Sheep Creek, Black Butte
Cu Co Ag
Super Porphyry Cu and Au|
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The Sheep Creek or Black Butte sedimentary rock-hosted copper (-cobalt-silver) deposit is located ~32 km north of White Sulphur Springs and ~180 km east of Missoula, Montana, western USA, and lies within the Helena Embayment, the eastern branch of the Mesoproterozoic Belt-Purcell Basin.
The Mesoproterozoic Belt-Purcell Basin covers an area of ~200 000 km2 with an ~750 km long, NNW trending main axis, extending from southern British Columbia in Canada to southern Idaho in the USA, with an eastern, east-west trending limb (Helena Embayment) representing a rift extending into the North American continent. The basin fill, collectively referred to as the Belt (Purcell in Canada) Supergroup, was deposited between ~1470 and ~1400 Ma (Anderson and Davis, 1995; Sears et al., 1998; Evans et al., 2000), and is divided into four groups. From the bottom to top (Graham et al., 2012):
Lower Belt - in the north-south oriented western axis of the basin consists of composed of an up to 12 km thick rift-fill sequence of siliciclastic turbidite-dominated sedimentary rocks of the Prichard Formation (and Canadian equivalent, Aldridge Formation; Cook and Van der Velden, 1995). Up to 40% of this unit may locally consist of synsedimentary mafic sills (Höy et al., 2000). The base of this sequence is only seen near the Sullivan deposit in British Colombia, where the Aldridge Formation overlies the 2,000 m thick Fort Steele Sandstone of fluvial or marine origin (Chandler, 2000).
The Helena Embayment is asymmetric, dominated by the coarse-grained LaHood Conglomerate submarine fans along its southern margin that give way to finer-grained lower Belt sedimentary rocks to the north. The LaHood Conglomerate is a pebbly arkose and arkosic paraconglomerate, including clasts of metamorphic rock derived from the south (McMannis, 1963). This sequence on the southern margin of the Helena Embayment dip toward the north and are unconformably overlain by Cambrian sandstone (McMannis, 1963).
On the northern margin of the embayment, near Sheep Creek, the Neihart Quartzite (Weed, 1899) unconformably overlies Palaeoproterozoic basement and is overlain by wavy laminated to bedded black shale, siltstone and thin sandstone of the Chamberlain Formation (Walcott, 1899). Toward the upper part of the Chamberlain Formation, carbonate beds containing distinct molar-tooth structures (crumpled calcite veins infilling gas-escape voids) are interbedded with silty shale (Godlewski and Zieg, 1984; Zieg, 1986).
The Chamberlain Formation grades upward into the Newland Formation, which is predominantly mixed siliciclastic-carbonate sedimentary rocks that become more carbonate rich up-section. In the Sheep Creek area, the Newland Formation is conformably overlain by the Greyson Formation siltstone and argillite (Walcott, 1899). The Newland Formation varies from ~610 to >2900 m, and the total thickness of basin fill in the Embayment has been estimated at ~6 km, and as such is appreciably thinner than that in the main basin to the west (Cook and Van der Velden, 1995; Höy et al., 2000).
Ravalli Group is composed of shale, siltstone and sandstone deposited in shallow-water environments, reflecting of a major regressive event (Connor et al., 1984).
Piegan Group or Middle Belt carbonate represent a transgressive carbonate-rich sequence with increased siliciclastic sedimentary rocks to the west (Grotzinger, 1986; Wallace et al., 1999) deposited on rocks of the Ravalli Group.
Missoula Group is composed of a sandy alluvial apron that gives way to shallow marine siliceous to carbonate mudstones to the east and north (Chandler, 2000).
The upper part of the Middle Belt carbonate and the entire Missoula Group are absent from the Helena Embayment.
Höy et al., 2000 regard the Belt-Purcell Basin as a rift basin, while Chandler (2000) interprets the basin as a "passive" rift, without pre-rift doming and volcanism but with the preservation of a basal quartzite. The sequence was deposited in <100 m.y., with average sediment accumulation rates of between 25 and 57 cm/ka in the main basin (Evans et al., 2000; Lydon, 2000).
