Regional Geologic Setting

Much of central Idaho is underlain by cryptic basement domains and a rift belt that overprinted the Laurentian basement (fig. 1A, B). The Paleoproterozoic juvenile crust of the Wallace terrane and western extension of the Great Falls tectonic zone compose the older basement elements (Lund and others, 2015). This basement was first modified by Mesoproterozoic extension, which resulted in the formation of a northwest-striking Lemhi depositional basin (Lund and others, 2004; Lund and Tysdal, 2007). The overprinted Neoproterozoic through early Paleozoic rift belt, which formed the western margin of Laurentia, resulted in north-northwest-striking rift structures, modification of basement by rift-related magmatism, and a broad rift-drift miogeoclinal basin (Lund, 2008; Lund and others, 2010). In south-central Idaho, discrete deformation events in the middle and late Paleozoic, which disturbed miogeoclinal deposition, resulted in successor depositional troughs and overlap successions. Diverse mineral deposit types are associated with the Neoproterozoic–Paleozoic sedimentary depositional environments and facies belts (Link and others, 1995; Lund, 2008; Lund and others, 2016).

The western edge of Laurentia is at the Salmon River suture in west-central Idaho (fig. 1B), along which late Mesozoic oblique convergence between Laurentia and various Pacific terranes caused accretion of allochthonous terranes. This accretion resulted in Late Cretaceous east-northeast-directed compression across central Idaho that telescoped rocks of the several sedimentary basins in this hinterland of the Cordilleran fold and thrust belt (Lund and Snee, 1988; Lund, 2004). Limited early-stage, accretion-related slab-wedge melting resulted in tonalite-granodiorite sheets, whereas associated crustal thickening and melting resulted in younger voluminous S-type granites; together, these formed the Late Cretaceous Atlanta lobe of the Idaho batholith (Lund and Snee, 1988; Lund, 2004; Gaschnig and others, 2010). Changes in plate motion in the Paleocene–Eocene (Lithgow-Bertelloni and Richards, 1998; Liu and others, 2008) resulted first in Paleocene amagmatic crustal relaxation in the Atlanta lobe area (and Idaho batholith magmatism switched to the north, forming the Bitteroot lobe) followed by extensional deformation that culminated in the Eocene Challis magmatism. The study area lies at the eastern edge of the Atlanta lobe and at the western extent of both the Challis volcanic field and Challis plutons.

Development of Central Idaho Mineral-Deposit Models

Because diverse metasedimentary, plutonic, and volcanic rocks host epigenetic deposits in the central Idaho mining districts (figs. 2, 3), there is a broad range of proposed deposit origins and a common linkage to a variety of proposed mineral belts. These mineral belts have been characterized using geographic groupings or, commodities, but the belts, so defined, resulted in intersecting linear arrays of deposits (for example, Bookstrom and others, 1998). Although not ascribed to geologic causes, these include (1) north-striking Florence-Stibnite gold belt, (2) north-northeast-striking Dixie-Thunder Mountain gold belt, (3) northeast-striking Marshall Lake-Elk City polymetallic belt (Green, 1972), (4) northeast-striking Idaho-Montana porphyry belt (Green, 1972; Rostad and others, 1978; Armstrong and others, 1978), (5) north-northwest-striking tungsten belt (Cook, 1956), and (6) northwest-striking Idaho cobalt belt (Hughes, 1983; Hahn and Hughes, 1984). Most of these are substantially modified or abandoned by new studies and so are not shown on maps for the present study.

In interpretations of their genesis, polymetallic deposits in central Idaho were initially interpreted as related to the Late Cretaceous Idaho batholith (Lindgren, 1904; Thompson and Ballard, 1924; Beckwith, 1928; Ross, 1931; Shenon and Reed, 1934). Later, Anderson (1951) argued the Cretaceous plutons of the Idaho batholith were barren intrusions that hosted mineral deposits only by means of magmatic recycling of metals from older sedimentary rocks or as host rocks through introduction of metals by Eocene magmatism. Thereafter, the interpretation of Eocene origins became most widely adopted (Bennett, 1980; Criss and Taylor, 1983; Fisher, 1985; Criss and others, 1991). Isotopic dating, undertaken to address the questions of origin, reveals that several central Idaho precious-polymetallic vein districts are temporally related to Late Cretaceous Idaho batholith two-mica granites (Lund and others, 1986; Gammons, 1988; Lund, 2004), whereas other Idaho vein districts are indeed related to Eocene magmatic and hydrothermal events (Leonard and Marvin, 1982; Fisher, 1985; Hardyman, 1985; Hardyman and Fisher, 1985; Kiilsgaard and Bennett, 1985). Dating of porphyry copper-molybdenum (Cu-Mo) deposits in the Idaho-Montana porphyry belt (Green, 1972; Rostad and others, 1978; Worthington, 2007) demonstrates that both Late Cretaceous and Eocene deposits are interspersed along the belt (Taylor and others, 2007). A corollary is that vast areas of central Idaho are underlain by Late Cretaceous plutons and Eocene igneous rocks that are texturally, compositionally, and temporally equivalent to those that host mineral deposits but that are barren. Thus, the distribution of mineral deposits bears little relation to the location of particular magmatic rocks, and there must be additional factors that were critical to mineral-deposit formation and the deposits’ present location.

Several other geologic features are correlated with mineral deposit occurrences. Deposits of the Idaho-Montana porphyry belt overlie the cryptic Paleoproterozoic juvenile crust of the Great Falls tectonic zone (fig. 1B), and that basement character has been suggested as an influence on subsequent metallogeny (Sanford and Wooden, 1995; Klein and Sims, 2007; Lund and others, 2015). Additionally, several northwest-striking, linear arrays of diverse epigenetic deposits transect the underlying Great Falls tectonic zone and adjacent basement domains (Klein and Sims, 2007). These arrays parallel the zone where older basement was reworked during Neoproterozoic–Paleozoic rifting episodes, and the arrays also parallel the long axis of coeval miogeoclinal (rift-drift) sedimentary deposits, which are present along the length of this western Laurentian rift system (fig. 1A; Lund, 2008). Aspects of the prolonged rifting events, including sedimentary facies belts and synsedimentary down-to-the-basin extensional faults, are both linked with characteristic metal endowments (Turner and Otto, 1995; Lund, 2008; Lund and others, 2015).

Complex questions and conflicting ideas about ages, processes, and origins of central Idaho mineral deposits are illustrated by a modern descriptive mineral-resource assessment, encompassing many central Idaho mining districts, and illustrating the descriptive complexities. That study identifies 17 discrete epigenetic deposit types, many co-located as part of a single deposit and many described in context of the proposed mineral belts that crisscross the province (Bookstrom and others, 1998), as described above. A recent study centered on deposits at Stibnite provides significant details for those deposits and integrates those detail with previous district and regional information (Dail and Zinsser, 2022). Modern mapping and geochronology presented in the present study help clarify framework elements for the mineral deposits in this complex metallogenic province.

Northwestern Mining Districts

Seven mining districts extend about 70 km northeast by 20 km east-west in a northwestern cluster of districts. They lie northwest of a series of normal faults (fig. 2). In the north, epigenetic deposits of the Elk City, Dixie, Buffalo Hump, and Orogrande districts are large quartz veins, wherein quartz has banded, massive, and open space textures. The veins are in persistent, north-northeast-striking fractures within the upper levels of Late Cretaceous (74 Ma, ⁴⁰Ar/³⁹Ar age determination from muscovite) muscovite-biotite granite and in overlying Mesoproterozoic roof pendants (figs. 2, 3; Lund and others, 1986). The precious-metal-rich, polymetallic-bearing assemblages for these districts (Thompson and Ballard, 1924; Beckwith, 1928; Shenon and Reed, 1934; Chauvot, 1986; Lund and others, 1986; Lund and Esparza, 1990) are relatively copper and base-metal rich for the province (tables 1, 2). Similarly, regional geochemical data exhibits anomalous cobalt-copper-gold associations for the area near these districts (Alminas, 1990). Veins of the Buffalo Hump district formed 72–71 Ma (⁴⁰Ar/³⁹Ar age determinations from white mica, Chauvot, 1986; Lund and others, 1986), were genetically related to the muscovite-biotite granite host rock and were emplaced at about 4 km depth based on cooling data (Lund and others, 1986). A single Eocene stock along a several-kilometer long silicified fault zone in the Orogrande district (figs. 2, 3) hosts a relatively gold-rich deposit that notably contains tungsten (table 2; Thompson and Ballard, 1924; Shenon and Reed, 1934).

Quartz-vein deposits in the Marshall Lake-Resort and Warren districts are gold-silver-rich, polymetallic bearing, and some contain scheelite (Reed, 1937; Lorain, 1938; May, 1984; Bookstrom and others, 1998; Mitchell, 2000). Quartz in veins has banded, massive, and open space textures and veins are undeformed. Most veins in the Marshall Lake-Resort district are hosted by fractures in roof-pendant country rocks consisting of Mesoproterozoic, Neoproterozoic, and probably some undated Paleozoic units; other veins in this district are at the top of Late Cretaceous granodiorite and granite. Veins of the Warren district are in biotite granodiorite several hundred meters below the base level of nearby Neoproterozoic or late Mesoproterozoic roof pendants. Granodiorite in the Warren and Marshall Lake-Resort districts is dated at 82.4 Ma (U-Pb age from zircon, this study), and quartz veins in the Marshall Lake-Resort district are dated at 74.5 Ma (⁴⁰Ar/³⁹Ar age determinations from white mica, Lund, 2004). Regional geochemical data show a gold-antimony-tungsten association for these districts (Alminas, 1990; Watts and King, 1999).

Significant gold placers of several ages and settings also characterize northwestern districts (tables 1, 2). Most of the placer districts are in Eocene–Miocene, fault-bounded, extensional basins (Capps, 1939, 1940; Reid, 1960). Some gold placers are in tilted Miocene(?) sediments indicating post-Eocene faulting. Most of the gold placers are in Pleistocene glacial-related deposits and in younger, active stream deposits (Shenon and Reed, 1934; Reed, 1937, 1939; Lorain and Metzger, 1938; Capps, 1939, 1940). Placers in the Warren and Marshall Lake-Resort districts (Marshall Lakes and Resorts districts are combined in this report because the production records were combined in some published accounts) bear significant monazite-bearing black sand deposits as well (Reed, 1937; Savage, 1961).

Southeastern Mining Districts

Six mining districts extend about 40 km northeast by 35 km east-west in a southeastern cluster of mining districts. The districts lie along and southeast of a southern set of large normal faults (fig. 2). At the northern margins of this southeastern district cluster, epigenetic deposits of the Ramey Ridge and Big Creek districts (figs. 2, 3; table 2) are structurally controlled quartz veins (Cater and others, 1973) in metamorphosed Mesoproterozoic Lemhi Group (Lund, 2004) and Neoproterozoic syenite-diorite suites (Lund and others, 2010). Deposits hosted in these Mesoproterozoic metasedimentary rocks are the most copper-rich in the region. These deposits also bear gold and silver as well as local cobalt (table 1). Deposits hosted in the syenite-diorite complex rocks are precious-metal rich and include the single stibnite occurrence in the district (Cater and others, 1973). Ages for these deposits range from 79 to 65 Ma, but it is unclear if these are ages for mineralizing events or for regional cooling (Lund, 2004). A few generally fresh, unaltered Eocene dikes cut the vein deposits, indicating that mineralizing events occurred prior to the Eocene (Shenon and Ross, 1936; Cater and others, 1973).

The Edwardsburg and Profile districts extend north-south for about 30 km along the Johnson Creek-Profile Gap shear zone and are bounded on the east by the Coin Mountain fault (figs. 2, 3; tables 1, 2). Deposits in these districts are hosted by Late Cretaceous granite-granodiorite as well as by Neoproterozoic and Paleozoic metasedimentary rocks. In the Edwardsburg and Profile districts, the metasedimentary rocks are present as screens and elongated blocks along the fault zones and, at the southeastern corner of the Edwardsburg district, they form the northern Stibnite roof pendant (fig. 3; Lund, 2004). The deposits are gold-rich and silver-polymetallic bearing quartz vein, skarn, stockwork, and disseminations within sheared, silicified, and altered host rocks (Shenon and Ross, 1936; Peterson, 1984; Gammons, 1988; Lund, 2004). Quartz in the veins and in silicified wall rock commonly exhibits plastic deformation characteristics (Peterson, 1984; Gammons, 1988; Lund, 2004). These veins lie along the fabrics in the Johnson Creek-Profile Gap shear zone, and some mineralized veins and zones are localized in minor discrete mineralized faults and shears (Peterson, 1984; Gammons, 1988) within the larger shear zone. Most of the early deposits were overprinted by breccia-fill tungsten and (or) antimony deposits (Shenon and Ross, 1936; Peterson, 1984; Gammons, 1988; Bookstrom and others, 1998). For tungsten, huebnerite is a constituent of some precious-polymetallic quartz veins within or adjacent to roof pendants, but scheelite is the constituent mineral in the overprinting breccia-type deposits (Peterson, 1984; Gammons, 1988; Bookstrom and others, 1998). Eocene dikes and elongated stocks intruded the sheared and mineralized rocks, and these intrusions are generally fresh and only locally fractured (Shenon and Ross, 1936; Peterson, 1984; Gammons, 1988; Bookstrom and others, 1998; Lund, 2004). Limited dating indicates that early mineralizing phases are about 79–73 Ma, closely following emplacement of the Late Cretaceous granite bodies. Younger mineralization phases produced ages dating from 69 to 57 Ma, associated with brittle deformation and changes in regional tectonics but not with magmatic activity. Locally, ages as young as 52–47 Ma, coeval with shallow emplacement of Eocene dikes, are also found (⁴⁰Ar/³⁹Ar age determinations from white mica, Gammons, 1988). Fluid-inclusion data document significant decrease in pressures at times of formation from initial mineralization settings with ductile signatures to later breccia-related settings (pressures 1.6–1.3 kilobars [kb] in early phases to 0.15–0.45 kb in brittle phases; temperatures of formation are in the 250–300 °C range; Gammons and others, 1985; Gammons, 1988). Regional geochemical data demonstrate gold-antimony-tungsten associations for these districts (Alminas, 1990; Watts and King, 1999).

