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that was developed within foreland thrust belts using observational data ... tension does not drive foreland shortening, as was earlier suggested (Price,. 1973) ..... by recovery processes, and exhibit undulatory extinction and subbasal de- formation ..... In the Grouse Creek and western Raft River Mountains, at least two pen-.
Alternating contraction and extension in the hinterlands of orogenic belts: An example from the Raft River Mountains, Utah

Michael L. Wells*

Department of Geosciences, University of Nevada, Las Vegas, Nevada 89154

ABSTRACT Combined macroscopic to microscopic structural analyses, detailed geologic mapping, and thermochronology were used to study the structural evolution of midcrustal rocks of the Sevier orogenic belt hinterland in northwestern Utah. These results, when combined with previous studies, provide new insight into the structural history of this region, and suggest alternating tectonic contraction and extension during Mesozoic to early Cenozoic time. Two allochthons form the upper plate of the Miocene Raft River detachment fault in the eastern Raft River Mountains. The lower allochthon comprises Neoproterozoic, Ordovician, and Pennsylvanian(?) strata, and is bounded below by the Raft River detachment fault and above by the middle detachment fault. Strata within the lower allochthon were dramatically attenuated by two episodes of ductile deformation. The first deformation (D1) took place at metamorphic temperatures of ≈500 °C, and resulted in penetrative fabrics throughout these rocks that record combined flattening and top-to-northeast shearing strains. The second deformation (D2) resulted in significant stratigraphic attenuation along discrete top-to-the-west shear zones that are generally parallel to lithologic contacts. Separates of synkinematic muscovite from the penetrative fabric yield 40Ar/39Ar cooling ages that indicate that D1 deformation occurred prior to cooling ca. 90 Ma. Both fabrics were subsequently folded about (D3) kilometerscale recumbent folds. The areally extensive middle allochthon, composed chiefly of Pennsylvanian and Permian rocks metamorphosed in the greenschist facies, was emplaced (D4) along the low-angle middle detachment fault. This fault cuts across various structural levels of the recumbently folded lower allochthon in its footwall, and juxtaposes greenschist facies over amphibolite-facies metamorphic rocks. The lower and middle allochthon were subsequently deformed (D5) into open folds with north-trending axes. Neogene extension (D6 ) produced an ≈200-mthick top-to-east ductile shear zone in Precambrian rocks, and formed the younger Raft River detachment fault, which forms the present upper contact of the ductile shear zone. Northeast-vergent D1 fabrics probably record shortening deformation, on the basis of fabric correlations with the Grouse Creek and Albion mountains, deformation kinematics, and synkinematic prograde metamorphism. D2 attenuation faults have been interpreted to record crustal extension of Late Cretaceous age. If D3 recumbent folds developed during extension, then the deformation sequence records early shortening followed by protracted extensional unroofing with variable structural styles. The favored alternative is that D3 recumbent folds *Email address: [email protected]

developed during shortening; in this case, D1 through D4 record episodic alternations of contraction and extension. Such alternations are consistent with observations from analog and theoretical models of contractional mountain belts and suggest that the hinterland of the Sevier orogenic belt underwent dynamic adjustments in crustal thickness and deformation kinematics in response to changes in the boundary conditions of the orogenic belt.

Figure 1. Generalized location and tectonic map of the northeastern Great Basin illustrating location of hinterland metamorphic rocks (shown in diagonal wavy pattern) and ranges mentioned in text (shown in stippled pattern). Barbed lines = thrusts of the Sevier belt, hachured lines = normal faults of the Wasatch system. Labeled geographic and structural elements: AL, Albion Mountains; BP, Black Pine Mountains; CRS, Confusion Range synclinorium; DC, Deep Creek Range; GC, Grouse Creek Mountains; GSL, Great Salt Lake; G-T, Toano-Goshute Mountains; N, Newfoundland Mountains; P, Pilot Range; PM, Pequop Mountains; R-EH, Ruby Mountains–East Humboldt Range; RR, Raft River Mountains; SR, Snake Range; SRPV, Snake River Plain volcanic rocks (shown in v pattern); SS, Sublet synclinorium; WH, Wood Hills.

GSA Bulletin; January 1997; v. 109; no. 1; p. 107–126; 16 figures; 1 table.

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INTRODUCTION AND CONCEPTUAL BACKGROUND A complete description of the structural evolution of a mountain belt includes temporal correlation between deformational events in its foreland and hinterland. Correlation between deformation events within the foreland and hinterland of the Mesozoic to early Cenozoic Sevier orogenic belt of the western United States (Fig. 1) and comparison of their respective kinematics have remained problematic, partly because Mesozoic structures have been fragmented during Cenozoic extension, and isotopic systems and metamorphic fabrics have been strongly overprinted during Tertiary metamorphism, uplift, and cooling. Observations made within the past decade from analog (Davis et al., 1983) and theoretical models (Platt, 1986; Molnar and Lyon-Caen, 1988) and field studies of contractional mountain belts (Platt, 1986; Boyer, 1992; Wallis et al., 1993), however, may provide a framework to interpret hinterland structural evolution and to link episodes of hinterland deformation to the evolution of the Sevier foreland fold and thrust belt. The forward-propagating sequence of deformation (toward the foreland) that was developed within foreland thrust belts using observational data (e.g., Armstrong and Oriel, 1965; Jones, 1971) and kinematic models (Boyer and Elliot, 1982) may be oversimplified. Although the time of initial contractional deformation may progress at the regional scale toward the foreland, it is very likely that the hinterlands of orogenic belts continue to thicken internally as a response to the lengthening of the orogenic belt by propagation of deformation toward the foreland (Davis et al., 1983; Platt, 1986; Boyer, 1992). Furthermore, active large-scale extension is observed within the hinterlands of modern contractional mountain belts (e.g., Dalmayrac and Molnar, 1981; Burchfiel and Royden, 1985). Ancient examples of simultaneous hinterland extension and foreland shortening have now been widely reported from many orogenic belts (e.g., Wells et al., 1990; Hodges et al., 1992a; Wallis et al., 1993). Although hinterland extension does not drive foreland shortening, as was earlier suggested (Price, 1973), syncontractional extension may be an integral part of the orogenic cycle. For example, changes in topographic slope, crustal thickness, rheology of the decollement, compressional boundary stresses, or erosion rates in a dynamically evolving orogen can lead to extension within the hinterland (Platt, 1986; Dahlen and Suppe, 1988). Factors that increase buoyancy forces arising from thickened and elevated crust (forces that tend to thin the crust) relative to compressional boundary forces, which thicken crust, may lead to syncontractional extension. An outcome of the dynamic adjustment of topography and crustal thickness is that alternating periods of contraction and extension can occur during regional contraction (Platt, 1986). The kinematic role of the Sevier belt hinterland during the late Mesozoic to early Cenozoic has been controversial, and the structural record is apparently contradictory (e.g., Hose and Danes, 1973; Allmendinger and Jordan, 1984; Snoke and Miller, 1988). For example, Mesozoic contractional structures (e.g., Miller and Gans, 1989; Snoke and Miller, 1988; Camilleri, 1992; Taylor et al., 1993) and Mesozoic extensional structures (e.g., Wells et al., 1990; Hodges and Walker, 1992) have been reported in the northeastern Great Basin. It is difficult, however, to compare the relative timing of many of these deformations from range to range because of the difficulty in tightly bracketing the ages for Mesozoic deformations in this region (e.g., Snoke and Miller, 1988; Miller et al., 1988). Nonetheless, a history of continued adjustments in hinterland crustal thickness (i.e., alternating contraction and extension) during contraction in the foreland would explain these apparent kinematic inconsistencies (Wells, 1991). Middle-crustal rocks exposed in metamorphic core complexes commonly record a protracted history of Mesozoic and Cenozoic deformations (e.g., Miller et al., 1988; Snoke and Miller, 1988). The Raft River Mountains are part of a large metamorphic core complex that extends across the Raft

