Shallow-Crustal Metamorphism During Late Cretaceous Anatexis in the Sevier Hinterland Plateau: Peak Temperature Conditions from the Grant Range, Eastern Nevada, U.S.A

Home , Anatexis, Grant Range, Pioche Shale

Shallow-crustal metamorphism during Late Cretaceous anatexis in the Sevier hinterland plateau: Peak temperature conditions from the Grant Range, eastern Nevada, U.S.A.

Sean P. Long1*, Emmanuel Soignard2 1SCHOOL OF THE ENVIRONMENT, WASHINGTON STATE UNIVERSITY, PULLMAN, WASHINGTON 99164, USA 2LEROY EYRING CENTER FOR SOLID STATE SCIENCE, ARIZONA STATE UNIVERSITY, TEMPE, ARIZONA 85287, USA

ABSTRACT

Documenting spatio-temporal relationships between the thermal and deformation histories of orogenic systems can elucidate their evolution. In the Sevier hinterland plateau in eastern Nevada, an episode of Late Cretaceous magmatism and metamorphism affected mid- and upper-crustal levels, concurrent with late-stage shortening in the Sevier thrust belt. Here, we present quantitative peak temperature data from the Grant Range, a site of localized, Late Cretaceous granitic magmatism and greenschist facies metamorphism. Twenty-two samples of Cambrian to Pennsylvanian metasedimentary and sedimentary rocks were analyzed, utilizing Raman spectroscopy on carbonaceous material, vitrinite reflectance, and Rock-Eval pyrolysis thermometry. A published reconstruction of Cenozoic extension indicates that the samples span pre-extensional depths of 2.5–9 km. Peak temperatures systematically increase with depth, from ~100 to 300 °C between 2.5 and 4.5 km, ~400 to 500 °C between 5 and 8 km, and ~550 °C at 9 km. The data define a metamorphic field gradient of ~60 °C/km, and are corroborated by quartz recrystallization microstructure and published conodont alteration indices. Metamorphism in the Grant Range is correlated with contemporary, upper-crustal metamorphism and magmatism documented farther east in Nevada, where metamorphic field gradients as high as ~50 °C/km are estimated. These data have implications for localized but significant thermal weakening of the plateau crust, including attaining temperatures for quartz plasticity at depths of ~5–6 km, and the potential for partial melting possibly as shallow as ~12–15 km. Thermal weakening may have contributed to a slowing of shortening rates documented in the Sevier thrust belt at this latitude at ca. 88 Ma, by locally inducing mid- and lower-crustal ductile thickening in the hinterland.

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INTRODUCTION necessary for quantitative thermobarometry. hyay, 1994). Semiquantitative peak temperature This, combined with a wealth of studies that parameters, such as conodont alteration index, Understanding the thermal history of an have focused on determining the pressure and only yield broad (~200 °C) bracketing ranges orogenic belt, and its relationship in space and temperature conditions of upper amphibolite for metasedimentary rocks that experienced time to the corresponding deformation history, facies rocks exhumed in the footwalls of core peak temperatures ≥300 °C (Königshof, 2003). can provide critical insights into the evolution complexes (e.g., McGrew et al., 2000; Cooper However, with the development of Raman spec- of orogenic systems. In the Jurassic-Paleogene et al., 2010; Wells et al., 2012; Hallett and Spear, troscopy on carbonaceous material (RSCM) U.S. Cordilleran orogenic belt, an episode of 2014), has led to a prevailing view that meta- thermometry (Beyssac et al., 2002), peak tem- Late Cretaceous (ca. 70–90 Ma) granitic mag- morphism in the Sevier hinterland dominantly peratures for greenschist facies metasedimen- matism and accompanying metamorphism has affected mid-crustal levels, despite earlier ef- tary rocks that contain organic matter can be been documented in the hinterland plateau of forts that presented field relations emphasizing determined within ±50 °C. This technique is the Sevier thrust belt in eastern Nevada and that high temperatures were locally attained in ideal for understanding the thermal conditions western Utah (e.g., Miller et al., 1988; Barton, the upper crust (Miller et al., 1988; Miller and of thick sedimentary sections such as the Cor- 1990), which produced greenschist and am- Gans, 1989). dilleran passive margin basin in eastern Nevada, phibolite facies metasedimentary rocks that are Part of the difficulty of determining the con- which is dominated by organic-rich carbonates now exposed within a series of highly extended ditions of shallow-crustal metamorphism is that, (e.g., Stewart, 1980). ranges (Fig. 1). However, the complex overprint until recently, quantitative techniques for ana- In this study, we present quantitative peak of Cenozoic extension across Nevada and Utah lyzing the peak temperature of sub-garnet grade temperature determinations from Paleozoic in many places hinders placing these rocks in rocks have not been available. Thermal matura- sedimentary and metasedimentary rocks in the an accurate pre-extensional structural and depth tion parameters utilized in the petroleum indus- Grant Range in eastern Nevada, which repre- context, in particular for greenschist facies rocks try, such as vitrinite reflectance, can estimate the sents the westernmost of a series of exposures that do not contain the mineral assemblages peak temperature of sedimentary rocks within of Late Cretaceous metamorphic rocks in the ±20–30 °C, but are applicable only up to peak Sevier hinterland in eastern Nevada (Figs. 1, 2). *[email protected] temperatures of ~200–300 °C (e.g., Mukhopad- The samples are placed in a detailed structural

