is the 180m that is outstanding has what got mr market excited...

is the 180m that is outstanding has what got mr market excited !?!
any geo's want to have a guess what the 180m is likely to contain ?
one of WA1 luni hits Hole 34 has 144m at 2.9% TiO2.



Ore Geology Reviews

Volume 61, September 2014, Pages 1-32


Review
The chemistry of hydrothermal magnetite: A review

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https://doi.org/10.1016/j.oregeorev.2013.12.013Get rights and content

Abstract

Magnetite (Fe₃O₄) is a well-recognized petrogenetic indicator and is a common accessory mineral in many ore deposits and their host rocks. Recent years have seen an increased interest in the use of hydrothermal magnetite for provenance studies and as a pathfinder for mineral exploration. A number of studies have investigated how specific formation conditions are reflected in the composition of the respective magnetite. Two fundamental questions underlie these efforts — (i) How can the composition of igneous and, more importantly, hydrothermal magnetite be used to discriminate mineralized areas from barren host rocks, and (ii) how can this assist exploration geologists to target ore deposits at greater and greater distances from the main mineralization? Similar to igneous magnetite, the most important factors that govern compositional variations in hydrothermal magnetite are (A) temperature, (B) fluid composition — element availability, (C) oxygen and sulfur fugacity, (D) silicate and sulfide activity, (E) host rock buffering, (F) re-equilibration processes, and (G) intrinsic crystallographic controls such as ionic radius and charge balance. We discuss how specific formation conditions are reflected in the composition of magnetite and review studies that investigate the chemistry of hydrothermal and igneous magnetite from various mineral deposits and their host rocks. Furthermore, we discuss the redox-related alteration of magnetite (martitization and mushketovitization) and mineral inclusions in magnetite and their effect on chemical analyses. Our database includes published and previously unpublished magnetite minor and trace element data for magnetite from (1) banded iron formations (BIF) and related high-grade iron ore deposits in Western Australia, India, and Brazil, (2) Ag–Pb–Zn veins of the Coeur d'Alene district, United States, (3) porphyry Cu–(Au)–(Mo) deposits and associated (4) calcic and magnesian skarn deposits in the southwestern United States and Indonesia, and (5) plutonic igneous rocks from the Henderson Climax-type Mo deposit, United States, and the un-mineralized Inner Zone Batholith granodiorite, Japan. These five settings represent a diverse suite of geological settings and cover a wide range of formation conditions.
The main discriminator elements for magnetite are Mg, Al, Ti, V, Cr, Mn, Co, Ni, Zn, and Ga. These elements are commonly present at detectable levels (10 to > 1000 ppm) and display systematic variations. We propose a combination of Ni/(Cr + Mn) vs. Ti + V, Al + Mn vs. Ti + V, Ti/V and Sn/Ga discriminant plots and upper threshold concentrations to discriminate hydrothermal from igneous magnetite and to fingerprint different hydrothermal ore deposits. The overall trends in upper threshold values for the different settings can be summarized as follows: (I) BIF (hydrothermal) — low Al, Ti, V, Cr, Mn, Co, Ni, Zn, Ga and Sn; (II) Ag–Pb–Zn veins (hydrothermal) — high Mn and low Ga and Sn; (III) Mg-skarn (hydrothermal) — high Mg and Mn and low Al, Ti, Cr, Co, Ni and Ga; (IV) skarn (hydrothermal) — high Mg, Al, Cr, Mn, Co, Ni and Zn and low Sn; (V) porphyry (hydrothermal) — high Ti and V and low Sn; (VI) porphyry (igneous) — high Ti, V and Cr and low Mg; and (VII) Climax-Mo (igneous) — high Al, Ga and Sn and low Mg and Cr.

