Black smokers and density currents: A uniformitarian model for the genesis of banded iron-formations

Ore Geology Reviews 32 (2007) 381 – 411 www.elsevier.com/locate/oregeorev

Black smokers and density currents: A uniformitarian model for the genesis of banded iron-formations Desmond F. Lascelles ⁎ Centre for Exploration Targeting, School of Earth and Geographical Sciences, University of Western Australia, Crawley, W.A., 6009 Australia Received 3 July 2004; accepted 21 November 2006 Available online 16 December 2006

Abstract Banded iron-formations (BIFs) were comparatively abundant and widespread marine sedimentary rocks in the Archean and Lower Proterozoic eras, but thereafter they appear to be restricted to the Neoproterozoic and Paleozoic eras, although there are indications of similar rocks forming at present. BIFs are important as the major source of iron ore for industry and have also been used to support hypotheses regarding the evolution of life, oceans, and the atmosphere in the Archean and Proterozoic. They apparently formed in deep water and consisted of a semi-regular alternation of quartz (chert) and iron-rich minerals with little or no terrigenous sediment and low (b 1 wt.%) alumina content, in contrast to Proterozoic to Phanerozoic oolitic ironstones (Clinton– Minette style) that formed in shelf environments, were comparatively small and rare and contain abundant iron aluminosilicates and hydroxides but little or no chert. Although ooliths may be locally abundant in granular iron-formations (GIFs) they differ from the Clinton–Minette style oolitic ironstones that formed in shelf seas during periods of slow sedimentation. A common mode of origin for marine deep-water iron-formations is proposed, in which hot fluids, consisting of marine and connate water leaching iron, silica and other elements from mafic and ultramafic rocks associated with mantle plumes or mid-oceanic ridges and active spreading centres are released into the ocean at underwater hot springs (black smokers). On contact with cold marine water, the least soluble elements are precipitated in the form of colloidal hydrous silicates (clay minerals) and hydroxides close to the hydrothermal vent. The hydrothermal fluids are high in silica and low in alumina causing the precipitation of alumina-poor iron silicates (nontronite) that dissociate into iron hydroxide and amorphous silica during diagenesis. The amorphous silica is typically entrapped by iron oxide laminae to form bands of chert. Breaches of the iron oxide laminae permitted the escape of the gelatinous amorphous silica during compaction and dewatering leaving a chert-free residue as the protore of non-hydrothermal sedimentary high-grade iron ore. The rapid deposition and abundant included water formed unstable mounds and chimneys around the vents. Slumping of the mounds caused by compaction, dewatering, gravity sliding, and seismic events produced turbidity currents forming proximal fans of GIF, but colloidal particles remained suspended in longer-lived density flows to deposit the ultra fine-grained BIF over vast areas of the ocean floor. Episodes of density current deposition were separated by intervals of slow pelagic sedimentation and silicification of the sea floor. The density currents commonly caused minor erosion of the sea floor with rip-out clasts incorporated in the new layers. The iron-rich hydrothermal fluids triggered the precipitation of dissolved ferrous iron accumulated by anoxic weathering to produce the huge deposits of BIF in the Archean and Paleoproterozoic until the rise in atmospheric oxygen stopped the accumulation of ferrous iron in the oceans leaving only the hydrothermal source for later deposits. © 2006 Elsevier B.V. All rights reserved. Keywords: Banded iron-formation; Black smokers; Turbidites; Iron ore; Precambrian environments

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1. Introduction As the third most abundant element in the Earth's crust (Mason, 1966), iron is widely distributed in sediments of all ages but typically forms less than 5% of the whole rock bulk composition. Iron-formation is defined as a stratigraphic unit with more than 15% Fe. The term ironstone is commonly used to denote Phanerozoic iron-rich rocks (James, 1966) and includes many types of iron-rich sedimentary rocks and soil formations that form in almost the whole gamut of sedimentary environments. Most post-Paleoproterozoic marine iron-formations are oolitic (Clinton–Minette style) and consist mainly of iron aluminosilicates and iron oxides without chert bands (James, 1966; Maynard, 1983), whereas ironformations with bands of chert and extremely low (N1%) alumina content are defined as banded ironformation (BIF). The origin of BIF has exercised many geologists over the past century, and numerous volcanigenic, replacement and sedimentary models in a variety of environments ranging from deep to shallow marine, paralic to terrestrial and placer to evaporitic concentration, have been proposed to explain the origin of these enigmatic rocks. Detailed analyses of the many concepts historically proposed for the origin of BIFs and their derived high-grade ore deposits are abundant in the voluminous literature of iron-formations. BIF is defined as a sedimentary rock with alternating mm- to cm-scale iron-rich and iron-poor layers, typically consisting of alternating bands of iron oxides, silicates or carbonates and chert, with a bulk chemistry containing more than 15% Fe (James, 1954). BIFs with recognizable detrital particles or ooliths are generally classed as granular iron-formation (GIF) and they may contain minor terrigenous particles and they may be interbedded with finely laminated iron-formation (Hall and Goode, 1978; Beukes, 1984). BIFs are typically extremely finegrained unless recrystallised by metamorphism and are devoid of recognizable detrital grains. Few recent BIFs are known and with minor exceptions both BIFs and cherty GIFs appear to be restricted to the Archean and Paleoproterozoic (Beukes, 1973; Klein and Beukes, 1992; Abbott and Isley, 2001; Huston and Logan, 2004). Since the Paleoproterozoic era, minor BIFs occur from the Mesoproterozoic to the Paleozoic eras (O'Rourke, 1961). Cherty BIF occurring in the Neoproterozoic (Gross, 1965; Klein and Ladeira, 2004) is classed as Rapitan type, Paleozoic BIFs as Lahn–Dill type (Quade, 1976; Kräutner, 1977), and there are some indications of similar deposits forming at present (Boström et al., 1969; Honnorez et al., 1981). The Rapitan and Lahn–Dill

laminated iron-formations typically contain abundant chert bands and appear to have formed under exceptional local environmental conditions. However, the high-grade hematite ore, most of the chert-free BIF at Mount Gibson, in Western Australia (Lascelles, 2006) and the majority of high-grade iron ore deposits in Western Australia show no evidence of the previous presence of chert bands, hydrothermal activity or major carbonate replacement (Fig. 1A–D). Attempts to find an explanation for the formation of nonhydrothermal chert-free BIFs led to the formulation of a new uniformitarian model for the origin of deep-water marine iron-formations. Positive evidence for the origin of BIFs is scant but the model is based on personal observations, published descriptions of BIF and iron ore deposits around the world and logical extrapolation from basic principles of geochemistry, sedimentology, diagenesis and weathering. Personal observations include over 20 years of extensive surface and underground mapping, diamond and percussion drilling, and petrography at numerous BIF and BIF-derived iron ore occurrences in both the Hamersley and Yilgarn provinces of Western Australia, as well as observations of many Tertiary iron ore deposits in Western Australia, and Phanerozoic iron ores in the U.K. In particular, the iron ore deposits at Paraburdoo (Lascelles, 1983), Mt. Tom Price, Rhodes Ridge, Hope Downs (Lascelles, in press-a), Brockman area, Jimblebar (Lascelles, 2001) and the Chichester Range (Lascelles, 2000) in the Hamersley Province, and Mt. Gibson (Lascelles, 2006) and Koolyanobbing (Lascelles, in press-b) in the Yilgarn Province were investigated over many years (Fig. 3). Since all known BIFs have undergone diagenesis and metamorphism (James, 1992), their primary nature and mineralogy must be deduced from ambiguous sedimentary structures, stratigraphy and theories of mineral paragenesis. While the origin of BIFs has been surmised to be analogous to modern iron-rich sediments precipitated from “deep-sea smokers” (Chukhrov, 1973; Isley, 1995; Krapez et al., 2003), BIFs show no similarity to the mounds of iron-rich mud forming around the hydrothermal vents and the reason for the essential differences has proved elusive. The recognition that BIFs were deposited as reworked Al-poor hydrous iron silicate and iron oxide mud (Krapez et al., 2003; Lascelles, in press-a), and that the chert bands in BIF were diagenetic, provides a link with these recent deposits. The great age of BIFs and their association with ancient cratons means that BIFs are typically found in ancient landscapes. Chemical weathering of the saprolite commonly extends hundreds of metres below the water

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Fig. 1. A. and B. Contact between cherty and chert-free BIF, Mt. Gibson (polished core sections, natural light); C. Cherty BIF saprolite, Hope Downs. Note pods of friable degraded chert; D. Chert-free BIF saprolite, Hope Downs. Ch = chert; Mt = magnetite.

table in these ancient landscapes and cycles of erosion that strip the regolith down to the water table expose the B-horizon not unweathered rock to the next cycle of soil formation. The changes that occur during weathering result in outcropping BIF appearing to be a simple, sequentially deposited, two-component system of chert and iron oxide. Outcrops of apparently fresh BIF are typically the result of the silicification of highly weathered BIF saprolite in which friable degraded chert and leached silicates and carbonates have been recemented by secondary silica. Drilling through these outcrops or exposure by cliff falls reveals the true underlying highly weathered state of the profile. Exposures of unweathered rock are extremely rare due to the rapidity of oxidation (Lascelles, in press-a) and should be viewed with suspicion, even in glaciated areas, since strongly cemented B-horizon may mimic fresh rock and may have been affected by multiple cycles of weathering. Typically, the “least weathered” samples collected from outcrops and mine workings are actually silicified B-horizon and many of the theories regarding the primary nature and origin of BIFs and the origin of

high-grade ore were based on weathered samples (Trendall and Blockley, 1970; Cloud, 1973; Trendall, 1973a,b; Morris, 1993; Klein and Beukes, 1993; Klein and Ladeira, 2004). Completely unweathered BIF obtained from deep drilling is very different and has a highly complex and variable mineralogy (Morris, 1980), from the complex diagenetic (Klein, 1983; Lascelles, 2006, in press-a) and metamorphic history (Gole, 1979, 1981; Klein, 1983; Miyano and Beukes, 1984). Recent studies of the Brockman, Weeli Wolli Iron (Krapez et al., 2003; Pickard et al., 2004) and Marra Mamba Iron Formations (Lascelles, in press-a) in the Hamersley Province (Table 1), and Mount Gibson Iron Formation in the Yilgarn Province Australia (Lascelles, 2002, 2006), of Western Australia have indicated that BIF was initially composed of hydrous iron silicates and hydroxides that were redeposited by density currents, and that the chert bands are the result of post-depositional diagenetic processes. High-grade iron ore deposits are identical in texture, mineralogy and chemistry to the adjacent cherty BIF

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except for the absence of chert bands. The discovery of significant intersections of chert-free BIF in completely unweathered diamond core (Fig. 2) from Mount Gibson (Fig. 3) led to a review of supergene processes in the origin of high-grade iron ore deposits (Lascelles, 2002, 2006). It was found that selective leaching of quartz from cherty BIF could not occur in the supergene environment and that high-grade hematite deposits derived solely from pre-existing chert-free BIF. Similarly the selective leaching of silica by hypogene processes cannot occur without leaving positive traces by affecting the silicate, carbonate and fine-grained iron oxides at the same time (Lascelles, 2006). Some of the chert-free BIF is formed by hydrothermal replacement of chert, silicates and finegrained iron oxide by carbonate (Fig. 4A–D), which is readily leached by weathering processes to form highgrade ore deposits (Barley et al., 1999; Taylor et al., 2001; Beukes et al., 2002; Lascelles, 2002; Dalstra et al., 2002; Dalstra and Guedes, 2004; Thorne et al., 2004; Lascelles, 2006; Lascelles, in press-b). At Mount Gibson and Koolyanobbing weathering of the carbonate replacement BIF gives rise to magnetite–goethite ore that is distinctive from the typical martite hematite ore but the distinction appears to have been lost during Proterozoic weathering of the Hamersley carbonate replacement ore. Mobilization of iron oxides (and silica) by hydrothermal fluids leads to distinctive veins and replacement deposits consisting of specular hematite (Miles, 1953; Griffin, 1980) (Fig. 4E, F). Many hypotheses regarding earth history have been postulated from the difference between Phanerozoic and

Precambrian iron-formations. In particular, variations in the concentration of dissolved iron, silica, oxygen and carbon dioxide in seawater through time, the probable development of photosynthetic organisms and the development of an oxygenated atmosphere have been deduced from the evolution of iron-formations (Lepp and Goldich, 1964; Cloud, 1973; Garrels et al., 1973; Holland, 1973; La Berge, 1973; Drever, 1974; Cloud, 1983; Ewers, 1983; James, 1983; Towe, 1983; Braterman and Cairns-Smith, 1987; Kump and Holland, 1992; Kuo et al., 1993; Morris, 1993; Kump et al., 2001; Holland, 2006). The subject has provided a fertile field for research and models of Earth history, but without a clear knowledge of the sedimentation, diagenesis and metamorphism of BIF, such models are speculative or even irrelevant. Since these models on the genesis and distribution of BIF relate to questions regarding the composition of the early oceans and atmosphere, weathering and transport conditions on the early land surface, volcanism and continental development in the Archean, and as they affect exploration for the largest volume and most basic of industrial metals, they have an importance beyond academic sedimentary and stratigraphic interest. Many previous models relied on assumptions that do not withstand close examination. The aim of this paper, therefore, is to present a model for the origin of BIF with global application that although still largely conjectural (Van Loon, 2004), is nevertheless free of illogical assumptions and does not conflict with field observations.

Table 1 Stratigraphy of the Hamersley Group Supergroup

Group

Formation

Turee Creek Group

Kungarra Formation Boolgeeda Iron Formation Woongarra Volcanics Formation Weeli Wolli Iron Formation Brockman Iron Formation

Hamersley Group

Mt. Bruce Supergroup

Mount McRae Shale Formation Mount Sylvia Iron Formation Wittenoom Formation

Marra Mamba Iron Formation

Fortescue Group

Jeerinah Formation

Member

Yandicoogina Shale Joffre Member Whaleback Shale Member Dales Gorge Member Colonial Chert Member Mt. McRae Shale Member Bee Gorge Member Paraburdoo Member West Angela Member Newman Member MacLeod Member Nammuldi Member Roy Hill Shale Member

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Fig. 2. Unweathered chert-free magnetite BIF, Mt. Gibson. Start of coring at bottom left. The first two trays are nearly pure magnetite with abundant iron silicate in third and fourth trays giving lighter gray colour. Chert bands (arrows) are only present in the last three rows of core at the extreme right, length of tray 1 m.

