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doi: 10.1130/0016-7606(1997)109<0107:ACAEIT>2.3.CO;2 1997;109;107-126Geological Society of America Bulletin

Michael L. Wells example from the Raft River Mountains, UtahAlternating contraction and extension in the hinterlands of orogenic belts: An

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ABSTRACT

Combined macroscopic to microscopic structural analyses, detailedgeologic mapping, and thermochronology were used to study thestructural evolution of midcrustal rocks of the Sevier orogenic belthinterland in northwestern Utah. These results, when combined withprevious studies, provide new insight into the structural history of thisregion, and suggest alternating tectonic contraction and extensionduring Mesozoic to early Cenozoic time.

Two allochthons form the upper plate of the Miocene Raft River de-tachment fault in the eastern Raft River Mountains. The lower alloch-thon comprises Neoproterozoic, Ordovician, and Pennsylvanian(?)strata, and is bounded below by the Raft River detachment fault andabove by the middle detachment fault. Strata within the lower alloch-thon were dramatically attenuated by two episodes of ductile defor-mation. The first deformation (D1) took place at metamorphic tem-peratures of ≈500 °C, and resulted in penetrative fabrics throughoutthese rocks that record combined flattening and top-to-northeastshearing strains. The second deformation (D2) resulted in significantstratigraphic attenuation along discrete top-to-the-west shear zonesthat are generally parallel to lithologic contacts. Separates of synkine-matic muscovite from the penetrative fabric yield 40Ar/39Ar coolingages that indicate that D1 deformation occurred prior to cooling ca.90 Ma. Both fabrics were subsequently folded about (D3) kilometer-scale recumbent folds.

The areally extensive middle allochthon, composed chiefly of Penn-sylvanian and Permian rocks metamorphosed in the greenschist fa-cies, was emplaced (D4) along the low-angle middle detachment fault.This fault cuts across various structural levels of the recumbentlyfolded lower allochthon in its footwall, and juxtaposes greenschist fa-cies over amphibolite-facies metamorphic rocks. The lower and mid-dle allochthon were subsequently deformed (D5) into open folds withnorth-trending axes. Neogene extension (D6) produced an ≈200-m-thick top-to-east ductile shear zone in Precambrian rocks, and formedthe younger Raft River detachment fault, which forms the present up-per contact of the ductile shear zone.

Northeast-vergent D1 fabrics probably record shortening deforma-tion, on the basis of fabric correlations with the Grouse Creek and Al-bion mountains, deformation kinematics, and synkinematic progrademetamorphism. D2 attenuation faults have been interpreted to recordcrustal extension of Late Cretaceous age. If D3 recumbent folds devel-oped during extension, then the deformation sequence records earlyshortening followed by protracted extensional unroofing with variablestructural styles. The favored alternative is that D3 recumbent folds

developed during shortening; in this case, D1 through D4 recordepisodic alternations of contraction and extension. Such alternationsare consistent with observations from analog and theoretical modelsof contractional mountain belts and suggest that the hinterland of theSevier orogenic belt underwent dynamic adjustments in crustal thick-ness and deformation kinematics in response to changes in the bound-ary conditions of the orogenic belt.

107

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

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

Figure 1. Generalized location and tectonic map of the northeasternGreat 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 geo-graphic and structural elements: AL, Albion Mountains; BP, BlackPine Mountains; CRS, Confusion Range synclinorium; DC, DeepCreek Range; GC, Grouse Creek Mountains; GSL, Great Salt Lake;G-T, Toano-Goshute Mountains; N, Newfoundland Mountains; P, Pi-lot Range; PM, Pequop Mountains; R-EH, Ruby Mountains–EastHumboldt Range; RR, Raft River Mountains; SR, Snake Range;SRPV, Snake River Plain volcanic rocks (shown in v pattern); SS, Sub-let synclinorium; WH, Wood Hills.*Email address: [email protected]



M. L. WELLS

108 Geological Society of America Bulletin, January 1997

INTRODUCTION AND CONCEPTUAL BACKGROUND

A complete description of the structural evolution of a mountain belt in-cludes temporal correlation between deformational events in its forelandand hinterland. Correlation between deformation events within the fore-land and hinterland of the Mesozoic to early Cenozoic Sevier orogenic beltof the western United States (Fig. 1) and comparison of their respectivekinematics have remained problematic, partly because Mesozoic structureshave been fragmented during Cenozoic extension, and isotopic systemsand metamorphic fabrics have been strongly overprinted during Tertiarymetamorphism, uplift, and cooling. Observations made within the pastdecade from analog (Davis et al., 1983) and theoretical models (Platt,1986; Molnar and Lyon-Caen, 1988) and field studies of contractionalmountain belts (Platt, 1986; Boyer, 1992; Wallis et al., 1993), however,may provide a framework to interpret hinterland structural evolution and tolink episodes of hinterland deformation to the evolution of the Sevier fore-land 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 ini-tial contractional deformation may progress at the regional scale toward theforeland, it is very likely that the hinterlands of orogenic belts continue tothicken internally as a response to the lengthening of the orogenic belt bypropagation of deformation toward the foreland (Davis et al., 1983; Platt,1986; Boyer, 1992). Furthermore, active large-scale extension is observedwithin the hinterlands of modern contractional mountain belts (e.g., Dal-mayrac and Molnar, 1981; Burchfiel and Royden, 1985). Ancient exam-ples of simultaneous hinterland extension and foreland shortening havenow been widely reported from many orogenic belts (e.g., Wells et al.,1990; Hodges et al., 1992a; Wallis et al., 1993). Although hinterland ex-tension does not drive foreland shortening, as was earlier suggested (Price,1973), syncontractional extension may be an integral part of the orogeniccycle. For example, changes in topographic slope, crustal thickness, rheol-ogy of the decollement, compressional boundary stresses, or erosion ratesin a dynamically evolving orogen can lead to extension within the hinter-land (Platt, 1986; Dahlen and Suppe, 1988). Factors that increase buoyancyforces arising from thickened and elevated crust (forces that tend to thin thecrust) relative to compressional boundary forces, which thicken crust, maylead to syncontractional extension. An outcome of the dynamic adjustmentof topography and crustal thickness is that alternating periods of contrac-tion and extension can occur during regional contraction (Platt, 1986).

The kinematic role of the Sevier belt hinterland during the late Mesozoicto early Cenozoic has been controversial, and the structural record is ap-parently contradictory (e.g., Hose and Danes, 1973; Allmendinger and Jor-dan, 1984; Snoke and Miller, 1988). For example, Mesozoic contractionalstructures (e.g., Miller and Gans, 1989; Snoke and Miller, 1988; Camilleri,1992; Taylor et al., 1993) and Mesozoic extensional structures (e.g., Wellset al., 1990; Hodges and Walker, 1992) have been reported in the north-eastern Great Basin. It is difficult, however, to compare the relative timingof many of these deformations from range to range because of the diffi-culty in tightly bracketing the ages for Mesozoic deformations in this re-gion (e.g., Snoke and Miller, 1988; Miller et al., 1988). Nonetheless, a his-tory of continued adjustments in hinterland crustal thickness (i.e.,alternating contraction and extension) during contraction in the forelandwould explain these apparent kinematic inconsistencies (Wells, 1991).Middle-crustal rocks exposed in metamorphic core complexes commonlyrecord a protracted history of Mesozoic and Cenozoic deformations (e.g.,Miller et al., 1988; Snoke and Miller, 1988). The Raft River Mountains arepart of a large metamorphic core complex that extends across the Raft

River, Albion, and Grouse Creek Mountains of northwestern Utah andsouthern Idaho (Figs. 1 and 2). This paper focuses on Mesozoic to earlyCenozoic deformation recorded in rocks in the upper plate of a Miocenelow-angle normal fault in the eastern Raft River Mountains. In contrast torocks exposed in the southern Albion and Grouse Creek Mountains, theserocks were far removed from Cenozoic plutons and were not thermallyoverprinted in Tertiary time, thus allowing isotopic dating of pre-Cenozoicdeformations. This paper synthesizes results of geologic mapping, macro-scopic and microscopic kinematic analysis, and thermochronologic stud-ies from the eastern Raft River Mountains. The results suggest that the hin-terland underwent alternating contraction and extension during lateMesozoic to early Cenozoic orogenesis.

REGIONAL TECTONIC SETTING OF THE RAFT RIVERMOUNTAINS

The Raft River Mountains are within the northeastern Great Basin, in thehinterland of the Mesozoic Sevier orogenic belt (Armstrong, 1968a), andeast of the outcrop belts affected by Paleozoic contractional orogeny. Post-Paleozoic deformation in the northeastern Great Basin can be broadly as-signed to Late Jurassic(?) to early Eocene development of the Sevier oro-genic belt, followed by middle(?) Eocene to recent crustal extension.

A belt of apparently surface-breaking contractional deformation in cen-tral Nevada (Ketner and Smith, 1974; Taylor et al., 1993) has been tenta-tively correlated with contractional structures in northeastern Nevada,northwestern Utah, and southern Idaho (Camilleri et al., 1992). Existingage brackets in central Nevada suggest that deformation both predated theearly Albian and was partly Late Cretaceous in age (Vandervoort andSchmidt, 1990; Taylor et al., 1993). In northeastern Nevada, the age con-straints vary from pre–Late Jurassic (Coats and Riva, 1983), to Late Juras-sic (Miller et al., 1987; Miller and Hoisch, 1992), to pre-Aptian (Thormanand Snee, 1988; Camilleri, 1992). Metamorphism and cleavage formationhave been interpreted to be of Jurassic and Late Cretaceous age in theSnake Range and vicinity (Miller et al., 1988; Miller and Gans, 1989). Inaddition, geobarometric studies of Mesozoic metamorphism from severaltracts 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 begunin Late Jurassic or Early Cretaceous time (e.g., Armstrong and Oriel, 1965;Wiltchko and Dorr, 1983; Heller et al., 1986, Royse, 1993). This definesseveral possible relations to the belt of contractional deformation in centraland northeastern Nevada and northwestern Utah. Hinterland contractionmay (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 theforeland thrust belt; or (4) have developed during several periods of defor-mation, both before and during tectonism in the foreland thrust belt.

Cretaceous extension of the Sevier belt hinterland, driven by lateral gra-dients in crustal thickness, was suggested based on studies in the easternRaft River and Black Pine Mountains (Wells and Allmendinger, 1990;Wells et al., 1990). Hodges and Walker (1992) outlined other localitieswhere Mesozoic extension is permissive (Hodges and Walker, 1992,Table 1), and suggested that late Mesozoic extension of the hinterland wasa fundamental process of Sevier orogenesis. The existing timing con-straints for Mesozoic hinterland extension do not permit temporal correla-tion to individual thrusting events within the well-established thrust se-quence in the foreland thrust belt (e.g., DeCelles, 1994), but do allow lateEarly to Late Cretaceous hinterland extension to be broadly bracketed asoccurring synchronous with overall shortening in the foreland.



CONTRACTION AND EXTENSION IN OROGENIC BELTS

Geological Society of America Bulletin, January 1997 109

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

Following regional contraction, crustal extension with associated plu-tonism and metamorphism affected the northeastern Great Basin from themiddle 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 orepisodic. Regionally, the age of initial Cenozoic extension decreases from

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



M. L. WELLS


Wells and Snee, 1993; Wells et al., 1994), suggesting that extension at thislatitude occurred only ≈5 to 10 m.y. after cessation of activity along thefrontal thrusts of the Sevier belt (Wiltschko and Dorr, 1983; Armstrong andOriel, 1965).

