An International Study Tour to Order
Organised for Vedanta Resources by Porter GeoConsultancy
Base Metal Mines of Northern Australia
7 to 10 November 2005
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Mount Isa
This tour was developed and organised by Porter GeoConsultancy Pty Ltd (PGC) to suit the specific requirements of the Hindustan Zinc and Exploration divisions of Vedanta Resources. All mine visit approvals, and the planning and organisation of logisitics (including ground transport and air charters), meals (including special dietary requirements) and accommodation were undertaken by PGC.

The tour visited the following mines in the State of Queensland and the Northern Territory of Australia:
A detailed itinery explaining the arrangements, what logistics items had been arranged, times, pick-up locations, maps, reporting places and times, etc., was presented to the tour members prior to the commencement of the tour to manage themselves.   A technical Literature Compilation and set of geological maps relevant to the deposits visited were also purchased/prepared and presented to the tour party prior to the commencement of the tour.   The group was based in Mount Isa throughout, travelling to the mines by either hire van or by charter aircraft.

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Ernest Henry .......... Monday 7 November, 2005 .......... Travel by hire van from Mt Isa.

The Ernest Henry IOCG style Cu-Au deposit is located 35 km NE of Cloncurry, 150 km east of Mt Isa and 750 km west of Townsville in north-west Queensland (#Location: 20° 26' 40"S, 140° 42' 21"E).

The deposit lies to the east of the Cloncurry Overthrust, within the Cloncurry-Selwyn zone of the Cloncurry Terrane, which comprises the eastern exposed margin of the Mount Isa Inlier of North-west Queensland. It contains IOCG deposits that are hosted by Palaeoproterozoic (1760-1660 Ma) silici-clastic metasedimentary and metavolcanic rocks that were deposited during periods of ensialic rifting.

The Ernest Henry deposit is hosted within the Eastern Succession of the Mount Isa Inlier, that consists of a poly-deformed Palaeo- and Mesoproterozoic volcano-sedimentary succession which is largely composed of evaporite-rich Cover Sequence 2 and silici-clastic-rich Cover Sequence 3 rocks (CS2 and 3). CS2 and 3 were deposited between 1790 and 1690 Ma and from 1680 to 1610 Ma respectively. To the west, these sequences overlie an older crystalline basement and a core of predominantly Cover Sequence 1 felsic volcanic and related intrusive rocks that correspond to the 1870 to 1850 Ma Barramundi Orogeny of northern Australia. Basement is not exposed in the Cloncurry district. Both CS2 and 3 were deposited in intracontinental rift settings, although the relationship between some parts of the sequence is obscured by the deformation history. Both sequences were also accompanied by the emplacement of various intrusive and volcanic rocks.

The first significant deformation to affect CS2 (but not CS3) was the 1750 to 1735 Ma Wonga extensional event. CS2 was extensively intruded by the 1750 to 1730 Ma Wonga Granite to the west, while the coeval Mount Fort Constantine volcanics are found to the NE. Minor tonalites, granitoids and diorite emplaced between CS2 and 3 have been dated at 1686 to 1660 Ma (including the Ernest Henry Diorite).

Thin skinned deformation of the ~1600 to 1520 Ma Isan Orogeny terminated deposition of Cover Sequence 3, and resulted in gross eastward tectonic transport, interleaving of major lithostratigraphic units, and a dominant north-south tectonic grain. This deformation has been divided into: a D1 event, which involved overall north-south compression, and is characterised by large-scale thrusts and isoclinal folds, thrust reactivation of large, km-scale, basin bounding extensional faults with CS3 rocks thrust over CS2, resulting in overturned limbs and a penetrative rock mass foliation; a D2 event, involving horizontal east-west compression producing major north-south upright to isoclinal folding of CS2 and 3 rocks, and a penetrative cleavage, which peaked at 1595 to 1580 Ma with a regional greenschist to upper amphibolite facies metamorphism and the development of anatectic pegmatites; and a D3 event which includes NW- and NE-trending brittle-ductile corridors of faulting, kinking and folding with steep plunges to the NW and NE, and dominantly north-south trending shear and fault zones and associated breccia formation.

Both CS2 and 3 were intruded by the voluminous Williams and Naraku granite batholiths at 1540 to 1500 Ma (including the 1530 Ma Mt Margaret Granite immediately to the east of the E1 deposits; Marshall and Oliver, 2007; Page and Sun, 1998). These represent the youngest felsic intrusions in the inlier, and have an outcrop exposure of >1500 km
2. They were emplaced in an intracratonic environment, and have a pre-, syn- and post-D3 timing, and are largely composed of alkaline to sub-alkaline, K-rich, A-type, magnetite-bearing granitoids. They range from diorite to syenogranite in composition and are typically more oxidised than similar older (~1670 Ma) granitoids in the Western Fold Belt of the Mount Isa Inlier. Sodic intrusions of similar age are rare.

A regionally extensive Na-Ca hydrothermal system in the Cloncurry district (>1000 km
2) affected all rock types, especially the resultant calc-silicate-rich lithologies of cover sequence 2. This alteration appears to have been formed by multiple periods of hydrothermal activity that locally overlapped and is most intense in breccia zones along large structural conduits and within calc-silicate-rich units. The bulk of the sodic-calcic alteration, dominantly regional albite and scapolite, was associated with fluids that were initially mostly sedimentary formation waters with lesser magmatic components, prior to and during peak metamorphism (Kendrick et al., 2008; Oliver et al., 2008; Baker et al., 2008). Subsequent more structurally controlled albite-actinolite-magnetite-titanite±clinopyroxene assemblages, were synchronous with major granite (e.g., Williams-Naraku batholiths) emplacement (Baker et al., 2008), involving a larger magmatic fluid component, and coincided with formation of the majority, but not all of the significant oxide Cu-Au deposits. These deposits may have some stratigraphic control, but are usually associated with brittle and brittle-ductile shear and fault structures which acted as conduits for the transport of high temperature (300 to 500°C) saline fluids into the host rocks (Williams, 1998).

The Ernest Henry deposit is concealed by 35 to 60 m of extensive Phanerozoic cover and does not outcrop. While the exact stratigraphic position of the host rocks is not known, they have been tentatively correlated with the 1730 ±10 Ma Mount Fort Constantine Meta-volcanics towards the top of Cover Sequence 2. The Mount Fort Constantine metavolcanics comprise dacite and andesite with subordinate metabasalts and calc-silicate metasedimentary rocks. The only other outcrop in the district is the 1480 Ma Mount Margaret granite some 12 km to the east. Within Cover Sequence 2, volcanism is common between 1790 and 1780 Ma, and 1760 to 1720 Ma, with later 1540 to 1450 Ma granitoids.

Within the immediate orebody area the principle lithologies encountered are: i). altered plagioclase phyric andesitic volcanic/hypabyssal rocks (ca 1740 Ma) which host the orebody where they are brecciated; ii). various siliciclastic, calc-silicate-rich and graphitic metasedimentary rocks that occur as <10 m thick intercalations within the metavolcanic rocks; and, iii). medium-grained metadiorite (ca 1660 Ma).

Structural analysis suggests that ore deposition accompanied reverse-fault movement between two northeast trending bounding shear zones and formed a pipe-like zone of dilation in the K-feldspathised metavolcanic rocks. The breccia pipe, plunges at ~45° to the SSESSE, nested between the ductile shear zones (Rusk et al., 2010). The orientation of this dilational zone is consistent with the shape and dip of the Ernest Henry ore breccia.

Ernest Henry magnetite-chalcopyrite ore

Image right - High grade Ernest Henry magnetite-chalcopyrite-bornite breccia mineralisation. Image by Mike Porter, 2013.

Four stages of alteration are recognised at Ernest Henry:
i). Regional pre-ore Na-Ca alteration, occurring mainly as albitic plagioclase-, magnetite-, clinopyroxene- and amphibole-rich veining and fault-related breccia-fill.
ii). Pre-mineralisation potassic-(manganese-barium) alteration which only contains minor sulphides, and is typified by multiple stages of K feldspar-, biotite-, amphibole-, magnetite-, garnet- and carbonate-bearing veins, and by fault-related breccia and alteration.
iii). Mineralisation associated alteration, characterised by K feldspar veining and alteration, although this may in part be hematite dusted feldspar. K feldspar alteration is most intense in the vicinity of copper-gold mineralisation, but forms a halo extending from several hundred metres up to 2 km beyond the ore body (Mark et al., 2006), although this outer halo may represent part of pre-ore regional alteration zone. Mineralisation is divided into two main stages, characterised by similar mineral assemblages. The first stage of economic Cu-Au mineralisation was the main ore-forming event, associated with a matrix-supported hydrothermal breccia that is enveloped by crackle veined K feldspar altered meta-volcanic rocks. The second stage of mineralisation occurs as a network of veins cutting earlier infill-supported ore-breccias, and contains a largely identical mineralogy to earlier stage. The ore-bearing assemblage dominantly comprises magnetite, pyrite, chalcopyrite, carbonate and quartz, with lesser apatite, barite, titanite, actinolite, biotite and fluorite. In the upper levels of the deposit, the bulk of the ore is present as hypogene chalcopyrite infilling between K feldspar-altered breccia clasts, while at greater depths, it both infills between, and replaces clasts. Electrum and native gold are closely associated with pyrite and chalcopyrite (Foster et al., 2007).
iv). Post-ore, volumetrically minor, multiple stage calcite-dolomite- and/or quartz-rich veining and alteration which lacks magnetite, and only carries a little gold. Deeper in the deposit, breccias include rounded clasts of previously mineralised breccias containing magnetite, pyrite and chalcopyrite, indicating multiple superimposed brecciation events (Rusk et al., 2010).

Ernest Henry magnetite-hematite-feldspar breccia

Image right - Ernest Henry magnetite and K feldspar-hematite breccia with chalcopyrite mineralisation. Sample collected by Richard Lilley; Image by Mike Porter, 2021.

Rusk et al. (2010) interpret the data from Ernest Henry to be consistent with the following genetic trend:
i). Rapid devolatilisation (of possibly both chloride-rich brines and CO
2-rich fluids) within the source magma chamber;
ii). Fluid over-pressuring in the roof of the magma chamber as a result of volatile exsolution and vapour expansion, assisted by a seal created by magma solidification, sodic-calcic alteration and/or contact metamorphism in the carapace of the igneous complex;
iii). Possible leakage of over-pressured magmatic fluid along structures controlling the location of the later breccia pipe, producing a pre-ore potassic alteration halo;
iv). The eventual failure of the seal and sudden release of fluid pressure, resulting in a high-energy fluid flow event driving brecciation and upward transported and milled clasts. The resultant breccia mass permitted the mixing and/or subsequent ingress of basinal brines circulating within fractured rocks several kilometres above the magma chamber. Fluid mixing, rapid depressurisation and resultant cooling led to ore precipitation within the matrix porosity between breccia clasts at the top of the orebody, where, as the fluid flow, temperature and pressure declined the breccia was sealed;
v). At depth, closer to the heat source, the temperature and pressure gradient degraded more slowly, allowing for fluid-rock reaction to be more protracted, such that prolonged chemical interaction between K feldspar-rich host rocks and ore fluids led to replacement style mineralisation within clasts, with the same mineral assemblage as observed in the shallower parts of the deposit.
vi). At the deepest levels, repetition of the cycle may have resulted in the release of a new pulse of fluids which brecciated and tapped earlier formed magnetite-chalcopyrite rich rocks, telescoping mineralised clasts upwards into the orebody along narrow channels, thereby upgrading ore.

Ernest Henry mineralised brecciated K feldspar

Image right - Ernest Henry brecciated hematite dusted K feldspar with chalcopyrite and magnetite veining. Sample collected by Richard Lilley; Image by Mike Porter, 2021.

The brecciated volcanic mass that hosts the ore forms a plunging elongate body, some 250 m thick, 300 m average length and extending at least 1000 m down plunge to the SSE. The breccia ranges from the unbrecciated volcanics, to crackle fracture veining to clast supported and matrix supported breccia to total clast digestion (massive matrix). The breccias typically contain 5-20 mm subrounded to rounded meta-volcanic and rare biotite altered meta-sedimentary clasts. The matrix is largely composed of magnetite, calcite, pyrite, biotite, chalcopyrite, K feldspar titanite and quartz. Accessory minerals include garnet, barite, molybdenite, fluorite, amphibole, apatite, monazite, arsenopyrite, a LREE fluorcarbonate, galena, cobaltite, sphalerite, scheelite, uraninite and tourmaline. The bulk of the economic mineralisation is restricted to breccia zones with more than 10% matrix.

The total reserve + resource prior to the commencement of mining in 1998 was 166 Mt @ 1.1% Cu, 0.54 g/t Au.
As of June 2003 the remaining resource totalled 117.9 Mt @ 1.13% Cu, 0.52 g/t Au.
At 30 June 2006, the reserves and resources were (Xstrata, 2007):
    Open cut proved reserves - 41 Mt @ 0.9% Cu, 0.5 g/t Au + probable reserves of 20 Mt @ 0.8% Cu, 0.4 g/t Au,
    Open cut measured + indicated resources were the same as, and included the proved and probable reserves,
    Open cut inferred resources - 1 Mt @ 0.4% Cu, 0.2 g/t Au,
    Underground indicated resources - 21 Mt @ 1.5% Cu, 0.7 g/t Au + inferred resources of 23 Mt @ 1.4% Cu, 0.7 g/t Au,

Open pit as at December, 2011 (Xstrata, 2012):
    Total resource and reserve - depleted during 2011 from 17 Mt @ 1.0% Cu, 0.5 g/t Au, 23% magnetite at December 31, 2010
Underground as at December, 2011 (Xstrata, 2012):
    Measured resource - 4 Mt @ 1.3% Cu, 0.7 g/t Au, 32% magnetite
    Indicated resource - 71 Mt @ 1.3% Cu, 0.7 g/t Au, 28% magnetite
    Inferred resource - 13 Mt @ 1.2% Cu, 0.6 g/t Au, 26% magnetite
    Total resource - 88 Mt @ 1.3% Cu, 0.7 g/t Au, 28% magnetite
    Total ore reserve (all probable) - 74 Mt @ 0.95% Cu, 0.5 g/t Au, 23% magnetite.

Remaining Mineral Resources and Ore Reserves as at 31 December 2020 were (Glencore Resources and Reserves Report, 2020):
    Measured resource - 4.7 Mt @ 0.93% Cu, 0.51 g/t Au
    Indicated resource - 55.2 Mt @ 1.16% Cu, 0.61 g/t Au
    Inferred resource - 15.5 Mt @ 1.17% Cu, 0.62 g/t Au, 26%
    Total resource - 75.4 Mt @ 1.15% Cu, 0.61 g/t Au
    Total ore reserve (Proved + Probable) - 38.5 Mt @ 0.95% Cu, 0.50 g/t Au.
NOTE: Resources include Reserves.

The operation is controlled by Ernest Henry Mining Pty Ltd, a subsidiary of Glencore plc. In October 2016 Evolution Mining purchased an economic interest in the copper and gold production from Ernest Henry Mining. In November, 2021, Glencore and Evolution Mining Limited entered into a binding agreement for the sale and purchase of Glencore’s 100% interest in Ernest Henry Mining Pty Ltd.
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Cannington .......... Tuesday 8 November, 2005 .......... Travel by private air charter from Mt Isa.

The Cannington lead - zinc - silver deposit is located 135 km SSE of Cloncurry, 200 km south-east of Mt Isa and 750 km west of Townsville in north-west Queensland (#Location: 21° 52' 9"S, 140° 55' 23"E).

The Cannington deposit occurs in the south eastern corner of the Mt Isa Inlier, to the east of the Cloncurry Overthrust, within the Eastern Fold Belt. The hosts belong to the strongly metamorphosed 1677±9 Ma Fullarton River Group, which have been extensively intruded by the 1560 to 1480 Ma granitoids. The deposit is overlain by 10 to 60 m of Cretaceous and more recent overburden, and was discovered as result of an aeromagnetic survey which identified a sharp magnetic high.

The deposit is hosted by a sequence of garnetiferous psammite within a migmatitic quartzo-feldspathic gneiss terrain. The sequence strikes north-south and is bounded by two major NNW trending structures, the Hamilton and Trepell Faults to the SW and NE respectively, and is cut by similar intervening faults. Four periods of deformation are recognised. The host migmatite gneiss contains intercalated bands of fine grained schistose biotite-sillimanite-quartz bands and pegmatitic quartz-feldspar, while a thick sequence of quartz-garnet-sillimanite and foliated garnet psammite (almandine) is developed in the hanging wall.

Mineralisation is crudely stratabound, occurring along the limbs of a large-scale, tight, recumbent D2 isoclinal synform with an easterly dip and a southerly plunge. The core of the synform is composed of amphibolites with encompassing silver-lead-zinc sulphide mineralisation. It is divided by faulting into a shallow, low-grade Northern Zone and a deeper, higher grade Southern Zone.

Within the Southern Zone, the isoclinal synform appears to control broad repetition patterns between ore lenses. Grade distribution within individual ore zones can also be related to zones of ductile strain and metasomatism influenced by strain partitioning around the termination of the Core Amphibolite. Within this Southern Zone, five main economic lode horizons and nine mineralisation types have been recognised. The mineralisation types are defined on the basis of distinctive zonations in Pb/Zn ratios, and Fe-rich versus siliceous gangue lithologies. The Fe-rich mineralisation types are characterised by coarse-grained, equigranular hedenbergite, Mn-Fe pyroxenoid, magnetite, olivine and fluorite mineralogies. Zones of extensive post-peak metamorphic metasomatism and retrogression contain assemblages of amphibole, almandine, ilvaite and pyrosmalite-dominant mineralogies with sulphide- and fluorite-rich ductile breccias. The siliceous mineralisation types represent late-stage metasomatism, with further modification of the mineralisation and retrogression of Fe silicates. These siliceous types have a distinctive low abundance of magnetite and fluorite (Walters and Bailey, 1998). Gangue minerals include pyroxmangite, manganese-fayalite, fluorapatite, fluorite and hedenbergite in the mafic associations, and blue-quartz, feldspar and carbonate in the siliceous lodes.