Clinton–Minette style oolitic ironstones have strong affinities with oolitic limestones, and were probably formed during anoxic periods of slow sedimentation (Young, 1989), either by the precipitation of siderite (± goethite) or by the replacement of calcium in limestone by iron during diagenesis (Siehl and Thein, 1989). The Clinton–Minette style oolitic ironstones, and paralic, lacustrine and terrestrial types of ironstones derive from a wide variety of processes and are specifically not included in this proposed new model for the origin of BIF. 2. Previous models The weathering of rocks in an anoxic atmosphere was suggested as a major source of iron and silica to the oceans (Cloud, 1973; Holland, 1973, 1984), leading to the precipitation of BIFs. However, iron-rich sediments are associated with hydrothermal vents on mid-ocean ridges, rifts, and active volcanic centres (Butuzova, 1966, 1968; Bischoff, 1972; Müller and Förstner, 1973; Butuzova et al., 1990; Binns, 2003; Yeats, 2003), and other models suggest hydrothermal activity associated with mantle plumes and ocean spreading centres and rifts as the source of iron and silica (Graf, 1978; Shegelsky, 1987; Kimberley, 1989; Isley, 1995; Isley and Abbott, 1999; Abbott and Isley, 2001; Ohmoto, 2003a), with or without additional input from fluvial sources. 2.1. Chemical precipitation The concept of BIFs as chemical precipitates is derived from its strong association with marine

sediments and typical absence of terrigenous detrital particles and granular sedimentary textures typically associated with transported sedimentary deposits, while GIFs were supposedly derived from the reworking of BIFs. Until recently, it was generally accepted that BIFs formed by alternating chemical precipitation of colloidal iron hydroxyoxides and silica (Lepp and Goldich, 1964; Garrels et al., 1973; Ewers, 1983; Garrels, 1987), as varves (Trendall and Blockley, 1970), or by periodic precipitation of iron with continuous evaporative precipitation of silica (Morris and Horwitz, 1983; Morris, 1993) in a marine shelf environment of open or barred ocean with the mesobanding supposedly due to compaction of varves. Various controlling mechanisms and cycles of deposition were postulated (Trendall and Blockley, 1970), generally based on annual, seasonal and/or Malenkovich cycles (Trendall and Blockley, 1970), but could typically only be applied to comparatively small sections of BIF sequences. The remarkable consistency down to the macroband stage of the Hamersley Group stratigraphy over the whole province, led to claims of matching microbands in drill cores hundreds of km apart (Trendall and Blockley, 1970; Trendall, 1983; Morris, 1993) that were purely coincidental and could not be demonstrated in field exposures. Little consideration was given to the origin of the iron silicates and carbonates, which were considered to be minor constituents in contrast to the chamosite-rich chert-poor Phanerozoic oolitic ironformations, apart from McConchie (1987). He suggested that colloidal iron oxides would only absorb anions from seawater at pH above 7.8, and that ancient seawater was more acidic than at present. Iron aluminosilicate-rich facies with rare iron oxide minerals

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Fig. 3. Simplified geological map of the Pilbara and Yilgarn blocks, Western Australia.

within the BIF were ascribed to the addition of airborne volcanic dust increasing the pH, alumina and phosphorous, and either reducing the photolytic oxidation of surface water or inhibiting the growth of photosynthetic organisms and production of biogenic O2 (Beukes and Klein, 1992; Morris, 1993). Evidence for this was produced by the identification of metamorphosed volcanic glass shards in some of the iron silicate-rich bands (La Berge, 1966). The assumption that oxygen was required for the precipitation of ferrous iron from solution as ferric hydroxide for the formation of oxide facies deposits, quite ignoring the low solubility of ferrous and ferrosoferric hydroxides, and notwithstanding evidence of anoxic conditions in the Archean, led to numerous hypotheses regarding atmospheric composition in the Archean Era. It was suggested that photosynthesis began at a very early age (Berkner and Marshall, 1964, 1965; Brinkman, 1969; Garrels et al., 1973; Drever, 1974; Braterman and Cairns-Smith, 1987) and, com-

bined with the photolytic breakdown of water and carbon dioxide, produced an oxygenated upper layer in the oceans (Degens and Stoffers, 1976; Towe, 1983). Iron was precipitated by the reaction of upwelling currents of anoxic water containing ferrous iron in solution with oxygenated surface water in a shelf or platform environment and the precipitation of silica was by simultaneous evaporative concentration (Drever, 1974; Degens and Stoffers, 1976; Ewers and Morris, 1981; Ewers, 1983; Morris and Horwitz, 1983; James, 1992; Morris, 1993). Archean oxygen levels in surface waters and the atmosphere were variously estimated at 0.1 PAL (present atmospheric level) to above PAL (Morris and Horwitz, 1983; Morris, 1993; Ohmoto, 1993; Beukes et al., 2000; Ohmoto, 2003a). Although there is evidence of considerable atmospheric oxygen during the early Proterozoic in the formation of oxidized sediments and palaeosols (Kasting, 1993; Rye and Holland, 1998; Beukes et al., 2000; Holland, 2002; Lascelles, 2002), so far there has been no indication of oxidized palaeosols or sediments apart from sulphates (Huston and Logan, 2004) in the Archean, and there is some evidence for a reducing atmosphere, (Roscoe, 1968; Fryer, 1977; Grandstaff, 1980; Kimberley et al., 1980; Rasmussen and Buick, 1999; England et al., 2002; Yang et al., 2002; Frimmel, 2005). Biogenic action has also been suggested as a factor in the precipitation of the iron oxides in BIF (Konhauser et al., 2002; Brocks et al., 2003; Fortin and Langley, 2005) and sedimentary organic matter as a dominating influence on the formation of Phanerozoic blackband ironstones and oolites (Spears, 1989). Many problems are unsolved by these scenarios, in particular, iron and silica in solution react to precipitate hydrous iron silicates such as nontronite (Moore and Maynard, 1929; Müller and Förstner, 1973; Yamamoto, 1988), chamosite or glauconite, and the microbands are typically grouped into iron oxide or silica mesobands of very variable thickness, which is difficult to reconcile with seasonal varves and Milankovitch cycles. Moreover, although chert horizons are common throughout the Precambrian there is little evidence to support evaporative precipitation in other sediments. Lepp and Goldich (1964) argued against direct precipitation of silica and suggested that the chert replaced other sediments during diagenesis. Simultaneous precipitation of various products must form unsorted deposits and only current deposition can produce laminated deposits (Paolo et al., 1989). Furthermore, Krapez et al. (2003), Pickard et al. (2004) demonstrated that chert is a diagenetic product replacing iron silicates and is not a primary precipitate.

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2.2. Origin of iron and silica Various models have been presented for the source of the iron and silica in BIF; from dissolved Fe2+ and SiO2 derived from surface weathering of continents (Cloud, 1973; Drever, 1974; Cloud, 1983; Garrels, 1987), from

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diagenetic or hydrothermal leaching of underlying sediments (Holland, 1973); or from volcanigenic sources (Van Hise and Leith, 1911; Trendall and Blockley, 1970; Morris and Horwitz, 1983; Kimberley, 1989; Butuzova et al., 1990; Gross, 1991; Isley, 1995; Isley and Abbott, 1999; Krapez et al., 2003), as the

Fig. 4. Hydrothermal ore deposits. A. Magnetite/carbonate BIF, Mt. Gibson, dark bands are magnetite, light bands are carbonate; B. Magnetite/ carbonate BIF, Mt. Gibson (plane polarized transmitted and reflected light); C. Replacement of chert by carbonate adjacent to small calcite vein, Mt. Gibson (plain transmitted light); D. Carbonate replacing iron silicates, fine-grained magnetite and chert, Mt. Gibson (plain transmitted light); E. Veins of specular hematite in cherty BIF, Weeli Wolli Formation, Paraburdoo (polished slab); F. Coarse-grained specular hematite, K deposit, Koolyanobbing. Ca = calcite; Carb = carbonate; Ch = chert; Chl = chlorite; Mt = magnetite; Qz = quartz; Sil = iron silicates.

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amount of ferrous iron in the oceans was typically calculated to have been inadequate (Trendall, 1965; Morris, 1993) to supply the vast quantities required for the Paleoproterozoic iron-formations. Others proposed that the supergene leaching of iron from continental landscapes in an anoxic atmosphere (Cloud, 1973; Holland, 1973; Fryer, 1977; Holland, 1984) would have provided sufficient iron in solution (Huston and Logan, 2004), and that increased photosynthesis and the change to an oxidizing atmosphere was the reason for the vast deposits of iron-formation in the Paleoproterozoic and their greatly diminished volume thereafter. However, the requirement for an oxygenated surface layer implies oxic weathering from 3.8 Ga (the age of the oldest known BIF; Appel, 1987) thus eliminating the land as a source of ferrous iron. Oxygen solubility is controlled by the partial pressure of O2 in the atmosphere (Kump and Holland, 1992) and an oxygenated surface layer cannot exist with a reducing atmosphere. However, this would not affect the input of ferrous iron from hydrothermal vents that might accumulate in a stratified anoxic sub layer, but such a layer would require continuous maintenance to offset diffusion of oxygen from the surface. In modern seas, anoxic conditions are maintained through the removal of oxygen by decay of organic material. Studies of the rare earth element geochemistry of BIF (Graf, 1978; Beukes and Klein, 1990; Derry and Jacobsen, 1990; Huston and Logan, 2004) suggested a bimodal origin for the iron from both oceanic and volcanic sources although Huston and Logan (2004) calculated that ample iron could have been present in solution. Barley et al. (1997) demonstrated the presence of a large igneous province that accompanied the deposition of the Hamersley Group, and other Archean and Proterozoic BIF sequences are similarly related to igneous provinces (Bayley and James, 1973; Button, 1976) even if not to volcanic centres. Nevertheless, debate still continues on the importance or otherwise of the oceanic accumulation of ferrous iron derived from anoxic weathering in the Archean (Huston and Logan, 2004). The development of deep ocean research over the second half of the 20th Century has demonstrated the abundance of iron- and silica-rich hydrothermal vents near mid-ocean ridges (Boström et al., 1969; Bischoff, 1972; Von Damm, 1990) and continental rifts (Müller and Förstner, 1973), mantle plumes (Karl et al., 1988) and back-arc basins (Butuzova, 1968; Yeats, 2003). Iron and silica leached from igneous and volcanic rocks by hydrothermal fluids (Bischoff and Dickson, 1975) is precipitated on contact with seawater; iron-rich sediments accumulate around the vents, and plumes of fine-grained

suspensions have been traced for thousands of kilometres across the ocean floor (Lonsdale, 1976; Lupton, 1996). Numerous authors have proposed these hydrothermal vents as the source of iron and silica for BIF (Von Damm, 1990; Isley, 1995; Barley et al., 1997; Abbott and Isley, 2001), and the relationship between hydrothermal systems and sediments appears convincing for deposits of the Lahn–Dill type (Chukhrov, 1973; Kräutner, 1977), but no such hydrothermal sources have yet been identified for BIF. The typically uniform laminated sedimentation over large areas is inconsistent with hydrothermal mounds and furthermore the modern iron-rich deposits from underwater hydrothermal vents differ considerably from BIF, consisting mainly of at best poorly bedded hydrous Al-poor iron silicate (nontronite) and ferric hydroxide (goethite) (Boström et al., 1969; Bischoff, 1972; Müller and Förstner, 1973). They typically contain abundant manganese, chert banding is generally minor and nothing on the scale of the large Precambrian deposits has been seen (James, 1983). 2.3. Depositional environment Previous authors proposed a variety of environmental characteristics required to alternately precipitate iron oxides and silica; whether anoxic open marine basins, continental shelves, isolated platforms (Morris and Horwitz, 1983), silled basins (Trendall and Blockley, 1970) or enclosed sabkhas (Garrels, 1987). Classifications of the various types of iron-formation have been used as a basis for models of deposition. Ironformations have been divided into oxide, silicate, sulphide and carbonate facies (James, 1954) supposedly relating to water depth or distance from source (Shegelsky, 1987), and Kimberley (1978) further classified laminated iron-formations into shallow-water or deep-water facies mainly based on sedimentary structures and associated rock types. However, the discovery of turbidites interbedded with Hamersley Group formations (Krapez et al., 2003; Pickard et al., 2004), has reopened the depth question since the sedimentary structures, such as rip-up clasts and ripples and the current bedding in peloidal and oolitic GIF, by which Kimberley (1978, 1989) defined laminated ironformation as being deposited in shallow water, are also produced by turbidity currents in deep water (Bouma, 1962). The issue is further confused by the interpretations of submarine fan deposits, which typically form from turbidity currents towards the base of the continental slope, as storm-derived mega-ripples (Hall and Goode, 1978; Goode et al., 1983; Pufahl and Fralick, 2004) and the interpretation of syneresis cracks

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Table 2 Uniformitarian model of marine iron-formations Stage

Minette type oolites

Precambrian BIF and GIF

Lahn–Dill type

Recent “black smokers”

1. Leaching and transportation of iron and silica to oceans from subaerial weathering 2. Leaching of hot mafic volcanic rocks by circulating seawater at spreading centres and over mantle plumes 3. Venting as hydrothermal smokers rich in silica, iron and manganese 4. Rapid cooling and reaction with sea-water precipitating chimneys and mounds of sediment around vent 5. Slumping of mounds generates turbidity currents 6. Coarse particles form proximal granular–oolitic iron-formations with interbedded fine-grained overbank deposits and toe mudrocks 7. Colloidal particles form density currents, transported and spread out over great distance, deposited as finely laminated hydroxides and silicates 8. Diagenesis and compaction

None (organic iron complexes and detrital clay)

Build up of Fe2+ and SiO2 in ocean

None

None

No apparent connection

Yes

Yes

Yes

No apparent connection

Yes

Yes

Yes

No apparent connection

Precipitation of Al-poor iron silicate and iron hydroxide

Oxidation and precipitation of iron and manganese silicates and hydroxides

Oxidation and precipitation of iron and manganese silicates and hydroxides

No apparent connection

Yes

Yes

Yes

No apparent connection

Yes

Yes

Yes

Epierogenic deposition of oolites during periods of slow sedimentation — iron aluminosilicates Clay crystallizes to chamosite; iron hydroxide to hematite–goethite None Typically none

Al-poor iron silicate and iron hydroxides

Al-poor iron and manganese silicates and hydroxides

Al-poor iron and manganese silicates and hydroxides

Clay crystallizes to iron oxide and amorphous chert; iron hydroxide to magnetite–hematite Yes Metamorphosed to magnetite and microquartz with minor minnesotaite

Clay crystallizes to iron oxide and amorphous chert; iron hydroxide to magnetite–hematite Yes May be metamorphosed

Poorly advanced

9. Chert banding 10. Metamorphism

as subaerial desiccation. Unfortunately diagnostic sedimentary structures other than the laminations and banding in BIF are typically observed only in associated and interbedded sedimentary rocks. The classification of iron-formations by Gross (1983) into Lake Superior and Algoma types was based on the various iron-formations associated with the Ungava craton of North America. It was intended to distinguish the areally widespread, apparently shallow water granular and oolitic textured Lake Superior type iron-formations, from the finely laminated or mesobanded Algoma type iron formations. The Lake Superior type were also associated with clastic and carbonate sediments com-

Rare None

monly containing stromatolites and deposited in basins on continental shelves adjacent to cratons, without obvious volcanic associations. The Algoma type iron-formations were typically associated with greywacke, turbidites, finegrained clastic sediments and volcanic rocks in greenstone belts within the cratons and supposedly deposited in narrow troughs on sea floors near volcanic centres (Gross, 1983). The contrast in depositional environments was used to explain the typically fine-grained, highly deformed and metamorphosed narrow lenticular bodies of the Algoma type BIFs compared to the relatively undeformed and major scale of the mainly granular Lake Superior iron-formations. Nevertheless, a gradual