STRATIGRAPHY OF THE EASTERN RAFT RIVERMOUNTAINS

Detailed studies of the stratigraphy, structural geology, and metamor-phism have been carried out by many workers at various localities withinthe 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 fo-cused 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-temperatureconditions during metamorphism (Hodges and McKenna, 1986; Saltzerand Hodges, 1988). These studies document the complex deformation his-tory recorded in these rocks.

Overprinting Mesozoic and Cenozoic deformations have produced anincomplete and greatly attenuated stratigraphic section in the Raft River,Albion, and Grouse Creek Mountains. A tectonically thinned sequence ofmetasedimentary and sedimentary units of Neoproterozoic to Triassic ageoverlies 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 beensubdivided 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 alloch-thon. Although these allochthons are bounded by major faults in manyareas of the Raft River, Albion, and Grouse Creek Mountains, the faults orshear zones that bound them are not necessarily correlative in age or originfrom one locality to another. Until further work is carried out, the alloch-thon designations imply only stratigraphic affinity and general structuralposition, rather than structural continuity. A brief description of the rockunits 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 gneissicmonzogranite (metamorphosed adamellite of Compton, 1972) that in-trudes schist and amphibolite (Compton, 1975; Compton et al., 1977), allof which are unconformably overlain by the Elba Quartzite. The ElbaQuartzite forms the lowermost part of the Raft River Mountains sequenceof Miller (1983), and has been assigned various ages including Paleozoic,Neoproterozoic (Armstrong, 1968a; Compton et al., 1977; Compton andTodd, 1979), and Paleoproterozoic (Crittenden, 1979). The structurallyoverlying 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

of Clarks Basin resemble those measured in other Neoproterozoic rock se-quences in western North America (Wells et al., 1996). The Precambrianrocks in the parautochthon record andesine-amphibolite-facies metamor-phism (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 corre-lated with those described within the lower allochthon to the west (Comp-ton, 1975; Compton et al., 1977; Compton and Todd, 1979), with one ex-ception: a highly tectonized marble unit, which everywhere occursadjacent to Ordovician rocks, was recognized and correlated by physicalstratigraphy with the lower part of the Pennsylvanian Oquirrh Formation.Metamorphic mineral assemblages within the Neoproterozoic schist ofMahogany Peaks include garnet + staurolite + muscovite + biotite + pla-gioclase + 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 petroge-netic grid for the KFMASH system (Spear and Cheney, 1989, Fig. 2), andindicate minimum temperature and pressure of ≈600 °C and 6.5 kbar. Ex-tensive retrogradation of staurolite to chloritoid and muscovite, and of bi-otite to chlorite, took place at greenschist-facies conditions. Metamorphictemperatures in Ordovician rocks were estimated to be at least 490 to520 °C by oxygen isotopic geothermometry of quartz-muscovite andquartz-biotite mineral pairs and conodont color alteration index (CAI) val-ues >7 (Wells et al., 1990).

The middle allochthon is separated from the lower allochthon by themiddle detachment fault (Fig. 3). The middle allochthon mainly comprisesrocks of the Pennsylvanian and Permian Oquirrh Formation; a thin (0–5 mthick) sliver of Chainman Shale and Diamond Peak Formation is locallypresent along its base. Conodonts from the lower Oquirrh Formation yieldCAI values of 5, indicating temperatures of 350 to 400 °C greenschist-fa-cies metamorphic conditions (Wells et al., 1990).

STRUCTURAL GEOLOGY OF THE EASTERN RAFT RIVERMOUNTAINS

The Raft River detachment fault separates Precambrian rocks of the pa-rautochthon from the overlying lower and middle allochthons of the upperplate (Figs. 3 and 5). The upper-plate rocks in the eastern Raft River Moun-tains provide a unique opportunity within the Raft River–Albion–GrouseCreek metamorphic core complex to study a protracted history of Meso-zoic to early Cenozoic deformations for several reasons. (1) The rocksachieved peak metamorphic conditions of greenschist to middle amphibo-lite facies and contain multiple penetrative deformation fabrics that recordsuperposed deformations. (2) Upper-plate Neoproterozoic to Permianrocks were not overprinted by Neogene mylonitization and metamorphismthat affected lower-plate Archean and Proterozoic rocks, and have residedat 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

TABLE 1. SEQUENCE OF DEFORMATION, EASTERN RAFT RIVER MOUNTAINS

Deformation Structure and kinematics Interpretation Timingevent

D1 Flattening and top-to-northeast shear Contraction Pre-90 MaD2 Top-to-west attenuation faulting Extension Ca. 90 MaD3 Recumbent folding ContractionD4 Top-to-west displacement along middle detachment Extension Late Eocene–Oligocene (?)D5 Open folding about north-south axesD6 Top-to-east displacement along Raft River detachment and footwall shear zone Extension Middle to early late MioceneD7 Doming





levels in Paleogene time and are structurally overprinted by penetrativetop-to-the-west Paleogene mylonitic fabrics and locally thermally over-printed 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 acomplex deformation history. The sequence of deformations (Table 1) de-termined by detailed geologic mapping (Figs. 6, 7, and 8) includes the fol-lowing: (D1) attenuation of rock units within the lower allochthon by earlyintrabed plastic flow; (D2) attenuation of units within the lower allochthonby top-to-west younger-over-older faulting; (D3) recumbent folding of at-

tenuated strata and faults; (D4) emplacement of the middle allochthon overthe lower allochthon along the middle detachment; (D5) upright, open fold-ing of the middle and lower allochthons; (D6) top-to-east normal-senseshearing along the shear zone at the top of the parautochthon and along theRaft River detachment fault, with concomitant high-angle normal faultingof the upper plate; (D7) doming of the Raft River detachment fault and ear-lier structures.

The Neoproterozoic through Pennsylvanian(?) units within the lowerallochthon were markedly attenuated during superposed deformations. Thesection varies from 50 to 500 m in thickness (Fig. 4) whereas, in neighbor-

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



M. L. WELLS


ing ranges, equivalent strata are 4 to 7 km thick. The attenuation has beenachieved by penetrative plastic thinning (D1) during amphibolite faciesconditions and later (D2) low-angle, younger-over-older ductile and brittlefaulting (referred to as attenuation faulting, after Hintze, 1978).

D1 Deformation

The most pervasive fabric in the lower allochthon is a penetrative folia-tion (S1) that is generally parallel to lithologic layering and inferred bed-ding. At the few locations where foliation is distinct from bedding, foliationis slightly inclined to the southwest relative to bedding, whether in uprightor overturned rocks (Fig. 9A). Lineation varies in degree of development,and is not ubiquitous (note few lineation measurements as compared to fo-liation). The sparse development of lineation precludes the analysis of de-formation superposition using lineation-orientation analysis. Lineation is

best developed in marbles and generally trends northeast in both upright andoverturned rocks (Fig. 10). The only recognized F1 folds are sparse meso-scopic southeastward-overturned folds with axes generally trending north-northeast. With the exception of a few asymmetric structures that indicateeast and northeast shearing, the majority of strained mesoscopic features aresymmetric with respect to S1 foliation (including pressure shadows aroundrigid objects and boudinage of quartzite and chert interbeds). This observa-tion, coupled with the locally apparent southwest dip of foliation relative tobedding, suggests that the D1 fabrics in carbonate rocks resulted from acombination of pure shear and northeast-directed simple shear. This kine-matic interpretation is supported by the microstructural investigations ofcalcitic marble and quartzite outlined as follows.

Calcitic Marbles. The lineation in marble is most commonly defined byelongate pressure shadows of quartz and calcite about pyrite grains, elon-gate calcite grains, and stretched and recrystallized fossil fragments. Thefoliation is microscopically defined by a grain-shape fabric of plasticallydeformed calcite and a parallel alignment of muscovite.

A large body of work has documented the development of crystallo-graphic preferred orientations in calcite-bearing rocks. Experimental stud-ies and computer simulation of coaxial (e.g., Turner et al., 1956; Casey etal., 1978; Wagner et al., 1982; Wenk et al., 1987) and noncoaxial (Rutterand Rusbridge, 1977; Kern and Wenk, 1983; Schmid et al., 1987; Wenk etal., 1987) deformation of limestone have determined expected crystallo-graphic preferred orientations (e.g., c-axes). In addition to the strain pathand the shear sense, the relative magnitudes of pure and simple shear in adeformation event can be derived from the crystallographic fabric (Dietrichand Song, 1984; Wenk et al., 1987; Schmid et al., 1987). The c-axis [0001]preferred orientation for two samples of Ordovician marble containing D1fabrics was determined by measurement of mutually perpendicular thinsections on the universal stage (Fig. 11, A and B). Both samples exhibit astrong preferred orientation of c-axes. The pole figures are slightly asym-metric with respect to the foliation normal (obliquities of ≈16°). The c-axisfabric obliquities (Wenk et al., 1987) suggest about equal components oftop-to-the-northeast simple shear and pure shear deformation, and a gen-eral shear strain path.

Quartzites. The microstructures in quartzites suggest deformation kine-matics similar to those of the calcitic marbles. Ordovician quartzite ex-hibits a weak macroscopic S1 foliation and typically no lineation. Neoprot-erozoic quartzite of Clarks Basin is more highly foliated and lineated andwas studied to elucidate the kinematics of D1 deformation. An east- tonortheast-trending stretching lineation (L1) is defined by elongate mus-covite and quartz grains, and foliation is defined by aligned muscovite anda grain-shape fabric of dynamically recrystallized quartz. Microscopically,quartz grain boundaries range from relatively straight with 120° triple junc-tions to curved to embayed (Fig. 9C). Many features similar to those de-scribed by Jessel (1987) indicate pinning of mobile quartz grain boundariesby muscovite and recrystallization by grain-boundary migration (Urai etal., 1986), consistent with deformation at amphibolite-facies metamorphicconditions. The more elongate recrystallized grains exhibit discrete extinc-tion domains with sharp domain boundaries, representing prismatic sub-grains (White, 1973). Subgrain boundaries lie at angles of 42° to 68° fromthe principal foliation, the large majority being inclined toward the south-west relative to principal foliation.

The preferred orientation of quartz c-axes was analyzed in three sam-ples. The c-axis fabrics are highly variable (Figs. 12 and 13, A–C). SampleRR91-14 exhibits a small circle distribution (with an opening angle of≈30°) about the pole to the foliation (Fig. 13A). Small circle distributionshave been produced under laboratory conditions of uniaxial flattening(Tullis et al., 1973) and measured in studies of naturally deformed rocksthat record coaxial flattening strains (e.g., Law et al., 1984). Sample

Figure 4. Comparison between attenuated stratigraphic thicknesseswithin the eastern Raft River Mountains and representative strati-graphic thicknesses from nearby localities in northwestern Utah (fromHintze, 1988). Of — Ordovician Fish Haven Dolomite, see Figure 3 forother abbreviations.





RR91-93 (Fig. 13B) exhibits a c-axis fabric intermediate between an asym-metric 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 asymmetricsingle girdle fabric (Fig. 13C) recording noncoaxial top-to-east-northeastshear. These samples are from different localities and different structuralhorizons within the outcrop of the quartzite of Clarks Basin, and no attemptwas made to determine the deformation partitioning with respect to struc-tural level. Elementary strain compatibility arguments (e.g., Law and Potts,1987), however, suggest that the differences in c-axis pole figures may re-late to a strain field with different amounts of northeast-directed noncoax-ial strain superimposed on uniform coaxial strain. For strain compatibilitybetween structural levels, the components of coaxial strain (to a first order)should be uniform for all structural levels, to avoid the development of ma-terial overlaps or gaps. Variability with structural level in the simple shearcomponent, however, will not result in development of material gaps oroverlaps. Taken together, the D1 fabrics within the carbonate and quartzitetectonites indicate components of coaxial flattening and noncoaxial top-to-northeast shearing strain.