Dominant sulphides are galena and sphalerite, with multiple generations and variable intergrowth relationships. Subordinate magnetite-pyrrhotite with minor marcasite and arsenopyrite-lollingite-chalcopyrite are characteristic of the Fe-rich mineralisation types. Pyrite is generally absent and is only locally associated with late structural and low-temperature metasomatic overprints. All of the mineralisation types in the Cannington deposit show a consistent extreme Ag enrichment, occurring as argentiferous galena with freibergite inclusions. High levels of Sb, Cd, As, Cu and F are also a feature of specific mineralisation types. Magnetite is found in some lodes (Walters and Bailey, 1998).

The total resource in May 2007 (Bailey, 1998) comprised - 43.8 Mt @ 11.6% Pb, 4.4% Zn, 538 g/t Ag.
Production in 2003-04 totalled 64 183 tonnes of Zn, 263 305 tonnes Pb and 1206.364 tonnes Ag.

Reserve and resource figures as at 30 June 2007, published by BHP Billiton (2008) include:
    Total measured + indicated + inferred resource - 44 Mt @ 383 g/t Ag, 8.9% Pb, 4.2% Zn, including
    Total proved + probable reserve - 22 Mt @ 402 g/t Ag, 9.3% Pb, 4.1% Zn.

Remaining JORC compliant mineral resources as at 30 June 2015, published by South32 (2015) include:
    Measured resource - 47 Mt @ 201 g/t Ag, 5.53% Pb, 3.66% Zn,
    Indicated resource - 14 Mt @ 127 g/t Ag, 3.91% Pb, 2.81% Zn,
    Inferred resource - 10 Mt @ 82 g/t Ag, 3.00% Pb, 1.95% Zn,
  Total underground resource - 71 Mt @ 170 g/t Ag, 4.86% Pb, 3.26% Zn, including
  Open pit
    Measured resource - 13 Mt @ 90 g/t Ag, 3.66% Pb, 2.21% Zn,
    Indicated resource - 7.9 Mt @ 58 g/t Ag, 2.51% Pb, 1.83% Zn,
  Total open pit resource - 20.9 Mt @ 78 g/t Ag, 3.23% Pb, 2.07% Zn.

The mine was originally operated by BHP Billiton, but was included in the South32 demerger from BHP Billiton in 2015.

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McArthur River .......... Wednesday 9 November, 2005 .......... Travel by private air charter from Mt Isa.

The McArthur River or HYC stratiform, sediment-hosted zinc-lead-silver deposit is located 50 km SSW of Borroloola and 700 km SE of Darwin in the Northern Territory, Australia. The similar Teena deposit is located 8 km to the west (#Location: McArthur River - 16° 24' 41"S, 136° 5' 41"E).

  Mineralisation has been known in the McArthur River district and exploited on a small scale since around 1880 when galena veining within carbonate rocks was first reported. Sporadic, but largely unsuccessful exploration was undertaken in the area between 1891 and 1912. During the 1952-53 season, Consolidated Zinc Pty Ltd (CZC) investigated the Bald Hills (or Bulburra) and Cooks (or Coxco) prospects that are 5 km west and 10 km SSE of McArthur River respectively. The results of this work was discouraging. Both are developed within the Reward Dolostone that overlies the carbonaceous shale host to the McArthur River deposit. The Bald Hills occurrence comprises coarse galena and sphalerite filling vugs and fractures, whilst at Cook's, colloform and crystalline crusts of sphalerite, galena and pyrite-marcasite are developed on fragments in carbonate breccias, and as veins and matrix to a crackle breccia. An earlier 1948 visit to the Bald Hills prospect by Mount Isa Mines Limited (MIM) staff first interested the company in the area. In 1954, following a conceptual study of the prospectivity for 'Mt Isa-style' zinc-lead mineralisation in northern Australia, aided by C L Knight of CZC who had been seconded to MIM, a 640 x 160 km area was secured by MIM, encompassing the known lead occurrences in the McArthur River area.
  The HYC deposit was first indicated by the discovery in 1955 of a small outcrop of jasper containing acicular crystals of hemimorphite by an MIM field assistant while taking stream sediment samples. After testing more promising targets in the district without significant result, a final wrap-up program of 2 or 3 shallow drill holes was proposed in 1959 to test this enigmatic outcrop. This drilling revealed the outcrop to be part of an otherwise barren breccia bed 20 m above the flat lying, zinc- and lead-bearing pyritic shales of the HYC orebody, which is largely concealed below alluvial cover. The initial two drill holes ended in extremely fine grained sphalerite that was not initially recognised. When this sphalerite was identified, the drill rig was returned to deepen one of the holes, and indicating the extent of the mineralised interval (Logan et al., 1990). The deposit was subsequently delineated, but due to the very fine grained nature of the sulphides that precluded an adequate metallurgical recovery, was not immediately developed. After extensive testing, MIM started mining and processing ore from an approximately 1.5 million tonnes of ore per annum underground mine in 1995. In August 2005, McArthur River Mining (MRM) commenced a test open pit that was subsequently expanded into full production in April 2007, when underground mining ceased. Annual production of ore from the open pit is approximately 2.5 million tonnes.
  The Teena deposit was discovered by drilling during 2013 to 2015 by Rox Resources Limited and Teck Australia, based on follow-up of historic drill intersections from deep reconnaissance drilling undertaken by Mount Isa Mines between 1976 and 1978 to test the potential of the Teena Sub-basin. These historic holes included an intersection of 9.2 m @ 3.5% Zn from 629.2 m, and two unreported follow up drill holes that had intersected narrow intervals of low-grade, broadly stratiform mineralisation. All were on the fringes of the core of the Teena sub-basin which is indicated by modern geophysics and was found to be the axis of the deposit.

The regional setting of the district and the geology and mineralisation of the McArthur River deposit are as follows, with descriptions on Teena and the other nearby deposits appended below that.

Regional setting

  The McArthur River deposit lies within the McArthur Basin, part of the NW-SE trending Carpentaria Zinc Belt, which extends for >1200 km, from Mount Isa in Queensland to the south, to Arnhem Land in the Northern Territory in the north.
  The Palaeo- to Mesoproterozoic McArthur Basin contains a 5 to 10 km package of mostly unmetamorphosed sedimentary and volcanic rocks deposited between ~1800 and 1575 Ma, and covers an area of ~180 000 km
2. This succession unconformably overlies 1890 to 1820 Ma metamorphosed and deformed igneous and metamorphic basement rocks of the Pine Creek and Arnhem Province to the NW and NE repectively, and the 1860 to 1840 Ma Murphy tectonic ridge to south. The latter was likely a palaeogeographic high, partially separating the McArthur Basin from the South Nicholson Basin and Lawn Hill Platform of the Isa Superbasin in Queensland (Plumb and Wellman 1987; Wygralak et al., 1988; Ahmad et al., 2013). Palaeozoic and younger sedimentary sequences of the Georgina, Arafura and Carpentaria basins unconformably overlie the McArthur Basin succession to the SW, north and SE repectively. The McArthur Basin may be continuous, beneath the Georgina basin, with the Tomkinson Basin in the Tennant Creek area, >300 km to the SW (Ahmad et al., 2013).
  The McArthur Basin is punctuated by two main, 50 to 100 km wide and ~300 km long asymmetric zones of faulting, the Walker and Batten fault zones that lie along the same arcuate, generally north-south trend, but are separated and marginally offset by the ~20 km wide, ESE-trending Urapunga Fault Zone. In detail, the Walker and Batten fault zones to the north and south, trend to the NNE and NNW respectively, curving to north-south towards their intersections with the Urapunga Fault Zone. This latter cross-cutting fault zone also divides the larger basin into the northern and southern McArthur basins. A deep seismic reflection survey (Rawlings et al., 2004) showed that the entire succession within these basins is essentially horizontal, with a thickness of ~8 km, that shows no significant variation on either side of, or within, the Walker and Batten fault zones (Rawlings 1999; Rawlings et al., 2004; Ahmad et al., 2013). This is contrary to the long held interpretation that these two zones represented grabens, the Walker and Batten troughs, flanked by thinner shelf sequences (e.g., Plumb and Derrick 1975; Plumb et al., 1980, 1990, Plumb and Wellman 1987).
  The McArthur River deposit is located immediately to the west of the major Emu Fault on the eastern margin of the Batten Fault Zone in the Southern McArthur Basin (Ahmad et al., 2013).
  The sequence within the McArthur Basin is part of the broader Mt Isa-McArthur succession, which is, in turn, part of the Northern Australian Platform cover. This 5 to 15 km thick volcano-sedimentary pile was deposited during the period 1800 to 1580 Ma in an intracontinental setting. Deposition took place in three super-basins which represent three nested cycles of deposition and exhumation, specifically the Leichhardt (1798 to 1738 Ma), Calvert (1728 to 1680 Ma) and Isa (1667 to 1575 Ma) super-basins, terminated by the 1590 to 1500 Ma Isan Orogeny, which was followed by the fourth and younger, ~1500 to ~1400 Ma Roper super-basin (Jackson et al., 1999, 2000; Betts et al., 2003; Withnall and Cranfield, 2013; Stewart, GeoScience Australia 2015). Each of these super-basins corresponds to a period of extension, and each is ended by a basin inversion, although less intense inversions are also recorded within the duration of these super-basins, affecting the stratigraphic packages deposited within them.
  All of the major stratiform Zn-Pb-Ag deposits of the Carpentaria Zinc Belt, including McArthur River, are hosted by 2 to 8 km thick successions of the Isa Superbasin, ranging in age from ~1660 to 1650 Ma (Dugald River, Lady Loretta), ~1650 Ma (Mt Isa), ~1640 (McArthur River) to 1595 Ma (Century; Queensland Department of Mines and Energy, 2000).
  Rawlings et al. (1997) and Rawlings (1999) divided the McArthur Basin succession into five basin-wide depositional 'packages', each of which is disconformity or unconformity bounded and characterised by similarities in age, stratigraphic position, lithofacies composition, style and composition of volcanism, and basin-fill geometry across the McArthur Basin. These packages can be related to the super-basins of Jackson et al. (1999, 2000).
  Within the Southern McArthur Basin, the packages, stratigrapic units and super-basins are as follows:
Redbank package which comprises the Tawallah Group and corresponds to the Leichhardt and Calvert super-basins succession. The Tawallah Group unconformably overlies basement metamorphic rocks of the east-west trending Murphy tectonic ridge, with the Leichhardt super-basin largely not being represented. Maximum and minimum ages are constrained by the basement (Scrutton Volcanics: 1850 Ma; Pietsch et al., 1991) and overlying McArthur Group ~1660 to 1610 Ma; Munson, 2019).
  The lower Tawallah Group comprises widespread fluvial to intertidal sandstones and conglomerates, and extensive flood basalts which were deposited between ~1790 and 1760 Ma (e.g., Yiyinti Sandstone, Westmoreland Conglomerate, Seigal Volcanics; Ahmad et al., 2013). These were succeeded between ~1760 and 1740 Ma by the middle Tawallah Group, composed of shallow marine, syn-extensional siliciclastic rocks overlain by post-extensional carbonate rocks (Ahmad et al., 2013) and a 1740 Ma mid-basin inversion that produced a regional unconformity at the base of the Wununmantyala Sandstone (Bull and Rogers, 1996). This unconformity and inversion event correspond to the boundary between the Leichhardt and Calvert super-basins (Blaikie and Kunzmann, 2020). Renewed extension from ~1730 to 1690 Ma produced widespread sedimentation and volcanism across the North Australian Craton. In the southern McArthur Basin, shallow marine to offshore siliciclastic rocks and minor carbonates of the upper Tawallah Group were deposited on an extensive low-relief platform, which included the Wollogorang Formation (Kunzmann et al., 2020), and the regionally extensive Settlement Creek Volcanics, which were predominantly emplaced as high-level sills, with the Gold Creek Volcanics inferred to be their eruptive equivalent (Ahmad et al., 2013; Rawlings, 2006). Available SHRIMP U-Pb zircon dates from upper Tawallah Group rocks near the centre of the succession range from 1708 to 1730 Ma (Page et al., 2000; Page and Sweet, 1998; Rawlings, 2002).
Goyder package of the upper Calvert super-basin, which is largely absent from the Southern McArthur Basin, but may include some sandstones included in the upper Tawallah Group (Ahmad et al., 2013; Rawlings, 1999).
Glyde package, which corresponds to the lower Isa super-basin. Within the Southern McArthur Basin, this package is represented by the McArthur Group that hosts the McArthur River deposit. In contrast to the unconformably underlying Tawallah Group, that is predominantly arenaceous, the up to 5 km thick McArthur Group comprises a succession of platformal stromatolitic dolostone and clastic sedimentary rocks with local pyritic and carbonaceous siltstone units (Winefield 1999). Deep seismic reflection data shows that in its central section, the McArthur Group gradually and systematically increases in thickness from ~1300 m to ~3200 m west to east respectively (Rawlings et al., 2004). Although exposures of the McArthur Group are confined to Batten Fault Zone, seismic data indicates that it continues at depth beyond the inferred limits of that fault zone (Rawlings et al., 2004). However, the one exception to the systematic westward thinning of the McArthur Group is around the Emu Fault, where the thickness appears to decrease to 2500 m immediately east of, and to >2700 m immediately west of the fault, whilst within the upward fanning flower structure of the Emu Fault, the thickness appears to locally increase to >3000 m.
  The McArthur Group is subdivided into the Umbolooga and overlying Batten subgroups, separated by a possible disconformity. The lower half of the Umbolooga Subgroup is characterised by alternating thick (200 to 650 m) units of sandstones, dolostones and dolomitic arenites. This succession fines upwards to dolostones and dololutites, including the 10 to 900 m thick, 1640±4 Ma Barney Creek Formation, composed of finely bedded to laminated, dolomitic, carbonaceous and pyritic siltstone, shale and dololutite, with locally abundant tuff beds and breccias, that hosts the McArthur River deposit within the HYC Pyritic Shale Member. This unit is underlain by the dololutites of the Teena Dolostone and is overlain by the uppermost unit of the Umbolooga Subgroup, the 30 to 350 m thick Reward Dolostone composed of dololutite, silty dololutite and lesser bedded to cross-bedded dolarenite. A local inversion and reactivation of faulting is recorded at ~1640 Ma, close to the time of deposition of the Barney Creek Formation and the McArthur River deposit. This event followed a longer interval of north-south directed extension and was succeeded by a return to extension. The overlying Batten Subgroup is largely composed of shallow marine to tidal and emergent dolostones, dololutite, dolomitic siltstones and lesser dolarenites, and a few beds of quartz sandstone. The uppermost unit of this subgroup has been dated at 1614±4 Ma (Ahmad et al., 2013).
Favenc package, which corresponds to the upper Isa super-basin. Within the Southern McArthur Basin, this package is represented by the up to 1600 m thick Nathan Group. This sequence, which unconformably overlies the McArthur Group, was deposited in shallow marine to terrestrial setting, and comprises a relatively thin and lenticular basal siliciclastic and often conglomeratic unit, overlain by thicker carbonate and siliciclastic rocks containing tuff beds that have yielded a date of 1589±3 Ma. The Nathan Group includes the quartz arenite and dolarenite with cross-beds, ripple marks and commons gypsum and halite casts of the widespread, up to 700 m thick Karns Dolostone (Ahmad et al., 2013).
Wilton package of the Roper super-basin, represented by the widespread Roper Group which covers an area of ~145 000 km
2 and comprises an upward-coarsening cyclic succession of mainly marine mudstone alternating with sandstone, with minor micritic and intraclastic limestone, and ooidal ironstone (Ahmad et al., 2013).

District and deposit geology

  Mineralisation at the McArthur River deposit is hosted by HYC Pyritic Shale Member lithofacies of the ~1640 Ma Barney Creek Formation, which occurs towards the top of the Umbolooga Sub-group of the McArthur Group. The immediate host sequence was deposited within a tectonically induced sub-basin that is bounded to the east by the major, north-south, syn-depositional, Emu Fault corridor, which lies on the eastern margin of the broader Batten Fault Zone.
  The Emu Fault is a wide, generally NNW trending corridor of faulting, that is commonly 1 to 2 km wide, but is up to 5 km in width when the Western fault splay to the south of McArthur River is included. It has been interpreted, from deep seismic data, to be an upward fanning 'flower structure'-like zone of dislocation, including east dipping to vertical strike-slip, and west vergent reverse faults, representing repeated activity over a long period, prior to, during and after mineralisation at McArthur River. The overall fault zone has a steep west dip, with only minor net displacement of the McArthur Group, generally of <500 m. Strike-slip displacement on the Emu Fault prior to deposition of the Roper Group was most likely sinistral, including during deposition of the McArthur Group, whilst the youngest movement appears to have been dextral (Rawlings et al., 2004; Rogers, 1996; Hinman, 1995).
  In contrast to the steep, overall strike-slip Emu Fault on the eastern margin of the broader Batten Fault Zone, the main, generally north-south structures toward the western margin, are east vergent thrust faults that post-date Roper River Group deposition and are splays off a deep detachment. The principal of these are the Tawallah and Scrutton faults that are ~35 to 50 and ~55 km to the west of the Emu Fault respectively. The sinuous Tawallah Fault merges with the Scrutton Fault to the south, and juxtaposes mainly Tawallah Group on the west/upper plate side over middle to upper McArthur Group to the east. In the north, the Tawallah Fault has an apparent 12 km of vertical displacement (Rawlings et al., 2004).

