Orogenic link ∼41°N–46°N: Collisional mountain building and basin closure in the Cordillera of western North America

Tectonic settings characterized by steady-state subduction are periodically disrupted by mountain-building events associated with arc-arc or arc-continent collision (Dewey and Bird, 1970; Howell et al., 1985; Clift et al., 2003; accretionary orogens of Cawood et al., 2009). In either case, collision-related orogenesis initiates with the attempted subduction of a bathymetric high (oceanic arc or basaltic plateau; Livaccari et al., 1981; Cloos, 1993; Liu and Currie, 2016) and continues until plate motions stabilize, e.g., subduction outboard of an accreted arc terrane (Hamilton, 1988; Moresi et al., 2014, and references therein). Andean-type orogens involving arc-continent collision over tens of millions of years (e.g., central North American Cordillera; Fig. 1) typically record pronounced crustal thickening, high-pressure silicic magmatism, and pervasive ductile deformation of regional extent (e.g., Oregon/Idaho Blue Mountains Province of Silberling et al., 1984; Hamilton, 1969a; Zen, 1985; Avé Lallemant, 1995; Žák et al., 2015). Late Mesozoic tectonic activity along/across the arc-continent boundary in west-central Idaho (Riggins region; Fig. 2) was contemporaneous with oblique subduction of the Farallon plate under ancestral North America (McKenzie and Morgan, 1969; Engebretson et al., 1985; Giorgis and Tikoff, 2004).

Detailed structural mapping and mineral age dating in the Riggins region (∼45°30′00″N) support a long-lived history of penetrative deformation, coeval metamorphism, and intermittent calc-alkaline magmatism (Snee et al., 1995; Gray et al., 2012; McKay et al., 2017). Tectonic activity is recorded across the Salmon River suture zone (SRSZ), which in our study extends from the northeastern Seven Devils Mountains in the west (∼116°30′00″W; Heavens Gate fault of Gray and Oldow, 2005) to the Late Cretaceous–Paleocene Idaho batholith in the east (∼116°00′00″W; western border zone of Taubeneck, 1971). Across the SRSZ, linear-planar (LS) tectonite fabrics are developed in accreted arc assemblages of the eastern Wallowa terrane (late Paleozoic–Mesozoic Riggins and Seven Devils Groups: Hamilton, 1963a; Vallier, 1977), Laurentian metasedimentary units (late Precambrian to early Paleozoic Belt and/or Windermere Supergroups; Lund et al., 2003), and intervening calc-alkaline intrusive complexes (Permian–Cretaceous plutons; Kauffman et al., 2014). Exceptional exposure and tectonic context allow for the assessment of disparate rock assemblages metamorphosed and deformed under mid- to upper-crustal conditions (e.g., Zen, 1985; Selverstone et al., 1992).

Despite its tectonic significance (cf. Brown and Ryan, 2011), few have attempted to map polyphase structures across this collisional orogen (Onasch, 1977; Blake, 1991) or relate regional tectonism (SRSZ) to the greater Cordillera (e.g., “tectonic escape” hypothesis of Wernicke and Klepacki [1988], “hit-and-run” Laramide model of Maxson and Tikoff [1996]). Since demarcation of the arc-continent boundary (Armstrong et al., 1977), studies combining structural geology and geochronology have focused on Cretaceous plutons exposed along the boundary (Manduca et al., 1993; Giorgis et al., 2008; Benford et al., 2010; Braudy et al., 2017; Montz and Kruckenberg, 2017), with little attention to country rock units involved in the contractional orogen. As a consequence, timing relations are inferred between structures recorded in accreted island-arc assemblages (F_1–5 fold elements of Onasch, 1977, 1987), autochthonous continental metasedimentary rocks (S_1r, L_m-1 fabrics; Blake et al., 2009, 2016), and intervening metaplutonic complexes (D_1–4 events; Manduca et al., 1993; “overprinting” of Giorgis et al., 2008; Gray, 2015). Nevertheless, Cordilleran-scale tectonic syntheses have been proposed to explain (1) arc-continent collision/terrane accretion (Selverstone et al., 1992; Lund et al., 2008), (2) postaccretion margin-parallel terrane translation (McClelland et al., 2000; Tikoff et al., 2001; Giorgis et al., 2005; Lewis et al., 2014; Schmidt et al., 2016a), and/or (3) collapse of a North American fringing volcanic arc–back-arc basin assemblage (Gray and Oldow, 2005). More recent workers attribute intense tectonic activity in west-central Idaho to hinterland evolution of the Sevier orogeny (LaMaskin et al., 2015).

In light of previous modeling, we have assessed the spatial distribution and development of mesoscopic structures in west-central Idaho. New across-strike structural mapping (scale = 1:24,000; 70 localities) and mineral crystallization ages (3 U-Pb zircon, 1 Lu-Hf garnet) refine the space-time evolution of LS tectonites across the arc-continent boundary (ca. 136–91 Ma fabric development). Given the coeval nature of east-west contraction in the Sevier fold-and-thrust belt (SFTB), midcrustal metamorphic tectonites (SRSZ) evolved together with more eastern structures of the SFTB: e.g., Paris, Willard, Basin-Elba, Meade, and Medicine Lodge thrust systems (DeCelles, 2004) and fabric elements along the Lewis and Clark Line (Sears et al., 2004). In this context, we propose that late Mesozoic terrane accretion and northward translation in the Cordilleran hinterland (∼41°N–46°N, present coordinates) drove shortening eastward into the continental interior (∼116°W–113°W). This history of kinematically linked hinterland-foreland contraction (SRSZ–SFTB orogens) followed Middle Jurassic collapse of late Paleozoic fringing volcanic arc assemblages, e.g., Intermontane composite terrane belt of British Columbia, Canada (Monger et al., 1982; cf. Martini et al., 2013). As an alternative to conventional modeling (noncollisional mountain building of Dewey and Bird, 1970; Burchfiel and Davis, 1975; Burchfiel, 1980; Price and Carmichael, 1986; Edelman, 1992; Taylor et al., 2000; Hyndman et al., 2005; Wells et al., 2008), we offer a collision-related tectonic origin and contractional evolution for central portions of the SFTB (Gray, 2016). Our hypothesis is compatible with cross-orogen structural, geophysical, and geochronological data derived from and/or applied to the North American Cordillera (Potter et al., 1986; Allmendinger et al., 1987; Oldow et al., 1989, 1990; Brown et al., 1992).

Long-standing convergent plate boundaries may experience significant periods of intense metamorphism and deformation leading to widespread contractional strains, particularly when the angle of convergence is high (e.g., Dewey, 1980; Tikoff and Teyssier, 1994). This type of margin characterized ancestral western North America, where eastward subduction of the Farallon plate persisted between at least Late Jurassic and early Cenozoic time (Hamilton, 1969b; Colpron and Nelson, 2009; Pavlis et al., 2019; for opposite polarity, see Johnston, 2008). Plate motion reconstructions between the northern Sierra Nevada and northwest Washington State (Atlantic and Pacific hotspot reference frames; Engebretson et al., 1985) indicate high-angle convergence and subduction in the latest Jurassic (ca. 150 Ma) progressing into right-oblique motion through late Early Cretaceous time (ca. 100 Ma). Analysis of the North American apparent polar wander path, however, shows left-oblique interaction with the Pacific basin ca. 150–135 Ma (May and Butler, 1986; May et al., 1989). Žák et al. (2015) interpreted magnetic fabric in northeast Oregon (Wallowa batholith; Fig. 1A) as recording dextral translation of the Blue Mountains Province ca. 140–125 Ma. Regional structural and stratigraphic relations to the south suggest northward terrane transport ca. 110 Ma (western Nevada shear zone; Wyld and Wright, 2001). In either case (sinistral or dextral motion), sustained plate convergence over the late Mesozoic (also Seton et al., 2012) resulted in closure of marginal marine basins and collapse of fringing volcanic arc assemblages, e.g., Stikinia of the Intermontane composite terrane belt (Monger et al., 1994; Hammer and Clowes, 2004). More outboard oceanic arc/plateau terranes (Wrangellia of the Insular belt; Jones et al., 1977; Nokleberg, 2005) and offshore tectonometamorphic assemblages (Franciscan forearc; Hsü, 1991; Dumitru et al., 2010) were subsequently attached to and/or translated northward along the Cordilleran margin (Coney et al., 1980; Saleeby, 1983; Jones et al., 1986; Kelley and Engebretson, 1994; Johnston, 1999, 2001, 2008; Wyld and Wright, 2001; Housen and Dorsey, 2005; Lee et al., 2007; Schmandt and Humphreys, 2011).

The Intermontane belt includes island arcs, marginal marine basins, and subduction-related accretionary complexes that formed proximal to North America in mid-Paleozoic to early Mesozoic time (e.g., Rusmore et al., 1988, 2013). Fringing arc–accretionary prism pairs include the Quesnellia–Cache Creek terranes of southeast British Columbia (Gabrielse, 1991; Johnston and Borel, 2007), Rattlesnake Creek–Stuart Fork terranes of northern California (Hacker et al., 1993; Ernst et al., 2017), and Olds Ferry–Baker(?) terranes of the southern Blue Mountains Province (Fig. 1A; Oldow et al., 1989; Kurz et al., 2012). Posttectonic sedimentary overlap assemblages indicate attachment to ancestral North America during the Middle Jurassic (e.g., Bowser basin of northern British Columbia [Ricketts et al., 1992; Logan et al., 2000]). Regional high-pressure metamorphism (>7 kbar: Ghent et al., 1979; Webster et al., 2017), pervasive ductile deformation (ca. 172–167 Ma fabric; Gibson et al., 2005), and intermittent magmatism (ca. 159 Ma anatectic plutons; Sevigny and Parrish, 1993) across southern portions of the Omineca belt are best explained by east-west contraction and crustal thickening associated with terrane accretion (Monger et al., 1972, 1982, 1994; Price, 1994; Murphy et al., 1995; Evenchick et al., 2007).

The Insular belt includes volcanic island arcs, oceanic plateaus, and overlying carbonate platforms—Wrangellia, Alexander, Peninsular terranes—which formed in deep-marine settings during early Paleozoic to middle Mesozoic time (Jones et al., 1977; Gehrels and Saleeby, 1987; Umhoefer and Blakey, 2006; Israel et al., 2014). This composite extends from southwestern Alaska through western British Columbia into the Blue Mountains region of Washington-Oregon-Idaho (Figs. 1A and 1B; Hillhouse et al., 1982; Oldow et al., 1989; Dickinson, 2004; Kurz et al., 2012) and was incrementally attached to the Intermontane belt during latest Jurassic to Paleocene time (Monger et al., 1982; Crawford et al., 1987; Hansen et al., 1989; Plafker et al., 1989; Rubin et al., 1990; Burchfiel et al., 1992; Chardon et al., 1999; Cole et al., 1999; Bergh, 2002; Ridgway et al., 2002; Evenchick et al., 2007; Trop and Ridgway, 2007; Schwartz et al., 2010, 2011). Integrated geophysical, geochronological, and mesoscopic structural data support a general model of early high-angle arc-continent collision with a progression toward dextral-oblique plate convergence and margin-parallel translation of outboard terranes, e.g., Wallowa arc of the northern Blue Mountains Province (Beck et al., 1981; Lund, 1984; Engebretson et al., 1985; Umhoefer, 1987; Wernicke and Klepacki, 1988; Hansen et al., 1989; Manduca et al., 1993; Irving et al., 1995; Wyld and Wright, 2001; Giorgis and Tikoff, 2004; Housen and Dorsey, 2005; Rusmore et al., 2013). Timing estimates on terrane accretion in the Pacific Northwest (arc/plateau assemblages between Idaho and Alaska; i.e., >45°N) suggest south-to-north closure of oceanic tracts separating the Wrangellian composite and Laurentian continental margin (e.g., Pavlis, 1982; Hampton et al., 2010).

The Blue Mountains region of northeastern Oregon, west-central Idaho, and southeastern Washington State includes oceanic fragments of the Wallowa and Olds Ferry island-arc terranes (Vallier and Batiza, 1978; Kays et al., 2006; Kurz et al., 2012, 2017), Baker terrane subduction-accretionary complex (Mullen, 1985; Schwartz et al., 2010, 2011), and Izee overlap basin assemblage (Mesozoic clastic terrane of Dickinson and Thayer, 1978; Dorsey and LaMaskin, 2007). Concealed largely by Miocene flood basalt (Columbia River Basalt Group; e.g., Hooper and Swanson, 1990), east-to northeast–trending terrane belts are discontinuously exposed over an area exceeding 50,000 km² (map coverage of Silberling et al., 1984; cf. Jagoutz and Schmidt, 2012, Kohistan arc complex: ∼55,000 km²). Controlled-source seismic refraction and wide-angle reflection profiling across the Baker, Izee, and Olds Ferry belts indicates a crustal thickness of ∼24–36 km (IDOR seismic line L of Stanciu et al., 2016; Davenport et al., 2017; Fig. 1A), and thus of sufficient age/buoyancy to have caused subduction zone jamming and collision-related orogenesis (Cloos, 1993; Dickinson, 2008). According to Vallier (1995), much of the Wallowa arc is missing. Explanations include westward-directed overthrusting of the Laurentian margin during terrane accretion and translation into northwest Washington, western British Columbia, and southern Alaska, i.e., south-to-north segmentation of Wrangellia (Jones et al., 1977; Saleeby, 1983; Johnston, 2001, 2008; figure 8b reconstruction of Dickinson, 2004).