D2 Attenuation Faults and Fabrics

Faults. Two major discontinuities in stratigraphy within the lowerallochthon mark faults. The structurally higher fault, the Emigrant Springfault (Wells, 1996; LAF1 of Wells, 1992; see Figs. 3 and 6), places Penn-sylvanian(?) calcitic marble over Ordovician and Silurian(?) dolomiticmarble throughout the mapped area. The Emigrant Spring fault removesabout 5 km of stratigraphic section, including Silurian, Devonian, and allbut centimeter-scale slivers of Mississippian rocks. This discontinuity is in-terpreted to be a fault, contrary to the earlier interpretation of an unconfor-mity (Armstrong, 1968b), because (1) Devonian and Mississippian strataare present in neighboring ranges, including the Black Pine, southernGrouse Creek, and Newfoundland Mountains, and the Pilot Range (Fig. 1;Hintze, 1988); (2) Pennsylvanian(?) marble lying above this discontinuitycontains a distinct mylonite zone parallel to the contact; (3) this contact islocally structurally discordant (Fig. 6); and (4) the removal of rock units bylow-angle faulting is characteristic of deformation at other stratigraphiclevels within the lower allochthon.

The second major discontinuity separates the schist of Mahogany Peaks

Figure 5. Generalized geologic map of the eastern Raft River Mountains indicating axial traces of F3 recumbent folds and F5 open folds. Notethat Pennsylvanian(?) marble commonly occurs within the cores of recumbent synclines. Stereoplots are equal area Pi-diagrams of poles to fo-liation from localities of F5 folds indicated by arrows; pole to best-fit great circle (fold axis) indicated by square. The outlined box indicates lo-cation of detailed map of Figure 6.



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eological Society of Am

erica Bulletin,January 1997

Figure 6. Detailed geologic map of the Crystal Peak area, eastern Raft River Mountains. The map relationships clearly demonstrate the struc-tural sequence of deformations. Location of cross section A–A′ indicated. Pennsylvanian(?) marble (IPot) crops out in cores of F3 recumbentfolds. See Figure 7 for explanation of lithologic symbols.

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and quartzite of Clarks Basin from the Ordovician carbonate rocks. Thefaulted nature of this contact has been clarified using carbon isotopes, con-firming a Neoproterozoic age for the quartzite of Clarks Basin, in contrastto the Cambrian age previously suggested (Compton et al., 1977; Comp-ton and Todd, 1979). This discontinuity, termed the Mahogany Peaks fault,places Ordovician over Neoproterozoic strata, omitting >2 km of strati-graphic section (Wells, 1992; Wells et al., 1996). The regional-scale geom-etry of these two large-displacement faults and the consistency in hanging-wall and footwall stratigraphic levels over large distances suggest that the

exposed segments of these faults are bedding-parallel flats in low-anglefaults with ramp and flat geometries.

In addition to the two major low-angle faults described above, the ma-jority of other formational contacts within the lower allochthon are low-angle faults that place younger rocks on older. These attenuation faults,with few exceptions, are subparallel to foliation and bedding in adjacentunits and are identified using both textural and stratigraphic criteria. Inplaces, there is moderate discordance between the fault zones and the adja-cent units, and between rock units on either side of the faults (Fig. 6). Sliv-ers of units are commonly strung out along the faults and, locally, units arecompletely omitted. In places, the faults are recognized by ductile high-strain zones localized along bedding-parallel contacts between strati-graphic units along which section is commonly missing. Where the faultednature of stratigraphic contacts is not evident by field inspection, signifi-cant lateral changes in map-scale unit thicknesses attest to their presence.

Kinematic Analysis. Within attenuation fault or shear zones, fabricsrange from ductile to brittle. In many localities, mylonitic fabrics are cut bysystems of normal faults of similar kinematics, suggesting a progressionfrom early ductile to later brittle shearing. Foliation within the shear zonesis subparallel to S1 foliation, except in rare localities where S1 foliation isclearly truncated by the D2 shear zones. This general parallelism of S1 andS2 foliations makes it difficult to define the margins of D2 shear zones.

A plastic shear zone is present within the lower part of the Pennsylvan-ian Oquirrh Formation(?) marble adjacent to the Emigrant Spring fault. Themarble is highly foliated and locally contains a well-developed east-trend-ing stretching lineation (Fig. 14, A and B) and isoclinal folds having hingelines parallel to lineation. A deformation gradient is present within this shearzone. The basal part is ultramylonitic, grading upward to more coarse-grained mylonitic marble. The ultramylonite is 1–2 m thick, and is charac-terized by extreme reduction in grain size to 8 to 20 µm, untwinned, equantcalcite grains. Grain size increases structurally upward to 2 mm within themylonite, and foliation and lineation are uniformly oriented across this tran-

Figure 7. Geologic cross section. See Figure 6 for location of cross-section line A–A′. Note D2 attenuation faults that omit the stratigraphic sec-tion are folded about F3 recumbent folds. D2 attenuation faults, F3 recumbent fold axial surfaces, and the middle detachment are folded aboutF5 open folds.

Figure 8. Photograph looking northeast toward Crystal Peak, illus-trating the overprinting relationships between various structures. Theprominent gently dipping white rock unit in the footwall of the struc-turally lowest fault (single barb, the Raft River detachment fault) isthe Elba Quartzite. The Raft River detachment truncates the broad F5upright anticline. The D2 attenuation faults (open teeth) and the D4middle detachment fault (double barbs) are folded by the broad F5fold. Field of view about 1.5 km across.



M. L. WELLS


sition. The marble ultramylonite contains sparse porphyroclasts of calcitethat are slightly elongate parallel to a foliation defined by a parallel align-ment of fine-grained muscovite. The deformation kinematics of the marbleultramylonite and more coarse-grained marble mylonite were investigatedby study of lattice-preferred orientations. Samples collected from over-turned fold limbs exhibit somewhat diffuse and complex c-axis orientationsand no systematic relationship between c-axes and macroscopic fabric ele-ments, suggesting significant intracrystalline straining after foliation for-mation, probably during folding. Because of this, only samples from uprightfold limbs that are probably less overprinted by folding strain are includedin the kinematic analysis. The lattice-preferred orientation of the ultramy-lonite was determined from X-ray pole figure goniometry. Calcite a-axesexhibit a strong preferred orientation (Fig. 11G). The a-axis girdle and in-ferred c-axis maxima are asymmetric with respect to foliation, indicatingnoncoaxial top-to-west shear. The more coarse-grained mylonitic marbleswere measured on the universal stage. All samples collected from uprightlimbs to D3 recumbent folds exhibit a strong preferred orientation of calcitec-axes at a high angle to foliation (Fig. 11, C–F). The asymmetries of the c-axis maxima relative to the normal to the grain-shape fabric are not uni-formly consistent and, therefore, analysis of all marble samples did not in-dicate a uniform shear sense. However, the samples were collected atdifferent structural levels above the ultramylonite, and the variations in thec-axis fabric may reflect deformation partitioning within the D2 shear zone,different degrees of D2 overprint on D1 fabric, or combinations of both. Dur-

ing e-twinning, the c-axis of the twin undergoes a change in orientation of52° 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 fo-liation development, c-axis reorientation during twinning (measured twinvolumes of 5% to 25% of the total grain) might have affected the earlier pre-ferred orientations. Nevertheless, the majority of measured samples exhibitlattice-preferred orientations consistent with a significant component ofwest-directed simple shear.

In addition to determination of the c-axis texture, a method of “dy-namic” analysis of calcite twin lamellae pioneered by F. J. Turner (1953)was employed. Although this “dynamic” method was originally interpretedto yield compressional and tensional stress axes, they are more aptly re-garded as incremental strain axes. Therefore they are referred to as short-ening and extension axes. The c-axis maxima are expected to correspondclosely to the compression direction that produced the grain-shape fabricand crystallographic preferred orientation (see Wells and Allmendinger,1990, for a review).

There is a strong deviation between shortening axes determined by theTurner method and the c-axis maxima in the samples of Pennsylvanian(?)and Ordovician marble. This discordance is interpreted to indicate over-printing deformations, and the twinning strain is interpreted to postdate thedevelopment of the predominant c-axis fabric and foliation. The late twinstrain is geometrically related to foliation (Fig. 15) rather than the presentgeographic reference frame (not shown), because of the significantly

Figure 9. D1 fabrics. (A) Highly transposed bedding within SwanPeak limestone in view looking north. Foliation is inclined west rela-tive to bedding. 12.5 cm eraser for scale. (B) Boudinaged quartz veinthat lies at low angle to foliation, illustrating large-magnitude intrabedplastic flow. Hammer for scale. (C) Photomicrograph of preserved D1fabric within quartzite of Clarks Basin. Muscovite defines S1 foliation.Note morphology of quartz grain boundaries, including embayed andstraight grain boundaries with 120° triple junctions, indicating grainboundary migration recrystallization. Photomicrograph 3 mm across.





greater consistency of kinematic axes between samples in the former pro-jection. The twin strain could have resulted from D2 westward shearing,which occurred along shear zones oriented subparallel to D1 foliation. Thekinematic axes are compatible with this interpretation; when viewed in anapproximately east-west projection perpendicular to foliation (Fig. 15), themajority of the samples exhibit Turner shortening axes consistent with fo-liation-subparallel westward shearing. The absence of twins in the marbletectonite that underwent dynamic recrystallization during D2 shearing sup-ports the interpretation that the twin strain in the more coarse-grained mar-ble is related to D2 deformation. Alternatively, the twin strain may recordwestward shearing related to westward translation of the middle allochthonduring D4, or foliation-parallel shear during F3 folding.

Other west-vergent shear indicators from the basal marble tectonite ofthe lower Oquirrh Formation include a secondary recrystallized grain-shape fabric present in coarser-grained marble that is oblique to and in-clined eastward relative to the mylonitic foliation, and deformed fold limbsthat are boudinaged where inclined eastward and folded where inclinedwestward relative to foliation. Common late brittle fault networks consis-tently indicate the same sense of shearing.

West-directed mylonitic shear zones, ranging in thickness from 1 to150 cm, are recognized within greatly attenuated Swan Peak and Eurekaquartzites. Mylonitic foliation is defined in thin section by an alignment of

elongate to ribbonlike quartz grains that are oblique to discontinuous tocontinuous zones of complete dynamic recrystallization and concentratedshear; this fabric is interpreted to represent an S-C fabric, and indicates top-to-west shearing (Fig. 14C). The ribbon quartz grains are not recrystallizedby recovery processes, and exhibit undulatory extinction and subbasal de-formation lamellae, indicating greenschist-facies metamorphic conditions.Within the subparallel zones of dynamic recrystallization defining C-sur-faces, 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 ex-tensional crenulation cleavage (Platt and Vissers, 1980) indicate westwardshearing.