McArthur River Geology
  The stratigraphy of the Umbolooga Sub-group, the lower of the two sub-groups that constitute the McArthur Group, may be summarised as follows, from the base (after Ahmad et al., 2013):
Masterton Sandstone, 40 to 650 m thick - a succession of fluvial to shallow marine, white, pink and red sandstone, minor mudstone and conglomerate that unconformably overlies various formations of the Tawallah Group.
Mallapunyah Formation, 100 to ~450 m thick - mudstone, siltstone, dolostone and sandstone, interpreted to have been deposited in a shallow-marine to sabkha environment, with pseudomorphs after halite and gypsum, and well developed diagnostic botryoidal quartz nodules (cauliflower chert).
Amelia Dolostone, 50 to 180 m thick - partly stromatolitic dolostone with local interbeds of dololutite containing diagenetic siderite.
Tatoola Sandstone, 80 to 350 m thick - comprising a lower, fine-grained, thinly bedded sandstone facies, deposited in a subtidal, deeper marine, storm-affected environment, and an upper, medium to coarse-grained, more thickly bedded sandstone and dolostone facies, interpreted to have been deposited in agitated shallow-water to intermittently emergent environment (Jackson et al., 1987).
Tooganinie Formation, ~200 m thick - a thick succession of dolostone and interbedded dolomitic sandstone, siltstone and shale.
Leila Sandstone, <10 to 30 m thick - shallow marine, fine to coarse-grained, dark grey weathering dolomitic sandstone with common cross bedding and ripple marks.
Myrtle Shale, 40 to 60 m thick - thinly bedded to laminated, lagoonal, commonly dolomitic siltstone, shale and fine-grained sandstone with common halite casts, and local stromatolitic dololutite.
Emmerugga Dolostone, 620 m thick - interpreted to have been deposited in a peritidal to deep marine environment on an extensive, relatively flat, intracratonic carbonate platform. It comprises a series of upward-shallowing cycles that deepen and thicken up-section (Winefield 1999), and has been subdivided into the:
  - Mara Dolostone Member - that comprises shallow marine to sabkha environment silicified stromatolitic units, alternating with silty or non-stromatolitic units, comprising dololutite, dolomitic siltstone, dolarenite and dolomitic breccia.
  - Mitchell Yard Dolostone Member - a non-stromatolitic, massive, grey featureless crystalline dolostone unit.
The package from the Myrtle Shale to Emmerugga Dolostone has been interpreted to represent a series of partly laterally equivalent and partly vertically distributed facies within what Winefield (1999) refers to as the Emmerugga Depositional Cycle, the first of three such cycles within the mid McArthur Group. They represent a phase of increasing accommodation within the host sub-basin, during which platformal carbonate lithofacies accumulated in a series of upward-shallowing cycles that thicken and deepen up-section, i.e., the deeper-water shale facies are regarded as being partially lateral equivalents of the shallow-water carbonate facies on the margins of the sub-basin (Winefield 1999).
Teena Dolostone, interpreted to have been deposited during basin subsidence, but infilled at a rate that resulted in upward shallowing to oolitic and oncolitic grainstone, imbricated flat-plate breccias, ripple-marked mixed carbonate-siliciclastic rocks, finely laminated dololutite and thin interbeds of dolomitic shale, siltstone, dolarenite and rare thin beds of K-metasomatised mudstone. It has been subdivided into the:
  - Lower Teena Dolostone, which is <60 m thick - thinly bedded to laminated dololutite, that is locally silicified, with dolomitic shale and sandstone, intraclast breccia and conglomerate, and dolarenite, deposited in a marine setting.
  - Coxco Dolostone Member, 15 to 70 m thick - composed of grey crystalline dololutite with rare conical stromatolite, differentiated on the basis of its radiating, needle-like aragonite accumulations developed normal to bedding, variously interpreted to have pseudomorphs of sea-floor (Walker et al., 1977), sea-floor aragonite (Brown et al., 1978), lacustrine trona (Jackson et al., 1987), and an abiogenic aragonite sea-floor cement (Winefield 1999). The latter author suggested they were sea-floor cements generated by upwelling of HCO3-, Fe2+ and Mn2+-rich anoxic bottom water during rapid changes in basin bathymetry.
McArthur River Stratigraphy   The contact between the Teena Dolostone and overlying Barney Creek Formation varies from a sharp transition, to a zone of brecciation, fracturing and fissuring of the Coxco Dolostone Member just below the contact, with sedimentary infill of open space.
Barney Creek Formation, 10 to 900 m thick - a largely recessive unit composed of finely laminated to thinly bedded dolomitic, carbonaceous and pyritic siltstone, shale and dololutite, with locally abundant tuff beds and rare breccias. Jackson et al (1987) defined Three members, or more accurately lithofacies, have been defined. These area the:
  - W-Fold Shale Member - which is composed of green and red dolomitic siltstone and shale, and green vitric tuff. Much of the siltstone has apparently been subjected to potassic alteration.
  - HYC Pyritic Shale Member - a recessive, thinly bedded, laminated, carbonaceous (to bituminous), dolomitic pyritic siltstone, sedimentary breccias and tuffaceous mudstone. The siltstones are composed of quartz, potassium feldspar, albite, ferroan dolomite, calcite, chlorite, illite and kaolinite, and have an average grain size between 15 and 20 µm. The sedimentary breccias, which are fingers of the laterally equivalent Cooley Dolostone (see below) range from chaotic slump breccias to well graded silt beds. The tuffaceous mudstone beds are up to several metres thick, comprising quartz, potassium feldspar, albite, dolomite and calcite being the major constituents, and have yielded dates of 1640±4, 1639±3 and 1638±7 Ma (Logan et al., 1990; Page and Sweet 1998; Page et al., 2000; SHRIMP zircons from tuffaceous sediments).
  - Cooley Dolostone Member - comprises one or more, locally-derived, overlapping, chaotic, fault-related intraformational slump breccias developed within the sub-basin on the western side of, and partially over, the uplifted Western Fault Block of the Emu Fault corridor. These breccias occupy the eastern fault controlled slope of the sub-basin and interfinger with the other lithofacies of the Barney Creek Depositional Cycle to the west. Clasts within these breccias range from millimetres to several tens of metres across, and were largely sourced from the Emmerugga and Teena dolostones (Pietsch et al., 1991) that are exposed in the uplifted Western Fault Block (Hinman, 1995). These breccias and megabreccias are also locally found at the upper contact of the Coxco Dolostone Member.
  There are three types of breccia, apparently related to the Cooley Dolostone Member lithofacies that interfinger with the HYC Pyritic Shale Member lithofacies (Walker et al., 1977). Type I breccias are the most abundant and are intimately related to the stratiform ores. These were apparently slumped into the sub-basin from the NNW from a very similar location and flow direction to the postulated ore fluids detailed later. Types II and III are localised in the upper stratigraphy of the Barney Creek Formation and are considered to have a different provenance and direction of movement into the sub-basin, with type II entering the basin from the SE (Walker et al., 1977). The type I sedimentary breccias and related dolomitic sediments are interpreted to be slump deposits and graded turbidites caused by deep seismic events and related movements along growth faults at the basin margin, including the NNW-SSE Emu Fault (e.g., Williams, 1978) or NW-SE to WSW-ESE trending faults to the north of the deposit (e.g., Walker et al., 1977; Logan, 1979).
  Within the main depocentre of the sub-basin, the ~80 m interval of the Barney Creek Formation above the W-Fold Shale is predominantly composed of HYC Pyritic Shale, with intercalations of tuffaceous mudstone beds and breccias, as described above. In this interval a suite of eight 2 to 12 m thick HYC Pyritic Shale Member lenses host the main ore deposit, and are separated by 2 to 5 m thick beds of sedimentary breccia related to the Cooley Dolostone Member. Above this interval, the remainder of the Barney Creek Formation is occupied by more 'thickly' intercalated beds of black dolomitic siltstones of the HYC Pyritic Shale and mass flow breccias of the Cooley Dolostone, with individual bands of each being 40 to 80 m thick. The black 'shales' in this upper interval only contain low-tenor Zn-Pb mineralisation. The differences in thickness and distribution of these breccias and intercalated 'shales' may be related to a changed tectonic regime and nature of seismic activity that also resulted in different breccia styes (types II and III) found in the upper part of the sequence, their provenance and rate of deposition. Type II breccias are inferred to have been emplaced from the SE corner of the sub-basin with a flow direction towards the NW, emanating from near the Ridge II discordant carbonate hosted and stratiform mineralisation (Large et al., 1998; see the Ridge II description below).
  Large et al. (1998) suggest that towards the close of the Barney Creek Formation deposition, the tectonic regime changed from extension to transpressive inversion (Hinman et al., 1994), with only minor iron-rich and reduced basinal brines contributed iron but only minor amounts of base metals into the sub-basin. They further suggest this inversion was accompanied by basin shallowing, and then deepening, corresponding to the concordant Ridge II Pb-Zn mineralisation, which is >300 m stratigraphically above the HYC ore, towards the top of the Barney Creek Formation (see the Ridge II description below), and finally, shoaling to the relatively shallow water Reward Dolomite facies (Bull, 1998).
Reward Dolostone, 30 to 350 m thick - comprising dololutite, silty dololutite, minor dolarenite, shale and local dolostone breccia, interpreted to have been deposited in an environment that ranged from peritidal, shallow subtidal to shallow open sea. The contact with the underlying Barney Creek Formation is generally conformable and gradational, although possible palaeo-regolith has been interpreted locally (Haines et al., 1993). The upper contact varies from being a transition from dolostone to siltstone and shale of the Caranbirini Member of the Lynott Formation, to disconformable where possible palaeoregolith is developed at the top of the Reward Dolostone (Haines et al., 1993, Pietsch et al., 1991).
  The Barney Creek Formation has been interpreted by Bull (1998) and Winefield (1999) as predominantly a deeper-water shaly carbonate facies that was deposited during maximum flooding related to a sea-level transgression. They suggest the carbonaceous pyritic shales are indicative of a quiet, anoxic, sub-wave base environment. These facies were deposited following an abrupt basinward shift at the top of the Emmerugga Depositional Cycle, marking the onset of tectonically-induced basin subsidence accompanied by what Winefield (1999) refers to as the Barney Creek Depositional Cycle. This latter cycle includes the Teena Dolstone, Barney Creek Formation and Reward Dolostone, interpreted to represent a series of partly laterally equivalent and partly vertically distributed facies, characterised by rapid lateral lithofacies variation. These lithofacies were deposited in numerous small sub-basins adjacent to reactivated, generally north-south trending faults (e.g., the Emu fault). According to Winefield (1999), facies architecture within the Barney Creek Depositional Cycle indicate that the thickest shale dominated facies appear to have been deposited and preserved adjacent to NNW to NW trending growth faults. Conversely, these lithofacies successions grade laterally into condensed sections adjacent to NNE to NE trending fault segments and laterally away from the thicker sub-basins. Winefield (1999) suggests this facies architecture is consistent with differential strike-slip movement along sinuous, generally north-south faults producing transtensional sub-basins adjacent to NNW to NW fault segments, and transpressive sub-basins or platform margins along NNE to NE trending fault sections. Adjacent to these sub-basin margins, gravitationally unstable slopes developed, characterised by coarse-grained clastic slope lithofacies and abundant slope-related deformation which include sediment (neptunian) dykes, liquifaction breccias, megabreccias and soft-sediment slump folds (Winefield 1999). The chaotic intraformational breccias of the Cooley Dolostone Member, which are developed along section of the Emu Fault and interfingers with the HYC Pyritic Shale Member away from the fault, represents one of the best examples of these slope lithofacies, as described above.
  The Umbolooga Sub-group is overlain by the Batten Sub-group, above an implied disconformity, as indicated by the interpreted partial palaeo-regolith described above. The lower units immediately following the Barney Creek Depositional Cycle are the:
Lynott Formation, 50 to 600 m thick - an evaporite-rich unit, mainly composed of dolomitic siltstone, dolarenite, stromatolitic dolostone and lesser dolomitic sandstone, split into three conformable members (Jackson et al., 1987), the:
  - Caranbirini Member, 0 to 400 m thick - thinly bedded, laminated fine-grained mudstone and dolomitic mudstone, which is locally a carbonaceous and/or pyritic, dololutite, with intraclast breccia and pink tuffaceous mudstone.
  - Hot Spring Member, 50 to 350 m thick - which is typically and mostly composed of dolomitic siltstone, silty dololutite and sandstone, with ubiquitous evaporite pseudomorphs, suggesting deposition on intratidal to supratidal flats (Pietsch et al., 1991).
  - Donnegan Member, 0 to 134 m thick - thinly to medium bedded, commonly rippled and cross-bedded red-brown to buff dolomitic siltstone and sandstone, with lesser dolarenite and thin to medium beds of chert altered stromatolitic dolostone. It is characterised by abundant small botryoidal quartz nodules, cauliflower and enterolithic chert, considered to be consistent with deposition in a supratidal to sabkha environment. Dating of a tuff band near the top of the member returned a date of 1636±4 Ma, within error of the Barney Creek Formation (Page et al., 2000).
Yalco Formation, <50 to 250 m thick - which conformably overlies the Donnegan Member and includes silicified, thinly bedded and interbedded, stromatolitic dololutite, silty dololutite, dolarenite and dolomitic sandstone, containing abundant intraclast chert breccias, laminae and nodules. Ripple marks, desiccation cracks, tepee structures and casts and moulds after evaporite minerals are evident, suggesting deposition in a shallow-marine environment that was subjected to periodic emergence.
  The Lynott Depositional Cycle, which extends from the Caranbirini Member to the top of the Yalco Formation, overlies the Barney Creek Depositional Cycle and is interpreted to represent a phase of renewed basin subsidence. According to Winefield (1999), the spatial distribution of the lithofacies of this new cycle is distinctly different to that of the previous, suggesting a subtle but important shift in the controlling structures.
Stretton Sandstone, <55 to 270 m thick - which conformably overlies the Lynott Depositional Cycle, and comprises a 1625±2 Ma sequence of thinly bedded, shallow marine to sub-wave base quartz arenites.
Looking Glass Formation, 30 to 70 m thick - silicified, peritidal to shallow marine stromatolitic dolostone and dolarenite.
Amos Formation, <70 m thick - which marks the top of the sub-group, composed of siltstone, sandstone and crystalline dolostone, representing red beds and palaeo-calcrete, dated at 1614±4 Ma.

