In the MSM area, Neoproterozoic successions were deposited in extensional regimes (Narbonne and Aitken 1995; Turner and Long 2008), possibly in the vicinity of what is now Australia, Siberia, South China, Antarctica, or West Africa (Hoffman 1991; Park et al. 1995; Rainbird et al. 1996; Harlan et al. 2003; Sears and Price 2003; Li et al. 2004; Li et al. 2007; Nelson and Colpron 2007; Evans 2009). The Neoproterozoic strata can be divided into two groups; the older Mackenzie Mountains Supergroup (Long and Turner 2012; Turner and Long 2012) and the younger Windermere Supergroup (Fig. 3A; Narbonne and Aitken 1995). The 4 to 6 km-thick Mackenzie Mountains Supergroup consists of, from oldest to youngest, sedimentary rocks of the “H1-unit” (now Tobacco Formation; Turner and Long 2012), the Tsezotene Formation, and the Katherine and Little Dal Groups (Fig. 3A; Long et al. 2008; Martel et al. 2011). Uranium-lead detrital zircon populations with ca. 1083 Ma and 1010 Ma dates from the Katherine Group provide a maximum age of deposition (Rainbird et al. 1996; Leslie 2009). A 778 Ma gabbro plug south of Coates Lake (Jefferson and Parrish 1989) and the cross-cutting ca. 780 Ma Tsezotene sills (Harlan et al. 2003) provide a minimum age for the Mackenzie Mountains Supergroup. The Tsezotene sills are associated with the “Little Dal basalts” and are part of the Gunbarrel event (Dudás and Lustwerk 1997; Harlan et al. 2003). Stratigraphy, sedimentology, and regional correlations indicate that the overall depositional setting for the Mackenzie Mountains Supergroup was likely an epicra-tonic basin (Narbonne and Aitken 1995; Long et al. 2008; Turner and Long 2008).

The Coates Lake Group is a sequence of sabkha- and lagoon-related sedimentary rocks that were deposited in a terrestrial rift environment (Jefferson and Parrish 1989). The Coates Lake Group is positioned unconformably above the Little Dal Group and Little Dal basalt and is unconformably overlain by the Rapitan Group of the Windermere Supergroup (e.g., Eisbacher 1978; Jefferson and Ruelle 1986; Dudás and Lustwerk 1997).

The 5 to 7 km-thick, rift-related Windermere Supergroup unconformably overlies the Coates Lake Group and the Mackenzie Mountains Supergroup (Fig. 3A; Eisbacher 1978; Narbonne and Aitken 1995). From the base upward, the Windermere Supergroup in the MSM includes the <755 Ma glacial marine/rift-related Rapitan Group (Yeo 1981; Ross and Villeneuve 1997), the deep-water, shaly Twitya Formation, shallower-water carbonate and siliciclastic strata of the Keele Formation, the glacial-related Ice Brook Formation, deep water shale and turbidites of the Sheepbed Formation, platformal carbonate strata of the Gametrail Formation, thick shallow-water, terrigenous, clastic strata of the Blueflower Formation, and the poorly exposed, shallow-water, carbonate-dominated rocks of the Risky Formation (Fig. 3A; Narbonne and Aitken 1995; MacNaughton et al. 2008).

The Neoproterozoic–Cambrian boundary is preserved in the main ranges of the Mackenzie Mountains, but sub-Cambrian erosion has cut progressively downward east to the mountain front, where Cambrian strata directly overlie lower units of the Mackenzie Mountains Supergroup. The Cambrian through Devonian (with minor Permian) strata were deposited on the Neoproterozoic strata in an evolving passive margin. In general, the sedimentary rocks are preserved as platformal facies in the northeast (Mackenzie Platform) and basinal facies to the southwest (Selwyn Basin; Figs. 2 and 3B; Morrow 1991; Gordey and Anderson 1993; Cecile et al. 1997).