Introduction

Magnetite (Fe₃O₄) is an important petrogenetic indicator and pathfinder mineral with a wide array of applications including geophysical studies, igneous petrology, provenance studies and mineral exploration (e.g., Dupuis and Beaudoin, 2011, Ghiorso and Sack, 1991b, Grant, 1984a, Lindsley, 1976a, McClenaghan, 2005, Razjigaeva and Naumova, 1992). The compositional variability of magnetite in response to varying formation conditions has been the focus of many studies over the last few decades. Recent years have seen particular interest in hydrothermal magnetite as a pathfinder mineral for exploration — facilitated by the development and improvement of analytical techniques such as laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) and trace mode electron microprobe analysis (EMPA), which allow in-situ measurements with increasingly lower detection limits (e.g., Dupuis and Beaudoin, 2011, Longerich et al., 1996, Nadoll and Koenig, 2011). We review compositional trends in magnetite from a variety of hydrothermal ore deposits and their host rocks and discuss how these compositional variations can be used as a geochemical fingerprint. We present previously published and unpublished minor and trace element data for hydrothermal and igneous magnetite from (1) banded iron formations (BIF) and related high-grade iron ore deposits in Western Australia, India, and Brazil, (2) Ag–Pb–Zn veins of the Coeur d'Alene district, United States, (3) porphyry Cu–(Au)–(Mo) deposits and associated (4) calcic and magnesian skarn deposits in the southwestern United Sates (Chino, Cobre, Copper Flat, Safford, Morenci) and Indonesia (Ertsberg district), and (5) plutonic igneous rocks from the Henderson Climax-type Mo deposit, United States, and the un-mineralized Inner Zone Batholith granodiorite, Japan.
Magnetite belongs to the space group Fd3m and has an inverse spinel structure with the general stoichiometry AB₂O₄ (Bragg, 1915, Fleet, 1981), where A represents a divalent cation such as Mg, Fe^{2 +}, Ni, Mn, Co, or Zn, and B represents a trivalent cation such as Al, Fe^{3 +}, Cr, V, Mn, or Ga (Lindsley, 1976a, Wechsler et al., 1984). Titanium, with a 4 + charge, can also occupy the B site when substitution is coupled with a divalent cation (Wechsler et al., 1984). Octahedral sites in the magnetite structure are randomly occupied by subequal numbers of ferric (Fe^{3 +}) and ferrous (Fe^{2 +}) iron atoms, whereas tetrahedral sites are exclusively occupied by the smaller ferric iron atoms Fe^{3 +}[Fe^{2 +}Fe^{3 +}]O₄ (Lindsley, 1976a, Waychunas, 1991, Wechsler et al., 1984) (Fig. 1A).
Magnetite displays many phase transitions with other spinel group minerals, including but not limited to, spinel (MgAl₂O₄), ulvöspinel (Fe₂TiO₄), chromite (FeCr₂O₄), galaxite (Mn,Mg)(Al,Fe)₂O₄, gahnite (ZnAl₂O₄), franklinite ((Zn, Fe,Mn)(Fe,Mn)₂O₄), and jacobsite (Mn,Fe,Mg)(Fe,Mn)₂O₄. Above 600 °C a continuous solid solution exists between magnetite and ulvöspinel (titanomagnetite_ss — Ti_xFe_{3 − x}O₄), and their oxidation products (titanomaghemite), with coupled substitution of Ti^{4 +} for Fe^{3 +} in the octahedral sites and Fe^{2 +} for Fe^{3 +} in the tetrahedral sites (Buddington and Lindsley, 1964, Waychunas, 1991, White et al., 1994). Below 600 °C thermodynamic data are poorly constrained and extensive miscibility gaps occur (Ghiorso and Sack, 1991b). A log fO₂–T diagram with relevant buffers for the Fe–Si–O system and a schematic phase diagram for the system Fe–O–S in fO₂–fS₂ space are shown in Fig. 2A and B (redrawn after Frost, 1991a, Frost et al., 1988). The minimum oxygen fugacity for any given temperature at which magnetite is stable is the iron-magnetite (IM) or magnetite–wüstite (MW) buffer (Buddington and Lindsley, 1964, Frost, 1991a). The fayalite–magnetite–quartz (FMQ) buffer marks the limit above which Fe is mostly incorporated into magnetite. Below the FMQ buffer, iron is predominantly present in silicates. The quartz–iron–fayalite (QIF) buffer marks the limit below which Fe occurs in its native state. The upper limit for magnetite stability is defined by the hematite–magnetite (HM) buffer, beyond which hematite is the dominant oxide mineral (Buddington and Lindsley, 1964, Frost, 1991a, Grant, 1984a). Cation substitution is generally more likely to take place at lower oxygen fugacity (Buddington and Lindsley, 1964, Frost, 1991a, Lindsley, 1991). Magnesium, Mn, Zn, and Ni may substitute Fe^{2 +}, whereas Fe^{3 +} can be replaced by