The Yellow Pine district, at the southern extent of the southeastern mining district cluster, is composed of three distinct geologic settings containing deposits of different characteristics, described from west to east. Deposits of the western Yellow Pine district are hosted along the Johnson Creek-Profile Gap shear zone. Country rocks are Late Cretaceous biotite granodiorite deformed together with screens of feldspathic metasandstone and minor carbonate and biotite schist of a late Mesoproterozoic to early Neoproterozoic package (fig. 3; Lund, 2004). Early deposits are shear-zone parallel, mesothermal gold-rich silver-polymetallic accumulations in quartz veinlets and disseminated in country rock, wherein quartz exhibits plastic deformation fabrics (Peterson, 1984). Those early deposits were overprinted by a range in brittle deformation from mildly brecciated rocks (broken but not rotated) to intensely brecciated rock (rotated clasts in a matrix of crushed country rocks). Deposits with this structural style host scheelite and stibnite in the breccia matrices as well as in lesser replacement of fragments. The breccia textures and associated mineralization are part of a younger, synextensional mineralization event, indicative of shallow, epithermal mineralizing processes (Cookro, 1985; Cookro and others, 1988; Gammons, 1988; Bookstrom and others, 1998). The regional geochemical signature of the western Yellow Pine district is characterized by a gold-antimony-tungsten association (Alminas, 1990; Watts and King, 1999).

The central Yellow Pine deposits (Stibnite mining area) have been mined and extensively explored most recently, so data for these deposits are more complete. Deposits are located along the throughgoing Meadow Creek fault and along several (mostly) northeast-striking spur faults on both sides of the main fault (fig. 3). Country rocks west of and along the Meadow Creek fault zone are foliated Late Cretaceous granodiorite, whereas predominant rocks east of the fault are Neoproterozoic–Paleozoic metasedimentary rocks of the Stibnite roof pendant. Early deposits are gold-silver, base-metal bearing quartz veins and veinlets that formed between 86 Ma (rhenium-osmium [Re-Os] age determinations from molybdenum, Konyshev and Muntean, 2016) and 78 Ma (⁴⁰Ar/³⁹Ar age determinations from white mica, Gammons, 1988) that were plastically deformed with foliated, silicified, and sericitized host granodiorite primarily along the Meadow Creek fault and its western spurs faults and are also hosted in alternating quartz-rich and carbonate-rich metasedimentary units and structures at the western edge of the Stibnite roof pendant (Gammons, 1988; Konyshev and Muntean, 2016). Disseminated gold as well as disseminated and breccia-hosted stockwork quartz-scheelite deposits overprinted early deposits at about 65–57 Ma (⁴⁰Ar/³⁹Ar age determinations from white mica and K-feldspar, Gammons, 1988; Gillerman and others, 2019), an age range unrelated to a magmatic event. These intermediate-age deposits are found in both granodiorite and metasedimentary rocks in the broad fault-damaged zone at the north end of the Meadow Creek fault (Yellow Pine pit area) and where fault splays intersected metasedimentary rocks in the western part of the Stibnite roof pendant (Cookro and others, 1988; Gammons, 1988; Bookstrom and others, 1998; Konyshev and Muntean, 2016; Gillerman and others, 2019). Gold-rich, epithermal, chalcedonic quartz veins overprinted both older vein types and are dated at 51–47 Ma, and scheelite is dated directly at about 45 Ma (Cookro and others, 1988; Bookstrom and others, 1998; U-Pb age determinations on scheelite, Gillerman and others, 2019); thus, this late event is coeval with Eocene magmatic activity. Both types of antimony deposits were cut by minor Eocene dikes. Fluid-inclusion data indicate pressures of formation decreased significantly between each of the mineralization phases (Gammons, 1988), with estimates of depth starting at about 5 km depth and culminating at near-surface depths (Marsh, E., and Bennett, M., 2021, U.S. Geological Survey oral commun.). Gold-antimony-tungsten geochemical associations are characteristic for this part of the district (Alminas, 1990; Watts and King, 1999).

The eastern Yellow Pine district contains gold and mercury deposits in an about 3-km wide zone hosted by Neoproterozoic to Paleozoic marble units of the Stibnite roof pendant, which overlies Cretaceous granitic rocks. Most ore zones are parallel to the strike of the marble lithologies, near the contacts between marble and quartzite, and some are in northeast-striking cross fractures (Larsen and Livingston, 1921; Schrader and Ross, 1926; Cookro and others, 1988; Bookstrom and others, 1998). Ores are in jasperoid zones and as disseminations that formed at near-surface conditions and interpreted as Eocene hot spring and chalcedonic fumarole deposits during the youngest hydrothermal activities (Larsen and Livingston, 1921; Cookro and others, 1988; Konyshev and Muntean, 2016; Gillerman and others, 2019).

Thunder Mountain deposits (near Sunnyside on figs. 2, 3) are in late-stage Eocene volcanic units preserved in a caldera collapse graben (Leonard and Marvin, 1982; Hardyman and Fisher, 1985; southern Big Creek graben, Stewart and others, 2013). These deposits from 43 Ma are hot-spring-related, disseminated gold and silver ore deposits (potassium-argon [K-Ar] dating of late-stage adularia, Adams, 1985). They are hosted in non-welded rhyolite and ash-flow tuff and in fault-juxtaposed volcaniclastic deposits (Umpleby and Livingston, 1920; Ross, 1933; Shannon and Reynolds, 1975; Adams, 1985; Parsley, 1997). Mineralizing fluids moved laterally along graben-related faults into permeable and carbon-rich layers (Shannon and Reynolds, 1975; Adams, 1985; Parsley, 1997; Bookstrom and others, 1998). The setting is similar to gold deposits of the Eocene Custer graben (30 km southeast, McIntyre and Johnson, 1985).

Mesoproterozoic Facies

The roof pendants in the many northern districts of the northwestern cluster plus the Ramey Ridge and Big Creek districts of the southern cluster are composed of fine-grained arkosic siltite and metasandstone, or gneissic equivalents of these rock types. Map units are thousands of meters thick. Based on the stratigraphic character of very thick, relatively fine-grained biotite-feldspar-quartz gneisses composing most country rocks in the northern part of the northwestern district cluster, these rocks are along-strike, western equivalents to upper parts of the Lemhi Group (fig. 3). Preserved sedimentary features in lower-metamorphic grade strata of the Big Creek district indicate large exposures of Mesoproterozoic Apple Creek and Gunsight Formations from the middle and upper parts of the Lemhi Group (Lund, 2004). A central band of rocks in the northern cluster and in the northern tip of the southeastern cluster are associated with thick calc-silicate gneiss and marble, biotite gneiss, and quartzite units that are western extensions of the Yellowjacket-Hoodoo association (fig. 3). Thus, although exposures are discontinuous because of intervening intrusions and off-setting structures, recent geologic mapping, stratigraphic studies (Lund, 2004; Lund and Tysdal, 2007; this study), and limited detrital zircon populations (Lund and others, 2008; Lewis and others, 2010, compare to Link and others, 2007; and xenocrystic zircon ages, Gaschnig and others, 2013) outline northwest-striking belts of Mesoproterozoic Lemhi Group and Yellowjacket-Hoodoo strata across central Idaho (fig. 4).

The presence of parallel Lemhi Group and Yellowjacket-Hoodoo stratigraphic belts is strengthened by the combination of regional cobalt-copper geochemical associations (Alminas, 1990) and gold-copper-polymetallic deposits along the belt and that this facies belt includes the Blackbird gold-cobalt-copper and related deposits of east-central Idaho (Klein and Sims, 2007; Lund and Tysdal, 2007; although not all deposits bear cobalt, Johnson and others, 1998). Based on Pb-isotopic data, the source of metals for the Blackbird cobalt-copper deposits was the Apple Creek Formation (Panneerselvam and others, 2012). However, as presently constituted, deposits in this facies belt are epigenetic, associated with Late Cretaceous thrust stacking of these facies, heat from crustal thickening, and from the circulation of resultant regional-metamorphic and magmatic fluids (Lund and others, 2011).

Neoproterozoic and Paleozoic Facies

In the Yellow Pine district of the southeastern cluster (fig. 2), some metasedimentary rocks were originally interpreted to be Paleozoic based on their lithologic and thickness characteristics (Larsen and Livingston, 1921; Schrader and Ross, 1926; Shenon and Ross, 1936; White, 1945; Smitherman, 1988) as discussed in one regional interpretation (Ross, 1934), but these same rocks are shown as Mesoproterozoic Yellowjacket-Hoodoo Formations on both local (Leonard, 1962) and regional maps (Fisher and others, 1992). In the western part of the northwestern cluster of districts, metamorphosed strata, which have similar characteristics to many in the southern cluster, were also correlated to Mesoproterozoic Yellowjacket-Hoodoo Formations (Knowles and Bennett, 1978) but were reinterpreted as Neoproterozoic and early Paleozoic (Lund and Esparza, 1990). The first two state geologic maps showed central Idaho metamorphosed strata first as Mesoproterozoic Belt Supergroup formations (Ross and Forrester, 1947) and later as Mesoproterozoic Yellowjacket-Hoodoo Formations (Bond, 1978), further complicating interpretations. The recent state map (Lewis and others, 2012) corrects correlations where interpretations had been updated.

Despite the divergent interpretations for the metamorphic rocks in these particular locations, many exposures of these strata have characteristics that allow them to be distinguished from known Mesoproterozoic rocks. These strata are in discrete packages characterized by abruptly contrasting lithodemic units that have thicknesses on scales of hundreds of meters. Among exposures containing rocks with these characteristics, packages of both Neoproterozoic and Paleozoic ages are documented. The Neoproterozoic section contains some distinctive units that have been dated, but otherwise, many similar lithologies are present in both packages. The map units of similar characteristics but possibly different ages are not uniquely distinguishable in the many small, fragmented, or complexly deformed exposures; thus, in important mining districts of the province, determining age and origin of metamorphosed strata is not straightforward.

A Neoproterozoic stratigraphic package is mapped in the Gospel Peaks roof pendant between the Buffalo Hump and Florence mining districts (figs. 2, 3; Lund and Esparza, 1990). This section is dated in the Big Creek roof pendant of the Edwardsburg district. A partial section is present in the Marshall Lake-Resort district and rocks of this age may also be present in the Yellow Pine, Profile, and Warren districts (Lund, 2004). The Neoproterozoic stratigraphic section is a documented succession of distinctive metavolcanic and metadiamictite units as well as quartzite and quartz-pebble conglomerate, marble and calc-silicate gneiss, and schist units. The section, with the metavolcanic and diamictite rocks in the middle part, is estimated to be as much as 3,300 meters (m) thick (Lund, 1984; Lund and Esparza, 1990; Lund, 2004). Neoproterozoic ages are confirmed by direct dates of about 685 Ma on metavolcanic rocks in the Edwardsburg district (U-Pb ages determined from zircon, Lund and others, 2003). The characteristic stratigraphic order, lithodemic compositions, unit thicknesses, and ages provide (1) direct correlation between the rock formations of central Idaho and those of the Windermere Supergroup of northern Idaho-northwestern Washington and southeastern Idaho; the correlations support related basin development in these areas (Lund and Cheney, 2016), and (2) outline the trace of an early phase of Rodinian continental rifting. Neoproterozoic syenite-diorite suites in the Ramey Ridge and Big Creek districts intruded Mesoproterozoic metasedimentary strata (figs. 3, 4; Lund, 2004), and based on ages of about 650 Ma and geochemical characteristics, they are interpreted as a younger pulse of Rodinian rifting along the same margin as the older Neoproterozoic volcanic rocks (Lund and others, 2010).

Another package with abruptly contrasting lithodemic units, including quartzite, marble, quartzite-pebble gneiss, and biotite schist, is locally associated with the dated Neoproterozoic strata. These are interpreted as probable Paleozoic metasedimentary strata based on stratigraphic characteristics (Shenon and Ross, 1936; White, 1945; Lund, 1984, 2004; Smitherman, 1988; Lund and Esparza, 1990; shown as Mesoproterozoic in Fisher and others, 1992). They are best identified where locally identified above the dated volcanic units but, elsewhere, this package remains difficult to distinguish from the similar lithologic types of the Neoproterozoic package. Probable Paleozoic units are mapped in the Profile district and at the eastern edge of the Florence district; these units are probably present as minor components of roof rocks in the Edwardsburg and Marshall Lake-Resort districts (Lund, 2004; figs. 2, 3). Several small exposures of probable Paleozoic rocks are also located at the southeastern corner of figure 3 (for example near sample number 06KL058). The largest of these exposures is preserved in the Stibnite roof pendant (Shenon and Ross, 1936; White, 1945; Smitherman, 1988; Lund, 2004; Stewart and others, 2016).

In the Stibnite roof pendant, the possible Paleozoic metasedimentary package is estimated at 750–1,000 m thick (Lund, 2004; Stewart and others, 2016). Post-Neoproterozoic ages for some of the lithodemic units are confirmed by detrital zircon analyses (Stewart and others, 2016; Isakson, 2017). However, because there are no directly datable units and compressional structures are complex, map units have been identified in different ways in different map versions. Mapping has been attempted based on lithologies (Shenon and Ross, 1936; White, 1945), on local mine geology lithologic units (Smitherman, 1988; Stewart and others, 2016), and on regionally based characteristics of stratigraphic packages (Lund, 2004). Each of these different approaches results in different structural interpretations. Furthermore, because correlations for the package are dependent on knowing the stratigraphic order, the age range and depositional setting(s) of the package as a whole are difficult to assess.

The amphibolite facies metamorphic grade and the absence of fossils or direct ages render many traditional correlation methods unavailable for the undated Neoproterozoic and Paleozoic exposures. Thus, other criteria need to be considered to constrain interpretations of original setting and sedimentary facies of the incomplete and scattered central Idaho exposures in comparison to better documented stratigraphic sections of south- and east-central Idaho (fig. 1B). Such alternative approaches are described following and include comparisons to (1) distinctive characteristics of stratigraphic packages, (2) detrital zircon populations as possible constraints on ages, provenances, or depositional settings, and (3) metal endowment assemblages based on regional geochemistry and mineral deposits.

Comparison to Known-Age Sedimentary Facies of South-Central Idaho

Neoproterozoic Packages

Neoproterozoic Windermere Supergroup volcanic and diamictite strata are identified discontinuously across a broad area of central Idaho and are constrained to be about 685 Ma by dating in the Edwardsburg district (SHRIMP U-Pb ages on zircon, Lund and others, 2003). A younger volcanogenic conglomerate unit in the southern Salmon River Mountains is dated at about 654 Ma from drillcore but is not exposed at the surface (SHRIMP U-Pb ages on zircon, Lund and others, 2010). Units in this stratigraphic package are correlated regionally to sections in southeastern and northern Idaho and along the western Laurentian margin (Lund and others, 2003, 2010; Lund and Cheney, 2016). Despite the Neoproterozoic package (Windermere Supergroup) being widely distributed along the ancient margin, no strata of this age are known to be exposed in south-central or east-central Idaho.