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transgression from shallow to deeper water and from the granular and oolitic facies to microbanded facies was envisaged with examples from the Ungava and Kaapvaal cratons (Beukes, 1980; Gross, 1983; Beukes, 1984). Australian BIF units occurring within greenstone belts on ancient cratons such as the Yilgarn and Pilbara Provinces are generally classed as Algoma type, and the GIF Frere Iron Formation in the Nabberu basin (Fig. 3) is considered to be Lake Superior type (Hall and Goode, 1978; Goode et al., 1983; Krapez and Martin, 1999). Petrologically the Hamersley Group matches the finegrained laminated Algoma classification but because of its great areal extent, supposedly shallow depth of deposition and underlying sialic crust (Morris and Horwitz, 1983) it was typically classified as the Lake Superior type or given a separate class (Hamersley type). Gross (1983) gave a separate classification to Neoproterozoic Rapitan type BIF due to the association with glaciogene rocks and the Paleozoic Lahn–Dill iron-formations apparently formed in small euxinic basins on the flanks of active volcanoes. 3. The new model 3.1. Introduction Current deep-sea iron-rich deposits appear to be formed initially in the vicinity of sea floor hydrothermal vents (Boström et al., 1969; Edmond et al., 1979) and later dispersed by turbidity and density currents to form laminated fine-grained iron-formations (Binns, 2003). A similar origin was proposed for the Devonian Lahn–Dill and similar deposits in Europe (Quade, 1970, 1976; Kräutner, 1977). The new model for marine iron-formations presented here (Table 2) groups the granular and banded ironformations as resedimented hydrothermal precipitates related to active margins and spreading centres in contrast to the oolitic ironstones formed on stable epeirogenic shelves without apparent hydrothermal or volcanigenic input (Young and Taylor, 1989). The model implies a deep-water environment for the formation of BIFs and GIFs that are transported and redeposited sediments derived from the initial precipitates around the hydrothermal vents. The model adheres to uniformitarian principles for the origin of BIFs since it is based on processes currently observed in marine sedimentary deposits. Unfortunately, no such processes are currently forming Clinton–Minette type oolitic ironstones. High concentrations of ferrous iron in solution (Huston and Logan, 2004) derived from anoxic

weathering in the Archean era became supersaturated with the input of ferrous iron from hydrothermal sea floor vents and were precipitated to form the vast ironformations of the Precambrian — a source that was unavailable to post-Paleoproterozoic BIF limited to hydrothermal input only (Table 2). The high-temperature hydrothermal fluids also leach minor ferric iron (McGuire et al., 1989), probably as chloride complexes that were probably sufficient to induce the precipitation of iron in the form of ferroso-ferric hydroxide (green rust) and some inclusion into the structure of the hydrous iron silicates. 3.2. Colloidal precipitation The formation of crystalline precipitates is the normal mode for ions in solution, but the near instantaneous supersaturation of ions with low crystallization potential, such as the transition elements and silica, by hot hydrothermal solutions venting into cold seawater causes the formation of colloidal suspensions (Yariv and Cross, 1979). Previous models for the chemical precipitation of BIF have typically considered in detail the ionic solubilities of iron and silica, but rarely consider the properties of colloidal suspensions, the precipitation mechanisms of colloids, the colloidal properties of the precipitates (Hiemenz, 1997), the sedimentary characteristics of ultra-fine suspensions or the diagenesis of such sediments. Silica in solution occurs mainly as Si (OH)4 monomer (Okamoto et al., 1957) and readily polymerises to form colloidal suspensions (Yariv and Cross, 1979) that combine with ferrous iron in solution to precipitate as hydrous iron silicate (nontronite) (Marshall, 1949; Hiemenz, 1997). Since ferrous iron is more soluble than silica at ambient temperatures (Moore and Maynard, 1929; Loughnan, 1969) excess ferrous iron may accumulate in solution. However, ferrous iron in solution will also form insoluble hydroxide, carbonate and sulphide, scavenges dissolved oxygen to precipitate as ferroso-ferric hydroxide or may have been converted to ferric iron by the photosynthetic or biogenic action of micro-organisms (Cloud, 1973; Konhauser et al., 2002). Residual concentrations of iron oxides (and carbonate?) are precipitated together with the iron silicates regardless of the redox state of the ocean (iron sulphide precipitation is controlled by the availability of sulphide ions). Colloidal suspensions behave somewhat differently from solutions in that colloidal silica is more stable than colloidal iron oxide (Hiemenz, 1997). Both iron oxides and hydrous iron silicates have very low crystallization

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Fig. 5. Deposition of iron-formations. A. “Black smoker” emits Fe2+ and H4SiO4 into ocean precipitating hydrous iron silicate and iron hydroxide around vent. Fe2+; B. slumping of unstable mud pile generates turbidity current from debris flow; C. coarse material is deposited close to source as lobes of granular iron-formation and fine-grained overbank deposits with colloidal particles remaining suspended in density current; D. colloidal particles eventually settle out as finely laminated sediments far from the source.

potentials at ambient temperatures and typically precipitate in the form of colloidal instead of crystalline aggregates (Yariv and Cross, 1979). Edge effects of the double layer charge on colloidal particles induce flocculated iron oxides to aggregate in thin sheets (Hiemenz, 1997). These hydrophobic oxide laminae display the property of semi-permeable membranes allowing slow diffusion of water and solutes but forming a barrier to hydrophilic colloids such as silica gel (Yariv and Cross, 1979; McConchie, 1987; Hiemenz, 1997) that accumulate in layers beneath continuous iron oxide laminae. Colloidal iron oxides are strongly attracted to crystalline silicate minerals to form adherent coatings (Hiemenz, 1997) but there is no trace of such coatings on the iron silicates and microquartz in BIF, which typically show clear separation of the iron oxides and siliceous minerals due to their contrasting colloidal properties (Hiemenz, 1997). 3.3. Deep-sea smokers Hydrothermal vents have been described from many places on the sea floor in particular from spreading ridges and recently active sea mounts (Boström et al., 1969; Von Damm, 1990; Kelley et al., 2001) as well as major continental rift systems (Müller and Förstner, 1973;

Butuzova et al., 1990). Apparently produced by oceanic waters circulating through lavas, sills, dykes and volcaniclastic sediments, the heated fluids, rich in both dissolved metallic ions and silica, are rapidly cooled on contact with cold ocean water, reducing the solubility of the entrained elements and producing supersaturation of the least soluble constituents, which are immediately precipitated as colloidal particles producing the typical plumes of black or white “smoke” (Parr et al., 2003). The colloidal particles of iron and silica combine as poorly crystalline Al-poor hydrous iron silicate (nontronite, Yariv and Cross, 1979), and since iron and manganese are more soluble than silica (Loughnan, 1969; Bischoff and Dickson, 1975) the excess is precipitated as hydroxides forming mounds of water-saturated ironrich fine mud (Fig. 5A) around the sulphidic chimneys (Edmond et al., 1979; Binns, 2003). Iron and manganese both have soluble divalent and insoluble trivalent forms but higher O2 fugacities are required for the precipitation of manganese (Anbar and Holland, 1992). Oxygen tends to be depleted by the oxidation of ferrous iron in the vicinity of the black smokers such that the sediments adjacent to the vent may contain a high proportion of ferrous iron, and the proportion and distance from the source of the manganese precipitation will depend on the amount of O2 available. Titanium tends to be precipitated

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Fig. 6. Frere Iron Formation. A. Lenticular bedding in the Frere Iron Formation; B. parallel bedding in the Frere Iron Formation. The photo in A was taken of the shaded face below the large shrub in the centre of the picture at right angles to the bedding shown in B; C. graded bedding in coarse peloidal iron formation. White spots are due to weathered voids; D. bands and lenticular pods of fine-grained pink jasper in GIF. (C and D are polished slabs of weathered surface grab samples); E. block diagram of Frere Range GIF; F. sketch diagram of GIF submarine apron of coalesced fans.

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with alumina and is comparatively abundant in iron aluminosilicates but rare in BIF. The chimneys and mounds of water-saturated fine mud are typically unstable (Binns, 2003) and debris flows (Fig. 5B) as seen in Lahn–Dill deposits (S. Hagemann, pers. comm., April 2006), probably triggered by seismic events (Rundquist and Sobolev, 2002), gave rise to turbidity currents (Hampton, 1972) carrying the sediments into deeper water. The coarser particles supported by the turbidity current are rapidly deposited in fans (Fig. 5D) as the currents decay forming lobes of graded coarse-grained deposits with laminated fine-grained levees and overbank deposits (Hesse and Chough, 1980; Bouma, 2000). Colloidal particles supported by Brownian motion remain in suspension after the turbidity current has decayed and may be transported as density currents and spread out over great distances in laminated deposits as the suspended particles slowly settle out (Fig. 5D). 3.4. Granular iron-formation Iron-formations in the Frere Range of the Nabberu Basin in Western Australia (Hall and Goode, 1978; Goode et al., 1983; Krapez and Martin, 1999) contain lenticular coarse peloidal hematite and chert gravels interbedded with jasper and fine-grained laminated iron mudrocks in the south (Fig. 6A–D) grading into laminated BIF in the northern parts of the basin. Previously interpreted as generated by storm waves, close examination of the formation revealed that each horizon consists of a strongly elongated lobe (Fig. 6E) of coarse sandy gravel passing laterally into fine-grained iron-rich mud that gradually diminishes with distance and is overlain by successive lobes. Graded and current bedding is present though poorly defined whereas the regular repetition of coarse lenses at each horizon (megaripples) characteristic of storm wave deposits (Tamura and Masuda, 2005) is absent. The Frere Iron-formation consists of submarine fans (Fig. 6F) deposited by turbidity currents with channels of coarse sediment bounded by levees and fine-grained overbank deposits (Hesse and Chough, 1980; Bouma, 2000). The underlying Yelma Formation contains shallow-water coarse clastic sediments and stromatolitic limestones but the iron-formations in the Frere Formation are separated by thick-bedded siltstones with rare ripple-bedded horizons. 3.5. Density current deposition Density currents formed from ultra fine-grained colloidal particles show the same depositional sequences

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as typical fine-grained sedimentary turbidites, but on a finer scale. Pelagic clay consisting mainly of iron aluminosilicate typically has larger particle sizes and settles faster than the colloidal particles in suspension and thus forms the base of flows that commonly show a limited erosive power. The absence of significant size differential between colloidal components of the density flow means that grading of sediments is by density and that fine-grained iron oxide particles are deposited before the less dense iron silicate and carbonate particles. Ripple layering and small-scale current bedding is rarely seen in small sections of BIF sequences, especially in drill core although it has been suggested that due to the gelatinous nature of the sediment (McConchie, 1987; Parr et al., 2003), ripple and convolute bedding (Fig. 7A, B(a)) could not be supported and would be obliterated by gravitational settling. Whether the mesobanding was formed by the same process as the microbands but on a coarser scale, or was due to compaction of microbands during diagenesis is uncertain as recrystallization commonly obscures the fine laminations in the oxide mesobands, and possibly either mechanism may apply locally. Mesobanding is typically less well defined in the chert-free parts of the BIF but is by no means absent (Fig. 2) although much of the apparent mesobanding in high-grade ore is due to secondary cementation. Individual density currents deposit a sequence of laminated iron hydroxyoxides and silicates, typically commencing with black shale and carbonate bands, and culminating in a silicified upper surface, producing the ideal macroband sequence of S band and BIF band. However, in many cases a second or third density current that may erode the semi-consolidated uppermost layer rapidly follows the first and the macrobands may consist of a sequence of several flows. Typically, but not always, a thin shale band or parting is formed at the base of each flow. S bands also commonly consist of turbidity–density current sequences but with a large component from a source of fine- to medium-grained sediments near the hydrothermal mounds. The absence of shallow water sedimentary characteristics and terrigenous sediment, thickness of sediment, and the presence of turbidites within BIF sequences (Hyde, 1980; Fralick, 1987; Krapez et al., 2003; Pickard et al., 2004; Pufahl and Fralick, 2004) suggest that while not necessarily abyssal the BIF were deposited in deep water beyond the reach of terrigenous sedimentation. No features that show the source area of the density currents have been identified nor are any of the usual indicators of transport direction typically present. Numerous compaction, dewatering, creep and slump structures indicate the highly water-charged initial

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Fig. 7. Sedimentary structures. A. Convoluted bedding in fine-grained carbonate caused by current drag, Hope Downs. B. Globular chert clast (a) deformed by current drag that has convoluted the overlying fine-grained carbonate layer (b) and an inverted sub-rectangular chert rip-out clast (c) in black shale matrix (d); C. anastomosing iron oxide laminae around chert clasts (Hope Downs). Ch = chert; He = hematite; Sh = shale.

sediment of flocculated colloids rather than solid particles (McConchie, 1987), in contrast to the granular iron-formation deposits. Turbidity currents decay within hours of formation but density currents may persist for several days and spread out over wide areas. 3.6. Phanerozoic iron-formation or BIF — changes in composition of the oceans Recent iron-rich mud accumulating around hydrothermal vents differs both chemically and mineralogically from Precambrian BIF (Table 1). To infer that BIF formed by the same processes of hydrothermal leachate venting onto the sea floor, then these differences must be explained. Iron oxides have very low solubility under the current strongly oxidizing weathering conditions, with mobilization restricted to complexes with organic humic acids or colloidal suspension, and although decaying organic matter with limited access to the

atmosphere may cause local anoxic conditions in which ferrous iron is stable, very little iron reaches the oceans from continental sources (Edmond et al., 1979). The build-up of free oxygen in the oceans, and consequently the atmosphere, was necessarily slow during the Archean and anoxic weathering transported abundant ferrous iron into the oceans. Ferrous hydroxide and carbonate are also highly insoluble in neutral solution and photolytic and photosynthetic oxygen would have been scavenged by ferrous iron to be precipitated as ferroso-ferric hydroxide (Yamamoto, 1988). Nevertheless, the concentration of ferrous iron in solution would continue to build up. It has also been assumed that the oceans were weakly acidic due to dissolved CO2 (McConchie, 1987), but the presence of Ca, Mg and Fe in solution as bicarbonates from weathering processes gives an alkaline reaction and the oceans were more likely to have been neutral to weakly alkaline. It seems clear that O2 levels increased

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sufficiently to have impacted subaerial weathering processes by the mid-Paleoproterozoic era (Rye and Holland, 1998; Beukes et al., 2000) and eventually sufficient O2 was produced to precipitate all dissolved ferrous iron and by oxic weathering prevent the further input of iron into the oceans. Of greater significance than oxidation to the difference between Precambrian and Phanerozoic iron-formations was the absence of the silica-secreting organisms that maintain dissolved silica at a very low level in modern ocean water (Heinen and Oehler, 1979). The solubility of silica is reduced in the presence of alumina and vice versa (Yariv and Cross, 1979) such that each act as scavengers of the other with the precipitation of authigenic hydrous aluminosilicate (Mason, 1966), until one is completely depleted. The high alumina content (N 10 wt.%) that is characteristic in both the aluminosilicate-rich Phanerozoic oolitic ironstones (Lepp and Goldich, 1964) and the shale bands in Precambrian iron-formations, is probably derived from fine dispersions of terrigenous or pelagic clay particles normally present even in oceanic waters far from land. Together with airborne pyroclastics, they accumulate during periods of slow deposition to form aluminosilicate horizons. Terrigenous and pelagic sediments and volcaniclastics may also be entrained in the initial debris flow from the hydrothermal mounds and transported by the turbidity currents. However, alumina values are characteristically very low (b1 wt.%) in BIFs rapidly deposited from density currents that are derived from the underwater hydrothermal mounds (Table 2). Biogenic processes play a very important role in the chemistry of modern oceans, through photosynthetic oxygenation to the extraction of silica and calcium carbonate (Heinen and Oehler, 1979). Phanerozoic chert is typically formed from the skeletons of silica-secreting organisms, and is not associated with oolitic ironstones, although non-skeletal chert is commonly present in Lahn–Dill and Rapitan iron-formations. No trace of recognizable silica-secreting organisms has been found in Precambrian chert (Lepp and Goldich, 1964) although La Berge (1973), describing vaguely spheroidal structures in certain chert, suggested that an early form of now extinct silica-secreting bacteria or algae might have existed. Certainly the Precambrian and Lahn–Dill cherts do not consist of the skeletons of diatoms. Phanerozoic carbonate rocks derive a large part of their content from skeletal material, but also from the precipitation of carbonate by removal of CO2 during photosynthesis. Early stromatolites may have precipitated calcium carbonate through removal of CO2 from solution and this is presumed to be evidence for photosynthesis (Hartman, 1984).

