Roles of Regional Structures and Country-Rock Facies in Defining Mineral Belts in Central Idaho Mineral Province with Detail for Yellow Pine and Thunder Mountain Mining Districts

Professional Paper 1884
Mineral Resources Program
By: , and 



A first draft of this manuscript was reviewed by K.V. Evans who helped clarify and organize ideas. His continued interest in this region is gratefully acknowledged. A second draft was reviewed by R.S. Lewis. Comments by both reviewers improved the final manuscript. Interactions with and assistance from many personnel of the U.S. Department of Agriculture Forest Service in the Red River and Grangeville Ranger Districts of the Nez Perce National Forest and in the Krassel District of the Payette National Forest made it possible to complete the field aspects of this research through the many years represented by this study.


The central Idaho metallogenic province hosts numerous mineral deposit types. These include Late Cretaceous precious-polymetallic vein deposits, amagmatic Paleocene–Eocene breccia-hosted gold-tungsten-antimony deposits, and Eocene mercury deposits in metasedimentary roof pendants and in Late Cretaceous granitoids. Hot-springs gold deposits in Eocene volcanic rocks are also included in the central Idaho province. New sensitive high resolution ion microprobe (SHRIMP) uranium-lead (U-Pb) ages for igneous rocks and for detrital zircon analyses of metasedimentary roof rocks along with geologic mapping clarify the geologic framework of the mineral deposits. This framework includes (1) structural controls for regional distribution of mining districts, (2) progressive structural development of individual districts, (3) regional sedimentary facies and their control of metals associations resulting in regional belts, and (4) influences of the several regional magmatic events.

In central Idaho, 15 mining districts form two clusters that are grouped about a 200-kilometer (km) long system of normal faults. The northwestern cluster is in the regional hanging wall west of large, west-side-down faults, and the mineral deposits are located along smaller faults and fractures that cut the regional hanging wall. The southeastern cluster is in the regional hanging wall east of a linked large east-side-down fault and along and controlled by related hanging wall faults. At the southern extent of the regional fault system, the Yellow Pine-Thunder Mountain districts span a nearly 24-km-wide, east-tilted crustal block of normal-fault dominoes, exposing original crustal depths from 5 to 10 km deep on the west in the Late Cretaceous to shallow-surface depths on the east in the Eocene.

Ore deposition in the northwestern district cluster was primarily Late Cretaceous and related to Idaho batholith plutons with only a single deposit related to a small Eocene intrusion; in the southeastern cluster, most deposits were initiated in the Late Cretaceous but with varying manifestations of overprinted Eocene mineralization activity. In the Yellow Pine-Thunder Mountain districts at the southern extent of the southern cluster, several mineralizing pulses occurred during hanging-wall collapse, such that (1) early deposits were multiply overprinted and (2) deposit depths, ages, and structural characteristics change progressively eastward. Originally deep-seated western Yellow Pine district deposits are Late Cretaceous viscoplastic mesothermal veins overprinted by Paleocene and Eocene breccia-hosted epithermal deposits. Central Yellow Pine district deposits contain early deeper vein systems but are primarily Paleocene and Eocene breccia-hosted epithermal deposits in Late Cretaceous plutonic rocks and Proterozoic–Paleozoic roof pendant rocks. Eastern district deposits are Eocene hot-springs-related deposits in the roof pendant. Thunder Mountain deposits farthest east are near-surface hot-springs deposits in Eocene volcanic and volcaniclastic rocks that overlie buried Cretaceous and older roof pendant rocks.

The mining district clusters are sited across several northwest-striking paleostratigraphic belts that are exposed in roof pendants and are offset by the regional normal fault system. A northeastern belt is Mesoproterozoic strata associated with gold-silver-copper±cobalt deposits. A central belt of Neoproterozoic rocks is not associated with mineral deposits in the central Idaho mineral province. A southwestern belt composed of probable Paleozoic deep-water miogeoclinal slope rocks and late Paleozoic epicratonic basinal rocks is thin and narrowly exposed but associated with gold-silver-antimony-tungsten±mercury deposits. These metasedimentary rocks (and their metal associations) are parts of regional mineral belts in which metal endowments are related to particular sedimentary facies belts and their Cretaceous thrust-fault juxtaposition and where these features have proximity to Late Cretaceous or Eocene igneous rocks. Offset and preservation or erosional stripping of these facies belts, thrust plates, igneous settings, and the associated regional mineral belts were controlled by the sense and magnitude of displacements across the regional normal-fault system.