In the southern Salmon River Mountains (fig. 1B; Bayhorse area southwest of Challis, fig. 4), strata in the middle levels of a folded thrust stack (Bayhorse antiform, Hobbs and others, 1991) have been recently interpreted to be Neoproterozoic based on a date of about 600 Ma from one of a number of small mafic bodies (Brennan and others, 2020). However, one of the units associated with the mafic rocks gradationally overlies a unit containing Late Cambrian to Early Ordovician fossils (Hobbs and Hays, 1990; Hobbs and others, 1991; although this is not acknowledged by Brennan and others, 2020). Examination of the mafic exposures reveals they are in a thin zone along a contact interpreted as a thrust fault for about 22 km (Hobbs and others, 1991), suggesting that the mafic exposures klippen from a Neoproterozoic remnant at the base of a thrust plate. The fossil evidence is compelling for the age of the strata in the core of the Bayhorse antiform, so herein the strata in the core of the Bayhorse antiform are interpreted as Cambrian–Ordovician following previous studies (Hobbs and Hays, 1990; Hobbs and others, 1991; Lund and others, 2010) and are thus discussed with other lower Paleozoic rocks. The only Neoproterozoic strata exposed in east-central Idaho and constrained by fossils and detrital zircon dating are in small exposures of upper Ediacaran and Lower Cambrian quartzite deposited on Mesoproterozoic rocks (McCandless, 1982; Carr and Link, 1999; Pearson and Link, 2021).

In the absence of known Windermere Supergroup strata exposed in south-central Idaho or in the intervening southern Salmon River Mountains, evaluation of any association between mineral deposits and the Neoproterozoic section is based on the dated stratigraphic section in central Idaho.

Paleozoic Packages

Correlating the probable Paleozoic metasedimentary rocks of central Idaho poses significant problems, given that stratigraphic successions and rock ages are largely unconstrained. The approach attempted in this study is to evaluate the probable Paleozoic metamorphosed strata of central Idaho as stratigraphic package(s) that help define potential original depositional settings. These package(s) are compared to four main Paleozoic facies belts that cross south- and east-central Idaho striking north-northwest and that were deposited in separate basinal settings. The four facies belts are numbered and described in the following paragraphs.

(1) An Ordovician to Pennsylvanian shelf facies composes the Paleozoic section in the eastern three tilt-block ranges of east-central Idaho (fig. 4). This greater than 1,500-m thick, primarily carbonate section (Ruppel and Lopez, 1988) was deposited in deepening water westward (Isaacson and others, 1983; Grader and Dehler, 1999). This facies is primarily associated with silver-lead-zinc vein and skarn deposits (fig. 4). Such a thick carbonate package with an accompanying metals assemblage is not present in central Idaho.

(2) Cambrian to Devonian shelf-break to slope strata are present in south-central Idaho but only locally exposed and found mostly southeast of Challis (these are the western extent of the facies on fig. 4). The section is composed of Cambrian to Ordovician siliciclastic rocks, Ordovician to Devonian siltstone, and Ordovician to Devonian carbonate rocks; of these, the Ordovician Kinnikinic Quartzite can be recognized from across the shelf and into this deeper water section (Isaacson and others, 1983; Hobbs and others, 1991; Grader and Dehler, 1999; Baar, 2009). The section is not associated with mineral deposits.

(3a) The Mississippian deep-sea fan turbidite (Antler flysch) facies of south-central Idaho (fig. 4) is composed of carbonaceous argillite and siltite and carbonate rocks and unconformably overlies Cambrian to Devonian shelf-break strata (facies number 2, above). (3b) Ordovician to Devonian deep-water facies rocks lie west of (and formed the highlands for) the Mississippian flysch basin and are of similar composition. Together these two different facies compose most of the Idaho “black shale belt.” They are structurally juxtaposed through their central extent and, combined, are more than 4,000-m thick (Link and others, 1995). The western strata of this “black shale belt” are associated with silver-lead deposits and the eastern are associated with gold deposits (Hall and others, 1978; Hall, 1985; Soulliere and others, 1995). If metamorphosed, these facies belts would amount to a comparably thick stack of carbonaceous schist, which is not present in central Idaho.

(4) The facies exposed farthest west in south-central Idaho is 3,000 m of Pennsylvanian to Permian strata complexly alternating carbonaceous siltstone, limestone, sandy limestone, sandstone, and conglomerate that were deposited in an epicratonic onlap basin (fig. 4; Mahoney and others, 1991; Link and others, 1995). Tungsten, silver-lead-zinc-tin, and antimony deposits are associated with these units; some mineral deposits are associated with depositional contacts, but most mineral deposits occur where these units were thrust faulted over early Paleozoic deep-water facies and where either Late Cretaceous or Eocene intrusions invaded the structures (Hall and others, 1978; Cookro, 1985; Hall, 1985). The compositional variations and the scale of compositional alternation of this facies are similar to, but generally thicker than, probable Paleozoic lithodemic units in the central Idaho mineral province.

The farthest north strata with fossil control for ages of some of the rocks are part of a folded thrust fault complex in the southern Salmon River Mountains (figs. 1B, 4). On the west is the primarily siliciclastic-facies of the Pennsylvanian–Permian onlap facies (Mahoney and others, 1991); these rocks overlie a thrust plate of Mississippian flysch facies that includes basal thrust slivers as old as Cambrian (Hall, 1985; Hobbs and others, 1991; Link and others, 1995). These two facies packages are along strike with similar facies in south-central Idaho. Deeper in the thrust stack are incompletely dated late Neoproterozoic–Cambrian(?), Cambrian(?), and dated Ordovician successions of quartzite, quartz-pebble conglomerate, siltstone, and carbonate, which seem to represent a broad, somewhat diverse succession that was telescoped. These are faulted over Cambrian(?) and Ordovician carbonate, slate, phyllite units. This complex structural stack was folded into two large antiforms (Hobbs and Hays, 1990; Hobbs and others, 1991). The Cambrian and Ordovician rocks in this area are a mostly older, thicker, and coarser-grained siliciclastic facies than the exposed Ordovician and younger deep-water carbonaceous-predominant facies in south-central Idaho (although both are included with the Paleozoic shelf-slope strata on fig. 4). This change in stratigraphy from south-central Idaho to the southern Salmon River Mountains reflects a transition in sedimentary successions across a north-northeast striking, syn-depositional paleostructure (fig. 4) called the “Salmon River lineament” by Hobbs and Hays (1990) and Hobbs and others (1991).

Provenance Characteristics

Another approach for evaluating the original facies of the metasedimentary rocks of central Idaho is to evaluate provenance information. Detrital zircon analyses can establish potential limits on ages but they can also provide possible matching of relative probabilities for detrital zircon populations among metasedimentary packages and some data about sources of detritus and stability of basins. Data from the mapped and dated Neoproterozoic volcanic rocks in the Edwardsburg district (Lund and others, 2003) and some map relations in the Yellow Pine district provide relative age information for some units. To provide detrital zircon population data for rocks with the best documented stratigraphic relations, five samples are analyzed in the present study (table 3, fig. 5A). Data for these samples are in Aleinikoff and others (2023). These data can be compared to existing detrital zircon data for other rocks of undetermined ages in this area (fig. 5B, C). Similarly, data from known-age Paleozoic facies belts in south-central Idaho can be compared to undated packages in central Idaho (fig. 5D–I).

Table 3.    

Uranium-thorium-lead (U-Th-Pb) data for detrital zircon from metasedimentary rock, central Idaho mineral province.

[Data are available in Aleinikoff and others (2023). Datum for latitude-longitude is World Geodetic System 1984. Propagated uncertainties include both reference material and instrumental uncertainties. Preferred ages determined according to ²⁰⁶Pb/²³⁸U age is used if grain is less 1,300 Ma and ²⁰⁷Pb/²⁰⁶Pb age is used if grain is more than 1,300 Ma. Discordant analyses (greater than about 10 percent discordant) are excluded from relative probability plots. ID, identification; Ma, million years; sig, sigma; abs, absolute; PDP, probability density plots]

Table 3.    Uranium-thorium-lead (U-Th-Pb) data for detrital zircon from metasedimentary rock, central Idaho mineral province.

