. 5E) or as medium grained (0.5–1 mm) stilpnomelaneake- and lath-shaped aggregates that are occasionally radiatingFig. 5F and G). The microgranules are pale to dark green to browndepending on the Fe2+/Fe3+ ratio). Some brown stilpnomelane-richands also show evidence of shrinkage texture possibly caused byehydration of a precursor phase to stilpnomelane (Fig. 5F). EDS
mages in Fig. 6D and E show that elements, such as Al and K, arenriched in the stilpnomelane microlaminae. As for the shards, thetilpnomelane microgranules may often only be preserved wherearly diagenetic chert formation prevented compaction of the gran-les (see Rasmussen et al., 2013). However, throughout parts of theore, very fine and diluted laminae of these microgranules can beound interbedded with chert and iron oxide microbands making
t difficult to separate true oxide-BIF from weak silicate dominatedIF. The composition of the microgranules and the ash matrix areresented in Table 1 and Fig. 7, wherein a large part of the mea-urements plot within the stilpnomelane field.
of the two symmetric wavy features. Note also the two wavy features with internalbands with chert and very fine hematite. The latter likely is a product of secondarytite, Hem = hematite, Rbk = riebeckite.
Ankeritic carbonate (Fe-rich dolomite) occurs as individualcrystals having predominantly a rhombic habit with clear inter-nal zoning visible in some of the larger (0.2–0.4 mm) crystals(Figs. 5F and 6B). As within the silicate–carbonate–oxide BIF, con-siderable recrystallization of the carbonate crystals can be seen bythe perfect euhedral outline to neighboring minerals. The large,randomly distributed rhombic carbonate crystals are normally dis-persed in very fine-grained chert cement, along with stilpnomelane(Fig. 6B). EDS images shows that these carbonates are both Fe- andMg-rich (Fig. 6F and G). The likeliness that the carbonates havegrown during burial late stage diagenesis and metamorphism canbe seen in Fig. 6H. It shows a trail of stilpnomelane grains togetherwith single chert grains engulfed by a rhombohedral carbonatecrystal, indicating the secondary growth of the carbonate crystalafter both stilpnomelane and chert. Different styles of carbonategrowth can be viewed in Figs. 5D and 6C; both showing excellentpreserved lamina of single prismatic carbonate crystals. In con-trast to the single rhombic ankerite crystals, EDS observations ofthese carbonates reveal a more dolomitic (Mg-rich) composition.The distinct symmetry of these dolomite crystals were also foundby Morris (1993) in the Marra Mamba BIF (see Fig. 1B), where it
was suggested they were pseudomorphs after swallowtail gypsymcrystals, and as such, they represent shallow water in this vicinity.
Another type of stilpnomelane sedimentation is the existence ofultra-thin microbeds consisting of stilpnomelane with fragments of
18 R. Haugaard et al. / Precambrian Research 273 (2016) 12–37
Fig. 4. Photos and photomicrographs illustrating the main petrographic characteristics of the silicate–carbonate–oxide BIF with representative core samples (A and B) andthin section images (C–J). This rock type is dominated by chert + magnetite + riebeckite + ankerite + crocidolite ± stilpnomelane. (A) and (B) Dense mesobands composed ofalternating microbands of chert + riebeckite + crocidolite ± oxide can be seen in the middle part of (A) and the upper half of (B). Dense, dark gray magnetite micro- andmacrobands and pale gray chert mesobands with magnetite microbands are evident in core (B), which furthermore have mesobands of riebeckite with chert and hematitemicrobands (lower part). The white arrows in (A) illustrate weak mobilization of riebeckite microbands. (C) Thin section PPL image (with inset close up image) of planarriebeckite microbanding. (D) Chert mesobands and nodules containing wavy-laminated chert and riebeckite + carbonate + crocidolite microbands. The distinct chert noduleswere likely formed during compaction of the BIF package illustrated with the condensed internal microbands at the white arrow. These internal laminae are further illustratedin (E) and (F). Note in (E) the finer grained chert fraction is restricted to the carbonate and riebeckite–crocidolite laminae only with coarser grained chert in between. (G)PPL thin section image of blue riebeckite microbands underlain by red hematite and opaque magnetite microbands. (H) Riebeckite and magnetite microbands in relationt PPL io ematit( rred to
qTi
pste
o very fine-grains of other oxides (presumably hematite and goethite). (I) Close upccur occasionally in relation with magnetite. Ch = chert, Mag = magnetite, Hem = hFor interpretation of the references to color in this figure legend, the reader is refe
uarts, chlorite and other sheet minerals (likely muscovite, Fig. 5I).he stilpnomelane is represented as both groundmass and as singlendividual flakes.
In terms of accessory phases, few randomly dispersed cubic
yrite crystals appear within stilpnomelane laminae, whereas aingle-crystal lamina of pyrite is seen in Fig. 5G. In betweenhe ash bed and the well-preserved bed of dolomite crystals,lectron microprobe EDS analysis reveals a 0.5 mm thin lamina
mage of the fibrous bluish crocidolite in a cherty groundmass. (J) Fe-talc alteratione, Rbk = riebeckite, Ank = ankerite, Cr = crocidolite, Ox = oxides (hematite/goethite).
the web version of this article.)
containing zircon, ilmenite, monazite, pyrite and apatite (Fig. 6Cand I).
4.1.4. Stilpnomelane mudrock
Stilpnomelane occurs also as massive, almost opaque,
mesobands (Fig. 5A). Detailed petrography shows that theapparent structure-less bands are plane-laminated on minutescale (<0.1 mm). The bands, which have a sharp base and top, vary
R. Haugaard et al. / Precambrian Research 273 (2016) 12–37 19
Fig. 5. Photos and photomicrographs illustrating the main petrographic characteristics of the stilpnomelane-rich tuffaceous mudrock and stilpnomelane mudrock. (A)Typical core section influenced by various brown to greenish stilpnomelane mudrock mesobands alternating with thick white chert bands with internal wavy microbandsof stilpnomelane + ankeritic dolomite. Note only relative sparse representation of iron oxides and magnetite. Riebeckite microbands are visible in the middle and lowerpart of the core section. (B) Whole thin section PPL image showing stilpnomelane-rich tuffaceous mudrock having a bottom greenish tuff bed grading up in to plane- andwavy lamination of stilpnomelane microgranules ending with a magnetite band intermixed with stilpnomelane. (C) Close up XPL image of the tuff bed in (B) containingwell-preserved shards in a stilpnomelane-rich matrix. (D) A PPL image of the wavy stilpnomelane laminae set from (B). Each laminae set ends with almost pure laminae ofchert. (E) A close up PPL image of the single crystal dolomite bed in (B) with the stilpnomelane rich tuff bed underneath and wavy stilpnomelane microgranular laminae above.(F) A close up PPL image of the stilpnomelane microgranules. (G) PPL image of chert with stilpnomelane laminae and randomly distributed euhedral ankerite crystals. Notestilpnomelane occur both as microgranules and as flakes. (H) Stilpnomelane and chert laminae containing radiating flakes and silt size shrinkage texture of stilpnomelane. Asingle laminae of pyrite crystals is seen in the middle part of the image. (I) Ultra-thin microbed (in PPL) of stilpmomelane rich bearing quartz + chlorite + other sheet silicates(possible muscovite) illustrating another style of volcanic input to the BIF basin than in (B). The bed is from the oxide BIF. Ch = chert, Mag = magnetite, Hem = hematite,Rbk = riebeckite, Ank = ankerite, Cr = crocidolite, Ox = oxides (hematite/goethite), Stp = stilpnomelane, Chl = chlorite, Qtz = quartz, Py = pyrite, Dol = dolomite.