The many mining districts of central Idaho compose a richly mineralized province with significant historic production and a broad variety of commodities (fig. 1, table 1). Diverse proposed mineralizing fluid sources include Cretaceous plutons, Eocene igneous rocks, Mesoproterozoic plutons, metamorphic events, or synsedimentary processes. Recent studies also suggest that cryptic basement character and an overprinted rift system may have influenced development of the metallogenic province. In this complicated geologic framework containing mineral deposits of such broad characteristics, widely accepted models for deposit genesis are generally lacking as is agreement about which geologic elements are critical to ultimately determining genesis.

Location of study area in northern Rocky Mountains.
Figure 1.

Index maps showing location of study area in northern Rocky Mountains. A, Map showing Neoproterozoic to early Paleozoic Laurentian rift features, location of miogeoclinal deposits, and related mineral deposits (Lund, 2008). B, Regional map showing cryptic basement domains from Lund and others (2015), distribution of epigenetic deposits from Klein and Sims (2007), and geographic areas discussed in text.

Table 1.    

Recorded metals production for lode and placer mining districts, central Idaho.

[--, not reported]

District name Lode production Placer production
aGold aSilver bCopper bLead bZinc bAntimony b Tungsten trioxide (WO3) aGold
Northwestern districts
Newsome1      1,236 298 -- -- -- -- -- 97,000
Tenmile1      18,866 12,906 4,235 10,302 707 -- -- 128,600
Orogrande1      20,827 11,698 2,871 20,941 -- -- present 32,000
Elk City1      27,646 11,315 750 847 400 -- -- --
Buffalo Hump1      33,534 33,579 11,118 49,253 -- -- -- --
Dixie1      2,408 1,534 927 2,483 -- -- -- 40,000–75,000
Florence1      738 615 -- -- -- -- -- 1,000,000
Marshall Lake and Resort1      21,102 13,204 13,204 10,302 707 present present 225,600
Warren1      58,234 123,337 12,947 26,585 294 -- present 906,000
Southeastern districts
Big Creek2,4 present present present present -- -- -- --
Ramey Ridge2 7,115 16,112 125,098 17,656 -- -- present --
Edwardsburg2,4 1,568 351 present -- -- -- present 3,500 estimated
Profile2,4 present present present present -- present -- --
Yellow Pine3 549,778 1,687,914 -- -- -- 75,618,556 17,306,110 --
Thunder Mountain3 117,993 153,681 -- -- -- -- -- --
Table 1.    Recorded metals production for lode and placer mining districts, central Idaho.

Lund and others, 1986; Lund and Esparza, 1990.


Cater and others, 1973.


Bookstrom and others, 1998.


Big Creek had a total production of $265,379; Edwardsburg district had $97,485; Profile district had $7,993 (Cater and others, 1973).


Troy ounces.



The present study endeavors to clarify settings of the diverse mineral deposits of the central Idaho mineral province by presenting and evaluating details from the districts, including (1) igneous rock relations to mineralization; (2) compositions, ages, provenances, and facies-belt origins for metasedimentary roof pendants using geologic mapping and uranium-lead (U-Pb) geochronology; (3) structural and kinematic histories of mineral deposit host rocks, crustal depths at times of mineralization, and the temporal relations between deformation and mineralization; and (4) the role of structural history as a factor in deposit formation as well as in deposit preservation and exposure. Yellow Pine and Thunder Mountain districts at the southern extent of the mineral province are the most complex and recently explored parts of the province. New details from these southern districts employ sensitive high resolution ion microprobe (SHRIMP) U-Pb ages of and emplacement settings for Late Cretaceous Idaho batholith phases and new structural mapping to evaluate crustal evolution and mineral deposit formation.

The goals of the study are to investigate potential common genetic factors among districts, test the verity and origin of a proposed regional tungsten belt and the geographic distributions of other metal endowments in deposits regionally, and advance a geologic framework useful for resource assessment of this mineral-rich province.

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.

Map showing major faults and northwestern cluster and southeastern cluster mining
                        districts of central Idaho.
Figure 2.

Mining districts and major faults of central Idaho, after Ross (1936) and Gustafson (1987). Light gray indicates those districts are referred to as northwestern cluster; dark gray indicates districts of southeastern cluster.

Generalized geologic map of central Idaho.
Figure 3.