The Olds Ferry terrane contains weakly metamorphosed Middle Triassic–Middle Jurassic arc-volcanogenic rocks of the Weatherby and Huntington Formations (Brooks and Vallier, 1978; Mann and Vallier, 2007; Tumpane and Schmitz, 2009). Intermediate-composition volcanic flows and hypabyssal intrusive rocks are interpreted to represent a deep-marine island arc (Pessagno and Blome, 1986; Vallier, 1995) or continent-fringing arc assemblage (Miller, 1987; Dickinson, 2004; isotopic enrichment of Kurz et al., 2017). In contrast, the Baker terrane includes variably metamorphosed ocean-floor fragments—pillowed basalt, radiolarian chert, serpentinite-matrix mélange—and arc-derived rocks of Middle Devonian to Early Jurassic age (e.g., Elkhorn Ridge Argillite and Burnt River Schist; Blome and Nestell, 1991; Schwartz et al., 2010). Early Permian faunas recovered from the Elkhorn Ridge Argillite support an exotic origin (Tethyan, McCloud affinity; e.g., Bostwick and Koch, 1962). Paleomagnetic data from turbidite deposits of the Ochoco basin (Cenomanian Gable Creek Formation; e.g., Wilkinson and Oles, 1971) suggest late Early Cretaceous proximity to the southern Sierra Nevada (∼39°N) and vertical-axis clockwise rotation (∼37°) with respect to stable North America (Fig. 1A; Housen and Dorsey, 2005). Overlying the Olds Ferry–Baker terrane couplet are fault-bounded clastic units of the Izee assemblage (Late Triassic Carnian–Norian Vester Formation: Dickinson and Thayer, 1978), which contain abundant detrital zircon of Precambrian to Paleozoic age (LaMaskin et al., 2011).

Lower-greenschist- to upper-amphibolite-facies volcanogenic, siliciclastic, and carbonate rocks in west-central Idaho (Riggins Group, Lucile Slate of Hamilton, 1963a, 1969a) are commonly correlated with Wallowa terrane exposures in the Seven Devils Mountains (Vallier and Fredley, 1972; Detra, 1980; Gray and Oldow, 2005), Snake River Canyon (Vallier, 1967; Goldstrand, 1987; White and Vallier, 1994), and Wallowa Mountains of northeast Oregon (Nolf, 1966; Whalen, 1988; Follo, 1994). Tropical marine fauna in the Martin Bridge Formation (reef carbonate facies; Stanley et al., 2008) and paleomagnetic data from the Wild Sheep Creek Formation (upper Seven Devils Group; Vallier, 1977; Hillhouse et al., 1982; Fig. 1B) indicate Middle–Late Triassic paleolatitudes ∼18° north of the equator. Wilson and Cox (1980) estimated ∼60° clockwise rotation of Early Cretaceous plutons (e.g., Wallowa batholith; Walker, 1986; Johnson et al., 2011; Fig. 1A) with respect to North America prior to Eocene time. Intra-oceanic rocks of the Wallowa terrane may represent a southern portion of Wrangellia (Jones et al., 1977; Saleeby, 1983; Hillhouse and Gromme, 1984; McGroder et al., 1989; Dickinson, 2004; Insular belt Epigondolella sp. of Orchard et al., 2006; Johnston, 2008; Kurz et al., 2017) or Stikinia (Sarewitz, 1983; Mortimer, 1986; Oldow et al., 1989) of the Alaskan and Canadian Cordillera.

In the Riggins region, westernmost exposures of ancestral North America (Laurentia) consist of garnet-sillimanite-biotite schist, calc-silicate schist, marble, and minor quartzite of unknown pre-Cretaceous age (Hamilton, 1969a; Lund, 1984; Kelly Mountain Schist of Blake, 1991). Generally speaking, fine- to medium-grained phyllosilicate minerals distinguish deeply eroded continental metasedimentary rocks on the east from coarsely crystalline hornblende-biotite schists and gneisses of island-arc origin to the west (eastern Blue Mountains Province; e.g., Blake et al., 2009). Along-strike correlations include high-grade units of the Mesoproterozoic Belt-Purcell (e.g., Winston, 1986) and/or Neoproterozoic Windermere Supergroups (Ross, 1991; Lewis et al., 2005), which extend southward from British Columbia, Canada, into northern and central Idaho (Fig. 1A). At this latitude (∼45°N), late Precambrian–Permian passive-margin strata are sparse due to exhumation and unroofing of the Idaho batholith ± truncation of the Laurentian continental margin (Lund et al., 2003, and references therein).

Middle Jurassic to late Early Cretaceous magmatism in the Blue Mountains Province is recorded by three populations of intrusive rocks (synthesis of Schwartz et al., 2011). The oldest plutons recognized are gabbroic to quartz dioritic in composition (ca. 162–154 Ma; Parker et al., 2008) and intrude southern portions of the Baker and Wallowa terranes. Intermediate ages are represented by tonalite-trondhjemite-granodiorite plutons within the same terranes (ca. 148–140 Ma; Tulloch and Kimbrough, 2003), e.g., syntectonic Pole Bridge pluton of the northwestern Wallowa batholith (Fig. 1A; Johnson et al., 2011; Žák et al., 2015). Youngest plutons include variably metamorphosed and/or deformed tonalitic bodies that intrude all terranes and the Izee basin assemblage (ca. 130–120 Ma; Johnson and Schwartz, 2009), e.g., ca. 130 Ma Fish Hatchery stock of the eastern Wallowa terrane (Aliberti, 1988; Gray and Isakson, 2016; figures 20 and 21 of Schmidt et al., 2016b; locality AM-5 of Mann, 2018; GSA Rapid River Stop 13 of Fig. 2A).

Calc-alkaline magmatism across the eastern Wallowa terrane and western Laurentian margin is recorded by variably deformed granitic rocks of the Hazard Creek complex (ca. 137–112 Ma; Armstrong et al., 1977; Manduca et al., 1993; Unruh et al., 2008; Schmidt et al., 2016a; Mann, 2018), Little Goose Creek complex (ca. 111–101 Ma; Manduca et al., 1993; Giorgis et al., 2008; Gray et al., 2012; Kauffman et al., 2014; McKay et al., 2017; Montz and Kruckenberg, 2017), and Payette River tonalite (ca. 93–85 Ma; Manduca et al., 1993; Snee et al., 1995; Gray et al., 2012; Idaho batholith border zone plutonic suite of Gaschnig et al., 2010). Magmatic epidote-bearing orthogneiss bodies (Zen, 1985) are elongated subparallel to the arc-continent boundary, with oldest plutons emplaced into volcanic arc assemblages in the west (e.g., ca. 114 Ma, Granite Mountain locality K92–8; Unruh et al., 2008), the youngest plutons emplaced into Laurentian metasedimentary rocks in the east (e.g., ca. 92 Ma, Salmon River Canyon locality LGp3; Gray et al., 2012), and plutons of intermediate age in between (e.g., ca. 110 Ma, Slate Creek locality C; Kauffman et al., 2014; Figs. 2 and 3B). Middle Cretaceous magmatism recorded by ca. 114–92 Ma plutons (sample localities K92–8, LGp3, and C) was coeval with east-west contractional deformation/crustal thickening in the SRSZ as indicated by westward-directed thrust/reverse faulting, tight-to-isoclinal folding, and tectonite fabric development (Snee et al., 1995; Gray et al., 2012; McKay et al., 2017).

Across the lithospheric boundary separating arc terranes of the Blue Mountains Province and western Laurentia (Stanciu et al., 2016), initial strontium ⁸⁷Sr/⁸⁶Sr ratios (Sr_i) recorded in Permian–Cretaceous calc-alkaline plutons show a change from Sr_i ≤ 0.704 in the west (oceanic crust) to Sr_i ≥ 0.708 in the east (continental crust; Armstrong et al., 1977; Manduca et al., 1992; Fleck and Criss, 2004). Width of the Sr_i isotopic transition (≤30 km, west-to-east; e.g., Fleck and Criss, 1985; Benford et al., 2010) has been attributed to post-accretion dextral-transpressional deformation along the Western Idaho shear zone (e.g., McClelland et al., 2000; Tikoff et al., 2001). East-west shortening is estimated between ∼28 and ∼112 km, with ∼50 km of right-lateral offset (Giorgis and Tikoff, 2004; Giorgis et al., 2005). This structure is proposed to follow the Sr_i 0.706 contour (i.e., best-fit curve) from the Owyhee Mountains in southwestern Idaho (Benford et al., 2010) through McCall and Riggins (e.g., Blake et al., 2009) before entering the Orofino-Ahsahka area in north-central Idaho (Figs. 1A and 2A; McClelland and Oldow, 2007; Lewis et al., 2014; Schmidt et al., 2016a). According to Giorgis et al. (2008), the main phase of deformation overlapped ca. 105–90 Ma magmatism and overprinted accretion-related contractional structures of the SRSZ (Lund and Snee, 1988): unspecified thrust/reverse faults, upright to overturned folds, and pervasive LS tectonite fabric.

Following the language of Lund et al. (2008), “Salmon River suture” refers to the ancient boundary separating accreted terranes of the Blue Mountains Province and western Laurentia as approximated by the Sr_i 0.706 isopleth. In contrast, “Salmon River suture zone” terminology considers the broad belt of metamorphism, magmatism, and penetrative deformation overlapping the Sr_i 0.706 isopleth and related to arc-continent collision (Lund, 1984; Giorgis et al., 2007). Although boundaries separating accretion- and non-accretion-related tectonic elements are not defined (field mapped at 1:24,000-scale), and superposition relations remain undocumented (overprinting shear zone fabric), structures assigned to the SRSZ and Western Idaho shear zone are explained by right-oblique oceanic-continental plate convergence (Giorgis et al., 2005; Lund et al., 2008). In addition to sharing a mid-Cretaceous kinematic history (dextral transpression), both elements are described as north-south–striking (McCall–New Meadows–Riggins–Slate Creek segment) and as comprising east-dipping LS tectonite fabric (figure 5 synoptic plot of McClelland et al., 2000; Tikoff et al., 2001). In this paper, we define through-going boundaries of the SRSZ based on across-strike structural mapping, polyphase fabric analysis (superposition relations), and the east–west limits of pervasive ductile deformation.

Fabric and fold element data were collected from 70 outcrops studied along an unbroken ∼48 km transect (Table 1; Fig. 2A) extending from the Oregon-Idaho border (Snake River: elevation ∼1400 ft [427 m]; ∼45°30′20″N, ∼116°40′00″W) east over the northern Seven Devils Mountains (He Devil Peak: elevation ∼9340 ft [2847 m]) into the lower Salmon River Canyon (French Creek: elevation ∼2000 ft [∼619 m]: ∼47°27′30″N, ∼116°00′00″W) near the village of Riggins, Idaho. Geologic strip maps and structural sections (1:24,000 scale, paper base mapping; Gray, 2013) were simplified to show major faults (Fig. 2B), representative fold geometries, and polyphase fabric elements with emphasis placed on regional synmetamorphic structures (S1–L1; Gray et al., 2012). Maps and sections are accompanied by stereographic projection plots compiled across four fault-bounded domains (I–IV; Fig. 3) bridging accreted island-arc rocks of the eastern Blue Mountains Province (Wallowa terrane + correlative units; Hamilton, 1963a; Vallier, 1977; Schmidt et al., 2016b) and the western margin of Laurentia (Hamilton, 1969a; Blake, 1991). From structurally lowest to highest levels (west to east across the SRSZ), major domain-bounding structures include the Heavens Gate fault (Gray and Oldow, 2005; Kauffman et al., 2014; Schmidt et al., 2016b), Rapid River thrust/Riggins normal fault (Hamilton, 1969a; Onasch, 1977, 1987; Hooper, 1982; Lund and Snee, 1988), and Pollock Mountain thrust (Aliberti, 1988; Blake, 1991; Selverstone et al., 1992; McKay et al., 2017).

Finite strain was evaluated on clast-supported volcanic conglomerate along the western edge of domain II in the northeastern Seven Devils Mountains (Heavens Gate Ridge; Fig. 2A). On a single rectangular block (∼25 m³), the long and short axes of pebbles and cobbles were measured across three intersecting faces (10 clasts/face). Mutually perpendicular faces and corresponding strain planes (cf. Ramsay and Huber, 1983) were defined by the S1 foliation (top surface of exposure: x-y plane) and its intersection with spaced cleavage (x-z plane) and systematic jointing (y-z plane). Lengths of long (x), intermediate (y), and short (z) principal axes were used to calculate axial ratios and construct a Flinn diagram. Pilot data were collected to (1) show changes in bulk strain across the Heavens Gate fault (located ∼1.5 km west), and (2) compare lithic clast geometries with porphyroclast shape fabrics determined along the arc-continent boundary (Giorgis and Tikoff, 2004).