20 R. Haugaard et al. / Precambrian Research 273 (2016) 12–37
Fig. 6. Backscatter electron images and microprobe elemental maps of various textures and minerals from the stilpnomelane-rich tuffaceous mudrock. (A) Backscatter image ofFig. 5C, showing dark K-feldspar shards in a stilpnomelane-rich matrix. (B) Well-crystallised Ankeritic dolomite with internal zoning distributed in chert with stilpnomelanemicrogranules. (C) Tuff bed overlain by perfect crystal-shaped dolomite laminae with homogenous stilpnomelane and dolomite laminae on top. The dominating wavy-laminae of chert and stilpnomelane microgranules are seen in the upper part. (D and E) Elemental maps showing the distribution of Al and K across wavy micro-laminaeof stilpnomelane and chert. Bright blue (D) and bright green to yellow (E) represent stilpnomelane-rich laminae having relative higher Al and K compare with darker chertrich laminae. (F and G) Shows distribution of Fe and Mg across micro-laminae of chert and stilpnomelane with euhedral ankeritic dolomite crystals. Note the zonation ofhigher Fe content (brighter bluish) in the crystals in (F) and the relative high content of Mg (brighter greenish colors) in (G). Note also brighter zones with respect to Feand Mg are evident for the stilpnomelane-rich laminae. (H) A single ankerite crystal that encapsulates a laminae of single-grain stilpnomelane proving a very late growthof the carbonate. Note also the engulfment of various amount of chert grains. (I) A close up backscatter image of the bright laminae in (C). EDS analyses show that thislaminae contains various amount of heavy minerals such as zircon, monazite, ilmenite, pyrite and also apatite (not shown). Ch = chert, Mag = magnetite, Rbk = riebeckite,Ank = ankerite, Stp = stilpnomelane, Py = pyrite, Dol = dolomite, Zrn = zircon, Mnz = monazite, Ilm = ilmenite. (For interpretation of the references to color in this figure legend,the reader is referred to the web version of this article.)
R. Haugaard et al. / Precambrian Research 273 (2016) 12–37 21
F showm entsa IF (Kle
iapmc
4
mgficmto
4
ofhtpaocib
F
ig. 7. Electron microprobe data from the stilpnomelane-rich tuffaceous mudrockicrogranules and 10 of the ash matrix (see also Table 4). Almost all of the measurem
nd greenalite fields are composed from microprobe data from the Marra Mamba B
n thickness from 0.5 to 2.0 cm, and contain disseminated quartznd K-feldspar fragments that seemingly indicate volcanogenicrovenance. The stilpnomelane mudrock bands are volumetricallyinor throughout the BIF but are often found in relation to
hert-rich and magnetite-poor sections of the core (Fig. 5A).
.1.5. Calcareous mudrockAt the top part of the core section, a noteworthy ∼5 cm thick
esoband of calcareous mudrock occurs. It contains very fine-rained white to pale-gray calcite–dolomite that grades into ane-grained greenish chloritized material. Angular fragments ofarbonate are dispersed within the former, whereas fragments ofostly quartz are dispersed in the latter. The precursor sediment to
he calcareous mudrock has erosionally truncated the underlyingxide BIF.
.2. Sedimentary structures
Primary sedimentary structures are generally absent through-ut the Joffre BIF. This is in full agreement with other core sectionsrom the Brockman Iron Formation (e.g., Trendall, 2002). However,ere we present possible current-generated sedimentary struc-ures developed prior to compaction and lithification of the BIFackage. Around the middle-part of the oxide BIF core (Fig. 3Bnd C), a chert rich mesoband has a wavy appearance composed
f two coherent and symmetric, concave-down, structures eacha. 1 cm thick. The interesting observation is that the planar lam-nae immediately above and within the trough is not disturbedy the underlying wavy bedding, making it difficult to interpret
ig. 8. Major bulk element data from the Joffre BIF plotted in ternary diagrams. (A) SiO2–
ing a SiO2–FeO + MgO–Al2O3 + K2O ternary diagram with 27 measurements of the fall within the stilpnomelane compositional field. The stilpnomelane, minnesotaitein and Gole, 1981) and crosschecked with mineral data from James (1954).
it as a post-depositional process. Rather, the concave-down struc-tures were generated prior to the deposition of the above laminae.Another, although weaker developed wavy structure can be seenin the upper left corner of Fig. 3C. Note the planar laminationjust above the wavy features. Weak sediment slumping is alsoseen between bedding planes of chert mesobands and iron oxidemesobands (white arrow Fig. 3A). In few of the magnetite bands,soft sediment deformation is evident by the presence of micro-flame structures (white arrow Fig. 3B). The latter two structuresare most likely of pre-diagenetic origin.
The chert bedding in the Joffre BIF is normally plane- to wavylaminated, but occasionally the chert forms lenses or nodules(Fig. 4A and C). These chert nodules (up to 1 cm thick) still con-tain internal wavy riebeckite–carbonate microbands. The featuresmost likely developed during burial metamorphism where the lat-eral termination happened by compaction of more iron-rich bandsabove and below the original chert layer (white arrow in Fig. 4C).The internal riebeckite–carbonate laminae get compacted from theinner part of the nodule to the outer part. As such, these chert nod-ules should be interpreted as a result of syn-compaction rather thanerosional features.
5. Bulk rock geochemistry
5.1. Major and trace elements
Geochemical data for major and trace elements are presentedin Tables 2 and 3. SiO2–Fe2O3–Al2O3 and SiO2–Fe2O3–CaO + MgOternary diagrams are shown in Fig. 8A and 8B. The evolution of
Fe2O3(t)–Al2O3 and (B) SiO2–Fe2O3(t)–CaO + MgO. See text for further explanation.
22 R. Haugaard et al. / Precambrian Research 273 (2016) 12–37
major
sFsvro((t
Fig. 9. Evolution with depth for selected
elected major and trace elements with depth are presented inig. 9A–J. Relative to chemically precipitated elements, such asilica and iron, the Al2O3 content in the BIF samples (Fig. 8A) isery low (<1 wt.%). In contrast, CaO and MgO are more elevated,eflecting the appearance of well-developed carbonates through-
ut the core (Fig. 8B). A constant low input of Al (Fig. 9B), TiFig. 9C) and high field strength elements such, as Zr and NbFig. 9K and L), is noteworthy. By contrast, elevated abundances ofhese insoluble elements are found within the calcareous mudrock,
and trace elements. See text for details.
the stilpnomelane-rich tuffaceous mudrock and the massive stilp-nomelane mudrock. More variable concentrations with depth areobserved for Fe, Na, Mn and REEs (Fig. 9A, G, I and J). Phospho-rous (Fig. 9H) shows low to moderate concentrations throughoutthe core except for two BIF outliers, which have up to four times
as much P as the other BIF samples likely as a result of very finegrained apatite.
In terms of elemental correlations (diagrams not shown), sig-nificant R-values are found for Al vs. Ti and Ti vs. Zr, indicating
R. Haugaard et al. / Precambrian Research 273 (2016) 12–37 23
Tab
le
1M
ajor
elem
ent
oxid
es
mea
sure
d
on
the
mic
rogr
anu
les
in
the
BIF
(DD
98-2
6A)
and
in
the
stil
pn
omel
ane-
rich
tuff
aceo
us
mu
dro
ck
(DD
98-6
)
and
the
com
pos
itio
n
of
the
shar
ds
and
the
shar
d
mat
rix
wit
hin
the
ash
bed
in
DD
98-6
.N
ote
fin
al
wt.
%
wit
h
H2O
calc
ula
ted
from
OH
con
ten
t.
The
com
pos
itio
n
of
the
mic
togr
anu
les
are
illu
stra
ted
in
Fig.
7.
Elem
ent
DD
98-2
6Aa
(mic
rogr
anu
les)
DD
98-6
b
(mic
rogr
anu
les)
DD
98-6
b
(sh
ard
s)D
D98
-6b
(sh
ard
mat
rix)
1
2
3
4
5
Ave
Ave
(n
=
22)
Ave
(n
=
40)
1
2
3
4
5
6
7
8
9
10
Ave
SiO
246
.24
47.0
2
45.5
7
45.6
5
45.3
6
45.9
7
46.5
4
64.3
4
47.7
7
46.4
1
48.5
2
47.5
7
47.4
0
47.3
8
49.1
4
50.9
4
48.1
2 46
.99
48.0
2Ti
O2
0.00
0.06
0.01
0.01
0.02
0.02
0.03
0.01
0.11
N.D
.
0.05
0.02
0.06
0.02
0.09
N.D
.
0.04
0.03
0.04
Al 2
O3
2.85
2.85
2.77
2.86
2.71
2.81
4.67
18.3
0
5.19
4.43
5.66
4.45
4.47
4.51
5.93
7.22
4.92
4.20
5.10
FeO
26.0
7
26.3
5
25.4
8
25.3
2
26.2
0
25.8
8
25.3
5
0.56
21.4
3
22.5
0
20.6
6
22.7
2
23.0
0
22.9
3
21.0
6
17.4
9 22
.50
23.2
6
21.7
5Fe
2O
37.