Al, V and Cr (Fig. 1B, C) (e.g., Barnes and Roeder, 2001, Lindsley, 1991, Ramdohr, 1955, Righter et al., 2006a). Vanadium, which has a range of possible oxidation states depending on prevailing oxygen fugacity conditions, is mainly present as V^{3 +} in titanomagnetites but commonly shows variable minor amounts of V^{4 +} (Balan et al., 2006, Bordage et al., 2011, Toplis and Carroll, 1995). Vanadium is incompatible at high oxygen fugacity levels due to its 5 + oxidation state. The strong oxygen fugacity-dependence of V can be used to help model the differentiation of magmatic intrusions (Toplis and Corgne, 2002). Similar applications are relevant for hydrothermal conditions. For example, magnetite and coexisting hematite have been used to define redox-potentials in hydrothermal fluids associated with ore deposits (Barnes, 1997, Otake et al., 2010).
Magnetite is one of the most abundant oxide minerals in the continental crust and has been recognized as an important indicator mineral for petrogenetic and geochemical studies since the early 20th century (Ramdohr, 1926). Indicator minerals such as magnetite are generally more resistant to weathering and transport than other coexisting mineral phases (McClenaghan, 2005). Magnetite is a widespread and easily identifiable accessory mineral in igneous, sedimentary, and metamorphic host rocks of a wide range of compositions (e.g., Annersten, 1968, Buddington and Lindsley, 1964, Prins, 1972, Ramdohr, 1955, Rumble, 1976b, Shcheka et al., 1980, Vincent and Phillips, 1954) and can incorporate a large number of foreign cations (Bowles et al., 2011, Lindsley, 1976a, Wechsler et al., 1984) (Fig. 1). In combination with isotopic studies, compositional patterns in magnetite can provide important constraints on petrogenetic factors such as temperature, pressure, and oxygen/sulfur fugacity (Anderson et al., 2008, Chiba et al., 1989, Frost, 1991b, Ghiorso and Sack, 1991b). Magnetite forms ideal pairs with silicates, carbonates, and other oxides such as hematite or ilmenite for geothermometry and geobarometry due to its distinct oxygen isotope fractionation factor (e.g., Buddington and Lindsley, 1964, Buddington et al., 1955, Chacko et al., 2001, Friedman and O'Neil, 1977, Ghiorso and Sack, 1991a, Powell and Powell, 1977, Sauerzapf et al., 2008, Zheng and Simon, 1991). By virtue of its magnetic properties, magnetite has also been the focus of many geophysical studies that investigated its use for geophysical mapping and exploration (e.g., Akimoto, 1955, Behn et al., 2001, Clark, 2001, Grant, 1984b, Malmqvist and Parasnis, 1972, McEnroe et al., 2001).
Minor and trace elements that have been reported in magnetite ranges from alkali and alkaline earth metals to transition metals, including REEs, and metalloid and non-metals (e.g., Borisenko and Lapin, 1972, Dare et al., 2012, Dupuis and Beaudoin, 2011, Klemm et al., 1985, Lindsley, 1991, McQueen and Cross, 1998, Nadoll et al., 2012, Nielsen and Beard, 2000, Righter et al., 2006a, Singoyi et al., 2006). However, a suite of elements, namely Mg, Al, Ti, V, Co, Ni, Zn, Cr, Mn, Ga and Sn are commonly present in magnetite of all origins at concentrations that can be detected by electron microprobe analysis (10 to > 1000 ppm). We refer to these cations as magnetite elements.
Concentrations of foreign cations in magnetite are responsive to a number of external factors that reflect formation conditions of the respective host rock. Comprehensive overviews of physicochemical properties of magnetite can be found in Bowles et al. (2011) and the Reviews in Mineralogy volumes on oxide minerals (Lindsley, 1991, Rumble, 1976a). Many classic studies that focused on igneous magnetite show extensive compositional variability in response to seven main controlling factors: (1) source rock or magma composition, (2) temperature, (3) pressure, (4) cooling rate, (5) oxygen fugacity, fO₂, (6) sulfur fugacity, fS₂, and (7) silica and sulfide activity (Buddington and Lindsley, 1964, Frost and Lindsley, 1991, Ghiorso and Sack, 1991a, Haggerty, 1991a, Mollo et al., 2013, Whalen and Chappel, 1988). Additionally, crystallographic factors such as ionic radius and overall charge balance are important aspects that put decisive constraints on possible substituting cations (Cornell and Schwertmann, 2003, Fleet, 1981, Goldschmidt, 1954, Wechsler et al., 1984) (Fig. 1). Magnetite is also an important accessory mineral in metamorphic rocks. The composition of metamorphic magnetite changes