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Hydrous iron silicates require the presence of alumina within the silicate lattice for stability and Alpoor amorphous hydrous iron silicate (nontronite) dissociates to colloidal iron oxides, silica and water during diagenesis (Caldwell and Marshall, 1942); i.e., goethite or ferrous hydroxide have higher activation energies of crystallization than nontronite. Minor amounts of Fe 3+ leached from the source rocks (McGuire et al., 1989), or produced by oxidation in post-Lower Proterozoic, may substitute for Al imparting greater stability to the silicate lattice of the Al-poor hydrous iron silicates. Rarely preserved primary Al-poor iron silicate within Precambrian BIF is typically adjacent to or within oxide laminae. 4. Discussion 4.1. Timing of density current formation The initial precipitate of colloidal hydrous iron oxides and silicates slowly crystallizes during diagenesis, but in most cases this had barely started before reworking and resedimentation from density currents. Crystallized iron oxides and aggregated particles are deposited rapidly from turbidity currents forming granular graded and fine laminated deposits, but density currents may transport colloidal material over large distances. The paucity of grain and sedimentary structures other than uniform fine laminations and banding within BIF, and the widespread uniform stratigraphy strongly supports the deposition of BIF as colloidal particles from slow-moving and longlived density currents. Diagenesis of the colloidal precipitates (Parr et al., 2003) commences slowly from the moment of settling, with flocculation of the colloidal particles, the excretion of pore water and the commencement of crystallization and corresponding dissociation of Al-poor iron silicates. The frequency of turbidity–density current formation due to slumping of the primary mounds (Hampton, 1972) would depend on the rate of accumulation of the hydrothermal mud around the source and the frequency of triggering events such as seismic disturbances (Rundquist and Sobolev, 2002), chimney collapse (Binns, 2003), etc., which would be abundant in active volcanic areas. Slumping of newly formed mounds gives rise to density currents consisting essentially of colloidal iron oxides and silicates. Older mounds in which diagenesis had commenced would produce clastic material of cryptocrystalline iron hydroxides and gelatinous to semi-lithified silica with minor iron silicates. This particulate material would be transported and deposited

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Fig. 8. Chert formation. A. Chert replacing iron silicate, remnants of minnesotaite remain in magnetite mesobands and adjacent to concentrations of magnetite grains in silicate mesoband, Mt. Gibson (crossed polarizers); B. dissociation of iron silicate to chert and fine-grained dusty magnetite, magnetite settles to base of chert band Mt. Gibson (plane polarized transmitted and reflected light); C. interbedded iron silicates and magnetite Hope Downs (plain transmitted light); D. secondary minnesotaite formed by reaction between magnetite and chert, Mt. Gibson (plain transmitted light); E. upper surface of chert layer showing pinch and swell and abundant microfaults in chert (length of pencil 15 cm), Hope Downs. F. Pattern of intersecting wave crests (solid lines) and microfaults (dashed lines) in E. Ch = chert; Mn = minnesotaite; Mt = magnetite; Sil = iron silicate (minnesotaite with traces of stilpnomelane).

as graded beds and laminated deposits of coarse to finegrained clasts by turbidity currents comparatively close to the source (Hesse and Chough, 1980), in addition to density currents of residual fine-grained colloidal material. Much of the silica would be deposited as

granular chert together with iron oxide and iron silicates, and some redispersed as hydrous colloidal particles, but numerous pods and spheroids of semi-gelatinous chert would be suspended within the turbidity current and due to their low specific gravity remain suspended in the

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subsequent density current. Deposition from the density currents of iron oxides and iron silicates forming typical finely laminated deposits would ensue with subsequent continued diagenesis and the formation of chert bands. Reaction between the sediment and seawater occurred during the sedimentation hiatus between flows with the formation of silicified hardgrounds that were commonly partly eroded by succeeding density currents producing shallow scouring and rip-outs. Individual horizons in BIF commonly display uniformity in the size and shape of clastic chert pods consistent with sorting and settling from slow-moving density currents contemporaneously with the flocculation and settling of iron oxides and silicates. Surface tension in separate portions of semi-gelatinous chert would tend to produce ellipsoidal to spheroidal shapes with a distinct surface “skin” (Fig. 7B(b)) during suspension and transportation in contrast to the bands of chert that formed in situ by diagenesis. Clastic chert pods commonly form distinct marker bands (e.g., ‘Potato Beds’) within the Hamersley Group Marra Mamba Iron Formation and show characteristic internal shrinkage with stellate and annular cracks due to hardening of the outer layers of the chert pods, whereas chert bands formed in situ typically show syneresis cracks extending from the surfaces of the chert (Lascelles, in press-a). Deposition of iron oxides simultaneously with deposition of transported silica gel produces the appearance of anastomosing iron oxide layers (Fig. 7D). However, localized slumping and intraformational deformation of the soft sediments may also form similar structures. It is apparent that the origin of structures involving the chert pods is highly complex and virtually every chert band requires individual interpretation concomitant with the enormous variety shown by the chert bands. Nevertheless, once BIF is accepted as deposited by density currents, the reasonable interpretation of the structures becomes possible. 4.2. Origin of the laminations and banding Krapez et al. (2003) demonstrated that the S bands in the Brockman BIF show characteristics of turbidity current deposition of fine-grained semi-colloidal mud, that the BIF is reworked and transported sediment and the chert bands are of diagenetic origin. Laminations and banding in sedimentary rocks are typically produced by current deposition with comparatively rare occurrences of banding produced by seasonal fluctuations in composition such as varved clays (Trendall and Blockley, 1970) and evaporite deposits in still lake waters (Garrels, 1987). However, all evidence from

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associated sediments points to a marine environment for laminated iron formations. Furthermore, marine environments are rarely still with slow oceanic currents circulating even in the deepest ocean trenches (Mackay et al., 2002), and are typically punctuated by rapid turbidity flows and density currents. Normal deposition from density currents would be expected to produce banding of the iron oxide and silicate (Bouma, 1962; Stow and Bowen, 1980; Bouma, 2000), by pulsations in the flow that are an inherent property of fluid movement (Hesse and Chough, 1980; Paolo et al., 1989). Three factors have prevented the general recognition of the deposition of BIF by density currents. Firstly, the weathering of BIF results in a two component rock consisting of alternating layers of iron oxide and chert that was generally accepted as the result of direct precipitation and secondly, the diagenetic dissociation of Al-poor hydrous iron silicates into colloidal silica and iron oxides (Fig. 8A, B) prevented the direct comparison with modern deep oceanic sediment samples collected from the ocean floor. Thirdly, the assumption that oxygenated surface water was required to precipitate ferric hydroxide implied that deposition would occur on continental shelves where the mixing of upwelling deep oceanic water with surface layers typically occurs. Rare features indicative of turbidity current deposition such as ripple and current bedding, shallow scouring and ripup clasts were presented as evidence for shallow water deposition (Kimberley, 1978) and submarine fans with lobes of coarse sediment in GIF (Fig. 6) were interpreted as mega-ripples formed by the winnowing action of storm waves (Goode et al., 1983; cf. Levell, 1980). The presence of ooliths and stromatolites in Lake Superior iron-formations (Barghoorn and Tyler, 1965; Goode et al., 1983; Walter and Hofmann, 1983; Simonson and Lanier, 1987) was also considered to be evidence of shallow water sedimentation, and together with granular iron-formations grading into laminated iron-formations and BIF (Hall and Goode, 1978; Beukes, 1980; Gross, 1983; Beukes, 1984; Beukes and Klein, 1990), combined to reinforce the shallow water origin of BIF (Morris and Horwitz, 1983). 4.3. Diagenetic origin of chert Colloidal silica is leached from surface rocks as a result of the weathering of silicate minerals and transported to the oceans by fluvial processes. Phanerozoic marine cherts are almost exclusively formed from the tests of silica secreting organisms, but no evidence has been found of such organisms during the Precambrian era although cherts are abundant and some contain

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Fig. 9. Diagenesis of BIF. A. Deposition of iron silicate with continuous and discontinuous laminas of iron oxides; B. dissociation of iron silicate into iron oxides and colloidal silica; C. differentiation into mesobands by settling of iron oxide grains and accumulation of silica below oxide laminae; D. dehydration and compaction. Loss of water by diffusion and escape of colloidal silica through fractures.

rare microfossils (Walter and Hofmann, 1983). Prior to the evolution of silica-secreting organisms, it was suggested that there was abundant silica available in Archean and Paleoproterozoic waters (Drever, 1982; Simonson, 1985; Maliva et al., 2005). Hydrous iron silicates (nontronite) are inherently less stable than hydrous iron aluminosilicates (Caldwell and Marshall, 1942; Yariv and Cross, 1979) and during diagenesis tend to dissociate into iron oxides and colloidal silica. The dissociation of Al-poor iron silicates does not appear to have been universal as remnants are common in microcrystalline quartz bands (Fig. 8A) and many primary sequences of magnetite interbanded with minnesotaite occur (Fig. 8B,C), where diagenetic dissociation of Al-poor iron silicate has not occurred. Secondary growth of minnesotaite from the reaction between quartz and magnetite during metamorphism is typically nucleated on magnetite grains and grows inward to chert bands (Fig. 8D), in contrast to the bedding parallel orientation of primary iron silicates. Some of the iron present in the source rocks and subsequently in the vent fluids will be Fe3+ (McGuire et al., 1989) and by substituting for Al in the silicate lattice increases the stability of Al-poor iron silicates. Magnetite microbands typically alternate with silicate microbands (Fig. 8C) and iron oxide–microquartz microband pairs are uncommon in most unweathered cherty BIFs. More commonly, the colloidal silica accumulates as distinct mesoband scale bands of chert since the hydrophilic colloidal silica is trapped by laminae of hydrophobic colloidal iron oxide (McConchie, 1987; Hiemenz, 1997) to form thin initially parallel layers. The chert bands typically include abundant internal fine

laminations and bands of disseminated dusty magnetite from the dissociation of the nontronite, with traces of unaltered accessory aluminosilicate, whereas aluminous iron silicate mesobands are typically unaffected and preserved as shale horizons. The crystallization of the chert bands from colloidal silica is evidenced by numerous syneresis cracks and many chert pods with stellar internal shrinkage cracks illustrate the large volume change between colloidal silica and microquartz, consistent with that expected for the crystallization of colloidal silica to form chert. Polymerisation and shrinking during syneresis tends to form the chert layers into spheroidal globules or pods that is accelerated by mechanical disturbances due to seismic shock or slumping, creep and the density differential between the iron oxides and the silica gel. The topology of chert bands is highly variable depending on the rates of sedimentation, diagenesis, slope of the sediments and microseismic disturbances. Calm periods of horizontal deposition produce relatively thick sequences of evenly banded chert layers. Microseismic shockwaves tend to form sublinear arcuate forms and moiré patterns formed by intersecting waves from separate centres are difficult to ascribe to other sedimentary processes (Fig. 8E, F) while creep and slumping may form rods and imbricated lenticular pods. Breaches of the colloidal oxide laminae through auto-brecciated zones or other disruptions of the layering allow the liquefaction and escape of the thixotropic gelatinous silica, with consequent collapse of the overlying oxide laminae causing further breaches in a domino effect that may give rise to large volumes of chert-free BIF (Fig. 9A–D). Ageing converts hydrophilic colloidal silica to hydrophobic gel as it

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polymerises through syneresis and dehydration (Okamoto et al., 1957). Consequently, there is a narrow time slot for the reversible hydrolysis and thixotropy of colloidal silica during diagenetic lithification and various textures in the BIFs may be produced depending on the timing of reworking of the initial precipitate. The formation of chert bands by the interaction with silicasaturated seawater that takes places at the sediment– water interface is a separate process, which occurs during intervals of low sedimentation and typically shows irregular lower chert contacts (Krapez et al., 2003) and may replace various sediments. The silicified surfaces may be eroded by later density flows into shallow scours and rip-outs producing sheets, semisolidified fragments broken along syneresis cracks (Fig. 7B(c)), or gelatinous globules of chert (Fig. 7B(b)), depending on degree of ageing and syneresis of the surface. 4.4. The distribution of iron-formations in time BIFs (including GIFs) are global in distribution and range in age from 3.8 Ga to 1.8 Ga with apparently peak development at ca. 2.4 Ga (Beukes, 1973; Klein and Beukes, 1992; Abbott and Isley, 2001; Huston and Logan, 2004). The earliest known BIF are dated at 3.8 Ga and appear to be increasingly abundant from 3.4 Ga to peak at 2.5 Ga (James, 1983). They rapidly decrease in abundance from 2.2 Ga to 1.6 Ga, after which they virtually disappear from the geological record with the exception of the Neoproterozoic Rapitan type BIF (Gross, 1965; Young, 1976; Klein and Beukes, 1993; Klein and Ladeira, 2004) and the Paleozoic Lahn–Dill type iron-formations (Quade, 1976; Kräutner, 1977). The apparent increasing abundance of banded iron-formations through the Archean to the enormous deposits of the Lower Proterozoic and their sudden demise thereafter has commonly been attributed to a gradual rise in photosynthetic production of O2 increasingly precipitating oceanic ferrous iron. The onset of oxidative weathering eventually terminated the supply from subaerial weathering and precipitated all remaining ferrous iron in solution (Holland, 1984; Holland and Beukes, 1990; Gutzmer et al., 2003; Holland, 2006). Thus it appeared that the oceans were exhausted as a major resource of dissolved iron after 1.8 Ga. Kump and Seyfried (2005) suggested the predominance of iron sulfide precipitation in modern black smokers is the result of high sulfate concentrations in modern oceans and that low oceanic sulfate concentrations in the Archean as a result of the low atmospheric oxygen enabled the precipitation of oxide BIFs. However, there are abundant instances of iron hydroxide and

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silicate precipitation associated with modern underwater hydrothermal vents and iron sulphides are locally abundant in BIFs, nevertheless the increase in stratabound sulphide deposits in the Mesoproterozoic is significant. At the low oxygen fugacity postulated for the Archean and Paleoproterozoic eras (Holland, 2006), iron silicates and iron hydroxyoxides are precipitated but manganese typically remains in solution although local concentrations may occur in particular associated with carbonate precipitation (Gross, 1987; Anbar and Holland, 1992). Abundant sedimentary manganese oxides in the Middle Proterozoic Bangemall Group sediments of Western Australia probably mark a further stage in the evolution of O2 in the hydrosphere. Manganese and iron are precipitated together in modern oceans and Phanerozoic iron-formations typically contain abundant manganese (Table 2). Factual data on the distribution and abundance of BIFs (including cherty GIFs) are heavily biased by abundant data on North American occurrences, less data on the Hamersley Province of Western Australia and the Kaapvaal Craton of South Africa with isolated and/or imprecise data on most of the other occurrences worldwide (based on tables in Huston and Logan, 2004). This is particularly noticeable when the worldwide abundance of iron-formations in greenstone belts is considered. The stratigraphic correlation of the supracrustal rocks in the Murchison Province of Western Australia (Watkins and Hickman, 1990) has shown that instead of the greenstone belts being narrow troughs, they are remnants of widespread depositional systems that covered the whole province (Gole, 1979, cf. Condie, 1979), and probably the Southern Cross Province as well. Similar correlation of the greenstone belts in the Pilbara Craton (Trendall, 1983), and the similarity in ages of greenstone belts in the Superior Province (Huston and Logan, 2004), suggests that iron-formations in Archean greenstone belts everywhere were equally as extensive as those in the Paleoproterozoic basins. In considering the volume of Archean BIF, the whole area of the cratonic province should be considered instead of just the disjointed remnants in the greenstone belts. This does not exclude the possibility that some BIF formed in smaller basins but it would appear that the relative abundances of Archean and Paleoproterozoic BIFs is due more to preservation than initial volume. Evidence for primary hematite in BIFs (Ayres, 1972; Han, 1978) was based on partly weathered (leached and recemented with secondary silica) core from DDH 44 at Paraburdoo (Fig. 3) and is not found in completely unweathered samples (Lascelles, 2002, 2006, in press-a).