24
7.32
7.08
7.03
7.28
7.19
7.04
–
5.95
6.25
5.74
6.31
6.39
6.37
5.85
4.86
6.25
6.46
6.04
Cr 2
O3
0.02
0.02
0.00
0.01
0.00
0.01
0.01
–
0.01
0.00
0.00
0.00
0.05
0.00
0.04
0.03
0.00
0.03
0.02
Mn
O
0.12
0.15
0.13
0.13
0.19
0.14
0.09
0.02
0.07
0.07
0.05
0.07
0.04
0.06
0.05
0.07
0.04
0.01
0.05
MgO
6.86
6.83
7.13
6.83
7.35
7.00
5.70
0.04
7.03
6.95
6.51
7.18
7.18
7.58
6.58
5.58
7.28
7.46
6.93
CaO
0.03
0.04
0.05
0.03
0.03
0.04
0.09
0.00
0.04
0.04
0.04
0.05
0.05
0.03
0.04
0.04
0.05
0.04
0.04
Na 2
O
0.21
0.15
0.32
0.54
0.20
0.28
0.37
0.05
1.28
0.89
0.76
0.69
0.98
1.12
0.95
0.84
0.55
0.64
0.87
K2O
2.07
1.85
2.48
3.62
1.99
2.40
2.83
16.5
9
3.85
2.94
4.21
2.77
2.66
2.85
4.10
5.43
3.70
3.26
3.58
H2O
calc
8.69
8.80
8.53
8.10
8.96
8.62
8.05
7.34
7.60
7.26
7.83
7.82
7.87
7.39
6.56
7.79
7.93
7.53
Tota
l
100.
40
101.
44
99.5
4
100.
13
100.
29
100.
36
100.
79
99.9
1
100.
08
98.0
7
99.4
5
99.6
6
100.
08
100.
72
101.
21
99.0
5
101.
25
100.
31
99.9
8
aB
IF.
bSt
ilp
nom
elan
e-ri
ch
tuff
aceo
us
mu
dro
ck.
Fig. 10. Shale normalized (PAAS) REY pattern from the Joffre BIF (gray area rep-resents 27 samples) showing pronounced fractionated HREE to LREE pattern alongwith a pronounced EuSN anomaly, (Eu/Eu*)SN. Note the high abundances of REY rela-
tive to the underlying Dales Gorge BIF. Average data for Dales Gorge BIF from Pecoitset al. (2009).
adsorption of these components to the fine-grained pelagic clayfraction. The absence of correlation between �REE and P or Zr sug-gests minor contribution from sedimentary apatite, monazite andzircon. Interesting, soluble elements, such as K and Ba, are mod-erately correlated with more conservative elements, such as Al, Tiand Nb. For example Al vs. K yields an R-value of 0.73. Sodium (Na),which shares a similar degree of mobility as K and Ba, is unrelatedto Al (R = −0.18). The Al vs. the REE is relatively uncorrelated. How-ever, the variation in LREE with Al, represented by the Al vs. Pr, arebetter correlated (R = 0.67) than the variation in HREE with Al, asrepresented by Al vs. Yb (R = 0.27). Also phosphorous, is unrelatedto Al2O3 (R = 0.24).
The maximum, minimum and average shale normalized REEpatterns for the 27 BIF samples are displayed in Fig. 10. Theabsolute concentrations of REEs in the Joffre BIF (average of17 ppm) are highly elevated relative to the undelaying DalesGorge BIF, although the REE patterns in both BIF are very similar(Pecoits et al., 2009). A general HREE to LREE enrichment (average(Pr/Yb)SN = 0.24), together with a pronounced EuSN anomaly (aver-age (Eu/Eu*)SN = 1.56) is observed (Fig. 10, Table 3). Part of the BIFpackage shows a positive YSN anomaly, (Y/Ho)SN, reflecting moreY to Ho in the seawater column (Fig. 10, Table 3). However, theYSN anomaly is not significant enough to be clearly evident in theaverage REY pattern.
In a primitive mantle-normalized spider diagram (Fig. 11), aver-age Joffre BIF is plotted with relevant associated lithologies. Thereis the same variation for many elements between the BIF and thethree Joffre lithologies and, although less pronounced with respectto the upper continental crust. Exceptions include the soluble ele-ments that tend to be concentrated in seawater, and hence in theBIF (e.g., P, Na and Sr). Notably, the P anomaly seen in the BIF is evenhigher than the average continental crust. The negative anomaliesof the Nb and Ti, two elements that are depleted in newly gener-ated continental crust, show an even larger negative anomaly inthe BIF likely due to the high insolubility of those elements in sea-water (Fig. 11). Similarly, insoluble elements, such as Th, Zr and Hf,which have positive anomalies for the associated lithologies, showa trough in the spider diagram for the BIF.
5.2. Fe-isotopes
Fe isotopes, as represented by �56Fe values and their standarddeviations, are presented in Table 4 and are plotted with the
24 R. Haugaard et al. / Precambrian Research 273 (2016) 12–37
Fig. 11. Primitive mantle normalized spider diagram showing the Joffre BIF and the intermixed massive stilpnomelane mudrock and the stilpnomelane-rich tuffaceousmudrock. Blue line shows the average upper continental crust from Rudnick and Gao (2003). For the BIF, note the opposite pattern with very low abundances among majorparts of the insoluble elements (Th, Nb, Zr, Hf, Ti) and the distinctive positive anomalies of soluble elements (etc., P, Na and Sr) relative to the stilpnomelane mudrock andstilpnomelane-rich tuffaceous mudrock. Note also the high positive plateau for the HREEs (Y, Yb, Lu) both for the BIF and for the intermixed tuffaceous detritus. A significantdrop in the trace metals (V, Cr and Ni) is seen in all of the lithologies. (For interpretation of the references to color in this figure legend, the reader is referred to the webversion of this article.)
Fig. 12. Shows the evolution of the Fe isotopes (�56Fe) throughout the Joffre BIF. Fractionation mechanism during Fe oxidation is represented by more positive �56Fe valuesthan the typical values for igneous rocks and mid ocean ridge (MOR) fluids. Interesting to note is that the stilpnomelane-rich tuffaceous BIF has a high positive �56Fe valueof 0.74. See text for further details. Field of igneous rocks and MOR fluids from Sharma et al. (2001) and Johnson et al. (2003).
R. Haugaard et al. / Precambrian Research 273 (2016) 12–37 25
Table 2The major elementsas oxides with core depth in the Joffre BIF. Note the three different lithologies.
Depth (m) Element oxide SiO2 (wt.%) Al2O3 Fe2O3 MnO MgO CaO Na2O K2O TiO2 P2O5 LOI TotalDetection limit 0.01 0.01 0.01 0.001 0.01 0.01 0.01 0.01 0.01 0.01
a Calcareous mudrock.b Stilpnomelane-rich tuffaceous mudrock.c Stilpnomelane mudrock.
tratigraphic depth in Fig. 12. Inserted, for comparison, are aver-ge �56Fe values for common igneous rocks and average mid oceanidge (MOR) fluids (Sharma et al., 2001; Johnson et al., 2003).sotopic fractionation that occurs during Fe(II) oxidation is repre-ented by more positive �56Fe values than the typical values forgneous rocks and hydrothermal fluids. At the top part of the core,rom 90 m to 220 m, a large amount of the samples have positive56Fe values averaging +0.33‰. Between ca. 220 m to 360 m the56Fe values have lower values averaging −0.21‰. In the bottomart of the core, from 360 m to 450 m, a change to more positivealues is seen by �56Fe values averaging +0.26‰.
. Discussion
.1. Post-depositional history
.1.1. Diagenetic mineral paragenesis and burial metamorphismDuring diagenesis and burial of the precursor BIF sediment,
he main factors controlling the flow of metasomatic fluids areependent on various conditions within the sedimentary basin,uch as a high water-to-rock ratio, the permeability, the temper-ture gradient, and the lithostatic pressure differential, amongstthers (e.g., Smith et al., 1982; Bau, 1993). Most primary hematiteas been replaced by dense magnetite bands and secondarily by
arger disseminated magnetite grains. The latter can partly obscurehe original fine-scale magnetite bedding as seen in Fig. 4H. Theeneration of magnetite in BIF could have occurred either throughissimilatory Fe(III) reduction in which primary ferric oxyhydrox-
des were reduced at the expense of organic carbon oxidationuring diagenesis (e.g., Konhauser et al., 2005; Li et al., 2011) or at
later stage during metamorphism, for instance through the reac-ion between hematite and a ferrous iron phase such as sideriteMiyano, 1987; Li et al., 2013).
Evidence for any siderite or primary hematite has not been
bserved in this work. Since siderite is considered to be of depo-itional or early diagenetic origin (e.g., Ayres, 1972), we infer thaturing progressive burial siderite reacted with hematite to formhe dominant magnetite phase.