Paleozoic platformal sedimentation was complex and influenced by variations in crustal thicknesses inherited from sub-Windermere and Cambrian extension (Cecile et al. 1997; Hadlari et al. 2012). This led to sedimentation on a number of segmented platforms, arches, and basins, resulting in numerous formations with complex lateral thickness variations and facies relationships. For simplicity, this is referred to collectively as the Mackenzie Platform (Fig. 2; Morrow 1991; Gordey and Anderson 1993; Cecile et al. 1997; Martel et al. 2011). The oldest Cambrian strata in the Mackenzie Platform consist of shallow-marine to flu-vial quartz sandstone (Backbone Ranges Formation). This is overlain by shallow-water Cambrian to Devonian dolostone and limestone assigned to various formations (e.g., Sekwi, Rock-slide, Broken Skull, Sunblood, Franklin Mountain, Mount Kindle, Tsetso, Camsel, Sombre, Arnica, and Bear Rock Formations; Fig. 3B), and capped by Middle Devonian open-marine fossiliferous limestone (e.g., Landry, Grizzly Bear, Headless, Nahanni, and Hume Formations; see Morrow 1991; Gordey and Anderson 1993; Pyle and Jones 2009; Martel et al. 2011). A significant feature of the Mackenzie Platform is the Mackenzie Arch, a northwest-trending locus of coalescing unconformities and formational thinning. East of the arch, formation thicknesses increase, but only modestly, and then farther east units thin again to merge with unfolded Paleozoic platform cover of the craton (Interior Platform; e.g., Cecile et al. 1997). West of the arch the platformal units increase substantially in thickness to the facies transition into the Selwyn Basin (Morrow 1991).

The relatively deeper water Selwyn Basin strata were deposited concomitantly with the sedimentary rocks on the Mackenzie Platform. The oldest exposed units in the Selwyn Basin are thick, turbiditic, clastic rocks of the Neoproterozoic through Cambrian Hyland Group. Lower and Middle Cambrian shales are sparsely preserved beneath the sub-Upper Cambrian unconformity. The Upper Cambrian is marked by basinal limestone of the Rabbitkettle Formation (Fig. 3B). Ordovician to Devonian strata are dominated by siltstones, shales, and cherts of the older Road River Group and younger Earn Group (Fig. 3B; Gordey and Anderson 1993; Pyle and Jones 2009; Martel et al. 2011).

During the Late Devonian to Mississippian there was an abrupt change in deposition style, reflecting complex tectonic events (e.g., Colpron et al. 2007; Lemieux et al. 2011). The lack of evidence of compressional deformation, the presence of local syn-sedimentary normal faults and of local felsic volcanic rocks are interpreted to reflect extension or transtension. As Devono–Mississippian tectonism waned, clastic sedimentation was succeeded by lower Mississippian mixed carbonate and siliclastic strata likely deposited on a muddy, shallow-marine shelf (Tsichu Group, Cecile 2000; Mattson Formation, Potter et al. 1993; Fig. 3B). Preservation of Carboniferous to Permian strata is fragmentary and limited to the westernmost and southern parts of the Mackenzie Mountains (Fig. 3B; Gordey and Anderson 1993; Martel et al. 2011).

Triassic and Cretaceous sedimentary units are only preserved locally (Figs. 2 and 3B). The Cretaceous strata are foredeep deposits that reflect the collisional evolution of the continental margin to the west (Coney et al. 1980; Mair et al. 2006a; Martel et al. 2011). Mesozoic igneous rocks are dominated by intermediate to felsic, ca. 110 to 90 Ma Cretaceous plutons that intruded Proterozoic and Paleozoic Selwyn Basin strata within the Selwyn Mountains (Figs. 2 and 3B; Hart et al. 2004a, b; Rasmussen et al. 2007). The plutons represent the southeastern extent of the Tombstone–Tungsten Suite (Hart et al. 2004a), and are interpreted to have formed in a distal, back-arc setting in response to collision of the Peninsular–Alexander–Wrangellia Superterrane with the western margin of North America (Mair et al. 2006a).