Analysis ID	Preferred Age (Ma)	2-sig (abs)	Included in relative probability plot
04KL102A (Placer lake ridge, siltite) [latitude 45.161, longitude −115.35667]
04KL102A-001	1697	74		Yes
04KL102A-002	1855	61		No
04KL102A-003	1414	74		Yes
04KL102A-004	1175	33		Yes
04KL102A-005	1368	90		Yes
04KL102A-006	1246	32		Yes
04KL102A-007	1849	64		Yes
04KL102A-008	1160	100		No
04KL102A-009	1585	66		Yes
04KL102A-010	1418	80		Yes
04KL102A-011	1014	33		No
04KL102A-012	1245	31		Yes
04KL102A-013	1852	64		No
04KL102A-014	2320	61		Yes
04KL102A-015	1474	65		Yes
04KL102A-016	1024	30		No
04KL102A-017	1692	68		Yes
04KL102A-018	1389	73		Yes
04KL102A-019	1475	65		Yes
04KL102A-020	1734	60		Yes
04KL102A-021	1382	71		Yes
04KL102A-022	1157	29		Yes
04KL102A-023	2305	58		Yes
04KL102A-024	1264	30		Yes
04KL102A-025	1282	29		Yes
04KL102A-026	1420	81		Yes
04KL102A-027	1259	67		Yes
04KL102A-028	1164	31		Yes
04KL102A-029	1206	32		Yes
04KL102A-030	1287	31		Yes
04KL102A-031	1785	66		Yes
04KL102A-032	1195	31		Yes
04KL102A-033	1041	42		Yes
04KL102A-034	1414	70		Yes
04KL102A-035	1259	58		No
04KL102A-036	1035	89		No
04KL102A-037	1754	66		Yes
04KL102A-038	1241	38		Yes
04KL102A-039	1336	73		Yes
04KL102A-040	871	40		No
04KL102A-041	1742	72		Yes
04KL102A-042	2506	64		Yes
04KL102A-043	1906	62		Yes
04KL102A-044	1455	73		Yes
04KL102A-045	1744	64		Yes
04KL102A-046	1667	65		Yes
04KL102A-047	1460	71		Yes
04KL102A-048	1193	66		No
04KL102A-049	1195	31		No
04KL102A-050	972	42		No
04KL102A-051	1731	71		Yes
04KL102A-052	1773	61		Yes
04KL102A-053	1457	66		Yes
04KL102A-054	520	51		No
04KL102A-055	2685	56		Yes
04KL102A-056	1187	37		Yes
04KL102A-057	730	29		No
04KL102A-058	1670	78		Yes
04KL102A-059	1740	64		Yes
04KL102A-060	1764	64		Yes
04KL102A-061	1066	19		No
04KL102A-062	1393	33		Yes
04KL102A-063	1790	32		Yes
04KL102A-064	1262	34		Yes
04KL102A-065	1260	23		Yes
04KL102A-066	1161	16		Yes
04KL102A-067	1231	17		Yes
04KL102A-068	1295	24		Yes
04KL102A-069	2633	29		Yes
04KL102A-070	1171	22		Yes
04KL102A-071	2331	27		Yes
04KL102A-072	1478	32		Yes
04KL102A-073	1297	41		Yes
04KL102A-074	2574	30		Yes
04KL102A-075	1342	40		No
04KL102A-076	1378	52		Yes
04KL102A-077	1326	47		Yes
04KL102A-078	1678	44		Yes
04KL102A-079	1708	38		Yes
04KL102A-080	1412	49		Yes
04KL102A-081	1244	27		Yes
04KL102A-082	1824	42		Yes
04KL102A-083	1745	37		Yes
04KL102A-084	1695	33		Yes
04KL102A-085	1872	33		Yes
04KL102A-086	1657	39		Yes
04KL102A-087	1732	38		Yes
04KL102A-088	1850	130		Yes
04KL102A-089	1695	43		Yes
04KL102A-090	2991	26		Yes
04KL102A-091	1398	44		Yes
04KL102A-092	1245	31		Yes
04KL102A-093	1870	46		Yes
04KL102A-094	2617	30		No
04KL102A-095	1393	34		Yes
04KL102A-096	985	97		No
04KL102A-097	1260	250		No
04KL102A-098	1687	32		Yes
04KL102A-099	1759	39		Yes
04KL102A-100	1280	26		Yes
04KL102A-101	1628	28		Yes
04KL102A-102	87	3		No
04KL102A-103	1617	44		Yes
04KL102A-104	1672	38		Yes
04KL102A-105	2700	35		Yes
04KL102A-106	1703	26		Yes
04KL102A-107	2640	33		Yes
04KL102A-108	1250	19		Yes
04KL102A-109	1671	42		Yes
04KL102A-110	1798	35		Yes
04KL102A-111	1321	35		No
04KL102A-112	1671	59		Yes
04KL102A-113	1839	39		Yes
04KL102A-114	1640	29		Yes
04KL102A-115	1753	52		Yes
04KL102A-116	1355	44		Yes
04KL102A-117	1850	100		Yes
04KL102A-118	1447	71		Yes
04KL102A-119	1827	36		Yes
04KL102A-120	1924	29		Yes
04KL107 (McFadden Point, metaquartzite) [latitude 45.17355, longitude −115.351117]
04KL107_001	1451	41		Yes
04KL107_002	1692	39		Yes
04KL107_003	1809	30		Yes
04KL107_004	1963	31		Yes
04KL107_005	2441	29		Yes
04KL107_006	1805	32		Yes
04KL107_007	1683	52		Yes
04KL107_008	1622	36		Yes
04KL107_009	1650	38		Yes
04KL107_010	1672	31		Yes
04KL107_011	1728	36		Yes
04KL107_012	2463	30		Yes
04KL107_013	2337	31		Yes
04KL107_014	1455	58		Yes
04KL107_015	1733	30		Yes
04KL107_016	2657	38		Yes
04KL107_017	1444	37		Yes
04KL107_018	1723	30		Yes
04KL107_019	1759	88		Yes
04KL107_020	1715	36		Yes
04KL107_021	1455	46		Yes
04KL107_022	1687	34		Yes
04KL107_023	1744	42		Yes
04KL107_024	1699	32		Yes
04KL107_025	2736	41		Yes
04KL107_026	1823	68		Yes
04KL107_027	1177	51		No
04KL107_028	1930	28		Yes
04KL107_029	1670	34		Yes
04KL107_030	1784	33		Yes
04KL107_031	1747	32		Yes
04KL107_032	1723	29		Yes
04KL107_033	1157	44		Yes
04KL107_034	2697	32		Yes
04KL107_035	1435	46		Yes
04KL107_036	2497	32		Yes
04KL107_037	1670	36		Yes
04KL107_038	1409	59		Yes
04KL107_039	1660	33		Yes
04KL107_040	1871	49		No
04KL107_041	1711	41		Yes
04KL107_042	1747	41		Yes
04KL107_043	910	23		No
04KL107_044	1747	46		Yes
04KL107_045	1469	40		Yes
04KL107_046	1627	44		Yes
04KL107_047	1687	44		Yes
04KL107_048	1720	40		Yes
04KL107_049	1756	43		Yes
04KL107_050	1721	42		Yes
04KL107_051	1419	48		Yes
04KL107_052	1860	46		Yes
04KL107_053	1712	39		Yes
04KL107_054	2611	35		Yes
04KL107_055	2312	40		Yes
04KL107_056	1445	61		Yes
04KL107_057	2187	61		Yes
04KL107_058	1443	42		Yes
04KL107_059	1762	39		Yes
04KL107_060	1690	42		Yes
04KL107_061	1804	40		Yes
04KL107_062	2362	46		Yes
04KL107_063	1712	42		Yes
04KL107_064	2494	35		Yes
04KL107_065	1670	39		Yes
04KL107_066	1685	53		Yes
04KL107_067	1707	40		Yes
04KL107_068	1701	46		Yes
04KL107_069	1720	38		Yes
04KL107_070	1412	44		Yes
04KL107_071	1721	42		Yes
04KL107_072	1718	46		Yes
04KL107_073	1469	65		Yes
04KL107_074	1698	40		Yes
04KL107_075	1719	43		Yes
04KL107_076	1767	44		Yes
04KL107_077	1771	49		No
04KL107_078	2335	40		No
04KL107_079	1767	41		Yes
04KL107_080	854	22		No
04KL107_081	1749	30		Yes
04KL107_082	1469	27		Yes
04KL107_083	1675	34		Yes
04KL107_084	1799	22		Yes
04KL107_085	1480	19		Yes
04KL107_086	2519	22		Yes
04KL107_087	1697	35		Yes
04KL107_088	1682	26		Yes
04KL107_089	1436	32		Yes
04KL107_090	1694	26		Yes
04KL107_091	1449	21		Yes
04KL107_092	1717	74		Yes
04KL107_093	1717	26		Yes
04KL107_094	1463	45		Yes
04KL107_095	1742	15		Yes
04KL107_096	1693	21		Yes
04KL107_097	1694	41		Yes
04KL107_098	1436	37		Yes
04KL107_099	1673	26		Yes
04KL107_100	1468	37		Yes
04KL107_101	1703	32		Yes
04KL107_102	1730	22		Yes
04KL107_103	1711	44		Yes
04KL107_104	2476	20		Yes
04KL107_105	1412	29		Yes
04KL107_106	3247	23		Yes
04KL107_107	1690	24		Yes
04KL107_108	1739	34		Yes
04KL107_109	1683	24		Yes
04KL107_110	1771	25		Yes
04KL107_111	1457	32		Yes
04KL107_112	1425	25		Yes
04KL107_113	1744	22		Yes
04KL107_114	1755	15		Yes
04KL107_115	1446	22		Yes
04KL107_116	1766	18		Yes
04KL107_117	1690	30		Yes
04KL107_118	1767	20		Yes
04KL107_119	1686	24		Yes
08KL005 (Jackass Gulch, metamorphosed quartz-pebble conglomerate) [latitude 44.903306, longitude −115.269472]
08KL005-001	1177	29		Yes
08KL005-002	1825	32		Yes
08KL005-003	2609	30		No
08KL005-004	2687	32		Yes
08KL005-005	1391	46		Yes
08KL005-006	1146	27		Yes
08KL005-007	1017	23		Yes
08KL005-008	1075	31		Yes
08KL005-009	832	38		No
08KL005-010	1205	26		No
08KL005-011	1117	28		Yes
08KL005-012	1376	35		Yes
08KL005-013	754	17		No
08KL005-014	432	12		No
08KL005-015	1127	62		Yes
08KL005-016	2546	28		Yes
08KL005-017	535	30		No
08KL005-018	950	25		Yes
08KL005-019	783	43		No
08KL005-020	1218	29		Yes
08KL005-021	997	28		Yes
08KL005-022	670	110		No
08KL005-023	1176	31		No
08KL005-024	2885	55		No
08KL005-025	1205	33		Yes
08KL005-026	883	34		No
08KL005-027	981	32		Yes
08KL005-028	2692	26		No
08KL005-029	1122	28		Yes
08KL005-030	587	16		No
08KL005-031	1222	33		Yes
08KL005-032	1336	45		Yes
08KL005-033	1692	47		Yes
08KL005-034	2658	26		No
08KL005-035	1159	40		No
08KL005-036	960	22		Yes
08KL005-037	1437	57		Yes
08KL005-038	1015	29		No
08KL005-039	1270	36		No
08KL005-040	1437	48		Yes
08KL005-041	504	39		No
08KL005-042	745	29		No
08KL005-043	880	110		Yes
08KL005-044	517	22		No
08KL005-045	1169	36		Yes
08KL005-046	1819	36		Yes
08KL005-047	1000	22		Yes
08KL005-048	762	30		No
08KL005-049	1771	45		No
08KL005-050	1016	31		No
08KL005-051	1771	35		Yes
08KL005-052	1348	46		Yes
08KL005-053	736	98		No
08KL005-054	963	47		No
08KL005-055	1833	34		Yes
08KL005-056	1136	29		Yes
08KL005-057	1032	28		Yes
08KL005-058	556	18		No
08KL005-059	1057	29		Yes
08KL005-060	563	43		No
08KL005-061	1788	53		Yes
08KL005-062	551	41		No
08KL005-063	1110	20		Yes
08KL005-064	2637	42		Yes
08KL005-065	919	29		No
08KL005-066	2511	43		Yes
08KL005-067	1164	27		Yes
08KL005-068	1687	49		Yes
08KL005-069	544	13		No
08KL005-070	1766	46		Yes
08KL005-071	1197	26		Yes
08KL005-072	1773	58		No
08KL005-073	1067	21		Yes
08KL005-074	583	15		No
08KL005-075	1171	27		Yes
08KL005-076	1110	21		Yes
08KL005-077	2580	45		No
08KL005-078	2658	46		Yes
08KL005-079	1124	20		Yes
08KL005-080	955	44		No
08KL005-081	778	41		No
08KL005-082	918	15		No
08KL005-083	1171	22		Yes
08KL005-084	949	32		Yes
08KL005-085	1038	24		Yes
08KL005-086	2707	42		Yes
08KL005-087	656	18		No
08KL005-088	719	25		No
08KL005-089	1688	47		Yes
08KL005-090	876	22		Yes
08KL005-091	460	100		No
08KL005-092	1679	49		Yes
08KL005-093	2545	41		Yes
08KL005-094	1756	48		Yes
08KL005-095	1022	23		Yes
08KL005-096	1177	25		Yes
08KL005-097	1012	37		No
08KL005-098	1699	52		Yes
08KL005-099	1766	44		Yes
08KL005-100	1094	23		Yes
08KL005-101	1044	28		Yes
08KL005-102	1218	24		Yes
08KL005-103	1331	75		Yes
08KL005-104	1165	30		Yes
08KL005-105	1067	29		Yes
08KL005-106	2852	40		No
08KL005-107	1697	55		Yes
08KL005-108	1753	46		Yes
08KL005-109	1001	24		Yes
08KL005-110	1724	60		Yes
08KL005-111	1796	47		Yes
08KL005-112	1220	130		No
08KL005-113	846	21		No
08KL005-114	1069	21		Yes
08KL005-115	1363	53		Yes
08KL005-116	1850	48		Yes
08KL005-117	1097	23		Yes
08KL005-118	1797	43		Yes
08KL005-119	480	150		No
08KL005-120	2584	48		Yes
20KL047 (Whiskey Creek, metasandstone gneiss) [latitude 44.95677, longitude −115.36973]
20KL047-001	124	6		Yes
20KL047-002	225	37		No
20KL047-003	228	33		No
20KL047-004	287	12		No
20KL047-005	289	20		No
20KL047-006	386	7		No
20KL047-007	415	33		No
20KL047-008	423	30		No
20KL047-009	433	59		No
20KL047-010	464	71		No
20KL047-011	500	53		No
20KL047-012	508	33		No
20KL047-013	557	30		No
20KL047-014	563	92		No
20KL047-015	595	14		No
20KL047-016	670	100		No
20KL047-017	708	12		No
20KL047-018	810	140		No
20KL047-019	861	19		No
20KL047-020	865	22		Yes
20KL047-021	897	31		Yes
20KL047-022	910	120		Yes
20KL047-023	975	9		Yes
20KL047-024	992	18		Yes
20KL047-025	1000	24		Yes
20KL047-026	1009	19		Yes
20KL047-027	1021	37		Yes
20KL047-028	1022	11		Yes
20KL047-029	1028	54		Yes
20KL047-030	1028	37		Yes
20KL047-031	1035	13		Yes
20KL047-032	1042	22		Yes
20KL047-033	1044	16		Yes
20KL047-034	1048	23		Yes
20KL047-035	1054	19		Yes
20KL047-036	1056	21		Yes
20KL047-037	1062	21		Yes
20KL047-038	1063	21		Yes
20KL047-039	1064	13		Yes
20KL047-040	1065	21		Yes
20KL047-041	1071	21		Yes
20KL047-042	1074	13		Yes
20KL047-043	1077	25		Yes
20KL047-044	1078	45		Yes
20KL047-045	1078	9		Yes
20KL047-046	1079	26		Yes
20KL047-047	1079	10		Yes
20KL047-048	1093	21		Yes
20KL047-049	1097	9		Yes
20KL047-050	1105	33		Yes
20KL047-051	1113	21		Yes
20KL047-052	1122	17		Yes
20KL047-053	1143	23		Yes
20KL047-054	1151	10		Yes
20KL047-055	1156	18		Yes
20KL047-056	1158	25		Yes
20KL047-057	1161	12		Yes
20KL047-058	1175	20		Yes
20KL047-059	1176	13		Yes
20KL047-060	1179	23		Yes
20KL047-061	1182	23		Yes
20KL047-062	1184	24		Yes
20KL047-063	1186	23		Yes
20KL047-064	1189	17		Yes
20KL047-065	1194	16		Yes
20KL047-066	1197	24		Yes
20KL047-067	1206	13		Yes
20KL047-068	1206	27		Yes
20KL047-069	1208	14		Yes
20KL047-070	1211	36		Yes
20KL047-071	1211	25		No
20KL047-072	1223	31		Yes
20KL047-073	1234	26		Yes
20KL047-074	1242	42		Yes
20KL047-075	1244	11		Yes
20KL047-076	1246	30		No
20KL047-077	1249	31		Yes
20KL047-078	1258	25		Yes
20KL047-079	1262	28		Yes
20KL047-080	1281	30		Yes
20KL047-081	1351	52		Yes
20KL047-082	2438	16		No
20KL047-083	1369	40		Yes
20KL047-084	1376	50		Yes
20KL047-085	1446	13		Yes
20KL047-086	1625	19		Yes
20KL047-087	1770	14		No
20KL047-088	1551	13		Yes
20KL047-089	1433	38		Yes
20KL047-090	1440	19		Yes
20KL047-091	1442	24		Yes
20KL047-092	1431	13		Yes
20KL047-093	1487	40		Yes
20KL047-094	1727	37		Yes
20KL047-095	1454	40		Yes
20KL047-096	1458	47		Yes
20KL047-097	1758	37		Yes
20KL047-098	1813	37		Yes
20KL047-099	1569	41		Yes
20KL047-100	1527	38		Yes
20KL047-101	1506	41		Yes
20KL047-102	1582	23		Yes
20KL047-103	1730	16		Yes
20KL047-104	1598	39		Yes
20KL047-105	1683	13		Yes
20KL047-106	1724	52		Yes
20KL047-107	1785	38		Yes
20KL047-108	1865	39		Yes
20KL047-109	1790	20		Yes
20KL047-110	1741	39		Yes
20KL047-111	1791	14		Yes
20KL047-112	1855	37		Yes
20KL047-113	1867	37		Yes
20KL047-114	1874	37		Yes
20KL047-115	1844	11		Yes
20KL047-116	1836	13		Yes
20KL047-117	2502	9		No
20KL047-118	2560	8		Yes
20KL047-119	2475	13		Yes
20KL047-120	3237	32		Yes
08KL058 (Monumental Summit, metamorphosed quartz-pebble conglomerate) [latitude 44.72394, longitude −114.966278]
08KL058-53.1	1229	47		Yes
08KL058-49.1	1304	26		No
08KL058-18.1	1377	28		Yes
08KL058-50.1	1425	17		Yes
08KL058-9.1	1466	22		Yes
08KL058-52.1	1693	16		Yes
08KL058-30.1	1704	12		Yes
08KL058-41.1	1705	16		Yes
08KL058-40.1	1730	13		Yes
08KL058-46.1	1737	14		Yes
08KL058-8.1	1788	15		Yes
08KL058-13.1	1799	14		Yes
08KL058-44.1	1834	17		Yes
08KL058-15.1	1838	20		Yes
08KL058-11.1	1853	18		Yes
08KL058-21.1	2162	20		No
08KL058-34.1	2226	30		No
08KL058-33.1	2261	12		Yes
08KL058-7.1	2328	47		Yes
08KL058-28.1	2401	16		No
08KL058-32.1	2411	14		Yes
08KL058-29.1	2416	23		No
08KL058-56.1	2422	7		No
08KL058-5.1	2457	6		No
08KL058-16.1	2461	12		Yes
08KL058-27.1	2462	9		Yes
08KL058-2.1	2469	11		Yes
08kl058-36.1	2474	27		Yes
08KL058-54.1	2477	8		Yes
08KL058-60.1	2482	28		Yes
08KL058-4.1	2484	14		No
08KL058-22.1	2486	15		Yes
08KL058-20.1	2486	10		Yes
08KL058-6.1	2498	21		Yes
08KL058-51.1	2509	8		Yes
08KL058-12.1	2510	13		No
08KL058-24.1	2532	11		No
08KL058-23.1	2535	15		Yes
08KL058-17.1	2541	11		Yes
08KL058-39.1	2544	18		Yes
08KL058-38.1	2545	16		Yes
08KL058-47.1	2546	10		Yes
08KL058-35.1	2550	9		Yes
08KL058-58.1	2552	40		Yes
08KL058-14.1	2552	11		Yes
08KL058-45.1	2554	11		Yes
08KL058-37.1	2556	7		Yes
08KL058-10.1	2558	6		Yes
08KL058-43.1	2562	7		Yes
08KL058-1.1	2564	7		Yes
08KL058-55.1	2579	7		No
08KL058-59.1	2602	7		No
08KL058-42.1	2626	22		Yes
08KL058-31.1	2637	12		Yes
08KL058-25.1	2688	24		Yes
08KL058-19.1	2715	5		No
08KL058-26.1	2724	9		Yes
08KL058-48.1	2737	10		No
08KL058-57.1	3126	8		Yes
08KL058-3.1	3252	7		Yes

The Edwardsburg district contains the about 685-Ma Edwardsburg Formation of Lund and others (2003). These dated rocks and a straight section in this locality provide context for post-volcanic units which are also part of the Neoproterozoic Windermere Supergroup, as well as for several closely associated pre-volcanic units (Gospel Peaks A) which are considered to be probably Neoproterozoic but older than the Windermere Supergroup (Lund and others, 2003; Lund, 2004). Two samples of pre-volcanic (probably pre-Windermere Supergoup) units display detrital zircon populations with a broad array of provenance signatures (Placer lake and Whiskey Creek samples, fig. 5A) including Paleoproterozoic and Mesoproterozoic peaks similar to those of the Ordovician units in the Clayton and Bayhorse antiforms (fig. 5D, E) and Pennsylvanian–Permian Wood River Formation (fig. 5F), Paleoproterozoic and Mesoproterozoic peaks similar to those of the Mesoproterozoic Lemhi Group strata (for example, Link and others, 2007; Lund and others, 2008), an unusual peak at about 1,250 Ma, and several peaks in the 1.2–1.0 billion years (Ga) range which is similar to many of the younger rocks in this region (below).