Generalized geologic map of central Idaho. Southern half modified from Lund (2004) and northern half compiled from new mapping, present study. Age data for metasedimentary rocks are in table 3 and plutonic samples are in table 4. (Ma, million years)

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.

Characteristics of Central Idaho Mining Districts

Fifteen mining districts were organized historically across much of central Idaho as shown on figure 2. Most of the placer districts were discovered in the early 1860s (Idaho State Historical Society, 1985). Much-publicized gold rushes were initially associated with successive discoveries of gold placer deposits in the different districts. Discovery of lode deposits lagged in the various mining districts, but most activities related to lode mining commenced from the late 1890s to about 1915 (Idaho State Historical Society, 1985). Because lode mining and production was more difficult in this era, associated lode-gold rushes were smaller and generally occurred several decades after initial gold rushes. Production details for the districts are in table 1.

The mining districts group into northwestern and southeastern clusters which are distributed across a regional-scale zone of normal faults (fig. 2). The geologic setting of each district is complex, and the variety of metal commodities in different districts is extensive. To ascertain possible patterns among districts across the province, the geologic and mineral-deposit characteristics and settings for each district are tabulated in table 2 and described following.

Table 2.    

Characteristics of mineral deposits for mining districts, central Idaho.

[m, meters; Ma, millions of years; Ar, argon; km, kilometers; U-Pb, uranium-lead; SHRIMP, sensitive high resolution ion microprobe; cm, centimeters; Re-Os, rhenium-osmium; --, not applicable]