Our westernmost sample locality (06KG15; 11KGHG01 of Gray, 2013) was selected due to its proximity to the Heavens Gate fault (∼75 km trace length), which represents both a domain boundary (I/II; Fig. 3A) and regional discontinuity (metamorphism, bulk strain) in the eastern Wallowa terrane (Fig. 2A). Sample localities SHd-04 and FC-05 were selected from central and eastern portions of the transect (domains III and IV, respectively; Fig. 3B), where higher-grade volcanic arc and continental margin assemblages are exposed (Hamilton, 1969a; Blake, 1991). These localities occupy critical areas of the SRSZ emphasized by previous workers (SHd-04: two-stage garnet of Selverstone et al., 1992; GSA Stop 2.7 shear zone flattening of Blake et al., 2009) and introduced in the present study (FC-05; Idaho batholith border zone of Taubeneck, 1971).

Zircon was isolated from oriented field samples using standard gravimetric and magnetic separation techniques at Boise State University (Boise, Idaho). Handpicked grains from sample separates were mounted on epoxy disks, polished to expose zircon centers, carbon-coated, and imaged by cathodoluminescence (CL) to determine spot locations for U-Pb isotopic analyses. Laser ablation–inductively coupled plasma–mass spectrometry (LA-ICP-MS) and chemical abrasion–isotope dilution–thermal ionization mass spectrometry (CA-ID-TIMS) analyses were also conducted at Boise State University. LA-ICP-MS used a New Wave Nd:Yag ultraviolet 213 nm laser coupled to a Thermo Scientific X-Series II quadrupole ICP-MS. Zircon dates were obtained using CA-ID-TIMS (on select grains) based on LA-ICP-MS results and morphology determined from CL. Zircons removed from mounts were subjected to a modified version of the chemical abrasion method of Mattinson (2005), consisting of a single step with concentrated HF (190 °C for 12 h). The U and Pb isotopic measurements were made on an IsotopX Isoprobe-T multicollector TIMS equipped with an ion-counting Daly detector and 9 Faraday cups. Analytical protocols, standard materials, and data reduction methods are outlined in the Supplemental Material¹.

Sample locality JV-003 in the lowermost Riggins Group (Figs. 2A and 3B; structural domain III; undated Fiddle Creek Schist; Hamilton, 1963a) was selected given its proximity to the Salmon River suture zone–Western Idaho shear zone boundary customarily assumed by previous workers (McClelland et al., 2000; Tikoff et al., 2001; Giorgis et al., 2005, 2008; GSA Stop 2.6: “transition” of Blake et al., 2009). At this location, felsic volcanic rocks of the Wallowa terrane (Early Permian Hunsaker Creek Formation equivalent of Kauffman et al., 2014) lie between the Rapid River and Pollock Mountain thrusts, i.e., east-west contractional structures bounding domains II/III and III/IV, respectively (Fig. 2B). Consistent with U-Pb zircon localities described above (06KG15: upper Seven Devils Group; SHd-04: Pollock Mountain Amphibolite), Lu-Hf garnet locality JV-003 represents a major lithotectonic assemblage (lower Riggins Group) and provides context for the continued assessment of overlapping(?) orogenic belts in west-central Idaho (Gray et al., 2012).

Several 200–250 mg garnet fractions were handpicked from our mechanically crushed and separated field sample. Fragments containing the least amount of mineral inclusions (e.g., quartz and ilmenite) were selected for analysis. Other inclusion types can possess different initial Hf isotopic compositions and were avoided (e.g., zircon; Scherer et al., 2000). Garnet fractions and one 200–250 mg whole-rock powder fraction were dissolved on a hotplate (∼110 °C) in a concentrated HF and HNO₃ acid mixture (10:1 ratio) for 2–3 d for garnet and 5–7 d for whole rock. A second 200–250 mg whole-rock powder fraction was dissolved in a high-pressure, steel-jacketed Teflon capsule (∼160 °C; 5–7 d) in concentrated HF and HNO₃ mixture (10:1 ratio). Following dissolution, samples were dried and converted from fluorides to chlorides using a H₃BO₃ and 6 M HCl mixture. Each sample was dried again, dissolved in 6 M HCl, spiked with a mixed Lu-Hf tracer, and allowed to equilibrate on a hotplate for at least 24 h. Chemical separations of Lu and Hf were carried out in clean laboratory facilities located at Washington State University (Pullman, Washington). Chemically separated Lu and Hf were dissolved in 2% HNO₃ and analyzed on a ThermoFinnigan Neptune multicollector (MC) ICP-MS at Washington State University. Our single Lu-Hf isochron age was calculated using Isoplot (Ludwig, 2003) and a ¹⁷⁶Lu decay constant of 1.867 × 10⁻¹¹ yr⁻¹ (Scherer et al., 2001; Söderlund et al., 2004). Cheng et al. (2008) and Zirakparvar et al. (2010) discussed the dissolution, spiking, and chemical separation methods used in this study.

Western foothills of the northern Seven Devils Mountains (Figs. 2A and 4A) rise out of Hells Canyon and Little Granite Creek drainage (45°30′20″N latitude) along a shallow (≤30°) west-facing dip slope in the Wallowa terrane (Fig. 3A, cross-section A–A′). Lower greenschist–facies (albitized) basaltic-andesite flows, interbedded volcaniclastic rocks, and carbonate lenses of the Middle to Late Triassic Wild Sheep Creek Formation (Figs. 1B and 4B−4E; Vallier and Batiza, 1978; Vallier, 1998) are deformed in north-south–trending upright symmetric open-to-close folds (outcrops 3 and 4, Table 1; cf. Fleuty, 1964). Broad-wavelength structures (λ ≥ 1 km; e.g., Devil’s Arch) are associated with steep, easterly dipping (≥70°) longitudinal fractures locally displaying strike-slip kinematics (Figs. 4B and 4G). In the far west, subvertical, north-south–striking fractures crosscut a quartz diorite pluton of late Early Cretaceous age (115.3 ± 2.5 Ma: hornblende, biotite; K-Ar locality V-3–86 of Vallier, 1995). Along this segment of the transect (domain I), arc supracrustal and plutonic rocks (e.g., V-3–86) preserve original igneous and sedimentary textures—glomeroporphyritic fabric, volcanic flow banding, lithic clast imbrication, graded bedding (Figs. 3A, 4C−4F)—and show no evidence of dynamic recrystallization or pervasive ductile strain (Vallier and Batiza, 1978).

In the Windy Saddle area (Figs. 5A and 5B), shallowly to moderately northeast-dipping volcanic sandstone and breccia are structurally overlain by upper-greenschist-facies (locally hornblende- and/or oligoclase-bearing) metamorphic tectonites across the Heavens Gate fault (outcrops 5–12, section A′–A″). This structure separates penetrative and nonpenetrative domains (II and I, respectively) and marks the limit of high-strain ductile deformation west of the arc-continent boundary (Fig. 2A, ∼116°30′00″W, “western limit of ductile fabric”). Three-point solutions determined on fault segments between Cannon Ball Mountain and Old Timer’s Mine suggest a shallow east- to southeast-dipping (≤30°) surface cutting synmetamorphic fabric in the hanging wall (outcrops 5, 11) and sedimentary layering in the footwall (outcrop 4). Local fault kinematics are unclear; however, the juxtaposition of greenschist tectonites (Figs. 5C, 5D, and 5E) with lower-grade rocks recording original textures (Figs. 4C, 4D, and 4E) supports local east-side-up/thrust motion. Schmidt et al. (2016b) interpreted normal displacement to the north (Lucile area; Fig. 2A). The Heavens Gate fault/tectonic front has a protracted deformation history that may also include strike-slip movement related to mid-Cretaceous dextral transpression in the lower Slate Creek–Riggins–northern McCall areas, i.e., oblique arc-continent collision (Lund and Snee, 1988; McClelland et al., 2000; cf. Murphy, 1997, transpressive thrusting, Yukon Territory).

Upper-plate protoliths include greenish-gray volcanic conglomerate, argillite, and minor basaltic-andesite flows of the Wild Sheep Creek Formation (Fig. 5B). Across Heavens Gate Ridge (Figs. 4B and 5A), supracrustal cover rocks are penetratively deformed with small (<100 m²) intermediate- to mafic-composition hypabyssal intrusive bodies of unknown age (T. Vallier, 1997, personal commun.). Coarsely crystalline feldspar porphyroclasts record a downdip (90° from strike azimuth) to steeply plunging mineral lineation and top-to-the-west/reverse sense of shear (Figs. 5D and 5E). Southeast of Windy Saddle (∼1.5 km), deeply eroded argillaceous rocks enclose a small tabular intrusion (locality 06KG15 of this study; outcrop 11) containing a moderately east-dipping gneissic foliation (S1; ∼50°) and downdip mineral lineation (L1) defined by aligned hornblende (Figs. 3A, 5B, 5F, and 5G). Smooth to moderately rough disjunctive foliation (microlithon spacing ≤5 mm: S2) cuts LS tectonite fabric at a shallow angle (≤30°; Fig. 5F), as described for second-generation structures in the Riggins region (figures 7C and 7D of Gray et al., 2012).

In clast-supported pebble-cobble conglomerate, S1 dips variably eastward (∼20°–60°) and is defined by triaxially deformed volcanic (mostly basaltic andesite) and minor carbonate clasts. Analysis of x/y and y/z axial ratios (cf. Flinn, 1962; Ramsay and Huber, 1983) revealed both flattening and constrictional strains (Figs. 6A and 6B), consistent with mixed oblate/prolate shape fabric reported from middle Cretaceous granitoids of the Little Goose Creek complex (Fig. 2A; Giorgis and Tikoff, 2004). Long-to-short axes viewed on the x-z strain plane reach 10-to-1 (Figs. 6C and 6D); aligned lithic clasts form a pervasive downdip to southeast-plunging stretching lineation on S1 (L1: ≤20° rake; Table 1). Shear sense indicators are rare. In most outcrops, clasts are symmetrically flattened on S1 (Fig. 6E), and the sense of shear is unclear (Figs. 6D and 6F); however, sigmoidal strain markers identified ∼1.5 km north of Heavens Gate lookout show top-to-the-west shear (Figs. 5C and 5E), as evidenced by stretched pebbles in structurally overlying Late Triassic/Norian carbonate rocks (outcrop 13, Fig. 3A; Martin Bridge Limestone of Hamilton, 1963a, 1969a; Epigondolella conodont sp. of Sarewitz, 1982). Asymmetric lithic clasts recording top-to-the-west/thrust kinematics (east-west crustal shortening; Figs. 2A, 3A, and 5C) combined with local strike-slip indicators (north-south subvertical shearing; Fig. 4G) suggest post-Norian transpressional deformation in the eastern Wallowa terrane (Žák et al., 2012, 2015).

Primary textures east of the Heavens Gate fault—flow banding, phenocryst clustering, clast imbrication, and graded bedding (S0)—are largely obscured by the S1 transposition foliation (S0 = S1; Fig. 7A) and overprinting postmetamorphic deformation (e.g., small-scale duplexing; Fig. 7B). Where recognized, S1 crosscuts S0 at a high angle (>70°); bedding/cleavage intersection relations suggest tight-to-isoclinal folding of S0, where S1 is axial-planar to overturned west-vergent nappe structures (southern Heavens Gate Ridge, section A′–A″; Fig. 3A). Depending on lithology, lithic clast size, and phyllosilicate content, superposed structures include penetrative spaced or crenulation cleavage (S2, S3 fabric) and associated intersection or crenulation lineations (L2, L3). Typically oriented subparallel to S1, S2 crosscuts synmetamorphic fabric at a shallow angle (<30°). In contrast, S3 strikes orthogonal to S1 and crosscuts at a high angle (Fig. 7C). Pencil cleavage formed by the intersection of S1 and S3 (argillaceous rocks) occupies an axial-planar position to shallow, easterly plunging, upright symmetric open-to-close map-scale folds (D3 of Gray and Oldow, 2005).

Greenschist-facies volcanic, volcaniclastic, and carbonate rocks in the west (Wild Sheep Creek and Martin Bridge units) are structurally overlain by higher-grade (biotite- to andesine-bearing) metamorphic tectonites across the Rapid River thrust fault (Figs. 2A and 3B, cross-sections A′′–A′′′ and B–B′; Hamilton, 1963b; Onasch, 1977; Lund and Snee, 1988). This east- to southeast-dipping, syn- to postmetamorphic structure (Hamilton, 1963a; Onasch, 1977; Aliberti, 1988; Snee et al., 1995) separates the Permian–Triassic Seven Devils and Riggins Groups (domains II and III, respectively). Along our transect, the Rapid River thrust is cut out by a major north-south–striking brittle normal fault (Riggins fault of Hooper, 1982; Hamilton, 1969a; Fitzgerald, 1984; Aliberti, 1988; Gray and Oldow, 2005) carrying gently west-tilted (≤25°) continental flood basalt in its hanging wall (Miocene Columbia River Basalt Group; Hooper and Swanson, 1990). The Riggins fault dips steeply east (≥75°), slices through ridge-and-valley topography at ∼90° (e.g., Papoose Saddle–Squaw Creek segment; Fig. 2A; Quarcoo and Gray, 2016), and relates to basin-and-range extensional deformation (Capps, 1941; Tikoff et al., 2001; Giorgis et al., 2006; Schmidt et al., 2016b) offsetting the crust-mantle boundary (Fig. 1A, cross-section A–A′; seismic line L of Stanciu et al., 2016).