In the immediated Sheep Creek area, Palaeoproterozoic granitic gneiss basement is unconformably overlain by 200 to 280 m of pinkish and white supermature quartzite of the Neihart Formation, possibly representing either an extensive fluvial to aeolian cratonic sand sheet deposited prior to initiation of the basin development, or a sand wedge filling the initial
Belt-Purcell rift basin with initial fluvial and subsequent shallow marine sedimentation. This sequence grades up into the 460 to 945 m thick Chamberlain Formation, composed of unevenly laminated, interbedded, black shale, siltstone and sandstone with some carbonate-rich beds near the top (Walcott, 1899; Keefer, 1972), interpreted as a near-shore mudflat deposit (Godlewski and
The Chamberlain Formation is overlain by the ~1100 m thick Newland Formation (Walcott, 1899), which has been subdivided into two members, the lower of which predominantly consists of laminated, variably carbonate bearing shale with locally interbedded debris flows, turbidites and carbonate beds. The upper member comprises variably carbonate rich shale and more abundant planar-bedded carbonate units (Zieg, 1986) that can be subdivided into two upward shoaling carbonate-terrigeneous
sequences. The Newland Formation was deposited in a subtidal setting, which Himes and Petersen (1990) estimated to have been deposited at water depths of >850 m on the basis of fluid inclusion data in diagenetic barite, representing a much greater depth than the underlying Chamberlain Formation (Graham et al., 2012).
The youngest Mesoproterozoic formation in the immediate Sheep Creek area is the Greyson Formation, a
<2000 m thick unit of parallel-bedded medium- to dark-grey siltstone with Bouma sequence bedding that grades upward
to greenish and pinkish siltstone, argillite and quartzite, with ripple marks and mud cracks becoming more frequent upwards (Connor et al., 1984). This formation represents the final stages of the second shoaling upward sequence of the upper Newland Formation, representing the transition from carbonate to siliciclastic facies in the Helena Embayment (Graham et al., 2012).
At Sheep Creek, the Newland Formation has a complex stratigraphic architecture interpreted to be controlled by syn-sedimentary faulting. Debris flow deposits (conglomeratic debrites) are locally abundant adjacent to the Volcano Valley and related faults and may fill palaeochannels or form aprons shed from fault scarps. No basement-rock, Neihart or Chamberlain Formation clasts have been recognized, although clasts of mineralised and hydrothermally altered rock of the Newland Formation are locally recognised. The lower member siliciclastic to mixed siliciclastic-carbonate shale and turbidites occur in the hanging wall of the Volcano Valley reverse fault and in the footwall as a fault bracketed structural compartment. The siliciclastic beds are generally dark, from medium to dark grey and black, while carbonate-rich beds are medium to light grey, giving the rock a zebra-striped appearance. Detrital silt-sized quartz, dolomite and white mica are the main components of the beds, with local minor chlorite and feldspar. Organic contents range from <0.1 to >3 wt.% (Lyons et al., 2000). The carbonate content of the beds is highly variable, locally, approaching pure dolomite (Graham et al., 2012).
The Newland Formation contains syngenetic to diagenetic pyrite-rich rock over an interval of at least a 24 km along the Volcano Valley Fault, and up to 8 km toward the south (Zieg et al., 1999). In the Sheep Creek area, pyrite-rich beds comprise five distinct, laterally continuous, to semi-continuous intervals that vary from several metres of cm-scale beds within shale and turbidites, to massive lenses of nearly 100% pyrite up to 90 m thick.
Two main sulphide zones are recognised. The lower comprises massive pyrite beds in the lowest Newland Formation in the footwall of the Volcano Valley Fault, contained mostly within debris flow-rich strata or slumped shale. The upper sulphide zone pyrite beds are in the hanging wall of the Volcano Valley Fault, and occur above the main debris flow deposits and form a ~2 km long blanket that extends from the Volcano Valley Fault south to near the Black Butte Fault at Sheep Creek.
Three localised massive, stylolitic dolomite to dolomitic limestone beds, each from ~15 to >40 m thick, are found above the lower sulphide zone in sections of the Sheep Creek area, while dolomite beds occur in the eastern part of the area beneath the basal debris flows of the upper sulphide zone (Graham et al., 2012).
The lower sulphide zone consists of one to four stratabound horizons of semimassive to massive pyrite with a cumulative thickness of 2 to 21 m in lowest Newland Formation in the footwall of the Volcano Valley Fault. The host sequence comprises (Graham et al., 2012):
Footwall un-silicified shale/silty shale, grades rapidly into underlying Chamberlain Formation, with rare, poorly developed replacement dolomite.
Footwall conglomerate (7 to 23 m), the upper portion of which is wholly silicified, and contains sporadic thin (2 to 4 mm) quartz and/or dolomite veins ± chalcopyrite.
Lower sulphide zone(s) (2 to 21 m), which are intensely silicified in the central part of the deposit, while dolomite dominates to west and east, containing massive, largely coarse-grained pyrite, interbedded with debris flows, with chalcopyrite usually concentrated near the base.