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After ca 2.4 Ga primary hematite appears to become of increasing importance, and magnetite is reportedly absent from the ca 0.8 Ga Rapitan and Urucum BIFs (Gross, 1965; Klein and Beukes, 1993; Klein and Ladeira, 2004), although it is abundant in the South Australian Razorback Ridge iron-formation of the same age (Whitten, 1970). The evidence of redbeds and palaeosols implies a marked change to oxygenic atmospheric and oceanic environments lacking in dissolved iron that is supported by the greatly reduced volume of iron-formations since the Lower Proterozoic (Holland and Beukes, 1990; Rye and Holland, 1998; Beukes et al., 2000; Canfield, 2005). Condie (1989) proposed a division of the evolution of the atmosphere into three stages. An initial reducing stage in which free oxygen occurs in neither the oceans or the atmosphere, an oxidizing stage in which small amounts of oxygen are present in the surface waters and the atmosphere but not in deep ocean waters and an aerobic stage in which free oxygen pervades the entire atmosphere and hydrosphere. This division would appear to match the mineralogy and distribution of BIF rather closely except for the magnetite-rich Razorback Ridge deposit in South Australia and reports of magnetite in Paleozoic BIFs (O'Rourke, 1961). The association of Neoproterozoic BIF with glacial sediments (Gross, 1965; Young, 1976; Klein and Beukes, 1992, 1993; Klein and Ladeira, 2004) suggests a possible origin due to reduced levels of biogenic oxygen combined with a high input of unweathered rock flour from glaciers that increased the levels of ferrous iron and silica in the oceans (Snowball Earth? Hoffman et al., 1998). 4.5. Redox state of original BIF sediments The Hamersley Group was gently folded and uplifted by the Ophthalmian event in the Lower Proterozoic and exposed to subaerial weathering with erosion products including microplaty hematite and high-grade ore occurring in the overlying Wyloo Group (Martin et al., 2000). Oxidative deep weathering produced microplaty hematite during the Proterozoic exposure (Lascelles, 2002) that first appears in the ca. 2030 Ga (Müller et al., 2005) Beasley River Quartzite and is abundant in the Mt. McGrath conglomerate, which also contains abundant clasts of high-grade hematite ore. Microplaty hematite is present in both cherty BIF and high-grade ore but is completely absent from BIFs that were not exposed to weathering at that time. Oxidized clasts are typically absent from Mt. Bruce Supergroup and older sediments of the Pilbara Region, although Beukes et al. (2000) argue for oxic atmospheres as far back as

2.45 Ga. BIF fragments derived from Archean greenstone belts occur within the Turee Creek Group at ca. 2440 Ga and euhedral grains of martite in Turee Creek Sandstone (Martin et al., 2000), develop perfect trellis texture indicating that they were deposited as magnetite and only oxidized under the current weathering environment. Pseudomorphs of minerals typically consist of finegrained aggregates, especially when formed at comparatively low temperatures such as supergene weathering and lower greenschist facies metamorphism. Pseudomorphs of magnetite, consisting of a mesh of fine crystallites (Fig. 10A), and oxidized euhedral siderite crystals are transformed to a mosaic of anhedral grains of hematite during metamorphism (Fig. 10B). Since the rate of reaction is proportional to the surface area of the reactants and inversely to the degree of perfection of the crystal lattice (Lepp, 1957: Langmuir, 1971), magnetite pseudomorphs, as well as colloform and other finegrained aggregates, oxidize more rapidly than wellcrystallized magnetite. Fine-grained primary laminae, inclusions and pseudomorphs of microplaty hematite (Fig. 10C) are thus the first to be oxidized by chemical weathering. Only well-crystallized magnetite grains develop the martite trellis texture (Fig. 10A) during subsequent oxidation and previously oxidized magnetite retains the texture of interlocking hematite grains (Fig. 10D). Fine-grained oxide microlaminae in BIF commonly show overgrowths of crystallized magnetite on lathshaped, bladed or hexagonal cores (Fig. 10E) that form internal hematite laminae and inclusions when partially oxidized (Han, 1978; Morris, 1980; Han, 1988; Lascelles, in press-a). Han (1988) documented cores and inclusions of hematite in magnetite grains from South Africa and North America, as well as from Western Australia, and from this it was inferred that the microlaminae consisted of hematite prior to metamorphic reduction. How much was originally hematite is debatable, since the wedgeshaped and rhombohedral examples are very similar to the bladed siderite crystals and microplaty hematite from Mt. Tom Price, Western Australia and the colloform, botryoidal and anhedral examples could equally be derived from poorly crystallized magnetite and lath-shaped inclusions from fragmented microlaminae. Very rarely, traces of martite trellis texture can be observed in the internal laminae at Hope Downs (Lascelles, in press-a). The oxide laminae in unweathered BIFs from the Hamersley Group (Brockman and Marra Mamba Iron formations) never show hematite overgrowths but always magnetite, which suggests that the primary iron hydroxide precipitates were incompletely oxidized

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Fig. 10. A. Well crystallized magnetite oxidized to martite and partly hydrated showing trellis pattern of hematite crystallites, Marra Mamba Iron Formation, Hope Downs (reflected light); B. euhedral siderite grains pseudomorphed by magnetite and partly oxidized showing aggregate of anhedral hematite and kenomagnetite grains, Mt. Tom Price, (plane polarized transmitted and reflected light); C. microplaty hematite along fractures and bedding planes in partly weathered BIF, Mt. Tom Price (plane polarized reflected light); D. mozaic textured hematite after magnetite (upper left clast) and microplaty hematite (lower right) in high-grade hematite clasts, Mt. McGrath Formation, Paraburdoo (reflected light with crossed polarizers); E. well-crystallized magnetite overgrowths on cryptocrystalline magnetite laminae, oxidized to martite and mozaic hematite respectively, Hope Downs (reflected light). F. unweathered magnetite, Mt. Tom Price (plane polarized transmitted and reflected light). Ch = chert; Go = goethite; He = hematite; Km = kenomagnetite, Lo = limonite; mpl = microplaty hematite; Mr = martite; Mt = magnetite; Py = pyrite; Vo = void. 3+ [(Fex2+, Fe1−x ) OH3−x, nH2O)]. Furthermore magnetite overgrowths occur on primary iron oxide laminae in both cherty and chert-free BIF but are not found on hematite in the McGrath Formation (Wyloo Group) that was

metamorphosed along with the Hamersley Group during the Capricorn Orogeny. The hematite formed during the Proterozoic subaerial weathering was converted to magnetite by metamorphism

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(Frost, 1979; Ohmoto, 2003b) during the Capricorn orogeny and all hematite currently present [with the possible exception of sub-microscopic hematite inclusions in microquartz at Mt. Gibson (Lascelles, 2006)] is due to the oxidation of metamorphic magnetite. Unweathered BIF from deep diamond drill cores at Mt. Gibson (Lascelles, 2006), Mt. Tom Price and Hope Downs (Lascelles, in press-a) contains no hematite at all (Fig. 10F), although magnetite pseudomorphs of rhomboidal hematite in hydrothermal veins at Mt. Gibson and microplaty hematite at Mt. Tom Price have been observed. Inclusions of hematite in partially oxidized magnetite are not evidence of primary hematite (Han, 1978; Morris, 1980; Han, 1988) if they are not present in completely unweathered BIF from the same area. Most jasper deposits are formed by the supergene oxidation of magnetite, siderite or pyrite in chert, however the submicroscopic hematite inclusions imparting the red colour to some unweathered microquartz bands at Mt. Gibson (Fig. 1A) are believed to be residual from traces of primary Fe3+ present in the initial silicate minerals and, unlike the abundant interstitial magnetite in the bands, were preserved from reduction during metamorphism by being completely enclosed in quartz. The question of the primary redox state of the original BIF sediments remains unresolved but in the author's opinion it is most probable that ferrous iron was abundant in both the silicates and the hydroxides. 4.6. Lake superior or Algoma type iron-formations The classification of BIFs into Lake Superior and Algoma types originated in North America (Gross, 1983). Its purpose was to distinguish the widespread Proterozoic, apparently shallow water, largely granular and oolitic iron-formations grading into laminated ironformations that were deposited in basins surrounding the Ungava Craton (Lake Superior), from the apparently deeper water Archean laminated BIFs deposited in greenstone belts within the cratonic areas (Algoma). Algoma type BIFs are typically highly deformed and metamorphosed, forming narrow lenticular and discontinuous outcrops associated with pillow lavas and greywacke whereas the Lake Superior type is typically relatively unmetamorphosed and undeformed, with widespread continuous outcrop and associated with conglomerate, quartzite and stromatolitic carbonate rocks. The widespread correlation of the Luke Creek Group within the Murchison Province of the Yilgarn Craton (Watkins and Hickman, 1990) and of greenstone belts in other cratons suggests that greenstone belts were

initially parts of widespread basins, with an original extent of the same order or greater than the Hamersley Province, that have been disrupted by large scale tectonic movements and granitic intrusions. The limited areal extent and variable thickness of the BIF units is typically due to tectonic thickening and thinning by isoclinal folding and faulting than by primary deposition and at least some Algoma type BIFs were originally as extensive as the Superior type (Condie, 1979; Gole, 1981; Trendall, 1983; Watkins and Hickman, 1990; Lascelles, 2006). The apparently smaller volume of iron-formations in the Archean may be related to the lesser preservation of Archean supracrustal rocks and not necessarily indicate smaller deposits although few if any reach the thickness of the Marra Mamba and Brockman Iron Formations. The Hamersley Group BIF is commonly classified as a Lake Superior type BIF due to its wide areal extent, gentle deformation and supposedly comparatively shallow deposition (Morris and Horwitz, 1983). Granular and oolitic iron-formations are, however, absent and recent studies have demonstrated volcanic and turbidite associations (Barley et al., 1997; Krapez et al., 2003), suggesting deposition at greater depth. The close similarity of banding and sedimentary structures in the BIF within the Archean greenstone belts of the Yilgarn and Pilbara Cratons to those of the Paleoproterozoic Hamersley Group indicates the same processes were involved in their deposition (Gole, 1981). GIFs are rarely found in greenstone belt iron-formations and furthermore, gradational lateral interfingering of Lake Superior GIF into BIF is reported from the Transvaal in South Africa, the Labrador Trough in Canada and the Frere Range in Western Australia. The Neoproterozoic GIFs show characteristics of Lake Superior type sedimentation yet are clearly interbedded with turbidites (Whitten, 1970; Klein and Beukes, 1992, 1993; Klein and Ladeira, 2004). 4.7. Silicate, carbonate and sulphide facies BIF Oxide facies BIF typically forms the largest and most widespread of BIF facies, but silicate and carbonate facies may be sufficiently extensive to be of economic significance, and if Sedex deposits are considered as sulphide facies iron-formation then their importance is greatly increased. Silicate facies commonly constitutes a minor proportion of BIFs where mesobands of undissociated Al-poor nontronite (minnesotaite) occur within oxide facies BIFs or may consist of bands of iron aluminosilicate-rich shale interbedded with chert and carbonate mesobands in Brockman Formation S bands

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or interbedded with GIF. Oxide laminae are typically rare in silicate facies BIF. Carbonate facies BIFs typically consisting of siderite, chert and minor silicate mesobands with very minor oxide laminae are found in Hamersley Group BIF in the MacLeod Member and in Dales Gorge Member S bands (Table 1). The carbonate is typically very fine-grained and apparently derived from sedimentary carbonate ooze that has reacted with Fe-rich solutions during transportation by turbidity and density currents (Lascelles, in press-a) but turbidity currents deposit some coarsegrained bands ranging from ferroan calcite to ankerite (Krapez et al., 2003). Coarser-grained granular to oolitic siderite interbedded with chert form significant parts of Lake Superior type BIFs (James, 1954; Beukes, 1984). Both types appear to be resedimented chemical precipitates. Black shales with pyrite and interbedded chert bands and nodules form the sulphide facies BIF of James (1954), but are typically of limited thickness and variable sulphide content, and thus rarely considered of economic interest for base metals. Commonly forming at the base of density current sequences, it would appear that flocculation of iron sulphide is more rapid than hydroxides and thus deposited along with detrital clay particles. The high carbon content could be produced from abundant growth of sulphur bacteria or the poisoning of pelagic biota by H2S. It is suggested that plumes of colloidal sulphides may be deposited with pelagic sediments during periods of slow sedimentation. Pyrite may be locally abundant in oxide facies BIFs but typically appears to have been introduced during diagenesis or metamorphism. 4.8. Origin of high-grade iron ore deposits As James (1954) observed, many BIFs contain parts that do not contain chert (iron-poor bands) and therefore do not fit in the current definition of BIF. This applies to some unweathered BIFs, as seen at Mt. Gibson and Koolyanobbing in the Yilgarn Craton and at Hope Downs in the Hamersley Province (Lascelles, 2002) (Fig. 3) and to all in situ high-grade iron ore deposits. The latter have been generally dismissed as supergene modifications of cherty BIFs, but recent study (Lascelles, 2002, 2006) has shown that chert-free BIF predates supergene processes and is a significant primary feature of BIFs. All BIFs are metamorphosed to the level that chert bands are converted to microquartz and although minor quartz may be leached by juvenile meteoric water from soil horizons above the water table (vertical flow), it is not leached during deep chemical

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weathering due to concentrations of dissolved amorphous silica in ground water (lateral flow) that are above the saturation level of quartz (Meinzer, 1949; Davis, 1964). In fact, iron oxides are leached from BIF during weathering and the upper parts of many BIF sequences consist solely of chert due to leaching of the iron. Many high-grade ore deposits extend to great depths below the water table with no evidence of vadose weathering and some deposits can be seen to pass into magnetite and unweathered chert-free BIF (Hashimoto Hill, Silvergrass, Giles Mini, Mt. Tom Price Southern Ridge and Mt. Gibson) or magnetite–carbonate BIF (Mt. Gibson, Koolyanobbing A deposit, and Mt. Tom Price North Deposit) at depth. Contacts between cherty BIF and chert-free BIF are typically sharp at all stages from completely unweathered BIF to surface outcrops and clearly the chert-free BIF is not due to more intensive weathering. The supergene leaching of quartz from cherty BIFs was an assumption that is totally unsupported by any physical or chemical evidence and is contrary to all experience of the chemical weathering of quartz-bearing rocks. Various hydrothermal (Roberts and Bartley, 1943; Dorr and Barbosa, 1963; Berge et al., 1977; Cabral et al., 2003) and metamorphic models (Li et al., 1993; Martin et al., 1998; Powell et al., 1999; Brown et al., 2004) have been proposed for selective leaching of quartz from BIFs based on the enhanced solubility of quartz at higher than ambient temperatures and pressures. However, hematite solubility is also enhanced and hematite veins (e.g., Paraburdoo; Fig. 4E) and replacement deposits of specular hematite (Francis Ck, Northern Territory, Australia; Roy Hill, W.A.; Fig. 3) clearly illustrate the mobility of hematite. In those deposits that show clear evidence of hydrothermal activity (Koolyanobbing, Mt. Tom Price, Mt. Gibson) little or no change in the relative proportions of the Fe and SiO2 is seen where both hematite and quartz are recrystallised with preservation of the original structure and there is no evidence for the selective removal of quartz. It is apparent that mobilization and recrystallization of the SiO2 and Fe has taken place simultaneously. To suggest that quartz can be selectively leached from BIFs on a giant scale by hydrothermal fluids that have left no other trace of their passage is ridiculous. The hydrothermal replacement of chert bands by carbonates produces a form of chert-free BIF (Fig. 4) that is up-graded by supergene leaching of the carbonates to leave a residue of iron oxides (Chown et al., 2000; Taylor et al., 2001; Dalstra et al., 2002; Lascelles, 2002; Dalstra and Guedes, 2004; Thorne et al., 2004;

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Lascelles, 2006, in press-b). In the carbonate replacement of the chert bands at Mt. Gibson, fine-grained magnetite and iron silicates are also replaced. However

the majority of deposits in Western Australia display original sedimentary textures identical to enclosing and adjacent cherty BIF, with no evidence of mobilization or

Fig. 11. Whole rock analyses of unweathered BIF (Mt. Gibson). A. Cherty BIF, B. silicate-rich BIF, C. chert-free BIF.