Amphibolite-facies metamorphism and attendant ductile deformation obscure original igneous and sedimentary textures in the Riggins Group (S1_R transposition foliation of Onasch, 1977, 1987). Bedding/cleavage intersection relations recorded in deep-marine metasedimentary rocks (Squaw Creek Schist; Hamilton, 1963a; Jurassic metaflysch of Lund et al., 2007) support early tight-to-isoclinal map-scale folding of S0 (outcrop 17; Table 1; Figs. 3B and 8A), where S1 is both axial-planar (fold hinges; Fig. 8B) and parallel to S0 (attenuated limbs; Fig. 8C). High-strain fabric (S1) is broadly deformed (λ ≥ 1.5 km) across upright symmetric to overturned southeast-plunging (∼20°–40°) postmetamorphic folds. Map-scale structures spanning lower portions of the Salmon River Canyon (Riggins town site to Ruby Rapids; Fig. 3B) include the upright Riggins synform and overturned Lake Creek antiform (outcrops 16–35, sections B–B′ and C–C′; Hamilton, 1969a; Onasch, 1977, 1987; Bruce, 1998; figure 10 of Blake et al., 2009; figure 16 of Schmidt et al., 2016b).

East of Berg Creek Ranch (Fig. 8A), synmetamorphic foliation is wrapped around the Lake Creek antiform (west end of section C–C′; Fig. 3B). Attitudes (S1) change abruptly across the hinge (∼1.5 km wide), where upright west-vergent mesoscopic folds and moderately east-dipping brittle reverse faults disrupt lower units of the Riggins Group (outcrops 25–33). Basal metaconglomerate of the Lightning Creek Schist defines a prominent marker horizon (Hamilton, 1963a); axial ratios reach 8-to-1 in triaxially deformed felsic volcanic clasts (outcrop 25; Fig. 8D). South of the Salmon River, arc-volcanogenic rocks are crosscut by an undated intermediate-composition pluton related(?) to the Chair Point plutonic complex, Permian–Triassic crystalline basement of the Wallowa terrane (Fig. 2A; Walker, 1986; Kurz et al., 2012; Kauffman et al., 2014). Local plutonic rocks show a steeply east-dipping (∼80°) gneissic foliation that is axial-planar to the Lake Creek antiform (S3: outcrop 27; Figs. 3B, 8E, and 8F). In adjacent mica-rich metasedimentary rocks (Lightning Creek Schist), S3 is associated with a shallow (≤20°) south-plunging crenulation lineation on S1 (L3: outcrop 24). Crenulations are orthogonal to steep downdip (∼75° west) chlorite mineral lineations (L1: outcrop 26) in the underlying Fiddle Creek Schist, i.e., silica-rich volcanic tuffs and flows defining the basal unit of the Riggins Group (Hamilton, 1963a).

West of Rough Creek (∼100 m; Fig. 3B), S1 dips steeply to the east (outcrop 28: Lu-Hf locality JV-003; Figs. 9A and 9B). Early fabric is subtle here and follows the trace of aligned phyllosilicates. Augen-shaped aggregate clusters of quartz ± feldspar enclosing millimeter-scale garnet are elongated in the downdip direction (L1; Fig. 9C). In this area, S2 is dominant and relates(?) to moderately east-dipping imbricate faults disrupting the antiformal hinge (Figs. 3B and 9A, inset). Approaching Ruby Rapids and the Berg Creek Amphibolite (Hamilton, 1963a), counterclockwise-rotated (viewing north) garnet porphyroblasts record top-to-the-northwest/reverse shear (uppermost Lightning Creek Schist; outcrops 33, 34, and 35). Deeply eroded biotite-chlorite schist containing ca. 112 Ma synkinematic garnet (Sm-Nd sample locality ID48; McKay et al., 2017) crops out in footwall exposures of the Pollock Mountain thrust (Blake, 1991; Selverstone et al., 1992), which strikes north to northeast across the Riggins region (>100 km trace length; Aliberti, 1988) and separates domains III and IV along our transect (Figs. 2A and 3B).

South of the Salmon River, the Pollock Mountain thrust places upper-amphibolite-facies metamorphic tectonites over lower-grade pelitic rocks of the Squaw Creek Schist (Fig. 2A; e.g., Ar-Ar locality R30, pre–118 Ma synkinematic hornblende of Snee et al., 1995; Aliberti and Manduca, 1988). Along our transect, west-northwest–directed thrusting cuts out the Squaw Creek Schist and carries the Pollock Mountain Amphibolite over the Berg Creek Amphibolite (eastern limb of Lake Creek antiform; section C–C′, outcrops 36–39). Metamorphic fabric dips variably southeast (∼45°–85°) in both the hanging wall and footwall (Gray et al., 2012; Blake et al., 2016); in the former, aligned hornblende needles define a downdip to obliquely (≤20° rake) southeast-plunging mineral lineation on S1 (L1: outcrops 37 and 39; L_m-1 of Blake et al., 2009). At locality SHd-04 (Fig. 10A), LS tectonites are deformed in west-northwest–vergent tight-to-isoclinal folds (SHd-04 host rocks). Mesoscopic structures plunge gently west-southwest (∼20°–30°) and possess a steep southeast-dipping (≥70°) axial-planar spaced and crenulation cleavage (tonalite dike and amphibolite host, respectively; Figs. 10B and 10C). Counterclockwise-rotated (viewing north) subhedral garnet porphyroblasts recording top-to-the-northwest shear (Fig. 10D), combined with isoclinally folded felsic stringers (Fig. 10E), support dextral-transpressional deformation in easternmost exposures of the Pollock Mountain Amphibolite.

Magmatic epidote-bearing plutons east of the Pollock Mountain Amphibolite (outcrops 40–47) are generally characterized by a northeast-striking gneissic foliation (S1: ∼45°–85° dip; Fig. 11A) and downdip to obliquely southeast-plunging hornblende ± biotite lineation (L1: ≤45° rake; Fig. 11B). South of the Salmon River, however, S1 attitudes vary from north-south–striking (∼15°–50° west-dipping) to east-west–striking (∼35° north-dipping) and thus effectively span a complete range of strike orientations (Onasch, 1977; Blake et al., 2016). Mineral and stretching lineations on S1 are rare (L1: ≤40° oblique plunge). Variably oriented metamorphic tectonites are mapped across the Van Ridge Gneiss (south of section B–B′ in Blake et al., 2016), which lies directly west of the arc-continent boundary (Sr_i 0.706 isopleth) akin to ca. 118–112 Ma granitic rocks of the Hazard Creek complex (Blake, 1991; Manduca et al., 1993; Mann, 2018; Unruh et al., 2008). North of the Salmon River, tight-to-isoclinally folded pegmatite stringers record opposing shear kinematics. Mesoscopic structures are observed across a single surface exposing the y-z strain plane (outcrop 48; Fig. 11C), where gneissic foliation (S1) strikes northeast (∼N10–15E), dips steeply southeast (∼80°), and contains a strong downdip stretching lineation (L1) defined by rodded quartz (Fig. 11B). From northeast to southwest across the exposure, minor folds display sinistral shear (S), dextral shear (Z), and neutral (M) shapes. Local ductile deformation/fold patterns revealing a combination of S, Z, and M geometries are consistent with east-west shortening and flattening perpendicular to S1/x-y (cf. Ramsay and Huber, 1983), as recorded by structures ∼30 km west on Heavens Gate Ridge (Wild Sheep Creek Formation; Figs. 6A and 6E).

Near the mouth of Kelly Creek (Fig. 3B), medium-grained calc-silicate metasedimentary rocks (undated Kelly Mountain Schist; Blake, 1991) mark the western edge of ancestral North America (Fig. 12A, cross-section C–C′, outcrop 49; Table 1). In this area, schistosity (S1) strikes subparallel to the arc-continent boundary (∼N10–25E trend). South of Kelly Creek, however, S1 orients ≤45° oblique to the boundary (e.g., outcrop 50). East of the Crevice pluton (figures 11 and 12c of Gray et al., 2012), tight-to-isoclinal folds deform fine-grained garnet-bearing mica schist (Fig. 12B, section D–D′, outcrops 54, 55, and 56). Second-order (parasitic) structures observed on shallow (∼20°–30°) east-northeast–dipping limbs support west-southwest tectonic transport (Fig. 12C), consistent with structures in accreted island-arc assemblages to the west (Seven Devils Group [Figs. 5B, 5C, 5E, and 7B], Riggins Group [Figs. 8A, 8B, 9A, and 9B], and Pollock Mountain Amphibolite [Figs. 10A, 10B, and 10D]). Mineral and stretching lineations (L1) are variably developed on S1, range from downdip to ∼45° obliquely plunging (Figs. 2A, 3B, and 12D, outcrops 49–54; Blake et al., 2016), and consist of aligned muscovite + biotite ± sillimanite (L_m-1 of Blake et al., 2009).

Granitic orthogneiss units east of the Sr_i 0.706 isopleth (undated Partridge Creek Gneiss; Blake, 1991) include tightly folded country rock screens of western Laurentia (Kelly Mountain Schist). Upright structures show steeply east-dipping axial surfaces (e.g., figure 9b of Gray et al., 2012). Across eastern portions of the Crevice pluton, east-northeast–dipping (≤45°) mylonitic shear zones (S2) offset S1 in the Kelly Mountain Schist (Fig. 12E). Local truncation relations are reminiscent of polyphase structures in the Riggins Group (e.g., Lightning Creek Schist), where shallow east-dipping brittle shears cut subvertical synmetamorphic fabric (Fig. 9A, inset). Ductile deformation is tracked eastward through continental metasedimentary rocks (outcrops 54–58) into border zone plutons of the Idaho batholith (outcrops 59, 60, and 61; Figs. 1A, 2A, and 3B; Taubeneck, 1971; Manduca et al., 1993; Gaschnig et al., 2010). Across the lower French Creek drainage, moderately developed north-northwest–dipping (∼50°–60°) gneissic fabrics are observed (outcrops 62–65); aligned biotite mats, elongate mafic masses (schlieren), and leucosomal layers combine to form migmatitic textures previously unrecognized along our transect.

Approximately 2 km east of French Creek, a strong north-south–striking solid-state fabric is observed in coarsely crystalline granitic orthogneiss (U-Pb locality FC-05, outcrop 67; Fig. 12G). Gneissic fabric is marginal to the French Creek migmatite zone, shares attitudes with S1 in the southeastern Crevice pluton/Kelly Mountain Schist, and resides >10 km east of the Sr_i 0.706 isopleth. Upriver (∼0.5 km), aligned mica flakes mark the trace of an east-dipping (∼65°) spaced foliation in medium-grained biotite tonalite (outcrop 69). Compared with locality FC-05, rocks record lower strain and resemble mildly deformed portions of the Crevice pluton (figure 12b of Gray et al., 2012) and Payette River tonalite (Fig. 2A; GSA Stop 5.3 of Giorgis et al., 2006). At this latitude (∼45°25′00″N), outcrop 69 marks the eastern limit of penetrative ductile deformation; farther east, massive medium-grained hornblende tonalite of the Idaho batholith is encountered (outcrop 70, Fig. 12H; peraluminous suite of Gaschnig et al., 2010). Our approximate boundary (Fig. 2A, “eastern limit of ductile fabric,” ∼116°00′00″W) lies ∼55 km north of the undeformed two-mica granite locality (GSA Stop 5.4) described by Giorgis et al. (2006).

Sample 06KG15 was collected from the southern terminus of Heavens Gate Ridge, ∼1 km east of Windy Saddle in the northeastern Seven Devils Mountains (Figs. 2A, 3A, 4A, 4B, 5A, and 5B). Medium-grained (≤5 mm) biotite-bearing hornblende diorite was sampled for ²³⁸U/²⁰⁶Pb zircon analysis (Figs. 5F and 5G). The majority of grains recovered are subhedral to anhedral with rounded edges, resulting in ellipsoidal geometries; prismatic and acicular crystal habits are less common. Most lack growth zoning and xenocrystic core-rim textures; however, many grains display possible remnant sector zoning in CL (Fig. 13A). LA-ICP-MS analysis yielded U-Pb dates ranging from ca. 150 to 127 Ma, and a few older dates between ca. 677 and 292 Ma. CA-ID-TIMS analysis of five zircon grains possessing rounded, prismatic, and acicular habits yielded a weighted mean U-Pb date of 135.84 ± 0.07 Ma and is interpreted here as a magmatic crystallization age (Figs. 13B and 13C; Table 2). Early Cretaceous zircon crystallization in the Heavens Gate area overlapped with syntectonic emplacement of ca. 145–120 Ma calc-alkaline plutons in northeastern Oregon, e.g., Wallowa and Bald Mountain batholiths, Cornucopia stock (Johnson et al., 1997; Žák et al., 2012, 2015). Given its age, location, and bulk chemical composition, sample 06KG15 may represent a small satellite body associated with weakly metamorphosed (greenschist) and deformed (local magmatic fabric) hornblende-bearing tonalite of the Fish Hatchery stock (130.0 ± 1.9 Ma; ²³⁸U/²⁰⁶Pb locality AM-05 of Mann, 2018; Aliberti, 1988; Snee et al., 1995; Gray and Isakson, 2016). Both intrusions are hosted by volcanogenic rocks of the Wild Sheep Creek Formation (Sarewitz, 1983; Schmidt et al., 2016b) and are proximal to east-northeast–dipping imbricate faults (U-Pb locality 06KG15, upper plate of Heavens Gate fault [Fig. 7B], and U-Pb locality AM-05, footwall of Morrison Ridge–Rapid River fault system [Fig. 2A]). At this latitude (∼45°21′00″N), locality 06KG15 provides the westernmost zircon data from intrusive rocks assigned to the SRSZ.