Mineralisation and alteration

  The McArthur River mineralisation is hosted within fine grained HYC Pyritic Shale Member lithofacies, and comprises 8 individual ore lenses that are from <2 to >12 m thick, each of which is separated by <2 to 5 m thick bands of interfingering Cooley Dolostone type sedimentary breccia, and lesser poorly bedded or fragmented pyritic and dolomitic shale interbeds. The deposit covers an area of ~2 km
2, and has an average composite thickness of 55 m. Its eastern margin is ~1.5 km west of the main Emu Fault, separated by the Cooley Dolostone breccia lithofacies and the Western Fault block, an up-faulted horst of sedimentary rocks older than the HYC Pyritic Shale Member within the broader Emu Fault corridor (Ahmad et al., 2013; Ireland et al., 2004; Logan et al., 1990). The western margin of the deposit has been folded and eroded. The northern margin interfingers with sedimentary slump breccia derived from older McArthur Group sediments to the north of the deposit, while the southern boundary is gradational into barren pyritic siltstone of the condensed Barney Creek Depositional Cycle (Walker et al., 1977).
  The stratiform ore lenses are characterised by regular, fine-grained bands of sphalerite and galena interlayered with beds of dark bituminous and tuffaceous shale on both the macro- and micro-scales (Croxford and Jephcott, 1972). The sulphide layers comprise a complex mixture, rather than mono-minerallic bands (Eldridge et al., 1993), although many of the metal sulphide rich bands are dominated by sphalerite and/or galena (Large et al., 1998). Whilst there are many different textural and compositional variations within these ore lenses, Large et al. (1998) recognised three end-member microband types:
• organic-rich mud layers, that are commonly <500 µm thick, are principally composed of illite, carbonaceous matter, pyrite and quartz, and are interpreted to be the product of normal pelagic sedimentation from the anoxic water column.
• quartz-illite-carbonate silt layers, which are usually 100 to 500 µm thick, although laminae up to several cm thick are also present. These thicker bands are normally graded, consistent with deposition by mass flow. Croxford and Jephcott (1972) regarded these bands to be turbidites.
• zinc-lead sulphide-rich layers, that are normally 200 to 500 µm thick and composed of uniform mixtures of sphalerite-galena±pyrite with lesser ankerite, quartz and minor chalcopyrite and arsenopyrite. Sphalerite is the dominant sulphide mineral, with sphalerite:galena ratio of from 2:1 to 12:1, based on observations of samples from ore lens 2.
  Whilst it is common for the three end-member microband types to be regularly interlayered, mixtures of these end members may also occur, e.g., layers composed of a mixture of organic muds and zinc-lead sulphides or zinc-lead sulphides and quartz-carbonate silts (Large et al., 1998).
Major sulphides Two generations of pyrite have been identified. The earlier, Py
1 comprises >80% of the pyrite at McArthur River and is present as 1 to 15 µm sized euhedral to subhedral, octahedral and cubic crystals, and as both loosely packed spherical framboids and tightly packed round to ellipsoidal macroframboids (Croxford and Jephcott, 1972). Py1 is a minor component of the entire deep-water facies of the Barney Creek Formation, but is abundant (5 to 60 vol.%) in and around the main McArthur River deposit.
1 is overgrown by Py2, which also occurs interstitially to the earlier pyrite crystals/spherules (Rye, 1974; Williams, 1978; Large et al., 1998) and is composed of fine to coarse (10 µm to 1 mm) anhedral to subhedral crystals, crystal aggregates and blebs, as well as the Py1 overgrowths. SHRIMP ion microprobe determinations of the sulphur isotope compositions of pyrite from the ore lenses has produced δ34S values for Py1 and Py2 that range from -13 to +15 and -5 to +45‰ respectively (Eldridge et al., 1993). This was interpreted by the same authors to indicate growth of pyrite likely involved H2S formed by biogenic processes. On average, the δ34S values of Py2 are about 15‰ higher than those of Py1, suggesting that the pyrite types may have formed sequentially from a single batch of sulphate in an environment which became closed with respect to sulphate supply through time. This has been taken to indicate an early biogenic diagenetic Py1 derived from sea water sulphate, and a later deeper diagenetic pyrite without direct access to sea water, only sulphate trapped in the host sediment (Eldridge et al., 1993).
  The occurrence of pyrite and zinc-lead sulphides differs, in that pyrite does not form discrete continuous layers, as do the zinc-lead and sediment bands, but instead occurs as patchy disseminations. Fine grained pyrite is disseminated in most, but not all, zinc-lead layers, within organic mud bands, or rarely, as discontinuous layers within quartz-carbonate silt bands. The more prominent coarse-grained pyrite, comprising irregular masses from 100 µm to >1 mm across, is common, erratically distributed within either the organic-rich mud layers or the zinc-lead bearing bands (Large et al., 1998).
  Sphalerite occurs as near monominerallic layers up to 1 mm thick, as elongate accumulations up to 0.2 mm in diameter, and as fine grained disseminations, whilst galena is also observed as near monominerallic layers up to 0.05 mm thick, as streaks in sphalerite layers and as fine-grained disseminations. Chalcopyrite is present as minute elongate inclusions in sphalerite (Ahmad et al., 2013).
  The differences between pyrite and base metals sulphides are also reflected by conventional sulphur isotope analyses of sulphide separates, which indicated that δ
34S values of pyrite (Py1 + Py2) increase from -3.9 to 14.3‰ from the footwall to hanging wall, whereas those of galena and sphalerite are constant at -1.2 to +5.7 and 3.3 to 8.9‰, respectively (Smith and Croxford 1973). These data are supported by Eldridge et al. (1993) who found that δ34S for base metal sulphides (sphalerite, galena and chalcopyrite) is restricted to a range from -5 to +14‰ and excluding the few late-stage veins, is further limited to a range of -5 to +8‰. Eldridge et al. (1993) interpreted these data to suggest the formation of the base metal sulphides was distinct from that of Py1 and Py2 precipitation, and that localisation of the base metal sulphides did not involve biogenic sulphide either directly through use of residual microbially generated H2S(aq.) or by replacement of Py1 or Py2. Eldridge et al. (1993) noted that while the ratio of Py1:Py2 fluctuates considerably, the temporal relationship is invariant, and the author's had not observed either pyrite type to have nucleated on, or to have overgrown, sphalerite or galena. These mutual textural and isotopic relationships they interpreted to suggest a paragenetic sequence at any one point or bed in the deposit, in which Py1 was followed by Py2, with galena and sphalerite forming later than pyrite (Eldridge et al., 1993; Ahmad et al., 2013).
  However, Ireland et al. (2004) observe that Py
2 nucleates in, permeates, and locally expands sphalerite rich laminae with associated recrystallisation of sphalerite to monominerallic aggregates free of inclusions (Large et al., 1998). It also displaces or overprints all other minerals associated with mineralisation at McArthur River and hence Ireland et al. (2004) consider it to be the latest sulphide phase in the paragenetic sequence.
  Eldridge et al. (1993) interpreted three modes of sphalerite-galena occurrence within the ore lenses:
• Mode 1 - fine-grained zinc-lead bands of sphalerite-galena distributed parallel to bedding. This mode constitutes >80% of the sphalerite-galena in the ore lenses. EIdridge et al. (1993) interpreted this sphalerite-galena to have formed as a grain cement within siltstone layers, enclosing many non-sulphide grains, initially by pore-space filling, but eventually leading to total replacement of the siltstone layers. Large et al. (1998) observe that these sulphide bands are intricately interlayered with, and have sharp conformable boundaries with both pelagic mudstone beds and quartz-carbonate turbidites. If the sphalerite-galena bands were purely replacive, the more reactive quartz-carbonate turbidite bands would be expected to be preferentially replaced, or at least have gradational boundaries with adjacent zinc-lead sulphide layers. They also point out that clasts within turbidite layers overlying sphalerite-galena bands cause compaction load features, while flame structures composed of sphalerite-galena extend up into overlying mass flow units indicating the presence of those sulphide bands before the turbidite introduction. In addition, the coarser sphalerite-galena bands contain physically rounded and subrounded clasts of base metal sulphides and pyrite, while layers composed of base metal sulphide microbreccias are common, comprising rounded to angular clasts of base metal sulphides in a matrix of mixed sediment and disseminated sulphides. Banded base metal sulphide clasts, from 1 to 70 cm across, have been observed in the inter-ore breccias (e.g., Croxford, 1968; Croxford and Jephcott, 1972; Scott and Lambert, 1979; Williams, 1979; Coutts, 1996), texturally identical to the banded zinc-lead sulphide ores, and these clasts appear to have undergone soft sediment deformation during incorporation into the breccia beds. Large et al. (1998) argue all of these observations suggest sphalerite-galena precipitation was either syn-depositional, or very early diagenetic, during deposition.
  Sphalerite within Mode 1 sulphide bands, Sp
1 of Ireland et al. (2004), is a very fine grained (1 to 10 µm) cryptocrystalline variety occurring as irregular, elongate and blebby 2 to 200 µm thick aggregates that coalesce to define sphalerite-dominated laminae that are 0.2 to 1 mm thick within carbonaceous siltstone. This sphalerite is characteristically associated with euhedral 100 to 500 µm ankerite crystals and euhedral 10 to 200 µm quartz crystals, and at a microscopic scale, bands of this sphalerite contain silt clasts, organic matter and other gangue minerals. Sp1 bands may include up to 50 vol.% Py1, and may coexist with and locally contain up to 50% euhedral quartz and ankerite crystals, but is commonly zoned from quartz-ankerite bearing Sp1 at the base, to quartz-ankerite poor varieties with abundant enclosed Py1 at the top. Laminations of Sp1 are locally disrupted by nodular carbonate developments and Py2 aggregates, suggesting Sp1 is a paragenetically early phase. The intimate relationships between Sp1, ankerite, and/or quartz and Py1 are considered by Ireland et al. (2004) to indicate co-deposition of these minerals, with quartz and ankerite forming immediately after the deposition of each Sp1 lamination.
  Zinc-lead sulphides have been considered by most authors to post date pyrite (e.g., Williams, 1978; Eldridge et al., 1993; Large et al., 1998), because Py
1 occurs as inclusions in all other minerals related to mineralisation. However, Ireland et al. (2004) note there is a consistent and complex relationship between Py1 and Sp1, which they imply, suggests that the formation of these minerals was contemporaneous.
  Laminated early Sp
1 has a δ34S values that range from 02.41 to ~12.1, and a mean of 3.8‰, zoned from 4 to 6 in the centre of the deposit, to 0 to 3‰ on the extreme lateral fringes to the north and south.
• Mode 2 coarser grained, patchy sphalerite-galena, intergrown with carbonate nodules, and within nodular carbonate layers (Eldridge et al., 1993; Large et al., 1998). Sphalerite associated with mode 2 zinc-lead sulphides, Sp
2, comprises 15 to 20% of all sphalerite at McArthur River, and may locally constitute up to 100% of the sphalerite in zones of abundant carbonate nodules. It comprises 1 to 10 mm crystalline aggregates that are associated with nodular, micritic or clastic carbonate, and usually occurs as variable replacement of sparry to micritic carbonate aggregates that are 0.5 to 4 mm thick. According to Ireland et al. (2004), Sp2 was formed late in the paragenetic sequence, following carbonate nodules that displace Sp1 bands, but are overprinted by Py2. These carbonate nodules and clasts usually have a rim of Sp2 and retain a core of carbonate, although they may locally be completely replaced or have internal textures selectively replaced by Sp2. Sp2 domains are usually virtually monominerallic with regard to sphalerite and devoid of inclusions, other than Py1, precursor carbonate and very rare euhedral quartz and ankerite. Sp2 and diffuse Sp1 that occur in adjacent laminae retain their textural differences and are both overprinted by Py2 (Ireland et al., 2004). Large et al. (1998) suggest this mineralisation is the result of replacement during diagenesis and recrystallisation associated with dolomite nodule growth in the presence of Zn-Pb saturated sediment pore fluids.
  The δ
34S values of Sp2 have a mean of 9.8‰ (ranging from 3.2 to 18.6), 6‰ greater than the deposit wide mean value for Sp1. These data represent a discrete population and are not a subset of the Sp1 δ34S population. Analyses of pairs of closely adjacent Sp1 and Sp2 samples, including in adjacent laminae, indicate that, despite the overlap of populations, δ34S values for Sp2 are consistently higher than those for Sp1.
• Mode 3 coarse-grained sphalerite-galena-chalcopyrite, occurring in discrete patches is considered to mainly have a clastic origin, being introduced during turbidite sedimentation (Large et al., 1998). These sulphides occur as rounded and subrounded base metal and pyrite clasts that are commonly found within the thicker graded turbidite beds with thicknesses of >1 cm. The rounded base metal sulphide clasts are interpreted to have been reworked from earlier formed sulphide layers on the sub-basin margin/floor and possibly by the erosion of chalcopyrite-rich vein-style mineralisation (similar to that at the Cooley deposit) from the faulted margins of the basin (Large et al., 1998). Others may have formed by the near-complete replacement of isolated carbonate nodules (Eldridge et al., 1993; Large et al., 1998). This mode of mineralisation is volumetrically insignificant (Large et al., 1998).
Other sulphides include galena, arsenopyrite and chalcopyrite which occur throughout the deposit. Galena is commonly found as very fine grained intergrowths in Sp
1 laminations, whilst chalcopyrite is occurs as fine intergrowths within some sphalerite-galena laminations and also as coarser aggregates replacing carbonate clasts in siltstone bands. Both are found in crosscutting deformation-related fractures. Arsenopyrite is very rare and occurs as euhedral prisms in micritic nodular, and brecciated carbonate (Ireland et al., 2004).
Alteration - The McArthur River orebody is surrounded by well developed alteration envelope composed of the following elements (Lambert and Scott, 1973; and Large et al., 2000):
• A Zn-Pb-Tl halo surrounds the deposit, characterised by values of >1000 ppm Zn, >100 ppm Pb and >4 ppm Tl, extending for up to 250 m into the hanging wall lithologies, up to 50 m into the footwall, and up to 15 km SW along the favourable pyritic siltstone host-rock facies of the condensed Barney Creek Depositional Cycle.
• A manganoan carbonate halo that is confined to the W-Fold Shale Member lithofacies that occurs immediately below the host HYC Pyritic Shale Member lithofacies, forming the most laterally extensive part of the composite HYC alteration system. The average MnO content of dolomite within the W-Fold shale increases systematically towards the deposit over a distance of at least 23 km from SW to NE (Large et al., 1998). According to Winefield (1999), rapid basin subsidence at th eclose of Teena Dolostone deposition, accompanied by deepening of the sub-basin and increasing water depth resulted in upwelling of HCO
3-, Fe2+ and Mn2+-rich anoxic bottom water. This resulted in Mn concentration deposited at the oxic/anoxic boundary transition below the developing margins of the sub-basin, forming an extensive Mn-halo in the footwall and along strike from the HYC mineralisation (Large et al., 1998).
• A ferroan dolomite-ankerite halo that is represented by a narrow zone of manganese-rich ferroan dolomite surrounding the deposit and persisting for ~23 km to the west, generally coextensively with the Zn-Pb-Tl envelope. The molar Fe/Mg ratio content of the dolomite which is generally >0.1, increases systematically toward the orebody footwall contact, both along and across strike (Large et al., 2000). Ankerite with pyrite accompanied deposition in the deepest water of the sub-basin at several hundred metres water depth (Large et al., 1998).
• An extensive C-O isotope halo within the dolomitic siltstones of the Barney Creek Depositional Cycle which extends for at least 15 km SW of the McArthur River deposit and approximately coincides with a the lithogeochemical halo of elevated Fe-Mn-Zn-Pb and Tl. Dolomite within this halo has an
18O enriched and 13C depleted isotope signature (δ18O = 23 to 26‰ SMOW, δ13C = -2 to -3.5‰ PDB), relative to normal Proterozoic sedimentary dolomite beyond the halo (δ18O = 20 to 23‰ and δ13C = 0 to -2‰). Dolomitic siltstone lamellae within the Zn-Pb-Ag ores have an isotopic range similar to that of the halo dolomites, suggesting the ore and halo equilibrated from the same hydrothermal fluid. Modelling of isotopic exchange associated with fluid-rock interaction suggests that the halo dolomites equilibrated with low-temperature fluids (50 to 120°C), which were enriched in 18O (δ18O = 5±5‰) but with an average crustal carbon isotope signature where δ13C = -6±1‰ (Large et al., 2001).
  These four halos described above occur on the south, southwest, northwest and northern fringes of the deposit and are considered to be related to the release of cool Fe- and Mn-bearing brines into developing sedimentary basins, both prior to and subsequent to the Zn-Pb mineralising episode (Large and McGoldrick 1998). Large et al. (2001) estimate that the brine pool temperatures were highest at 40 to 70°C in and adjacent to the McArthur River deposit, decreasing to values of 17 to 30°C remote from the deposit.
  The chemistry of the primary hydrothermal fluids responsible for this alteration and accompanying mineralisation, prior to their release into the brine pool, has been constrained by isotopic and thermodynamic modelling to be mildly acid to near-neutral with temperatures of 100 to 240°C, salinity of ~25 wt.% NaCl
equiv. and Zn-Pb concentrations below sphalerite and galena saturation (Rye and Williams, 1981; Cooke et al., 2000; Large et al., 2001).
• Hydrocarbons that are diagnostic of systems in which organic matter is rapidly heated during the passage of hot water through sediments, in particular higher temperature transported polycyclic aromatic hydrocarbons (PAHs), were differentially deposited in higher concentrations within ore bearing beds, and are positively correlated with base metal abundance. In contrast, less permeable beds, e.g., mudstones within centimetres to tens of centimetres above, below or within the ore lenses, carry a higher proportion of overprinted hydrocarbons interpreted to have formed during low-temperature diagenesis or normal in situ burial maturation (Chen et al., 2003).
  These high temperature aromatic hydrocarbons have not been encountered at other locations in the McArthur Basin, away from the orebodies.
  Experimental data indicate the hydrothermally generated PAHs form at temperatures of between 250 and 400°C (Simoneit, pers. comm., 2001 reported in Chen et al., 2003), higher than the estimates of Large et al. (2001) and other authors detailed above. These hydrocarbons are distributed over a length of >2 km, and to maintain these temperatures would require insulation from cool sea water by at least 10 to 20 m of sediment below the sea floor. Chen et al. (2003) envisage brine flow to have been largely parallel to bedding, concentrated in the silty layers because of rapid loss of permeability in the enclosing muddy units during shallow burial. This suggest that channeling of hot fluids along permeable horizons led to the interaction with organic matter and sulphates within the silty turbidites. Ore-forming reactions would produce acid that, in turn, would have generated additional porosity through carbonate dissolution. Examination of ultrathin sections from PAH rich siltstone units indicates loss of dolomite and formation of micro-stylolites in the ore zones, and re-deposition of Fe- and Mn-enriched nodular carbonates downstream.
  As the brines carrying PAHs cooled and the solubility of the PAHs dropped, they were precipitated from solution in order of their molecular weight. The resultant pattern of distribution of PAHs is consistent with the existence of a large thermal gradient that suggests the PAHs were generated upstream to the NNE and moved laterally to the SSW with the mineralising brine, cooling as they went, passing from high to lower grade sulphide accumulations to distal zones of nodular carbonates (Chen et al., 2003). The same authors envisage a proportion of these brines escaped across the stratigraphy to the surface and formed pyrite by thermochemical sulphate reduction in the hotter zones, but most likely by bacterial sulphate reduction in cooler zones closer to the sea floor, accompanied by 'nodular carbonate' and carbonate 'crusts', and silica gel at the sediment-water interface.
Conclusions - Large et al. (2001) and Ireland et al. (2004) and previous authors they draw from, mount a convincing argument that much of the pyrite-sphalerite-galena within the McArthur River deposit, particularly Py
1 and Sp1, was emplaced at the sediment water interface from a stratified brine pool developed in the deepest part of a fault controlled sub-basin adjacent to the Emu fault corridor. Hydrothermal fluid are postulated to enter the lower brine pool as a series of pulses related to seismic activity along the feeder fault systems (Large et al., 1998). This seismic activity was responsible for initiating mass flow turbidites and breccias and for pumping hydrothermal fluids into the sub-basin. Each new pulse of brine has a similar salinity but higher temperature (120 to 240°C, Rye and Williams, 1981) than the cooled waters in the lower brine pool, and consequently rose rapidly to the top of water column, before beginning to cool and follow a convective path downward. As each new pulse of hydrothermal fluid cooled, it would sink and mix with the lower brine pool fluids and precipitate sphalerite, galena and pyrite to be mixed with subwave base sedimentation of mass flow clastic carbonates and pelagic organic-rich fines to form the laminated sedimentary ores (Large et al., 2001).
  Large et al. (1998) suggest that, on average, more than 10 000 pulses of dense fluid were released to generate the intricate layering of base metal sulphides and sediment layers common to each ore lens. Ireland et al. (2004) suggest that each new pulse resulted in a layer of Sp
1 and Py1 and that entrapped pore fluids, due to permeation into the underlying sediment, resulted in the diagenetic overprinting by replacement Sp2 and Py2. They argue that rather than a paragenetic sequence that relates to the entire deposit, each new pulse produced its own paragenetic sequence on a laminae/bed scale of Sp1/galena and Py1 → nodular carbonate, followed by subsurface diagenetic Sp2 and Py2 replacement, to be repeated by multiple pulses. This would account for the textural complexity and isotopic disequilibrium between sulphide phases. Possibly the Sp1 and Py1 crystallised from different levels in the stratified brine pool column where different δ34S levels were stable, to settle together on the sea floor, but have different S compositions.
  Chen et al. (2003) also put forward a convincing argument for shallow replacement by higher temperature hydrothermal fluids as described above. Possibly both processes were represented, with those fluids that were released into the sea and brine pool accounting for the bulk of the Sp
1 and Py1, while brine from the same source was also transmitted along permeable strata in the 'soggy' sediment pile below the sub-basin floor to overprint earlier stratiform sulphide accumulations to produce Sp2 and Py2 by replacement and/or remobilisation.