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River, Albion, and Grouse Creek Mountains of northwestern Utah and southern Idaho (Figs. 1 and 2). This paper focuses on Mesozoic to early Cenozoic deformation recorded in rocks in the upper plate of a Miocene low-angle normal fault in the eastern Raft River Mountains. In contrast to rocks exposed in the southern Albion and Grouse Creek Mountains, these rocks were far removed from Cenozoic plutons and were not thermally overprinted in Tertiary time, thus allowing isotopic dating of pre-Cenozoic deformations. This paper synthesizes results of geologic mapping, macroscopic and microscopic kinematic analysis, and thermochronologic studies from the eastern Raft River Mountains. The results suggest that the hinterland underwent alternating contraction and extension during late Mesozoic to early Cenozoic orogenesis. REGIONAL TECTONIC SETTING OF THE RAFT RIVER MOUNTAINS The Raft River Mountains are within the northeastern Great Basin, in the hinterland of the Mesozoic Sevier orogenic belt (Armstrong, 1968a), and east of the outcrop belts affected by Paleozoic contractional orogeny. PostPaleozoic deformation in the northeastern Great Basin can be broadly assigned to Late Jurassic(?) to early Eocene development of the Sevier orogenic belt, followed by middle(?) Eocene to recent crustal extension. A belt of apparently surface-breaking contractional deformation in central Nevada (Ketner and Smith, 1974; Taylor et al., 1993) has been tentatively correlated with contractional structures in northeastern Nevada, northwestern Utah, and southern Idaho (Camilleri et al., 1992). Existing age brackets in central Nevada suggest that deformation both predated the early Albian and was partly Late Cretaceous in age (Vandervoort and Schmidt, 1990; Taylor et al., 1993). In northeastern Nevada, the age constraints vary from pre–Late Jurassic (Coats and Riva, 1983), to Late Jurassic (Miller et al., 1987; Miller and Hoisch, 1992), to pre-Aptian (Thorman and Snee, 1988; Camilleri, 1992). Metamorphism and cleavage formation have been interpreted to be of Jurassic and Late Cretaceous age in the Snake Range and vicinity (Miller et al., 1988; Miller and Gans, 1989). In addition, geobarometric studies of Mesozoic metamorphism from several tracts of metamorphic rocks indicate substantial localized structural burial (Hodges et al., 1992b; Lewis et al., 1992). Thrusting in the foreland fold and thrust belt to the east may have begun in Late Jurassic or Early Cretaceous time (e.g., Armstrong and Oriel, 1965; Wiltchko and Dorr, 1983; Heller et al., 1986, Royse, 1993). This defines several possible relations to the belt of contractional deformation in central and northeastern Nevada and northwestern Utah. Hinterland contraction may (1) predate the development of the foreland fold and thrust belt (e.g., Allmendinger et al., 1984; Thorman et al., 1990; Miller and Hoisch, 1992); (2) be broadly synchronous with early activity in the Sevier thrust belt (Bartley and Gleason, 1990); (3) be “out-of-sequence” with respect to the foreland thrust belt; or (4) have developed during several periods of deformation, both before and during tectonism in the foreland thrust belt. Cretaceous extension of the Sevier belt hinterland, driven by lateral gradients in crustal thickness, was suggested based on studies in the eastern Raft River and Black Pine Mountains (Wells and Allmendinger, 1990; Wells et al., 1990). Hodges and Walker (1992) outlined other localities where Mesozoic extension is permissive (Hodges and Walker, 1992, Table 1), and suggested that late Mesozoic extension of the hinterland was a fundamental process of Sevier orogenesis. The existing timing constraints for Mesozoic hinterland extension do not permit temporal correlation to individual thrusting events within the well-established thrust sequence in the foreland thrust belt (e.g., DeCelles, 1994), but do allow late Early to Late Cretaceous hinterland extension to be broadly bracketed as occurring synchronous with overall shortening in the foreland.

Geological Society of America Bulletin, January 1997

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Figure 2. Tectonostratigraphic map of the Raft River, Black Pine, Albion, Grouse Creek, and Matlin Mountains. Box in eastern Raft River Mountains indicates location of more detailed geologic map shown in Figure 5. Modified from Compton (1975), Compton et al. (1977), Todd (1980, 1983), Miller (1983), and author’s mapping.

Following regional contraction, crustal extension with associated plutonism and metamorphism affected the northeastern Great Basin from the middle to late Eocene to the present (Compton, 1983; Dallmeyer et al., 1986; Lee et al., 1987; Mueller and Snoke, 1993; Smith and Sbar, 1974; Wells and Snee, 1993). It is not clear whether extension was continuous or episodic. Regionally, the age of initial Cenozoic extension decreases from

early Eocene in the Pacific Northwest to middle Miocene at the latitude of Las Vegas, Nevada (e.g., Armstrong and Ward, 1991; Axen et al., 1993; Duebendorfer and Wallin, 1991). South of the Snake River Plain, there is rare evidence for Eocene extension (Miller et al., 1987; Armstrong and Ward, 1991). In the Raft River, Albion, and Grouse Creek Mountains, however, middle to late Eocene extension is likely (Saltzer and Hodges, 1988;

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M. L. WELLS TABLE 1. SEQUENCE OF DEFORMATION, EASTERN RAFT RIVER MOUNTAINS Deformation event D1 D2 D3 D4 D5 D6 D7

Structure and kinematics Flattening and top-to-northeast shear Top-to-west attenuation faulting Recumbent folding Top-to-west displacement along middle detachment Open folding about north-south axes Top-to-east displacement along Raft River detachment and footwall shear zone Doming

Wells and Snee, 1993; Wells et al., 1994), suggesting that extension at this latitude occurred only ≈5 to 10 m.y. after cessation of activity along the frontal thrusts of the Sevier belt (Wiltschko and Dorr, 1983; Armstrong and Oriel, 1965). STRATIGRAPHY OF THE EASTERN RAFT RIVER MOUNTAINS Detailed studies of the stratigraphy, structural geology, and metamorphism have been carried out by many workers at various localities within the Raft River, Albion, and Grouse Creek Mountains (Armstrong, 1968b, 1976; Compton, 1972, 1975; Compton et al., 1977; Miller, 1980, 1983; Todd, 1980, 1983; and Wells et al., 1996). Several investigations have focused on deformation kinematics (Sabisky, 1985; Malavielle, 1987a, 1987b; Saltzer and Hodges, 1988; Wells et al., 1990; Manning and Bartley, 1994; Wells and Struthers, 1995), and on quantifying pressure-temperature conditions during metamorphism (Hodges and McKenna, 1986; Saltzer and Hodges, 1988). These studies document the complex deformation history recorded in these rocks. Overprinting Mesozoic and Cenozoic deformations have produced an incomplete and greatly attenuated stratigraphic section in the Raft River, Albion, and Grouse Creek Mountains. A tectonically thinned sequence of metasedimentary and sedimentary units of Neoproterozoic to Triassic age overlies an Archean basement complex over an area greater than 4000 km2 (Figs. 2, 3, and 4) (Armstrong, 1968b; Compton et al., 1977, Wells et al., 1996). Within the eastern Raft River Mountains, these rocks have been subdivided into three low-angle fault-bounded tectonostratigraphic units (Figs. 3 and 5) (Compton et al., 1977; Miller, 1980, 1983; Todd, 1980, 1983; Wells, 1992). From structurally lowest to highest (Fig. 3), these are: (1) the parautochthon, (2) the lower allochthon, and (3) the middle allochthon. Although these allochthons are bounded by major faults in many areas of the Raft River, Albion, and Grouse Creek Mountains, the faults or shear zones that bound them are not necessarily correlative in age or origin from one locality to another. Until further work is carried out, the allochthon designations imply only stratigraphic affinity and general structural position, rather than structural continuity. A brief description of the rock units that compose the parautochthon and allochthon follows; see Wells (1992, 1996) for more detailed lithologic descriptions. The Green Creek complex (Armstrong and Hills, 1967; Armstrong, 1968b), exposed in the parautochthon, is composed of ca. 2.5 Ga gneissic monzogranite (metamorphosed adamellite of Compton, 1972) that intrudes schist and amphibolite (Compton, 1975; Compton et al., 1977), all of which are unconformably overlain by the Elba Quartzite. The Elba Quartzite forms the lowermost part of the Raft River Mountains sequence of Miller (1983), and has been assigned various ages including Paleozoic, Neoproterozoic (Armstrong, 1968a; Compton et al., 1977; Compton and Todd, 1979), and Paleoproterozoic (Crittenden, 1979). The structurally overlying quartzite of Clarks Basin has been shown to be Neoproterozoic, suggesting a Proterozoic age for the Elba Quartzite (Wells et al., 1996). Anomalously high δ13C values from marble interbeds within the quartzite

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Interpretation

Timing

Contraction Extension Contraction Extension

Pre-90 Ma Ca. 90 Ma

Extension

Middle to early late Miocene

Late Eocene–Oligocene (?)