Quartz lattice studies from mylonitic Eureka Quartzite substantiate thekinematics determined from the microscopic- and mesoscopic-scale fea-tures. Two samples were measured on a universal stage for determination ofthe preferred orientation of c-axes (Fig. 13, D and E). Sample RR38 showsan asymmetric single girdle fabric recording noncoaxial top-to-west shear(Fig. 13D). The c-axis maxima at positions close to the primitive circle sug-gest dislocation glide along basal planes in <a>, consistent with greenschist-facies metamorphic conditions (Schmid and Casey, 1986). SampleRR91-28 (Fig. 13E) exhibits an asymmetric type II crossed girdle (Lister,1977) with a leading edge (Behrmann and Platt, 1982) interpreted to record

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



M. L. WELLS


a significant component of noncoaxial, top-to-west shear (Schmid andCasey, 1986).

Relative timing relations between D1 fabric and D2 attenuation faults areclearer in the eastern Raft River Mountains than in areas to the west. TheD2 attenuation faults are interpreted to postdate penetrative foliation devel-opment (D1) for the following reasons: (1) attenuation faults locally trun-cate D1 foliation; (2) quartzite fabrics within attenuation faults are dis-

tinctly lower temperature than D1 fabrics and are unrecovered; and (3) D1fabrics have distinct kinematics from the attenuation fault zones.

D3 Recumbent Folding (F2)

The Ordovician through Pennsylvanian(?) rocks of the lower allochthonare deformed into tight to isoclinal recumbent folds with amplitudesgreater than 1.5 to 2 km. The complex superposition of later deformations,however, precludes determination of original fold geometry. The youngestunit involved in recumbent folding is the Pennsylvanian(?) marble, whichoccupies the cores of recumbent synclines (Figs. 6 and 16A). The D1 foli-ation consistently dips more steeply southwestward than transposed bed-ding on both upright and overturned limbs, a relationship that is inconsis-tent with the foliation forming synchronously with folding, but permits thefoliation to either predate or postdate folding. The D1 foliation is inter-preted to predate folding because F3 recumbent folds deform D2 attenua-tion faults, and D2 faults locally truncate and overprint the D1 foliation. Thelack of a recognizable foliation associated with F3 folding in the easternRaft River Mountains may be due to relatively dry, postmetamorphic con-ditions that were unfavorable for low-temperature deformation mecha-nisms, such as fluid-assisted grain boundary diffusion.

The F3 recumbent folds are bracketed as younger than D2 attenuationfaults and older than the D4 middle detachment. Attenuation faults occur inboth upright and overturned limbs and remove the same amount of strati-graphic section on both limbs (e.g., the Emigrant Spring fault, Figs. 3, 6,and 7); therefore, recumbent folding is interpreted to postdate attenuationfaulting within the lower allochthon. The middle allochthon rests on vari-ous structural levels of both upright and overturned fold limbs (Figs. 6 and7), indicating that F3 recumbent folding preceded emplacement of the mid-dle allochthon.

D4 Emplacement of the Middle Allochthon Along the MiddleDetachment

Greenschist-facies, Pennsylvanian to Permian Oquirrh Formation rocksof the middle allochthon overlie amphibolite-facies, Neoproterozoic toPennsylvanian(?) rocks of the lower allochthon along the middle detach-ment fault. The transport direction for this fault is determined from study ofa thin sliver of Chainman Shale and Diamond Peak Formation that is lo-cally present along the base of the middle allochthon. A 5-m-thick layer ofblack shale is exposed at Bald Knoll (Fig. 2). Pressure shadows aroundpyrite are ubiquitous and impart a strong west-southwest–trending stretch-ing lineation in this shale. The pressure shadows are strongly asymmetricrelative to foliation as viewed parallel to foliation and perpendicular to lin-eation (Fig. 16B). The morphologies of the fibers within the pressure shad-ows are similar to those described by Etchecopar and Malavieille (1988)for rigid, face-controlled fibers growing around euhedral pyrite during non-coaxial flow. These pressure shadows are interpreted as recording top-to-the-west-southwest shear. Because this shale occurs as a thin sliver at thebase of the middle allochthon, the kinematics of deformation within theshale are thought to record distributed simple shear along the middle de-tachment. Westward translation of the middle allochthon is consistent withdata 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 atmetamorphic conditions no higher than lower greenschist facies. This isevident from the metamorphic discontinuity across this fault in the studyarea and farther to the west (Compton et al., 1977), and the deformationmechanisms within the sheared base of the middle allochthon: pressure so-lution and brittle fracturing of quartz and pressure solution and low-tem-

Figure 11. Calcite pole figures for marbles within the lower alloch-thon. Diagrams A through F are c-axis pole figures determined frommeasurement on a universal stage. (G) is X-ray derived a-axis pole fig-ure. The east-west great circle represents the macroscopic foliationplane and the lineation is indicated by the filled circle. RR112 (A) andRR32 (B) are Ordovician marbles; all other samples are from Penn-sylvanian(?) marble. RR3 (G) is a marble ultramylonite. Contours forc-axis pole figures are 2, 4, 6, 8, 10, and 12 sigma; contours for a-axispole figure (G) are of 1, 1.5, 2.0, and 2.5 of random distribution.





perature plasticity of calcite. The displacement along the middle detach-ment was, therefore, probably late metamorphic to postmetamorphic.

The middle detachment fault can be traced discontinuously across theeast-west extent of the Raft River Mountains and into the Grouse Creek andAlbion Mountains. This fault was interpreted by Compton et al. (1977) torepresent the principal detachment in the Raft River–Albion–Grouse Creekmetamorphic core complex (RR-A-GC MCC), and have top-to-east dis-placement. This fault, as exposed in the eastern Raft River Mountains, is in-terpreted to represent an extensional detachment fault with westward trans-lation, because it removes stratigraphic section, juxtaposes lower grade onhigher grade metamorphic rocks (Compton, 1977; Wells, 1992), and locallyrecords top-to-the-west kinematics. This fault may be related to deeper leveltop-to-west extensional plastic shear of Eocene–Oligocene age in the west-ern 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 ofCrystal Peak, an F5 open fold with a wavelength greater than 1500 m iswell 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 apparentlytruncated by the Raft River detachment.

The D5 folds could have formed contemporaneous with D6 extensionalshear. However, the occurrence of D5 folds in fold trains (as opposed tosingle folds), upright axial surfaces, and small wavelengths for some ofthese folds suggests that they are not folds related to the geometry of un-derlying normal faults such as reverse-drag or fault-bend folds. Further-more, these folds are cut by high-angle faults (Fig. 7) that in turn are trun-cated by the Raft River detachment fault, indicating that the folds formedprior 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 pa-rautochthon there is an ≈200-m-thick shear zone that is parallel to beddingin the Proterozoic units and the underlying unconformity with Archeanrocks. The shear zone accommodated large-scale top-to-the-east displace-ment (Compton, 1980; Sabisky, 1985; Malavieille, 1987a; Wells, 1992),and strain intensity progressively increases from west to east. Myloniticfabrics are the most highly developed within the Elba Quartzite and itsschist member, and fabric intensity related to eastward shearing typicallydies out abruptly downward within the Green Creek complex. Northward-overturned recumbent folds are developed within the extensional shear

Figure 12. Location map for microstruc-tural samples of Figures 11, 13, and 15. SeeFigure 5 for explanation of patterns andsymbols.



M. L. WELLS


zone. Fold hinge lines are parallel to the extension direction and greaterthan 20 km in length (Compton, 1980; Malavieille, 1987a).

Superimposed on the shear zone in the parautochthon is the brittle RaftRiver detachment fault. This major structural discontinuity forms the up-per contact of the schist member of the Elba Quartzite. The detachmentfault is concordant with foliation and bedding within the lower plate, buttruncates structures within the upper plate.

There is a close spatial and kinematic association between the detach-ment fault and the underlying shear zone. Fabrics within the uppermostshear zone are commonly retrograde and, where a cataclasite is present, themylonite is progressively overprinted by cataclastic deformation struc-turally upward toward the detachment fault. Both plastic and brittle struc-tures exhibit the same top-to-the-east shear sense.

The upper plate of the Raft River detachment fault is cut by numeroushigh-angle faults with displacements from centimeters to about 2 km. Thestrikes of the faults vary greatly, but the majority strike roughly north. Onlytwo faults, both of small displacement, were noted that cut the Raft Riverdetachment fault. The upper-plate high-angle faults, therefore, formed ei-ther prior to or during movement of the Raft River detachment fault.

Thermochronological results from within and beneath the shear zone inthe parautochthon contrast markedly with the Late Cretaceous muscoviteages from the lower allochthon (Wells et al., 1990). Late-early to early-lateMiocene cooling ages (40Ar/39Ar from muscovite, biotite, and microcline,Wells and Snee, 1993; fission track in apatite, Wells et al., 1994) collectivelyindicate rapid cooling during the Miocene, during footwall unroofing re-lated to extensional displacement. The mylonite zone has been previouslyinterpreted to represent a Mesozoic thrust-sense shear zone (Malavielle andCobb, 1986; Snoke and Miller, 1988), and Cenozoic normal-sense shear

zone (Malavielle, 1987a; Wells, 1992). Our studies indicate that plasticshearing occurred within the temperature window spanned by rapid cool-ing, as indicated by the deformation conditions recorded in the microstruc-tures (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 structuralrelief. The shear zone, the strata in the parautochthon, and the detachmentfault outline this large structure. The axis of the elongate anticline is sub-parallel to the transport direction of both the footwall shear zone and thebrittle detachment fault. Nonplanar detachment faults of comparablegeometry are present within many other core complexes, in particular corecomplexes 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 thenorthern Albion Mountains (Miller, 1980). The earlier exhibits a northeast-trending lineation (L1 of Miller, 1980), with a component of top-to-northeastshear (Malavielle, 1987b), and the later a northwest-trending lineation (L2 ofMiller, 1980) and top-to-northwest shear (Malavieille, 1987b). Rocks con-taining these fabrics yielded Mesozoic and Cenozoic conventional K-Arcooling ages (biotite, muscovite, and hornblende), eight of which range from

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 EurekaQuartzite. The east-west great circle represents the macroscopic foliation plane; the filled circle indicates lineation.





66 to 81 Ma (Armstrong, 1976). Hodges and Walker (1992) interpreted theL2 fabric to record Cretaceous down-to-the-northwest extensional shear. Athird lineation, trending west-northwest, is present in the southwestern Al-bion Mountains within the Middle Mountain shear zone, a west-dipping ex-tensional shear zone with top-to-west-northwest shear sense (Miller et al.,1983; Saltzer and Hodges, 1988). The Middle Mountain shear zone isyounger than L2 of Miller (1980), and metamorphic minerals that grew dur-ing shearing yield late Eocene to Oligocene (K-Ar and 40Ar/39Ar) coolingages (Armstrong, 1976; Miller et al., 1983; Saltzer and Hodges, 1988).

Grouse Creek and Western Raft River Mountains

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

to northwest–trending lineation in the Grouse Creek and western RaftRiver Mountains are problematic. They exhibit similarity in trends andkinematics to the Mesozoic(?) L2 of Miller (1980) in the northern AlbionMountains, the late Eocene to early Oligocene west-northwest–trendinglineation of the Middle Mountain shear zone, and west-trending lineationthat affects the late Oligocene Red Butte Canyon stock in the southernGrouse Creek Mountains (Compton et al., 1977; Todd, 1980), and mayrepresent two or more separate colinear fabrics (Miller et al., 1983).