District name Metasedimentary host associations Late Cretaceous plutonic host associations Relation to Eocene igneous rocks Deposit characteristics Ore minerals (see table 1 for commodities and production) Age(s) of mineralization Sources of data
Northwestern mining districts
Newsome and Tenmile Mesoproterozoic, fine-grained, biotite-feldspar-quartz gneiss of Lemhi Group and local younger carbonate and quartzite gneiss Minor Late Cretaceous granite and migmatite -- Quartz veinlets and veins in silicified, faulted wall rock. Post-vein brecciation, gouge, and offsets of veins Pyrite, chalcopyrite, arsenopyrite, galena, covellite, chalcocite, native gold, gold telluride undetermined Thompson and Ballard, 1924; Shenon and Reed, 1934; Reid, 1960
Placers in Miocene–Holocene fault graben, as basin fill as much as 100 m above river level and active stream channels. As much as 25 m thick. Some eluvial placers Gold undetermined Shenon and Reed, 1934; Capps, 1941
Orogrande Mesoproterozoic, fine-grained, biotite-feldspar-quartz gneiss of Lemhi Group, Mesoproterozoic augen gneiss, and Mesoproterozoic feldspar-quartz metasandstone and sillimanite-biotite schist Minor Late Cretaceous granite and granodiorite Few post-ore andesite, porphyritic dacite dikes Pinch and swell quartz veins (as much as 2 m wide) and disseminations in silicified country rock in shear zones. Some late-stage brecciation Pyrite, chalcopyrite, galena, tetrahedrite, sphalerite, native gold, late-stage covellite undetermined Thompson and Ballard, 1924; Shenon and Reed, 1934; present study
Eocene granite porphyry stock in north-striking Late Cretaceous shear zone Ore minerals as disseminations and quartz stringers in Eocene granite porphyry Pyrite, chalcopyrite, galena, molybdenite, tetradymite, free gold, gold telluride, wolframite, scheelite Eocene Thompson and Ballard, 1924; Shenon and Reed, 1934; Lund and Esparza, 1990; present study
-- Placers in Miocene-Holocene fault graben occupied by active stream channel Gold Miocene to Holocene Shenon and Reed, 1934; present study
Elk City Mesoproterozoic, fine-grained, biotite-feldspar-quartz gneiss of Lemhi Group and Mesoproterozoic augen gneiss Minor Late Cretaceous granite and migmatite Small post-vein porphyritic andesite dikes noted along faults in several mines Quartz veins and lenses in fracture zones. As much as 6 m wide and 100 m long. Subsequent fault offsets, parallel gouge, and brecciation Pyrite, sphalerite, chalcopyrite, galena, native gold undetermined Shenon and Reed, 1934; present study
Placers in Miocene–Holocene clay, sand, gravel, cobble deposits (as much as 35 m thick) located in fault graben and active stream channels Gold undetermined Reed, 1934; Shenon and Reed, 1934; Capps, 1941; Reid, 1959, 1960; present study
Buffalo Hump Mesoproterozoic calc-silicate gneiss of the Yellowjacket-Hoodoo package structurally over feldspar-quartz gneiss of Lemhi Group Late Cretaceous 84 Ma tonalite (40Ar/39Ar on hornblende), granodiorite, and 74 Ma muscovite-biotite granite (40Ar/39Ar on muscovite) -- North- to north-northeast-striking, comb and vuggy textured, multi-phase quartz veins, as much as 4 m wide, some for more than 0.5 km long, in fractures. Some shorter quartz offshoots Pyrite, sphalerite, chalcopyrite, and tetrahedrite, galena, arsenopyrite, molybdenite, native gold, telluride, argentite, stibnite, covellite 71–72 Ma (40Ar/39Ar on white mica) Thompson and Ballard, 1924; Beckwith, 1928; Shenon and Reed, 1934; Chauvot, 1986; Lund and others, 1986; Lund and Esparza, 1990; present study
Dixie Mesoproterozoic, fine-grained, biotite-feldspar-quartz gneiss of Lemhi Group and Mesoproterozoic augen gneiss Late Cretaceous muscovite-biotite granodiorite and granite Few porphyritic rhyodacite dikes in district Quartz veinlets localized, few large veins. Placer deposits in glacial outwash and modern stream channels. Central to a Miocene(?) graben structure Pyrite, galena, sphalerite, chalcopyrite, native gold undetermined Capps, 1939, 1941; present study
Florence Late Cretaceous tonalite and granodiorite -- Quartz veinlets, short quartz veins, and silicified country rock Native gold, telluride, cerargyrite, silver sulfides undetermined Reed, 1939; Capps, 1941
Placers in older graben-fill lignite-peat, clay, sand, gravel, and cobble deposits, eluvial settings, modern stream channels Gold, uranium undetermined
Marshall Lake and Resort Neoproterozoic metavolcanic rocks and probably younger quartzite, marble, andalusite-biotite schist. Mesoproterozoic biotite-feldspar quartz gneiss Late Cretaceous muscovite-biotite granodiorite and granite -- West-northwest striking, pinch and swell, quartz veins, as wide as 1.2 m Gold, cerargyrite, galena, sphalerite, tetrahedrite, magnetite, scheelite 74.5 Ma (40Ar/39Ar on white mica) Lorain, 1938; Capps, 1940, 1941; Savage, 1961; May, 1984; Lund and Esparza, 1990; Lund, 2004
Pleistocene-Holocene glacial moraines and outwash reworked by modern streams. Deposits overlie tilted Miocene basin-fill in northwest-striking fault graben Gold, corundum, monazite-bearing black sands, garnet, cinnabar, rutile, zircon, magnetite, illmenite undetermined
Warren Neoproterozoic quartzite, feldspar-quartz metasandstone, sillimanite-biotite schist roof rocks above plutonic rocks at district margins; eroded over district Late Cretaceous porphyritic muscovite-biotite granodiorite 82.4 Ma (U-Pb SHRIMP, this study) -- East-striking, lenticular quartz veins, generally 30–60 cm wide, persisting for as much as 1 km in length. Located several hundred meters below projected base of Neoproterozoic roof rocks. In fractures that were reactivated after vein formation Pyrite, sphalerite, tetrahedrite, arsenopyrite, stibnite, pyrargyrite, native gold, argentite, native silver, mottramite, scheelite, mimetite undetermined Reed, 1937; Lorain, 1938; Lorain and Metzger, 1938; Capps, 1941; Lund and Esparza, 1990; Lund, 2004; present study
Placer deposits in successive Pleistocene-Holocene glacial moraines, interglacial outwash, modern stream channels. Overlying tilted Miocene basin fill Gold, monazite-bearing black sands, zircon, garnet undetermined
Southeastern mining districts