The oldest recognized mineral deposit type in the MSM is stratabound copper hosted by the <780 Ma Coates Lake Group (Fig. 3A). This region is commonly referred to as the Redstone copper belt, and the mineralization is generally preserved in the hanging wall of the Plateau fault (Fig. 4). The Coates Lake Group is divided into three formations; alluvial-fan and playa-lake deposits of the Thundercloud Formation grade into the overlying sabkha deposits of the Redstone River Formation, which is overlain by marine carbonate rocks of the Coppercap Formation (Figs. 3 to 6; Jefferson 1983; Jefferson and Ruelle 1986).

Exposures of the Coates Lake Group are restricted to the central part of the Mackenzie Mountains and this unit is thickest at Coates Lake (Fig. 4; Aitken 1977; Jefferson and Ruelle 1986; Colpron and Jefferson 1998; Jefferson and Colpron 1998). In the northern Mackenzie Mountains, the succession pinches out rapidly and may have been eroded prior to deposition of the overlying Rapitan Group; scattered remnants are found as far north as 64º45’ N (Fig. 4; Eisbacher 1981; Jefferson and Parrish 1989; Martel et al. 2011). South of Coates Lake, the Coates Lake Group may be present in the subsurface, but its extent is unknown. Thickness and lithofacies patterns in the Coates Lake Group define a series of depositional embayments (Fig. 4; Jefferson 1983; Jefferson and Parrish 1989) caused by syn-depositional normal faulting that formed a series of half-grabens (Eisbacher 1981; Jefferson and Ruelle 1986; Jefferson and Parrish 1989). This extensional activity was probably related to incipient opening of the paleo-Pacific Ocean and the first stage of Rodinia break-up (e.g., Colpron et al. 2002; Li et al. 2007).

At least 25 Cu (± Ag) occurrences are recognized (Fig. 4), of which the deposit at Coates Lake is the most significant. Historic resource estimates suggest that it contains 33.6 million tonnes grading 3.9% Cu and 9 g/t Ag (Table 1; Hitzman et al. 2005). A thorough review of the deposit geology is provided by Chartrand and Brown (1985), Jefferson and Ruelle (1986), and Chartrand et al. (1989). The Jay prospect to the north is currently the second-most-significant deposit, with a historic resource estimate of 1.2 million tonnes at 2.7% Cu (Hitzman et al. 2005).

Very low-grade disseminated copper sulphide mineralization is widespread, but the most noteworthy copper occurrences are generally stratabound or stratiform lenses in the “Transition Zone” between the Redstone River and Coppercap Formations (Figs. 3A, 5A, and 6; Jefferson and Ruelle 1986). The relationship between ore bodies and host rocks indicates that these deposits are of the Kupferschiefer-type (Hitzman et al. 2005). Ore mineralization in the district is believed to be the result of fluid mobilization in an actively extending basin (Fig. 6; Jefferson and Ruelle 1986; Chartrand et al. 1989), a concept that is supported by sedimentary and ore textures, which suggest that mineralization took place during very early diagenesis (Jefferson and Ruelle 1986; Chartrand et al. 1989). Copper and associated metals were liberated by oxidation of metals present in aquifer rocks, transported in either compaction-derived or meteoric, recharge-related chloride brines through sulphate-bearing red-beds, which then interacted with a reducing interface to precipitate sulphides (Fig. 6; Chartrand et al. 1989; Brown 2005; Hitzman et al. 2005). This interface is the Transition Zone between the Redstone River and Coppercap Formations, and an upward zoning from copper-rich to copper-poor ore minerals indicates sub-vertical fluid flow (Fig. 6B; Jefferson and Ruelle 1986). The copper source is suggested by Chartrand et al. (1989) to be the now red and oxidized evaporitic beds of the Redstone River Formation, or the volcaniclastic conglomerates of the underlying Thundercloud Formation. It is likely that the underlying “Little Dal basalt” (Dudás and Lustwerk 1997) provided most of the copper, which was either leached directly from the basalt, or liberated from volcanic detritus in overlying units (e.g., Thundercloud Formation; Figs. 3A and 6A).