of Clarks Basin resemble those measured in other Neoproterozoic rock sequences in western North America (Wells et al., 1996). The Precambrian rocks in the parautochthon record andesine-amphibolite-facies metamorphism (e.g., Winkler, 1976). The lower allochthon consists of Neoproterozoic quartzite and schist; Ordovician calcitic marble, phyllite, quartzite, and dolomite; Silurian(?) dolomite; and Pennsylvanian(?) marble. These units are generally correlated with those described within the lower allochthon to the west (Compton, 1975; Compton et al., 1977; Compton and Todd, 1979), with one exception: a highly tectonized marble unit, which everywhere occurs adjacent to Ordovician rocks, was recognized and correlated by physical stratigraphy with the lower part of the Pennsylvanian Oquirrh Formation. Metamorphic mineral assemblages within the Neoproterozoic schist of Mahogany Peaks include garnet + staurolite + muscovite + biotite + plagioclase + quartz and staurolite + kyanite + biotite + muscovite + zoisite + quartz, and indicate peak amphibolite-facies metamorphic conditions. These assemblages are confined to the fields “l-7 and m-7” of the petrogenetic grid for the KFMASH system (Spear and Cheney, 1989, Fig. 2), and indicate minimum temperature and pressure of ≈600 °C and 6.5 kbar. Extensive retrogradation of staurolite to chloritoid and muscovite, and of biotite to chlorite, took place at greenschist-facies conditions. Metamorphic temperatures in Ordovician rocks were estimated to be at least 490 to 520 °C by oxygen isotopic geothermometry of quartz-muscovite and quartz-biotite mineral pairs and conodont color alteration index (CAI) values >7 (Wells et al., 1990). The middle allochthon is separated from the lower allochthon by the middle detachment fault (Fig. 3). The middle allochthon mainly comprises rocks of the Pennsylvanian and Permian Oquirrh Formation; a thin (0–5 m thick) sliver of Chainman Shale and Diamond Peak Formation is locally present along its base. Conodonts from the lower Oquirrh Formation yield CAI values of 5, indicating temperatures of 350 to 400 °C greenschist-facies metamorphic conditions (Wells et al., 1990). STRUCTURAL GEOLOGY OF THE EASTERN RAFT RIVER MOUNTAINS The Raft River detachment fault separates Precambrian rocks of the parautochthon from the overlying lower and middle allochthons of the upper plate (Figs. 3 and 5). The upper-plate rocks in the eastern Raft River Mountains provide a unique opportunity within the Raft River–Albion–Grouse Creek metamorphic core complex to study a protracted history of Mesozoic to early Cenozoic deformations for several reasons. (1) The rocks achieved peak metamorphic conditions of greenschist to middle amphibolite facies and contain multiple penetrative deformation fabrics that record superposed deformations. (2) Upper-plate Neoproterozoic to Permian rocks were not overprinted by Neogene mylonitization and metamorphism that affected lower-plate Archean and Proterozoic rocks, and have resided at high structural levels since Late Cretaceous time (Wells et al., 1990). (3) Stratigraphically equivalent strata within the western Albion Range, western Raft River, and Grouse Creek Mountains were at deeper structural

Geological Society of America Bulletin, January 1997

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Figure 3. Tectonostratigraphic column for the eastern Raft River Mountains. Thicknesses are approximate maxima and highly variable. Positions of low-angle faults are indicated.

levels in Paleogene time and are structurally overprinted by penetrative top-to-the-west Paleogene mylonitic fabrics and locally thermally overprinted adjacent to Paleogene plutons (Compton et al., 1977; Todd, 1980; Saltzer and Hodges, 1986; Wells et al., 1994; Wells and Struthers, 1995). The upper plate of the Miocene Raft River detachment fault has had a complex deformation history. The sequence of deformations (Table 1) determined by detailed geologic mapping (Figs. 6, 7, and 8) includes the following: (D1) attenuation of rock units within the lower allochthon by early intrabed plastic flow; (D2) attenuation of units within the lower allochthon by top-to-west younger-over-older faulting; (D3) recumbent folding of at-

tenuated strata and faults; (D4) emplacement of the middle allochthon over the lower allochthon along the middle detachment; (D5) upright, open folding of the middle and lower allochthons; (D6) top-to-east normal-sense shearing along the shear zone at the top of the parautochthon and along the Raft River detachment fault, with concomitant high-angle normal faulting of the upper plate; (D7) doming of the Raft River detachment fault and earlier structures. The Neoproterozoic through Pennsylvanian(?) units within the lower allochthon were markedly attenuated during superposed deformations. The section varies from 50 to 500 m in thickness (Fig. 4) whereas, in neighbor-

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Figure 4. Comparison between attenuated stratigraphic thicknesses within the eastern Raft River Mountains and representative stratigraphic thicknesses from nearby localities in northwestern Utah (from Hintze, 1988). Of — Ordovician Fish Haven Dolomite, see Figure 3 for other abbreviations. ing ranges, equivalent strata are 4 to 7 km thick. The attenuation has been achieved by penetrative plastic thinning (D1) during amphibolite facies conditions and later (D2) low-angle, younger-over-older ductile and brittle faulting (referred to as attenuation faulting, after Hintze, 1978). D1 Deformation The most pervasive fabric in the lower allochthon is a penetrative foliation (S1) that is generally parallel to lithologic layering and inferred bedding. At the few locations where foliation is distinct from bedding, foliation is slightly inclined to the southwest relative to bedding, whether in upright or overturned rocks (Fig. 9A). Lineation varies in degree of development, and is not ubiquitous (note few lineation measurements as compared to foliation). The sparse development of lineation precludes the analysis of deformation superposition using lineation-orientation analysis. Lineation is

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best developed in marbles and generally trends northeast in both upright and overturned rocks (Fig. 10). The only recognized F1 folds are sparse mesoscopic southeastward-overturned folds with axes generally trending northnortheast. With the exception of a few asymmetric structures that indicate east and northeast shearing, the majority of strained mesoscopic features are symmetric with respect to S1 foliation (including pressure shadows around rigid objects and boudinage of quartzite and chert interbeds). This observation, coupled with the locally apparent southwest dip of foliation relative to bedding, suggests that the D1 fabrics in carbonate rocks resulted from a combination of pure shear and northeast-directed simple shear. This kinematic interpretation is supported by the microstructural investigations of calcitic marble and quartzite outlined as follows. Calcitic Marbles. The lineation in marble is most commonly defined by elongate pressure shadows of quartz and calcite about pyrite grains, elongate calcite grains, and stretched and recrystallized fossil fragments. The foliation is microscopically defined by a grain-shape fabric of plastically deformed calcite and a parallel alignment of muscovite. A large body of work has documented the development of crystallographic preferred orientations in calcite-bearing rocks. Experimental studies and computer simulation of coaxial (e.g., Turner et al., 1956; Casey et al., 1978; Wagner et al., 1982; Wenk et al., 1987) and noncoaxial (Rutter and Rusbridge, 1977; Kern and Wenk, 1983; Schmid et al., 1987; Wenk et al., 1987) deformation of limestone have determined expected crystallographic preferred orientations (e.g., c-axes). In addition to the strain path and the shear sense, the relative magnitudes of pure and simple shear in a deformation event can be derived from the crystallographic fabric (Dietrich and Song, 1984; Wenk et al., 1987; Schmid et al., 1987). The c-axis [0001] preferred orientation for two samples of Ordovician marble containing D1 fabrics was determined by measurement of mutually perpendicular thin sections on the universal stage (Fig. 11, A and B). Both samples exhibit a strong preferred orientation of c-axes. The pole figures are slightly asymmetric with respect to the foliation normal (obliquities of ≈16°). The c-axis fabric obliquities (Wenk et al., 1987) suggest about equal components of top-to-the-northeast simple shear and pure shear deformation, and a general shear strain path. Quartzites. The microstructures in quartzites suggest deformation kinematics similar to those of the calcitic marbles. Ordovician quartzite exhibits a weak macroscopic S1 foliation and typically no lineation. Neoproterozoic quartzite of Clarks Basin is more highly foliated and lineated and was studied to elucidate the kinematics of D1 deformation. An east- to northeast-trending stretching lineation (L1) is defined by elongate muscovite and quartz grains, and foliation is defined by aligned muscovite and a grain-shape fabric of dynamically recrystallized quartz. Microscopically, quartz grain boundaries range from relatively straight with 120° triple junctions to curved to embayed (Fig. 9C). Many features similar to those described by Jessel (1987) indicate pinning of mobile quartz grain boundaries by muscovite and recrystallization by grain-boundary migration (Urai et al., 1986), consistent with deformation at amphibolite-facies metamorphic conditions. The more elongate recrystallized grains exhibit discrete extinction domains with sharp domain boundaries, representing prismatic subgrains (White, 1973). Subgrain boundaries lie at angles of 42° to 68° from the principal foliation, the large majority being inclined toward the southwest relative to principal foliation. The preferred orientation of quartz c-axes was analyzed in three samples. The c-axis fabrics are highly variable (Figs. 12 and 13, A–C). Sample RR91-14 exhibits a small circle distribution (with an opening angle of ≈30°) about the pole to the foliation (Fig. 13A). Small circle distributions have been produced under laboratory conditions of uniaxial flattening (Tullis et al., 1973) and measured in studies of naturally deformed rocks that record coaxial flattening strains (e.g., Law et al., 1984). Sample