Shrunken Head dike rises above the Salmon River ∼4 km west of the Sr_i 0.706 isopleth (Gus Creek confluence; Figs. 3B, 10A, and 10B). Fine- to medium-grained (≤5 mm) biotite tonalite was collected for ²³⁸U/²⁰⁶Pb zircon analysis (Fig. 10C). Most grains are prismatic; however, crystal habits range from acicular to equant. Internal structures vary from simple/unzoned to showing well-developed growth rings; some grains possess inherited cores with remnant and resorbed sector zoning (Fig. 13A). LA-ICP-MS analysis yielded the following U-Pb zircon dates: ca. 94–83 Ma, ca. 120–97 Ma, and ca. 390–140 Ma. The youngest and most consistent population produced a weighted mean date of 90.1 ± 1.2 Ma. We interpret older populations as xenocrysts incorporated from enclosing island-arc–supracrustal rocks, i.e., late Paleozoic to mid-Mesozoic volcanic cover and/or associated hypabyssal intrusions (Vallier, 1977; Schmidt et al., 2016b). CA-ID-TIMS analysis on six zircon grains showing unzoned to well-developed growth rings yielded a weighted mean U-Pb date of 90.62 ± 0.03 Ma, which is interpreted here as a magmatic crystallization age (Figs. 13B and 13C; Table 2). Late Cretaceous dike injection/zircon crystallization (ca. 91 Ma) overlapped with silicic magmatism in the Looking Glass pluton (91.7 ± 2.4 Ma; ²³⁸U/²⁰⁶Pb locality LGp03 of Gray et al., 2012), Payette River tonalite (91.5 ± 1.1 Ma; ²³⁸U/²⁰⁶Pb locality 01–53; Giorgis et al., 2008; see Fig. 2A), and other syntectonic plutons in west-central Idaho (e.g., 90 ± 5 Ma; ²³⁸U/²⁰⁶Pb locality 83z11 of Manduca et al., 1993; Idaho batholith border zone suite of Gaschnig et al., 2010). To our knowledge, all previously published dates within this age range (U-Pb zircon: 90 ± 5 Ma) were derived from calc-alkaline intrusive rocks located east of the Sr_i 0.706 isopleth (Fig. 2A).

Sample locality FC-05 resides in the western gneissic border zone of the Idaho batholith (Taubeneck, 1971), ∼10 km east of the Sr_i 0.706 isopleth (southern banks of Salmon River; Figs. 2A, 3B, and 12F). Medium to coarsely crystalline biotite tonalite (Fig. 12G) was collected for ²³⁸U/²⁰⁶Pb zircon analysis. Crystal habits range from prismatic to subequant, and internal structures are variable. Most grains show convoluted, embayed, and/or rounded xenocrystic cores or are CL dark with bright overgrowth textures; others exhibit well-developed growth zoning (Fig. 13A). LA-ICP-MS zircon analysis yielded distinct populations of ca. 97–85 Ma (equant/CL-dark) and ca. 165–123 Ma. CA-ID-TIMS analysis of seven grains selected from our younger population yielded concordant results between ca. 93 and 86 Ma (Figs. 13B and 13C; Table 2), with no resolvable age of zircon crystallization. If dates reflect magmatic crystallization, then FC-05 overlaps in age with SHd-04, situated ∼15 km west—in addition to LGp03, 01–53, and 83z11—and also belongs to the border zone suite (e.g., Gaschnig et al., 2010). At this latitude (∼45°25′00″N), locality FC-05 provides the easternmost zircon data from intrusive rocks assigned to the SRSZ.

Sample JV-003 was collected from siliceous metavolcanic rocks (lowermost Riggins Group; Hamilton, 1963a) exposed along the Salmon River ∼6 km west of the Sr_i 0.706 isopleth (Lake Creek bridge area; Figs. 2A, 3B, 9A, and 9B). Medium-grained garnet porphyroblasts are subhedral to anhedral in form, surrounded by augen-shaped clusters of quartz ± feldspar (Ca-diffusion halos), and elongated subparallel to stretching in the downdip direction (Figs. 3B and 9C). Linear regression using three garnet fractions and two whole-rock analyses yielded a date of 119.8 ± 6.7 Ma (Table 3). Regression including all garnet fractions yielded an overlapping date of 117.9 ± 7.0 Ma, with a higher mean square of weighted deviation (MSWD). Linear regressions result in part from the spread of points (¹⁷⁶Lu/¹⁷⁷Hf >8) combined with small uncertainties, but may also represent complexities in the ages. Early Cretaceous crystallization falls within analytical error of 121.8 ± 3.2 Ma garnet growth in the upper Slate Creek drainage (hornblende gneiss; ¹⁷⁶Lu/¹⁷⁷Hf locality DW-02 of Wilford, 2012; Fig. 2A) and 124.3 ± 5.8 Ma growth in lower Squaw Creek drainage (biotite schist; ¹⁴⁷Sm/¹⁴⁴Nd locality 26 of McKay et al., 2017; Figs. 2A and 3B). Amphibolite-facies metamorphism/ductile deformation across the central SRSZ (e.g., locality JV-003) is also compatible with syntectonic magmatism reported along strike to the north (ca. 120–110 Ma tonalite-trondhjemite plutonic suite of Lee, 2004; McClelland and Oldow, 2007; Snee et al., 2007; 115.8 ± 1.3 Ma ²³⁸U/²⁰⁶Pb zircon locality 06RL401 of Schmidt et al., 2016a) and to the south (118 ± 5 Ma ²³⁸U/²⁰⁶Pb locality 83z9 of Manduca et al., 1993; 114.4 ± 2.2 Ma locality K92–8 of Unruh et al., 2008; 112.2 ± 1.6 Ma locality AM-4 of Mann, 2018; Fig. 2A).

In this section, we discuss late Mesozoic tectonic activity in west-central Idaho (post-150 Ma, Table 3), which followed ca. 170 Ma terrane accretion in the central Klamath Mountains (Rattlesnake Creek suture belt: Irwin and Wooden, 1999; figures 1 and 3 of Dickinson, 2008; Fig. 1A) and offshore(?) amalgamation of the Blue Mountains Province (ca. 160–155 Ma; Schwartz et al., 2011). The discussion proceeds in a time-sliced fashion, with emphasis placed on mesoscopic structural, coeval metamorphic, and intermittent magmatic activity east of the Heavens Gate fault. Our tectonic model linking the SRSZ and SFTB orogens is introduced (Fig. 14).

Along central portions of the Cordilleran margin (∼40°N–48°N: northern Sierra Nevada to northwest Washington; Fig. 1A), latest Jurassic–middle Cretaceous high-angle convergence (ca. 150–125 Ma; Figs. 14A and 14B) resulted in underthrusting of oceanic lithosphere beneath western North America (Engebretson et al., 1985; Burchfiel and Davis, 1975). Over this interval, east-west contractional strains likely accumulated along low-angle imbricate faults dipping eastward into the continental margin. Progressive thrust stacking and tectonic burial of underplated arc-volcanogenic rocks—Blue Mountains Province—conceivably setup the pressure-temperature-depth conditions necessary for achieving upper-amphibolite-facies metamorphism (∼7–11 kbar, ∼550–675 °C, ∼15–20 km; e.g., Zen, 1985; Selverstone et al., 1992). In this scenario, loading of buoyant oceanic crust (mature arc of Hamilton, 1988; Cloos, 1993; Stancin et al., 2016) was possibly associated with structurally overriding (undocumented thrust?) Neoproterozoic–Paleozoic passive-margin strata. Westward-vergent deformation across the incipient Salmon River suture was coeval with intracontinental contraction in the Sevier hinterland (e.g., Albion Mountains, ≤250 km east), including ca. 140 Ma thrust-loading of stratal assemblages above and west of the Basin-Elba fault (Miller, 1983; Kelly et al., 2015).

In the Riggins region of west-central Idaho, midcrustal metamorphic conditions are represented by ca. 144–135 Ma core garnet growth in the Pollock Mountain Amphibolite, i.e., structurally highest thrust sheet (locality 422 of Getty et al., 1993; ID03b/23 of McKay et al., 2017). Along our transect, high-strain synmetamorphic deformation of probable Early Cretaceous age is recorded along southern Heavens Gate Ridge (Fig. 3A). As described, ca. 136 Ma intrusive rocks (locality 06KG15; Figs. 5F and 5G) were deformed together with volcanic arc cover of the Wild Sheep Creek Formation (Figs. 5B, 5C, and 6C–6F). Based on regional mapping and U-Pb age data (northern Heavens Gate Ridge–Lucile area; Kauffman et al., 2014; Quarcoo and Gray, 2016), we interpret low-volume silicic magmatism (06KG15) as late syntectonic to local fabric development (Figs. 5C–5E) and possibly related to ca. 130 Ma imbricate thrust stacking/anatectic melting across western portions of the SRSZ (Figs. 2A, 7A, and 7B; figures 19, 20A, and 20B of Schmidt et al., 2016b). Tectonic activity in the northern Seven Devils/Heavens Gate area (eastern Wallowa terrane) overlapped with calc-alkaline magmatism and coeval dextral-transpressional deformation in northeast Oregon, e.g., ca. 140 Ma emplacement of the Pole Bridge and associated syntectonic plutons (Wallowa batholith; Fig. 1A; Vallier, 1995; Johnson et al., 1997, 2011; Žák et al., 2015).

Sustained high-angle plate convergence (Fig. 14B; ∼41°N–44°N), crustal thickening, and uplift/cooling west of the Sr_i 0.706 isopleth are recorded by synkinematic mineral growth in the Pollock Mountain Amphibolite (Fig. 2; ca. 128 Ma/119 Ma garnet/hornblende from sample locality 598 of Selverstone et al., 1992, Getty et al., 1993; ca. 124 Ma garnet ID23 of McKay et al., 2017) and structurally underlying rocks of the Riggins Group (Fig. 2; ca. 124 Ma/112 Ma garnet ID26/48 of McKay et al., 2017; ca. 120 Ma garnet JV-003 of this study; ca. 117 Ma hornblende R7/30 of Snee et al., 1995). In each case, metamorphism was accompanied by pervasive flattening deformation (Figs. 8D and 9C), linear-planar fabric development (outcrops 14–39; Table 1; Fig. 3B), and/or top-to-the-west porphyroblast rotation (Fig. 10D; see figure 15b of Gray et al., 2012; figure 4d of McKay et al., 2017). Dynamic recrystallization persisted through late Early Cretaceous time, as volcanogenic and overlying carbonate platform assemblages of the eastern Wallowa terrane—Riggins/Seven Devils Groups, Martin Bridge Formation—were progressively overridden by major westward-advancing thrust sheets: ca. 144–124 Ma thrusting of continental margin strata onto the Pollock Mountain Amphibolite (cryptic miogeoclinal structure proposed here), followed by ca. 124–112 Ma thrusting of the Pollock Mountain Amphibolite over the Riggins Group (Hamilton, 1969a; Aliberti, 1988; Selverstone et al., 1992; McKay et al., 2017; Figs. 14B and 14C).

Syntectonic mineral growth overlapped with emplacement of magmatic epidote-bearing tonalitic plutons east of the Pollock Mountain Amphibolite (Round Valley pluton of Zen, 1985; ca. 118–112 Ma Hazard Creek complex of Manduca et al., 1993; Unruh et al., 2008,Schmidt et al., 2016a; Mann, 2018). High-pressure granitic magmatism was coeval with east-west shortening across the arc-continent boundary, as evidenced by contractional structures at localities AM-6 (ca. 130 Ma), JV-003 (ca. 120 Ma), 83z9 (ca. 118 Ma), ID58 (ca. 116 Ma), K92–8 (ca. 114 Ma), AM-4 (ca. 112 Ma), and 83z14 (ca. 110–105 Ma). Structures are broadly time-transgressive and track ductile deformation through the Riggins Group, Pollock Mountain Amphibolite, and suture zone plutonic suite (Fig. 2A). By 112 Ma, high-angle ocean-continent (Farallon–North America) convergence had transitioned into right-oblique plate motion (Fig. 14C), thus adding a strike-slip component of deformation to the Cordilleran margin, i.e., 120–90 Ma interval of McClelland et al. (2000). As a consequence, boundary-normal shortening/crustal thickening (Pollock Mountain, Rapid River, Ahsahka thrust systems; Fig. 10A) and associated high-pressure silicic magmatism (Hazard Creek complex + equivalent units; e.g., Manduca et al., 1993) were accompanied by right-lateral transcurrent shearing (Figs. 4B, 4G, and 10E) and margin-parallel northward translation of the Blue Mountains block (Lund and Snee, 1988; Wyld and Wright, 2001; Giorgis et al., 2005).