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recrystallization by hydrothermal activity (cf. Dorr and Barbosa, 1963; Rosière and Rios, 2004). Contacts between ore (chert-free BIF) and cherty BIF are sharp and typically subvertical and, after taking into account the greatly reduced silica due to the absence of chert bands, the chert-free BIF is chemically and texturally indistinguishable from typical cherty BIF (Fig. 11). Pressure solution during intense folding was suggested for the Krivoy Rog deposits in the Ukraine that form shoots of high-grade ore along fold axes (Belevtsev et al., 1983) but there is no evidence for such a process in the Hamersley Range deposits. Similar shoots of high-grade microplaty hematite ore that replace unfolded black shale and chert-free BIF in the eastern Chichester Range of the Hamersley Province (Hannon et al., 2005; Lascelles, 2005, Hancock Prospecting P/L, unpublished company report) appear to be of hydrothermal origin. An alternative mechanism of deformational partitioning of iron oxides and chert (Findlay, 1994; Hippertt et al., 2001) suggested that the high-grade ore deposits were enormous boudinage necks formed by pulling apart of chert bands during diagenesis and leaving the oxide bands to form the necks. The model is attractive, as it does not require any selective solution of quartz. Abundant evidence of boudinaged chert bands can be seen in cherty BIFs (Fig. 7C) ranging from the microscopic to mesoscopic associated with syneresis (podding) and slumping (intraformational folding) but evidence for metre scale and larger boudinage is absent and the boudinage horizons are typically restricted to a single density current sequence with unboudinaged flat lying BIF above and below. The origin of the chert-free BIFs is speculative but it is suggested that loss of the chert bands occurred during early diagenesis when the silica was still in a thixotropic colloidal state and the gelatinous chert bands were sufficiently mobile to escape through fractures in the iron oxide layers. Collapse of the overlying oxide layers into the voids left by escape of the chert bands both seals the ends of chert bands preventing further lateral loss (Fig. 1A, B) and by initiating further fractures above the escaping chert layer extends the chert-free BIF to the sediment/water interface (Fig. 9). 5. Conclusions By combining the low alumina content of BIFs compared to Phanerozoic oolitic iron-formations with the diagenetic instability of low-Al nontronite, a model is suggested (Table 2) for the genesis of Archean and Phanerozoic marine iron-formations and a diagenetic

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origin for the chert bands in BIFs that is based on the difference in chemistry between Archean and modern oceans. The new model for the genesis of banded ironformations proposes that hydrothermal solutions, containing abundant iron and silica derived from the reaction of subsurface water with hot volcanic rock, enter the oceans as “black smokers” at mid-ocean ridges, continental margins or other recently active volcanic sites. On cooling and dispersion, ferrous iron, in solution or as colloidal hydroxide particles, reacts with dissolved silica to form hydrous Al-poor iron-rich silicates, and together with excess iron precipitated as hydrous oxides form mounds around vent chimneys. According to the concentration of dissolved silica, iron, and available oxygen and sulphur, a variety of final products ensue. During the Archean and Paleoproterozoic eras, anoxic weathering prior to the rise of atmospheric oxygen introduced vast amounts of ferrous iron to the oceans; there to be precipitated as silicate, sulphide, and carbonate, but much would accumulate in solution. The input of iron (and silica) from hydrothermal vents supersaturated the oceans in their vicinity with respect to ferrous iron and thus triggered the deposition of the dissolved iron as diagenetically unstable Al-poor hydrous iron silicate and hydroxide to form the extensive deposits of the Archean and Paleoproterozoic eras. Manganese typically remained in solution with only minor amounts deposited locally. At high oxygen concentrations, such as exist in modern oceans; little or no iron in solution reaches the oceans from subaerial weathering and Phanerozoic iron formations are typically low volume deposits. High O2 fugacities precipitate iron and manganese in trivalent form close to the vent and Phanerozoic BIFs have comparatively high concentrations of manganese and may be associated with extensive manganese deposits. Deposition of iron-formations as colloidal iron silicate and hydroxide mud in mounds with abundant entrapped water leads to instability and debris flows that develop into turbidity currents. It is suggested that granular ironformations show characteristics of proximal deposition particulate loads by turbidity currents and density currents of suspended colloidal particles dispersed by the debris flow deposit finely laminated iron-formation over large areas of the ocean floor. Thus deposition by density currents may take place at a great distance from the source of the iron and silica (Fig. 5). Abundant evidence of turbidity current deposition is seen in adjacent and interbedded strata, and the uniformly fine grain, laminations and banding in typical oxide facies BIFs is characteristic of density flow deposition.

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Dissociation of the Al-poor hydrous iron silicates during diagenesis into iron oxides, colloidal silica and water gives rise to the bands of chert, while loss of thixotropic colloidal silica during early diagenesis produces volumes of chert-free BIF (Fig. 1) that forms the protore for high-grade iron ore deposits. Normal pelagic sedimentation of fine-grained aluminosilicate clay particles deposited interbedded shale bands but only minor traces of aluminosilicate grains were occluded in BIF from the contained seawater in the turbidity and density flows. Silicification of the surface of sediments commonly occurred during intervals of nondeposition to form chert bands at the top of density flow sequences. The sedimentology of BIFs makes sequential sense with interbedded sediments when considered as density current deposits and particularly when the association of GIF grading into BIF is considered. The close association of Neoproterozoic BIFs with glaciogene sediments may explain their anomalous time distribution since the silica and ferrous iron content of glaciofluval streams, and consequently seawater, would be high due to the suspended load of finely ground unweathered rock and the greatly reduced metabolic activity of silica secreting organisms due to the low water temperature. Proponents of the “Snowball Earth” theory further suggest greatly diminished biogenic processes and that photosynthesis and metabolic extraction of silica practically ceased at this time and that the pO2 of the oceans was greatly reduced. Modern hydrothermal vents on the sea floor (‘black smokers’) form mounds of iron hydroxide and hydrous silicates (Al-poor nontronite) from which debris flows and plumes of colloidal particles redeposit the mound material as transported sediments. To date, these sediments have not been penetrated to sufficient depth to show diagenetic changes. Phanerozoic oolitic iron-formations (Table 2) apparently form slowly during transgressive periods of low sediment input (Young and Taylor, 1989) as slowly accumulated pelagic deposits of suspended terrigenous fine quartz and aluminosilicate clay particles. The chamosite thus formed does not dissociate during diagenesis, and although chert is rare, low-Al highgrade hematite deposits do not form. The origin of the iron is debatable but it has been suggested that it was transported in the form of organic complexes, or perhaps from the interaction of seawater with buried sediments. Note that silica is not precipitated to form chert bands at any time (Table 2) although rare concretions may form. Although numerous details for specific iron-formations and facies remain to be elucidated, it is proposed that this model forms a uniformitarian explanation of BIFs

from the Archean to the present including Phanerozoic laminated iron-formations, granular iron formations and modern deep sea sedimentation. It provides a reasonable interpretation of the sedimentary environment in which BIFs were formed and may provide a useful basis for the interpretation of iron-formations and their associated stratigraphic sequences. Acknowledgements This study was undertaken as part of a PhD thesis on the origin of BIF and in situ derived iron ore deposits at the University of Western Australia. The assistance of Hamersley Iron Ltd. Hancock Prospecting Pty. Ltd., Hope Downs Management Services Ltd., Koolyanobbing Iron Pty. Ltd., Mt. Gibson Iron Pty. Ltd., and Portman Mining Ltd. in providing access to diamond core, company reports and records; and the constructive criticism by M.E. Barley and the journal reviewers is gratefully acknowledged. References Abbott, D., Isley, A., 2001. Oceanic upwelling and mantle-plume activity: Palaeomagnetic tests of ideas on the source of the Fe in early Precambrian iron formations. In: Ernst, R.E., Buchan, K.L. (Eds.), Mantle Plumes: Their Identification through Time: Boulder, ColoradoGeological Society of America Special Paper 352, 323–339. Anbar, A.D., Holland, H.D., 1992. The photochemistry of manganese and the origin of BIF. Geochimica et Cosmochimica Acta 56, 2595–2603. Appel, P.W.U., 1987. Geochemistry of the Early Archean Isua IronFormation, West Greenland. In: Appel, P.W.U., LaBerge, G.L. (Eds.), Precambrian Iron-Formations. Theophrastus Publications, Athens, pp. 31–68. Ayres, D.E., 1972. Genesis of iron-bearing minerals in banded iron formation mesobands in the Dales Gorge Member, Hamersley Group, Western Australia. Economic Geology 67, 1214–1233. Barghoorn, E.S., Tyler, S.A., 1965. Microorganisms from the Gunflint Chert. Science 147, 563–577. Barley, M.E., Pickard, A.L., Sylvester, P.J., 1997. Emplacement of a large igneous province as a possible cause of banded iron formation 2.45 billion years ago. Nature 385, 55–58. Barley, M.E., Pickard, A.L., Hagemann, S.G., Folkert, S.L., 1999. Hydrothermal origin for the 2 billion year old Mount Tom Price giant ore deposit, Hamersley Province, Western Australia. Mineralium Deposita 34, 784–789. Bayley, R.W., James, H.L., 1973. Precambrian iron-formations of the United States. Economic Geology 68, 934–959. Belevtsev, Y.N., Belevtsev, R.Y., Siroshtan, R.I., 1983. The Krivoy Rog Basin. In: Trendall, A.F., Morris, R.C. (Eds.), Iron Formation: Facts and Problems. Elsevier, Amsterdam, pp. 211–251. Berge, J.W., Johansson, K., Jack, J., 1977. Geology and origin of the hematite ores of the Nimba Range, Liberia. Economic Geology 72, 582–607. Berkner, L.V., Marshall, L.C., 1964. The history of growth of oxygen in the Earth's atmosphere. In: Brancazio, P.J., Cameron, A.G.W.

D.F. Lascelles / Ore Geology Reviews 32 (2007) 381–411 (Eds.), The Origin and Evolution of Atmospheres and Oceans. John Wiley and Sons, Inc., New York, pp. 102–126. Berkner, L.V., Marshall, L.C., 1965. On the origin and rise of oxygen concentration in the earth's atmosphere. Journal of the Atmospheric Sciences 22, 225–261. Beukes, N.J., 1973. Precambrian iron-formations of Southern Africa. Economic Geology 68, 960–1004. Beukes, N.J., 1980. Lithofacies and stratigraphy of the Kuruman and Griquatown iron-formations, Northern Cape Province, South Africa. Transactions of the Geological Society of South Africa 83, 69–86. Beukes, N.J., 1984. Sedimentology of the Kuruman and Griquatown iron-formations, Transvaal Supergroup, Griqualand West, South Africa. Precambrian Research 24, 47–84. Beukes, N.J., Klein, C., 1990. Geochemistry and sedimentology of a facies transition – from microbanded to granular iron-formation – in the early Proterozoic Transvaal Supergroup, South Africa. Precambrian Research 47, 99–139. Beukes, N.J., Klein, C., 1992. Models for iron-formation deposition. In: Schopf, J.W., Klein, C. (Eds.), The Proterozoic Biosphere: A Multidisciplinary Study. Cambridge University Press, pp. 147–152. Beukes, N.J., Gutzmer, J., Dorland, H., Szabo, A., 2000. Paleoproterozoic Laterites and Iron Ores: International Geological Congress 2000, Rio de Janeiro, Brazil. Abstract volume. Beukes, N.J., Gutzmer, J., Mukhopadhyay, J., 2002. The geology and genesis of high-grade hematite iron ore deposits. Proceedings Iron Ore 2002. Australasian Institute for Mining and Metallurgy, Perth, pp. 23–29. Reprinted in Transactions of the Institute of Mining and Metallurgy, Section B: Applied Earth Science, 112. p. 18–25. Binns, R.A., 2003. The PACMANUS hydrothermal field, Eastern Manus Basin, Papua New Guinea: a decade of seafloor investigation and the first deep drilling of an active, felsic-hosted, submarine hydrothermal field. In: Yeats, C.J. (Ed.), Seabed Hydrothermal Systems of the Western Pacific: Current Research and New Directions, June 2003. . CSIRO Division of Exploration and Mining Report, vol. 1112F. CSIRO Division of Exploration and Mining, Sydney. Bischoff, J.L., 1972. A ferroan nontronite from the Red Sea geothermal system. Clays and Clay Minerals 20, 217–223. Bischoff, J., Dickson, F.W., 1975. Seawater–basalt interaction at 200 °C and 500 bars: implications as to the origin of sea floor heavy metal deposits and regulation of seawater chemistry. Earth and Planetary Science Letters 25, 385–397. Boström, K., Peterson, M.N.A., Joensuu, O., Fisher, D.E., 1969. Aluminum-poor ferromanganoan sediments on active oceanic ridges. Journal of Geophysical Research 74, 3261–3270. Bouma, A.H., 1962. Sedimentology of Some Flysch Deposits. Elsevier Publishing Company, Amsterdam. 168 pp. Bouma, A.H., 2000. Fine-grained mud-rich turbidite systems; model and comparison with coarse-grained sand-rich systems. In: Bouma, A.H., Stone, C.G. (Eds.), Fine-Grained Turbidite Systems. American Association of Petroleum Geologists Memoir, vol. 72, pp. 9–19. Braterman, P.S., Cairns-Smith, A.G., 1987. Iron photoprecipitation and genesis of the banded iron-formations. In: Appel, P.W.U., LaBerge, G.L. (Eds.), Precambrian Iron-Formations. Theophrastus, Athens, pp. 215–245. Brinkman, R.T., 1969. Dissociation of water vapor and evolution of oxygen in the terrestrial atmosphere. Journal of Geophysical Research 74, 5355–5368. Brocks, J.J., Buick, R., Summons, R.E., Logan, G.A., 2003. A reconstruction of Archean biological diversity based on molecular fossils from the 2.78 to 2.45 billion-year-old Mount Bruce

407

Supergroup, Hamersley Basin, Western Australia. Geochimica et Cosmochimica Acta 67, 4321–4355. Brown, M.C., Oliver, N.H.S., Dickens, G.R., 2004. Veins and hydrothermal fluid flow in the Mt Whaleback Iron Ore District, eastern Hamersley Province, Western Australia. Precambrian Research 128, 441–474. Button, A., 1976. Transvaal and Hamersley basins—review of basin development and mineral deposits. Minerals Science and Engineering 8, 262–293. Butuzova, G.Y., 1966. Iron-ore sediments of the fumarole field of the Santorrini volcano, their compositions and origin: Doklady Academi Science, U.S.S.R. Earth Science Sections 168, 215–217. Butuzova, G.Y., 1968. Recent volcano-sedimentary iron-ore process in the Santorini volcano caldera (Aegean Sea) and its effect on the geochemistry of sediments: Academy Science, U.S.S.R. Geological Institute Transactions 194, 215–217. Butuzova, G.Y., Drits, V.A., Morozov, A.A., Gorschkov, A.I., 1990. Processes of formation of iron–manganese oxyhydroxides in the Atlantis-II and Thetis Deeps of the Red Sea. Sediment-hosted Mineral Deposits. Special Publication of the International Association for Sedimentology, vol. 11, pp. 57–72. Cabral, A.R., Filho, R., Jones, R.D., 2003. Hydrothermal origin of soft hematite ore in the Quadrilátero Ferrífero of Minas Gerais, Brazil: petrographic evidence from the Gongo Soco iron ore deposit. Transactions of the Institution of Mining and Metallurgy (Section B, Applied Earth Science) 112, 279–285. Caldwell, O.G., Marshall, C.E., 1942. A study of some chemical and physical properties of the clay minerals nontronite, attapulgite and saponite. Missouri University College Agricultural Research Bulletin, vol. 354. 51 pp. Canfield, D.E., 2005. The early history of atmospheric oxygen: homage to Robert M. Garrels. Annual Review of Earth and Planetary Sciences 33, 1–36. Chown, E.H, N'Dah, E., Mueller, W.U., 2000. The relation between iron-formation and low temperature hydrothermal alteration in an Archean volcanic environment. Precambrian Research 101, 263–275. Chukhrov, F.V., 1973. On the genesis problem of thermal sedimentary iron ore deposits. Mineralium Deposita 8, 138–147. Cloud, P., 1973. Paleoecological significance of the banded iron formation. Economic Geology 68, 1135–1143. Cloud, P., 1983. Banded iron-formation — a gradualist's dilemma. In: Trendall, A.F., Morris, R.C. (Eds.), Banded Iron-Formation: Facts and Problems. Elsevier, Amsterdam, pp. 401–416. Condie, K.C., 1979. Archean greenstone belts. Developments in Precambrian Geology, vol. 3. Elsevier, New York. 434 pp. Condie, K.C., 1989. Plate Tectonics and Crustal Evolution. Pergamon Press, Oxford. 476 pp. Dalstra, H., Guedes, S., 2004. Giant hydrothermal hematite deposits with Mg–Fe metasomatism: a comparison of the Carajas, Hamersley, and other iron ores. Economic Geology 99, 1793–1800. Dalstra, H., Harding, T., Riggs, T., Taylor, D., 2002. Banded ironformation hosted high-grade hematite deposits, A coherent group? Proceedings Iron Ore 2002. Australasian Institute for Mining and Metallurgy, Perth, pp. 57–62. Davis, S.N., 1964. Silica in streams and groundwater. American Journal of Science 262, 870–871. Degens, E.T., Stoffers, P., 1976. Stratified waters as a key to the past. Nature 263, 22–27. Derry, L.A., Jacobsen, S.B., 1990. The chemical evolution of Precambrian seawater: evidence from REEs in banded iron formations. Geochimica et Cosmochimica Acta 54, 2965–2977.