Geological Society of America Bulletin, January 1997

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Figure 5. Generalized geologic map of the eastern Raft River Mountains indicating axial traces of F3 recumbent folds and F5 open folds. Note that Pennsylvanian(?) marble commonly occurs within the cores of recumbent synclines. Stereoplots are equal area Pi-diagrams of poles to foliation from localities of F5 folds indicated by arrows; pole to best-fit great circle (fold axis) indicated by square. The outlined box indicates location of detailed map of Figure 6. RR91-93 (Fig. 13B) exhibits a c-axis fabric intermediate between an asymmetric single girdle and an asymmetric type II crossed girdle (Lister, 1977) and records a significant component of noncoaxial, top-to-northeast shear (Schmid and Casey, 1986). The third sample, RR67, shows an asymmetric single girdle fabric (Fig. 13C) recording noncoaxial top-to-east-northeast shear. These samples are from different localities and different structural horizons within the outcrop of the quartzite of Clarks Basin, and no attempt was made to determine the deformation partitioning with respect to structural level. Elementary strain compatibility arguments (e.g., Law and Potts, 1987), however, suggest that the differences in c-axis pole figures may relate to a strain field with different amounts of northeast-directed noncoaxial strain superimposed on uniform coaxial strain. For strain compatibility between structural levels, the components of coaxial strain (to a first order) should be uniform for all structural levels, to avoid the development of material overlaps or gaps. Variability with structural level in the simple shear component, however, will not result in development of material gaps or overlaps. Taken together, the D1 fabrics within the carbonate and quartzite tectonites indicate components of coaxial flattening and noncoaxial top-tonortheast shearing strain.

D2 Attenuation Faults and Fabrics Faults. Two major discontinuities in stratigraphy within the lower allochthon mark faults. The structurally higher fault, the Emigrant Spring fault (Wells, 1996; LAF1 of Wells, 1992; see Figs. 3 and 6), places Pennsylvanian(?) calcitic marble over Ordovician and Silurian(?) dolomitic marble throughout the mapped area. The Emigrant Spring fault removes about 5 km of stratigraphic section, including Silurian, Devonian, and all but centimeter-scale slivers of Mississippian rocks. This discontinuity is interpreted to be a fault, contrary to the earlier interpretation of an unconformity (Armstrong, 1968b), because (1) Devonian and Mississippian strata are present in neighboring ranges, including the Black Pine, southern Grouse Creek, and Newfoundland Mountains, and the Pilot Range (Fig. 1; Hintze, 1988); (2) Pennsylvanian(?) marble lying above this discontinuity contains a distinct mylonite zone parallel to the contact; (3) this contact is locally structurally discordant (Fig. 6); and (4) the removal of rock units by low-angle faulting is characteristic of deformation at other stratigraphic levels within the lower allochthon. The second major discontinuity separates the schist of Mahogany Peaks

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Geological Society of America Bulletin, January 1997 Figure 6. Detailed geologic map of the Crystal Peak area, eastern Raft River Mountains. The map relationships clearly demonstrate the structural sequence of deformations. Location of cross section A–A′ indicated. Pennsylvanian(?) marble (IPot) crops out in cores of F3 recumbent folds. See Figure 7 for explanation of lithologic symbols.

CONTRACTION AND EXTENSION IN OROGENIC BELTS

Figure 7. Geologic cross section. See Figure 6 for location of cross-section line A–A′. Note D2 attenuation faults that omit the stratigraphic section are folded about F3 recumbent folds. D2 attenuation faults, F3 recumbent fold axial surfaces, and the middle detachment are folded about F5 open folds.

Figure 8. Photograph looking northeast toward Crystal Peak, illustrating the overprinting relationships between various structures. The prominent gently dipping white rock unit in the footwall of the structurally lowest fault (single barb, the Raft River detachment fault) is the Elba Quartzite. The Raft River detachment truncates the broad F5 upright anticline. The D2 attenuation faults (open teeth) and the D4 middle detachment fault (double barbs) are folded by the broad F5 fold. Field of view about 1.5 km across.

and quartzite of Clarks Basin from the Ordovician carbonate rocks. The faulted nature of this contact has been clarified using carbon isotopes, confirming a Neoproterozoic age for the quartzite of Clarks Basin, in contrast to the Cambrian age previously suggested (Compton et al., 1977; Compton and Todd, 1979). This discontinuity, termed the Mahogany Peaks fault, places Ordovician over Neoproterozoic strata, omitting >2 km of stratigraphic section (Wells, 1992; Wells et al., 1996). The regional-scale geometry of these two large-displacement faults and the consistency in hangingwall and footwall stratigraphic levels over large distances suggest that the

exposed segments of these faults are bedding-parallel flats in low-angle faults with ramp and flat geometries. In addition to the two major low-angle faults described above, the majority of other formational contacts within the lower allochthon are lowangle faults that place younger rocks on older. These attenuation faults, with few exceptions, are subparallel to foliation and bedding in adjacent units and are identified using both textural and stratigraphic criteria. In places, there is moderate discordance between the fault zones and the adjacent units, and between rock units on either side of the faults (Fig. 6). Slivers of units are commonly strung out along the faults and, locally, units are completely omitted. In places, the faults are recognized by ductile highstrain zones localized along bedding-parallel contacts between stratigraphic units along which section is commonly missing. Where the faulted nature of stratigraphic contacts is not evident by field inspection, significant lateral changes in map-scale unit thicknesses attest to their presence. Kinematic Analysis. Within attenuation fault or shear zones, fabrics range from ductile to brittle. In many localities, mylonitic fabrics are cut by systems of normal faults of similar kinematics, suggesting a progression from early ductile to later brittle shearing. Foliation within the shear zones is subparallel to S1 foliation, except in rare localities where S1 foliation is clearly truncated by the D2 shear zones. This general parallelism of S1 and S2 foliations makes it difficult to define the margins of D2 shear zones. A plastic shear zone is present within the lower part of the Pennsylvanian Oquirrh Formation(?) marble adjacent to the Emigrant Spring fault. The marble is highly foliated and locally contains a well-developed east-trending stretching lineation (Fig. 14, A and B) and isoclinal folds having hinge lines parallel to lineation. A deformation gradient is present within this shear zone. The basal part is ultramylonitic, grading upward to more coarsegrained mylonitic marble. The ultramylonite is 1–2 m thick, and is characterized by extreme reduction in grain size to 8 to 20 µm, untwinned, equant calcite grains. Grain size increases structurally upward to 2 mm within the mylonite, and foliation and lineation are uniformly oriented across this tran-

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Figure 9. D1 fabrics. (A) Highly transposed bedding within Swan Peak limestone in view looking north. Foliation is inclined west relative to bedding. 12.5 cm eraser for scale. (B) Boudinaged quartz vein that lies at low angle to foliation, illustrating large-magnitude intrabed plastic flow. Hammer for scale. (C) Photomicrograph of preserved D1 fabric within quartzite of Clarks Basin. Muscovite defines S1 foliation. Note morphology of quartz grain boundaries, including embayed and straight grain boundaries with 120° triple junctions, indicating grain boundary migration recrystallization. Photomicrograph 3 mm across.

sition. The marble ultramylonite contains sparse porphyroclasts of calcite that are slightly elongate parallel to a foliation defined by a parallel alignment of fine-grained muscovite. The deformation kinematics of the marble ultramylonite and more coarse-grained marble mylonite were investigated by study of lattice-preferred orientations. Samples collected from overturned fold limbs exhibit somewhat diffuse and complex c-axis orientations and no systematic relationship between c-axes and macroscopic fabric elements, suggesting significant intracrystalline straining after foliation formation, probably during folding. Because of this, only samples from upright fold limbs that are probably less overprinted by folding strain are included in the kinematic analysis. The lattice-preferred orientation of the ultramylonite was determined from X-ray pole figure goniometry. Calcite a-axes exhibit a strong preferred orientation (Fig. 11G). The a-axis girdle and inferred c-axis maxima are asymmetric with respect to foliation, indicating noncoaxial top-to-west shear. The more coarse-grained mylonitic marbles were measured on the universal stage. All samples collected from upright limbs to D3 recumbent folds exhibit a strong preferred orientation of calcite c-axes at a high angle to foliation (Fig. 11, C–F). The asymmetries of the caxis maxima relative to the normal to the grain-shape fabric are not uniformly consistent and, therefore, analysis of all marble samples did not indicate a uniform shear sense. However, the samples were collected at different structural levels above the ultramylonite, and the variations in the c-axis fabric may reflect deformation partitioning within the D2 shear zone, different degrees of D2 overprint on D1 fabric, or combinations of both. Dur-