in response to mainly two factors: temperature and oxygen fugacity (Frost, 1991c). The partitioning behavior and distribution of trace elements in magnetite depends on the metamorphic grade (Evans and Frost, 1975, Skublov and Drugova, 2003, Van Baalen, 1993). Low-grade metamorphic magnetite is compositionally homogeneous and has very low trace element concentrations compared to most other types of hydrothermal magnetite (Frost, 1991c, Nadoll et al., 2012). This has been interpreted as a reflection of the low formation temperatures (Nadoll et al., 2012). Although elements such as Ti are essentially immobile under low temperatures, Van Baalen (1993) showed that Ti can be mobilized on a length scale of meters even in low-grade metamorphic rocks. Verlaguet et al. (2006) came to similar conclusions for the comparably immobile element Al. Furthermore, Al contents in spinel minerals from metamorphic rocks containing silicates are governed by pressure-, temperature-, and water fugacity-sensitive equilibria involving chlorite (Evans and Frost, 1975). Oxide–silicate, oxide–oxide, and intra-oxide re-equilibration reactions are further important controls of the magnetite stability during cooling of plutonic and metamorphic rocks (e.g., Frost, 1991a, Frost, 1991b, Frost, 1991c, Frost and Lindsley, 1991, Nadoll et al., 2012, Wones, 1989).
Recent studies put further constraints on compositional trends of igneous magnetite and reinforced its petrogenetic significance. For example, the significance of magnetite from massive sulfide Ni–Cu–PGE deposits as a sensitive gauge for fractionation processes in sulfide melts has been investigated by Dare et al. (2012). They found that minor and trace element concentrations in magnetite are governed by the successive depletion of lithophile elements in the sulfide melt and the uptake of siderophile elements such as Co, Mo, Ni, Pb, Sn, and Zn in co-crystallizing sulfides. Jenner et al. (2010) recently emphasized the significant role that the onset of magnetite crystallization plays for the saturation of sulfides in a melt. Righter et al. (2006a) presented new values for Ni, Co, and V partition coefficients between Cr-rich spinels and silicate melt. Igneous magnetite from carbonatites has been investigated by a number of authors who found indicative minor and trace element signatures that can be attributed to controlling factors such as oxygen fugacity and interdependencies among substituting cations (Bagdasarov, 1989, Bailey and Kearns, 2002, Reguir et al., 2008). Reguir et al. (2008) observed V and Mn concentrations to be directly controlled by oxygen fugacity. Furthermore, they present data that Zn in igneous magnetite from carbonatites shows a covariation with Mn, suggesting a link between the two elements controlled by coupled substitution.
Ryabchikov and Kogarko (2006) constructed fO₂–T diagrams for the Russian Khibina magmatic system using Fe–Ti-oxide equilibria. Varying Mg, Al, Mn, Ti, and Cr concentrations in magnetite have been linked with the different intrusive stages at the Sokli complex in northeastern Finland by Lee et al. (2005). Laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) trace element analyses of magnetite from Kiruna, Sweden, showed that V and Mn are the most abundant trace elements in igneous magnetite, whereas Cu, Zn, and Pb are not commonly incorporated (Müller et al., 2003).
The identification of characteristic minor and trace element signatures in magnetite of predominantly hydrothermal origin commenced with the pioneering work of Borisenko and Lapin, 1971, Borisenko and Lapin, 1972 and Shcheka et al. (1988), who were also among the first to apply multivariate statistics to magnetite minor and trace element data in order to reveal underlying compositional trends that could be used to discriminate sample populations. Recent studies have presented data that suggest that magnetite from magmatic–hydrothermal deposits displays systematic variations in minor and trace element concentrations that can be used to fingerprint ore deposits (Beaudoin and Dupuis, 2009, Dupuis and Beaudoin, 2011, Kamvong et al., 2007, Nadoll et al., 2012, Rusk et al., 2009, Singoyi et al., 2006). Carew et al. (2006) and Rusk et al. (2010) presented evidence that Ti, V, and Mn concentrations in hydrothermal magnetite from the Ernest Henry iron oxide copper gold (IOCG) deposit in Australia's Cloncurry region can be used to discriminate barren from mineralized host rocks. Singoyi et al. (2006) observed