The timing of the ankerite–ferroan dolomite crystals seams to beof a later stage than magnetite formation. By engulfing both earlychert and very late stage burial-to-low metamorphic stilpnomelane(Fig. 6H), these euhedral rhombic carbonates were formed very latein the post-depositional story.
Riebeckite in BIF is believed to be of pre- to syn-metamorphicorigin that either completely or partially replaced chert andchert–magnetite bands (e.g., Beukes, 1973), while the crocodoliteformation is ascribed to later regional folding generating overpres-sured zones during vertical extension (e.g., Krapez et al., 2003).Timing of riebeckite growth in the Joffre BIF was before the for-mation of the chert nodules (Fig. 4C and D), indicating an earlierformation stage than the main compaction in the basin. However,mobilization of some riebeckite after the main compaction hasresulted in thin veins as seen in Fig. 4A. The abundant riebeckite islikely related to the migration of alkali-bearing solutions with highNa+ activity (see also Miyano and Klein, 1983). Na is uncorrelatedwith immobile elements (e.g., Al2O3, TiO2, Nb, La, Zr, Hf and Th)found in the fine-grained ash and stilpnomelane. This is a result ofNa being hosted within riebeckite and crocidolite only. The averageNa2O concentration is 1.1 wt.% and shows that the sodium contentin the Joffre BIF lies well above other Hamersley BIF, such as theDales Gorge and the upper and lower parts of the Marra Mamba(Fig. 13). In fact, K and Ba are moderately correlated to the afore-mentioned immobile elements, as well as with various trace metals,such as V, Cr, Ni and Co, suggesting a minimum degree of K- andBa-mobilization during burial metamorphism. This indicates thateven with the activity of Na-bearing fluids, the degree of alterationof the primary elements has been minor. Interestingly, the highriebeckite content in Joffre BIF is not seen reflected in a lower silicacontent (see Fig. 14 and Section 6.1.2), which would be expected ifchert was being replaced by riebeckite. Alternatively, the formationof laminated microbands of dense to fibrous riebeckite restricted tochert laminae suggest that the precursor of the chert was magadi-ite (NaSi7O13(OH)3), a sodium rich silica gel (e.g., Eugster and Chou,1973; Drever, 1974; Miyano and Klein, 1983; Morris, 1993). Thus,it cannot be excluded that Na could have been part of the originallyprecipitated components of the BIF.
6.1.2. Supergene enrichmentA second means by which the BIF sediment can be altered is
via supergene weathering. For instance, the downward flow ofmeteoric oxidative fluids would cause oxidation of any reduced
26
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Table 3Selected trace elements and relevant REY anomalies with core depth in the Joffre BIF. Note the three different lithologies.
Depth (m) Element Al (wt.%) Fe Mn Na P K Ti (ppm) V Cr Ni Zn Ge Rb Sr Y Zr Nb Mo BaDetection limit 0.2 (ppm) 3.7 0.030 0.5 5 6 0.09 0.05 0.05 0.06 0.08 0.02 0.04 0.03 0.02 0.09 0.04 0.02 0.03
a Calcareous mudrock.b Stilpnomelane-rich tuffaceous mudrock.c Stilpnomelane mudrock.
Fig. 13. Average major element plot for some of the least altered BIF from theHamersley Group. Wittenoom BIF from Webb et al. (2003); Marra Mamba BIF fromKt
pnmpt2pHaoBDcBBI
volcanic sources, which are represented in the three Joffre sam-
lein and Gole (1981). Notice the high amount of Na2O in Joffre BIF compared withhe other BIFs.
hases in the BIF sediments, resulting in phase changes of mag-etite into high-grade hematite. Depending on pH, those fluidsay also have led to the dissolution of carbonate minerals (low
H) or silicate minerals (high pH). Low pH solutions would causehe loss of MgO, CaO, and perhaps even Al2O3 (e.g., Webb et al.,003). This, however, is not evident in the Joffre BIF when com-ared to the other associated unaltered BIF (as shown in Fig. 13).igh pH solutions would dissolve silica resulting in chert depletionnd the concomitant formation of martite and high-grade hematitere formation as evident in the Mt. Whaleback and Mt. Tom PriceIF from the Hamersley Group (Ewers and Morris, 1981; Taylor andalstra, 2001; Webb et al., 2003). These high-grade-iron and low-hert-content iron formations are not features seen in the Joffre
IF chemistry. In fact, the Joffre BIF resembles the more unalteredIF from the Marra Mamba and Wittenoom formations (Fig. 14).
ndeed, as seen in Fig. 14, the Joffre BIF has SiO2 and Fe2O3 values
within the range of expected values for relatively non-enriched,Hamersley style BIF. With that stated, late-diagenetic magnetite isobserved locally in the Joffre core to have been partly oxidized topost-metamorphic fine-grained hematite (Fig. 4H). Furthermore,very fine-grained hematite–goethite grains (possible a variety ofmartite) are developed sporadically throughout (Fig. 4I).
6.2. Seawater chemistry
6.2.1. The REE budgetIt is generally accepted that REEs measured in Archaean and
Paleoproterozoic BIF can potentially mimic the REE composition inthe contemporaneous ocean water at the time of precipitation (e.g.,Dymek and Klein, 1988; Bau and Dulski, 1996; Bolhar et al., 2004).If true, it implies that (1) all the REEs were dissolved in seawa-ter prior to precipitation, and (2) post-depositional metamorphismdid not remove or add any of the elements. The REE patterns allhave the characteristic fractionated HREE to LREE enriched patterns(Fig. 10), and this suggests an overall minor contribution from ter-rigenous sources on the REE budget. The �REE vs. the degree ofcrustal contamination represented by the (Pr/Yb)SN ratios is plot-ted in Fig. 15A. The diagram reflects a higher concentrations ofREEs in the seawater and hence, an increase in the REE contribu-tion to the Joffre BIF basin relative to the underlying Dales GorgeBIF (see Pecoits et al., 2009). A large portion of the samples fromthe ca. 2.5 Ga Kuruman BIF (South Africa) also contain higher �REEthan the Dales Gorge BIF, but still less than a large portion of theJoffre BIF samples (Fig. 15A). Furthermore, the Marra Mamba BIF(not shown here) contains lower overall REE concentrations thanthe Joffre BIF (see Alibert and McCulloch, 1993). Whether this isdue to a higher input of submarine hydrothermal fluids and/orvolcanogenic input to the basin is unknown. However, the REEsystematics show that while all samples exhibit the characteris-tic fractionated shale-normalized ((Pr/Yb)SN < 1) seawater pattern,the portion of the Joffre BIF samples that have higher amount oftotal REEs (Fig. 15A) appear to be related to a weak increase in Aland LREEs (Fig. 15B). This, in turn, may be controlled by the same
ples with intermixed volcanogenic detritus (Fig. 15B). This revealsthat small proportions of pelagic ash particles may have had animpact on the overall REE signature and the total REE content of the
R. Haugaard et al. / Precambrian Research 273 (2016) 12–37 29
F e claso al andH ).
suitc
FJ(sD
ig. 14. Bar charts illustrating the same BIFs as in Fig. 6 but compared with one of thutcome during supergene enrichment. This is not seen for the other BIFs in gener2O in some of the BIFs. Data from Mt. Tom Price BIF from Taylor and Dalstra (2001
eawater. Petrographically, evidence of stilpnomelane microgran-
les associated with both the shard bearing ash bed, as well as
ntermixed with some of the silica- and iron-oxide microbands inhe BIF (Table 1), may explain the weakly higher �LREE and Alontent displayed in Fig. 15B.
ig. 15. (A) The �REE as a function of (Pr/Yb)SN for both the Joffre BIF, the Dales Gorge Boffre BIF samples have high input of REE and are weakly elevated in (Pr/Yb)SN values. Tstilpnomelane-rich tuffaceous mudrock, stilpnomelane mudrock and calcareous mudroamples have an increase in Al and LREE likely as a consequence of ash particles mixed wulski (1996).