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ing e-twinning, the c-axis of the twin undergoes a change in orientation of 52° relative to the host grain (Turner et al., 1956; Handin and Griggs, 1961). If the twinning strain postdates the initial c-axis preferred orientation and foliation development, c-axis reorientation during twinning (measured twin volumes of 5% to 25% of the total grain) might have affected the earlier preferred orientations. Nevertheless, the majority of measured samples exhibit lattice-preferred orientations consistent with a significant component of west-directed simple shear. In addition to determination of the c-axis texture, a method of “dynamic” analysis of calcite twin lamellae pioneered by F. J. Turner (1953) was employed. Although this “dynamic” method was originally interpreted to yield compressional and tensional stress axes, they are more aptly regarded as incremental strain axes. Therefore they are referred to as shortening and extension axes. The c-axis maxima are expected to correspond closely to the compression direction that produced the grain-shape fabric and crystallographic preferred orientation (see Wells and Allmendinger, 1990, for a review). There is a strong deviation between shortening axes determined by the Turner method and the c-axis maxima in the samples of Pennsylvanian(?) and Ordovician marble. This discordance is interpreted to indicate overprinting deformations, and the twinning strain is interpreted to postdate the development of the predominant c-axis fabric and foliation. The late twin strain is geometrically related to foliation (Fig. 15) rather than the present geographic reference frame (not shown), because of the significantly

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Figure 10. Lineations from the eastern end of the Raft River Mountains. (A) Lineation map with stereographic projections of lineations from the parautochthon and the lower allochthon. (B) Rose diagrams of azimuths from four units of the lower allochthon. Because of the gradational nature of many D2 shear zones and the near parallelism of associated foliation, D1 and D2 lineations and foliations were not differentiated in the field. The predominance of east-west lineations within the Pennsylvanian(?) marble reflects the greater development of lineations related to westward D2 attenuation faulting. (C) All foliations, S1 and S2 combined.

greater consistency of kinematic axes between samples in the former projection. The twin strain could have resulted from D2 westward shearing, which occurred along shear zones oriented subparallel to D1 foliation. The kinematic axes are compatible with this interpretation; when viewed in an approximately east-west projection perpendicular to foliation (Fig. 15), the majority of the samples exhibit Turner shortening axes consistent with foliation-subparallel westward shearing. The absence of twins in the marble tectonite that underwent dynamic recrystallization during D2 shearing supports the interpretation that the twin strain in the more coarse-grained marble is related to D2 deformation. Alternatively, the twin strain may record westward shearing related to westward translation of the middle allochthon during D4, or foliation-parallel shear during F3 folding. Other west-vergent shear indicators from the basal marble tectonite of the lower Oquirrh Formation include a secondary recrystallized grainshape fabric present in coarser-grained marble that is oblique to and inclined eastward relative to the mylonitic foliation, and deformed fold limbs that are boudinaged where inclined eastward and folded where inclined westward relative to foliation. Common late brittle fault networks consistently indicate the same sense of shearing. West-directed mylonitic shear zones, ranging in thickness from 1 to 150 cm, are recognized within greatly attenuated Swan Peak and Eureka quartzites. Mylonitic foliation is defined in thin section by an alignment of

elongate to ribbonlike quartz grains that are oblique to discontinuous to continuous zones of complete dynamic recrystallization and concentrated shear; this fabric is interpreted to represent an S-C fabric, and indicates topto-west shearing (Fig. 14C). The ribbon quartz grains are not recrystallized by recovery processes, and exhibit undulatory extinction and subbasal deformation lamellae, indicating greenschist-facies metamorphic conditions. Within the subparallel zones of dynamic recrystallization defining C-surfaces, an oblique grain-shape fabric (Lister and Snoke, 1984; Law et al., 1984) is present that also indicates the same sense of vorticity (Fig. 14D). Elsewhere, in finely laminated quartzite and phyllite, shear bands or extensional crenulation cleavage (Platt and Vissers, 1980) indicate westward shearing. Quartz lattice studies from mylonitic Eureka Quartzite substantiate the kinematics determined from the microscopic- and mesoscopic-scale features. Two samples were measured on a universal stage for determination of the preferred orientation of c-axes (Fig. 13, D and E). Sample RR38 shows an asymmetric single girdle fabric recording noncoaxial top-to-west shear (Fig. 13D). The c-axis maxima at positions close to the primitive circle suggest dislocation glide along basal planes in , consistent with greenschistfacies metamorphic conditions (Schmid and Casey, 1986). Sample RR91-28 (Fig. 13E) exhibits an asymmetric type II crossed girdle (Lister, 1977) with a leading edge (Behrmann and Platt, 1982) interpreted to record

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tinctly lower temperature than D1 fabrics and are unrecovered; and (3) D1 fabrics have distinct kinematics from the attenuation fault zones. D3 Recumbent Folding (F2) The Ordovician through Pennsylvanian(?) rocks of the lower allochthon are deformed into tight to isoclinal recumbent folds with amplitudes greater than 1.5 to 2 km. The complex superposition of later deformations, however, precludes determination of original fold geometry. The youngest unit involved in recumbent folding is the Pennsylvanian(?) marble, which occupies the cores of recumbent synclines (Figs. 6 and 16A). The D1 foliation consistently dips more steeply southwestward than transposed bedding on both upright and overturned limbs, a relationship that is inconsistent with the foliation forming synchronously with folding, but permits the foliation to either predate or postdate folding. The D1 foliation is interpreted to predate folding because F3 recumbent folds deform D2 attenuation faults, and D2 faults locally truncate and overprint the D1 foliation. The lack of a recognizable foliation associated with F3 folding in the eastern Raft River Mountains may be due to relatively dry, postmetamorphic conditions that were unfavorable for low-temperature deformation mechanisms, such as fluid-assisted grain boundary diffusion. The F3 recumbent folds are bracketed as younger than D2 attenuation faults and older than the D4 middle detachment. Attenuation faults occur in both upright and overturned limbs and remove the same amount of stratigraphic section on both limbs (e.g., the Emigrant Spring fault, Figs. 3, 6, and 7); therefore, recumbent folding is interpreted to postdate attenuation faulting within the lower allochthon. The middle allochthon rests on various structural levels of both upright and overturned fold limbs (Figs. 6 and 7), indicating that F3 recumbent folding preceded emplacement of the middle allochthon. D4 Emplacement of the Middle Allochthon Along the Middle Detachment

Figure 11. Calcite pole figures for marbles within the lower allochthon. Diagrams A through F are c-axis pole figures determined from measurement on a universal stage. (G) is X-ray derived a-axis pole figure. The east-west great circle represents the macroscopic foliation plane and the lineation is indicated by the filled circle. RR112 (A) and RR32 (B) are Ordovician marbles; all other samples are from Pennsylvanian(?) marble. RR3 (G) is a marble ultramylonite. Contours for c-axis pole figures are 2, 4, 6, 8, 10, and 12 sigma; contours for a-axis pole figure (G) are of 1, 1.5, 2.0, and 2.5 of random distribution. a significant component of noncoaxial, top-to-west shear (Schmid and Casey, 1986). Relative timing relations between D1 fabric and D2 attenuation faults are clearer in the eastern Raft River Mountains than in areas to the west. The D2 attenuation faults are interpreted to postdate penetrative foliation development (D1) for the following reasons: (1) attenuation faults locally truncate D1 foliation; (2) quartzite fabrics within attenuation faults are dis-

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Greenschist-facies, Pennsylvanian to Permian Oquirrh Formation rocks of the middle allochthon overlie amphibolite-facies, Neoproterozoic to Pennsylvanian(?) rocks of the lower allochthon along the middle detachment fault. The transport direction for this fault is determined from study of a thin sliver of Chainman Shale and Diamond Peak Formation that is locally present along the base of the middle allochthon. A 5-m-thick layer of black shale is exposed at Bald Knoll (Fig. 2). Pressure shadows around pyrite are ubiquitous and impart a strong west-southwest–trending stretching lineation in this shale. The pressure shadows are strongly asymmetric relative to foliation as viewed parallel to foliation and perpendicular to lineation (Fig. 16B). The morphologies of the fibers within the pressure shadows are similar to those described by Etchecopar and Malavieille (1988) for rigid, face-controlled fibers growing around euhedral pyrite during noncoaxial flow. These pressure shadows are interpreted as recording top-tothe-west-southwest shear. Because this shale occurs as a thin sliver at the base of the middle allochthon, the kinematics of deformation within the shale are thought to record distributed simple shear along the middle detachment. Westward translation of the middle allochthon is consistent with data from the Black Pine Mountains, 15 km northeast of the study area (Fig. 1) (Wells and Allmendinger, 1990). Movement along the middle detachment fault probably occurred at metamorphic conditions no higher than lower greenschist facies. This is evident from the metamorphic discontinuity across this fault in the study area and farther to the west (Compton et al., 1977), and the deformation mechanisms within the sheared base of the middle allochthon: pressure solution and brittle fracturing of quartz and pressure solution and low-tem-