408

D.F. Lascelles / Ore Geology Reviews 32 (2007) 381–411

Dorr, J.v.N., Barbosa, A.L.M., 1963. Geology and ore deposits of the Itabira district, Minas Gerais, Brazil. U.S. Geological Survey Professional Paper 341C (110 pp.). Drever, J.I., 1974. Geochemical model for the origin of Precambrian banded iron formations. Geological Society of America Bulletin 35, 1099–1106. Drever, J.I., 1982. The Geochemistry of Natural Waters. Prentice– Hall, Inc. 388 pp. Edmond, J.M., Measures, C., Mangum, B., Grant, B., Sclater, F.R., Collier, R., Hudson, A., Gordon, L.I., Corliss, J.B., 1979. On the formation of metal-rich deposits at Ridge crests. Earth and Planetary Science Letters 46, 19–30. England, G.L., Rasmussen, B., Krapez, B., Groves, D.I., 2002. Palaeoenvironmental significance of rounded pyrite in siliciclastic sequences of the Late Archean Witwatersrand Basin: oxygendeficient atmosphere or hydrothermal alteration? Sedimentology 49, 1133–1156. Ewers, W.E., 1983. Chemical factors in the deposition and diagenesis of banded iron-formation. In: Trendall, A.F., Morris, R.C. (Eds.), Iron-Formation: Facts and Problems. Elseviers Science Publishers, pp. 491–512. Ewers, W.E., Morris, R.C., 1981. Studies of the Dales Gorge Member of the Brockman Iron Formation, Western Australia. Economic Geology 76, 1929–1953. Findlay, D., 1994. Diagenetic boudinage, an analogue model for the control on hematite enrichment iron ores of the Hamersley Iron Province of Western Australia, and a comparison with Krivoi Rog of Ukraine, and Nimba Range, Liberia. Ore Geology Reviews 9, 311–324. Fortin, D., Langley, S., 2005. Formation and occurrence of biogenic iron-rich minerals. Earth-Science Reviews 72, 1–19. Fralick, P., 1987. Depositional environment of Archean iron-formation: inferences from layering in sediment and volcanic hosted end members. In: Appel, P.W.U., LaBerge, G.L. (Eds.), Precambrian Iron-Formations. Theophrastus, Athens, pp. 251–266. Frimmel, H.E., 2005. Archean atmospheric evolution: evidence from the Witwatersrand gold fields, South Africa. Earth-Science Reviews 70, 1–46. Frost, B.R., 1979. Metamorphism of iron-formation: paragenesis in the system Fe–Si–C–O–H. Economic Geology 74, 775–785. Fryer, B.J., 1977. Rare earth evidence in iron formation for changing Precambrian oxidation states. Geochimica et Cosmochimica Acta 41, 361–367. Garrels, R.M., 1987. A model for the deposition of the microbanded Precambrian iron formations. American Journal of Science 287, 81–106. Garrels, R.M., Perry Jr., E.A., Mackenzie, F.T., 1973. Genesis of Precambrian iron-formations and the development of atmospheric oxygen. Economic Geology 68, 1173–1179. Gole, M.J., 1979. Metamorphosed banded iron-formations in the Archean Yilgarn Block, Western Australia. Unpublished PhD thesis, University of Western Australia. 255 pp. Gole, M.J., 1981. Archean banded iron-formations, Yilgarn Block, Western Australia. Economic Geology 76, 1954–1974. Goode, A.D.T., Hall, W.D.M., Bunting, J.A., 1983. The Nabberu Basin of Western Australia. In: Trendall, A.F., Morris, R.C. (Eds.), Iron-Formations: Facts and Problems. Elsevier, Amsterdam, pp. 295–323. Graf, J.L., 1978. Rare earth elements, iron formations and seawater. Geochimica et Cosmochimica Acta 42, 1845–1850. Grandstaff, D.E., 1980. Origin of uraniferous conglomerates at Elliot Lake, Canada, and Witwatersrand, South Africa: implications for

oxygen in the Precambrian atmosphere. Precambrian Research 13, 1–26. Griffin, A.C., 1980. Structural geology and sites of iron ore deposition in Koolyanobbing Range, Western Australia. Second International Archean Symposium, Archean Geochemistry Project, Perth, pp. 64–65. Gross, G.A., 1965. Geology of iron deposits in Canada, Vol. 1. General geology and evaluation of iron deposits. Canadian Geological Survey, Economic Geology Report. Gross, G.A., 1983. Tectonic systems and the deposition of ironformation. Precambrian Research 20, 171–187. Gross, G.A., 1987. Mineral deposits on the deep seabed. Marine Mining 6, 109–119. Gross, G.A., 1991. Genetic concepts for iron formation and associated metalliferous sediments. Economic Geology Monographs 8, 51–81. Gutzmer, J., Schneiderhan, E., Beukes, N.J., Strauss, H., Mezger, K., 2003. Unravelling chemostratigraphic signatures of sedimentation and diagenesis in Palaeoproterozoic iron and manganese formations. Transactions of the Institute of Mining and Metallurgy (Section B, Applied Earth Science) 112, 159–161. Hagemann, S.G. University of Western Australia, pers comm, April 2006. Hall, W.D.M., Goode, A.D.T., 1978. The early Proterozoic Nabberu basin and associated iron formations of Western Australia. Precambrian Research 7, 129–184. Hampton, M.A., 1972. The role of subaqueous debris flows in generating turbidity currents. Journal of Sedimentary Petrology 42, 775–793. Han, T.-M., 1978. Microstructures of magnetite as guides to its origin in some Precambrian iron-formations. Fortschritte der Mineralogie 56, 105–142. Han, T.-M., 1988. Origin of magnetite in Precambrian iron-formations of low metamorphic grade. In: Zachrisson, E. (Ed.), Proceedings of the Seventh Quadrennial IAGOD Symposium, E. Schweizerbart'sche Verlagsbuchhandlung (Nagele u. Obermiller), Stuttgart, pp. 641–656. Hannon, E., Kepert, D.A., Clark, D., 2005. From target generation to two billion tonnes in 18 months — the re-invention of the Chichester Range. Iron Ore Conference 2005. The Australasian Institute of Mining and Metallurgy Publication Series No 8/2005, pp. 73–77. Hartman, H., 1984. The evolution of photosynthesis and microbial mats: a speculation on the banded iron formations. In: Cohen, Y., Castenhole, R.W., Halverson, H.O. (Eds.), Microbial Mats: Stromatolites. Alan R. Liss, Inc., New York, pp. 449–453. Heinen, W., Oehler, J.H., 1979. Evolutionary aspects of biological involvement in the cycling of silica. In: Trudinger, P.A., Swaine, D.J. (Eds.), Biochemical Cycling of Mineral-Forming Elements. Elsevier, Amsterdam, pp. 431–465. Hesse, R., Chough, S.K., 1980. The northwest Atlantic mid-oceanic channel of the Labrador Sea: 11. Deposition of parallel laminated levee-muds from the viscous sublayer of low-density turbidity currents. Sedimentology 27, 697–711. Hiemenz, P.C., 1997. Principles of Colloid and Surface Chemistry, 3rd edition. Marcel Dekker, New York. 650 pp. Hippertt, J., Lana, C., Takeshita, T., 2001. Deformation partitioning during folding of banded iron formation. Journal of Structural Geology 23, 819–834. Hoffman, P.F., Kaufman, A.J., Halverson, G.P., Schrag, D.P., 1998. A Neoproterozoic snowball Earth. Science 281, 1342–1346. Holland, H.D., 1973. The oceans: a possible source of iron in ironformations. Economic Geology 68, 1169–1172.

D.F. Lascelles / Ore Geology Reviews 32 (2007) 381–411 Holland, H.D., 1984. The Chemical Evolution of the Atmosphere and Oceans. Princeton University Press, Princeton, New Jersey. 582 pp. Holland, H.D., 2002. Volcanic gases, black smokers and the great oxidation event. Geochimica et Cosmochimica Acta 66, 3811–3826. Holland, H.D., 2006. The oxygenation of the atmosphere and oceans. Philosophical Transactions of the Royal Society B 361, 903–915. Holland, H.D., Beukes, N.J., 1990. A palaeoweathering profile from Griqualand West, South Africa: evidence for a dramatic rise in atmospheric oxygen between 2.2 and 1.9 BP. American Journal of Science 290-A, 1–34. Honnorez, J., Von Herzen, R.P., Barrett, T.J., Becker, K., Bender, M.L., Borella, P.E., Hubberten, H.-W., Jones, S.C., Karato, S., Laverne, C., Levi, S., Migdisov, A.A., Moorby, S.A., Schrader, E.L., 1981. Hydrothermal mounds and young ocean crust of the Galapagos: preliminary deep sea drilling results, Leg 70. Geological Society of America Bulletin 92, 457–472. Huston, D.L., Logan, G.A., 2004. Barite, BIFs and bugs: evidence for the evolution of the Earth's early hydrosphere. Earth and Planetary Science Letters 220, 41–55. Hyde, R.S., 1980. Sedimentary facies in the Archean Timiskaming Group and their tectonic implications, Abititbi Greenstone Belt, Northeastern Ontario, Canada. Precambrian Research 12, 161–195. Isley, A.E., 1995. Hydrothermal plumes and the delivery of iron to banded iron formation. Journal of Geology 103, 169–185. Isley, A.E., Abbott, D.H., 1999. Plume-related mafic volcanism and the deposition of iron-formation. Journal of Geophysical Research 104, 15,461–15,477. James, H.L., 1954. Sedimentary facies of iron formation. Economic Geology 49, 235–293. James, H.L., 1966. Chemistry of the iron-rich sedimentary rocks. United States Geological Survey Professional Paper 440-W (60 pp.). James, H.L., 1983. Distribution of banded iron-formation in space and time. In: Trendall, A.F., Morris, R.C. (Eds.), Banded Iron-Formation: Facts and Problems. Elsevier, Amsterdam, pp. 471–490. James, H.L., 1992. Precambrian iron-formations: Nature, origin, and metamorphic evolution from sedimentation to metamorphism. In: Wolf, K.H., Chilingarian, G.V. (Eds.), Developments in Sedimentology 47. Diagenesis, vol. 111. Elsevier Science Publishers, Amsterdam, pp. 543–589. Karl, D.M., McMurtry, G.M., Malhoff, A., Garcia, M.O., 1988. Loihi Seamount, Hawaii: a mid-plate volcano with a distinctive hydrothermal system. Nature 335, 532–535. Kasting, J.F., 1993. Earth's early atmosphere. Science 259, 920–926. Kelley, D.S., Karson, J.A., Blackman, D.K., Fruh-Green, G.L., Butterfield, D.A., Lilley, M.D., Olson, E.J., Schrenk, M.O., Roek, K.K., Lebonk, G.T., Rivizzigno, P., 2001. An off-axis hydrothermal vent field near the Mid-Atlantic Ridge at 30°8′N. Nature 412, 145–149. Kimberley, M.M., 1978. Paleoenvironmental classification of iron formations. Economic Geology 73, 215–229. Kimberley, M.M., 1989. Exhalative origin of iron-formations. Ore Geology Reviews 5, 13–145. Kimberley, M.M., Tanaka, R.T., Farr, M.R., 1980. Composition of Middle Precambrian uraniferous conglomerate in the Elliot Lake– Agnew Lake area of Canada. Precambrian Research 12, 375–392. Klein, C., 1983. Diagenesis and metamorphism of Precambrian banded iron-formations. In: Trendall, A.F., Morris, R.C. (Eds.), Iron-Formations: Facts and Problems. Elsevier, Amsterdam, pp. 417–469.

409

Klein, C., Beukes, N.J., 1992. Proterozoic iron-formations. In: Conde, K.C. (Ed.), Proterozoic Crustal Evolution. Developments in Precambrian Geology, vol. 10. Elsevier Science Publishers, Amsterdam, pp. 383–418. Klein, C., Beukes, N.J., 1993. Sedimentology and geochemistry of the glaciogenic Late Proterozoic Rapitan iron-formation in Canada. Economic Geology 88, 542–565. Klein, C., Ladeira, E.A., 2004. Geochemistry and mineralogy of Neoproterozoic banded iron-formations and some selected siliceous manganese formations from the Urucum District, Matto Grosso do Sul, Brazil. Economic Geology 99, 1233–1244. Konhauser, K.O., Hamade, T., Raiswell, R., Morris, R.C., Ferris, F.G., Southam, G., Canfield, D.E., 2002. Could bacteria have formed the Precambrian banded iron formations? Geology 30, 1079–1082. Krapez, B., Martin, D. McB., 1999. Sequence stratigraphy of the Palaeoproterozoic Nabberu Province of Western Australia. Australian Journal of Earth Sciences 46, 89–103. Krapez, B., Barley, M.E., Pickard, A.L., 2003. Hydrothermal and resedimented origins of the precursor sediments to banded iron formation: sedimentological evidence from the Early Palaeoproterozoic Brockman Supersequence of Western Australia. Sedimentology 50, 979–1011. Kräutner, H.G., 1977. Hydrothermal–sedimentary iron ores related to submarine volcanic rises. In: Klemm, D.D., Schneider, H.J. (Eds.), Time and Stratabound Ore Deposits. Springer Verlag, Berlin, pp. 232–253. Kump, L.R., Holland, H.D., 1992. Iron in Precambrian rocks: implications for the global oxygen budget of the ancient Earth. Geochimica et Cosmochimica Acta 56, 3217–3223. Kump, L.R., Seyfried Jr., W.E., 2005. Hydrothermal fluxes during the Precambrian: effect of low oceanic sulfate concentrations and low hydrostatic pressure on the composition of black smokers. Earth and Planetary Science Letters 235, 654–662. Kump, L.R., Kasting, J.F., Barley, M.E., 2001. Rise of atmospheric oxygen and the “upside-down” Archean mantle. Geochemistry, Geophysics, Geosystems 2 (Paper number 2000GC000114). Kuo, P.H., Holland, H.D., Danielson, A., 1993. Late Archean atmospheric CO2. Annual Meeting of Geological Society of America, p. A89. Abstract volume. La Berge, G.L., 1966. Altered pyroclastic rocks in iron formation in the Hamersley Range. Western Australia. Economic Geology 61, 147–161. La Berge, G.L., 1973. Possible biological origin of Precambrian ironformations. Economic Geology 68, 1098–1109. Langmuir, D., 1971. Particle size effect on the reaction goethite = hematite + water. American Journal of Science 271, 147–156. Lascelles, D.F., 1983. Inliers of banded iron-formation in the Proterozoic Wyloo Group sediments near Paraburdoo, Western Australia. Journal of the Geological Society of Australia 30, 161–165. Lascelles, D.F., 2000. Marra Mamba Iron Formation stratigraphy in the eastern Chichester Range, Western Australia. Australian Journal of Earth Sciences 47, 799–806. Lascelles, D.F., 2001. The Ferro Gully North Mine, Hamersley Province, Western Australia. Australasian Institute of Mining and Metallurgy Proceedings, vol. 306, pp. 47–52. Lascelles, D.F., 2002. A new look at old rocks: a non-supergene origin for BIF-derived in situ high-grade iron ore deposits. Proceedings Iron Ore 2002. Australasian Institute for Mining and Metallurgy, Perth, pp. 107–126. Lascelles, D.F., 2005. Roy Hill Project, Chichester Range, W.A.: Hancock Prospecting P/L, unpublished company report.