sic altered BIF from Mt. Tom Price. A gain of iron and a loss of silica will be a natural for the Joffre BIF in particularly. Note the “sum-to-100” problem due to CO2 and
Many Precambrian chemical sediments are also evaluated on
the basis of their shale-normalized LaSN- and CeSN-anomalies. Thecause for the anomalous behavior of La reflects enhanced stabilityof La in solution and, accordingly, may be related to the absenceof inner 4f electrons (e.g., De Baar et al., 1985). CeSN anomalies
IF and the same style 2.5 Ga old Kuruman BIF (South Africa). A large portion of thehe field of intermixed volcanogenic detritus is defined by the three Joffre samplesck). (B) Shows the similar trend as in (A) but here with Al vs. �LREE. The sameith the BIF. Dales Gorge BIF from Pecoits et al. (2009); Kuruman BIF from Bau and
30 R. Haugaard et al. / Precambrian Research 273 (2016) 12–37
F Most
p 96). (Bd the b
rtslaCharFwCnsbi
6
P+tSiuaMieBct
ig. 16. (A) Graph illustrating shale-normalized depletion/enrichment of La and Ce.ositive La anomaly and no Ce anomaly (diagram modified from Bau and Dulski, 19uring precipitation of Joffre BIF. Note the steady increase of the EuSN anomaly from
eflect the redox state of the water column from which the par-icles precipitated. In general, oxygenated marine settings show atrong negative CeSN anomaly, whereas suboxic and anoxic watersack large negative CeSN anomalies (e.g., German et al., 1991; Byrnend Sholkovitz, 1996). Oxidation of Ce(III) to Ce(IV) greatly reducese solubility, resulting in preferential removal onto Mn–Fe oxy-ydroxides, organic matter, and clay particles. In contrast, suboxicnd anoxic waters lack significant negative CeSN anomalies due toeductive dissolution of settling Mn–Fe-rich particles. As shown inig. 16A, all the samples plot in the two fields of no Ce anomaly,hich likely reflects oxygen levels too low to oxidize Ce(III) toe(IV), with the concomitant scavenging of Ce(IV) and a resultingegative CeSN anomaly in the BIF. In contrast, a large fraction of theamples show a positive LaSN anomaly, which suggests that La haseen stabilized and weakly fractionated relative to the other LREE
n the seawater column prior to precipitation.
.2.2. The hydrothermal inputBy examining Sm–Nd isotopes in the Joffre BIF, Jacobsen and
imentel-Klose (1988) obtained an average depleted εNd value of2.1 (n = 4), and, therefore, linked a large portion of the REEs inhe Joffre BIF to submarine hydrothermal alteration of the seafloor.imilarly, Alibert and McCulloch (1993) reported a gradual changen the εNd values from the lower Marra Mamba BIF with εNd val-es of −0.6 to more depleted εNd values of +1 for the Dales Gorgend Joffre BIFs (see Fig. 1B for stratigraphic position). Alibert andcCulloch (1993) additionally made Nd mass balances suggest-
ng that mid-ocean hydrothermal fluids mixed with seawater could
xplain around 50% of the sourced Nd (and hence the REEs) in JoffreIF. In this regard, the hydrothermal evolution of the Joffre BIF basinan be discerned by the variation of the EuSN anomaly (Eu/Eu*SN)hroughout the BIF sedimentary succession.
of the BIF samples plot within the field of (Ce/Ce*)SN < 1 and (Pr/Pr*)SN ∼ 1, meaning) Shows the hydrothermal evolution (represented by the EuSN anomaly (Eu/Eu*)SN)ottom to the top of the core with a peak of ∼2.1 around 250 m depth.
The (Eu/Eu*)SN in modern seawater is identical to thePost-Archaean Average Shale, whereas in modern submarinehydrothermal solutions Eu is enriched with (Eu/Eu*)SN values > 1(e.g., Danielson et al., 1992; Kato et al., 1998). Fig. 16B shows theevolution of (Eu/Eu*)SN with depth in the Joffre BIF. The graphshows that the mixture of hydrothermal fluids with seawaterresulted in Eu anomaly well above 1 throughout the entire coredepth. However, the input of hydrothermal fluids affected the BIFto varying degrees. An increase in dissolved Eu2+ resulted in theincrease of the (Eu/Eu*)SN, that peaked between 100 m and 155 mof the BIF sequence (Fig. 15B). This section reflects a larger hydro-thermal input, with a maximum EuSN anomaly of ∼2.1.
6.3. Volcanic activity
6.3.1. Ash-fall tuff and microgranulesRelict shards within the tuff bed of the Joffre BIF (see
Figs. 5C and 6A) indicate a volcanic provenance most likely fromexplosive magmatic eruptions (e.g., Fisher and Schmincke, 1984).Well-preserved wavy lamina and lamina sets of microgranulesdeposited on top of the graded tuff bed have been found in the upperpart of the Joffre BIF. The granular texture and the well-preservedshards can only exist if early formation of diagenetic silica pre-vented later compaction of the lamina. Thin section and microprobeanalysis of the microgranules and the tuff matrix clearly suggest astilpnomelane phase. Thus, the stilpnomelane microgranules arelinked to the ash bed underneath, and as such, the microgran-ules most likely represent ash material that has been reworked by
unknown processes within the water column or at the seafloor,e.g., by density currents or contourites (e.g., Krapez et al., 2003;Rasmussen et al., 2013). Direct evidence for the latter processes arerare due to burial and metamorphic overprinting but the preserved
R. Haugaard et al. / Precambrian Research 273 (2016) 12–37 31
Fig. 17. TiO2–Zr, log–log plot showing the Joffre BIF and the relation of these elements to other voluminous lithologies that may have had an influence on the seawaterchemistry. In conjunction with Joffre tuffaceous mudrock (from Pickard, 2003), the stilpnomelane-rich tuffaceous mudrock, the calcareous mudrock and the stilpnomelanemudrock, the Joffre BIF produce a (power) regression that is linked to a rhyolite-only-source represented by the Woongarra rhyolites. In contrast, a more bimodal TiO2–Zrcontribution, best represented by the Dales Gorge S-bands, which clearly have affinities to average continental crust. The Joffre tuffaceous mudrock and Woongarra rhyolitesfrom Barley et al. (1997) and Pickard (2003); original linear regression line of Dales Gorge S-bands (Zr = 244TiO2(wt.%) – 2.1) from Ewers and Morris (1981); field of submarinekomatiite (3.2 Ga Ruth Well Fm.), submarine basalt (2.72 Ga Kylena basalt), subaerial basalt on cont. crust (2.69 Ga Medina basalt) from Arndt et al. (2001); field of pillowb . (199( ental
M
cCsegrTsa
npocdageccttegcWgtfBugpt
6
os
asalt (upper Fortescue Group), flood basalt (2.78 Ga Mt. Roe Fm.) from Nelson et al2003); Dales Gorge tuff (S13 and S15) from Pickard (2003); average upper contin
cLennan (1981).
urrent generated structure developed in the oxide-BIF (Fig. 3B and) is most likely a result of density currents on the seafloor duringedimentation rather than a post-depositional feature (e.g., Krapezt al., 2003). Those authors linked the origin of locally preservedranules to hydrothermal mud that was transported, deposited ande-sedimented on submarine volcanic flanks by density currents.hey suggested that the granular muds may have been the precur-or for the BIF, rather than direct precipitates of amorphous silica-nd iron oxides from seawater.
Recently, an interesting model on the formation of the stilp-omelane microgranules within the Dales Gorge BIF has beenroposed by Rasmussen et al. (2013) who suggested flocculationf an Fe(III)-rich and Al-poor hydrous silicate either in the waterolumn or on the seabed, which was subsequently reworked byensity currents to form lamina sets with a basal granular bed and
more granular diluted-amorphous mud lamina on top. From aeochemical perspective, the high Fe(III)-oxide content and thenriched HREE (Y, Yb, Lu) relative to both primitive mantle andontinental crustal values (Fig. 11) strongly support that the pre-ursor to the stilpnomelane mudrock and the stilpnomelane-richuffaceous mudrock in the Joffre BIF reacted with the seawater prioro settling, a finding in agreement with the model by Rasmussent al. (2013). In addition to iron and silica, the stilpnomelane micro-ranules in the Joffre BIF also contain various amounts of otheromponents, such as Al2O3, TiO2, K2O, Zr, Nb, Th and trace metals.
e thus propose that the precursor to the stilpnomelane micro-ranules in Joffre BIF may have been very fine, reactive ash particleshat chemically interacted with seawater, thereby stripping ironrom seawater prior to their deposition at the seafloor. In the JoffreIF, stilpnomelane also exists with other textures than microgran-les. For example, the very thin microbed containing stilpnomelaneroundmass and flakes alongside quartz fragments and varioushyllosilicates (Fig. 5I) may represent settling of volcanic detritushrough the water column.