The few Neoproterozoic Windermere Supergrpoup equivalent rocks that have been analyzed for detrital zircon are from the Edwardsburg section (Isakson, 2017; this study) and have distinctly different detrital zircon populations than Mesoproterozoic units. Several samples of this type, which depositionally overlie the Neoproterozoic Edwardsburg Formation metavolcanic units but are also correlated with the Windermere Supergroup (Lund and others, 2003; Lund, 2004), contain detrital zircon derived from the about 685 Ma metavolcanic rocks but the main peak ranges from about 1,000-1,600 Ma and samples contain some Paleoproterozoic zircon as well (Isakson, 2017). Strata that structurally overlie the 685 Ma metavolcanic units do not include zircon sourced from the volcanic rocks but only zircon of Mesoproterozoic and older ages (see McFadden Point sample, fig. 5A); however, based on the compositional and sedimentological features, these rocks are interpreted as post-Mesoproterozoic ages but probably pre-Windermere Supergroup (Gospel Peak A; Lund and others, 2003; Lund, 2004). There are no known Neoproterozoic strata equivalent to the Windermere Supergroup or the Gospel Peak A units exposed in south-central Idaho or in the southern Salmon River Mountains (see Neoproterozoic Packages section above) for provenance comparison.

The maximum ages from a number of lithologic types in the roof pendants of the southern cluster of districts, as provided by youngest zircon populations (Stewart and others, 2016; Isakson, 2017; table 3 and fig. 5A, B, C), corroborate the stratigraphic interpretations that they are post-Mesoproterozoic. The age data require that several of them are younger than the 685 Ma volcanic event and that a few are younger than Cambrian (fig. 5B). Many of the unknown-age samples from the Yellow Pine district contain zircon derived from the 685 Ma volcanic source or 1.2–1.0 Ga sources, although some contain evidence for both sources. Other samples from interlayered or structurally interleaved units contain only older zircon populations in the 1.7 Ga range although interpreted as probably Paleozoic units(fig. 5C). The lower Monumental Summit sample (figs. 5B) has similar zircon population to that of the Jackass Gulch sample (fig. 5A) which is located farther southeast (fig. 3) and between the Stibnite roof pendant and the southern Salmon River area; these samples contain a broad range of zircon ages between about 900 and 1300 Ma and are quite similar to the structurally lower schist sample (fig. 5C) and Whiskey Creek sample (fig. 5A) that are interpreted as older than Windermere Supergroup. With the available data, the presence or absence of the Neoproterozoic or late Mesoproterozoic detrital zircon do not seem to uniquely identify the Neoproterozoic rocks or discriminate them from younger strata.

The oldest Paleozoic strata from which detrital zircon data have been acquired in south-central Idaho are the Middle Ordovician Kinnikinic Quartzite (fig. 5I). Samples from this unit produced consistent Paleoproterozoic detrital zircon populations with few or no younger grains. In addition to the regional correlation of rocks of this unit, its zircon population further documents a Middle Ordovician sand wedge that spread from shelf to deep-water settings across from east- to south-central Idaho and that had a characteristic provenance (Baar, 2009). Samples of Devonian deep-water (lower part of the Milligen Formation, fig. 5G) and shelf-break (Jefferson Formation, fig. 5I), plus the Mississippian flysch strata (Copper Basin Group samples, fig. 5G), display the same Paleoproterozoic peak as the predominant signature. The detrital zircon age distribution for the Salmon River assemblage from the southern Salmon River Mountains (fig. 5D) is much the same as the Mississippian Copper Basin Group, with which it is correlated (Link and others, 1995). The thrust-sliver quartzite from the southern Stibnite roof pendant (fig. 5C) has a main Paleoproterozoic peak similar to this group. The upper quartzite from the Monumental Summit quartzite-conglomerate unit of the southern Stibnite roof pendant has some similar grains but also contains a larger, younger Paleoproterozoic population (fig. 5B) making it more similar to the Cambrian samples from the outer core of the Clayton antiform (fig. 5D). The Middle Pennsylvanian Hailey Member of the Wood River Formation (fig. 5F) contains the Paleoproterozoic detrital zircon peak similar to the Kinnikinic Quartzite but also contains significant amounts of younger zircon peaks, indicating mixed provenances (fig. 5I). These data indicate that a particular source terrain remained intermittently available, that sediment transport from it was active during much of the span from the Ordovician to Pennsylvanian, and that the sand wedge from it also spread westward into central Idaho.

Devonian slope-break and deep-water facies belts of south-central Idaho (fig. 4) both produced the two youngest peaks at about 500–450 Ma (Jefferson Formation [fig. 5I] and the upper part of the Milligen Formation [fig. 5G]). Samples from the Pennsylvanian–Permian epicratonic basin facies belt and from Upper Devonian deep-water facies of western south-central Idaho (fig. 4) also produced young peaks at about 500–450 Ma (fig. 5F). These young peaks from some of the Devonian strata—and those from the Wood River Formation—match peaks in several of the samples from the Yellow Pine district as in the Missouri Ridge, western facies sample, of the northern Stibnite roof pendant (fig. 5B). However, except for matching youngest grains, the overall relative-probability patterns (“barcodes”) from the known-age strata of south-central Idaho and Stibnite samples are not well matched.

The Middle to Upper Devonian Jefferson and Devonian Milligen Formations (fig. 5I, G) and the Pennsylvanian–Permian Wood River Formation (fig. 5F) contain late Mesoproterozoic detrital zircon (about 1.2–1.0 Ga). These zircon are also present in three of the curves shown from the southern Salmon River Mountains (fig. 5D, E), in two of the curves shown from the Yellow Pine district (both the basal conglomerate of the Monumental Creek unit of the northern Stibnite roof pendant, fig. 5B), in the structurally lowest schist unit of the southern Stibnite roof pendant, fig. 5C), in the unknown-age Jackass Gulch sample from southeast of the Yellow Pine district, and in the Neoproterozoic Placer lake sample (older than the 685 Ma volcanic rocks) from the Edwardsburg district (fig. 5A). Many other zircon peaks from these curves differ but it is clear that through time and across this entire region, the provenance switched repeatedly such that a late Mesoproterozoic source was only intermittently included.

Samples from south-central Idaho have fairly uniform intra-facies detrital zircon populations, indicating relatively consistent provenances and sediment transport systems. However, the metasedimentary rocks from the Yellow Pine district (Stewart and others, 2016; Isakson, 2017) produced six types of relative-probability patterns from 12 samples and additional patterns from samples analyzed for the present study (fig. 5A–C). These data suggest that the central Idaho rocks were derived from a variety of provenances and that the sources changed markedly during the depositional history. Thus, there are strikingly different zircon populations (that is, provenance and sediment transport factors) between south-central and central Idaho samples, although the rocks are along strike and parallel to the Neoproterozoic-Paleozoic continental margin (Lund, 2008; Lund and others, 2010).

Mineral Deposit Associations

The most well documented and complete of the central Idaho Neoproterozoic sections, located in the Gospel Peak and Big Creek roof pendants, are not associated with mineral deposits. Likewise, the probable Neoproterozoic rocks in the Stibnite roof pendant do not host the mineralized zones. The only Neoproterozoic rocks in the central Idaho mineral province that host deposits are in the Marshall Lake-Resort district, where carbonate-bearing biotite schist correlated with the upper part of the Neoproterozoic section (Lund, 2004) hosts gold-silver±lead-zinc±antimony±tungsten vein deposits.

The Paleozoic miogeoclinal facies belts, which cross from east- to south-central Idaho, are linked to mineral belts in a manner that matches a generalized model (Lund, 2008), although there is added complication where middle Paleozoic flysch basin and late Paleozoic onlap basin successions were telescoped by thrust faulting onto the miogeoclinal belts. The east-central Idaho shelf to slope facies strata are associated with silver-lead-zinc deposits of Mississippi Valley Type (MVT) and Sedex types, respectively. The Ordovician–Mississippian outer shelf, deep-water facies are overlapped by Mississippian deep-sea fan, turbidite (Antler flysch) basin strata, and gold is added as a significant commodity to deposits. Farther west, mines and prospects associated with Pennsylvanian–Permian epicratonic onlap basin facies are characterized by gold-silver-tungsten-antimony enrichments (Worl and Johnson, 1995). The setting associated with the most complicated metals assemblages in western south-central Idaho is where the Ordovician–Devonian deep-water facies were depositionally or structurally overlapped by the Pennsylvanian–Permian epicratonic-basin onlap facies (fig. 4; Hall and others, 1978; Cookro, 1985; Hall, 1985; Link and others, 1995).

Farther north, the Cambrian–Ordovician package on the western side of the southern Salmon River Mountains contains a greater proportion of coarse siliciclastic strata than more southerly sections do, and it hosts silver-lead and tungsten occurrences (Hobbs, 1985; Hobbs and others, 1991). Those deposits are located where the Cambrian–Ordovician package was structurally overridden from the west by the thrust plate of Mississippian deep-water and Pennsylvanian–Permian onlap facies and where Late Cretaceous or Eocene magmas intruded along the structures (Hall and others, 1978; Cookro, 1985; Hall, 1985). Each of these facies belts is matched with associated regional geochemical anomalies for the relevant metal signatures (Erdman and others, 1995). Lower Cambrian–Middle Ordovician phyllite and dolostone in the eastern part of the southern Salmon River Mountains, which may represent a significant facies transition from (or are not exposed in) south-central Idaho, host significant silver-lead-zinc vein deposits.

The central Idaho districts in and adjacent to roof pendants that contain most of the probable Paleozoic strata are those linked to gold-antimony-tungsten geochemical associations (Alminas, 1990) and to deposits containing these metals (table 2, fig. 1). This places the central Idaho tungsten-gold deposits and their host rocks north-northwest of tungsten-endowed mineral deposits of south-central Idaho.

Neoproterozoic–Paleozoic Facies Belts Reconstruction

Although the western Laurentian rift system and developing miogeocline were geographically continuous in Idaho (Lund, 2008; Lund and others, 2010), the details (see Neoproterozoic Packages and Paleozoic Packages sections, above) present along-strike contrasts among stratigraphic packages in central Idaho (Lund, 2004; Stewart and others, 2016), the southern Salmon River Mountains (Hobbs and Hays, 1990; Brennan and others, 2020), and south-central Idaho (Isaacson and others, 1983; McFadden and others, 1988; Grader and Dehler, 1999). Information from the stratigraphic packages indicate a thick Neoproterozoic section is present in central Idaho but absent in south-central Idaho, whereas a broader, thicker, and more complete set of Paleozoic miogeoclinal facies belts is present in east- and south-central Idaho compared to central Idaho (figs. 1B, 4).

Available detrital zircon analyses from probable Paleozoic sections in central Idaho indicate diverse provenances and some unique populations (which can be evaluated independent of stratigraphic stacking or structural complications) in comparison to the known-age strata in south-central Idaho (compare fig. 5A–C to F–I and D–E). Together, the stratigraphic and detrital zircon comparisons reveal that central Idaho had the more tectonically active Paleozoic basin histories, as shown by (1) the more abrupt changes of sedimentary character—quartz-rich siliciclastic sheets alternating with carbonate and carbonaceous-silt deposition—as well as (2) distinctly variable provenances (sediment sources).

A significant change in Paleozoic stratigraphic packages and provenances occurs in the southern Salmon River Mountains (fig. 1B) where Hobbs and others (1991) describe the juxtaposition of coeval stratigraphic sections with contrasting stratigraphic and compositional characteristics. Similarly, the Neoproterozoic rocks in central Idaho are a thick section that is different from the only occurrence documented in a deep drillhole in the eastern part of the southern Salmon River Mountains area (Lund and others, 2010) and from the thin upper Neoproterozoic unit locally present in east-central Idaho (McCandless, 1982; Hobbs and others, 1991). The details suggest that there was a long-lived transition zone in Neoproterozoic–Paleozoic depositional basins in this area (as suggested by Hobbs and others, 1991). This transition zone probably originated as a northeast-striking transform in the Neoproterozoic through middle Paleozoic miogeoclinal rift system(s) (for example, Lund and others, 2003, 2010) and is herein named the Salmon River transform (fig. 4; Salmon River lineament of Hobbs and Hays, 1990; Hobbs and others, 1991).