Big Creek and Ramey Ridge Mesoproterozoic Lemhi Group locally overlain by Yellowjacket-Hoodoo package klippen -- Eocene volcanic rocks depositional on top; potentially underlain by Eocene granite. Eocene porphyritic dikes cut veins East- or north-northeast-striking, lenticular quartz veins, as much a 1 m wide, in shear zones Pyrite, chalcopyrite, chalcocite, galena, and magnetite; geochemical anomalies are gold, silver, copper, iron, and cobalt undetermined Shenon and Ross, 1936; Cater and others, 1973; Lund, 2004; Lund and others, 2010
Neoproterozoic syenite-diorite complex Northwest-striking, lenticular quartz veins in shear zone, persistent for more than 650 m Pyrite, chalcopyrite, pyrrhotite, galena, local stibnite late in vein formation 79–65 Ma (40Ar/39Ar on white mica)
Edwardsburg Mesoproterozoic Lemhi Group structurally overlain by Yellowjacket-Hoodoo package Silicified, sericitized, sheared Late Cretaceous muscovite-biotite granodiorite and granite Eocene granite porphyry intruded and underlies; Eocene porphyritic rhyolite and rhyodacite dikes parallel to and cutting veins North-northeast-striking quartz veins in shear zones; at eastern margin of Johnson Creek-Profile Gap shear zone Pyrite, chalcopyrite, sphalerite, galena, tetrahedrite, gold undetermined Shenon and Ross, 1936; Cater and others, 1973; Lund, 2004
Neoproterozoic and Paleozoic quartzite, marble, and sillimanite-biotite schist North-northeast-striking quartz veins and silicified disseminations in sheared and altered Late Cretaceous granitic rocks and metasedimentary screens. Early phases plastically deformed quartz, later phases associated with tensional openings and brecciation. Located along Johnson Creek-Profile Gap shear zone Pyrite, arsenopyrite, stibnite, huebnerite, sphalerite, chalcopyrite, galena, tetrahedrite, pyrrhotite, telluride, scheelite, cinnabar, boulangerite, covellite Early phases 79–73 Ma; younger phases 69–67 Ma; late phase and Ar loss 52–47 Ma (40Ar/39Ar) Shenon and Ross, 1936; Gammons and others, 1985; Gammons, 1988; Lund, 2004
Profile Neoproterozoic and Paleozoic quartzite, marble, and sillimanite-biotite schist; at south, Neoproterozoic–Mesoproterozoic feldspar-quartz metasandstone Silicified, sericitized, sheared Late Cretaceous muscovite-biotite granodiorite and granite Eocene porphyritic rhyolite and rhyodacite dikes parallel to and cutting veins in northern district; sparse andesite dikes in southern district North-northeast-striking quartz veins and silicified stockwork, disseminations in sheared and altered granitic rocks and metasedimentary screens. Early tabular quartz veins, plastically deformed; later calcite filling in tensional openings and breccia. Located within Johnson Creek-Profile Gap shear zone Pyrite, arsenopyrite, pyrrhotite, sphalerite, molybdenite, gold; late scheelite undetermined Shenon and Ross, 1936; Peterson, 1984; Gammons, 1988; Lund, 2004
Yellow Pine Neoproterozoic and Paleozoic quartzite, marble, and sillimanite-biotite schist in central and east Neoproterozoic–Mesoproterozic feldspar-quartz metasandstone and minor calc-silicate marble of unknown age in west Silicified, sericitized, sheared Late Cretaceous tonalite and granodiorite Few fine-grained andesite to rhyodacite dikes, some altered West: early tabular quartz veins, plastically deformed; late quartz and calcite filling in tensional openings and breccia; within Johnson Creek-Profile Gap shear zone West: early pyrite, arsenopyrite, pyrrhotite, molybdenite, sphalerite, gold; intermediate age scheelite undetermined Larsen and Livingston, 1921; Schrader and Ross, 1926; Shenon and Ross, 1936; White, 1940; Cooper, 1951; Lasmanis, 1981; Lewis, 1984; Peterson, 1984; Cookro, 1985; Cookro and others, 1988; Gammons, 1988; Bookstrom and others, 1998; Lund, 2004; table 4-3 of Gillerman and others, 2019
Central (Stibnite area): quartz veins and silicified stockwork on north-striking fault segment in sheared and altered granitic rocks primarily precious-polymetallic deposits. Veins, disseminations, and replacements in sheared and altered granitic rocks and metasedimentary screens along fault conjugates and splays are gold rich. Intermediate age stockwork and disseminated antimony in hanging wall damage zone. Late-stage, breccia-related quartz-scheelite veins central to main structural pathways in hanging wall damage zone. Within damage zone of Meadow Creek fault and at junction of Meadow Creek and Sugar Creek faults, including along main fault, splay faults, and conjugate faults Central (Stibnite area): early pyrite, arsenopyrite, pyrrhotite, molybdenite, sphalerite, gold; intermediate age stibnite, scheelite; late stibnite, scheelite, cinnabar, silver sulfosalts Early phases at 86 Ma (Re-Os), main veins at 77.9 Ma, tungsten and antimony at about 57 and 51–47 Ma (40Ar/39Ar, U-Pb)
East: primarily stratabound parallel to strike of marble units near contacts with quartzite, few in north-northeast-striking gouge zones. In silicified zones, jasperoids, chalcedonic fumeroles; interpreted as shallow, epithermal systems. Within Stibnite roof pendant across 3 km in hanging wall of Meadow Creek fault, footwall of Thunder Mountain cauldron complex boundary fault East: cinnabar with limited pyrite, realgar, orpiment, scheelite, stibnite undetermined
Thunder Mountain Neoproterozoic and Paleozoic quartzite, marble, and sillimanite-biotite schist roof pendants underlie Thunder Mountain cauldron complex Late Cretaceous granodiorite underlies Thunder Mountain cauldron complex Eocene rhyolite welded and ash-flow tuffs 46.7 Ma and lacustrine beds with lignite and plant debris in north-northeast-striking caldera-collapse graben Hot-spring-related, disseminated; in shallow, porosity- and composition-controlled layers and structures. Limited to units in young units within caldera-collapse graben. Graben-related, steeply dipping, silicified and argillized fault zones served as conduits for mineralizing fluid into permeable non-welded tuff layers and volcaniclastic debris-flow rocks Native gold, electrum, pyrite 43 Ma (40Ar/39Ar) Umpleby and Livingston, 1920; Ross, 1933; Cater and others, 1973; Shannon and Reynolds, 1975; Adams, 1985; Parsley, 1997; D. John, U.S. Geological Survey, oral commun.
Table 2.    Characteristics of mineral deposits for mining districts, central Idaho.