Numerous red-bed or Kupferschiefer-type copper deposits are known globally (Hitzman et al. 2005), but North American deposits that are temporal equivalents of the Coates Lake Group copper deposits remain unrecognized. On a global scale, the most similar occurrence may be the Zambian Copperbelt (e.g., Hitzman et al. 2005). This copper belt is Neoproterozoic, has a similar rift-like environment of deposition and mineralization, and is overlain by glaciogenic rocks (e.g., Hitzman et al. 2005; Selley et al. 2005). A recent paleo-reconstruction of Rodinia by Evans (2009), an alternative to that of Li et al. (2007), places Africa adjacent to northwestern Laurentia at 780 Ma; according to this model, the Zambian Copperbelt may have formed adjacent to the Coates Lake Group, on the opposite side of a rift. Conversely, recent geochronological constraints on the Zambian Copperbelt, and possible constraints presented in this study, indicate that it may be ca. 50–100 million years older than the Coates Lake Group (Johnson et al. 2007).

In the MSM, stratiform iron deposits consist of hematite-jasper layers (Fig. 5B; e.g., Clout and Simonson 2005; Bekker et al. 2010) within the Neoproterozoic Rapitan Group, a glacial-marine sedimentary succession near the base of the Windermere Supergroup (Fig. 3A; Eisbacher 1978, 1981; Yeo 1981, 1986; Klein and Beukes 1993; Hoffman and Halverson 2011; Hoffman et al. 2011; Baldwin et al. 2012). The Rapitan Group unconformably overlies either the Coates Lake Group, or Little Dal Group and was deposited after 750 Ma, the age of a granitic dropstone in tillite (Ross and Villeneuve 1997) and prior to 685 Ma, deduced by correlation to glacial equivalents in Idaho (Lund et al. 2003; Fanning and Link 2004). Park (1997) used paleomagnetic poles and ages of the Franklin Sills to indicate deposition at low paleolatitudes at ca. 725 Ma for the Rapitan Group. Using regional correlations with the Shaler Supergroup, Young (1992) suggested that the iron formation is younger than 723 Ma (also see Rainbird et al. 1996), which is supported by U–Pb zircon ages of interbedded tuffs and dropstones in correlative strata in Idaho (<717 Ma; >685 Ma; Lund et al. 2003; Fanning and Link 2004) and a ca. 716 Ma interbedded tuff from correlated rocks in Yukon (Macdonald et al. 2010). On a grand scale, these glacial marine rocks are associated with the contested Sturtian snowball Earth episode (e.g., Hoffman and Schrag 2002; Young 2002; Eyles and Januszczak 2004; Har-land 2007; Macdonald et al. 2010), the nature of which has been questioned as the result of recent radiometric agedating (e.g., Lund et al. 2003; Kendall et al. 2006; Azmy et al. 2008).

The Rapitan Group consists of three formations. Diamictite of the Mount Berg Formation is overlain by the Sayunei Formation, a dominantly maroon-weathering siltstone with lesser conglomerate, sandstone, and diamictite. The uppermost Shezal Formation is dominated by green-weathering (with lesser maroon) diamictite (Fig. 3; Eisbacher 1981; Yeo 1981, 1986; Klein and Beukes 1993). Eisbacher (1981), Klein and Beukes (1993), and Hoffman and Halverson (2011) place the iron formation at the top of the Sayunei Formation (Fig. 3A), whereas Yeo (1981, 1986) and Baldwin et al. (2012) place this unit predominantly within the Shezal Formation. It is possible that the thick iron formation in the northern Mackenzie Mountains may not be stratigraphically equivalent to other, thinner iron-formation exposures farther south (Fig. 7) within the Sayunei Formation.

Although some iron formation is locally present along the strike length of the exposed Rapitan Group (Fig. 7; Eisbacher 1981; Yeo 1986; Klein and Beukes 1993; Hoffman and Halverson 2011; Baldwin et al. 2012), the thickest and most exciting exposure, from a mineral deposit perspective, is located in the very northwest part of the Mackenzie Mountains in both Yukon and NWT (Fig. 7), in the Cranswick River–Iron Creek sub-basin. In this area, the iron formation is exposed over a strike length of 40 km, with thicknesses up to 100 m (Yeo 1986; Klein and Beukes 1993), and is referred to as ‘Crest-IF’, ‘Rapitan iron formation’, or the ‘Snake River deposit’. The iron formation is composed of rhythmically bedded hematite containing chert-jasper nodules (Fig. 5B) with lesser interbedded and banded hematite-jasper (i.e., banded iron formation).