.3.2. The source of the detritusSeawater composition is inevitably linked to the composition
f surrounding continental crust and to changes in the degree ofubmarine hydrothermal alteration of the seafloor. As a result of
2); field of dolerite and tuff (Weeli Wolli Fm.) from Barley et al. (1997) and Pickardcrust from Taylor and McLennan (2009); average Fortescue shale from Taylor and
a change in the overall heat regime and style of plate tectonics,major geochemical changes took place at the Archaean-Proterozoicboundary (e.g., Taylor and McLennan, 1981, 1985, 2009). For exam-ple, the upper continental crust and their sedimentary derivativesbecame enriched in Large Ion Lithophile Elements (LILE) (e.g., K,Rb), LREE, Zr, Th and Hf, while they became relatively depletedin various transition metals (e.g., Cr, Ni), Fe and Mg (Taylor andMcLennan, 1985; Condie, 1993). Stilpnomelane hosts the dominantpart of the detritus in the Joffre BIF. The occurrence of stilpnome-lane is likely the replacement product of greenalite (a ferrous–ferricphyllosilicate of the kaolinite–serpentine group), and suggested tobe a key indicator of mafic volcanogenic derived material (Winkler,1979; Pickard et al., 2004). Unfortunately, the nature of the volcanicprecursor for the Joffre BIF, and other BIF in general, is difficult todeduce since the ash has been deposited in a basin influenced byFe(II)-rich seawater which subsequently interacted with the highlyreactive ash particles (see Section 6.3.1). However, the primitivemantle normalized spider diagram in Fig. 11 shows that apart fromoverall lower abundances, a large part of the insoluble elementsin the Joffre BIF follow the same pattern as that for the two associ-ated volcanic deposits (stilpnomelane-rich tuffaceous mudrock andstilpnomelane mudrock) and on a broader scale, the upper conti-nental crust. Therefore, it is important to look in detail at thoseelements to be able to characterize the provenance of the detritus.
The low concentrations of insoluble elements (e.g., Nb, Ti, Th,Zr, Hf) can be used as a fingerprint for what lithologies wereexposed to weathering on the continents proximal to the BIF depo-sitional basin. In Fig. 17 (TiO2 vs. Zr), two insoluble elements thatremain adsorbed to the pelagic detritus are plotted for a range oflocal lithologies that all potentially could have had an influenceon the seawater chemistry. A clear dominance from rhyolite typevolcanics (here represented by the Woongarra rhyolites) is seenfor the Joffre BIF. The regression line is made on data from theJoffre sequence, including the BIF, stilpnomelane-rich tuffaceousmudrock, stilpnomelane mudrock, calcareous mudrock (this study)
and 4 tuffaceous mudrock samples from Pickard (2003). All of thosesamples plot intermediate between the igneous lithologies and theJoffre BIF, indicating a mixture of volcanogenic material and chemi-cal sediment. Similar to the Dales Gorge S-bands (dashed regression
3 brian Research 273 (2016) 12–37
lmTfscWt(it
mmm(1tpcswae
6
az(tvteaff(G�iorfsp2nptadsa(co
gowbbbh
Fig. 18. �56Fe histogram of Joffre BIF compared with that of Dales Gorge BIF. The Jof-fre BIF is 42 analyses of whole rock while the Dales Gorge is 40 analyses on magnetitealone. Despite that fact, the Joffre BIF shows a more skewed distribution toward more
2 R. Haugaard et al. / Precam
ine), the massive stilpnomelane mudrock seems to be related to aore intermediate (average upper continental crust) than basaltic
iO2–Zr composition. Interesting, none of the TiO2 and Zr in the Jof-re BIF is sourced from ultramafic sources or any of the mafic rockuites presented in Fig. 17. This could also explain the low con-entrations of trace metals in the Joffre basin (etc., Ni and Cr). The
eeli Wolli tuff, which represents volcanic activity after deposi-ion of Joffre BIF, but before the occurrence of Woongarra rhyolitessee Fig. 1B), also represents a more felsic TiO2–Zr composition andndicates the continuing dominance of felsic volcanic activity onhe later stages of Hamersley deposition.
In this study, no direct evidence of any shelf-derived epiclasticaterial to the Joffre BIF basin has been found throughout the 355eter of core section. This is in agreement with other workers whoade similar observations for the Hamersley Group BIF as a whole
e.g., Ewers and Morris, 1981; Morris and Horwitz, 1983; Morris,993). Instead, it seems likely that most of the continental inputo the Joffre basin during BIF precipitation was through volcanicathways in the form of pyroclastic input and not from an erosiveontinent. If that is the case, then the lack of terrigenous clasticsuggests that the Joffre BIF sequence was either formed in a deep-ater setting or in shallower water but within a continental starved
nd transgressed shelf margin with only minor contribution to thelemental budget from an erosive continental crust.
.4. Iron isotopes and Fe(II) oxidation
The stratigraphic variations in �56Fe in the Joffre BIF (Table 4nd Fig. 12) suggest that the Fe-cycle in the early Paleaoprotero-oc seawater was affected by different fractionation mechanismse.g., Steinhofel et al., 2010). For igneous rocks, the 56Fe/54Fe iso-ope ratio (expressed as �56Fe) is generally unfractionated withalues around 0 ± 0.15‰, while for mid-ocean hydrothermal fluidshe �56Fe is slightly negative, ranging from −0.3 to −0.6‰ (Sharmat al., 2001; Johnson et al., 2003). However, at low temperatures,nd influenced by redox processes, it is possible to significantlyractionate the heavy and lighter Fe-isotopes, yielding a range of dif-erent �56Fe values as evident in the various BIF facies analyzed hereFig. 12). The probability histogram for Joffre BIF (Fig. 18A) and Dalesorge BIF (Fig. 18B) shows the former displaying higher positive56Fe values than the Dales Gorge BIF. This is surprising consider-
ng that the measured �56Fe (n = 40) for the Dales Gorge BIF wasn magnetite only, while for the Joffre BIF the measured �56Fe rep-esents bulk analyses (n = 42). In contrast to Fe-carbonates, whichrequently exhibits negative �56Fe values (Johnson et al., 2003),econdary hematite and magnetite are often the only minerals dis-laying positive �56Fe values (Johnson et al., 2003; Rouxel et al.,005). This implies that the main control on the bulk �56Fe sig-ature measured in Joffre BIF is likely from a ferric oxyhydroxiderecursor to magnetite (e.g., ferrihydrite) since magnetite is by farhe most dominating mineral. Taking all the ankerite into account,
simple mass balance consideration also suggests a more skewedistribution toward the positive excursion if magnetite was mea-ured only. Bulk BIF samples with overall positive �56Fe values canlso be found in many other Archaean and Palaeoproterozoic BIFsPlanavsky et al., 2012) and show that these were a sink for isotopi-ally heavy Fe. The role of secondary riebeckite on the �56Fe budgetf BIF is currently unknown.
This Fe-isotope fractionation can be attributed to either inor-anic or organic processes where hydrothermal controlled Fe(II) isxidized to Fe(III) resulting in precipitation of ferric oxyhydroxidesith positive �56Fe values (Johnson et al., 2003). Experimental work
y Croal et al. (2004) showed that the above fractionation is possi-le by the existence of anaerobic photoautotrophic Fe(II)-oxidizingacteria that use Fe2+ as an electron donor and precipitate ferri-ydrite enriched in �56Fe by up to +1.5‰. However, since direct
positive �56Fe values, which indicate higher rate of iron oxidation in the contempo-raneous seawater. See text for further interpretation. Values for Dales Gorge BIF arefrom Rouxel et al. (2005).
oxidation of Fe2+ by free O2 would generate similar positive �56Fepatterns, this “fingerprint” of anaerobic biogenic fractionation canonly be predicted if there is independent evidence for an anoxy-genic ocean-atmosphere (Croal et al., 2004). We suggest that duringdeposition of the Joffre BIF submarine hydrothermal injected fluids,together with pyroclastic detritus, played a greater role on seawa-ter chemistry compared to continental derived epiclastic material.This is not unexpected given that the Brockman Iron Formation as awhole has been linked to major plume breakouts (e.g., Barley et al.,1997) that developed in association with the emergence of newcontinental crust during supercontinent assembly (e.g., Condie,2005). A plume breakout would cause shallower mid-ocean ridges(e.g., Ernst et al., 2004), which in turn, would lead to transgres-sive events that submerged the shallow shelf in waters directlyinfluenced by submarine mantle degassing and hydrothermal alter-ation of the oceanic crust (seen through the (Eu/Eu*)SN evolution inFig. 16B). A combination of high Fe2+, abundant reduced gases fromthe mantle (e.g., CH4, H2, H2S), and alteration of new oceanic crustwould have acted as O2 sinks and likely promoted marine anoxiain the Joffre basin. This scenario could explain that the fraction-ated positive Fe-isotopes were a result of anaerobic photosyntheticFe(II)-oxidizing bacteria consuming lighter Fe-isotopes faster thanheavier. For that to happen, a stratified ocean having a large Fe(II)-
rich deep water pool and a shallower upper water pool whereFe(II) oxidation and enrichment of Fe isotopes seems plausible. Itis important to note that during diagenesis and pore water inter-action, it has been shown by Busigny et al. (2014) that loss of light
rian R
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e from the pore waters is unlikely to generate positive Fe isotopealues in the sediment.