Reserves, resources and production

The total geological resources prior to mining - 227 Mt @ 9.2% Zn, 4.1% Pb, 41 g/t Ag, 0.2% Cu (Logan et al., 1990).

Reserves and resources as of mid 2004 (Xstrata Zinc) totalled:
  Proven reserve - 5.2 Mt @ 31.0% Zn, 5.3% Pb, 53 g/t Ag,
  Probable reserve - 26.0 Mt @ 11.0% Zn, 5.1% Pb, 53 g/t Ag
  Measured resource - 80.0 Mt @ 13.0% Zn, 5.8% Pb, 57 g/t Ag,
  Indicated resource - 41.0 Mt @ 12.0% Zn, 5.5% Pb, 57 g/t Ag,
  Inferred resource - 0.7 Mt @ 17% Zn, 5% Pb, 60 g/t Ag.

Production from McArthur River in the 12 months to June 2004 totalled - 1.59 Mt @ 13.1% Zn, 5.6% Pb, 55 g/t Ag, representing the final stages of the underground mine before commencement of the open pit operation in 2005.

According to Ahmad et al. (2013), ore reserves and mineral resources at McArthur River at 30 June 2007 were:
  Total reserves - 46.3 Mt @ 9.6% Zn, 4.2% Pb, 43 g/t Ag;
  Total resources - 144 Mt @ 11.2% Zn, 4.8% Pb, 48 g/t Ag.

Remaining JORC compliant ore reserves and mineral resources as at 31 December, 2016 (Glencore, 2017) were:
  Ore reserves
    Proved reserve - 71.2 Mt @ 10.6% Zn, 5.00% Pb, 50.1 g/t Ag;
    Probable reserve - 45.0 Mt @ 7.4% Zn, 3.6% Pb, 37 g/t Ag;
    Total reserve - 117.0 Mt @ 9.4% Zn, 4.5% Pb, 45 g/t Ag.
  Mineral resources which are inclusive of reserves
    Measured resource -123 Mt @ 9.94% Zn, 4.64% Pb, 46.9 g/t Ag;
    Indicated resource - 64 Mt @ 8.9% Zn, 4.1% Pb, 43 g/t Ag;
    Measured + indicated resource - 190 Mt @ 9.6% Zn, 4.5% Pb, 46 g/t Ag;
    Inferred resource - none reported.
The current mine capacity is 5 Mt of ore per annum.

The McArthur River operation is owned by McArthur River Mining, a subsidiary of Glencore.

Other related deposits/mineralisation

  A number of other deposits and occurrences, both stratiform within the Barney Creek Depositional Cycle and transgressive within other lithologies are known within a 20 km radius to the south and west of McArthur River, including:

  This deposit is located 8 m west of the McArthur River zinc-lead mine, and occurs within the east-west elongated Teena sub-basin, a half-graben that is bounded to the north by the ENE-WSW trending growth structure, the Jabiru Fault. Mineralisation is hosted within rocks interpreted to equate with the Barney Creek Formation. The Barney Creek Formation rests on the Teena Dolostone, and is thickest adjacent to, and to the south of, the Jabiru Fault, thinning to the south. The sequence comprises the basal W-Fold Shale Member, overlain progressively by the HYC Shale Member which hosts the ore deposit, Pyritic Barney Creek Formation and the Upper Barney Creek Formation, which transgresses the sequence to the south to rest directly on the Teena Dolostone by ~1 km south of the Jabiru Fault. The Reward Dolostone and Caranbirini Formation overlie the Barney Creek Formation. To the north of the Jabiru Fault, Rox Resources sections show the Upper Barney Creek Formation resting directly on the Teena Dolostone. Those same sections indicate south down displacement on the fault of ~250 to 300 m.
  The deposit is concealed, being shallowest to the west at depths of ~500 m, deepening to be at ~1000 m depth, ~1300 m to the east. It extends for ~400 to 300 m outward from the Jabiru Fault, narrowing to the east.
  The zinc-lead mineralisation at the Teena deposit is interpreted to have been precipitated as stratiform sulphide minerals within fine grained carbonaceous sediments that were accumulating at the base of an anoxic brine pool during sub-basin development, similar to those described above at the McArthur River mine. At a mesoscopic scale, the mineralisation occurs as bedded massive sulphides intercalated with carbonaceous shales and calcareous siltstones. Several phases of mineralisation have been observed ranging from near-syngenetic depositional, to late stage hot influx events during diagenesis and remobilisation and replacement during basin inversion. In the mineralised lodes the principal sulphide minerals (in abundance order) are sphalerite, pyrite, pyrrhotite and galena along with trace arsenopyrite. The main gangue minerals are silicates (orthoclase, quartz and muscovite), ankerite and traces of barite.
  Zinc-lead mineralisation occurs as two sub-parallel, conformable lenses, known as the Lower and the Upper Lodes. The Upper Lode is thicker (up to ~30 m) and higher grade than the Lower (up to 10 m thick). At a 6% Zn+Pb cut-off grade, the JORC compliant inferred mineral resources for each lode are estimated to be, Upper Lode - 45 Mt @ 12.0% Zn, 1.8% Pb, Lower Lode - 14 Mt @ 8.2% Zn, 1.2% Pb, for a total inferred resource of 58 Mt @ 11.1% Zn, 1.6% Pb (This summary is based on Rox Resources Ltd, ASX Announcement,1 June 2016).

  This prospect is located 19 km SW of McArthur River, and contains significant stratiform zinc-lead mineralization within the Barney Creek Depositional Cycle in the Myrtle Sub-basin, at depths of up to 500 m. In initial drilling, the dominant sulphide mineralisation encountered was pyrite, typically occurring as laminae within carbonaceous shale, with sphalerite and galena evident as monominerallic veinlets or as fine to medium grains within carbonate veins (Shannon et al., 1980). Subsequent testing intersected higher grade mineralisation occurring as sphalerite, galena and pyrite in a well bedded calcareous shale, interpreted to be the HYC Pyritic Shale, or an equivalent unit, in a similar style to that of the McArthur River deposit, but with a coarser grain size. A JORC-compliant resource estimate was 43.6 Mt @ 4.09% Zn, 0.95% Pb, at a cut-off grade of 3% Zn+Pb; or 15.3 Mt @ 5.45% Zn and 1.4% Pb at a cut-off grade of 5% Zn+Pb, (Rox Resources Ltd, ASX Announcement,15 March 2010). Current resources include: indicated: 5.8 Mt @ 3.6% Zn, 0.9% Pb; inferred: 37.8 Mt @ 4.2% Zn, 1.0% Pb (Rox Resources Ltd, ASX Announcement,1 June 2016).

  This occurrence is located about 5 km west of McArthur River (Murray 1952; Fricker 1962; Rawlins, 1967, Walker et al., 1977) and contains stratiform Pb-Zn concentrations in the HYC Pyritic Shale Member. Two deep diamond drillholes, 1500 m apart, encountered ~200 m of HYC Pyritic Shale Member and 70 m of the W-Fold Shale Member. The first of these intersected a 30 m @ 3% Zn, including 2 m @ 9% Zn; the second cut 40 m @ 2.2% Zn, which included a basal 3 m @ 9.5% Zn (Murray 1975). A subsequent hole intersected 7.46 m @ 7.76% Zn, 1.7% Pb and 5 g/t Ag. The mineralisation is apparently similar to that at McArthur River and comprises of very fine-grained pyrite, sphalerite and galena, occurring as delicately laminated to massive bands, concordant with the shale bedding. Much of the sphalerite occurs in up to 0.5 mm thick pale straw-yellow laminae, interbedded with dolomitic shale. Coarse-grained red sphalerite with scattered galena crystals is also reported in a few thin beds, which may also host concretionary base metal sulphides (Murray 1975). No inter-ore breccias of the type found at McArthur River have been encountered (Ahmad et al., 2013).

Emu Plains
  This prospect is located ~3 km north of the McArthur River deposit. Two diamond drill holes encountered mineralisation averaging 2% Zn and 0.7% Pb over widths of 50 m (c). Another encountered HYC Pyritic Shale Member lithofacies from 54 to 140 m below the collar, recording assays of 1% Zn over an interval of 7 m. Mineralisation is very similar in style to that at the McArthur River deposit and is at the same stratigraphic level (Ahmad et al., 2013).

Barney Creek Sub-basin
  This east-west elongated sub-basin is located ~17 km southwest of the McArthur River deposit. Diamond drilling intersected ~10 m @ 0.42% Zn, 0.12% Pb as stratiform base metal mineralisation within pyritic shale of the Barney Creek Formation (Ahmad et al., 2013).

Cooley I, II and III
  These deposits are located ~1 km to to the east of the McArthur River deposit within the Emu fault corridor. Outcropping mineralisation has only been found at Cooley I, while the other two have only been intersected in drillholes. All are situated within the Emu Fault corridor, hosted by brecciated Emmerugga Dolostone, which has been attributed to the Mara Dolostone Member (Walker et al., 1977).
  Cooley II is closest to the main Emu Fault, and is principally a copper deposit, with subordinate lead and zinc, whilst Cooley I and III are mainly lead and zinc deposits.
  The outcropping Emmerugga Dolostone at Cooley I strikes north-south and dips moderately to the west, where it comprises well bedded stromatolitic dolostone. This succession is cut by NW striking, steeply NE dipping dolomite 'dykes' that are up to 1 m thick. These 'dykes' carry galena-bearing sparry dolomite veins running parallel to the 'dyke' walls. Some crosscutting galena-dolomite stringers and pods are also observed. Locally, the Emmerugga Dolostone is brecciated and carries galena-sphalerite pods as breccia fill. Two types of breccia have been defined in drill core (Williams, 1978), namely, Br
1 which consists of angular dolostone clasts, up to 1 m across set in a dark matrix of dolomite grains, pyrite and carbonaceous material; and Br2, composed of angular clasts of Br1 in a matrix of dolomite and sulphides which occur as veins or disseminations, with interlocking laths of barite, or dolomite pseudomorphs after barite, and carbonaceous material. The Br1 breccias are interpreted to be the result of movement of faults within the Emu Fault corridor, as well as debris slumping and/or solution collapse. Exposed breccias at Cooley I are interpreted to belong to this phase of brecciation.
2 breccias are interpreted to be due to solution collapse following the introduction of mineralising fluids into porous Br1 breccias. The dolomite 'dykes' may represent channelways for the mineralising solutions, or may be neptunian dykes of redistributed carbonate.
  The sulphides within these deposits are relatively coarse grained and were deposited sequentially in the order of pyrite-marcasite-barite-dolomite → Cu sulphides → galena-sphalerite, with the Cu-bearing sulphides closest to the main Emu Fault, mostly at Cooley II. Sulphur isotopic ratios of pyrite, sphalerite and galena have a range and occurrence similar to those at McArthur River, with pyrite having large variations, whilst galena and sphalerite are restricted to a narrower range. Galena-sphalerite sulphur isotope geothermometry suggest a temperature of 290°C for Cooley I, and 275°C for Cooley II (Williams 1978).
  The δ
13C and δ18O values of ore-stage dolomite define a linear trend, whilst individual deposits yield distinctive values that become heavier away from the Emu Fault. Modelling of δ13C and δ18O data indicates temperatures of 310°C for the main Emu Fault zone, 300 to 290°C for Cooley I, 275 to 250°C for Cooley II and 240 to 185°C for Cooley III.
  Williams (1978) and Rye and Williams (1981) considered the mineralisation to be epigenetic and suggested the hydrothwermal fluids emanated from the Emu Fault and flowed to the west, consistent with the observed decrease in Cu/Pb+Zn ratios and temperatures away from the Emu Fault.
  The above is paraphrased from Ahmad et al. (2013), whi, in turn, summarised detailed descriptions by Williams (1978) and Rye and Williams (1981).

Ridge I and II
  These occurrences are located east of the McArthur River and west of the Cooley deposits, largely within the Emu fault corridor, although the Ridge II stratiform mineralisation extends to the west of the main fault zone, immediately to the south of the McArthur River deposit. At Ridge I, mineralisation is hosted within Cooley Dolostone Member sedimentary breccia, whereas at Ridge II, the upper half is within HYC Pyritic Shale Member lithofacies and the lower half is in brecciated Cooley Dolostone Member lithofacies.
  At Ridge I and the lower half of Ridge II, coarse grained sulphides occur as open-space fillings within the breccia matrix the Cooley Dolostone lithofacies, whereas in the upper half of Ridge II, sulphides are fine grained and similar in texture to those of the McArthur River stratiform ore lenses. The principal ore minerals are pyrite, galena, sphalerite and minor chalcopyrite, and the paragenetic sequence is similar to that in the Cooley deposits, and as in those deposits, two types of breccias, Br
1 and Br2 have been recognised.
  In the upper half of the Ridge II occurrence, stratiform mineralisation occurs directly above the breccia-fill sulphide accumulations and within the upper part of the HYC Pyritic Shale lithofacies that is developed above the main Cooley Dolostone. This mineralisation is >300 m above the main HYC ore zone, and extends to the west into the southern section of the main HYC sub-basin.
  Stable isotope systematics at the Ridge occurrences are similar to those in the Cooley deposits, although calculated temperatures are lower. Galena-sphalerite sulphur isotope geothermometer indicated a temperature range 180 to 120°C for the Ridge II deposit, whilst modelling of carbon and oxygen isotope data from ore-stage dolomite gave temperatures of 190 to 175°C for Ridge I, and 225 to 190°C for Ridge II (Rye and Williams 1981). The preceding is based on Ahmad et al. (2013), which in turn summarises Murray (1975), Williams (1978), Logan (1979) and Rye and Williams (1981).

  The Coxco deposit is located ~10 km SE of the McArthur River mine, within the Emu fault corridor. It derives its name from the adjacent mines, Cox and Cooks, which were first discovered in the late nineteenth century, when they were worked for high-grade oxide ore from a number of shallow shafts and pits. Production records are not available.
  Extensively drilling by the Carpentaria Exploration Company Pty Ltd (CEC), and later a CEC and North Ltd joint venture, resulted in an unpublished 'reserve' estimates of 'several million' tonnes @ 2.5% Zn, 0.5% Pb (Walker et al., 1983).
  Mineralisation is hosted by the Reward Dolostone, which is unconformably overlain by the Lynott Formation and unconformably underlain by the Mara Dolostone Member of the Emmerugga Dolostone. The Reward Dolostone-Lynott Formation contact is silicified, comprising quartz, void-filling chalcedony and goethite, possibly representing a palaeo-regolith surface. Adjacent to the Mara Dolostone-Reward Dolostone unconformity, the Mara Dolostone Member contains fine grained, laminated, discordant quartz-dolomite veins (Ahmad et al., (2013).
  Lenses and dykes of breccia are found within the Reward Dolostone and upper part of the Mara Dolostone Member. These breccia lenses are composed of angular pieces of host rock set in a fine-grained matrix of dolomite, quartz, clay, organic matter, feldspar, chert, iron sulphides, collophane, mica and dolomite pseudomorphs after barite. They have a sharp contacts with the host rock and are interpreted to represent solution cavity fill developed during an episode of karst development. A second breccia type, interpreted to be crackle breccia, is preferentially developed within the Reward Dolostone, immediately below the silicified zone. It is clast supported, with rounded to angular fragments, without any exotic lithologies, and originated by brittle deformation during hydraulic fracturing (Walker et al., 1983).
  Two stages of Pb-Zn mineralisation have been recognised: Stage I, which comprises sphalerite, marcasite, pyrite and galena, occurring as colloform and crystalline crusts and fragments within the karst breccia matrix; and Stage II composed of sphalerite, marcasite, pyrite and galena, present as veins and as matrix in the crackle breccia.
  Comb structures are common in stage II mineralisation, with pyrite and marcasite near the base → sphalerite → galena. Stage II mineralisation crosscuts the Reward Dolostone, the silicified zone, void fill, stage I mineralisation and the basal Lynott Formation, and hence formed following lithification of the Lynott Formation.
  Two-phase (water+vapour) primary liquid inclusions from stage II sphalerite and dolomite have freezing temperatures in the range -22.2 to -27.8°C, suggesting high-salinity brines, with homogenisation temperatures in the range 169 to 100°C, confirmed by galena-sphalerite sulphur isotope geothermometry estimates of 191 to 128°C. Lead isotope ratios are similar to those of the McArthur River deposit and are interpreted to reflect ore-fluid derivation from basinal brines, although some samples have a more radiogenic component, suggesting possible leaching of the McArthur Group sedimentary rocks. Carbon and oxygen isotope data from bituminous matter and dolomite show that the former to be mainly biogenic, with δ
13C values of -37.7 to -32.5‰. The host dolomite has a wide spread of δ18O values ranging from 19.8 to 24.7‰, but a narrow range of δ13C values of -3.0 to 0.1‰. Two samples of stage II dolomite gave δ18O values of 18.2 and 20.4‰ and δ13C values of 0.8 and -1.3‰. These values are not comparable to those of ore stage mineralisation in the Cooley and Ridge deposits. Walker et al. (1983) suggested stage I mineralisation took place in karst cavities resulting from the mixing of metal-bearing brines with reduced sulphur bearing groundwater. In contrast, the same author implied stage II fluids were basinal brines at temperatures of 170 to 100°C, with precipitation being due to a biological sulphate reduction by hydrocarbons at the site of metal deposition. The preceding is based on Ahmad et al. (2013).