Black Pine Mountains

A Cretaceous to Miocene deformation sequence is present in Devonianto Permian rocks in the Black Pine Mountains (Smith, 1982; Wells and All-mendinger, 1990), within the upper plate of the Raft River detachmentfault, 15 km northeast of the eastern Raft River Mountains (Fig. 2). Thesestrata record a structural history similar to that of Paleozoic strata withinthe Raft River Mountains described here (Wells et al., 1990). The oldestdeformation produced a bedding-parallel foliation and east-trending elon-gation lineation. Strain and microstructural studies of this fabric document160% east-west layer-parallel extension, layer-perpendicular shortening,

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



M. L. WELLS


plane strain, and a coaxial strain path (Wells and Allmendinger, 1990;Wells et al., 1990). Microstructural and 40Ar/39Ar studies indicate that fo-liation formation and greenschist-facies metamorphism were of late EarlyCretaceous age (Wells et al., 1990). West-facing recumbent folds deformthe bedding-parallel foliation, and a second cleavage is locally developedwithin the fold hinge zones. Low-angle faults with top-to-west kinematics

truncate the recumbent folds, and generally occur at low angles to strataand attenuate the stratigraphic section. For example, a low-angle faultplaces Mississippian Chainman–Diamond Peak Formation over DevonianGuilmette Formation and removes all but 2-m-thick structural lenses ofMississippian (Kinderhookian) limestone.

DISCUSSION

Significance of D1 Fabric

Similar cooling ages and deformation kinematics favor correlation of D1in the eastern Raft River Mountains with D1 in the Albion Mountains. Thenortheast trend of L1 in the eastern Raft River Mountains, in spite of muchvariability, resembles that of L1 in the Albion Mountains (Miller, 1980) andL1 in the Grouse Creek Mountains (Compton et al., 1977; Todd, 1980).Fabrics associated with these lineations in the Albion and Grouse CreekMountains 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

Figure 15. XZ principal plane representation of shortening axes de-rived from Turner’s method for marbles within the lower allochthon.All samples from upright limbs of folds. RR112 and RR32 are Ordovi-cian marbles; all other samples are from Pennsylvanian(?) marble.Great circle represents macroscopic foliation and the lineation is indi-cated by the filled circle. Contours of 2, 4, 6, 8, 10, and 12 sigma. Notethat an alternative method to the Turner method (Dietrich and Song,1984) was applied to several samples. The results of both methods yieldsimilar orientations of kinematic axes, and the Turner method was cho-sen because of greater ease of presentation for a large number of mea-surements. (G) Representation of quadrants of infinitesimal shorteningand extension in sinistral simple shear deformation.

Figure 16. D3 and D4 structures. (A) Photograph of folded D2 my-lonitic foliation within Pennsylvanian(?) marble. Butt of hammer, 2 cmacross for scale, in lower left. (B) Photomicrograph of pressure shadowfrom Chainman Shale at base of middle allochthon at Bald Knoll.View looking north. Asymmetric quartz-calcite pressure shadowsaround pyrite are ubiquitous within the shale and indicate westwardshearing. Photomicrograph 5.6 mm across.





(Miller, 1980; Malavieille, 1987b). Seven Late Cretaceous K-Ar ages havebeen reported from rocks containing northeast-trending L1 lineations fromthe northern Albion Mountains (Armstrong, 1976; Miller, 1980). Three40Ar/39Ar muscovite cooling ages from marble and schist from the lowerallochthon in the eastern Raft River Mountains containing D1 fabrics areLate Cretaceous (Wells et al., 1990). The D1 fabrics are interpreted to havedeveloped prior to cooling ca. 90 Ma.

The D1 fabric in the lower allochthon of the eastern Raft River Moun-tains is interpreted as a prograde metamorphic fabric. Mineral-pair δ18Ogeothermometry of muscovite, biotite, and quartz that microscopically de-fine the S1 foliation from three Ordovician samples consistently yield tem-peratures between 490 and 520 °C (Wells et al., 1990). The depth of strati-graphic burial for Middle Ordovician strata in this region is estimated to beabout 10 to 11 km. Assuming conservative geothermal gradients of 25 to30 °C/km, an additional 6 to 10 km of structural burial is required. Themetamorphic assemblages from the Neoproterozoic schist of MahoganyPeaks (as outlined earlier) suggest pressures and temperatures in excess of6.5 kbar and 600 °C, respectively (Spear and Cheney, 1989; Bohlen et al.,1991), also suggesting a doubling of the stratigraphic section. Thus, themetamorphic conditions indicate tectonic loading, not thermal metamor-phism at stratigraphic burial depths (in contrast to metamorphism in the Pi-lot Range to the south; Miller and Hoisch, 1992). Work is in progress to de-termine more precisely the peak pressures and thus provide a betterestimate of the amount of Mesozoic tectonic burial.

The kinematics of general shear, with components of layer-perpendicularflattening and northeast-directed shear, together with the prograde nature ofthe deformation, are most compatible with deformation during and result-ing from crustal contraction and thickening. The middle detachment faultplaces greenschist-facies rocks of the middle allochthon over middle am-phibolite facies rocks of the lower allochthon, omits a significant amount ofstructural section, and is interpreted to have removed the bulk of the thrustsheet responsible for burial. The D1 fabric within the lower allochthon mayrecord distributed simple shear and vertical flattening resulting from com-bined emplacement and gravitational spreading at the base of a thrust nappe.The northeast component of shear permits the possibility that there is a di-rect link between this deformation and eastward translation of rocks withinthe Sevier foreland.

A preserved thrust relationship in the northern Albion Mountains mayrepresent the remnants of a once areally extensive thrust nappe responsiblefor much of the burial evident in the eastern Raft River Mountains. AtMount Harrison in the northern Albion Mountains, an inverted sequence ofNeoproterozoic rocks (quartzite assemblage, Fig. 1), unlike Neoproterozoicstrata within the lower part of the Raft River Mountain sequence, struc-turally overlies Ordovician and older rocks of the Raft River Mountain se-quence (Miller, 1980, 1983; Armstrong, 1968b). Thermobarometric studiesof hanging-wall and footwall rocks (Hodges and McKenna, 1986) suggestthat the fault places rocks with mineral assemblages yielding estimates of4.80–5.4 kbar and 532–562 °C over rocks with pressure-temperature con-ditions 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 thesethermobarometric estimates, suggest that this structure is a Mesozoic thrustfault (Hodges and McKenna, 1986; Malavieille, 1987b). Hodges andWalker (1992) suggested that the stratigraphic juxtaposition was initially ofthrust sense, and that the latest movement along the fault was of normalsense in Late Cretaceous time.

D2 Extension

The west-directed D2 attenuation faults that place younger over olderstrata within the lower allochthon are interpreted to represent normal faults

(Wells et al., 1990). Although movement of thrusts across previously de-formed strata can produce local younger over older relationships, the ubiq-uitous attenuation of strata together with the common progression of earlyplastic and later brittle shearing of similar kinematics is most consistentwith the kinematics of extension. The age of normal faulting is constrainedby 40Ar/39Ar muscovite cooling ages (82 to 90 Ma) from Ordovician rocksin the footwall of the Emigrant Spring fault, and an 40Ar/39Ar muscovitecooling age (88.5 Ma) from Pennsylvanian(?) marble tectonite within theD2 Emigrant Spring fault (Wells et al., 1990). The D2 fabrics are interpretedto have developed prior to or synchronous with cooling recorded by themuscovite cooling ages (Wells et al., 1990).

The principal D2 attenuation faults in the eastern Raft River Mountainsare apparently present throughout most of the RR–A–GC MMC and arethus of large displacement. A stratigraphic juxtaposition similar to that ofthe Emigrant Spring fault, including the distinctive mylonitic Pennsylva-nian rocks, has been recognized on the western side of the Grouse CreekMountains 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 for70 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 Moun-tains, in contrast to the Emigrant fault, do not unequivocally demonstratewhether it is deformed by D3 recumbent folds. Farther west, recumbentfolds 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 struc-ture solely due to its similarity in structural style to other D2 faults, but itmay be of an entirely different age.

Recumbent Folding and Kinematic Reversals

The case for alternations in contraction and extension principally relieson the interpretation of the F3 folds. If the F3 folds developed during sub-horizontal extension, then deformations D2 through D4 record a protractedhistory of extension, probably during progressive uplift and cooling, beforethe development of the Miocene Raft River detachment fault (Table 1). Thealternative interpretation is that the folds developed during crustal shorten-ing. In this case, the D1 fabrics and the D3 recumbent folds record shorten-ing, whereas the D2 attenuation faults and the D4 middle detachment recordextension (Table 1).

Recumbent folds developed during crustal extension have been reportedfrom several mountain belts (e.g., Malavielle, 1987a; Froitzheim, 1992;Mancktelow, 1992; Fletcher and Bartley, 1994). In the majority of exam-ples, synextensional folds are developed within extensional shear zones,and most have fold hinge lines subparallel to the extension direction. Thesefolds are generally interpreted to have developed either due to progressiverotation of fold hinge lines into parallelism with the shearing direction(Cobbold and Quinquis, 1980), or with hinge lines initially parallel to theextension direction, thereby recording a component of horizontal shorten-ing orthogonal to the extension direction (Fletcher and Bartley, 1994;Manktelow and Pavlis, 1995). A tectonic environment for the formation ofrecumbent folds during crustal extension has also been postulated that doesnot require formation within high-strain shear zones. Froitzheim (1992), onthe basis of observations of second-generation Alpine recumbent folds,suggested that synextensional folds can be developed within a strain fieldof subvertical shortening and subhorizontal extension if mechanical layer-ing is initially steeply inclined. Such folds may form during either a coax-ial or a noncoaxial strain path, and their hinge-line orientations may reflectthe initial strike of layering.

In the Raft River Mountains, kilometer-scale recumbent folds are devel-oped within the parautochthon extensional shear zone and have hinge lines



M. L. WELLS


parallel to the extension direction (Compton, 1980; Sabisky, 1985; Mala-vieille, 1987a). These folds were suggested by Malavieille (1987a) to haveformed as a result of large-magnitude top-to-east noncoaxial strain acrosslayering inclined gently northward, combined with rotation of fold axesduring noncoaxial shearing. The F3 recumbent folds within the lowerallochthon, however, cannot have formed by the aforementioned foldingmechanisms because they lack the necessary coeval fabrics recording highshear strains, and were therefore not developed within an extensional shearzone. In addition, there is no evidence for the presence of steeply dippingstrata following D2 extensional shearing, as required for the folding mech-anism proposed by Froitzheim (1992). On the contrary, the large tract ofmetamorphic rocks within the Raft River, Albion, and Grouse CreekMountains, including many areas that have not been subjected to Cenozoicductile shearing, is characterized by shallowly dipping lithologic layeringand foliation. The exception to this is rocks within the upper plate of theRaft River detachment fault in the central and eastern Raft River Moun-tains that were significantly tilted during Miocene extension. Another po-tential mechanism for F3 fold formation during extension is large-scaledrag folding related to movement on upper and/or lower bounding low-an-gle normal faults. The middle detachment fault forms an upper boundingsurface to these folds. However, the middle detachment cuts across variousstructural levels within the F3 folds, and the fold limbs immediately be-neath the middle detachment are neither consistently upright nor over-turned (Fig. 6), suggesting that slip on the middle detachment was not co-eval with fold formation. These folds do not have the appropriategeometries to have formed by reverse-drag or fault-bend fold mechanismsas a result of movement on structurally lower normal faults with listric orramp-flat geometries (e.g., Gibbs, 1984; Ellis and McClay, 1988).