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Figure 12. Location map for microstructural samples of Figures 11, 13, and 15. See Figure 5 for explanation of patterns and symbols.

perature plasticity of calcite. The displacement along the middle detachment was, therefore, probably late metamorphic to postmetamorphic. The middle detachment fault can be traced discontinuously across the east-west extent of the Raft River Mountains and into the Grouse Creek and Albion Mountains. This fault was interpreted by Compton et al. (1977) to represent the principal detachment in the Raft River–Albion–Grouse Creek metamorphic core complex (RR-A-GC MCC), and have top-to-east displacement. This fault, as exposed in the eastern Raft River Mountains, is interpreted to represent an extensional detachment fault with westward translation, because it removes stratigraphic section, juxtaposes lower grade on higher grade metamorphic rocks (Compton, 1977; Wells, 1992), and locally records top-to-the-west kinematics. This fault may be related to deeper level top-to-west extensional plastic shear of Eocene–Oligocene age in the western Raft River, Grouse Creek, and Albion Mountains (Saltzer and Hodges, 1988; Wells and Snee, 1993; Wells and Struthers, 1995). D5 OPEN FOLDING (F5) Both the lower and middle allochthon are folded into broad, upright, open folds, the axes of which trend roughly north-south (Fig. 5). South of Crystal Peak, an F5 open fold with a wavelength greater than 1500 m is well exposed (Figs. 5 and 8). Other folds have wavelengths of 30 to 150 m. The upright F5 folds clearly deform both the middle detachment and un-

derlying low-angle faults within the lower allochthon, and are apparently truncated by the Raft River detachment. The D5 folds could have formed contemporaneous with D6 extensional shear. However, the occurrence of D5 folds in fold trains (as opposed to single folds), upright axial surfaces, and small wavelengths for some of these folds suggests that they are not folds related to the geometry of underlying normal faults such as reverse-drag or fault-bend folds. Furthermore, these folds are cut by high-angle faults (Fig. 7) that in turn are truncated by the Raft River detachment fault, indicating that the folds formed prior to latest movement on the Raft River detachment. D6 Footwall Shear Zone and Raft River Detachment Fault Structurally beneath the Raft River detachment and within the parautochthon there is an ≈200-m-thick shear zone that is parallel to bedding in the Proterozoic units and the underlying unconformity with Archean rocks. The shear zone accommodated large-scale top-to-the-east displacement (Compton, 1980; Sabisky, 1985; Malavieille, 1987a; Wells, 1992), and strain intensity progressively increases from west to east. Mylonitic fabrics are the most highly developed within the Elba Quartzite and its schist member, and fabric intensity related to eastward shearing typically dies out abruptly downward within the Green Creek complex. Northwardoverturned recumbent folds are developed within the extensional shear

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Figure 13. Quartz c-axis pole figures for the D1 and D2 fabric. (A–C) D1 fabric in quartzite of Clarks Basin. (D and E) D2 fabric in Eureka Quartzite. The east-west great circle represents the macroscopic foliation plane; the filled circle indicates lineation.

zone. Fold hinge lines are parallel to the extension direction and greater than 20 km in length (Compton, 1980; Malavieille, 1987a). Superimposed on the shear zone in the parautochthon is the brittle Raft River detachment fault. This major structural discontinuity forms the upper contact of the schist member of the Elba Quartzite. The detachment fault is concordant with foliation and bedding within the lower plate, but truncates structures within the upper plate. There is a close spatial and kinematic association between the detachment fault and the underlying shear zone. Fabrics within the uppermost shear zone are commonly retrograde and, where a cataclasite is present, the mylonite is progressively overprinted by cataclastic deformation structurally upward toward the detachment fault. Both plastic and brittle structures exhibit the same top-to-the-east shear sense. The upper plate of the Raft River detachment fault is cut by numerous high-angle faults with displacements from centimeters to about 2 km. The strikes of the faults vary greatly, but the majority strike roughly north. Only two faults, both of small displacement, were noted that cut the Raft River detachment fault. The upper-plate high-angle faults, therefore, formed either prior to or during movement of the Raft River detachment fault. Thermochronological results from within and beneath the shear zone in the parautochthon contrast markedly with the Late Cretaceous muscovite ages from the lower allochthon (Wells et al., 1990). Late-early to early-late Miocene cooling ages (40Ar/39Ar from muscovite, biotite, and microcline, Wells and Snee, 1993; fission track in apatite, Wells et al., 1994) collectively indicate rapid cooling during the Miocene, during footwall unroofing related to extensional displacement. The mylonite zone has been previously interpreted to represent a Mesozoic thrust-sense shear zone (Malavielle and Cobb, 1986; Snoke and Miller, 1988), and Cenozoic normal-sense shear

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zone (Malavielle, 1987a; Wells, 1992). Our studies indicate that plastic shearing occurred within the temperature window spanned by rapid cooling, as indicated by the deformation conditions recorded in the microstructures (Wells et al., 1994), and confirm a Miocene age. D7 Broad Folding The present structure of the Raft River Mountains is an east-trending, doubly plunging anticline of 26 km length and 1.2 km of exposed structural relief. The shear zone, the strata in the parautochthon, and the detachment fault outline this large structure. The axis of the elongate anticline is subparallel to the transport direction of both the footwall shear zone and the brittle detachment fault. Nonplanar detachment faults of comparable geometry are present within many other core complexes, in particular core complexes of the Colorado River trough region (e.g., Spencer, 1982). ROCK FABRICS ELSEWHERE IN THE RAFT RIVER– ALBION–GROUSE CREEK METAMORPHIC CORE COMPLEX Albion Mountains Two presumed Mesozoic deformations have been documented in the northern Albion Mountains (Miller, 1980). The earlier exhibits a northeasttrending lineation (L1 of Miller, 1980), with a component of top-to-northeast shear (Malavielle, 1987b), and the later a northwest-trending lineation (L2 of Miller, 1980) and top-to-northwest shear (Malavieille, 1987b). Rocks containing these fabrics yielded Mesozoic and Cenozoic conventional K-Ar cooling ages (biotite, muscovite, and hornblende), eight of which range from

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Figure 14. D2 attenuation faults and fabrics. (A and B) Tectonized Pennsylvanian(?) marble in immediate hanging wall of Emigrant Spring fault. (A) Mylonitic foliation. (B) Stretching lineation viewed on foliation surface. (C and D) Greenschist facies mylonitic fabrics within Eureka Quartzite indicating top-to-west shear. (C) Elongate quartz ribbons and mica-rich zones of concentrated shear, respectively, define an S-C fabric indicating top-to-west (sinistral) shear. (D) Detail of oblique grain-shape fabric within zones of concentrated shear, indicating dextral shear.

66 to 81 Ma (Armstrong, 1976). Hodges and Walker (1992) interpreted the L2 fabric to record Cretaceous down-to-the-northwest extensional shear. A third lineation, trending west-northwest, is present in the southwestern Albion Mountains within the Middle Mountain shear zone, a west-dipping extensional shear zone with top-to-west-northwest shear sense (Miller et al., 1983; Saltzer and Hodges, 1988). The Middle Mountain shear zone is younger than L2 of Miller (1980), and metamorphic minerals that grew during shearing yield late Eocene to Oligocene (K-Ar and 40Ar/39Ar) cooling ages (Armstrong, 1976; Miller et al., 1983; Saltzer and Hodges, 1988).

to northwest–trending lineation in the Grouse Creek and western Raft River Mountains are problematic. They exhibit similarity in trends and kinematics to the Mesozoic(?) L2 of Miller (1980) in the northern Albion Mountains, the late Eocene to early Oligocene west-northwest–trending lineation of the Middle Mountain shear zone, and west-trending lineation that affects the late Oligocene Red Butte Canyon stock in the southern Grouse Creek Mountains (Compton et al., 1977; Todd, 1980), and may represent two or more separate colinear fabrics (Miller et al., 1983). Black Pine Mountains

Grouse Creek and Western Raft River Mountains In the Grouse Creek and western Raft River Mountains, at least two penetrative fabric sets are present. north-northeast– to north-trending lineation (with top-to-the-north-northeast shear, Malavieille, 1987b) is probably correlative with the Mesozoic northeast-trending lineation in the Albion Mountains. This lineation is overprinted by a younger west-northwest– to northwest-trending lineation that yields a top-to-west-northwest and northwest shear sense (Compton et al., 1977; Todd, 1980; Compton, 1983; Malavieille, 1987b). The ages and correlation of fabrics containing west-

A Cretaceous to Miocene deformation sequence is present in Devonian to Permian rocks in the Black Pine Mountains (Smith, 1982; Wells and Allmendinger, 1990), within the upper plate of the Raft River detachment fault, 15 km northeast of the eastern Raft River Mountains (Fig. 2). These strata record a structural history similar to that of Paleozoic strata within the Raft River Mountains described here (Wells et al., 1990). The oldest deformation produced a bedding-parallel foliation and east-trending elongation lineation. Strain and microstructural studies of this fabric document 160% east-west layer-parallel extension, layer-perpendicular shortening,