Sedimentological and geo-chemical data were used by Yeo (1986) to attribute the source of iron to rift-related hydrothermal plumes. That study suggested that, although glaciers were present, deposition of the iron formation was not related to glaciation (Fig. 8A). This view is supported by Young (2002) who pointed out that the thick glacial deposits of the Shezal Formation overlie the iron formation. Klein and Beukes (1993), on the basis of Eu and other rare-earth elements, trace elements, and stable isotope data inferred deposition of the Rapitan iron formation in water chemically more similar to modern ocean water, than was the case for earlier Proterozoic and Archean iron formations (Klein and Beukes 1993; Klein and Ladeira 2004; Klein 2005). This suggests that although rift-related hydrothermal plumes may have been the source of the iron, the iron was concentrated and then diluted in ocean water due to sub-glacial anoxic conditions during the Sturtian snowball Earth (Fig. 8B). Coupled with sedimentological insights, Klein and Beukes (1993) suggest that the iron was then precipitated during glacial retreat and subsequent oxygenation of the ocean (Fig. 8B). A more complex mechanism of iron formation precipitation is suggested by Baldwin et al. (2012). Using REE+Y and Mo–U geochemical constraints and bedrock observations they suggest that the iron was derived from glacier-related, terrigenous sediment rather than from hydrothermal fluids. In this model dissolved iron created increasingly ferruginous conditions in an ice-capped semi-restricted basin. Iron precipitation occurred during glacial retreat and was related to upwelling of dissolved iron, which interacted with the top layer of progressively oxygenated waters (Baldwin et al. 2012). We have observed dropstones in the iron formation as well as in an underlying glacially-derived diamictite and the overlying Shezal tillite. This likely indicates a more dynamic ice environment of glacial retreat and advance, including meltout from calved icebergs and possibly multiple periods of open water and water column overturning. Overall, it seems unquestionable that the iron formation was related to glaciation, but the source of iron and its mechanism of deposition may require further study.

The Crest iron formation has a drilled historical resource estimate of 5.6 billion tonnes of ~47% Fe (Iron Creek area, Yukon; Fig. 7; Table 1), with a regional resource estimated at >10 billion tonnes (Gross 1970). This iron deposit is probably the largest in western North America and the largest undeveloped deposit in North America (Gross 1970; Klemic 1970). It remains uneconomic to exploit due to the current lack of infrastructure in the region. Possible analogues to the Rapitan iron formation are summarized by Yeo (1986), Young (2002), Klein (2005), and Hoffman et al. (2011).

Over 230 Zn–Pb (± Cu, Ag, Ba, Sb, Ga) prospects and deposits are present throughout the MSM area in the Paleozoic Mackenzie Platform that formed east of the Selwyn Basin (Fig. 12; Martel et al. 2011). The mineralogy within this ‘Mackenzie Mountains Zinc District’ commonly includes sphalerite, smithsonite, and galena, with variable amounts of barite and copper-sulphide/sulphosalts. Fluid-inclusion data indicate temperatures of formation from 160° to >200°C (Fraser 1996; Carrière and Sangster 1999; Gleeson 2006; Paradis 2007; Wallace 2009). Such deposits are referred to here as ‘carbonate-hosted’ rather than MVT due to these elevated temperatures of formation (cf. Hewton 1982; Nelson et al. 2002; Goodfellow 2007; Leach et al. 2005; Paradis et al. 2007; Lund 2008). As Paradis (2007) and Wallace (2009) point out, however, many other aspects of these mineralized zones are consistent with the MVT classification (Leach et al. 2005; Paradis et al. 2007).