The alternative and simplest interpretation for the formation of the F3recumbent folds is that they record subhorizontal shortening related tocrustal contraction. Recumbent folds having axes perpendicular to thetransport direction are well recognized from the internal parts of most foldand thrust belts (e.g., Coward et al., 1989). In addition to the F3 folds inthe eastern Raft River Mountains, recumbent folds that deform the twoprincipal attenuation faults and may represent D3 structures are present inthe northern Albion Mountains (Miller, 1980), at Vipont Mountain in thenorthern 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 theeastern Raft River Mountains remains unclear because of their poorly un-derstood geometry. Foliation and bedding generally dip westward, suggest-ing east-vergent folds. However, much of this westward tilt results from ro-tation concurrent with down-to-the-east normal faulting within the upperplate of the Raft River detachment. In addition, the lack of minor structuresdeveloped during folding precludes their use to elucidate folding geometryand kinematics. Determination of the vergence of these folds, however, isnot crucial to their interpretation of contractional origin. If these folds areeast vergent, their kinematics are easily explained within the context of aneastward-propagating orogenic belt. East-facing recumbent folds that de-form 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 maybe west vergent, on the basis of the similarity between the Raft River andBlack Pine structural sequences and the clear west vergence in the BlackPine recumbent folds (Wells et al., 1990; Wells and Allmendinger, 1990).West-verging contractional structures within the hinterland of an overalleast-verging orogen, although seemingly problematic, can be explained inthe context of back thrusting. Back thrusts are common in thrust belts andcommonly occur in association with ramps (Serra, 1977; Butler, 1982).

Geologic relations suggest a major subsurface Mesozoic ramp near theBlack Pine Mountains and the possibility of a link between west-vergentcontraction and ramp location. The Sublett synclinorium (Armstrong,1982) occurs just east of the Black Pine Mountains and may represent thenorthern continuation of the Confusion Range synclinorium (Hose, 1977)of east-central Utah. The western limb of the regional synclinoria probablyrepresents the eastern limb of a large ramp anticline (Armstrong, 1982; VonTish et al., 1985). An additional Mesozoic west-verging structure, the Wa-ter 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 anotherreason: the presence of Mesozoic extensional structures in close proximityto a major ramp permits a causal link. Crustal thickening and topographicuplift related to hanging-wall displacement over the ramp and probablefootwall imbrication and duplex formation may lead to syncontractionalextension, analogous to the effect of underplating on the dynamics of ac-cretionary prisms (Platt, 1986).

The kinematic succession recorded within the eastern Raft River Moun-tains suggests alternating extensional and contractional episodes during theMesozoic to early Cenozoic evolution of the Sevier belt hinterland. This in-terpretation is attractive for several reasons. The documentation of exten-sion of Late Cretaceous age in this region, as well as elsewhere within theCordilleran hinterland (e.g., Hodges and Walker, 1992), suggests that theinternal dynamics of the mountain belt were responsive to changes in theboundary conditions of the orogenic wedge, such as the width of the moun-tain belt, convergence rate, and degree of underplating and consequent top-ographic development (Platt, 1986; Dahlen and Suppe, 1988). This beingthe case, we would expect alternations from extension to contraction to ac-company any further changes that upset the balance between compres-sional boundary stresses and gravitational buoyancy stresses. In addition,there are many conflicting reports of probable extensional and contrac-tional structures of Mesozoic age in the northeastern Great Basin (e.g.,Miller and Gans, 1989; Snoke and Miller, 1988; Wells et al., 1990). Amodel invoking reversals of contraction and extension would explain theseapparent kinematic inconsistencies. Other localities within the Sevier belthinterland where alternations in contraction and extension may be recordedinclude the Deep Creek Mountains of eastern Nevada. In the Deep CreekMountains, post-peak metamorphic low-angle faults that attenuate thestratigraphic section are folded about the west-vergent Water Canyon re-cumbent anticline, which is in turn intruded by the latest CretaceousTungstonia granite (Rodgers, 1987; Miller et al., 1988; Nutt and Thorman,1992, 1993).

CONCLUSIONS

The structural history recorded in Neoproterozoic to Permian rocks ofthe eastern Raft River Mountains suggests alternating contraction and ex-tension in this region during late Mesozoic to early Cenozoic time. Alter-nating contraction and extension are predicted by theoretical and analogmodels of contractional orogenic development, and although previouslyunrecognized, may be common within the Sevier belt and other orogenichinterlands.

It remains to be demonstrated whether this kinematic sequence is moreregionally applicable within the Sevier belt hinterland. Such a historywould explain apparently conflicting reports of probable extensional andcontractional structures of Mesozoic age in the northeastern Great Basin.Documentation of the timing of these kinematic alternations may providea framework with which to link hinterland deformation with the develop-ment of the foreland thrust belt.





ACKNOWLEDGMENTS

This research was supported by the National Science Foundation (grantEAR-9317387 to Wells and grant EAR-8720952 to R. W. Allmendinger),the Geological Society of America, Sigma Xi, the Graduate School at Cor-nell University, and the National Research Council Postdoctoral Fellow-ship Program. Kenneth Tillman and Trenton Cladouhos provided very ableassistance in the field. R. W. Allmendinger, P. A. Camilleri, D. M. Miller, I.Lucchitta, and V. R. Todd provided useful comments and suggestions on anearlier version of this manuscript. This paper was significantly improvedby 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 theconclusions.

REFERENCES CITED

Allmendinger, R. W., and Jordan, T. E., 1984, Mesozoic structure of the Newfoundland Mountains, Utah: Hor-izontal shortening and subsequent extension in the hinterland of the Sevier belt: Geological Society ofAmerica Bulletin, v. 95, p. 1280–1292.

Allmendinger, R. W., Miller, D. M., and Jordan, T. E., 1984, Known and inferred Mesozoic deformation in thehinterland of the Sevier belt, northwestern Utah, in Kerns, G. J., and Kerns, R. L., eds., Geology of north-western Utah, southern Idaho and northeast Nevada: Utah Geological Association Publication 13,p. 21–34.

Armstrong, F. C., and Oriel, S. S., 1965, Tectonic development of the Idaho-Wyoming thrust belt: American As-sociation of Petroleum Geologists Bulletin, v. 49, p. 1847–1866.

Armstrong, R. L., 1968a, Sevier orogenic belt in Nevada and Utah: Geological Society of America Bulletin,v. 79, p. 429–458.

Armstrong, R. L., 1968b, Mantled gneiss domes in the Albion Range, southern Idaho: Geological Society ofAmerica Bulletin, v. 79, p. 1295–1314.

Armstrong, R. L., 1976, The geochronology of Idaho (Part 2): Isochron/West, no. 15, p. 1–33.Armstrong, R. L., 1982, Cordilleran metamorphic core complexes—From Arizona to southern Canada: Annual

Reviews of Earth and Planetary Science, v. 10, p. 129–154.Armstrong, R. L., and Hills, F. A., 1967, Rb-Sr and K-Ar geochronologic studies of mantled gneiss domes, Al-

bion Range, southern Idaho, USA: Earth and Planetary Science Letters, v. 3, p. 114–124.Armstrong, R. L., and Ward, P., 1991, Evolving geographic patterns of Cenozoic magmatism in the North Amer-

ican Cordillera: The temporal and spatial association of magmatism and metamorphic core complexes:Journal of Geophysical Research, v. 96, p. 13201–13224.

Axen, G. J., Taylor, W. J., and Bartley, J. M., 1993, Space-time patterns and tectonic controls of Tertiary exten-sion and magmatism in the Great Basin of the western United States: Geological Society of America Bul-letin, v. 105, p. 56–76.

Bartley, J. M., and Gleason, G., 1990, Tertiary normal faults superimposed on Mesozoic thrusts, Quinn Canyonand Grant Ranges, Nye County, Nevada, in Wernicke, B. P., ed., Basin and Range extensional tectonicsnear the latitude of Las Vegas, Nevada: Geological Society of America Memoir 176, p. 195–212.

Behrmann, J. H., and Platt, J. P., 1982, Sense of nappe emplacement from quartz c-axis fabrics; an example fromthe Betic Cordilleras (Spain): Earth and Planetary Science Letters, v. 49, p. 208–215.

Bohlen, S. R., Montanas, A., and Kerrick, D. M., 1991, Precise determination of the equilibria kyanite-silli-manite and kyanite-andalusite and a revised triple point for Al2SiO5 polymorphs: American Mineralogist,v. 76, p. 677–680.

Boyer, S. E., 1992, Geometric evidence for synchronous thrusting in the Alberta and northwest Montana thrustbelts, in McClay, K. R., ed., Thrust tectonics: London, Chapman & Hall, p. 377–390.

Boyer, S. E., and Elliot, D., 1982, Thrust systems:American Association of Petroleum Geologists Bulletin, v. 66,p. 1196–1230.

Burchfiel, B. C., and Royden, L. H., 1985, North-south extension within the convergent Himalayan region: Ge-ology, v. 13, p. 679–682.

Butler, R. W., 1982, The terminology of structures in thrust belts: Journal of Structural Geology, v. 4, p. 239–245.Camilleri, P. A., 1992, Mesozoic structural and metamorphic features in the Wood Hills and Pequop Mountains,

northeastern Nevada, in Wilson, J. R., ed., Field guide to geologic excursions in Utah and adjacent areasof Nevada, Idaho, and Wyoming: Utah Geological Survey Miscellaneous Publication 92-3, p. 93–105.

Camilleri, P. A., Miller, D. M., Snoke, A. W., and Wells, M. L., 1992, Structure and fabric of metamorphic ter-rains in the northeastern Great Basin: Implications for Mesozoic crustal shortening and extension, in Wil-son, J. R., ed., Field guide to geologic excursions in Utah and adjacent areas of Nevada, Idaho, andWyoming: Utah Geological Survey Miscellaneous Publication 92-3, p. 19–26.

Casey, M., Rutter, E. H., Schmid, S. M., Siddans, A. W. B., and Whalley, J. S., 1978, Texture development in ex-perimentally deformed calcite rocks, in Proceedings of the 5th International Conference on Textures ofMaterials: Berlin, Springer-Verlag, p. 231–240.

Coats, R. R., and Riva, J. F., 1983, Overlapping overthrust belts of late Paleozoic and Mesozoic ages, northernElko County, Nevada, in Miller, D. M., Todd, V. R., and Howard, K. A., eds., Tectonic and stratigraphicstudies in the eastern Great Basin: Geological Society of America Memoir 157, p. 305–327.

Cobbold, P. R., and Quinquis, H., 1980, Development of sheath folds in shear regimes: Journal of Structural Ge-ology, v. 5, p. 383–399.

Compton, R. R., 1972, Geologic map of Yost quadrangle, Box Elder County, Utah, and Cassia County, Idaho:U.S. Geological Survey Miscellaneous Geologic Investigations Map I-672, scale 1:31 680.

Compton, R. R., 1975, Geologic map of Park Valley quadrangle, Box Elder County, Utah, and Cassia County,Idaho: U.S. Geological Survey Miscellaneous Geologic Investigations Map I-873, scale 1:31 680.

Compton, R. R., 1980, Fabrics and strains in quartzites of a metamorphic core complex, Raft River Mountains,Utah, in Crittenden, M. D., Jr., Coney, P. J., and Davis, G. H., eds., Cordilleran metamorphic core com-plexes: Geological Society of America Memoir 153, p. 385–398.

Compton, R. R., 1983, Displaced Miocene rocks on the west flank of the Raft River–Grouse Creek core com-plex, Utah, in Miller, D. M., Todd, V. R., and Howard, K. A., eds., Tectonic and stratigraphic studies in theeastern Great Basin: Geological Society of America Memoir 157, p. 271–279.