Over this interval, dextral-oblique consumption of the Farallon plate (∼45° convergence angle: Giorgis and Tikoff, 2004; Fig. 14D) continued producing magma of the suture zone suite (ca. 112 Ma: locality AM-4 of Mann, 2018), and was possibly associated with early construction of the Idaho batholith (ca. 100 Ma metaluminous phase; Gaschnig et al., 2010). Mid-Cretaceous tectonic activity included emplacement/deformation of the Crevice and age-equivalent plutons (ca. 110–104 Ma; locality Cp02 of Gray et al., 2012; locality C of Kauffman et al., 2014), Looking Glass pluton/Payette River tonalite (ca. 93–85 Ma; e.g., LGp03 of Gray et al., 2012), and smaller calc-alkaline bodies dated herein (ca. 91 Ma: SHd-04; ca. 93–86 Ma: FC-05; Figs. 10A–10C, 12F, 12G, and 13C). Pervasive ductile deformation recorded in easternmost exposures of the Riggins Group (ca. 112 Ma; locality ID48 of McKay et al., 2017) may relate to westward-directed movement along the Pollock Mountain thrust and/or syntectonic magmatism in the western Hazard Creek complex (ca. 114–112 Ma; AM-4 of Mann, 2018; K92–8 of Unruh et al., 2008). If previous whole-rock geochemical correlations are correct (Manduca et al., 1993; Blake et al., 2009), then ca. 114–112 Ma silicic magmatism/contractional deformation in the Keating Ridge Gneiss, Spring Creek Gneiss, and/or Van Ridge Gneiss (Fig. 3B; units of Blake, 1991) temporally overlapped with thrust-related garnet growth in the Lightning Creek Schist (ID48; McKay et al., 2017). In this context, dextral transpression west of the Sr_i 0.706 isopleth was expressed by horizontal shortening in island-arc crust (westward-vergent folding, thrusting, fabric development; Figs. 7, 8, 9, and 10A–10D) and vertical lengthening/extrusion of adjacent intrusive rocks (eastern Van Ridge Gneiss; Fig. 11). East of the Sr_i 0.706 isopleth, right-oblique convergence was accompanied by ca. 105–90 Ma extrusion (localities Cp02, LGp03, FC-05 granitoids; Figs. 2A, 3B, and 12G) and westward-vergent folding, and tectonite fabric development (western Laurentia: Kelly Mountain Schist + correlative units; Figs. 12A–12D; Blake, 1991; Lund and Snee, 1988).

As described, Laurentian metasedimentary rocks were deformed together with Cretaceous granitoids east of the arc-continent boundary (orthogneiss screens; Fig. 12E). Local parallelism of solid-state fabric (S1: southeastern Crevice pluton/Kelly Mountain Schist; cross-section D–D′ in Fig. 3B) supports syn– to post–104 Ma penetrative deformation (Gray et al., 2012; West Mountain locality 10NB376 of Braudy et al., 2017). Timing constraints are compatible with ca. 111–105 Ma structures documented north and south of our transect (e.g., locality C of Kauffman et al., 2014; 83z14/99MG of Manduca et al., 1993; Giorgis et al., 2008; Fig. 2A). Age-equivalent calc-alkaline magmatism and east-west contractional deformation in the Partridge Creek Gneiss/Little Goose Creek complex (Blake et al., 2009, 2016) would extend ca. 111–104 Ma structures westward into the Van Ridge Gneiss/Hazard Creek complex (Figs. 2A and 11), i.e., easternmost plutons of the Blue Mountains Province. Subsequent magmatism and deformation within the Idaho batholith border zone—syntectonic emplacement ca. 93–85 Ma—are recorded by the Looking Glass pluton (Fig. 3B) and U-Pb sample localities of this study (Figs. 10A and 12F).

Based on the available age constraints and unit correlations reported from the Salmon River Canyon, we propose that LS tectonite fabric (strain) accumulated ca. 115–86 Ma east of the arc-continent boundary (Sr_i 0.706 isopleth). Otherwise stated, ductile deformation persisted between the onset of Partridge Creek Gneiss/Little Goose Creek complex magmatism (ca. 115 Ma; upper age bound of Manduca et al., 1993; 110–120 Ma tectonism of McClelland et al., 2000; Lee, 2004; Snee et al., 2007; Kauffman et al., 2014) and the end of intrusive activity recorded by locality FC-05 (ca. 86 Ma: lower age-bound of present study; Fig. 13). This interval overlapped with tectonism west of the Sr_i 0.706 isopleth, as supported by ca. 112 Ma synkinematic garnet growth (Riggins Group locality ID48; McKay et al., 2017) and ca. 118–100 Ma magmatism in the Hazard Creek–Little Goose Creek complexes, i.e., age-equivalent portions of the Keating Ridge Gneiss, Spring Creek Gneiss, and/or Van Ridge Gneiss (unit correlations of Blake et al., 2009, 2016). West of locality ID48, late Early Cretaceous shortening was accommodated by east-dipping imbricate structures of the Rapid River thrust–Morrison Ridge fault system (small-scale duplexing in Schmidt et al., 2016b; Shingle Creek structures of Onasch, 1977), which crosscut ca. 113–111 Ma garnet-bearing metamorphic tectonites (Wilford, 2012; McKay et al., 2017; see truncated isograd of Hamilton, 1963b, 1969a; figure 2 of Gray et al., 2012) and ca. 118–107 Ma hornblende-defined fabric in the upper Riggins Group (Squaw Creek Schist, Berg Creek Amphibolite; Figs. 2A, 3B).

Given the orogen-scale continuity of LS tectonite fabric (∼25–40 km across strike; Fig. 2A; Table 1; cf. continuity concept of Sander, 1930), intensity of ductile deformation (Figs. 5C, 5D, 5G, 6D, 8C, 9C, 10A, 11B, 12B, and 12G), and involvement of island-arc–intrusive rock–continental margin assemblages along our transect, we attribute the formation of ca. 136–91 Ma synmetamorphic structures (Table 4) to arc-continent collision and associated evolution of the dextral-transpressional Salmon River suture. Although the age and extent of deformation differ (see chronology of Snee et al., 1995; map compilation of Lund, 2004), our interpretation of fabric (origin) is consistent with Lund and Snee (1988) and others supporting post–160 Ma terrane accretion in west-central Idaho. More recent workers suggest otherwise.

In the model of LaMaskin et al. (2015), Early Cretaceous (ca. 144–128 Ma) upper-amphibolite-facies metamorphism, thrust-related crustal thickening, and pervasive ductile deformation (Hamilton, 1963a; Onasch, 1977, 1987; Lund and Snee, 1988; Blake, 1991; Selverstone et al., 1992; Getty et al., 1993; Snee et al., 1995; Gray et al., 2012; McKay et al., 2017) postdates terrane accretion and relates to noncollisional tectonism of the Sevier orogeny (Armstrong, 1968; Burchfiel, 1980). Tectonic activity occurred in response to a through-going Andean-type subduction zone system emerging along the Cordilleran margin. The geochemical analysis of Late Jurassic volcanogenic rocks in the Pittsburg Landing area (Fig. 2A; ca. 160–150 Ma Coon Hollow Formation) led authors to conclude that arc-continent collision is not recorded by penetrative structures in the Riggins region (outcrops 5–38; Fig. 3). According to LaMaskin et al. (2015), exposed rocks of the Wallowa terrane occupied a fluvial to deep-marine setting during pre–160 Ma terrane accretion. Because the sedimentary assemblages at Pittsburg Landing have not produced detrital zircon of pre–late Paleozoic age (3% of grains older than 260 Ma; LaMaskin et al., 2015) and were sourced from crystalline basement/volcanic cover rocks of the Wallowa arc (Walker, 1986; Vallier, 1995; Schmidt et al., 2016a, 2016b; Kurz et al., 2017), we question terrane accretion prior to 160 Ma sedimentation (Coon Hollow basin [CHB] of Figs. 2A and 14). Furthermore, Middle–Late Triassic rocks are strongly deformed in northwest-vergent fold nappes and imbricate faults (Klopton Creek fold-thrust system; White and Vallier, 1994; Kauffman et al., 2014) reminiscent of contractional structures across Heavens Gate Ridge (Fig. 3A, section A′–A″; Fig. 7B). More importantly, the sedimentary response supporting pre–160 Ma accretion is simply not recorded by this fluvial to deep-marine depocenter (Morrison, 1963; Vallier, 1977; Goldstrand, 1987; Lewis et al., 2014). For age-equivalent and older arc basins in the western U.S. Cordillera that do record proximity to—and host sediment from—the Laurentian continental margin (e.g., northern Sierra Nevada and Klamath Mountains), readers are referred to LaMaskin (2012).

In our view, the Coon Hollow basin (Pittsburg Landing area) contains no direct evidence of sediment sourced outside of the Wallowa terrane (mud rock ε_Nd = 1.9–5.0; LaMaskin et al., 2015). The absence of a clear North American source (i.e., detrital zircon population) is difficult to reconcile with Coon Hollow sedimentation subsequent to terrane accretion (versus a precollisional marine setting). In a postcollisional scenario, basin fill would include mixed detritus shed from the collisional orogen (∼25–40-km-wide SRSZ) and breakdown of shielding topographic barriers, i.e., clastic material sourced from the island-arc assemblage and ancestral North America (e.g., Boghossian et al., 1996). Given this basin’s proximity to the Sr_i 0.706 isopleth (Laurentia) and the Idaho batholith (∼35 km east; Fig. 2A), an influx of terrigenous sediment would be expected in response to arc accretion, midcrustal exhumation, and unroofing of proximal continental margin strata (Lund et al., 2003; Giorgis et al., 2008; Gaschnig et al., 2010). Amato et al. (2013) presented a similar argument for terrane accretion in southern Alaska. Detrital zircon dating in the Mesozoic Chugach accretionary complex (e.g., Berg et al., 1972) suggested that Wrangellia was isolated from Laurentian sources until middle Late Cretaceous time (ca. 89–85 Ma). According to Amato et al. (2013), the lack of continentally derived sediment in the oldest accretionary unit (ca. 164 Ma Potter Creek flysch assemblage; “mesomélange” of Amato and Pavlis, 2010) is incompatible with a mid-Jurassic collision event (McClelland et al., 1992a) because detritus shed from that orogen should have been carried into the coeval trench. Unlike the outboard Chugach accretionary complex (shielded by Peninsular-Alexander-Wrangellia terranes), the Coon Hollow basin occupied an inboard (intra-arc) position and was protected from western Laurentia by the easternmost Wallowa terrane. In this context, postcollisional sedimentation in west-central Idaho (160–150 Ma interval of LaMaskin et al., 2015) would have transected a narrow (∼35 km-wide), north-south tract of pre-Cenozoic topography separating the Pittsburg Landing area and ancestral North America.

Orogenic components in west-central Idaho are characteristic of major arc-arc and arc-continent collision zones in the North American Cordillera and other well-studied accretionary orogens (Moores, 1970, 1998; Brown and Ryan, 2011; Xiao and Santosh, 2014), e.g., Trans-European suture zone/Avalonia composite arc accretion (Pharoah, 1999; Nance et al., 2002), Kohistan arc obduction/pre-Oligocene collision in northern Pakistan (Bard, 1983; Khan et al., 2009), and Grampian-Taconic orogeny of western Ireland–New England/episodic island-arc–ophiolite accretion (van Staal et al., 2007; Hollis et al., 2012). As described for these orogens, lithotectonic assemblages in west-central Idaho record the products of oblique plate convergence culminating with post–160 Ma arc-continent collision: thrust-induced regional metamorphism, intermittent high-pressure silicic magmatism, and pervasive ductile deformation. In this context, the Seven Devils–Riggins–Salmon River Canyon transect (Figs. 2 and 3; Gray, 2013) provides a basis for comparison with other high-strain zones involving terrane accretion.

The Omineca belt of southeastern British Columbia overlaps the lithospheric boundary separating easternmost allochthonous terranes of the Canadian Cordillera and the Laurentian continental margin (Monger et al., 1972, 1982, 1994; Archibald et al., 1983; Wheeler and McFeely, 1991). Positioned along strike of the SRSZ (Fig. 1A), the belt is characterized by widespread transpressional deformation, granitic magmatism, and greenschist- to upper-amphibolite-facies metamorphism (Gabrielse and Reesor, 1974; Price, 1981; Avé Lallemant and Oldow, 1988; Carr, 1991, 1992; Parrish, 1995). In the south (∼51°N), the Selkirk fan records multiple generations of structures (Brown and Tippett, 1978; Evenchick et al., 2007). Early phases of fan development (ca. 172–163 Ma; Gibson et al., 2005, 2008) are represented by kilometer-scale bivergent contractional structures (e.g., Carnes nappe; Brown and Lane, 1988), outcrop-scale tight-to-isoclinal folds, and high-strain transposition foliation defined by peak metamorphic mineral alignment. Early fabric elements (D1 and D2) are overprinted by lower-strain coaxial folds (D3) constrained to ca. 104–84 Ma (U-Th-Pb zircon/monazite; Gibson et al., 2008). According to Monger et al. (1982), regional synmetamorphic structures evolved in response to the progressive collision of composite arc terranes with ancestral North America (Crawford et al., 1987; Rubin et al., 1990; Journeay and Friedman, 1993; Chardon et al., 1999; McClelland and Mattinson, 2000; Gehrels et al., 1990, 2009; Simony and Carr, 2011).