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Figure 16. D3 and D4 structures. (A) Photograph of folded D2 mylonitic foliation within Pennsylvanian(?) marble. Butt of hammer, 2 cm across for scale, in lower left. (B) Photomicrograph of pressure shadow from Chainman Shale at base of middle allochthon at Bald Knoll. View looking north. Asymmetric quartz-calcite pressure shadows around pyrite are ubiquitous within the shale and indicate westward shearing. Photomicrograph 5.6 mm across. Figure 15. XZ principal plane representation of shortening axes derived from Turner’s method for marbles within the lower allochthon. All samples from upright limbs of folds. RR112 and RR32 are Ordovician marbles; all other samples are from Pennsylvanian(?) marble. Great circle represents macroscopic foliation and the lineation is indicated by the filled circle. Contours of 2, 4, 6, 8, 10, and 12 sigma. Note that an alternative method to the Turner method (Dietrich and Song, 1984) was applied to several samples. The results of both methods yield similar orientations of kinematic axes, and the Turner method was chosen because of greater ease of presentation for a large number of measurements. (G) Representation of quadrants of infinitesimal shortening and extension in sinistral simple shear deformation. plane strain, and a coaxial strain path (Wells and Allmendinger, 1990; Wells et al., 1990). Microstructural and 40Ar/39Ar studies indicate that foliation formation and greenschist-facies metamorphism were of late Early Cretaceous age (Wells et al., 1990). West-facing recumbent folds deform the bedding-parallel foliation, and a second cleavage is locally developed within the fold hinge zones. Low-angle faults with top-to-west kinematics

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truncate the recumbent folds, and generally occur at low angles to strata and attenuate the stratigraphic section. For example, a low-angle fault places Mississippian Chainman–Diamond Peak Formation over Devonian Guilmette Formation and removes all but 2-m-thick structural lenses of Mississippian (Kinderhookian) limestone. DISCUSSION Significance of D1 Fabric Similar cooling ages and deformation kinematics favor correlation of D1 in the eastern Raft River Mountains with D1 in the Albion Mountains. The northeast trend of L1 in the eastern Raft River Mountains, in spite of much variability, resembles that of L1 in the Albion Mountains (Miller, 1980) and L1 in the Grouse Creek Mountains (Compton et al., 1977; Todd, 1980). Fabrics associated with these lineations in the Albion and Grouse Creek Mountains record a component of top-to-the-northeast and north shearing (Malavieille, 1987b) prograde metamorphism (Miller, 1980; Todd, 1980) and are thought to have developed in a contractional tectonic regime

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(Miller, 1980; Malavieille, 1987b). Seven Late Cretaceous K-Ar ages have been reported from rocks containing northeast-trending L1 lineations from the northern Albion Mountains (Armstrong, 1976; Miller, 1980). Three 40Ar/39Ar muscovite cooling ages from marble and schist from the lower allochthon in the eastern Raft River Mountains containing D1 fabrics are Late Cretaceous (Wells et al., 1990). The D1 fabrics are interpreted to have developed prior to cooling ca. 90 Ma. The D1 fabric in the lower allochthon of the eastern Raft River Mountains is interpreted as a prograde metamorphic fabric. Mineral-pair δ18O geothermometry of muscovite, biotite, and quartz that microscopically define the S1 foliation from three Ordovician samples consistently yield temperatures between 490 and 520 °C (Wells et al., 1990). The depth of stratigraphic burial for Middle Ordovician strata in this region is estimated to be about 10 to 11 km. Assuming conservative geothermal gradients of 25 to 30 °C/km, an additional 6 to 10 km of structural burial is required. The metamorphic assemblages from the Neoproterozoic schist of Mahogany Peaks (as outlined earlier) suggest pressures and temperatures in excess of 6.5 kbar and 600 °C, respectively (Spear and Cheney, 1989; Bohlen et al., 1991), also suggesting a doubling of the stratigraphic section. Thus, the metamorphic conditions indicate tectonic loading, not thermal metamorphism at stratigraphic burial depths (in contrast to metamorphism in the Pilot Range to the south; Miller and Hoisch, 1992). Work is in progress to determine more precisely the peak pressures and thus provide a better estimate of the amount of Mesozoic tectonic burial. The kinematics of general shear, with components of layer-perpendicular flattening and northeast-directed shear, together with the prograde nature of the deformation, are most compatible with deformation during and resulting from crustal contraction and thickening. The middle detachment fault places greenschist-facies rocks of the middle allochthon over middle amphibolite facies rocks of the lower allochthon, omits a significant amount of structural section, and is interpreted to have removed the bulk of the thrust sheet responsible for burial. The D1 fabric within the lower allochthon may record distributed simple shear and vertical flattening resulting from combined emplacement and gravitational spreading at the base of a thrust nappe. The northeast component of shear permits the possibility that there is a direct link between this deformation and eastward translation of rocks within the Sevier foreland. A preserved thrust relationship in the northern Albion Mountains may represent the remnants of a once areally extensive thrust nappe responsible for much of the burial evident in the eastern Raft River Mountains. At Mount Harrison in the northern Albion Mountains, an inverted sequence of Neoproterozoic rocks (quartzite assemblage, Fig. 1), unlike Neoproterozoic strata within the lower part of the Raft River Mountain sequence, structurally overlies Ordovician and older rocks of the Raft River Mountain sequence (Miller, 1980, 1983; Armstrong, 1968b). Thermobarometric studies of hanging-wall and footwall rocks (Hodges and McKenna, 1986) suggest that the fault places rocks with mineral assemblages yielding estimates of 4.80–5.4 kbar and 532–562 °C over rocks with pressure-temperature conditions of 3.5–4.1 kbar and 477–527 °C. Late Cretaceous K-Ar cooling ages (Armstrong, 1976) from rocks within this fault zone, combined with these thermobarometric estimates, suggest that this structure is a Mesozoic thrust fault (Hodges and McKenna, 1986; Malavieille, 1987b). Hodges and Walker (1992) suggested that the stratigraphic juxtaposition was initially of thrust sense, and that the latest movement along the fault was of normal sense in Late Cretaceous time. D2 Extension The west-directed D2 attenuation faults that place younger over older strata within the lower allochthon are interpreted to represent normal faults

(Wells et al., 1990). Although movement of thrusts across previously deformed strata can produce local younger over older relationships, the ubiquitous attenuation of strata together with the common progression of early plastic and later brittle shearing of similar kinematics is most consistent with the kinematics of extension. The age of normal faulting is constrained by 40Ar/39Ar muscovite cooling ages (82 to 90 Ma) from Ordovician rocks in the footwall of the Emigrant Spring fault, and an 40Ar/39Ar muscovite cooling age (88.5 Ma) from Pennsylvanian(?) marble tectonite within the D2 Emigrant Spring fault (Wells et al., 1990). The D2 fabrics are interpreted to have developed prior to or synchronous with cooling recorded by the muscovite cooling ages (Wells et al., 1990). The principal D2 attenuation faults in the eastern Raft River Mountains are apparently present throughout most of the RR–A–GC MMC and are thus of large displacement. A stratigraphic juxtaposition similar to that of the Emigrant Spring fault, including the distinctive mylonitic Pennsylvanian rocks, has been recognized on the western side of the Grouse Creek Mountains in the vicinity of Vipont Mountain (Fig. 2). The Mahogany Peaks fault (Compton 1972, 1975; Compton and Todd, 1979; Crittenden, 1979; Wells et al., 1996) crops out discontinuously for 70 km north to south, and for 50 km west to east in the RR–A–GC MMC. Exposures of the Mahogany Peaks fault in the eastern Raft River Mountains, in contrast to the Emigrant fault, do not unequivocally demonstrate whether it is deformed by D3 recumbent folds. Farther west, recumbent folds that deform the Mahogany Peaks fault may be of D3 fold generation, but this has not been clearly demonstrated. It is grouped here as a D2 structure solely due to its similarity in structural style to other D2 faults, but it may be of an entirely different age. Recumbent Folding and Kinematic Reversals The case for alternations in contraction and extension principally relies on the interpretation of the F3 folds. If the F3 folds developed during subhorizontal extension, then deformations D2 through D4 record a protracted history of extension, probably during progressive uplift and cooling, before the development of the Miocene Raft River detachment fault (Table 1). The alternative interpretation is that the folds developed during crustal shortening. In this case, the D1 fabrics and the D3 recumbent folds record shortening, whereas the D2 attenuation faults and the D4 middle detachment record extension (Table 1). Recumbent folds developed during crustal extension have been reported from several mountain belts (e.g., Malavielle, 1987a; Froitzheim, 1992; Mancktelow, 1992; Fletcher and Bartley, 1994). In the majority of examples, synextensional folds are developed within extensional shear zones, and most have fold hinge lines subparallel to the extension direction. These folds are generally interpreted to have developed either due to progressive rotation of fold hinge lines into parallelism with the shearing direction (Cobbold and Quinquis, 1980), or with hinge lines initially parallel to the extension direction, thereby recording a component of horizontal shortening orthogonal to the extension direction (Fletcher and Bartley, 1994; Manktelow and Pavlis, 1995). A tectonic environment for the formation of recumbent folds during crustal extension has also been postulated that does not require formation within high-strain shear zones. Froitzheim (1992), on the basis of observations of second-generation Alpine recumbent folds, suggested that synextensional folds can be developed within a strain field of subvertical shortening and subhorizontal extension if mechanical layering is initially steeply inclined. Such folds may form during either a coaxial or a noncoaxial strain path, and their hinge-line orientations may reflect the initial strike of layering. In the Raft River Mountains, kilometer-scale recumbent folds are developed within the parautochthon extensional shear zone and have hinge lines