The key carbonate-hosted, base-metal deposits in the MSM area are Gayna River, Bear–Twit, Prairie Creek, and potentially Wrigley–Lou (Figs. 5D-E and 12; Heal 1976). Gayna River is in the northern Mackenzie Mountains (Fig. 12), and is characterized by breccia-cemented, gallium-rich sphalerite (Fig. 5D) with lesser pyrite and hematite, hosted in the Neoproterozoic Little Dal Group (Hewton 1982; Dewing et al. 2006; Wallace 2009). An historical resource estimate of 50 million tonnes with over 5% combined Zn–Pb is reported by Hewton (1982; Table 1), although the overall tonnage is probably exaggerated (Wallace 2009). The Bear–Twit prospect, southeast of Gayna River (Fig. 12), contains Zn–Pb–Cu–Ag (sphalerite, galena, tetrahedrite) in veins that cross-cut Paleozoic strata (Dewing et al. 2006). Bear–Twit has an historical resource estimate of 7 million tonnes @ 5.4% Zn, 2.6% Pb, and 17.1 g/t Ag, although the overall tonnage may be over-estimated (Dewing et al. 2006). Prairie Creek, in the southern Mackenzie Mountains (Fig. 12), contains stratabound, sphalerite-galena-bearing ore (Fig. 5E) hosted in the Ordovician Whittaker Formation as well as late, Ag-rich, Pb–Zn veins that crosscut early Paleozoic carbonate rocks and shale (Paradis 2007). Canadian Zinc Corporation is in the process of developing a mine plan for the Prairie Creek Zn–Pb–Ag deposit, and report a measured resource of 5.8 million tonnes @ 20% Zn–Pb and 212 g/t Ag (Table 1).

Comparisons among the deposits and prospects in the MSM area are difficult owing to differences in host-rock age (Fig. 3B), mineralogy, and alteration. There is an array of different mineralization styles, even within individual occurrences (cf. Nelson et al. 2002; Lund 2008). An emerging hypothesis is that some of the carbonate-hosted Zn–Pb occurrences, such as Gayna River, were influenced by northeast-trending faults (present co-ordinates) that were active during deposition of Neoproterozoic strata (Aitken and Cook 1974; Turner and Long 2008. Later fluids may have used these structures to access favourable structural or lithostratigraphic depot-zones (Turner and Long 2008). This relationship may be similar to that of the Pine Point MVT – MacDonald Fault relationship to the east (Hannigan 2007).

In contrast, the Prairie Creek deposit contains a stratabound Zn–Pb component, hosted by the Ordovician–Silurian Whittaker Formation, overprinted by both MVT-like style and later, Ag-rich vein style mineralization (Fraser 1996; Paradis 2007). Prairie Creek, therefore, embodies several of the various carbonate-hosted Zn–Pb types present throughout the MSM area. The MVT-like mineralization at Prairie Creek and the main stage ore minerals at Gayna River do, however, have very similar fluid chemistries (Wallace 2009). The common threads within this diverse group of prospects and deposits are a similarity in Zn–Pb metal budget (± Ba and other base metals), and their location in Neoproterozoic or Paleozoic carbonate rocks. The depositional setting for the Paleozoic carbonate rocks was a passive margin (Mackenzie Platform), with the Selwyn Basin marking the transition to the proto-Pacific Ocean and peri-cra-tonic terranes to the west. The tectonic environment of mineralization, however, is more complex, as discussed below.

The main difficulty in linking these diverse deposits to a tectonic ‘event’ is the lack of age constraints (cf. Nelson et al. 2002; Goodfellow 2007; Lund 2008), and controversy persists over the timing of mineralization at many of the occurrences. Generally, the arguments revolve around Cambrian to Mississippian versus Cretaceous–Tertiary metal emplacement, or possibly both (also see Fraser 1996; Morris and Nesbitt 1998; Nelson et al. 2002; Paradis et al. 2006; Paradis 2007; Wallace 2009). Considering the ages of the sedimentary units that host these Zn–Pb deposits, their mineralizing events must be younger than Cambrian for Gayna River and a number of other occurrences in the northern Mackenzie Mountains, younger than Silurian for Prairie Creek, and younger than Devonian for a number of other prospects. Field relationships indicate that the deposits are structurally constrained, generally as breccias or veins. Cretaceous–Tertiary deformation provides a minimum age for carbonate-hosted Pb–Zn, and some of the other base-metal occurrences may also be related to this event (Godwin et al. 1982; Fraser 1996; Morris and Nesbitt 1998; Dewing et al. 2006; Wallace 2009). However, deformation may differ in timing from one deposit to another and may have taken place in a few pulses over a ~100 m.y. interval (e.g., Morris and Nesbitt 1998; Mair et al. 2006a; Martel et al. 2011). Other workers have suggested a mineralization event during the Devonian to Mississippian (e.g., Nelson et al. 2002), possibly related to SEDEX deposits in the Selwyn Basin (Goodfellow 2007), which Nelson et al. (2002, 2006) suggest took place in a distal back-arc setting. The precise timing and regional relationships are outstanding problems in linking deposit types, a subject also raised by Hannigan (2007) with respect to the Pine Point MVT deposit.