  The small Reward deposit is located about 15 km west of the McArthur River deposit and was exploited as a historic underground mine, producing a few hundred tonnes of ore, probably during the late 1950s. Mineralisation is hosted by the Reward Dolostone and underlies a silicified zone of massive to brecciated, pervasive and generally structureless chert/jasper, with some local remnant bedding. The silicified zone resembles that at Coxco and other similar small occurrences in the district, and may represent a palaeo-regolith. The silicified zone is underlain, across an irregular, but sharp contact, by an ~2 m thick yellow ochreous zone. The yellow ochreous zone is, in turn, underlain by red-brown, ferruginous gossanous material, containing cerussite, pyromorphite and minor malachite. At one location in the deposit area, a small outcrop of unsilicified Reward Dolostone is heavily fractured, containing fracture fill galena, dolomite and sphalerite. Galena and dolomite stringers are also evident in unweathered Reward Dolostone elsewhere in close proximity to the mine (Ahmad and Ferenczi 1993).

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George Fisher Zn-Pb-Ag .......... Thursday 10 November, 2005 .......... Travel by hire van from Mt Isa.

The George Fisher South (previously Hilton or P49) and George Fisher North (L72) are located approximately 20 km north of the city of Mt Isa, in north-western Queensland, Australia and 975 km west of Townsville (#Location: George Fisher North - 20° 32' 56"S, 139° 28' 6"E; George Fisher South - 20° 34' 4"S, 139° 28' 35"E).

  Whilst lead-zinc mineralisation had been discovered at what is now Mount Isa in March 1923 by prospector John Campbell Miles, and mining commenced in 1931, the Hilton deposit was not discovered until 1947, when S R Carter, a Mount Isa Mines (MIM) geologist, recognised outcropping cerussite at surface. Diamond drilling to test this discovery was commenced in August 1948, with the first drill hole intersecting a narrow interval of zinc mineralisation. From then until 1957, a significant follow-up drilling program was undertaken, and by 1950, the Hilton ore reserves stood at 26 Mt. The drilling program was curtailed in 1957 due to a fall in metal prices and heavy capital expenditure at the main Mount Isa operation. In 1966, MIM consolidated its mining lease holdings by taking up all the ground between Hilton and the Mount Isa operations within a single Special Mining Lease and diamond drilling recommenced at Hilton. The Hilton reserve was subsequently increased to 37 Mt. MIM decided to develop Hilton in 1969 but market factors delayed start-up of the 1 Mt/y mine and concentrator until 1990. In 1981, similar mineralisation was located 2 km further north at Hilton North, later to be renamed George Fisher North. Mining commenced at George Fisher in 2000.

Continental to Regional Setting

  The George Fisher and Mount Isa orebodies lie within the Leichhardt River Fault Trough of the Western Fold Belt, part of the Mount Isa Inlier (or Domain) of Northwest Queensland. The Mount Isa Inlier comprises three major elements from west to east: the Western Fold Belt, the Kalkadoon-Leichhardt Belt and the Eastern Fold Belt, which are predominantly north-south trending sedimentological and structural domains (Blake and Stewart, 1992; O'Dea et al., 1997). The Western Fold belt is bounded to the east by 1870 to 1850 Ma felsic volcanic and coeval granitoid rocks of the Kalkadoon-Leichhardt Belt, interpreted to represent a remnant of a magmatic arc related to the Palaeoproterozoic Barramundi Orogeny. To the NW, the Western Fold belt is separated from the McArthur Basin in the Northern Territory by the 1853±4 Ma Murphy Metamorphics basement exposed as the narrow ENE-WSW trending Murphy Inlier.
  The Western Fold belt is further divided into the narrow Leichhardt River Fault Trough immediately to the west of the Kalkadoon-Leichhardt Belt, and the broader Lawn Hill Platform further to the west, each separated from its neighbour by a major north-south trending terrane boundary fault zone (Blenkinsop et al., 2008; Foster and Austin, 2008). In the northern half of the belt, this structure is the major Mount Gordon Fault, while to the south it is in part the Mount Isa Fault.
  The 1300 x 60 to 200 km Mount Isa Inlier-McArthur Basin succession comprise Late Palaeo- to Mesoproterozoic sequences that are, in turn, part of the more extensive 1800 to 1580 Ma Northern Australian Platform cover, a 5 to 15 km thick volcano-sedimentary succession draped across the northern third of the continent. Deposition within the Mt Isa-McArthur basin system took place in three super-basins which represent three nested cycles of deposition and exhumation, specifically the Leichhardt (1798 to 1738 Ma), Calvert (1728 to 1680 Ma) and Isa (1667 to 1575 Ma) super-basins, terminated by the 1590 to 1500 Ma Isan Orogeny, which was followed by the younger Roper super-basin.
  All of the major stratabound Zn-Pb-Ag deposits of northern Australia, including the George Fisher and Mount Isa deposits, are hosted by the Isa Super-basin in the McArthur Basin (e.g., McArthur River), Lawn Hill Platform (e.g., Century and Lady Loretta), Leichhardt River Fault Trough (e.g., George Fisher, Mount Isa Zinc) and Eastern Fold Belt (e.g., Dugald River).
Western Fold Belt Mt Isa Inlier   A ~25 to 30 m.y. hiatus marked the end of deposition and inversion of the Calvert superbasin within the Mount Isa inlier. During this period the Calvert Super-basin sedimentary rocks were uplifted and incised. Towards the end of this hiatus, the linear, north-south trending, voluminous, lopolithic 1671±8 Ma Sybella Granite, a foliated coarse porphyritic biotite granite, was intruded along the boundary between the Leichhardt River Trough and the Lawn Hill platform immediately to the west of Mount Isa. The exposed and sub-outcropping batholith occupies an area of ~220 x 15 to 30 km and defines the Sybella Domain of Withnall and Cranfield (2013). The northern extremity of this batholith passes into the NNE trending Mount Gordon Fault Zone which separates the Leichhardt River Trough from the Lawn Hill platform in the northern part of the Mt Isa Domain.
  The depositional hiatus during Calvert super-basin inversion ended with renewed extension, development of the Isa Superbasin and recommencement of sedimentation from 1670 to 1590 Ma (Jackson et al., 2000). This superbasin is interpreted to have resulted from thermal subsidence of the lithosphere, the sag phase of Etheridge and Wall, 1994, and of Betts and Lister, 2001. This subsidence was most intense to the west and NW of the Leichhardt River Domain, where an extensive, thick blanket of carbonaceous shale, stromatolitic dolostone, and turbiditic sandstone and siltstone were deposited on the Lawn Hill platform the ~10 km thick McNamara Group (Krassay et al., 2000; Southgate et al., 2000). A similar sequence was deposited in the McArthur basin to the north, the 5.5 km thick McArthur Group, which hosts the McArthur River deposit in the Northern Territory (Rawlings, 1999). Within the narrower (40 to 90 km wide) Leichhardt River Trough, the Isa Superbasin sequence is thinner and and spans a narrower temporal range, represented by the up to 7.5 km thick, ~1670 to ~1647 Ma Mt Isa Group, which hosts the Mount Isa, George Fisher South (previously Hilton) and George Fisher North deposits close to its western margin, immediately adjacent to the regional-scale Mount Isa Fault. The Lady Loretta deposits are located within a similar sequence in the lower sections of the overlying McNamara Group on the eastern margin of the Lawn Hill platform. Century is located further to the north, also within the Lawn Hill Platform, adjacent to the major Termite Range Fault, which may merge with the Mount Isa Fault to the SE.
  The regional structure of the Western Fold Belt of the Mount Isa Domain is dominated by faults and shear zones striking north-south (e.g. Mt Isa Fault), NW to NNW (e.g., Riversleigh and Termite Range faults) and NNE (e.g., Mount Gordon Fault), with subsidiary east-west, NW–SE or NE–SW structures.
  Subsequent to the 1870 to 1850 Barramundi Orogeny, the rocks of the Mount Isa Domain were subjected to at least three phases of extensional faulting and rifting, and at least as many episodes of intervening inter- and post-rift compressional deformation and basin inversion spanning the three superbasin cycles. These events culminated in the most intense episode, the 1590 to 1500 Ma Isan Orogeny that terminated deposition of the Mount Isa Group (e.g. Blake, 1987; O'Dea et al., 1997; Betts et al., 2006; Gibson et al., 2012). Thrusting during inversion of the superbasins is interpreted, in many cases, to have reactivated listric faults active during the preceding extensional events.
  The intensity of deformation and metamorphism varies markedly across the Mount Isa Inlier/Domain, although, in contrast to the more strongly deformed Eastern Fold Belt, it rarely exceeds greenschist facies in the rocks of the Lawn Hill Platform and Leichhardt River Fault Trough (Gibson et al., 2016). Metamorphism was largely low-pressure sub-greenschist facies during the Isan Orogeny (Jacques et al., 1982; Rubenach 1992; Oliver 1995; Rubenach and Barker 1998), although higher metamorphic grades to upper amphibolite facies are exposed in the hanging wall of the Mt Isa Fault, adjacent to the Sybella Batholith (Betts et al., 2006).
  During deposition in the Leichhardt Superbasin, the Eastern Creek Volcanic pile and the thicker syn-rift sedimentary rocks, accumulated in an elongate, fault-bounded basin which, following deformation, is preserved as a 50 to 80 km wide belt (Bain et al., 1992; Blake, 1987; Derrick, 1982; Eriksson et al., 1983; Gibson et al., 2012; Jackson et al., 2000; Scott et al., 2000), largely restricted to the area of the Leichhardt River Fault Trough. This basin was bounded by faults typically striking NNW and dipping steeply to both the east and west, with displacements varying from a few hundred metres to several kilometres. Variations in this pattern indicate this is not a single basin but a series of highly asymmetric, overlapping, half-grabens, each oriented north-south and up to 70 to 130 km long, controlled by pre-existing fabric in the underlying basement (Gibson et al., 2016). Rifting ceased by no later than ~1740 Ma, with intrusion of post-kinematic granites and gabbro, whilst the regional north-south trending Leichhardt Anticline developed between 1740 and 1710 Ma within the Leichhardt River Fault Trough during east-west shortening inversion. Compression was followed by an episode of thermally induced subsidence (Gibson et al., 2016).
  The western margin of the northern half of the Leichhardt River Fault Trough, separating it from the Lawn Hill Platform, is defined by the major, up to 7 km wide Mount Gordon Fault Zone, along which, the most recent activity has been dextral displacement of up to 10 km cutting all other fault sets. Towards its southern extremity, the Mount Gordon Fault curves north-south and then SSE to merge with and become part of the north-south Mt Isa fault corridor. Further south, the latter swings to the SSW again to become the Rufus and Mount Annabele fault zones.
  The intensity of deformation within the Leichhardt River Fault Trough increases and is more complex in the vicinity of the Mt Isa Fault Zone where at least five deformation events have been interpreted based on overprinting relationships (Betts et al., 2006), including faulting and folding and overprinting of different strands within the zone (Connors and Lister 1995), three of which are indicated by post crystallisation dating of metamorphic minerals from the Sybella Granite at 1610±13, 1544±12 and 1510±13 Ma (Bell 1986). These deformational events are detailed in the 'Structure' and 'District Scale Structure and Metamorphism' sections below.
  The Lawn Hill Platform differs from the Leichhardt River Fault Trough in that much of the Calvert and Isa superbasin successions are preserved in contrast to the latter, which is more deeply eroded, with only remnants of these sequences remaining. In the northern Lawn Hill Platform, immediately west of the Century deposit, a series of major NW-SE to NNW-SSE trending, east dipping faults occur over a width of ±20 km, including, from SW to NE, the Riversleigh, Little Range and Termite Range fault zones. These structures appear to have controlled deposition as listric faults bounding asymmetric, SW thickening half grabens. They also accommodated subsequent basin inversion as SW vergent thrusts, mainly during the Isan Orogeny. Depositional thickening and thrust displacement is most intense across the Riversleigh Fault zone, influencing the Leichhardt, Calvert and Isa Superbasin sequences. These structures approach, but do not appear to intersect the Mount Gordon-Mount Isa fault zones. Other NNW-SSE structures on the Lawn Hill Platform include the possibly sinistral May Downs Fault further to the south, which curves south to follow the western margin of the Sybella Granite just to the west of Mount Isa. To the east of these structures, over the platform, there appears to be little variation in thickness of the Calvert Superbasin sequence across other faults, suggesting little structural activity within the interior of the superbasin during deposition (Gibson et al., 2016), although ENE-WSW, mainly Calvert-age structures have been variously attributed to NW-SE crustal extension on the Lawn Hill Platform (Betts et al., 1998, Southgate et al, 2006).
  At least two composite, regional Isan Orogeny basin inversion events are recorded in the Western Fold Belt, including:
 i). ~north-south shortening and inversion of pre-existing east-west rift related structures between 1595 and 1570 Ma to produce east-west oriented axial-planar foliations in the immediate hangingwall of inverted extensional faults in the Leichhardt River Fault Trough and ENE-WSW structures in the Lawn Hill Platform (O'Dea and Lister 1995; O'Dea et al., 1997; Lister et al. 1999);
 ii). a major phase of ~east-west shortening that developed crustal-scale north-south trending, shallowly plunging folds, and north-south trending reverse faults and thrusts across the Leichhardt River Fault Trough, and NE-trending folds throughout the Lawn Hill Platform between 1570 and 1550 Ma (Betts and Lister, 2002; Betts et al., 2004; Blaikie et al., 2017). Between ~1550 and 1540 Ma, the structural regime changed to strike-slip and oblique faulting and regional wrenching (O'Dea et al., 1997; Lister et al., 1999). This late east-west shortening was accommodated along NW trending sinistral, and NE trending dextral strike-slip faults. During inversion, zones of anomalous strain and buttressing occurred near normal faults and granitic plutons (Betts et al., 2006).
  The Mount Isa terrane is considered to have been cratonised after the Isan Orogeny and the architecture of the region has not changed since the Late Mesoproterozoic.
  It has been suggested that the Mount Isa Fault Zone is a Palaeoproterozoic Barramundi Orogen terrane-bounding suture (e.g., Hobbs et al., 2000). However, geophysical and geochemical data studied by Bierlein and Betts (2004) showed that basement rocks on either side of the Mount Isa Fault have similar densities, which is consistent with geochemical observations and Sm-Nd data that suggest basement lithologies on either side of the structure are geochemically and isotopically indistinguishable from each other, and that the Mount Isa Fault is unlikely to represent a suture zone that separates different Palaeoproterozoic terranes. Their data also indicates the crustal blocks on both sides must have been in close proximity of each other since the Palaeoproterozoic, and that the Western Fold Belt was part of the ancestral North Australian Craton well before the Barramundi Orogeny.