Hanging wall, debris flow-rich sediments with irregular dolomite veins and porphyroblasts, commonly
more pyritic near base, where alteration is up to 30m thick.
The paragenesis of the lower sulphide zone began with laminated to massive, syngenetic to early diagenetic, fine-grained pyrite and associated net-textured pyrite, with rare barite crystals. The crosscutting relationship of associated dolomitisation, Fe sulphide deposition, and silicification of host rocks indicate formation at depth below the sediment-water interface. Dolomite alteration was interspersed with intervals of secondary iron sulphide intergrowth and replacement forming coarse, commonly grossly tabular and crustiform crystals. Apatite and local associated recrystallizsation of clays to illite, muscovite and chlorite are most abundant within coarsely silicified rock. Silica and dolomite were likely introduced in multiple stages. Chalcopyrite occurs within alteration dolomite indicating that at least weakly cuprous fluids were present during carbonate growth, although the majority of chalcopyrite and ore-stage pyrite was deposited during and after silicification, replacing dolomite, occurring in dolomite veinlets and fractures that crosscut silicified rock. Quartz, dolomite, and chalcopyrite all replaced early barite (Graham et al., 2012).
The upper sulphide zone contains chalcopyrite-bearing massive pyrite and consists of one to four stacked pyrite horizons near the top of the lower Newland Formation in the hangingwall of the Volcano Valley Fault, and ~50 m below the carbonate marker bed at the transition from the lower to upper member of the Newland Formation. The pyritic intervals are separated by several metres of shale and/or debris flows with rare thin silicified dolomite or chert beds. The basal pyritic interval is generally the thickest and is usually immediately underlain by unsilicified to weakly silicified shale and, rarely, debris flows. The host sequence comprises (Graham et al., 2012):
Footwall sedimentary rocks, largely composed of unaltered shale and debris flow sedimentary rocks, with local dolomite and barite replacement, and occasional pyrite in debris flow matrix.
Upper sulphide zone (7 to 63 m), occurring as 1 to 5 horizons of barite-rich, massive pyrite with sedimentary interbeds. Silicification is most common near the base of the zone, with chalcopyrite mostly in the lowest part of the zone, generally associated with barite or quartz.
Mixed debris flows, turbidite, shale (35 to 80 m).
Lower Newland to Upper Newland Formation transition which immediatley overlies a ~5 m thick marker dolomite unit.
The paragenesis of the upper sulphide zone is similar to that of the lower sulphide zone, with early diagenetic microcrystalline to framboidal, locally cobaltiferous pyrite formation grossly coincident with extensive barite and minor dolomite crystal growth. Secondary pyrite and marcasite overgrew and cemented earlier pyrite and replaced barite, and was sometimes accompanied by inclusions of galena, sphalerite and chalcopyrite. Quartz also replaced barite, with the most intense silicification in the north of the deposit, although the intensity of silicification also generally decreases upward and outward. Thin siliceous layers may be either original chert beds or due to replacement of carbonate beds. As in the lower sulphide zone, chalcopyrite and minor pyrite deposition was associated with silicification late in the paragenetic sequence. Chalcopyrite and tennantite replaced fine-grained pyrite, barite, and locally, crystalline dolomite, and overgrew and formed veinlets in later coarse-grained pyrite. Textures suggest that silicification continued in some areas after chalcopyrite deposition (Graham et al., 2012).
Cominco American Inc. outlined geologic resources to 1993 (Zieg et al., 1999) of:
Lower sulphide zone - 4 Mt @ 4% Cu;
Upper sulphide zone - 4.5 Mt @ 2.5% Cu, 0.1% Co
Tintina Resources Inc. (2012) has delineated the following mineral resources in the Upper Zone:
Measured resource - 2.66 Mt @ 2.99% Cu, 0.12% Co, 16.3 g/t Ag;
Indicated resource - 6.52 Mt @ 2.77% Cu, 0.13% Co, 15.5 g/t Ag;
Inferred resource - 1.26 Mt @ 2.52% Cu, 0.10% Co, 15.2 g/t Ag.
The most recent source geological information used to prepare this summary was dated: 2012.
This description is a summary from published sources, the chief of which are listed below.
© Copyright Porter GeoConsultancy Pty Ltd. Unauthorised copying, reproduction, storage or dissemination prohibited.
Graham G, Hitzman M W and Zieg J, 2012 - Geologic Setting, Sedimentary Architecture, and Paragenesis of the Mesoproterozoic Sediment-Hosted Sheep Creek Cu-Co-Ag Deposit, Helena Embayment, Montana : in Econ. Geol. v.107 pp. 1115-1141|
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