Compton, R. R., and Todd, V. R., 1979, Oligocene and Miocene metamorphism, folding, and low-angle faultingin northwestern Utah: Reply: Geological Society of America Bulletin, v. 90, p. 307–309.

Compton, R. R., Todd, V. R., Zartman, R. E., and Naeser, C. W., 1977, Oligocene and Miocene metamorphism,folding, and low-angle faulting in northwestern Utah: Geological Society of America Bulletin, v. 88,p. 1237–1250.

Coward, M. P., Dietrich, D., and Park, R. G., 1989, eds., Alpine tectonics: London, Geological Society SpecialPublication 45, 450 p.

Crittenden, M. D., Jr., 1979, Oligocene and Miocene metamorphism, folding, and low-angle faulting in north-western Utah: Discussion: Geological Society of America Bulletin, v. 90, p. 305–306.

Dahlen, F. A., and Suppe, J., 1988, Mechanics, growth, and erosion of mountain belts, in Clark, S. P., Jr., Burch-fiel, B. C., and Suppe, J., eds., Processes in continental lithospheric deformation: Geological Society ofAmerica Special Paper 218, p. 161–178.

Dallmeyer, R. D., Snoke, A. W., and McKee, E. H., 1986, The Mesozoic-Cenozoic tectonothermal evolution ofthe Ruby Mountains, East Humboldt Range, Nevada: A Cordilleran metamorphic core complex: Tecton-ics, v. 5, p. 931–954.

Dalmayrac, B., and Molnar, P., 1981, Parallel thrust and normal faulting in Peru and the constraints on the stateof stress: Earth and Planetary Science Letters, v. 55, p. 473–481.

Davis, D., Suppe, J., and Dahlen, F. A., 1983, Mechanics of fold-and-thrust belts and accretionary wedges: Jour-nal of Geophysical Research, v. 88, p. 1153–1172.

DeCelles, P. G., 1994, Late Cretaceous–Paleocene synorogenic sedimentation and kinematic history of the Se-vier thrust belt, northeast Utah and southwest Wyoming: Geological Society of America Bulletin, v. 106,p. 32–56.

Dietrich, D., and Song, H., 1984, Calcite fabrics in a natural shear environment, the Helvetic nappes of westernSwitzerland: Journal of Structural Geology, v. 6, p. 19–32.

Duebendorfer, E. M., and Wallin, E. T., 1991, Basin development and syntectonic sedimentation associated withkinematically coupled strike-slip and detachment faulting, southern Nevada: Geology, v. 19, p. 87–90.

Ellis, P. G., and McClay, K. R., 1988, Listric extensional fault systems—Results of analogue model experiments:Basin Research, v. 1, p. 55–70.

Etchecopar, A., and Malavieille, J., 1988, Computer models of pressure shadows: A method for strain measure-ment and shear-sense determination: Journal of Structural Geology, v. 9, p. 667–677.

Fletcher, J. M., and Bartley, J. M., 1994, Constrictional strain in a non-coaxial shear zone: Implications for foldand rock fabric development, central Mojave metamorphic core complex, California: Journal of StructuralGeology, v. 16, p. 555–570.

Froitzheim, N., 1992, Formation of recumbent folds during synorogenic crustal extension (Austroalpine nappes,Switzerland): Geology, v. 20, p. 923–926.

Gibbs, A. D., 1984, Structural evolution of extensional basin margins: Geological Society of London Journal,v. 141, p. 609–620.

Handin, J. W., and Griggs, D. T., 1961, Deformation of the Yule marble, part II: Geological Society of AmericaBulletin, v. 62, p. 863–886.

Heller, P. L., Bowdler, S. S., Chambers, H. P., Coogan, J. C., Hagen, E. S., Shuster, M. W., Winslow, N. S., andLawton, T. F., 1986, Time of initial thrusting in the Sevier orogenic belt, Idaho-Wyoming and Utah: Ge-ology, v. 14, p. 388–391.

Hintze, L. F., 1978, Sevier orogenic attenuation faulting in the Fish Springs and House ranges, western Utah:Brigham Young University Geology Studies, v. 25, part 1, p. 11–24.

Hintze, L. F., 1988, Geologic history of Utah: Brigham Young University Geology Studies, Special Publica-tion 7, 202 p.

Hodges, K. V., and McKenna, L., 1986, Structural and metamorphic characteristics of the Raft River-Quartziteallochthon juxtaposition, Albion Mountains, southern Idaho: Geological Society of America Abstractswith Programs, v. 18, p. 117.

Hodges, K. V., and Walker, J. D., 1992, Extension in the Cretaceous Sevier orogen, North American Cordillera:Geological Society of America Bulletin, v. 104, p. 560–569.

Hodges, K. V., Parrish, R. R., Housh, T. B., Lux, D. R., Burchfiel, B. C., Royden, L. H., and Chen, Z., 1992a, Si-multaneous Miocene extension and shortening in the Himalayan Orogen: Science, v. 258, p. 1466–1470.

Hodges, K. V., Snoke, A. W., and Hurlow, H. A., 1992b, Thermal evolution of a portion of the Sevier hinterland:The northern Ruby Mountains–East Humboldt Range and Wood Hills, northeastern Nevada: Tectonics,v. 11, p. 154–164.

Hose, R. K., 1977, Structural geology of the Confusion Range, west-central Utah: U.S. Geological Survey Pro-fessional Paper 971, 9 p.

Hose, R. K., and Danes, Z. F., 1973, Development of the late Mesozoic to early Cenozoic structures of the east-ern Great Basin, in de Jong, K. A., and Scholten, R., eds., Gravity and tectonics: New York, J. Wiley andSons, p. 429–441.

Jessel, M. W., 1987, Grain boundary migration microstructures in a naturally deformed quartzite: Journal ofStructural Geology, v. 9, p. 1007–1014.

Jones, P. B., 1971, Folded faults and sequence of thrusting in Alberta Foothills: American Association of Petro-leum Geologists Bulletin, v. 55, p. 292–306.

Jordan, T. E., 1983, Structural geometry and sequence, Bovine Mountain, northwestern Utah, in Miller, D. M.,Todd, V. R., and Howard, K. A., eds., Tectonic and stratigraphic studies in the eastern Great Basin: Geo-logical Society of America Memoir 157, p. 215–227.

Kern, H., and Wenk, H. R., 1983, Texture development in experimentally induced ductile shear zones: Contri-butions to Mineralogy and Petrology, v. 83, p. 231–236.

Ketner, K. B., and Smith, J. F., 1974, Folds and overthrusts of Late Jurassic or Early Cretaceous age in northernNevada: U.S. Geological Survey Journal of Research, v. 2, p. 417–419.

Law, R. D., and Potts, J. D., 1987, The Tarskavaig nappe of Skye, northwest Scotland: A re-examination of thefabrics and their kinematic significance: Geological Magazine, v. 124, p. 231–248.

Law, R. D., Knipe, R. J., and Dayan, H., 1984, Strain path partitioning within thrust sheets: Microstructural andpetrofabric evidence from the Moine thrust zone at Loch Eriboll, northwest Scotland: Journal of Struc-tural Geology, v. 6, p. 477–498.

Lee, J., Miller, E. L., and Sutter J. F., 1987, Ductile strain and metamorphism in an extensional tectonic setting—A case study from the northern Snake Range, Nevada, U.S.A, in Coward, M. P., Dewey, J. F., and Han-cock, P. L., eds., Continental extensional tectonics: Geological Society of London Special Publication 28,p. 267–298.

Lewis, C. J., Selverstone, J., Wernicke, B., and Bartley, J. M., 1992, Deep burial depths for metapelites from thelower plate of the northern Snake Range decollement, Nevada: Geological Society of America Abstractswith Programs, v. 24, no. 6, p. 23.

Lister, G. S., 1977, Crossed girdle c-axis fabrics in quartzites plastically deformed by plane strain and progres-sive simple shear: Discussion: Tectonophysics, v. 39, p. 51–54.

Lister, G. S., and Snoke, A. W, 1984, S-C mylonites: Journal of Structural Geology, v. 6, p. 617–639.Malavieille, J., 1987a, Extensional shearing and kilometer-scale “a” type folds in a Cordilleran metamorphic

core complex (Raft River Mountains, northwestern Utah): Tectonics, v. 6, p. 423–448.Malavieille, J., 1987b, Kinematics of compressional and extensional ductile shearing deformation in a meta-

morphic core complex of the northeastern Basin and Range: Journal of Structural Geology, v. 9,p. 541–554.

Malavieille, J., and Cobb, F., 1986, Cinematique des deformations ductiles dans trois massifs metamorphiquesde l’Ouest des Etats-Unis: Albion (Idaho), Raft River et Grouse Creek (Utah): Geological Society ofFrance Bulletin, p. 885–898.

Mancktelow, N. S., 1992, Neogene lateral extension during convergence in the Central Alps: Evidence from inter-related faulting and backfolding around the Simplonpass (Switzerland): Tectonophysics, v. 215, p. 295–317.

Mancktelow, N. S., and Pavlis, T. L., 1995, Fold-fault relationships in low-angle detachment systems: Tectonics,v. 13, p. 668–685.



M. L. WELLS


Manning, A. H., and Bartley, J. M., 1994, Postmylonitic deformation in the Raft River metamorphic core com-plex, northwestern Utah: Evidence of a rolling hinge: Tectonics, v. 13, p. 596–612.

Miller, D. M., 1980, Structural geology of the northern Albion Mountains, south-central Idaho, in Crittenden,M. D., Jr., Coney, P. J., and Davis, G. H., eds., Cordilleran metamorphic core complexes: Geological So-ciety of America Memoir 153, p. 399–423.

Miller, D. M., 1983, Allochthonous quartzite sequence in the Albion Mountains, Idaho, and proposed Protero-zoic Z and Cambrian correlatives in the Pilot Range, Utah and Nevada, in Miller, D. M., Todd, V. R., andHoward, K. A., eds., Tectonic and stratigraphic studies in the eastern Great Basin: Geological Society ofAmerica Memoir 157, p. 191–213.

Miller, D. M., and Hoisch, T., 1992, Mesozoic structure, metamorphism, and magmatism in the Pilot Range andthe Toano Range, in Wilson, J. R., ed., Field guide to geologic excursions in Utah and adjacent areas ofNevada, Idaho, and Wyoming: Utah Geological Survey Miscellaneous Publication 92-3, p. 79–92.

Miller, D. M., Armstrong, R. L., Compton, R. R., and Todd, V. R., 1983, Geology of the Albion–RaftRiver–Grouse Creek Mountains area, northwestern Utah and southern Idaho: Utah Geological and Min-eral Survey Special Studies, v. 59, p. 1–59.

Miller, D. M., Hillhouse, W. C., Zartman, R. E., and Lanphere, M. A., 1987, Geochronology of intrusive andmetamorphic rocks in the Pilot Range, Utah and Nevada, and comparison with regional patterns: Geo-logical Society of America Bulletin, v. 99, p. 880–885.

Miller, E. L., and Gans, P. B., 1989, Cretaceous crustal structure and metamorphism in the hinterland of the Se-vier thrust belt, western U.S. Cordillera: Geology, v. 17, p. 59–62.

Miller, E. L., Gans, P. B., Wright, J. E., and Sutter, J. F., 1988, Metamorphic history of east-central Basin andRange Province, in Ernst, W. G., ed., Metamorphism and crustal evolution of the western United States:Englewood Cliffs, New Jersey, Prentice-Hall, Rubey Volume VII, p. 649–682.