One interesting aspect is the �56Fe value of +0.59‰ for theIF dominated by volcanic ash-fall. This value is significantlyigher than igneous rock values (Fig. 12), reflecting contempora-eous Fe(II) oxidation during volcanic activity. A portion of thesh formed microgranules consisting of stilpnomelane, and asuch, the iron in the granules must have been fractionated in theater column before deposition and reworking. The depositional
ink between ash-fall, the stilpnomelane microgranules, and evi-ence of oxidation fractionating the Fe is interesting in that pulsesf fine-grained volcanic ash may have promoted higher bacte-ia production throughout the upper water column (photic zone),peeding up Fe(II) oxidation and precipitation of the ferric oxyhy-roxides that subsequently mixed with the ash-particles duringeposition. Concomitantly, increased biomass may have settled tohe seafloor, allowing for the needed reductants to dissimilatorye(III) reduction and magnetite formation to occur during diagene-is and metamorphism (Konhauser et al., 2005; Kolo et al., 2009; Lit al., 2011, 2013). In contrast, the massive stilpnomelane band inhe bottom of the core, having affinities to a more basaltic composi-ion, shows a high negative �56Fe value of −0.74, possibly mirroringiagenetic pyrite with high negative �56Fe values. Today, the fer-ilizing effect of volcanic ash and the importance of volcanism forhe marine biogeochemical cycle has been described (for a reviewee Duggen et al., 2010).
.5. The GOE: sulfur and trace metals
The late Archaean to early Palaeoproterozoic (∼2.7–2.4 Ga)arks one of the most important periods in Earth history with a
umber of major interlinked environmental events. These include,mongst others, (1) major episodes of mantle plume activity andontinental assembly (Eriksson et al., 2001; Condie, 2004; Barleyt al., 2005), (2) a peak in BIF deposition (Isley and Abbott, 1999;ekker et al., 2010); (3) development of an oxic layer in the oceansReinhard et al., 2009; Kendall et al., 2010), (4) evolution of aerobi-ally respiring organisms (Eigenbrode and Freeman, 2006; Godfreynd Falkowski, 2009; Konhauser et al., 2011), and ultimately (5)he oxygenation of the atmosphere, the so-called Great Oxidationvent (GOE) (Holland, 2002; Bekker et al., 2004).
The GOE represents a transition in time from an atmospheressentially devoid of free oxygen (O2 < 10−5 times the presenttmospheric level, PAL) to oxygen concentrations higher than 10−5
AL (Pavlov and Kasting, 2002; Kopp et al., 2005; Buick, 2008). TheOE is best defined by a loss of mass-independent sulfur isotope
ractionations (S-MIF) in sedimentary rocks (Farquhar et al., 2001),ith data from various locations worldwide showing that S-MIF
ontinued to persist in the rock record until sometime between.45 and 2.32 Ga (Guo et al., 2009; Canfield and Farquhar, 2009).his rise of free atmospheric oxygen facilitated the onset of oxida-ive continental weathering reactions and increased the fluxes ofulfate and redox-sensitive trace elements to the oceans (Canfield,005; Anbar et al., 2007; Frei et al., 2009; Reinhard et al., 2009;onhauser et al., 2011). Collectively, these studies show that theOE was a protracted process that took hundreds of millions ofears (Lyons et al., 2014).
The picture of how the GOE emerged has recently gained morelarity through a compilation of Cr concentrations in BIF throughime, which showed a significant enrichment beginning at 2.45 Gan the Weeli Wolli Formation (Konhauser et al., 2011). Given thensolubility of Cr minerals, its mobilization and incorporation into
IF indicates enhanced chemical weathering at that time, most
ikely associated with the evolution of aerobic continental pyritexidation. Interestingly, evidence for a ‘whiff’ of oxygen was pre-iously noted for the underlying 2.5 Ga Mt. McRae shale (Fig. 1B)
esearch 273 (2016) 12–37 33
where molybdenum (Mo) concentrations increase from <5 ppm(near the crustal level) to 40 ppm, and then decrease back to<10 ppm (Anbar et al., 2007). These patterns were interpretedas reflecting a transient oxygenation event in the atmosphere,although this view has been complicated by the suggestion thatfeatures indicative of oxidative weathering in the pre-GOE rockrecord may instead stem from localized O2 production in asso-ciation with biological soil crusts and freshwater microbial matscovering riverbed, lacustrine, and estuarine sediments (Lalondeand Konhauser, 2015).
The rise in oxygen between 2.4 and 2.3 Ga permitted theincreased delivery of sulfate to the oceans via enhanced oxida-tive sulfide weathering on land. Once in seawater, the sulfate hadtwo major sinks, (1) iron sulfide precipitation as a consequenceof bacterial sulfate reduction in the water column, or (2) evapor-itic precipitation of gypsum. In the first instance, Canfield (1993)proposed that increased levels of sulfide in the oceans effectivelytitrated out any remaining Fe2+ in seawater, leading to the end ofBIF deposition. Evidence in support of higher sulfide productioncomes from increasing fractionation between sulfur isotopes; val-ues for 34S/32S (�34S) are centered on mantle values (0‰) priorto around 2.45 Ga but then increase to around 25‰ after 2.45 Ga(Canfield and Farquhar, 2009). In the second instance, primary sul-fate evaporites are rarely reported before 2.45 Ga (Schröder et al.,2008), confirming insufficient dissolved sulfate availability beforethat time. In the Joffre BIF, neither petrographic nor geochemi-cal data support the presence of major sulfide or sulfate mineralphases; pyrite has only been documented as a small mineral con-stituent in the associated stilpnomelane-rich tuffaceous mudrock.
The overall low abundances of trace metals in the Joffre BIF rela-tive to upper continental crust values (Fig. 11) reflects either (1) lowsolubility related to a low degree of oxidative continental weather-ing of exposed mineral sulfides; (2) that the adjacent land massesto the Joffre basin were of a composition that did not amply supplythose metals – for instance, a lack of ultramafic–mafic sources; or(3) continental weathering had only a subordinate control on theseawater chemistry within the Joffre basin.
In the case of Mo and Cr, their exceedingly low concentrationsin the Joffre BIF (Fig. 9M and O) suggests that their parent miner-als, such as molybdenite (MoS2) and chromite ([FeCr]2O4), werenot significantly dissolved. As demonstrated by Anbar et al. (2007),increased concentration of Mo in the 2.5 Ga Mount McRae shale ofWestern Australia reflected a ‘whiff’ of oxygen before the GOE; theMo was sourced from oxidative weathering of Mo-bearing sulfidesin crustal rocks. By contrast, low levels of Ni (Fig. 9N) in the JoffreBIF is in full agreement with the findings of Konhauser et al. (2009)who suggested that the cooling of the upper mantle led to decreasederuption of komatiite lavas (with high Ni content), reduced supplyof Ni to seawater, and thus less incorporation into marine chemicalprecipitates, such as BIF after 2.7 Ga. Despite the predicted low lev-els of Mo, Cr and Ni in seawater at 2.46 Ga, those metals are mainlycontrolled by the stilpnomelane suggesting that those elementswere, to some extent, controlled by the volcanic input best rep-resented by the tuff material in the stilpnomelane-rich tuffaceousmudrock.