This along-paleomargin transition in the Neoproterozoic–Paleozoic sedimentary rocks frustrates correlation of units from south-central to central Idaho but also clarifies the problem. The evidence from stratigraphic, compositional, and provenances for Neoproterozoic–Paleozoic strata of central Idaho compared to south-central Idaho suggests that the Paleozoic miogeoclinal basin tectonic history diverged across the Salmon River transform, such that, for the northern Paleozoic rocks, the miogeoclinal rift basin narrowed, the stratigraphic sections thinned, and the deposits were of a coarser nature. However, there are indications of continuity for stratigraphic packages, such as the detrital zircon data that partially constrain ages and provide some indications of corresponding sediment sources, the stratigraphic characteristics (compositions and their thicknesses and alternation styles) that are similar along strike, and the presence of gold-silver-antimony-tungsten deposits. Together, continuity of these characteristics along strike of the paleomargin and across the transition suggests some continuity in the essential characteristics of sedimentary deposits and perhaps in the underlying crust. These data particularly suggest that some proportion of the metasedimentary rocks in the Yellow Pine district (and potentially to lesser degrees in the Profile, Edwardsburg, and to Marshall Lake-Resort districts) likely include northern correlatives to the western Pennsylvanian–Permian epicratonic facies and to early Paleozoic deep-water facies.

Mesoproterozoic Salmon River Mountains Orthogneiss

Mesoproterozoic foliated porphyritic granite and granodiorite and augen gneiss is discontinuously exposed in a northwest-striking belt across central Idaho. They are dated at 1380–1360 Ma (U-Pb ages determined from zircon, Evans and Zartman, 1990; Aleinikoff and others, 2012) and classified as peraluminous calc-alkalic, ferroan granitoids (du Bray and others, 2018). They intruded the middle formations of the Mesoproterozoic Lemhi Group and gneisses correlated with those strata. Gold-polymetallic vein deposits in the Orogrande, Elk City, and Dixie mining districts (fig. 2) are locally hosted in these granitoids but are also in Mesoproterozoic metasedimentary and Cretaceous granitic rocks (table 2). Although an early phase of xenotime in gold-cobalt-copper deposits of the Blackbird mining district of east-central Idaho is dated at about the same age as a structurally juxtaposed Mesoproterozoic granitoid (Aleinikoff and others, 2012), most of the datable minerals related to the deposits produced Cretaceous ages (Lund and others, 2011; Aleinikoff and others, 2012). Similarly, the deposits that are hosted in Mesoproterozoic augen gneiss of the central Idaho districts are structurally controlled, and available dates indicate mineralizing events were Cretaceous (Lund and others, 1986) and not related to the Mesoproterozoic magmatism.

Neoproterozoic Syenite-Diorite Suites of Big Creek-Beaverheads Belt

Neoproterozoic syenite-diorite suites in the Ramey Ridge and Big Creek districts (figs. 2, 3) are dated at about 650 Ma and help mark the strike and general location of a second pulse of Neoproterozoic continental rifting that followed the initial phase identified by the nearby about 685-Ma volcanic rocks (Lund and others, 2003, 2010). The syenite-diorite complexes and the Mesoproterozoic strata they intruded both host the structurally controlled mineral deposits. Those in syenite-diorite complexes are precious-metal rich, one bearing stibnite, whereas the deposits in the older country rocks are copper rich (Cater and others, 1973). Because of their structural control, deposits are presumed to be Cretaceous or possibly Eocene in age and preliminary dating of the deposits hosted in the syenite-diorite suites resulted in ages of 79–65 Ma (Lund, 2004).

Late Cretaceous Idaho Batholith

Based on geologic-map geometries and ages of intrusive units, the northern and central Atlanta lobe of the late Cretaceous Idaho batholith (which underlies the central Idaho mineral province) is interpreted to have age- and composition-defined shells (see Lund, 2004). Most mining districts in the province are located near exposures of metasedimentary roof rocks (in three dimensions) and are seldom found in similar age or similar composition plutons where those are not adjacent to the metasedimentary rocks. Our combined geologic mapping and geochronology are applied to test the geometry and veracity of the semi-concentric shells and most particularly to identify crustal depths related to fault offsets (that is, uplifted or tilted blocks) and the relative offset amounts, timing, and duration of such structural events. This is another way to test the controls for mineral deposit localization and also potentially for controls on preservation.

U-Pb Ages across Atlanta Lobe

Many of the mining districts are hosted in Late Cretaceous plutons of the Idaho batholith and the interpretation of timing, amount of offset, and sense of movement on bounding faults can only be gained through details relating to emplacement ages and geometry of plutons. Preliminary ages for Late Cretaceous granites of the northern Atlanta lobe are presented in Unruh and others (2008, TIMS U-Pb method [thermal ionization mass spectrometry] and multigrain analyses). However, many of those ages are relatively imprecise because of zircon with inherited cores or overgrowths caused by multiple events. To mitigate these problems, samples from across the Atlanta lobe, extending from adjacent to the Salmon River suture on the west to the central Yellow Pine district, are reanalyzed by the SHRIMP U-Pb method (fig. 6, table 4). Data for these samples are in Aleinikoff and others (2023). These new emplacement ages help constrain some of the mineralization and fault movement questions for the mineral province. Results complement ages for Stibnite area granitoids and for mineral deposits in the central Yellow Pine to Thunder Mountain districts (tabulated in Gillerman and others, 2019).

Table 4.    

Sensitive high resolution ion microprobe (SHRIMP) uranium-thorium-lead (U-Th-Pb) data for zircon from intrusive rocks, west-central Idaho.

[Data are available in Aleinikoff and others (2023). Datum for latitude-longitude is WGS 84%, percent; ppm, parts per million; Ma, million years; err, error; ---, no data]

Table 4.    Sensitive high resolution ion microprobe (SHRIMP) uranium-thorium-lead (U-Th-Pb) data for zircon from intrusive rocks, west-central Idaho.

sample*	measured 204Pb/206Pb	measured 207Pb/206Pb	% common 206Pb	U (ppm)	Th/U (ppm)	206Pb†/238U (Ma)	err§ (Ma)	238U†/206Pb	err§ (%)	207Pb†/206Pb	err§ (%)
MC13-91 (Upper Payette Lake foliated hornblende-biotite tonalite) [latitude 45.1040, longitude −116.0316]
MC13-1.1	---	0.0427	−0.64	246	0.36	88.7	2.1	72.16	2.3	0.0427	9.6
MC13-2.1	---	0.044	−0.48	237	0.43	88.8	2.8	72.06	3.1	0.044	9.6
MC13-3.1	---	0.0447	−0.38	310	0.25	87.6	1.9	73.05	2.2	0.0447	8.6
MC13-4.1	---	0.0483	0.06	188	0.42	89.2	2.2	71.81	2.5	0.0483	10.3
MC13-5.1	---	0.0453	−0.31	232	0.39	87.2	2.1	73.43	2.4	0.0453	9.6
MC13-6.1	0.001526	0.0508	0.39	308	0.38	85.2	2.2	75.13	2.6	0.0278	44.5
MC13-7.1	0.001947	0.0556	0.99	254	0.42	86.3	2.2	74.17	2.5	0.0261	38.9
MC13-8.1	0.002608	0.0501	0.25	203	0.53	96.4	3.6	66.38	3.8	---	---
MC13-9.1	---	0.049	0.1	587	0.16	110.6	1.8	57.81	1.7	0.049	5
MC13-10.1	0.000472	0.0482	0.05	907	0.52	89.9	1.5	71.23	1.7	0.0412	10
MC13-11.1	---	0.0462	−0.20	294	0.26	91.6	2	69.87	2.2	0.0462	8.9
MC13-12.1	---	0.0515	0.48	364	0.2	82.5	1.6	77.6	2	0.0515	6.8
MC13-13.1	---	0.0492	0.18	573	0.15	87.1	1.6	73.49	1.8	0.0492	6.1
MC1-91 (Paddy Flat locally foliated, biotite granodiorite) [latitude 44.7844, longitude −115.8744]
MC1-1	0.000079	0.0487	0.15	2582	0.67	75.1	0.8	85.32	1.1	0.0476	2.5
MC1-2	---	0.0499	0.3	934	0.27	77.8	0.9	82.33	1.2	0.0499	2.8
MC1-3	0.000459	0.0469	−0.09	372	0.31	80.8	1.2	79.27	1.4	0.0401	9.3
MC1-4	0.00026	0.047	−0.07	391	0.81	78.4	1.1	81.72	1.4	0.0431	6.6
MC1-5	0.000053	0.0485	0.1	2592	0.73	86	0.9	74.46	1.1	0.0477	1.9
MC1-6	---	0.0472	−0.05	349	0.87	79	1.1	81.14	1.4	0.0472	7.9
MC1-7	0.000367	0.0521	0.6	360	0.94	70.7	1	90.67	1.5	0.0467	7.4
MC1-8	---	0.048	0.04	858	0.34	79.9	0.9	80.23	1.2	0.048	2.9
MC1-9	0.000158	0.0481	0.06	655	0.19	77.5	1	82.64	1.3	0.0457	4.4
MC1-10	0.000061	0.0489	0.15	1983	0.73	83.5	0.9	76.69	1.1	0.048	2.2
MC1-11	---	0.0457	−0.25	389	0.64	81.7	1.2	78.42	1.5	0.0457	5
MC1-12	0.000155	0.0475	−0.02	801	0.11	80.3	1	79.78	1.2	0.0452	4.3
MC1-13	---	0.0509	0.41	447	0.2	84.4	1.1	75.87	1.3	0.0509	3.9
MC1-14	---	0.0477	0.01	385	0.72	81.2	1.2	78.87	1.4	0.0477	5
MC1-15	---	0.0458	−0.22	446	1.55	78.1	1.1	82.05	1.4	0.0458	4.4
91KL139 (Chimney Rock muscovite-biotite granite) [latitude 45.3328, longitude −115.7778]
91KL139-1.1	0.000783	0.0514	0.45	807	0.25	90.4	3.2	71.21	3.6	0.0397	10.6
91KL 139-2.1	0.000316	0.0493	0.2	803	0.17	85.8	3	75.03	3.6	0.0446	6.2
91KL 139-3.1	0.000351	0.0479	0.04	1808	0.11	81.5	2.9	79.06	3.5	0.0427	4.9
91KL 139-4.1	0.000487	0.057	1.18	387	0.27	83.4	3.1	77.23	3.7	0.0499	8.9
91KL 139-5.1	0.001567	0.0656	2.27	1219	0.16	79.5	2.9	81.17	3.7	0.0423	15.8
91KL 139-6.1	0.000466	0.0484	0.09	497	0.17	82.3	3	78.29	3.6	0.0415	8.7
91KL 139-7.1	---	0.0448	−0.36	600	0.55	81.2	2.9	79.31	3.6	0.0448	3.8
91KL 139-8.1	---	0.0488	0.15	410	0.45	81.8	3	78.79	3.7	0.0488	4.4
91KL 139-9.1	---	0.0485	0.1	1558	0.13	84.2	3	76.45	3.5	0.0485	2.5
91KL 139-10.1	0.000646	0.0443	−0.41	294	0.68	78.3	2.9	82.31	3.8	0.0346	16.3
91KL 139-11.1	-0.000532	0.0507	0.38	474	0.28	82.5	3	78.02	3.7	0.0584	9.5
91KL 139-12.1	0.000337	0.0495	0.23	1114	0.11	82.8	2.9	77.79	3.6	0.0445	5.4
91KL 139-13.1	-0.000017	0.048	0.04	2549	0.05	84.6	3	76.16	3.5	0.0483	1.7
91KL 139-14.1	0.000337	0.0491	0.19	860	0.49	80.7	2.9	79.86	3.6	0.0441	6.3
91KL 139-15.1	0.000328	0.0501	0.26	1376	0.26	99.8	3.5	64.49	3.5	0.0452	5.1
MC15-91 (East Fork South Fork muscovite-biotite granite) [latitude 45.0056, longitude −115.7444]
MC15-1.1	0.000153	0.0459	−0.22	1272	0.26	80.2	0.9	79.88	1.2	0.0436	3
MC15-2.1	0.000127	0.0491	0.2	1087	0.15	77.7	0.8	82.44	1	0.0472	3.2
MC15-3.1	0.000581	0.0488	0.14	2043	1.35	82.6	0.9	77.56	1	0.0402	8.1
MC15-4.1	0.001108	0.0506	0.37	214	0.28	81	1.5	79.12	1.8	0.034	28
MC15-5.1	---	0.0492	0.22	139	0.33	74.6	1.3	85.89	1.8	0.0492	7.2
MC15-6.1	−0.000269	0.0485	0.08	1008	0.11	93.3	1	68.59	1.1	0.0525	4.8
MC15-7.1	---	0.0502	0.31	216	0.24	87	1.2	73.6	1.4	0.0502	4.9
MC15-8.1	0.000373	0.0501	0.32	297	0.4	78.7	1.4	81.39	1.7	0.0446	8.1
MC15-9.1	0.000279	0.0475	−0.01	876	0.71	79.3	0.9	80.83	1.1	0.0434	5.2
MC15-10.1	0.001145	0.0566	1.15	97	0.38	70.9	1.6	90.48	2.3	0.0395	24.3
MC15-11.1	---	0.0483	0.1	604	0.47	77.2	0.9	82.98	1.1	0.0483	3.1
MC15-12.1	0.000835	0.0522	0.59	186	0.22	73.4	1.3	87.34	1.8	0.0398	18.8
MC15-13.1	0.000026	0.0475	−0.04	14579	0	92.6	0.8	69.13	0.9	0.0471	0.7
MC15-14.1	0.000065	0.0492	0.2	4060	0.08	79	0.7	81.1	0.9	0.0482	1.7
MC15-15.1	0.000119	0.047	−0.07	2072	0.04	79.7	0.8	80.41	1	0.0453	1.9
MC15-16.1	---	0.0459	−0.24	309	0.21	90.1	1.1	71.01	1.3	0.0459	4.1
MC15-17.1	---	0.0539	0.8	163	0.57	74.4	1.2	86.15	1.6	0.0539	5.8
MC15-18.1	---	0.0481	0.07	2495	0.1	78.9	0.7	81.22	0.9	0.0481	1.4
MC15-19.1	0.000068	0.0472	−0.05	2637	0.11	77.8	0.8	82.39	1	0.0462	1.8
MC15-19.2	0.000076	0.0465	−0.18	1957	0.28	94.4	1.1	67.82	1.1	0.0454	2.2
MC15-20.1	0.000478	0.0479	0.06	859	0.22	73.7	0.8	86.99	1.1	0.0408	8.1
MC15-20.2	---	0.1771	6.95	348	0.06	2023	18.2	2.71	1	0.1771	0.4
MC15-21.10	---	0.0482	0.07	2110	0.19	79.9	0.8	80.16	1	0.0482	2.1
MC15-22.10	0.000083	0.0472	−0.03	1201	0.4	73.7	0.7	86.99	1	0.046	2.6
MC92-36 (Johnson Creek foliated porphyritic epidote-hornblende-biotite granodiorite) [latitude 44.7861, longitude −115.5285]
MC92-36-2.1	---	0.0497	0	715	0.39	93.6	0.9	68.4	0.9	0.0497	2.2
MC92-36-3.1	0	0.0467	0	887	0.56	93.7	0.6	68.3	0.6	0.0467	2.1
MC92-36-5.1	0.000141	0.0481	0.26	845	0.61	92.7	0.6	69	0.7	0.0461	3.2
MC92-36-6.1	0.000257	0.0492	0.48	600	0.43	92.8	0.6	68.9	0.6	0.0454	4.7
MC92-36-7.1	0.000308	0.0498	0.57	803	0.5	93.5	1.1	68.4	1.2	0.0452	4.1
MC92-36-8.1	---	0.0504	0	790	0.53	91.3	1.1	70.1	1.2	0.0504	2.2
MC92-36-9.1	0.000035	0.0495	0.07	875	0.53	92.2	1.3	69.4	1.5	0.049	2.4
MC92-36-10.1	0.000206	0.0487	0.38	746	0.48	94.2	0.5	67.9	0.6	0.0456	5.3
MC92-36-11.1	−0.000039	0.0491	−0.07	759	0.47	93.3	0.9	68.6	1	0.0497	2.7
MC92-36-12.1	0.000316	0.0483	0.59	778	0.43	91.9	0.7	69.7	0.7	0.0436	4.6
MC92-36-13.1	0.000062	0.0487	0.12	1359	1.03	95.5	0.9	67	0.9	0.0478	3
MC92-36-14.1	0.000215	0.0483	0.4	856	0.55	92.4	0.7	69.3	0.8	0.0451	3.7
MC92-36-15.1	0.000159	0.0499	0.3	592	0.46	92.5	1.9	69.2	2.1	0.0476	4
MC92-36-16.1	0.000112	0.0486	0.21	830	0.4	93.3	0.8	68.6	0.9	0.0469	3.1
MC92-36-17.1	0.000061	0.0475	0.11	969	0.49	94.8	0.5	67.5	0.6	0.0466	2.5
MC92-36-18.1	0.000029	0.048	0.05	1044	0.42	94.3	0.6	67.9	0.6	0.0475	2.2
MC92-36-19.1	0.000184	0.048	0.34	1026	0.35	93.9	0.8	68.1	0.9	0.0452	3.2
MC92-36-20.1	−0.000051	0.0492	−0.09	1142	0.68	94.4	0.7	67.8	0.8	0.0499	2.1
MC92-36-21.1	0.000178	0.0481	0.33	867	0.55	93.1	0.8	68.8	0.9	0.0454	3.5
MC92-36-22.1	0.000128	0.0493	0.24	960	0.66	91.6	0.6	69.9	0.7	0.0474	2.9
MC92-36-1.1r	0.000072	0.0457	0.134	582	0.18	86.4	1.2	74.1	1.4	0.0446	3.7
MC92-36-2.1r	0.001107	0.0494	2.055	292	0.24	88.3	1	72.5	1.1	0.0328	22.1
MC92-36-3.1r	0.001524	0.0652	2.83	560	0.18	88.7	1.8	72.2	2.1	0.0426	14.3
MC92-36-4.1r	0.00008	0.051	0.149	515	0.18	87.2	0.8	73.5	0.9	0.0499	5.5
MC92-36-5.1r	0.002055	0.0505	3.817	324	0.26	83.5	1.6	76.7	2	0.0192	48.4
MC92-36-6.1r	0.001197	0.0478	2.224	410	0.17	87.6	1.7	73.1	1.9	0.0298	21.5
MC92-36-8.1r	0.000458	0.0512	0.851	519	0.18	86.4	1.6	74.1	1.9	0.0444	7.7
MC92-36-10.1r	0.000101	0.0553	0.188	370	0.39	85.2	0.9	75.2	1	0.0538	4.2
MC92-36-11.1r	0.002719	0.0614	5.053	176	0.3	80.8	2.1	79.2	2.6	0.0199	69.3
MC92-36-13.1r	0.000301	0.0543	0.559	563	0.19	84	2.4	76.2	2.8	0.0499	5.3
MC92-36-14.1r	0.003755	0.0567	6.976	365	0.18	81.5	1.9	78.6	2.4	0.0021	559.1
MC92-35 (Pepper Creek foliated hornblende-biotite granodiorite) [latitude 44.9019, longitude −115.3910]
MC92-35-1.1i	0.000058	0.0458	0.11	917	0.21	95.1	0.5	67.3	0.6	0.0449	2.5
MC92-35-2.1i	0.000267	0.0514	0.5	537	0.11	94.6	0.8	67.7	0.9	0.0475	4.7
MC92-35-3.1i	0.000098	0.0475	0.18	586	0.14	94.8	0.9	67.5	0.9	0.0461	3.4
MC92-35-4.1i	0.000236	0.0457	0.44	492	0.12	94.2	1.1	67.9	1.1	0.0421	8.7
MC92-35-5.1i	0.000268	0.0502	0.5	962	0.17	97.7	0.8	65.5	0.8	0.0462	3.6
MC92-35-6.1i	0.000108	0.0484	0.2	816	0.86	94.8	0.7	67.5	0.8	0.0468	3
MC92-35-7.1i	0.000057	0.0458	0.11	1007	0.24	95.4	1.1	67.1	1.1	0.0449	2.6
MC92-35-9.1i	−0.000092	0.0496	−0.17	949	0.17	97.4	1	65.7	1.1	0.051	5.3
MC92-35-10.1i	0.000064	0.0498	0.12	1320	0.28	98.2	0.7	65.2	0.8	0.0489	2
MC92-35-11.1i	0.000147	0.0488	0.27	1359	0.17	98.1	0.5	65.2	0.6	0.0466	2.5
MC92-35-12.1i	0.000371	0.0492	0.69	958	0.22	95.4	0.7	67.1	0.8	0.0438	4.3
MC92-35-13.1i	−0.000063	0.0469	−0.12	951	0.07	95.9	0.5	66.7	0.6	0.0479	2.4
MC92-35-14.1i	---	0.0469	0	620	0.1	93.7	1	68.3	1.1	0.0469	2.5
MC92-35-16.1i	0.000028	0.0497	0.05	1018	0.19	94.8	0.5	67.5	0.5	0.0493	3.4
MC92-35-17.1i	−0.000126	0.0496	−0.23	943	0.14	95	0.7	67.3	0.8	0.0515	2.6
MC92-35-18.1i	−0.000212	0.0486	−0.39	283	1.32	94.7	0.7	67.6	0.7	0.0517	7.6
MC92-35-19.1i	0	0.0478	0	1663	0.01	95.8	0.5	66.8	0.5	0.0478	1.5
MC92-35-20.1i	0.000122	0.0479	0.23	244	0.85	93.3	1.2	68.6	1.3	0.0461	5.8
MC92-35-21.1i	0.0001	0.0479	0.19	1245	0.11	97	0.6	66	0.6	0.0465	3.6
MC92-35-3.1o	---	0.0461	0	829	0.19	97.7	1.1	65.5	1.1	0.0461	2.1
MC92-35-6.1o	0.000039	0.0476	0.072	932	0.18	96.9	1.3	66	1.3	0.047	2.4
MC92-35-7.1o	−0.000026	0.0485	−0.049	1335	0.26	95.9	0.7	66.7	0.7	0.0489	3.4
MC92-35-8.1o	0.000466	0.0526	0.865	880	0.12	92	2.2	69.5	2.4	0.0457	5.1
MC92-35-9.1o	0.000181	0.046	0.336	444	0.12	93.3	1.6	68.6	1.7	0.0433	5.5
MC92-35-10.1o	0.000049	0.0475	0.091	1228	0.26	98.1	0.6	65.2	0.7	0.0468	1.9
MC92-35-11.1o	0.000163	0.0514	0.302	983	0.19	96.1	1.3	66.6	1.4	0.049	2.9
MC92-35-13.1o	0.00054	0.0518	1.002	835	0.16	97.5	1.2	65.6	1.2	0.0438	5.9