Radiogenic Pb isotope data from galena in a number of carbonate-hosted Zn–Pb occurrences shed some light on the evolution of these deposits when compared with the shale curve for SEDEX deposits (Fig. 11: Godwin et al. 1982; Godwin and Sinclair 1982; Morrow and Cumming 1982; Fraser 1996; Gleeson et al. 2007; Paradis 2007). The Pb isotopic data indicate a range of mineralizing styles. At Prairie Creek, the Pb isotopic data mimics lead from Silurian through Devonian SEDEX deposits (Fig. 11). Paradis (2007) used this to indicate that the stratabound deposits at Prairie Creek formed simultaneously with SEDEX deposits in the Selwyn Basin. The homogeneity is taken to infer either a similar fluid source, or significant mixing of Pb (Morrow and Cumming, 1982). Data in Morrow and Cumming (1982), Fraser (1996), and Paradis (2007) have been used to suggest that the Ag-bearing veins at Prairie Creek are related to Cretaceous deformation (referred to as the ‘Cadillac’ deposit), although the lead isotopic data collectively cross the shale curve suggesting a Triassic age (Fig. 11). In contrast, data from Gayna River (Gleeson et al. 2007) do not match either style of mineralization at Prairie Creek, and may represent the less radiogenic end-member on a mixing line with the more radiogenic, previously defined, “Cordilleran carbonate-lead” trend (Fig. 11A; Nelson et al. 2002; Paradis et al. 2006). Isotopic Pb data on galena from a variety of other prospects fall within the range of Gayna River or Prairie Creek data, or between them (Fig. 11A). The galena Pb isotopic data from the Wrigley–Lou prospects are too radiogenic to plot on Figure 11 (Heal 1976), but fall within the Cordilleran carbonate-lead trend. Notably, the signature from the Pine Point MVT deposit is significantly less radiogenic than that of lead from the carbonate-hosted Zn–Pb prospects in the MSM area, likely indicating distinct and separate mineralizing events, fluid pathways, and/or fluid sources (Fig. 11; Hannigan 2007).

Regional relationships and Pb isotopic data suggest that a minimum of two generations, but more likely three or four, of carbonate-hosted Zn–Pb (+ other base-metal) deposits are present in the MSM area. These include: (1) stratabound ore at Prairie Creek that is similar to, and possibly related to, the Cambrian–Mississippian SEDEX type in the Selwyn Basin to the west (Figs. 10 and 12); (2) vein and breccia-hosted Zn–Pb deposits, such as those at Gayna River and Bear–Twit, which may or may not be related to SEDEX mineralization (Figs. 10 and 12); (3) possible MVT deposits, and; (4) ca. 250 to 60 Ma, vein-hosted, Ag-Zn–Pb deposits, such as those at Prairie Creek (Figs. 10 and 12).