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).

Metasedimentary Country-Rock Characteristics

Although epigenetic deposits in the central Idaho mineral province are of both Cretaceous and Eocene ages and associated with hosting or adjacent igneous rocks (table 2), vast expanses of identical igneous rocks in the province are not associated with mineral deposits. Geochemical signatures and different metal assemblages in different districts suggest metal endowments patterns in the mineral province that are not explained by magmatic characteristics. The inconsistent relations of igneous rocks to deposits suggest that pre-magmatic country rocks need to be considered as influencing factors in the location of deposits in the metallogenic province.

Although country rocks across the province were shown as Mesoproterozoic in two generations of state geologic maps (Ross and Forrester, 1947; Bond, 1978) and in most descriptions of mining districts, additional geologic mapping and the ability to date some units demonstrated that younger stratigraphic packages underlie large parts of the area (Lund, 1984; Lund and Esparza, 1990; Lund and others, 2003; Lund, 2004), and this change is represented on the recent state map (Lewis and others, 2012). Modern data for central Idaho country-rock packages—including their compositions, ages, provenances, facies trends, and metal endowments—help identify their basin settings and facies origins and test their potential influences on metal endowments and, consequently, on the location of mineral belts.

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).

Map showing the location of epigenetic mineral deposits in south-central to central
Figure 4.

Generalized geologic map of south-central to central Idaho (modified from Lewis and others, 2012; Lund and others, 2016) showing location of epigenetic mineral deposits (modified from Klein and Sims, 2007). Salmon River transform is after the Salmon River lineament of Hobbs and others (1991). (Au, gold; Co, cobalt; Cu, copper; Hg, mercury; Sb, antimony; Mo, molybdenum)

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. 5DI).

Detrital zircon relative probability plots for samples of probable Neoproterozoic
                              and Paleozoic metasedimentary rocks of undetermined ages in central Idaho and of unmetamorphosed
                              rocks of known ages in south-central Idaho.
Figure 5.