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parallel to the extension direction (Compton, 1980; Sabisky, 1985; Malavieille, 1987a). These folds were suggested by Malavieille (1987a) to have formed as a result of large-magnitude top-to-east noncoaxial strain across layering inclined gently northward, combined with rotation of fold axes during noncoaxial shearing. The F3 recumbent folds within the lower allochthon, however, cannot have formed by the aforementioned folding mechanisms because they lack the necessary coeval fabrics recording high shear strains, and were therefore not developed within an extensional shear zone. In addition, there is no evidence for the presence of steeply dipping strata following D2 extensional shearing, as required for the folding mechanism proposed by Froitzheim (1992). On the contrary, the large tract of metamorphic rocks within the Raft River, Albion, and Grouse Creek Mountains, including many areas that have not been subjected to Cenozoic ductile shearing, is characterized by shallowly dipping lithologic layering and foliation. The exception to this is rocks within the upper plate of the Raft River detachment fault in the central and eastern Raft River Mountains that were significantly tilted during Miocene extension. Another potential mechanism for F3 fold formation during extension is large-scale drag folding related to movement on upper and/or lower bounding low-angle normal faults. The middle detachment fault forms an upper bounding surface to these folds. However, the middle detachment cuts across various structural levels within the F3 folds, and the fold limbs immediately beneath the middle detachment are neither consistently upright nor overturned (Fig. 6), suggesting that slip on the middle detachment was not coeval with fold formation. These folds do not have the appropriate geometries to have formed by reverse-drag or fault-bend fold mechanisms as a result of movement on structurally lower normal faults with listric or ramp-flat geometries (e.g., Gibbs, 1984; Ellis and McClay, 1988). The alternative and simplest interpretation for the formation of the F3 recumbent folds is that they record subhorizontal shortening related to crustal contraction. Recumbent folds having axes perpendicular to the transport direction are well recognized from the internal parts of most fold and thrust belts (e.g., Coward et al., 1989). In addition to the F3 folds in the eastern Raft River Mountains, recumbent folds that deform the two principal attenuation faults and may represent D3 structures are present in the northern Albion Mountains (Miller, 1980), at Vipont Mountain in the northern Grouse Creek Mountains, in the western Raft River Mountains (Compton, 1972), and in the southern Grouse Creek Mountains (Todd, 1980; Jordan, 1983). The vergence of the F3 recumbent folds within the lower allochthon in the eastern Raft River Mountains remains unclear because of their poorly understood geometry. Foliation and bedding generally dip westward, suggesting east-vergent folds. However, much of this westward tilt results from rotation concurrent with down-to-the-east normal faulting within the upper plate of the Raft River detachment. In addition, the lack of minor structures developed during folding precludes their use to elucidate folding geometry and kinematics. Determination of the vergence of these folds, however, is not crucial to their interpretation of contractional origin. If these folds are east vergent, their kinematics are easily explained within the context of an eastward-propagating orogenic belt. East-facing recumbent folds that deform attenuation faults have been mapped in the southern Grouse Creek (Jordan, 1983; Todd, 1980) and western Raft River Mountains (Compton, 1972). Alternatively, the F3 folds in the eastern Raft River Mountains may be west vergent, on the basis of the similarity between the Raft River and Black Pine structural sequences and the clear west vergence in the Black Pine recumbent folds (Wells et al., 1990; Wells and Allmendinger, 1990). West-verging contractional structures within the hinterland of an overall east-verging orogen, although seemingly problematic, can be explained in the context of back thrusting. Back thrusts are common in thrust belts and commonly occur in association with ramps (Serra, 1977; Butler, 1982).

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Geologic relations suggest a major subsurface Mesozoic ramp near the Black Pine Mountains and the possibility of a link between west-vergent contraction and ramp location. The Sublett synclinorium (Armstrong, 1982) occurs just east of the Black Pine Mountains and may represent the northern continuation of the Confusion Range synclinorium (Hose, 1977) of east-central Utah. The western limb of the regional synclinoria probably represents the eastern limb of a large ramp anticline (Armstrong, 1982; Von Tish et al., 1985). An additional Mesozoic west-verging structure, the Water Canyon anticline exposed in the Deep Creek Mountains (Rodgers, 1987), occurs along strike of this ramp. The position of the inferred crustal ramp may be relevant for another reason: the presence of Mesozoic extensional structures in close proximity to a major ramp permits a causal link. Crustal thickening and topographic uplift related to hanging-wall displacement over the ramp and probable footwall imbrication and duplex formation may lead to syncontractional extension, analogous to the effect of underplating on the dynamics of accretionary prisms (Platt, 1986). The kinematic succession recorded within the eastern Raft River Mountains suggests alternating extensional and contractional episodes during the Mesozoic to early Cenozoic evolution of the Sevier belt hinterland. This interpretation is attractive for several reasons. The documentation of extension of Late Cretaceous age in this region, as well as elsewhere within the Cordilleran hinterland (e.g., Hodges and Walker, 1992), suggests that the internal dynamics of the mountain belt were responsive to changes in the boundary conditions of the orogenic wedge, such as the width of the mountain belt, convergence rate, and degree of underplating and consequent topographic development (Platt, 1986; Dahlen and Suppe, 1988). This being the case, we would expect alternations from extension to contraction to accompany any further changes that upset the balance between compressional boundary stresses and gravitational buoyancy stresses. In addition, there are many conflicting reports of probable extensional and contractional structures of Mesozoic age in the northeastern Great Basin (e.g., Miller and Gans, 1989; Snoke and Miller, 1988; Wells et al., 1990). A model invoking reversals of contraction and extension would explain these apparent kinematic inconsistencies. Other localities within the Sevier belt hinterland where alternations in contraction and extension may be recorded include the Deep Creek Mountains of eastern Nevada. In the Deep Creek Mountains, post-peak metamorphic low-angle faults that attenuate the stratigraphic section are folded about the west-vergent Water Canyon recumbent anticline, which is in turn intruded by the latest Cretaceous Tungstonia granite (Rodgers, 1987; Miller et al., 1988; Nutt and Thorman, 1992, 1993). CONCLUSIONS The structural history recorded in Neoproterozoic to Permian rocks of the eastern Raft River Mountains suggests alternating contraction and extension in this region during late Mesozoic to early Cenozoic time. Alternating contraction and extension are predicted by theoretical and analog models of contractional orogenic development, and although previously unrecognized, may be common within the Sevier belt and other orogenic hinterlands. It remains to be demonstrated whether this kinematic sequence is more regionally applicable within the Sevier belt hinterland. Such a history would explain apparently conflicting reports of probable extensional and contractional structures of Mesozoic age in the northeastern Great Basin. Documentation of the timing of these kinematic alternations may provide a framework with which to link hinterland deformation with the development of the foreland thrust belt.

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CONTRACTION AND EXTENSION IN OROGENIC BELTS

ACKNOWLEDGMENTS This research was supported by the National Science Foundation (grant EAR-9317387 to Wells and grant EAR-8720952 to R. W. Allmendinger), the Geological Society of America, Sigma Xi, the Graduate School at Cornell University, and the National Research Council Postdoctoral Fellowship Program. Kenneth Tillman and Trenton Cladouhos provided very able assistance in the field. R. W. Allmendinger, P. A. Camilleri, D. M. Miller, I. Lucchitta, and V. R. Todd provided useful comments and suggestions on an earlier version of this manuscript. This paper was significantly improved by the collective detailed and critical reviews of J. M. Bartley, R. Fletcher, J. Fletcher, and B. John, although they may not agree with some of the conclusions. REFERENCES CITED Allmendinger, R. W., and Jordan, T. 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