At least 100 mineral occurrences in the study area have been shown to be directly related to intermediate to felsic, ca. 100 to 90 Ma (mid-Cretaceous) plutons and batholiths that are widespread throughout the northern Cordillera (Lang 2000; Selby et al. 2003; Hart et al. 2004a, b; Hart 2007; Rasmussen et al. 2007). In particular, tungsten skarn (e.g., Meinert et al. 2005) has historically been the most significant source of mining-generated revenue in this area (e.g., Dick and Hodgson 1982; Hart et al. 2004b). The plutons were emplaced in the Selwyn Basin area near a lithological transition to shallower-water platformal, sedimentary rocks (Mackenzie Platform; Fig. 13; e.g., Gordey and Anderson 1993); contiguous with a belt of plu-tons that extends from central Alaska through to British Columbia (Hart et al. 2004a, b; Hart 2007). These intrusions are mainly high-K, calc-alkaline to alkaline, biotite ± hornblende ± clinopyroxene ± muscovite-bearing granitoids comprising mainly granitic (senso stricto) to granodioritic compositions, with minor diorite to quartz diorite dykes. These are separated into four geographically, temporally, petro-logically, and metallogenically distinct suites. These subdivisions are the older Tay River suite (99–96 Ma), the Tungsten suite (98–96 Ma), the Mayo suite (97–93 Ma), and the Tombstone suite (93–89 Ma; Lang 2000; Hart et al. 2004a, b; Rasmussen et al. 2007, 2010). Most of the Tombstone suite intrusions, and many of the Mayo suite intrusions, follow an alkalic to subalkalic, monzonitic to monzogranitic compositional trend that contrasts with the dioritic-granodioritic-granitic trend exhibited by the majority of ‘mid’-Cretaceous plutons in the MSM area (Rasmussen et al. 2010). Regional geologic reconstructions (Nelson and Colpron 2007) and geochemical profiles (e.g., Lang 2000; Hart et al. 2004b; Rasmussen et al. 2007) indicate that the magmatism occurred in an area undergoing dextral transpression, with localized extension, hundreds of kilometres inboard of a post-collisional continental margin (Hart et al. 2004b; Rasmussen et al. 2010). The overall trend of the mid-Cretaceous plutonic belts is thought to parallel the North American Neoproterozoic rifted margin, the Paleozoic passive margin, and perhaps the transition from the Mackenzie Platform to the Selwyn Basin (Cecile et al. 1997; Hart et al. 2004b; Hart 2007). The distinctive metal budgets of each plutonic suite, as defined by Hart et al. (2004b) and Rasmussen et al. (2007) are interpreted to reflect their magmatic sources. Magmatic segregation and differentiation processes may have also played a role in the formation of tungsten skarns (Rasmussen et al. 2011).

Intrusions in the MSM area belong to the Tay River suite and the southeast Tombstone–Tungsten belt (TTB) that comprises the Tungsten, Mayo, and Tombstone suites (Fig. 13; Lang 2000; Selby et al. 2003; Hart et al. 2004a, b; Rasmussen et al. 2007, 2010). The predominant forms of intrusion-related deposits associated with these plutonic rocks are divided into two broad categories (Figs. 13 and 14; Hart et al. 2004a, b). The first group consists of skarn-related deposits that are proximal to, and typically genetically related with, mainly the Tungsten suite and highly evolved phases of the Tay River suite. These skarns are mostly tungsten-rich (± minor Cu and/or Zn, Mo, Bi, and/or Au; Dick and Hodgson 1982; Bowman et al. 1985; Selby et al. 2003; Hart et al. 2004a, b; Rasmussen et al. 2007; Yuvan et al. 2007; Rasmussen et al. 2010), although Au-(Ag)-, or Sb-rich skarns are also associated with some Mayo and Tombstone suite intrusions. More distal base metal-rich deposits (e.g., Cu, Zn, Pb, Ag, Au) associated with all of the intrusive suites occur in the form of replacements or mantos and veins (Lang 2000; Hart et al. 2004a; Rasmussen et al. 2007). The second category of deposits associated with mid-Cretaceous plutons consists of pegmatite and quartz vein-hosted gemstones (emerald, Marshall et al. 2004; Tomb-stone suite-hosted, semi-precious tour-maline, Ercit et al. 2003). Pegmatite-hosted, rare-element deposits in the Little Nahanni Pegmatite Group (Li, Cs, Ta, Nb, and Sn; Figs. 13 and 14) are included here, although no direct link with the mid-Cretaceous plutons has been established and the pegmatites may be younger (Groat et al. 2003; Barnes et al. 2007).