^*

Analytical sessions: All samples analyzed on the USGS/Stanford SHRIMP-Reverse Geometry. Dates of isotopic measurements: MC13-91 (3/2002), MC1-91 (1/2001), MC91-28 and MC15-91 (10/2002), MC92-36 and MC92-35 cores (10/2017), MC92-36 and MC92-35 rims (12/2017). Sample names abbreviations: r (rim), i (inner area), o (outer area); all other data obtained from analyses of cores.

^†

Corrected for common Pb using the ²⁰⁴Pb-correction method.

^§

1-sigma error.

Eocene Challis Volcanic-Plutonic Complex

With the exception of the Thunder Mountain district, the central Idaho mineral province is hosted by older rocks west of exposures of Eocene Challis volcanic and plutonic rocks. Emplacement ages are not available for most central Idaho Eocene intrusive rocks (plutons, stocks, or dikes) but, regionally, cooling ages range from about 51 to 37 Ma and dates for the volcanic rocks fall in the same range (Leonard and Marvin, 1982; Fisher and others, 1992).

For the area near the northern cluster of mining districts, a large body of moderately shallowly emplaced, Eocene monzogranite plutons forms most of the Chamberlain Basin pluton east of the Bargamin Creek and Johnson Creek-Profile Gap faults (fig. 3). Closely related, shallowly intruded, epizonal granites are between the en echelon parts of the Bargamin Creek and Johnson Creek-Profile Gap faults. Farther west between the Bargamin Creek and Dixie Meadows-Blanco Creek faults, shallowly emplaced, granite porphyry to andesite dikes intruded the older country rocks. These relations indicate stepwise changes in exposed crustal levels across the faults—deeper exposures are on the east and shallower on the west. For the southern cluster of mining districts, the Profile Gap dike swarm (Leonard and Marvin, 1982; Lund, 2004), along the central part of the Johnson Creek-Profile Gap fault zone, is composed of such a large volume of coalescing granite porphyry dikes that the country rocks are difficult to locate. These and the adjacent Savage Point granite porphyry pluton, which lies west of the shear zone, contain a significant proportion of miarolitic cavities, indicating very shallow emplacements (probably 2–4 km depth, Lund, 2004). The Pistol Creek dike swarm ranges in composition from rhyolite to diorite, extends southwest along the trace of the Thunder Mountain/Big Creek graben, and is structurally controlled (fig. 3).

The Thunder Mountain cauldron complex (Fisher and others, 1992) of the Eocene Challis Volcanic Group is exposed east of the Coin Mountain fault, a structure east of and related to the Johnson Creek-Profile Gap shear zone (figs. 3, 7). Three eruptive calderas are proposed for the main stage of volcanic activity as part of the Thunder Mountain cauldron complex (Leonard and Marvin, 1982; Stewart and others, 2016). The main units (“Tcrd, rhyolite and dacite, undifferentiated” on fig. 7) are an estimated 1,500-m thick, latite and rhyolite ash-flow tuff succession. These were overprinted by the caldera-collapse graben of the Thunder Mountain district (Leonard and Marvin, 1982; Hardyman and Fisher, 1985; named Big Creek graben at its northern extent by Stewart and others, 2013, 2016) that preserved the about 300-m thick, late-stage, rhyolite flows, tuffs, and volcanoclastic deposits (“Tcs, rhyolite tuffs and sedimentary rocks of Sunnyside” on fig. 7). These deposits in the collapse graben were erupted from a local vent, and the upper layers were interlayered with volcaniclastic conglomerate to sandstone, mudflows, volcanic breccia, and intercalated lacustrine sediments that include lignite and plant debris (the “Dewey beds”; Umpleby and Livingston, 1920; Ross, 1933; Shannon and Reynolds, 1975; Leonard and Marvin, 1982; Adams, 1985; Parsley, 1997). Ages for main phases of the Thunder Mountain cauldron complex are 49–48 Ma and for late-stage caldera-collapse units are approximately 46.7 Ma (D. John, U.S. Geological Survey, 2019, oral commun.).

The western boundary of the Eocene volcanic-plutonic Thunder Mountain cauldron complex is mapped in most places as a moderately to steeply dipping, faulted margin (Lund, 2004; Stewart and others, 2016), although younger graben structures cut the basal contact locally. The graben structures controlled and localized the deposition of younger volcanic units southwest of and outside of the main cauldron complex (figs. 3, 7; Stewart and others, 2016). Using the structures and elevation differences between highest and lowest volcanic exposures, estimates for thickness of the main volcanic deposits are a minimum thickness of about 800 m near the western margin and of about 300 m for the deposits in the graben.

Northern Area Faults

The northwestern cluster of mining districts is bounded on the east by a subparallel set of north-northeast-striking steep faults (fig. 2), primarily the Bargamin Creek fault zone (Weis and others, 1972; Cater and others, 1973; Greenwood and Morrison, 1973) plus the Blanco Creek (Lewis and others, 1990; 1998) and en echelon Dixie Meadows faults (named herein). West of the Bargamin Creek and northern Johnson Creek-Profile Gap faults, the country rocks are Late Cretaceous granitic rocks plus Mesoproterozoic and Neoproterozoic metasedimentary roof rocks (fig. 3). Westward from those faults, the metasedimentary rocks become progressively younger and (or) structurally higher. In the area between the Bargamin Creek and Blanco Creek-Dixie Meadows faults, the older rocks were intruded by shallowly emplaced Eocene epizonal stocks and dike-swarms. East of the Bargamin Creek fault, the Eocene Chamberlain Basin pluton is exposed to deep enough levels that no chilled upper margin, older roof rocks, or dike swarms are preserved above it. That different depths of Eocene intrusive systems are expressed across the faults, and that rocks as young as Eocene were deformed locally, indicate that faulting continued until after emplacement of the Eocene intrusive rocks.

The country-rock relations indicate down-to-the-west offsets across this set of faults. Mylonitic fabrics in Mesoproterozoic orthogneiss and in a few Eocene dikes along the Bargamin Creek fault (at about 45°30’ on fig. 3) also record down-to-the-west senses of displacement. The amounts of offset across the Bargamin Creek fault after most of the Eocene offsets would be the difference between emplacement depths greater than about 6 km (below the chill-zone level) in the Chamberlain Basin pluton on the east side of the fault and emplacement of less than 2 km (very shallow) for Eocene intrusive rocks and even shallower overlying roof rocks on the west side of the fault. These provide an estimate of more than 4 km of offset across the Bargamin Creek fault. The metamorphic rocks in the Buffalo Hump roof pendant are a 600-m thick layer above the mesozonal Late Cretaceous granites that are determined to have been emplaced at about 9 km depth (Lund and others, 1986). In contrast, epizonal Eocene granites in the central zone between the Dixie Meadows-Blanco Creek and Bargamin faults were emplaced at depths of about 4 km.