Regional to Deposit Scale Geology

The host sequence in the Mount Isa-George Fisher area is the Mount Isa Group, which unconformably overlies the Surprise Creek Formation of the Calvert Superbasin, and older rocks of the Leichhardt Superbasin, as follows:
Leichhardt Superbasin
May Downs Gneiss - which is found ~5 km to the west of Mount Isa, between the overlying Mount Guide Quartzite and an intrusive contact with the 1671±8 Ma Sybella Granite to the west. It comprises plagioclase (microcline)-quartz gneiss and schist containing minor biotite, sillimanite and muscovite. It is reported to be younger than 1789±4 Ma and may be a more intensely metamorphosed equivalent of the Bottletree Formation (Australian Stratigraphic Units Database, viewed April, 2019).
Bottletree Formation - up to 3000 m of porphyritic, rhyolitic to dacitic lava flows and ash-flow deposits, with interlayered greywacke and greywacke conglomerate, meta-arenite, quartzite, epidotic quartzite and grit. Sheared, schistose, amygdaloidal to massive metabasalt is common at or near the base and top of the formation. It unconformably overlies granites and volcanics rocks of the Kalkadoon-Leichhardt Belt, and is conformably followed by the Mount Guide Quartzite, the basal unit of the Haslingden Group;
Mount Guide (and Leander) Quartzite - of the Guide Supersequence which comprises a >4000 m package of alluvial sheet, braided river and marginal lacustrine and marine deposits composed of a lower sequence of mainly greywacke type metasediments, including conglomerate and micaceous quartzite, and an upper suite of mostly ridge-forming quartz-rich sandstone metamorphosed to quartzite (Blake, 1987).
  Deposition of the Bottletree Formation and Guide Supersequence was in an east-west extensional regime. The overlying sequences were deposited in a north-south directed extensional environment (Blaikie et al., 2017).
Eastern Creek Volcanics - a 1790 to 1740 Ma, 6 to 15 km thick tholeiitic continental volcanic suite, principally composed of amygdaloidal to massive metabasalt, with interbedded clastic sedimentary rocks, deposited in an extensional setting. It has been subdivided into the Cromwell Metabasalt, Lena Quartzite and overlying Pickwick Metabasalt. These volcanic rocks were extruded under subaerial or shallow water conditions and are part of a larger bimodal igneous province represented in the Eastern Fold Belt by mafic and felsic volcanic rocks, including 1780 Ma rhyolites and ignimbrites of the Argylla Volcanics and the 1760 Ma Bulonga Volcanics (Neumann et al., 2009; Withnall and Hutton, 2013; Gibson et al., 2016);
Myally Subgroup - formerly known as the Judenan Beds. This sequence constitutes the Myally Supersequence and is an up to 4000 m thick succession, from base to top, principally of quartzite, feldspathic quartzite to siltstone with associated pebbly to conglomeratic sandstone, mudstone, arkose, shale, quartzite, schist, dolomitic sandstone, oolitic and stromatolitic dolostone. Some felsic tuff occurs at the top, and metabasalt near the base of the sequence (Blake, 1987). This unit is the uppermost member of the Haslingden Group;
Quilalar Formation - which is largely only exposed on the eastern margin of the Leichhardt River Fault Trough, distal to the Mount Isa area. It constitutes the Quilalar Supersequence and represents a transition to ~1750 to 1740 Ma post-rift sedimentation characterised by storm-, tide- and wave-dominated marine shelf and continental facies quartzite composed of feldspathic quartzite, orthoquartzite, conglomerate, arkosic grit, shale, siltstone, minor limestone and dolostone.
Calvert Superbasin
Bigie Supersequence
Bigie Formation, a dominantly fluvial succession in a NW-SE directed extensional basin setting that formed as a SE thickening half graben between ~1710 and 1690 Ma. The sequence comprises up to 800 m of purple-brown hematitic sandstone, pebbly sandstone, tuff, conglomerate and red-brown siltstone.
Fiery Creek Volcanics - a suite of bimodal volcanic rocks dated at 1708±2 Ma, comprising up to 750 m of rhyolite, agglomerate and amygdaloidal/vesicular altered basalt, interbedded with sandstone, conglomerate and siltstone.
Prize Supersequence
Surprise Creek Formation - a 2500 m thick, generally upward fining sequence of alternating sandstone and quartzite sheets and wedge shaped siltstone horizons.
Isa Superbasin
Gun Supersequence - which is separated from the top of the Prize Supersequence by an unconformity and emplacement of the 1671±8 Ma Sybella Granite and 1678±2 Ma Carters Bore Rhyolite. It comprises, from the base to top:
Mount Isa Group which constitutes the Gun Supersequence in the Leichhardt River Fault Trough, and the lower sections of the Loretta Supersequence on the Lawn Hill Platform.
Warrina Park Quartzite - a thin, discontinuous basal unit resting unconformably on rocks of the Calvert and Leichhardt superbasins. It comprises an up to 300 m thick, white conglomeratic orthoquartzite to blocky feldspathic sandstone with abundant ripple marks and cross-bedding;
Mondarra Siltstone - which is generally 300 to 1000 m thick with a maximum of 1700 m. It is a thinly laminated, weakly micaceous, dolomitic siltstone with fine sandstone and black dolostone, overlain by dolomitic sandstone and micaceous slate.
Breakaway Shale - up to 300 m of thinly laminated, siliceous siltstone/shale, locally chert, typically bleached at surface, but carbonaceous at depth. It is principally composed of quartz, albite, chlorite and muscovite with accessory carbon, microcline pyrite and tourmaline.
Native Bee Siltstone - a carbonaceous, dolomitic shale and siltstone sequence up to 800 m thick, with rare, thin, vitrophyric tuff bands. Two main fancies have been recognised: i). a featureless carbonaceous shale; and ii). a rock type consisting of alternating white calcite and dark carbonaceous quartz-dolostone layers about 2 to 20 mm thick known locally as the 'zebra-shale'. It contains stromatolites, halite pseudomorphs, flat pebble conglomerates and cross-bedded channel deposits.
Urquhart Shale - which has been dated at ~1655 Ma and has a maximum thickness of 1050 m. Where unmineralised, it comprises a carbonaceous, dolomitic shale and siltstone, closely resembling the Native Bee Siltstone. It contains numerous potassium rich marker beds, some of which have been positively identified as tuffs. Five facies have been recognised within the unit at George Fisher, each intercalated as bands ranging from a few mm to 25 m in thickness (after Chapman, 2004):
    - Medium-bedded stylolitic mudstone - occurring as units with sharp boundaries that range from 5 to 15 m in thickness, composed of individual planar, laminated to massive beds that are 5 to 10 cm thick. This facies occurs as sheet-like bodies with consistent widths across the deposit. The mineralogy comprises quartz, calcite, and/or ferroan dolomite with accessory white mica, K feldspar and carbonaceous matter;
    - Banded mudstone - occurring as units that range from 2 to 4 m in thickness with gradational boundaries. It is composed of discontinuous 10 to 100 m long lenses to sheets that extend across the deposit. Beds are more commonly massive than laminated or graded, and are 2 to 5 cm thick. The mineralogy is the same as for the medium-bedded stylolitic mudstone. White, layer parallel carbonate bands are common.
    - Carbonaceous Siltstone - occurring as units with gradational boundaries that range from 2 to 50 cm in thickness, composed of mm to sub-millimetre scale alternating light grey and dark wavy laminations with intermittent millimetre thick planar laminations. The light grey laminations are predominantly composed of ferroan dolomite and quartz. The dark laminations have a similar composition, but in addition, contain abundant carbonaceous material. These bands are interleaved with banded mudstones.
    - Pyritic Siltstone - occurring as units with gradational boundaries that range from 1 to 5 m in thickness, composed of mm to sub-millimetre scale alternating light grey and dark wavy laminations with intermittent millimetre thick planar laminations. The thickest intervals occur as sheets that are continuous across the deposit. However, at gradational boundaries with banded mudstone, individual laminations may be discontinuous. The light laminations have the same composition as those of the carbonaceous siltstone, while the darker laminations are composed of fine grained spheroidal pyrite and minor carbonaceous material. White, layer parallel carbonate bands and nodular carbonate are common to abundant. Pyritic Siltstone beds tend to be interbedded with banded mudstone to form units from a few to >100 m in thickness and constitute >80% of the central and upper sections of the Urquhart Shale.
    - Tuffaceous Marker Beds - occurring as 0.5 to 10 cm thick beds that are continuous across the deposit. They have sharp contacts and are composed of quartz and K feldspar with accessory chlorite and white mica.
  A different suite of constituent facies of the Urquhart Shale have been mapped at Mount Isa, although the variations are largely nomenclatural.
Spear Siltstone - composed of rhythmically laminated siliceous dolomitic siltstone and siliceous dolostone up to 160 m thick. It contains stromatolites, halite pseudomorphs, flat pebble conglomerates and cross-bedded channel deposits.
Kennedy Siltstone - which is very similar to the underlying Spear siltstone, differing only in its course bedded to massive character over much of the sequence which has a maximum thickness on 320 m. Sedimentary brecciation is common to both units.
Magazine Shale - a carbonaceous siliceous shale unit that is the uppermost formation of the Mount Isa Group. It's upper section, along its entire strike length, is truncated by the Paroo Fault, with a maximum preserved thickness of 220 m. Outcrops have a characteristic reddish colour due to contained iron oxides.


  A succession of deformation events has been recognised within the Leichhardt River Fault Trough, best reflected in the vicinity of the Mount Isa Fault, although not all are evident in the broader Western Fold Belt. These are as follows:
D1, as defined by Bell (1983) is a pre- to early Isan Orogeny subhorizontal, north-south shortening that occurred between ~1610 and 1590 Ma. It produced local thrusting, east-west trending folds, variably developed bedding-parallel carbonaceous seams (S1), and a north-south trending mineral lineation at the Mount Isa mine (Bell, 1991), but is not recognised at the George Fisher deposit (Chapman, 2004). It is generally characterised by east-west trending, near vertical foliation parallel to bedding, with a north-south trending mineral elongation (Bell and Hickey, 1998).
D2 which produced the bulk of the regional folds with steeply dipping north-south trending axial planes. It was caused by a east-west shortening event during the Isan Orogeny at ~1544 Ma (Page and Bell, 1986), which at George Fisher was responsible for upright, tight folds with steeply dipping, subvertical, north-south striking axial planes and an associated S2 foliation. It is also defined by concentrations of carbonaceous material. It accompanied folding of and reverse displacement on the Paroo and Hanging Wall faults and dyke intrusion (Chapman, 2004). S2 is the dominant regional schistosity, commonly with a steeply south pitching L2/2 stretching lineation (Bell and Hickey, 1998). The Mount Isa and George Fisher deposits are located on the western limb of a regional D2 anticline (Davis, 2004).
D3, previously known as D2.5 of Bell and Hickey 1996; Perkins 1996; Bell and Hickey 1998; Chapman 2004 and others. This deformation was first recognised after D2 and D3 had been established in the literature, and hence was designated as D2.5 by Bell and Hickey (1996). However, subsequent authors (e.g., Davis, 2004) referred to it as D3 and then referred to the the subsequent D3 event as D4, etc. (c.f., Wilde et al., 2006; Long, 2010). This deformation, which is evident at both Mount Isa and George Fisher, was the result of NE-SW to ENE-WSW compression (McLellan et al., 2014). It rotated D2 structures (specifically S2) into gently dipping microfolds (or crenulations) to macroscopic folds with subhorizontal axial planes and east vergence, but no recorded foliation. These folds may either die out along their axial planes into mineralised layers, or curve to merge into S3 (Perkins, 1996).
D4 of Davis (2004) and subsequent authors, (previously D3 of earlier authors as listed above), which is another shortening event, but reflects SE-NW shortening and took place at ~1510 Ma. It resulted in further development of folds with steeply dipping, subvertical axial planes and gently plunging axes that trend NW-SE to NNW-SSE and have an axial planar schistosity, S4 (Page and Bell, 1986; O'Dea et al., 1997; Bell and Hickey, 1998). These folds are more localised than those of D2. The largest D4 structure in the Mount Isa deposit area is the Mount Isa Fold, a syncline-anticline pair with a wavelength of 200 to 400 m, amplitude of 130 m, and a near vertical common short limb. Other D4 folds have wavelengths no wider than 10 to 20 m.
Post D4, described by Chapman (2004) as Post D3, which is interpreted to have produced local NW-SE trending folds; brittle NNW-SSE splays from the Mount Isa-Paroo faults, including the Transmitter and Gidyea Creek faults between George Fisher and Mount Isa; and reactivation of the Mount Isa and Paroo fault zone. This deformation is assumed to encompasses the D5 and/or D6 listed below, but may well represent the late stages of D4.
D5, which was previously described by Bell and Hickey (1998) as D4, is only manifested in the Mount Isa District as flat lying kinks, mainly to the west of the Mount Isa Fault in more schistose rock types. These kinks may be distinguished from D3 crenulations in that the latter have not been refolded by D4. The sense of shearing is generally east-vergent (Bell and Hickey, 1998).
D6, is only very weakly and sparsely developed, forming crenulations or poorly developed crenulation cleavage. Where observed to the west of the Mount Isa Fault, it is associated with local retrogression. While not significant in the Western Fold Belt, it is more important in the Eastern Succession, being related to gold mineralisation at Starra (Bell and Hickey, 1998).

District Scale Structure and Metamorphism

  In the Mount Isa-George Fisher area, the Mount Isa Group rocks occurs as a west-dipping, north-south striking sequence on the western limb of a regional D2 anticline. The stratigraphic package is ~4000 m thick, decreasing north of the Transmitter fault (~2 km south of Hilton) where sections of the group have been removed by faulting (Valenta, 1994). The formations of the Upper Mount Isa Group are truncated to the west by the Paroo and Mount Isa faults and are locally fault-bounded against the lower Mount Isa Group to the east across the Barkly shear zone. The latter structure roughly separates the Urquhart Shales and Native Bee Siltstone.
  The Paroo fault is interpreted to have originally formed as a listric extensional fault during deposition of the Kennedy Siltstone and Magazine Shale (Long, 2010), but may be a reactivated predecessor structure that bounded a half graben during Leichhardt Superbasin deposition (Lister, 2002; Betts et al., 2003). The related Mount Isa Fault was active by D2. Whilst both the Paroo and main Mount Isa faults were reactivated by brittle deformation during D2 and D3 to D4, Long (2010) and McLellan et al. (2014) propose that the Paroo Fault was folded during these deformations, while also undergoing reverse shearing. Long (2010) suggests this folding was influenced by uplift of the Sybella batholith which acted as a buttress during D2 shortening (Betts and Lister, 2002). The main Mount Isa and Paroo Faults are broadly parallel at surface to the west of the Mount Isa deposit, northward to beyond George Fisher. Over this interval, the Paroo Fault is in the footwall, and immediately east, of the steeply west dipping Mount Isa Fault. Down dip at Mount Isa however, the Paroo Fault diverges eastward from the Mount Isa Fault where it takes the form of gently north plunging, fault dislocated synform and then an antiform, before dipping steeply east to depth. As a result, the 60 to 65°W dipping Mount Isa Group was truncated at depth by the flat lying section of the folded Paroo Fault and juxtaposed with the underlying similarly steeply dipping Eastern Creek Volcanics and Guide Quartzite at 1655±4 Ma (Southgate et al., 2000; Valenta, 1994; Perkins et al., 1999). To the south of the Mount Isa deposit, this fold is evident where it intersects the surface, and the Paroo Fault diverges sharply from the main Mount Isa Fault juxtaposing the same two sequences. Brittle deformation related to reactivation of these two faults is the major manifestation of D4 brittle deformation in deposits of the Mount Isa district (Chapman, 2004). D2 and D3 shearing and reverse displacement along the folded Paroo Fault is regarded as being important in producing zones of dilation exploited by mineralising fluids (Long, 2010).
  Peak metamorphism apparently occurred between D2 and D3 to the west of the Mount Isa and Paroo faults, where metamorphic grades are significantly higher, increasing from biotite to sillimanite-K feldspar grade westward toward the Sybella batholith (Connors and Page, 1995). Metamorphic temperatures within the Mount Isa Group declined northward from ~350 to 300°C at Mount Isa (Rubenach, 1992) to ~200°C at George Fisher (Chapman,1999, 2001). The age of the peak of metamorphism has been estimated at 1550 to 1570 Ma (SHRIMP U-Pb of zircons in folded pegmatite - 1554 ± 10 Ma; Connors and Page, 1995; chemical Th-U-total Pb isochron method or CHIME of monazite - 1570 Ma with no errors quoted; Hand and Rubatto, 2002; and Pb-Pb step-leach dating of tourmaline from quartz-tourmaline veins -1528±51 and 1577±48 Ma; Duncan et al., 2006).

Deposit Scale Structure

  At George Fisher North, at the deposit and ore lens scale, mudstone and siltstone units generally dip at 30 to 80° west, with small domains extending for ~50 to 100 m down-dip characterised by near-vertical to slightly overturned bedding. These and other observations are interpreted to imply the presence of open folds with both steep and gently dipping axial planes (Chapman, 2004).
  Small, centimetre to tens of centimetre scale tight to open folds are evident in some of the pyritic siltstone units, particularly where these units are mineralised with sphalerite and galena. These small scale folds have a wide range of fold profiles and orientations, with refolded folds and several cleavages evident. The distribution and geometry of these folds have no obvious parasitic relationship to the larger scale open folds detailed above and contrast to the more simple structure of the deposit at the larger scale. Also, in contrast to the Zn-Pb mineralised pyritic units, the largely barren mudstone beds lack similar small scale structures (Chapman, 2004).
  At the mineralised unit scale, Chapman (2004) recognises four foliations at George Fisher, denoted as GFS1 to 4 inclusive. The earliest observed cleavage is usually the bedding-parallel GFS1, characterised by anastomosing penetrative seams of carbonaceous material in mudstones and a slaty cleavage in pyritic siltstones. The earliest folds are tight to isoclinal, east-vergent and north-south striking with 0.5 to 2 m amplitudes, and an axial plane cleavage, GFS2, developed in their hinges. These structures are refolded by younger GFF3 recumbent, and upright GFF4 folds with tight to isoclinal profiles around north-striking axes. These structures commonly produce complex interference patterns at 0.5 to 2 m scales. Chapman (2004) noted that it was not clear whether these small scale folds and multiple GFS1 to GFS4 cleavages were developed during a single continuous deformation or as successive, discrete events comparable to the regional D1 to D4.