7. The palaeoenvironment
In the epiclastic-starved basin, seawater composition duringJoffre BIF deposition was controlled by two supply systems: (1)convective upwelling of deep, hydrothermally enriched seawater
to the outer continental platform, and (2) volcanic pyroclastic detri-tus from distal volcanic centers (see Fig. 19 for a simplified model).During convective upwelling, Fe(II) rich deeper waters floodedthe platform. The upwelling was likely discontinous because this
34 R. Haugaard et al. / Precambrian Research 273 (2016) 12–37
Fig. 19. A simplified palaeoenvironmental model for the formation of the Joffre BIF. (A) Represent the situation during formation of ferric-hydroxide precipitation from thephotic zone. This scenario has nutrient-rich upwelling onto the platform from deeper waters influenced by hydrothermal activity. Together with diluted fine-grained ashp oxidati ) forma ration
ptMraAlmeltTtFstta(e
pdtl
articles this increase the nutrient level in the surface water speeding up the Fe(III)
n the sediment. (B) Show the general situation during silica and carbonate (sideritend FeCO3 from the upper seawater. See text for further explanations. The concent
rocess depends on seasonal variations in the oceanic currents inhe surface water as previously suggested by Morris (1993) for the
arra Mamba BIF (Fig. 1B). Evidence that upwelling varied in time iseflected by the relatively large variations in both the EuSN anomalynd the thickness and distribution of magnetite bands in the core.longside Fe(II), the concentration of REY, including Eu(II), were
ikely to be enriched in the deeper water. If seafloor plumes and sub-arine volcanism facilitated BIF deposition (as proposed by Barley
t al., 1997; Isley and Abbott, 1999), then the deeper water was alsoikely enriched in reduced gases from the mantle (e.g., CH4, H2, H2S)hat could act as a sink for oxygen and thus promote marine anoxia.he upwelling, along with pyroclastic material, brought impor-ant nutrients to the photic zone speeding up the photoautotrophice(II) oxidation. When upwelling ceased (Fig. 19B), the precursorediment to chert then precipitated from the surface waters dueo the high concentrations of dissolved silica in the oceans at thatime (e.g., Konhauser et al., 2007). In addition, dissolved bicarbon-te or silica reacted with Fe(II) and formed fine-grained sideriteAyres, 1972; Klein and Beukes, 1989) or greenalite (Rasmussent al., 2013), respectively.
Superimposed on this internal marine dynamic system is the
yroclastic source that provided small amounts of fine-grainediluted ash particles. In the photic zone the ash intermixed withhe BIF precipitates leading to the different styles of stilpnome-ane mineralogy, that is, the two stilpnomelane-rich rock facies
ion through photoautotrophic Fe(II) oxidation and leaves a positive �56Fe signatureation. The shut off of the upwelling resulted in relative pure deposition of Si(OH)4
values of Fe(II) in the upper and deeper seawater is taken from Morris (1993).
and those disseminated within the more typical oxide facies BIF.From the above two sources a large amount, relative to the under-lying Dales Gorge BIF, of REY became incorporated into the ironand silica precipitates on the clastic-starved continental platform(Fig. 19). The small rise in total LREE and Al in about half of thesamples (Fig. 15B) could thus reflect the influence of very dilutedash particles on the REY budget of the general BIF precipitate (seealso Table 1).
In the photic zone, enhanced productivity of anaerobic pho-toautotrophic Fe(II)-oxidizing bacteria led to heavier iron isotopefractionation in the precipitated ferric-hydroxides and subse-quently in the various style of magnetite bands (Fig. 19A). Inclastic-starved environments, without input of bioessential nutri-ents through erosion of the nearby landmasses, other sources musthave controlled biological productivity. Perhaps the delivery offine-grained ash particles, coupled to upwelling of hydrothermallyenriched seawater, were those nutrient sources? In the case of theformer, experiments in seawater have shown that airborne volcanicash particles have soluble coatings containing important micronu-trients (Cu, Zn etc.) and macronutrients (P, K and NH4+) that uponcontact with seawater will be released within minutes. This implies
that most of the released nutrients are accessible in the photic zone(see Duggen et al., 2010 and references therein).
This source is hard to quantify since distal tephra input can bevery dispersed in the atmosphere depending on the size of the
rian R
edbtmasafbE
8
rtriasrntp(cgd
esvtbmntucwTfmlitvactt((biatvsbf
ra
R. Haugaard et al. / Precamb
ruption and numerous meteorological factors. However, seafloorrilling carried out to quantify the tephra input to the Pacific Oceanasin, the largest and oldest (174 Ma) ocean basin on Earth, showshat about 25 vol.% of the existing oceanic sediments are tephra
aterial, half of which comes from subaerial arc volcanism (Straubnd Schmincke, 1998). A major portion of this material is not neces-ary deposited as distinct ash layers but instead occurs as dispersedsh particles in the marine sediments (Duggen et al., 2010). There-ore, the role of volcanic ash acting as a fertilizing agent for oceaniciota may have been an underestimated factor on the Precambrianarth.
. Summary
The ca. 2.45 Ga old Joffre BIF can be subdivided into two majorock types (oxide-facies and silicate–carbonate–oxide facies) andhree minor rock types (stilpnomelane mudrock, stilpnomelane-ich tuffaceous mudrock and calcareous mudrock). The oxide-faciess dominated by chert, magnetite and hematite, with lessermount of riebeckite, carbonate, crocidolite and stilpnomelane. Theilicate–carbonate–oxide facies is dominated by chert, magnetite,iebeckite and ankerite, with minor hematite, crocidolite and stilp-omelane. A clear lithological boundary between these two faciesypes is not feasible; rather it is gradational. Although the dominantarts of the sedimentary structures found are of secondary originchert nodules, flame structures, etc.), in rare cases the oxide-faciesontains primary syn-depositional features presumably of current-enerated origin. This illustrates the existence of weak bottom (etc.,ensity) currents within the BIF basin.
There is no evidence for epiclastic material, sourced from anrosive continent, within the Joffre BIF. Instead, petrographicaltudies show that the three minor rock types all have detritus ofolcanogenic origin. Furthermore, the chemostratigraphy showshat soluble elements, such as K2O and Ba, co-vary with insolu-le elements, such as Al2O3, Ti, Zr and Nb, suggesting only minorodification of the geochemistry. This relationship supports the
otion that adjacent volcanoes, delivering pyroclastic sediment tohe Joffre BIF basin, were the main source of both insoluble and sol-ble elements. Indeed, the stilpnomelane-rich tuffaceous mudrockonsists of volcanic tuff with well-preserved shards overlain byavy laminae and laminae sets of stilpnomelane microgranules.
hese granules most likely originated from re-worked volcanic ashormed either on the seafloor or in the water column. Since the
atrix within the tuff bed is of stilpnomelane composition, it isikely that the felsic ash particles reacted with ferric oxyhydrox-des in the water column during settling, thereby gaining the irono form stilpnomelane. This notion is supported by the high �56Fealue of +0.59‰ for this rock type. In fact, detailed geochemistrynd petrography shows that small proportions of pelagic ash parti-les (now in the form of stilpnomelane) have had a minor impact onhe overall REE signature. While all BIF samples exhibit the charac-eristic fractionated shale normalized (SN) seawater pattern withPr/Yb)SN < 1, a large portion of the samples have high total REEsrelative to other similar BIF) and weakly elevated, although stillelow 1, (Pr/Yb)SN ratios. These correspond to a slight increase
n Al and LREEs, which is directly linked to the volcanic sourcesdjacent to the Joffre BIF basin. The TiO2–Zr ratio of Joffre BIF andhe mudrocks indicates a felsic source related to the same style ofolcanics as the slightly younger Woongarra rhyolites. We demon-trate here that the precursor to stilpnomelane does not have toe an indicator of mafic volcanism but instead it could have been
elsic volcanic ash that interacted with Fe-rich seawater.
In addition to the volcanic contribution, the input of subma-ine hydrothermal fluids to the seawater played a significant roles a source of solutes to the BIF. This is most clearly evidenced
esearch 273 (2016) 12–37 35
by the EuSN anomaly, (Eu/Eu*)SN, which is above 1 throughoutthe entire succession, with a peak value of ∼2.1 between 100and 155 m of core depth. Within the water column, a large frac-tion of the Fe(II) sourced from the mid-ocean ridge environmentunderwent heavy isotopic fractionation where Fe(II)-oxidation andsubsequently precipitation of ferric oxyhydroxides resulted in highpositive �56Fe values ranging between +0.04‰ and +1.21‰ (aver-age +0.46‰). This process seems to have been more pronouncedrelative to the underlying Dales Gorge BIF.
Evidence for elevated abundances of sulfur and redox sensitivetrace metals (e.g., Mo, Cr) have not been found in the BIF, presum-ably implying a low degree of oxidative continental weathering.This is not surprising given the lack of epiclastic components andthe fact that the dominant control on the detritus was from felsicvolcanics. Correspondingly, it is clear that the Joffre BIF is poorlysuited as a chemical proxy for the study of atmospheric oxygenand its weathering impact on local landmasses.
Acknowledgements
We are grateful to Rio Tinto in Australia for providing accessto their core samples. A number of colleagues at the University ofAlberta, including Mark Labbe, Martin Von Dollen, Ilona Ranger,Tom Chacko, Andrew Locock and Igor Jakab, are highly appreci-ated for their help, as well as Birger Rasmussen for instructivediscussions concerning BIF petrology. The Natural Sciences andEngineering Research Council of Canada (NSERC) supported thiswork.
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