Detrital zircon relative probability plots for samples of probable Neoproterozoic and Paleozoic metasedimentary rocks of undetermined ages in central Idaho and of unmetamorphosed rocks of known ages in south-central Idaho. A, Relative probability plots for five samples of metasedimentary rocks in the central Idaho mineral province from the present study. Placer lake ridge siltite (04KL102a, “Placer lake” is informal location name used in Cater and others, 1973, fig. 73) from below dated volcanic rocks of Edwardsburg Formation (685 million years [Ma], Lund and others, 2003; Lund, 2004) from Big Creek roof pendant; Whiskey Creek metasandstone gneiss (20KL047), having angular upper contact with probable Paleozoic units, from central Stibnite roof pendant; McFadden Point quartzite (04KL107) structurally above Neoproterozoic units from Big Creek roof pendant (Lund and others, 2003); Monumental Summit metamorphosed quartz-pebble quartzite (08KL058) from southern Stibnite roof pendant; Jackass Gulch metamorphosed quartz-pebble quartzite (08KL005) from roof pendant along Middle Fork of the Salmon River 34 km east-southeast of Stibnite and without context with other strata. Locations shown on figure 3. Analytical data is presented in table 3 and available in Aleinikoff and others (2023). B, Relative probability plots for units in the Stibnite roof pendant including a sample from Missouri Ridge Formation of Lund (2004; named in Lund and others, 2003) in the probable youngest facies in the northern part of the roof pendant (14DS12, Stewart and others, 2016; Isakson, 2017) and two samples from the upper and lower (denoted as 12DS19 and QPC, Stewart and others, 2016) parts of the quartzite and conglomerate unit of Monumental Summit that composes about a third of the roof pendant (Lund, 2004; Stewart and others, 2016); these quartzite and conglomerate samples from the Stibnite roof pendant (Stewart and others, 2016; Isakson, 2017) are from the same map unit as Monumental Summit sample in figure 5A. C, Relative probability plots for biotite-feldspar-quartz schist that is the structurally lowest unit in the southern Stibnite roof pendant and for a quartzite unit from a thrust sliver (Lund, 2004) internal to the southern Stibnite roof pendant (14RL011a and 12DS14, respectively, from Stewart and others, 2016; Isakson, 2017). D, Relative probability plots for samples from the southwestern part of southern Salmon River Mountains area (fig. 1B) including, from west to east: Ordovician to Mississippian Salmon River assemblage at the northern extent of recognized flysch basin facies rocks (Beranek and others, 2016) in a thrust plate over three quartz-rich sandstones (8NK15, 4TA09, 3TA09, Krohe and others, 2020) which are located from structurally shallower to deeper in the west limb of Clayton antiform and have ages that may be Cambrian to Ordovician, Cambrian, and late Neoproterozoic to Cambrian, respectively (Hobbs and others, 1991). E, Relative probability plots from southeastern part of Salmon River Mountains area, including units on west and east limbs of Bayhorse antiform (DTB17-11 and DTB17-19, respectively, Brennan, 2020) of Ordovician or older age in the thrust sheet overlying the fold core (Hobbs and others, 1991) and in the core of Bayhorse antiform (10NK15, Krohe and others, 2020) from Ordovician slate overlying Ordovician fossil-bearing rocks (Hobbs and others, 1991). F, Relative probability plots for the epicratonic basin facies, Wood River Formation of the Pennsylvanian–Permian Sun Valley Group from Link and others (2014). G, Relative probability plots for deep-water facies, Devonian Milligen Formation from Beranek and others (2016). H, Relative probability plot for flysch basin facies, Mississippian Copper Basin Group from Beranek and others (2016). I, Relative probability plots for shelf and slope facies, Ordovician Kinnikinic Quartzite and Devonian Jefferson Formation from Beranek and others (2016).

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]

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

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. 5AC). 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. 5AC to FI and DE). 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.

Igneous Events in Relation to Crustal and Deposit Settings

Four magmatic episodes occurred in the central Idaho mineral province, in the Mesoproterozoic, Neoproterozoic, Late Cretaceous, and Eocene. Each magmatic episode provides information on the crustal setting at the time of magma genesis and on a number of factors including intrusive histories and rock fabrics. Present settings provide details of upper-crustal structural evolution. All of these were potential drivers of mineralizing systems.

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).

Sensitive high resolution ion microprobe uranium-lead results showing representative
                           images of dated zircon and SHRIMP isotopic data.
Figure 6A–L.

Sensitive high resolution ion microprobe (SHRIMP) uranium-lead (U-Pb) results showing representative images of dated zircon and SHRIMP isotopic data. For each sample, representative cathodoluminescence images of three grains are shown, followed by Concordia plots showing U-Pb analyses of dated detrital zircon (two-sigma error ellipses) and weighted averages plots (two-sigma error bars). White-filled error ellipses and error bars indicate data used for age calculations; gray-filled ellipses and error bars indicate data excluded from age calculations. Data are available in Aleinikoff and others (2023). (µm, micrometers; Ma, million years; MSWD, mean squared weighted deviation)

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]

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