Dextral-transpressional deformation west of the Omineca belt is recorded along the ∼1500 km-long, arc-arc collisional boundary separating the Insular and Intermontane composite terranes (Coast plutonic complex of Monger et al., 1982; Coast Mountains orogen of Gehrels et al., 1990; see also Journeay and Friedman, 1993). As described by Crawford et al. (1987), syntectonic magmatism along this east-west compressional regime involved underthrusting of the Alexander-Wrangellia (Insular) composite belt beneath the Intermontane belt (Stikinia/Yukon-Tanana terranes). West-directed folding/thrusting (Sumdum-Fanshaw fault system; McClelland et al., 1992b) was accompanied by emplacement of magmatic epidote-bearing plutons at mid- to lower-crustal levels (e.g., ca. 100 Ma syntectonic Ecstall pluton; Hutchison, 1982; Zen, 1985). According to McClelland et al. (2000), metamorphism increases from subgreenschist- to upper-amphibolite-facies conditions across the contractional belt (west to east). Mesoscopic structures associated with regional-scale thrust faults and folds (nappes) are dominated by east-dipping transposition fabric observed in arc-supracrustal rocks (Crawford and Hollister, 1982; Stowell and Crawford, 2000). When taken together, the overall structural style (imbricate thrusts, tight-to-isoclinal folds, linear-planar fabrics), kinematics of contractional deformation (westward-vergent), metamorphic trends (easterly increase in grade), and deep-seated silicic magmatism (epidote-bearing plutons) are remarkably similar to orogenic components of the SRSZ. In both regions, Early Cretaceous (120–90 Ma) calc-alkaline magmatism and prograde metamorphism record pronounced crustal thickening historically attributed to terrane accretion (Davis et al., 1978; Monger et al., 1972, 1982, 1994; Zen, 1985; Crawford et al., 1987; Blake, 1991; Selverstone et al., 1992; Getty et al., 1993; Manduca et al., 1992; Avé Lallemant, 1995; Snee et al., 1995; McClelland and Mattinson, 2000; Stowell and Crawford, 2000; Gray et al., 2012; McKay et al., 2017).

The North Cascade Mountain system and magmatic arc of north-central Washington mark the southern continuation of the Coast Mountains orogen, i.e., Insular-Intermontane suture of Whitney and McGroder (1989) and North Cascades–southeastern Coast belt of Hurlow (1993), Miller and Paterson (2001), and Gehrels et al. (2009). Late Paleozoic to Cretaceous arc assemblages occupying the Cascades metamorphic core (Fig. 1A; Chelan Mountains, Nason, and Swakane terranes; Tabor et al., 1989; Journeay and Friedman, 1993; Wyld et al., 2006) were intruded ca. 115–85 Ma synchronous with regional contraction (e.g., transpressional Pasayten fault zone of Hurlow, 1993; Miller et al., 2016). Over this interval, the northern Cascades experienced westward-directed thrusting, dextral strike-slip displacement, and pervasive ductile deformation together with amphibolite-facies metamorphism (Brandon et al., 1988; Hurlow, 1993; Paterson and Miller, 1998). According to Miller and Paterson (2001), the magmatic arc and core regions of the orogen reached a crustal thickness of ≥55 km (cf. SRSZ: >50 km; Zen and Hammarstrom, 1984; Selverstone et al., 1992). Mid-Cretaceous tectonism is interpreted to record major intra-arc shortening and final suturing of the Insular composite terrane to western Laurentia (Whitney and McGroder, 1989; McGroder, 1991; van der Heyden, 1992; Dickinson, 2004).

Western Ireland is possibly the most well-studied region exposing the late Early–Middle Ordovician Grampian orogeny (Fig. 15; Lambert and McKerrow, 1976; Soper et al., 1999; Dewey and Ryan, 2015), which is broadly equivalent to the Taconic arc-continent collisional event in eastern North America (Dewey and Shackleton, 1984). In the North Mayo tectonic zone, Neoproterozoic strata of the Dalradian Supergroup (Condon and Prave, 2000) are interpreted to record ca. 720–595 Ma passive-margin sedimentation related to the rifting/breakup of Rodinia (Strachan and Holdsworth, 2000). To the south, continentally derived mélange and dismembered ophiolite (Cambrian–Ordovician Clew Bay Complex; e.g., Dewey and Mange, 1999) combine to form the Northwestern composite terrane (e.g., Murphy et al., 1991). High-pressure metasedimentary rocks of the Dalradian and Clew Bay assemblages are separated by a strong magnetic lineament (Fair Head–Clew Bay Line; e.g., Max et al., 1983) marking the westward continuation of the Highland Boundary fault system (Scotland). In both Britain and Ireland, this northeast-striking zone of contractional/strike-slip deformation overlaps the boundary separating autochthonous Laurentian margin strata and accreted island-arc rocks of the Grampian orogeny (e.g., van Staal et al., 1998). Polyphase structural mapping in the Irish Caledonides shows that the Dalradian Supergroup and Clew Bay Complex shared a Middle Ordovician tectonic history. According to Chew (2003), contractional deformation is represented by shallow east-plunging isoclinal folds (nappes) and tectonite fabric (S1–L1) recording pronounced flattening/stretching strains, i.e., triaxially deformed lithic clasts (cf. Heavens Gate; Figs. 3A, 6, 7, and 8D). Regionally developed synmetamorphic structures are attributed to oblique plate convergence—northwest-directed sinistral transpression—and collision involving ancient volcanic arcs of the Iapetus Ocean (Fig. 15, inset) and Laurentian continental margin (Harris, 1995; Ryan and Dewey, 2011).

In their pioneering study, Getty et al. (1993) attributed ca. 144 Ma garnet growth in west-central Idaho (locality 422; Fig. 2A) to outboard amalgamation of the Blue Mountains Province or its onset of suturing to ancestral North America. McKay et al. (2017) refined this date (ca. 141 Ma; locality ID03a) and reported additional Sm-Nd data (ca. 135–112 Ma crystallization) from the Pollock Mountain Amphibolite and Riggins Group (localities ID23 and ID48, respectively). These workers interpreted garnet growth as recording progressive thrust stacking associated with prolonged arc-continent collision (∼30 m.y.; McKay et al., 2017). In this context, suturing was under way by ca. 141 Ma and followed closure of an intervening oceanic tract (syncollisional flysch basin of Pavlis, 1982; ca. 155–50 Ma suture of Sigloch and Mihalynuk, 2017) separating the Wallowa/Baker/Izee/Olds Ferry terranes and western margin of Laurentia. We interpret long-lived tectonic activity in the SRSZ (ca. 145–90 Ma) as recording orthogonal to right-oblique terrane accretion, margin-parallel translation, and clockwise rotation of the Blue Mountains block. The onset of suturing followed ca. 160 Ma amalgamation (Schwartz et al., 2010, 2011) and coincided with final closing of the basin (minimum age ca. 141 Ma); subsequent metamorphism, contractional deformation, and strike-slip displacement were associated with sustained arc-continent collision/convergence between the Farallon and North American plates. Syntectonic magmatism dominated late-stage, accretion-related activity in the SRSZ (Manduca et al., 1992, 1993; Giorgis et al., 2005, 2008; Figs. 2A, 10A, 10C, 12F, and 12G) and was focused along the thermally weakened arc–continent collisional boundary (Tikoff et al., 2001; Giorgis et al., 2006).

Basin closure by ca. 141 Ma suggests that outboard terranes (Insular belt) accreted from south to north between Idaho and Alaska (>45°N; present coordinates) over latest Jurassic to Paleocene time (Getty et al., 1993; Ridgway et al., 2002; Gehrels et al., 2009). In this scenario, an ocean tract of indeterminate width—narrow backarc or broader basin?—closed incrementally northward as transpressional terranes were successively attached to the Laurentian margin. Initial collision of the Blue Mountains block occurred east-northeast of the Sierra Nevada–Franciscan subduction zone system (Fig. 1A; Hamilton, 1969b; Anczkiewicz et al., 2004; suture belt of Dickinson, 2008). Final accretion and entrapment of amalgamated terranes took place in the right-angle bend (Sr_i 0.706 isopleth)/Syringa embayment region of north-central Idaho (Fig. 1A; Yates, 1968; Schmidt et al., 2016a). Timing estimates for terrane accretion in west-central Idaho (McKay et al., 2017; this study), the Coast Mountains orogen (Monger et al., 1982, 1994; Miller et al., 2016), and central/eastern Alaska Range (Trop and Ridgway, 2007) are compatible with basin closure and suturing in the south (Idaho: ca. 144–141 Ma garnet growth) followed by more northern events over time (Alaska: end of suturing ca. 60 Ma; e.g., Ridgway et al., 2002). Overlapping ages of middle Cretaceous tectonism (115–90 Ma) in west-central Idaho (SRSZ), north-central Washington (Cascades), and westernmost Canada (Coast Mountains) support north-south fragmentation and subsequent addition of the Wrangellian composite to the Cordilleran collage (Coney et al., 1980). Our correlation of the Wallowa terrane with Wrangellia (Jones et al., 1977; Dickinson, 2004; Kurz et al., 2017) is based on similarities in structural style, rock associations, and the timing of accretion along the SRSZ–northern Cascades–Coast Mountains orogen (e.g., Selverstone et al., 1992; Vallier, 1995; Crawford et al., 1987; Whitney and McGroder, 1989; Miller and Paterson, 2001; McKay et al., 2017). Having established along-strike connections with the northern Cordillera (>46°N; Fig. 1A), and compared orogenic components with western Ireland (Fig. 15), we call attention to areas east of the SRSZ and consider terrane accretion in the context of age-equivalent mountain building on the continental interior (cf. Burchfiel, 1980; Edelman, 1992; DeCelles, 2004).

Areas extending from northwestern Utah through northeastern Nevada into south-central Idaho (∼38°N–43°N, ∼113°W–116°W) occupy the Sevier hinterland region of the Cordilleran orogen (Fig. 1A; e.g., Misch, 1960; Allmendinger et al., 1984). While the location of initial thrusting is unclear (southern Idaho—Royse et al., 1975; Miller, 1980; northeast Nevada—Hodges et al., 1992; Hudec, 1992; west-central Utah—Lawton et al., 1997), compressional deformation propagated eastward with local hindward-directed and out-of-sequence events over late Mesozoic time (Morley, 1988; DeCelles, 1994, 2004; Long et al., 2014). Kinematic reconstructions reported from the Sevier type area in west-central Utah (Canyon and Pavant Ranges; Fig. 1A) indicate east-west shortening exceeded 160 km (DeCelles and Coogan, 2006). Displacement estimates are lower in the north (∼100–160 km: eastern Idaho–western Wyoming–northern Utah salient [e.g., Royse, 1993]; ∼50 km: southwest Montana–southeast Idaho segment [e.g., Skipp, 1988]), but exceed 200 km along the northwestern Montana–Alberta, Canada, segment (Price et al., 2000). Here, we discuss the timing of tectonic activity across the SFTB (hinterland-foreland regions; ∼116°W–113°W) and recall coeval accretion-related events in the SRSZ (west-central Idaho; ∼117°W–116°W).

Latest Jurassic–middle Early Cretaceous contractional deformation, greenschist- to upper-amphibolite-facies metamorphism, and granitic magmatism are recorded in northeastern Nevada (Pilot, Ruby, East Humboldt Ranges; e.g., Miller, 1984; Hudec and Wright, 1990), northwestern Utah (Silver Islands, Newfoundland Mountains; e.g., Miller and Allmendinger, 1991; Allmendinger and Jordan, 1984), east-central Nevada/west-central Utah (Snake Range; e.g., Miller et al., 1988), and south-central Idaho (Albion Mountains; Snoke and Miller, 1988; Wells et al., 1997). While many of these areas (∼41°N–42°N) record Cenozoic extensional deformation associated with core-complex development (Coney and Harms, 1984; MacCready et al., 1997), an earlier history of east-west contraction is also recognized (Fig. 1A; ca. 150–145 Ma metamorphic tectonites, central Ruby Mountains [e.g., Hudec, 1992]; east-southeast–vergent folds/penetrative fabric, northern Albion Mountains [e.g., Miller, 1980][DeCelles, 2004]). Smith et al. (1993) noted the coincidence of ca. 165–150 Ma contraction with an increased rate of plate convergence (Farallon–North America, >8 cm/yr; Gordon et al., 1984; Engebretson et al., 1985), as compared to the ca. 145–130 Ma interval (∼5 cm/yr; Humphreys, 1995; Madsen et al., 2006). This apparent decrease in plate convergence rate overlapped in time with westward-directed thrusting/crustal thickening/core garnet growth (Pollock Mountain Amphibolite; Getty et al., 1993), pronounced flattening deformation (Wild Sheep Creek Formation; Fig. 6; Aliberti, 1988), and pervasive linear-planar fabric development (figure 3A of McKay et al., 2017) across western portions of the SRSZ (Figs. 2A, 14A, and 14B). Coeval calc-alkaline magmatism and contractional deformation are recorded by ca. 145–130 Ma stocks, sills, and satellite plutons emplaced into the eastern Wallowa terrane (Heavens Gate Ridge U-Pb zircon localities D and 06KG15 of Kauffman et al., 2014; this study; Figs. 5 and 13; Tables 1 and 4), magmatic epidote-bearing plutons south of Riggins (Armstrong et al., 1977; Zen, 1985; Manduca et al., 1993; Unruh et al., 2008), and syntectonic granitoids in northeastern Oregon (Johnson et al., 1997, 2011; Žák et al., 2012, 2015).