three distinct types of magnetite (referred to as types A, B, and C) from selected volcanic-hosted massive sulfide (VMS), skarn, IOCG, and Broken Hill-type clastic-dominated Pb–Zn deposits in Australia. Type A commonly and consistently incorporates Mg, Al, Ti, V, Mn, Co, Ni, Zn, Ga and Sn at levels above LA-ICP-MS detection limits, whereas type B displays lesser concentrations and more heterogeneous distribution across different samples for Cr, As, Zr, Nb, Mo, REE, Ta, W and Pb. Their type C magnetite hosts no detectable concentrations of Cu, Ag, Se, Tl, Te, Bi and Au. Singoyi et al. (2006) further suggest that Sn/Ga and Al/Co ratios can help to distinguish magnetite from VMS, skarn, IOCG, and Broken Hill-type clastic-dominated Pb–Zn deposits; an approach that was later adopted by Kamvong et al. (2007). McQueen and Cross (1998) demonstrated how trace element variations in hydrothermal magnetite hosted in contact metasomatic skarn deposits can provide characteristic signatures and consequently, help to target magnetite-bearing ore deposits. Most recently, Duan et al. (2012) documented systematic compositional variations among different types of hydrothermal magnetite, tracing the evolution of the magmatic–hydrothermal mineralization in the Washan porphyry-style Fe deposit, East China. Huang et al. (2013) described the compositions of hydrothermal magnetite from the magmatic–hydrothermal Cihai Fe deposit in China. They compared their results with the magnetite classification scheme suggested by Dupuis and Beaudoin (2011) and showed that magnetite from the Cihai deposit is depleted in V and Ti compared to magnetite from Fe–Ti–V deposits and has Mg, Al, Ti, V, Cr, Co, Ni, Mn, Zn, Ga, and Sn concentrations that exceed those commonly found in magnetite from skarn deposits.
Mining and exploration companies have recognized the great potential of employing the geochemistry of magnetite as an exploration tool for many years. There is a large body of unpublished work that investigated the use of magnetite as a pathfinder or indicator mineral. One of the main challenges in mineral exploration is to detect geochemical signatures of ore deposits at greater and greater distances from the main mineralization. Magnetite, among other detrital heavy minerals such as ilmenite, rutile, garnet, and zircon has been a valuable asset to mineral exploration and provenance studies (e.g., Dupuis and Beaudoin, 2011, Nadoll et al., 2012, Nadoll 2011, Razjigaeva and Naumova, 1992, Shcheka et al., 1982). Hydrothermal and igneous magnetite can provide vectors to mineralized areas where the primary mineralogical context is lost, deeply covered, or highly altered, similar to the use of rutile as a resistate indicator mineral (Anand and Butt, 2010, Borisenko and Lapin, 1972, Buddington and Lindsley, 1964, Hutton, 1950, McQueen and Cross, 1998, McQueen and Whitbread, 2002, Rumble, 1976a and references therein, Shcheka et al., 1982, Triebold et al., 2007, Zack et al., 2004). Razjigaeva and Naumova (1992) used variations in Ti, Mn, Cr, V, Ni, Co, Zr, Sn, Zn, Pb, and Cu concentrations in detrital magnetite to trace the source rocks of sediments, comparing them to a set of standard igneous magnetites first described by Shcheka et al. (1980). Similarly, Grigsby (1990) used petrographic and chemical variations in detrital magnetite to discriminate among felsic, intermediate, and mafic volcanic and plutonic source rocks.
The term hydrothermal magnetite can be rather ambiguous especially when considering the complex geological and mineralogical relationships found in many hydrothermal ore deposits (Fig. 3). Multiple vein generations, large- to small-scale overprints of multiple alteration stages, fluid/rock interaction, and secondary weathering processes translate into a variety of different types of hydrothermal magnetite. These are commonly synonymously addressed in publications. However, it is important to consider the different types of hydrothermal magnetite when attributing specific minor and trace element concentrations to a specific deposit or type of mineralization. One of the main challenges for researchers that investigate hydrothermal magnetite is to diligently record the large variety of minor and trace element compositions in magnetite that reflect a diverse formation and alteration history, and at the same time invoke diagnostic trends that can be used to reliably discriminate different ore deposits, as well as barren zones from mineralized zones.