These data indicate significant (about 5 km) uplift in the interval between the Late Cretaceous and Eocene. The increasing amounts of Eocene intrusive rocks east of the Dixie Meadows-Blanco Creek fault and particularly the Bargamin Creek and northern Johnson Creek-Profile Gap faults suggest that the faults exerted control on the western extent of Eocene magmatism. Further, the location of rock types across the northern part of figure 3 describe a post-middle Eocene, regional west-dipping crustal block. The mineral deposits of the northwestern district cluster all lie west of the Bargamin Creek fault in minor fractures that are generally parallel to the main north-northeast striking structures (Chauvot, 1986; Lund and Esparza, 1990).

At its southwestern end, movement on the Bargamin Creek fault transferred into the South Fork and Smith Saddle faults (named herein; figs. 2, 3), which are both characterized by deformation fabrics in Late Cretaceous granitic rocks. Where offset of metasedimentary rocks can be determined across the South Fork fault, general movement was down to the west. Offset is not determined for the Smith Saddle fault, but because it cut and deformed Late Cretaceous granitic rocks, the time of fault movement is constrained as Late Cretaceous or younger. Here, the presence of sparse metasedimentary roof pendants on the west side of the Bargamin Creek fault compared to more widespread and generally older strata on the east side could suggest there was also activity on this part of the fault prior to Cretaceous intrusion as well as after. Motion across the Bargamin Creek fault also transferred southward to the northern extent of the regional Johnson Creek-Profile Gap shear zone (Lund, 2004). At this northern end of the Johnson Creek-Profile Gap shear zone, rocks as young as Eocene dikes locally display deformation fabrics. The relations among Mesoproterozoic roof rocks and Late Cretaceous intrusive rocks on the west and among Eocene intrusive rocks on the east indicate down-to-the-west offset of several kilometers after Eocene intrusions, similar to offsets determined for the Bargamin Creek fault.

Central Scissoring Zone

South of about 45°15’, mining districts of the southeastern cluster lie within or east of the Johnson Creek-Profile Gap shear zone (fig. 2). A 3-km long segment in the western Edwardsburg district is the only part of this fault where Mesoproterozoic roof pendants, Late Cretaceous intrusive rocks, and Eocene intrusive rocks are all present on both sides. This segment was intruded by the dikes and small stocks of the Eocene Profile Gap dike swarm, and the fault segment also controlled the location of a subvolcanic complex (fig. 3; the Wolf Fang complex, Lund, 2004). The Late Cretaceous granitic rocks are intensely silicified in a more than 5-km wide zone and discrete shear zones display local, outcrop-scale, dextral ductile structures (Lund, 2004; Montz and others, 2013). North of this segment, the Johnson Creek-Profile Gap has west-side-down offset (See section Northern Area Faults, above) but, south of the segment, the fault had east-side-down offset (See section Southern Area Faults, below). That megascopic evidence indicates that this segment is a zone where dip-slip motion reversed and is thus a zone of scissoring on the fault. Rotational movement would be expected and is probably reflected by the local dextral fabrics.

The silicification and shear fabrics in Late Cretaceous granitic rocks along the fault zone indicate significant activity after emplacement of the Late Cretaceous intrusive rocks. At present, there is no known evidence from older rocks that the fault had pre-Late Cretaceous history. The emplacement of the Eocene Profile Gap dike swarm and subvolcanic complex into the fault zone indicates that the fault controlled Eocene igneous activity and was active during the Eocene.

Southern Area Faults

The southeastern cluster of districts in the central Idaho metallogenic province lies east of a set of large faults (Fisher and others, 1992; Lund, 2004). These faults are the southern extension of a large regional fault system, which includes the Bargamin Creek and related faults at the north and which extend for nearly 100 km from north to south. In the southeastern district cluster, the Johnson Creek-Profile Gap shear zone is the westernmost of the major faults and extends through the Edwardsburg, Profile, and Yellow Pine mining districts (fig. 2). In the southern part of the mineral province, the Meadow Creek and Coin Mountain faults east of the Johnson Creek-Profile Gap shear zone also control and bound mineral deposits in the Yellow Pine-Thunder Mountain districts (figs. 2, 3), which are the most recently mined and explored of the province. Faults at both regional and district scale are distinctly imaged by geophysical methods (Anderson and others, in press). The faults are described separately (see sections “Johnson Creek-Profile Gap Shear Zone”, “Meadow Creek Fault Zone”, and “Coin Mountain Fault”)

Yellow Pine and Thunder Mountain Districts Structural Model

Leonard (1985; Leonard and Christian, 1987) presents a structural and mineral deposit model for southeastern districts in which the mineral-hosting major structures (including Johnson Creek-Profile Gap shear zone and Meadow Creek fault system) are interpreted as Eocene “extra-caldera” subsidence features (ring fractures), although those authors acknowledge deposits are not zoned outward from proposed source calderas. That model is abandoned because the mineral-deposit-hosting faults of the Yellow Pine and Profile districts all display evidence of ductile deformation preceding and associated with 79–78 Ma silicification, alteration, and mineralization activities—significantly prior to the about 49 Ma onset of Eocene volcanism. In another model, Cookro and others (1988) interpret the faults as part of a dextral transcurrent system and identify the Meadow Creek fault as a dextral strike-slip fault along which Stibnite area deposits were localized at an extensional jog. However, details from the metasedimentary roof rocks do not support strike-slip offset. Furthermore, observations that specific mineral deposit types formed at different crustal depths (Cooper, 1951) and that younger, epithermal deposit types overprinted older, mesothermal types (Cookro and others, 1988) cannot be explained by strike-slip faulting.

Observed progressions from viscoplastic to brittle structural conditions, and from deep-seated to shallow ore formation conditions in the mineralized rocks at the same locations, are best explained by normal offset on the bounding structural systems for the following reasons:

Facies Belt Reconstruction as Potential Metallogenic Influences

In the central Idaho mineral province, different lines of evidence reveal linkages between country rock facies belts and metals endowments in ore deposits, allowing the regional distribution of both to be assessed. For Mesoproterozoic strata (Lemhi basin), exposures trace northwest from the east-central Idaho gold±cobalt±copper deposits through the Big Creek and Ramey Ridge districts (southeastern cluster), maintain the northwest strike between the Johnson Creek-Profile Gap shear zone and the Smith Saddle fault although the exposure has a north trending extent, and maintain the northwest strike into districts in the Elk City area west of the Bargamin Creek fault (northwestern cluster, figs. 3, 8). The Neoproterozoic–Paleozoic facies belts are parallel to but lie south of (above) the Mesoproterozoic belts. The Neoproterozoic facies belt contains gold-silver±lead-zinc±antimony±tungsten deposits in the Marshall Lake-Resort district but otherwise does not seem to be correlated with mineral deposits. The Paleozoic facies belts of the shelf to slope—which are characterized primarily by carbonate rocks and by lead-zinc deposits in a broad, northwest-striking belt in east-central Idaho—are apparently absent along strike in central Idaho. The stratigraphic and detrital zircon data indicate that parts of the central and western Paleozoic facies belts of south-central Idaho (the deep-water slope, orogenic-basin, and epicratonic-basin strata) are locally present along strike into central Idaho.

At that northwestern extent of central Idaho, the successions may be much thinner, belts much narrowed, and structures complicated but, as with the deep-water slope, orogenic-basin, and epicratonic-basin strata of western south-central Idaho, the central Idaho correlatives are characterized by gold-silver, antimony, and tungsten deposits. From the Yellow Pine and Profiles districts, some parts of these facies belts with their associated metals assemblages trace north to the Edwardsburg district along the Johnson Creek-Profile Gap shear zone and, from where the fault displacement reverses in the Edwardsburg district, rocks that are probably part of this facies belt trace northwest through the Warren and Marshall Lake-Resort districts on the west side of the fault system. The belt apparently terminated at the Salmon River suture as it is not present in the Gospel Peaks roof pendant.

Evidence from deposits in western south-central Idaho indicates that sulfur and metals were sourced from the Pennsylvanian–Permian epicratonic-basin onlap facies in the presence of regional thrust faults that initially juxtaposed those units over the Ordovician–Devonian deep-water facies and that secondarily provided plumbing during successive younger magmatic-hydrothermal events (Hall and others, 1978). Similar genetic linkages seem evident in central Idaho. In addition to equivalent facies and crustal conditions, additional geologic elements are associated with mineral deposits in south-central Idaho. These are northwest-striking synsedimentary basin-deepening structures and Cretaceous thrust faults that juxtaposed facies belts (Hall and others, 1978; Hall, 1985; Hobbs, 1985; Lund and others, 2016). Both of these structural elements probably extend along strike and across the Salmon River transform among the discontinuous exposures into central Idaho (fig. 8), although individual structures would be expected to differ across the transform. Linkages of northwest-striking linear arrays of structures and specific metal endowments with associated facies belts, particularly the belt of gold-antimony-tungsten deposits, support a significant mineral belt from south-central Idaho and diagonally across central Idaho, albeit across the depositional basin and structure transition (transform) at the southern Salmon River Mountains (fig. 8).

The combined lines of evidence about the central Idaho deposits and their host rocks place them at the northwestern extent of a north-northwest striking linear array of tungsten-endowed mineral deposits. Based on the locations of metals-deposit occurrences (but not necessarily on geologic factors), this trend was previously proposed as a “tungsten belt” by Cook (1956; Savage, 1970; the single significant tungsten occurrence east of the identified tungsten belt, the Ima mine [fig. 4; Anderson, 1948; Cook, 1956] is in the area of Paleozoic-basin deepening faults [Lund and others, 2016]). The presence of tungsten endowments in this trend is strengthened by more recent mineral exploration, regional geochemical data, and mineral-deposit research (Peterson, 1984; Cookro, 1985; Hall, 1985; Hobbs, 1985; Cookro and others, 1988; Gammons, 1988; Skipp and others, 1989; Fisher and others, 1992; Link and others, 1995; Chang and Meinert, 2008; Skipp and Kuntz, 2009). However, at the time that the “tungsten belt” was proposed, the metasedimentary rocks across central Idaho were all mapped as Mesoproterozoic (Ross and Forrester, 1947; Bond, 1978). As presented herein, the newer geologic mapping, dating, and detrital zircon constraints allow the identification of discrete stratigraphic belts of discretely different ages and (or) settings. These facies belts directly correspond to the tungsten belt trend as well as to other types of metals belts that cross the central Idaho mineral province (fig. 8).

In the central Idaho mineral province, the combined Mesoproterozoic through Paleozoic facies belts themselves are southwest dipping (the facies packages young to the southwest; fig. 3). Along the regional-scale Johnson Creek-Profile Gap/Bargamin Creek fault system, the facies belts display a northward step in the area of the reversal of movement on the Johnson Creek-Profile Gap fault. This northward separation could have resulted from some transcurrent displacement prior to Late Cretaceous plutonism, but direct evidence for that is absent. Most, if not all, of the separation across the fault system can be attributed to the large amounts of dip-slip displacement of a thick, southwest-tilted set of the several-aged facies belts across the normal faults.

Role of Regional Faults in Mining District Distribution

The mining districts of central Idaho are linked by the nearly 200-km long, broad zone of subparallel, north-northeast-striking, steep normal faults (primarily the Bargamin Creek and Johnson Creek-Profile Gap faults), which mostly had opposite offsets (figs. 2, 3, 4). Without exception, northern mining districts are located west of the large normal faults, but southern districts are located both within and east of the southern faults.

The timing and magnitudes of faulting affected the ages and styles of mineralization in hanging-wall settings and may have influenced whether the fault zones themselves were mineralized or barren. In the southern part of the mineral province, deposits formed during multiple Late Cretaceous to Eocene structural openings and fluxes of mineralizing fluids related to the Johnson Creek-Profile Gap, and Meadow Creek, and Coin Mountain faults; the continued fault activity also preserved the Thunder Mountain cauldron complex, subvolcanic plutonic rocks, and Eocene gold deposits at near-surface crustal levels in the easternmost hanging-wall fault domino.

In contrast, in the northern part of the province, the Bargamin Creek and linked faults are related to significant Late Cretaceous mineralizing activity and only minor Eocene mineralizing activity but are related to significant gold placer development in Miocene–Holocene fault basins; all placers are in hanging-wall settings. Along with their roles in siting mineral deposits, the regional-scale, steep, Johnson Creek-Profile Gap and Bargamin Creek normal-fault systems controlled the regional positions for stratigraphic facies and mineral belts of central Idaho. Furthermore, these faults controlled the uplift and erosion of deep rocks versus the preservation of shallow crustal levels and, thus, controlled the presence or absence of Late Cretaceous plutonic roof zones and related ore deposits. The offsets of facies belts, localization of associated mineral deposits, and controls of exposed crustal depths across this regional fault system provide guidance for future mineral resource assessments and for exploration of specific metal endowments, whether in igneous or metasedimentary rocks.

The distribution of placer districts with respect to the faults is dissimilar between the two district clusters (compare details of the northwestern and southeastern mining districts in table 2). Large and historically important gold placer districts are a significant characteristic of the hanging wall west of the northern part of the normal-fault system. The gold placers in the Elk City, Newsome-Tenmile, Orogrande, Dixie, Florence, Marshall Lake-Resort, and Warren districts, plus heavy-mineral placers in the Warren and Marshall Lake-Resort districts, all lie west of the down-to-the-west Bargamin Creek-Dixie Meadows-Blanco Creek-Smith Saddle-South Fork faults and the northern part of Johnson Creek-Profile Gap shear zone. Although shallow, near-surface, mostly Eocene deposits are present in the eastern Yellow Pine and Thunder Mountain districts, there are no significant placers associated with the hanging-wall (eastern) side of the down-to-the-east Johnson Creek-Profile Gap, Meadow Creek, and Coin Mountain faults. Farther east, several gold placer districts were historically significant, and these are along a northeast-striking swath near the Salmon River transform. The placer districts in the northwestern district cluster are locally as old as Pleistocene but mostly Holocene. The ages of the placer deposits indicate that structural activity on the northern part of the normal-fault system, including uplift with erosion on one side and basin formation with deposition on the other side, continued into the Holocene. The absence of placer basins east of the southern normal-fault system may reflect that significant fault activity in the southeastern cluster of districts mostly ended in the Eocene.