According to DeCelles (2004), evidence of Early Cretaceous (142–112 Ma) magmatism, metamorphism, and contractional deformation is sparsely scattered across the Sevier hinterland; geochronological data constrain eastward displacement on aerially extensive thrust sheets to ca. 145–132 Ma (Canyon Range thrust [apatite fission track]—Stockli et al., 2001; Willard thrust [⁴⁰Ar/³⁹Ar muscovite]—Yonkee et al., 1989, [fission track]—Burtner and Nigrini, 1994; Paris thrust [K-Ar illite]—Burtner and Nigrini, 1994; Basin-Elba thrust [⁴⁰Ar/³⁹Ar muscovite]—Wells et al., 2008, [¹⁷⁶Lu/¹⁷⁷Hf garnet]—Kelly et al., 2015). Wyld and Wright (2001) emphasized the lack of subduction-related magmatism and regional shortening ca. 140–120 Ma, as compared to Late Jurassic (Camilleri et al., 1997) and post–120 Ma time (e.g., Burchfiel et al., 1992). Orogenic quiescence is explained by (1) extension in the Cordilleran magmatic arc and forearc regions (e.g., Smith et al., 1993), and (2) partitioning of contractional deformation in the arc edifice and frontal portions of the orogen (Wyld and Wright, 2001). In the latter, contractional strains accumulated along major north-northeast–striking transpressional belts (e.g., western Nevada shear zone; Fig. 1A) and more eastern thin-skinned structures of the SFTB (DeCelles, 2004), i.e., stage 1 deformation of Yonkee and Weil (2015). We attribute tectonism in the Western Nevada shear zone (northern segment, ca. 140–130 Ma) to dextral-oblique collision and translation of the Blue Mountains Province.

Middle Cretaceous dextral transpression in the Cordilleran hinterland (pre–115 Ma to ca. 108 Ma: Pueblo Mountains, southeast Oregon; Wyld and Wright, 2001) coincided with emplacement of the Meade thrust sheet (onset ca. 115 Ma: DeCelles et al., 1993; Mitra, 1997) and layer-parallel shortening in the foreland region (e.g., pressure-solution cleavage; Fig. 1A, cross-section C–C′; Protzman and Mitra, 1990). Deformation was contemporaneous with west-southwest–directed thrusting in the Orofino-Ahsahka belt (onset ca. 116 Ma: Schmidt et al., 2016a; Yates, 1968; Strayer et al., 1989; McClelland and Oldow, 2007) and displacement on the Pollock Mountain–Rapid River–Morrison Ridge fault systems in the SRSZ (ca. 116–111 Ma; Figs. 2, 3, 14B, and 14C; Hamilton, 1963b, 1969a; Aliberti, 1988; Gray et al., 2012; McKay et al., 2017). Subsequent activity in the SFTB included ca. 100–85 Ma slip along the Medicine Lodge, Tendoy, and Sapphire thrusts (southwest Montana–east-central Idaho segment; Fig. 1A, section A–A′; Schmitt et al., 1995) and sinistral transpression across the northwest-southeast–trending Lewis and Clark Line (Fig. 1A, section B–B′; e.g., flexural-slip fabric of Sears et al., 2004). Late Early Cretaceous tectonism north of the Snake River Plain (SFTB: ca. 100–85 Ma; Idaho and Montana) overlapped with contraction in the Orofino-Ahsahka belt (active through ca. 90 Ma: McClelland and Oldow, 2007; Lewis et al., 2014; Schmidt et al., 2016a) and ductile deformation in the Coolwater culmination (SRSZ: onset ca. 100 Ma; “orogenic welt” of Lund et al., 2008; Fig. 14D).

As the angle of dextral-oblique plate convergence steadily increased (ca. 120–90 Ma; ∼30°–60°; see inset Fig. 14D; Engebretson et al., 1985; Giorgis and Tikoff, 2004), and the magnitude of boundary-normal shortening decreased (>100–50 km; Skipp, 1988; Royse, 1993; DeCelles, 2004), contractional deformation shifted from north-south–striking structures along the Idaho-Wyoming-Utah salient to northwest-southeast–striking thrusts (e.g., Meade and Medicine Lodge) and sinistral-transpressional elements (Lewis and Clark Line) of southwestern Montana (Fig. 1A). Based on the age relations (northward diachroneity) and attitudes of structures outlined above, this transition was possibly related to early high-angle arc-continent collision (east-west compression ca. 150–125 Ma; ∼41°N–43°N) followed by northward translation and clockwise rotation (dextral transpression ca. 125–90 Ma; ∼43°N–46°N) of the Blue Mountains block (ca. 150–90 Ma accretion-related tectonism in the SRSZ [Lund and Snee, 1988; Selverstone et al., 1992; Lund et al., 2008; cf. Fossen and Tikoff, 1998]). In our tentative model, age-equivalent tectonic activity in the SFTB (Fig. 1; Table 4; ca. 140–90 Ma; ∼41°N–46°N) was associated with oceanic-continental (Farallon–North American) plate coupling enhanced by terrane accretion: subduction zone jamming (positive buoyancy; Cloos, 1993; Ranalli et al., 2000), oceanward-jumping (plate reorganization; Hamilton, 1988), and resultant end-loading of the Laurentian continental margin (Oldow et al., 1990; Yonkee and Weil, 2015; Blakey and Ranney, 2018).

In oblique subduction settings where thick oceanic tracts are attached to the structurally overriding plate, addition of buoyant crust is balanced by shortening in the colliding mass and continental margin assemblage (e.g., Quesnellia arc/Neoproterozoic strata of the Omineca belt; Monger et al., 1982; Scholl et al., 1986; Ross, 1991). As subduction/accretion continue, terrane fragments are progressively transferred to the overriding (accommodating) plate, satisfying mass conservation/balance requirements, i.e., hinterland displacement within newly conjoined crustal units (Struik, 1988; Oldow et al., 1990; Wilson, 1990). Tectonic stresses are transmitted away from the subduction zone, across an emergent arc-continent collision zone (Brown and Ryan, 2011), into the continental interior (Dickinson et al., 1978). Sustained high-angle convergence results in widespread transpressional deformation, thrust-related crustal thickening, and intermittent magmatism across the contractional orogen (e.g., Trans-European suture zone; Pharoah, 1999). Given that the SRSZ and SFTB are spatially (117°W–113°W latitude; Fig. 1), temporally (ca. 145–90 Ma tectonism; Figs. 2 and 14; Table 4), and kinematically linked (shared basal-décollement; cf. Bally, 1984; Fig. 16), a causal relation is considered here.

Most geologists accept that formation of the Cordilleran foreland fold-and-thrust belt (Sevier orogen; Armstrong, 1968) followed accretion of composite terranes in the absence of major collisional events (Allmendinger and Jordan, 1984; Burchfiel et al., 1992; DeCelles, 1994; Taylor et al., 2000; English and Johnston 2004; ribbon continent alternative of Johnston, 2008). Models favoring noncollisional mountain building include (1) rapid oceanic-continental plate convergence with a thermally weakened back-arc region (Hyndman et al., 2005), (2) flat-slab subduction of a seafloor spreading center, seamount, and/or oceanic plateau (Bird, 1988; Murphy et al., 2003), and (3) dextral strike-slip movement along the Northern Rocky Mountain–Tintina trench fault system with an eastward component of displacement (Price and Carmichael, 1986). Regardless of the mechanism, large-magnitude crustal shortening across the central SFTB (Fig. 1A) was coeval with collision-related tectonism in the SRSZ (Figs. 2A and 3). Although their structural styles and conditions of deformation differ (thick-skinned mid/lower-crustal levels in SRSZ versus thin-skinned, shallow/midcrustal levels in SFTB; cf. hinterland-foreland of Mexican orogeny; Fitz-Díaz et al., 2017), overlapping ages of contraction argue for kinematic coordination of respective structures (southern Canadian Rocky Mountains [Brown et al., 1992]; northern U.S. Rocky Mountains [McClelland and Oldow, 2004]). Moreover, seismic reflection data collected across northern Idaho and Washington (∼33–35 km depth, 48.5°N, COCORP seismic line WTF-82–1; Fig. 1A) show prominent west-dipping sequences (crust-penetrating thrusts) linking accretion-related shortening with contractional structures of western Montana (Potter et al., 1986, 1987). Accordingly, we propose that long-lived terrane accretion, translation, and rotation in the Cordilleran hinterland—i.e., tectonic evolution of the SRSZ—drove crustal shortening eastward into the foreland region.

In our model (Fig. 14), contact between arc-volcanogenic rocks of the Blue Mountains Province and passive-margin sedimentary strata initiated north of the Mesozoic marine province (Speed, 1978) near the juncture of Idaho, Oregon, and Nevada (Fig. 1A, schematic cross-section C–C′). We suggest that northern portions of the Western Nevada shear zone (southern counterpart of SRSZ?; figure 1 of Wyld and Wright, 2001) accommodated terrane accretion and served in transmitting early eastward-directed tectonic stresses. Late Mesozoic contractional structures of the SRSZ and SFTB (bivergent thrust faults, tight-to-isoclinal folds, tectonite fabrics; Figs. 1–3; Tables 1 and 4) evolved together as arc terranes impacted (∼41°N: high-angle collision; Fig. 14A), migrated northward along (dextral translation/clockwise rotation; Figs. 14B and 14C), and were ultimately incorporated into the continental margin (46°N, Syringa embayment; Fig. 14D; Schmidt et al., 2016a). Large-scale horizontal shortening of miogeoclinal strata was compensated (volumetrically mass-balanced) by the addition of buoyant oceanic crust (Scholl et al., 1986; Cloos, 1993; Stanciu et al., 2016) to the leading edge of ancestral North America. Hinterland tectonites that equilibrated under midcrustal metamorphic conditions (SRSZ [Selverstone et al., 1992]; SFTB [Kelly et al., 2015]) were kinematically linked to upper-crustal foreland structures along a gentle westward-dipping basal-décollement system underlying the Cordilleran orogen (seismic profiling—Bally et al., 1966; Allmendinger et al., 1987; Cook et al., 1992; orogenic float concept—Oldow et al., 1989, 1990; hinterland-foreland linkage—Brown et al., 1992). In this framework, latest Jurassic (ca. 145 Ma) to late Early Cretaceous (ca. 90 Ma) composite terrane accretion, translation, and rotation in the Cordilleran hinterland played a dynamic role in transmitting displacements across the SRSZ into the SFTB (∼41°N–46°N, ∼117°W–113°W; Gray, 2016).

Structural and geochronological studies conducted in west-central Idaho indicate cross-orogen linkages between tectonic elements of the western U.S. Cordillera: Salmon River suture zone and Sevier fold-and-thrust belt. A causal link with suturing of volcanic arc terranes (Blue Mountains Province) to ancestral North America (western Laurentia) is proposed for development of east-west contractional structures in the Sevier orogen: mid- to upper-crustal thrust faults, regional folds, and related fabric elements between the modern latitudes of Riggins, Idaho (∼45°N), and Winnemucca, Nevada (∼41°N). Early structures associated with late Mesozoic island-arc–continent collision include the Western Nevada shear zone (southeast Oregon), Willard–Basin-Elba–Paris–Meade thrust systems of southeastern Idaho, and midcrustal metamorphic tectonites of northeastern Nevada: contractional structures south of the Snake River Plain. Later structures include the Medicine Lodge thrust, Sapphire thrust, and sinistral-transpressional elements along the Lewis and Clark Line north of the Snake River Plain (southwestern Montana, east-central Idaho). In response to right-oblique oceanic-continental plate convergence, subduction zone jamming, terrane transfer, and accretion (end-loading of Laurentian margin), compressional deformation propagated >250 km eastward across the SRSZ (hinterland: ∼117°W–116°W) into the SFTB (foreland: ∼115°W–113°W). This long-lived history of temporally overlapping (ca. 145–90 Ma tectonic activity) and kinematically linked (shared basal décollement) hinterland-foreland contraction followed pre–145 Ma collapse of fringing arc assemblages and stabilization of the North American margin. In our view, orogenic components in west-central Idaho—pervasive ductile deformation (>25-km-wide tectonite belt), pronounced thrust-related crustal thickening (>50 km: Pollock Mountain–Rapid River–Morrison Ridge–Heavens Gate fault systems), and high-pressure syntectonic magmatism (∼8–11 kbar: suture zone plutons)—fit a collisional model explanation for intracontinental mountain building in the Cordillera of western North America (∼41°N–46°N). Late Mesozoic terrane accretion supports inclusion of the Wallowa arc with Wrangellia and implies progressive south-to-north ocean closure between Idaho and Alaska (ca. 145–60 Ma).

The first author thanks G. Ernst, A. Snoke, and R. Cole for discussing Cretaceous accretion in the western U.S. Cordillera (California, Idaho, and Alaska, respectively), P. Umhoefer and T. Pavlis for in-depth technical reviews, and Science Editor D. Fastovsky for additional comments/encouragement. Correspondence with G. Harlow is greatly appreciated. M. Mavec is acknowledged for field assistance (2017). We recognize LaMaskin et al. (2015) for stimulating part of this research, which has its origins in academic studies providentially guided by J.S. Oldow at the University of Idaho (1996–1997). Structural mapping in the south-central Kessler Creek quadrangle was funded by U.S. Geological Survey EDMAP Program award G16AC00102 (2016) to Gray. This paper is dedicated to Tracy L. Vallier.