  The George Fisher mines are some 20 km to the north along strike from the Mt Isa orebodies, and are located immediately to the east of the major, north-south trending Mt Isa fault system. Both George Fisher South (Hilton) and North are hosted by similar facies in the upper sections of the same host Urquhart Shale which also embrace the Mt Isa Zn-Pb ores, although the orebodies are generally thinner and more disrupted by faulting than the similar ores at Mt Isa. Little copper is known in association with the zinc lead ores at both Hilton and George Fisher.
  Economic grade mineralisation at George Fisher occurs within a north-south striking, west dipping stratigraphic package that is 350 m thick, has a strike length of 1200 m, and extends down dip for >1000 m. It is obliquely truncated by the NNW-SSE trending Spring Creek and Gidyea Creek faults to the north and south respectively, and to the west at depth by the Hanging Wall Fault.
  The George Fisher North deposit comprises 11 mineralised stratigraphic intervals that occur as stacked stratabound lenses, hosted by the Pyritic Siltstone and interbedded Pyritic Siltstone and Banded Mudstone units of the Urquhart Shale as described above. These units are separated by thicker bands of barren Banded Mudstone.
  Mineralised intervals are denoted, from the stratigraphic base upward, as 1, 2 and A to I inclusive, with internal subdivisions denoted by a numerical subscript. The C and D orebodies contain almost half of the known George Fisher resource at grades of around 10.7% Zn, 5.9% Pb and 111g/t Ag. Each of the mineralised interval has distinctive bedding characteristics and tuffaceous marker beds, enabling detailed stratigraphic correlations throughout the ore-bearing sequence (Johnston et al., 1998, Tolman et al., 2002).
  Each of these mineralised intervals is segmented along strike by NE-SW to NNE-SSW striking faults which Chapman (2004) sugget are post-D4, but may be late D4. These faults contain pods of coarse calcite ±dolomite ±quartz ±sphalerite ±fluorite ±pyrite, localised in jogs, linked by graphitic fault gouge.
  The George Fisher South (Hilton) deposit has 7 to 10 stacked ore lenses within a 250 m stratigraphic interval, which have been complicated by intense shortening associated with the Isan Orogeny. The two deposits are separated by a 2 km strike length of intensely faulted, barren shale, mudstone and siltstone.
  These Pyritic Siltstone and interbedded Pyritic Siltstone and Banded Mudstone units carry 10 to >50% fine-grained diagenetic pyrite, distributed over a thickness of 800 m, enveloping both the George Fisher North and South deposits and extending over a strike length of >10 km.
  The George Fisher North deposit has undergone a long and complex history of hydrothermal activity. The the earliest base metal sulphide-bearing assemblages of the deposit were preceded by four distinct gangue-forming events that included:
  i). early dolomite, ankerite and ferroan dolomite cement replacing primary sedimentary grains and rock matrix, such that the rock is enriched in Fe and Mn relative to background Urquhart Shale;
  ii). nodular and banded calcite that predates stylolite formation, manifested as as pervasive bleaching of barren mudstone while nodular calcite is rhythmically interwoven with the fine-grained diagenetic pyrite in finely laminated siltstone and finely banded mudstone;
  iii). widely developed, fine-grained pyrite precipitated within some siltstone units and along carbonaceous stylolite surfaces;
  iv). development of celsian, hyalophane, K feldspar, calcite ±ferroan dolomite ±quartz veins and alteration halos.
  The subsequently introduced economic sulphide mineralization predominantly comprises sphalerite and galena with traces of native silver, tetrahedrite and chalcopyrite. Some of the individual ore lenses can be distinguished on the basis of their metal and gangue mineral assemblages. Alteration that is spatially associated with the late, weak, syntectonic Cu mineralisation includes pyrrhotite, siderite, ferroan ankerite, biotite, chlorite, muscovite and magnetite.
  Four distinct styles of strata-bound Zn-Pb mineralisation are distinguished at George Fisher, based on the dominant sulphide minerals, grain size, structural setting and paragenesis (from Chapman, 2004):
Layer-parallel disseminated sphalerite, which is conformable to bedding at a mm to cm scale, and stratabound at all other levels. The sphalerite is very fine grained, with µm to mm widths, and is subhedral to anhedral with a honey to light brown and red-brown colour. It is disseminated in mm to cm thick diffuse bands subparallel to host rock layering, and is very variable, comprising between 5 and 45% of a particular band. Individual sphalerite grains and aggregates mimic, or are finer grained than the host rock. However, at a mining scale, mineralisation of this style alone is not economic.
  Where developed within siltstone, disseminated sphalerite is dispersed within quartz-calcite-ferroan dolomite laminations, but is rare in pyritic and carbonaceous bands. Where it is found within nodular carbonate layers, it occurs as isolated interstitial grains, infill after rhombohedral calcite, and as mm sized crystal aggregates with irregular grain contacts with carbonates. In mudstone, sphalerite of this type, locally with traces of galena, occurs as grains that are interstitial to carbonate and quartz and are variably intergrown with bitumen.
  Sulphides occur as irregular 'flecks', aggregates, fine disseminations and concentrations that follow and mimic the sedimentary beds, including crossbeds. Some tuffaceous marker beds and white carbonate bands also contain disseminated sphalerite with characteristics that are similar to the mudstone hosted sphalerite.
  These observations indicate disseminated sphalerite mineralisation was precipitated in open space within older nodular carbonate material (which itself transgresses stratigraphy at a district scale) and occurs as interstitial pore fillings in mudstones with bitumen. As such, zones of disseminated sphalerite are also discordant to bedding, as well as being discontinuous at the deposit scale, and replace layers that define current produced bedding at finer scales.
Layer-parallel vein-hosted sphalerite, occurring as fine to medium grained, typically red-brown, semi-massive crystalline sphalerite in discrete layer-parallel bands that are interpreted to be veins. It is coarser grained than sphalerite in the disseminated layer-parallel style, and occurs interstitial to, and surrounds, gangue minerals that have preserved euhedral to subhedral outlines typical of material deposited in open fractures. These veins vary from 1 to 30 mm in thickness, averaging <10 mm, and either occur alone or in stacks of 5 to 10 veins over a width of 10 to 30 cm. They are usually found rhythmically intercalated with pyritic siltstones and at contacts between thin-banded pyritic siltstone and mudstone units. They have planar-parallel or folded margins, and variably sharp to diffuse boundaries. The latter occur where they are bounded by sedimentary layers with abundant disseminated sphalerite. These veins, which comprise 30 to 95 vol.% interstitial sphalerite also carry various combinations of paragenetically associated quartz, calcite ±K feldspar ±hyalophane [(K,Ba)Al(Si,Al)Si
2O8] ±celsian [Ba(Al2Si2O8)] ±hydrophlogopite ±ferroan dolomite.
  Silicate and carbonate vein filling minerals have significantly coarser grain sizes and are readily distinguishable from those of the wall-rocks. Some of the quartz and carbonate have relict euhedral crystal outlines, whilst all other associated minerals, except hydrophlogopite, have brecciated and corroded grain contacts with sphalerite. These relationships imply sphalerite and hydrophlogopite were deposited after carbonate, quartz and feldspar during a discrete but closely related episode of open-space vein fill. These textures are further interpreted to suggest earlier precipitated silicate and carbonate vein infill was brecciated on a small scale, dissolved and/or replaced by sphalerite during later vein opening and mineral precipitation. Layer-parallel disseminated, sphalerite and hydrophlogopite in the adjacent wall rocks are interpreted to be an alteration selvage to the veins (Chapman, 2004). Crenulation of sphalerite vein margins by GFS2 and intensification of the GFS2 cleavage along vein margins are taken to imply formation of sphalerite veins prior to GFD2.
Breccia-hosted sphalerite, is characterised by a finer grain size and greater internal structural complexity compared to the vein-hosted sphalerite, although both have an overall strata-bound character. The breccia matrix is composed of very fine grained sphalerite with minor pyrrhotite and galena, enclosing 5 to 50 vol.% clasts. The clasts include irregular, mm to cm size subrounded fragments of folded and boudinaged mudstone and pyritic siltstone, sugary ferroan dolomite vein fragments, and aggregates of calcite ±ferroan dolomite ±K feldspar ±hyalophane ±quartz, similar to the infill in the layer-parallel sphalerite veins.
  This brecciated mineralisation postdates sugary ferroan dolomite veining, which crosscuts both vein-hosted and layer-parallel disseminated sphalerite, but occur as clasts in the breccia. Textural relationships suggest breccia formation was synchronous with, or postdated GFF2 fold development.
  Breccia-hosted sphalerite is interpreted to be the product of mechanical deformation of pre-existing vein-hosted sphalerite, resulting from strain partitioning in sphalerite veins due to rheological differences between sulphides, mudstone layers and ferroan dolomite veins at cm to tens of cm scales (Chapman, 2004).
Vein and breccia hosted galena, comprises:
 i). Discordant vein hosted galena composed of fine to coarse grained galena ±pyrite ±sphalerite ±pyrrhotite ±carbonate ±chlorite ±quartz. Subhedral galena is interstitial to euhedral pyrite, forming irregularly interlocking crystal aggregates with sphalerite and pyrrhotite. Veins generally have planar to irregular margins, are a few mm thick, and are discontinuous with tapering extremities. They typically form orthogonal networks perpendicular to bedding and are structurally continuous with galena breccias.
 ii). Breccia hosted galena occurs as the following variants:
  - Coarse-grained galena-rich breccias with a matrix principally composed of ~40 to 70% galena and accessory pyrite, enclosing fragments that vary from a few mm to a few cm in size to larger blocky to subrounded clasts that predominantly comprise mudstone and pyritic siltstone as well as layer-parallel, disseminated and vein-hosted sphalerite mineralisation. Additionally there are microscopic clasts of wall rocks and irregular, subrounded monominerallic silicate, carbonate or sphalerite. These breccias vary from 1 to 100 cm, but are mostly 2 to 10 cm thick, occurring as concordant bands with planar, parallel and discordant margins.
  - Fine-grained galena-rich breccias, where the contained galena has a dull lustre and occurs as a fine grained 40 to 75% matrix enclosing microscopic to cm sized clasts similar to those of the coarse-grained galena breccias described above. Clasts usually have oval to spheroidal to elongate shapes and are rounded, sometimes with a bedding-parallel preferred orientation. These breccias occur as conformable bands with planar-parallel margins and range from 1 to 15 cm in thickness.
  - Mixed sulphide-rich breccias comprise ~70% matrix composed of variable amounts of fine-grained galena, sphalerite, pyrrhotite and pyrite. These sulphides are irregularly distributed, producing mineralogically distinct bands to wavy and flamelike domains within the breccia matrix at a mm to cm scale. Clasts range from sub-mm to a few cm in size, and are typically well rounded to ovoid to irregular in shape, and include fragments of variably folded host rock, layer parallel, disseminated and vein hosted sphalerite mineralisation, as well as monomineralic calcite, ferroan dolomite and quartz. These breccias are typically 3 to 10 cm thick and have sharp to wavy margins, subparallel to bedding.
  There is a consistent timing relationship of galena rich mineralisation compared to all zinc mineralisation styles on a deposit scale, indicating it represent the youngest ore-forming event at George Fisher. The discordant galena veins are structurally continuous with the breccias, indicating the two are coeval. The veins commonly occur in NW-SE to NE-SW trending conjugate sets that crosscut GFF2 and GFF3 folds, with an orientation compatible with their formation during GFF4 fold development.

  After extensive Ph.D. investigations of the George Fisher North ore deposit and hosts at all scales, Chapman (1999; 2001; 2004) concluded that the earliest sphalerite mineralisation is hosted by bedding parallel veins. Sphalerite in these veins occurs as infill with hydrophlogopite, and as an apparent alteration product of precursor carbonate-quartz-Ba-K feldspar vein fill. Layer-parallel, disseminated sphalerite occurs as infill after the replacement of carbonate in siltstones, mudstones and nodular carbonate layers and is interpreted to represent a halo to the veins. It was further observed that the earliest sphalerite mineralisation postdated the emplacement of nodular carbonate, which transgresses stratigraphy at a district scale, and occurred prior to district-wide folding marked by the onset of GFD2 at George Fisher. The emplacement of galena in late tectonic structural sites post-dated all sphalerite mineralisation and was emplaced during GFD4 deformation. There is an intimate deposit scale spatial association and grade variation between galena and sphalerite, suggesting a common origin for Zn and Pb and remobilisation in the late stages of deformation, supported by the observation of significant textural reconstitution of ore during regional folding. Mineralisation styles indicative of sea-floor deposition were not observed, although their complete obliteration and remobilisation cannot be entirely excluded.

George Fisher North is concealed and is partially connected to George Fisher South (Hilton) on its northern margin. Both mines are underground operations and are owned and operated by Glencore's Mt Isa Zinc Division.

Sub-economic mineralisation follows the strike of the Urquhart Shale for 4 km to the north and south of the two economic orebodies.

The original resources were estimated as follows (Forrestal 1990, Valenta 1994, Chapman 2004):
    George Fisher South (Hilton):  120 Mt @ 10.2% Zn, 5.5% Pb, 100 g/t Ag,
    George Fisher North:               108 Mt @ 11.1% Zn, 5.4% Pb, 93 g/t Ag.

JORC compliant Ore Reserves and Mineral Resources (at June 2006) were as follows (X-Strata 2007):
  George Fisher South underground
      Proved reserves - 12.5 Mt @ 8.3% Zn, 5.7% Pb, 127 g/t Ag
      Probable reserves - 5.9 Mt @ 7.8% Zn, 5.8% Pb, 126 g/t Ag
      Measured resources - 25.3 Mt @ 9.7% Zn, 6.9% Pb, 150 g/t Ag
      Indicated resources - 10.6 Mt @ 9.2% Zn, 6.6% Pb, 139 g/t Ag
      Inferred resources - 10 Mt @ 10% Zn, 6% Pb, 100 g/t Ag
  George Fisher North underground
      Proved reserves - 11.3 Mt @ 8.9% Zn, 4.7% Pb, 91 g/t Ag
      Probable reserves - 15.1 Mt @ 8.3% Zn, 3.9% Pb, 75 g/t Ag
      Measured resources - 14.5 Mt @ 10.4% Zn, 5.2% Pb, 101 g/t Ag
      Indicated resources - 27.9 Mt @ 9.5% Zn, 4.0% Pb, 74 g/t Ag
      Inferred resources - 45 Mt @ 9% Zn, 4% Pb, 80 g/t Ag

Mine production in 2005 totalled 2.7 Mt @ 8.3% Zn, 5.0% Pb, 115 g/t Ag.

JORC compliant Ore Reserves and Mineral Resources (at 31 December, 2011) were as follows (X-Strata 2012):
  George Fisher South underground
      Proved + probable reserves - 19.6 Mt @ 6.5% Zn, 4.4% Pb, 96 g/t Ag
      Measured + indicated resources - 49.8 Mt @ 8.6% Zn, 5.9% Pb, 125 g/t Ag
      Inferred resources - 23 Mt @ 8% Zn, 5% Pb, 113 g/t Ag
  George Fisher North underground
      Proved + probable reserves - 61.8 Mt @ 7.5% Zn, 3.6% Pb, 62 g/t Ag
      Measured +indicated resources - 106.2 Mt @ 8.6% Zn, 3.7% Pb, 63 g/t Ag
      Inferred resources - 65 Mt @ 8% Zn, 4% Pb, 69 g/t Ag
  Handlebar Hill open pit primary
      Proved + probable reserves - 2.1 Mt @ 7.6% Zn, 3.4% Pb, 54 g/t Ag
      Measured +indicated resources - 6.9 Mt @ 6.9% Zn, 2.3% Pb, 43 g/t Ag
      Inferred resources - 1 Mt @ 5% Zn, 2% Pb, 30 g/t Ag
  Handlebar Hill open pit oxide
      Proved + probable reserves - 0.5 Mt @ 0.4% Zn, 8.5% Pb, 89 g/t Ag
      Measured +indicated resources - 0.6 Mt @ 0.4% Zn, 7.8% Pb, 85 g/t Ag
      Inferred resources - Nil

JORC compliant Ore Reserves and Mineral Resources (at 31 December, 2018) were as follows (Glencore Resources and Reserves report, 2018):
  George Fisher South underground
      Proved + probable reserves - 16.7 Mt @ 6.2% Zn, 4.4% Pb, 99 g/t Ag
      Measured + indicated resources - 56 Mt @ 8.3% Zn, 5.0% Pb, 106 g/t Ag
      Inferred resources - 27 Mt @ 8% Zn, 4% Pb, 87 g/t Ag
  George Fisher North underground
      Proved + probable reserves - 72 Mt @ 7.0% Zn, 3.1% Pb, 52 g/t Ag
      Measured +indicated resources - 168 Mt @ 8.9% Zn, 3.5% Pb, 55 g/t Ag
      Inferred resources - 53 Mt @ 9% Zn, 4% Pb, 57 g/t Ag
  Handlebar Hill open pit primary
      Proved + probable reserves - Nil
      Measured +indicated resources - 5.2 Mt @ 6.6% Zn, 2.2% Pb, 37 g/t Ag
      Inferred resources - 0.8 Mt @ 5% Zn, 2% Pb, 30 g/t Ag
  Handlebar Hill open pit oxide
      Proved + probable reserves - 0.5 Mt @ 0.4% Zn, 8.5% Pb, 89 g/t Ag
      Measured +indicated resources - 0.6 Mt @ 0.4% Zn, 7.8% Pb, 85 g/t Ag
      Inferred resources - Nil
  NOTE: Mineral Resources are inclusive of Ore Reserves.

Mine production at George Fisher North and South for the period 1 January 2018 to 31 December 2018 totalled:
    2.9 Mt @ 7.3% Zn, 3.9% Pb, 68g/t Ag.

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Vedanta Resources - background

Vedanta Resources plc is a London listed diversified and integrated FTSE 250 metals and mining group with annual sales of USD 1.9 billion. The company's principal operations are located in India, where it has a major market share in each of its main metals: aluminium, copper, zinc and lead. Its principal zinc assets include three main mines, Rampura-Agucha (reserves+resources approx. 75 Mt @ 12.7% Zn, 1.8% Pb), Rajpura-Dariba (reserves+resources approx. 42.5 Mt @ 6.3% Zn, 2.1% Pb) and Zawar (reserves+resources approx. 42 Mt @ 4.4% Zn, 2.3% Pb), all in Rajasthan. These mines supply the company's three smelters, two in Rajasthan and one in Andhra Pradesh (with a combined capacity of 400 000 tpa Zn and 35 000 tpa Pb). The company's annual mine production has been around 210 000 t. of contained Zn, but is currently being expanded to 400 000 tpa, and is managed through its subsidiary Hindustan Zinc Ltd. Vedanta's Indian copper interests are principally smelting (300 000 tpa Cu metal capacity) through its subsidiary Sterlite Industries. There are also substantial copper operations in Zambia (51% of Konkola Copper Mines 250 000 tpa Cu) and 2 copper mines in Australia (Thalanga and Mt Lyell). The company's aluminium mining and smelting capacity is currently being expanded to 400 000 tpa, supported by its own power stations.

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For more information contact:   T M (Mike) Porter, of Porter GeoConsultancy   (

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T M (Mike) Porter of Porter GeoConsultancy Pty Ltd on behalf, and to the specification, of the client.

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