Molnar, P., and Lyon-Caen, H., 1988, Some simple physical aspects of the support, structure, and evolution ofmountain belts, in Clark, S. P., Jr., Burchfiel, B. C., and Suppe, J., eds., Processes in continental lithos-pheric deformation: Geological Society of America Special Paper 218, p. 197–207.

Mueller, K. J., and Snoke, A. W., 1993, Progressive overprinting of normal fault systems and their role in Ter-tiary exhumation of the East Humboldt–Wood Hills metamorphic core complex, northeastern Nevada:Tectonics, v. 12, p. 361–371.

Nutt, C. J., and Thorman, C. H., 1992, Mesozoic to early Tertiary attenuation and Tertiary uplift on the GoshuteIndian Reservation, southwestern Deep Creek Range, Nevada-Utah: Geological Society of America Ab-stracts with Programs, v. 24, no. 6, p. 55.

Nutt, C. J., and Thorman, C. H., 1993, Eocene sedimentary and volcanic rocks and their use in dating Mesozoicand Tertiary structures in the southern Deep Creek Range, Nevada and Utah: Geological Society of Amer-ica Abstracts with Programs, v. 25, no. 5, p. 128.

Platt, J. P., 1986, Dynamics of orogenic wedges and uplift of high-pressure metamorphic rocks: Geological So-ciety of America Bulletin, v. 97, p. 1037–1053.

Platt, J. P., and Vissers, P. L., 1980, Extension structures in anisotropic rocks: Journal of Structural Geology, v. 2,p. 397–410.

Price, R., 1973, Large-scale gravitational flow of supracrustal rocks, southern Canadian Rockies, in de Jong,K. A., and Scholten, R., eds., Gravity and tectonics: New York, Wiley-Interscience, p. 491–502.

Rodgers, D. W., 1987, Thermal and structural evolution of the southern Deep Creek Range, west-central Utahand east-central Nevada [Ph.D. dissert.]: Stanford, California, Stanford University, 149 p.

Royse, F., 1993, Case of the phantom foredeep: Early Cretaceous in west-central Utah: Geology, v. 21,p. 133–136.

Rutter, E. H., and Rusbridge, M., 1977, The effect of non-coaxial strain paths on crystallographic preferred ori-entation development in the experimental deformation of marble: Tectonophysics, v. 39, p. 73–86.

Sabisky, M. A., 1985, Finite strain, ductile flow, and folding in the central Raft River Mountains, northwesternUtah [M.S. thesis]: Salt Lake City, University of Utah, 69 p.

Saltzer, S. D., and Hodges, K. V., 1988, The Middle Mountain shear zone, southern Idaho: Kinematic analysisof an early Tertiary high-temperature detachment: Geological Society of America Bulletin, v. 100,p. 96–103.

Schmid, S.M., and Casey, M., 1986, Complete fabric analysis of some commonly observed quartz c-axis pat-terns, in Hobbs, B. C., and Heard, H. C., eds., Mineral and rock deformation: Laboratory studies: Ameri-can Geophysical Union Geophysical Monograph 36, p. 263–296.

Schmid, S. M., Panozzo, R., and Bauer, S., 1987, Simple shear experiments on calcite rocks: Rheology and mi-crofabric: Journal of Structural Geology, v. 9, p. 747–778.

Serra, S., 1977, Styles of deformation in the ramp regions of overthrust faults, in Lawson, E. L., Norwood, E. R.,Wach, P. H., and Hale, L. A., eds., Rocky Mountain thrust belt geology and resources: Casper, Wyoming,Wyoming Geological Association Guidebook, 29th Field Conference, p. 487–498.

Smith, J. F., 1982, Geologic map of the Strevell 15 minute quadrangle, Cassia County, Idaho: U.S. GeologicalSurvey Bulletin 1536, 36 p.

Smith, R. B., and Sbar, M. L., 1974, Contemporary tectonics and seismicity of the western United States withemphasis on the intermountain seismic belt: Geological Society of America Bulletin, v. 85, p. 1205–1218.

Snoke, A. W., and Miller, D. M., 1988, Metamorphic and tectonic history of the northeastern Great Basin, inErnst, W. G., ed., Metamorphism and crustal evolution of the western United States: Englewood Cliffs,New Jersey, Prentice-Hall, Rubey Volume VII, p. 606–648.

Spear, F. S, and Cheney, J. T., 1989,A petrogenetic grid for peliteic schists in the system SiO2-A12O3-FeO-MgO-K2O-H2O: Contributions to Mineralogy and Petrology, v. 101, p. 149–164.

Spencer, J. E., 1982, Origin of folds of Tertiary low-angle fault surfaces, southeastern Californian and westernArizona, in Frost, E. G., and Martin, D. L., eds., Mesozoic-Cenozoic tectonic evolution of the ColoradoRiver region, California, Arizona, and Nevada: San Diego, California, Cordilleran Publishers, p. 123–134.

Taylor, W. J., Bartley, J. M., Fryxell, J. E., Schmitt, J. G., and Vandervoort, D. S., 1993, Tectonic style and re-gional relations of the Central Nevada thrust belt, in Lahren, M. M., Trexler, J. H., and Spinosa, C., eds.,Crustal evolution of the Great Basin and the Sierra Nevada: Reno, University of Nevada, Department ofGeological Sciences, Cordilleran/Rocky Mountain Section, Geological Society of America Guidebook,p. 57–96.

Thorman, C. H., and Snee, L. W., 1988, Thermochronology of metamorphic rocks in the Wood Hills and PequopMountains, northeastern Nevada: Geological Society of America Abstracts with Programs, v. 20, no. 7,p. A18.

Thorman, C. H., Ketner, K. B., and Peterson, F., 1990, The Elko orogeny—Late Jurassic orogenesis in theCordillean miogeocline: Geological Society of America Abstracts with Programs, v. 22, no. 3, p. 88.

Todd, V. R., 1980, Structure and petrology of a Tertiary gneiss complex in northwestern Utah, in Crittenden,M. D., Jr., Coney, P. J., and Davis, G. H., eds., Cordilleran metamorphic core complexes: Geological So-ciety of America Memoir 153, p. 349–383.

Todd, V. R., 1983, Late Miocene displacement of pre-Tertiary and Tertiary rocks in the Matlin Mountains, north-western Utah, in Miller, D. M., Todd, V. R., and Howard, K. A., eds., Tectonic and stratigraphic studies inthe eastern Great Basin: Geological Society of America Memoir 157, p. 239–270.

Tullis, J. A., Christie, J. M., and Griggs, D. T., 1973, Microstructures and preferred orientations of experimen-tally deformed quartzites: Geological Society of America Bulletin, v. 84, p. 297–314.

Turner, F. J., 1953, Nature and dynamic interpretation of deformation in calcite of three marbles:American Jour-nal of Science, v. 251, p. 276–298.

Turner, F. J., Griggs, D. T., Clark, R. H., and Dixon, R. H., 1956, Deformation of the Yule marble, part VII: Geo-logical Society of America Bulletin, v. 67, p. 1259–1294.

Urai, J. L., Means, W. D., and Lister, G. S., 1986, Dynamic recrystallization of mimerals, in Hobbs, B. C., andHeard, H. C., eds., Mineral and rock deformation: Laboratory studies: American Geophysical Union Geo-physical Monograph 36, p. 161–199.

Vandervoort, D., and Schmidt, J. G., 1990, Cretaceous to early Tertiary paleogeography in the hinterland of theSevier thrust belt, east-central Nevada: Geology, v. 18, p. 567–570.

Von Tish, D. B., Allmendinger, R. W., and Sharp, J. W., 1985, History of Cenozoic extension in central SevierDesert, west-central Utah, from COCORP seismic reflection data: American Association of PetroleumGeologists Bulletin, v. 69, p. 1077–1087.

Wagner, F., Wenk, H. R., Kern, H., Van Houtte, P., and Esling, C., 1982, Development of preferred orientationin plane strain deformed limestone: Experiment and theory: Contributions to Mineralogy and Petrology,v. 80, p. 132–139.

Wallis, S. R., Platt, J. P., and Knott, S. D., 1993, Recognition of syn-convergence extension in accretionarywedges with examples from the Calabrian Arc and the eastern Alps: American Journal of Science, v. 293,p. 463–495.

Wells, M. L., 1991, Episodic Mesozoic-early Cenozoic contraction and extension in the Cordilleran hinterland,Raft River Mountains, northwestern Utah: Geological Society of America Abstracts with Programs, v. 23,no. 5, p. A131.

Wells, M. L., 1992, Kinematics and timing of sequential deformations in the eastern Raft River Mountains, inWilson, J. R., ed., Field guide to geologic excursions in Utah and adjacent areas of Nevada, Idaho, andWyoming: Utah Geological Survey Miscellaneous Publication 92-3, p. 59–78.

Wells, M. L., 1996, Interim geologic map of the Kelton Pass 7.5′ quadrangle, Box Elder County, Utah, and Cas-sia Counry, Idaho: Utah Geological Survey Open-file Report, 67 p., scale 1:24 000.

Wells, M. L., and Allmendinger, R. W., 1990, An early history of pure shear in the upper plate of the Raft Rivermetamorphic core complex; Black Pine Mountains, southern Idaho: Journal of Structural Geology, v. 12,p. 851–868.

Wells, M. L., and Snee, L. W., 1993, Geologic and thermochronologic constraints on the initial orientation ofthe Raft River detachment and footwall shear zone: Geological Society of America Abstracts with Pro-grams, v. 25, no. 5, p. 161–162.

Wells, M. L., and Struthers, J. R., 1995, Influence of rheological constraints on mid-crustal extensional shearzone geometry: Geological Society of America Abstracts with Programs, v. 27, no. 6, p. A283.

Wells, M. L., Dallmeyer, R. D., and Allmendinger, R. W., 1990, Late Cretaceous extension in the hinterland ofthe Sevier thrust belt, northwestern Utah and southern Idaho: Geology, v. 18, p. 929–933.

Wells, M. L., Snee, L. W., and Blythe, A. E., 1994, Mesozoic and Cenozoic extensional history of the north-eastern Great Basin: Insights from the Raft River and Grouse Creek mountains: Geological Society ofAmerica Abstracts with Programs, v. 26, no. 2, p. 103.

Wells, M. L., Peters, M. T., and Miller, D. M., 1996, Neoproterozoic age for the lower part of the Raft River Se-quence: Implications for Mesozoic evolution of the Sevier Belt hinterland: Geological Society of Amer-ica Abstracts with Programs, v. 28, no. 5, p. 122.

Wenk, H. R., Takeshita, T., Bechler, E., Erskine, B. G., and Matthies, S., 1987, Pure and simple shear calcite tex-tures: Comparison of experimental, theoretical and natural data: Journal of Structural Geology, v. 9,p. 731–745.

White, S., 1973, Deformation lamellae in naturally deformed quartz: Nature, v. 245, p. 26–28.Wiltschko, D. V., and Dorr, J. A., 1983, Timing of deformation in overthrust belt and foreland of Idaho,

Wyoming, and Utah: American Association of Petroleum Geologists, v. 67, p. 1304–1322.Winkler, H., 1976, Petrogenesis of metamorphic rocks: Berlin, Springer-Verlag, 334 p.

MANUSCRIPT RECEIVED BY THE SOCIETY MAY 9, 1995REVISED MANUSCRIPT RECEIVED APRIL 6, 1996MANUSCRIPT ACCEPTED MAY 7, 1996

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