410

D.F. Lascelles / Ore Geology Reviews 32 (2007) 381–411

Lascelles, D.F., 2006. The Mt Gibson banded iron-formation hosted magnetite deposit: two distinct processes for the origin of highgrade iron ore deposits. Economic Geology 101, 651–666. Lascelles, D.F., in press-a. The genesis of the Hope Downs iron ore deposit, Hamersley Province, Western Australia. Economic Geology. Lascelles, D.F., in press-b. Genesis of the Koolyanobbing iron ore deposits, Yilgarn Province, Western Australia: Applied Earth Science, Transactions of the Institution of Mining and Metallurgy. Lepp, H., 1957. Stages in the oxidation of magnetite. American Mineralogist 42, 679–681. Lepp, H., Goldich, S.S., 1964. Origin of Precambrian iron-formations. Economic Geology 59, 1025–1060. Levell, B.K., 1980. Evidence for currents associated with waves in Late Precambrian shelf deposits from Finnmark, North Norway. Sedimentology 27, 153–166. Li, Z.X., Powell, C.M., Bowman, R., 1993. Timing and genesis of Hamersley iron-ore deposits. Exploration Geophysics 24, 631–636. Lonsdale, P., 1976. Abyssal circulation of the south-east Pacific and some geological interpretations. Journal of Geophysical Research 81, 1163–1176. Loughnan, F.C., 1969. Chemical Weathering of the Silicate Minerals. American Elsevier Publishing Co. Inc., New York. 154 pp. Lupton, J.E., 1996. A far-field hydrothermal plume from Loihi Seamount. Science 272, 976–979. Mackay, D.J., OSullivan O'Sullivan, J.E., Watson, R.J., 2002. Iron in the western Pacific. A riverine or hydrothermal source for iron in the equatorial undercurrent. Deep Sea Research (Part 1) 49, 887–993. Maliva, R.G., Knoll, A.H., Simonson, B.M., 2005. Secular change in the Precambrian silica cycle: insights from chert petrology. Geological Society of America Bulletin 117, 835–845. Marshall, C.E., 1949. The Colloid Chemistry of the Silicate Minerals. Academic Press, New York. 195 pp. Martin, D. McB., Li, Z.X., Nemchin, A.A., Powell, C.McA., 1998. A pre 2.2 Ga age for giant hematite ores of the Hamersley Province, Australia? Economic Geology 93, 1084–1090. Martin, D.M., Powell, C.M., George, A.D., 2000. Stratigraphic architecture and evolution of the early Paleoproterozoic McGrath Trough, Western Australia. Precambrian Research 99, 33–64. Mason, B., 1966. Principles of Geochemistry. John Wiley and Sons, Inc., New York. 329 pp. Maynard, J.B., 1983. Geochemistry of Sedimentary Ore Deposits. Springer Verlag, New York. 305 pp. McConchie, D., 1987. The geology and geochemistry of the Joffre and Whaleback Shale Members of the Brockman Iron Formation, Western Australia. In: Appel, P.W.U., La Berge, G.L. (Eds.), Precambrian Iron-Formations. Theophrastus Publications, Athens, pp. 541–597. McGuire, A.V., Dyar, M.D., Ward, K.A., 1989. Neglected Fe 3+/Fe2+ ratios — a study of Fe3+ content of megacrysts from alkali basalts. Geology 17, 687–690. Meinzer, O.E., 1949. Hydrology. Dover Publications Inc., New York. 712 pp. Miles, K.R., 1953. Koolyanobbing iron ore. In: Edwards, A.B. (Ed.), The Geology of Australian Ore Deposits, 5th Empire Mining and Metallurgical Congress, Melbourne. The Australasian Institute of Mining and Metallurgy, pp. 172–176. Miyano, T., Beukes, N.J., 1984. Phase relations of stilpnomelane, ferriannite and riebeckite in very low-grade metamorphosed iron-formations. Transactions of the Geological Society of South Africa 87, 111–124. Moore, E.S., Maynard, J.E., 1929. Solution, transportation and precipitation of iron and silica. Economic Geology 24, 272–303, 365–402 and 506–527.

Morris, R.C., 1980. A textural and mineralogical study of the relationship of iron ore to banded iron-formation in the Hamersley Iron Province of Western Australia. Economic Geology 75, 184–209. Morris, R.C., 1993. Genetic modeling for banded iron-formation of the Hamersley Group, Pilbara Craton, Western Australia. Precambrian Research 60, 243–286. Morris, R.C., Horwitz, R.C., 1983. The origin of the iron-formationrich Hamersley Group of Western Australia — deposition on a platform. Precambrian Research 21, 273–297. Müller, G., Förstner, U., 1973. Recent iron ore formation in Lake Malawi, Africa. Mineralium Deposita 8, 278–290. Müller, S.G., Krapez, B., Barley, M.E., Fletcher, I.R., 2005. Giant iron-ore deposits of the Hamersley Province related to the breakup of Paleoproterozoic Australia: new insights from in situ SHRIMP dating of baddeleyite from mafic intrusions. Geology 33, 577–580. Ohmoto, H., 1993. The banded iron-formations in the Hamersley Basin: products of the oxygen-rich Archean atmosphere? Annual Meeting of the Geological Society of America, Abstracts with Program, p. A89. Ohmoto, H., 2003a. Banded iron-formations and the evolution of the atmosphere, hydrosphere, biosphere, and lithosphere. Transactions of the Institution of Mining and Metallurgy (Section B, Applied Earth Science) 112, 161–162. Ohmoto, H., 2003b. Nonredox transformation of magnetite–hematite in hydrothermal systems. Economic Geology 98, 157–161. Okamoto, G., Okura, T., Goto, K., 1957. Properties of silica in water. Geochimica et Cosmochimica Acta 12, 123–132. O'Rourke, J.E., 1961. Paleozoic banded iron-formations. Economic Geology 56, 331–361. Paolo, C., Wiele, S.M., Reinhart, M.A., 1989. Upper regime parallel lamination as the result of turbulent sediment transport and lowamplitude bed forms. Sedimentology 36, 47–59. Parr, J., Yeats, C., Binns, R., 2003. Petrology, trace element geochemistry and isotope geochemistry of sulfides and oxides from the PACMANUS hydrothermal field, Eastern Manus Basin, Papua New Guinea. In: Yeats, C.J. (Ed.), Seabed Hydrothermal Systems of the Western Pacific. CSIRO Exploration and Mining Report, vol. 112F. CSIRO Division of Exploration and Mining, North Ryde, Australia, pp. 53–57. Pickard, A.L., Barley, M.E., Krapez, B., 2004. Deep-marine depositional setting of banded iron formation: sedimentological evidence from interbedded clastic sedimentary rocks in the early Palaeoproterozoic Dales Gorge Member of Western Australia. Sedimentary Geology 170, 37–62. Powell, C.M., Oliver, N.H.S., Li, Z.X., Martin, D.M., Ronaszeki, J., 1999. Synorogenic hydrothermal origin for giant Hamersley iron oxide ore bodies. Geology 27, 175–178. Pufahl, P.K., Fralick, P.W., 2004. Depositional controls on Palaeoproterozoic iron formation accumulation, Gogebic Range, Lake Superior region, USA. Sedimentology 51, 791–808. Quade, H., 1970. Der Bildungraum und die genetische Problematik der vulkanosedimentären Eisenerze. Clausthaler Hefte 9, 27–65. Quade, H., 1976. Genetic problems and environmental features of volcano–sedimentary iron-ore deposits of the Lahn–Dill type. In: Wolf, K.H. (Ed.), Handbook of Strata-Bound and Stratiform Ore Deposits, vol. 7. Elsevier, Amsterdam, pp. 255–294. Rasmussen, B., Buick, R., 1999. Redox state of the Archean atmosphere: evidence from detrital heavy minerals in ca. 3250– 2750 Ma sandstones from the Pilbara Craton, Australia. Geology 27, 115–118.

D.F. Lascelles / Ore Geology Reviews 32 (2007) 381–411 Roberts, H.M., Bartley, M.W., 1943. Hydrothermal replacement in deep-seated ores of the Lake Superior region. Economic Geology 38, 1–24. Roscoe, S.M., 1968. Huronian rocks and uraniferous conglomerates in the Canadian Shield. Geological Survey of Canada Paper 68-40 (205 pp.). Rosière, C.A., Rios, F.J., 2004. The origin of hematite in high-grade iron ores based on infrared microscopy and fluid inclusion studies: the example of the Conceição Mine, Quadrilátero Ferrífero, Brazil. Economic Geology 99, 611–621. Rundquist, D.V., Sobolev, P.O., 2002. Seismicity of mid-oceanic ridges and its geodynamic implications: a review. Earth-Science Reviews 58, 143–161. Rye, R., Holland, H.D., 1998. Palaeosols and the evolution of atmospheric oxygen: a critical review. American Journal of Science 298, 621–672. Shegelsky, R.J., 1987. The depositional environment of Archean iron formations, Sturgeon–Savant greenstone belt, Ontario, Canada. In: Appel, P.W.U., LaBerge, G.L. (Eds.), Precambrian IronFormations. Theophrastus Publications, Athens, pp. 329–344. Siehl, V.A., Thein, J., 1989. Minette type ironstones. In: Young, T.P., Taylor, W.E.G. (Eds.), Phanerozoic Ironstones. Special Publication, vol. 46. The Geological Society, London, pp. 175–193. Simonson, B.M., 1985. Sedimentology of cherts in the Early Proterozoic Wishart Formation, Quebec-Newfoundland, Canada. Sedimentology 32, 23–40. Simonson, B.M., Lanier, W.P., 1987. Early silica cementation and microfossil preservation in iron-formation stromatolites, early Proterozoic of Canada. In: Appel, P.W.U., LaBerge, G.L. (Eds.), Precambrian Iron-Formations. Theophrastus Publications, Athens, pp. 187–214. Spears, D.A., 1989. Aspects of iron incorporation into sediments with special reference to the Yorkshire Ironstones. In: Young, T.P., Taylor, W.E.G. (Eds.), Phanerozoic Ironstones. Special Publication, vol. 46. The Geological Society, London, pp. 19–30. Stow, D.A.V., Bowen, A.J., 1980. A physical model for the transport and sorting of fine-grained sediment by turbidity currents. Sedimentology 27, 31–46. Tamura, T., Masuda, F., 2005. Bed thickness characteristics of inner shelf storm deposits associated with a transgressive to regressive Holocene wave-dominated shelf, Sendai coastal plain, Japan. Sedimentology 52, 1375–1396. Taylor, D., Dalstra, H.J., Harding, A.E., Broadbent, G.C., Barley, M.E., 2001. Genesis of high-grade hematite orebodies of the Hamersley Province, Western Australia. Economic Geology 96, 837–873. Thorne, W.S., Hagemann, S.G., Barley, M.E., 2004. Petrographic and geochemical evidence for hydrothermal evolution of the North Deposit, Mt Tom Price, Western Australia. Mineralium Deposita 39, 766–783. Towe, K.M., 1983. Precambrian atmospheric oxygen and banded iron formations: a delayed ocean model. Precambrian Research 20, 161–170. Trendall, A.F., 1965. Progress report on the Brockman Iron Formation in the Wittenoom–Yampire area. Western Australia Geological Survey Annual Report for 1964, pp. 55–65.

411

Trendall, A.F., 1973a. Precambrian iron-formations of Australia. Economic Geology 68, 1023–1034. Trendall, A.F., 1973b. Varve cycles in the Weeli Wolli Formation of the Precambrian Hamersley Group, Western Australia. Economic Geology 68, 1089–1097. Trendall, A.F., 1983. The Hamersley Basin. In: Trendall, A.F., Morris, R.C. (Eds.), Iron Formation: Facts and Problems. Elsevier Science Publications, Amsterdam, pp. 69–129. Trendall, A.F., Blockley, J.G., 1970. The iron formations of the Precambrian Hamersley Group, Western Australia, with special reference to the associated crocidolite. Western Australia Geological Survey, Bulletin 119 (353 pp.). Van Hise, C.R., Leith, C.K., 1911. The Geology of the Lake Superior Region. United States Geological Survey Monograph 52 (641 pp.). Van Loon, A.J., 2004. From speculation to model: the challenge of launching new ideas in the earth sciences. Earth-Science Reviews 65, 305–313. Von Damm, K.L., 1990. Seafloor hydrothermal activity: black smoker chemistry and chimneys. Annual Review of Earth and Planetary Sciences 18, 173–204. Walter, M.R., Hofmann, H.J., 1983. The palaeontology and palaeoecology of Precambrian iron-formations. In: Trendall, A.F., Morris, R.C. (Eds.), Iron-Formation: Facts and Problems. Developments in Precambrian Geology, vol. 6. Elsevier, pp. 373–400. Watkins, K.P., Hickman, A.H., 1990. Geological evolution and mineralisation of the Murchison Province. Geological Survey of Western Australia, Bulletin 137 (267 pp.). Whitten, G.F., 1970. The investigation and exploitation of the Razorback Ridge iron deposit. Geological Survey of South Australia, Report on Investigations 33 (165 pp.). Yamamoto, S., 1988. Ferromagnesian and metalliferous pelagic clay minerals in oceanic sediments. In: Chilingarian, G.V., Wolf, K.H. (Eds.), Diagenesis, ll. Developments in Sedimentology, vol. 43. Elsevier, Amsterdam, pp. 115–146. Yang, W., Holland, H.D., Rye, R., 2002. Evidence for low or no oxygen in the late Archean atmosphere from the ∼2.76 Ga Mt Roe #2 paleosol, Western Australia. Geochimica et Cosmochimica Acta 66, 3707–3718. Yariv, S., Cross, H., 1979. Geochemistry of Colloid Systems. Springer–Verlag, New York. 450 pp. Yeats, C.J., 2003. Mineralogy and geochemistry of alteration at the PACMANUS hydrothermal field, Eastern Manus Basin, Papua New Guinea: vertical and lateral variation at low and hightemperature vent sites. In: Yeats, C.J. (Ed.), Seabed Hydrothermal Systems of the Western Pacific. CSIRO Exploration and Mining Report, vol. 1112F, pp. 53–57. North Ryde, Australia. Young, G.M., 1976. Iron-formation and glaciogenic rocks of the Rapitan Group, Northwest Territories, Canada. Precambrian Research 3, 137–158. Young, T.P., 1989. Eustatically controlled ooidal ironstone deposition. In: Young, T.P., Taylor, W.E. (Eds.), Phanerozoic Ironstones. Geological Society Special Publication, vol. 46, pp. 51–63. Young, T.P., Taylor, W.E., 1989. Phanerozoic Ironstones. Geological Society Special Publication 46 (251 pp.).