Section snippets

Electron microprobe

Electron microprobe analysis (EMPA) were carried out at North Ryde (Sydney) on a Cameca Camebax™ automated wave length dispersive spectroscopy scanning electron microprobe with an acceleration voltage of 15 kV and a beam current of 20 nA. The analytical conditions for each element are given in Table 1. Line overlap corrections were performed for the overlap of Ti Kβ on VKα, V Kβ on CrKα, Cr Kβ on MnKα and Mn Kβ on FeKα. ZAF corrections were applied for all elements. Microprobe results were
Geologic background

This section provides a geologic overview of the investigated areas and deposits. Locality maps are shown in Fig. 4, Fig. 5. A complete list of samples including host rock and alteration types are shown in Table 3.
Magnetite chemistry

A total of 1415 EMPA and LA-ICP-MS analyses, of which 1103 have not been previously published, represent the database for our geochemical evaluation. Summary statistics are shown in Table 4 and the corresponding box and whisker plots are shown in Fig. 7. Probability plots (Fig. 8) and an upper threshold value radar plots (Fig. 9) are essentially derivates of the box and whisker plots, but help to identify compositional trends by displaying different aspects of the data. Probability plots help
Inclusions

Mineral inclusions are common features in igneous and hydrothermal magnetite; they reflect magnetite formation conditions and give a direct insight into the evolution of the corresponding host rock. Mineral inclusions are of particular interest when geological and mineralogical context have been removed or significantly altered, which is the case for stream sediments or regolith cover. Some inclusions, such as Cu sulfide minerals, can be a direct pointer towards mineralized areas. Wang et al.
Concluding remarks

One of the main challenges for mineral exploration is the recognition of mineralized areas at increasingly greater distance from primary mineralization. Minor and trace element data for magnetite from a variety of mineral deposits are becoming more available, and help to further elucidate trends that can be used as geochemical signatures for provenance studies and mineral exploration. Published and previously unpublished magnetite minor and trace element data presented here give evidence for a
Acknowledgements

We thank the iron ore research group at UWA (CET) — Paul Duuring, Ana-Sophie Hensler, and Warren Thorne — for kindly providing data for this project, and Steffen Hagemann for helpful feedback. Thomas Angerer thanks the LabMaTer group at UQAC, Sarah-Jane Barnes, Sarah Dare, Dany Savard, Sadia Mehdi, and the CODES laser ablation team at UTAS, Sarah Gilbert and Leonid Danyushevsky, for granting access to their laser systems and helping with analyses and data reduction. Further acknowledgments go