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Figure 1. Map of Cordilleran OR ID Chapman et al., 2015), which supported surface 120°W 117°W 114°W 111°W retroarc region in Nevada elevations as high as ~3 km (Cassel et al., 2014; RMT (NV), Utah (UT), and Cali- Salt Snell et al., 2014). fornia (CA) (modified from GT Lake WY Elko City 41°N During Cordilleran orogenesis, two epi- Long, 2015). Deformation sodes of granitic magmatism, one in the Jurassic fronts of Cordilleran thrust Area of systems shown in blue, Fig. 7 (ca. 155–165 Ma) and one in the Late Creta- with spatial extents shaded ceous (ca. 70–90 Ma), have been documented Sierra Nevada Renomagmatic ar gray. Late Paleozoic thrust t in eastern Nevada and western Utah, and each is LFTB old- systems shown in brown. Ely associated spatially and temporally with meta- 39°N Areas of exposed Jurassic morphism and ductile, contractional deforma- and Cretaceous metamor- WUTB Sevier f thrust bel

phic rocks are highlighted, tion (e.g., Barton et al., 1988; Miller et al., 1988; CNTB including metamorphic N Miller and Hoisch, 1995; Wells and Hoisch, 2008). The Jurassic episode, which in most core complexes in red and Area of Cedar highly extended ranges Fig. 2A City 100 km places is characterized by spatially localized in orange. AZ—Arizona; c UT 37°N metamorphism in proximity to upper-crustal in- AZ CNTB—Central Nevada CA NV trusions (Miller et al., 1988; Miller and Hoisch, thrust belt; GT—Golconda 1995), has been interpreted as heating and thrust; ID—Idaho; LFTB—Luning-Fencemaker thrust belt; OR—Oregon; RMT—Roberts Mountains thrust; WUTB—Western Utah thrust belt; WY—Wyoming. low-magnitude contractional deformation that accompanied a pulse of back-arc magmatism (Miller and Hoisch, 1995). The more region- context, obtained from a recently published Cretaceous (ca. 130 Ma) foreland basin sub- ally significant Late Cretaceous magmatic event structural reconstruction of the Grant Range sidence indicates the onset of shortening (e.g., produced the peak metamorphic conditions re- (Long and Walker, 2015), which shows that Jordan, 1981; Currie, 2002), while some argue corded in much of eastern Nevada (Miller and they were distributed between crustal depths of for emplacement of thrust sheets as early as the Gans, 1989; Barton, 1990), and affected both ~2.5–9 km prior to Cenozoic extension. Three latest Jurassic (ca. 145 Ma) (e.g., DeCelles and mid- and upper-crustal levels. This event oc- different thermometry techniques are utilized, Coogan, 2006). The onset of crustal thickening curred late during the Sevier shortening history, including RSCM on Cambrian, Ordovician, in parts of the Sevier hinterland is interpreted to while the Nevadaplano was approaching its and Silurian rocks, and vitrinite reflectance and have begun as early as the Late Jurassic (ca. 150 maximum crustal thickness (e.g., DeCelles and Rock-Eval pyrolysis on Devonian, Mississip- Ma), on the basis of initial prograde metamor- Coogan, 2006), and was locally associated with pian, and Pennsylvanian rocks. The temperature phism recorded within rocks now exhumed in significant crustal thickening (Camilleri and data are used to quantify a metamorphic field the footwalls of metamorphic core complexes Chamberlain, 1997; McGrew et al., 2000). Late gradient, which is interpreted to reflect the peak (Hoisch et al., 2014; Cruz-Uribe et al., 2015; Cretaceous magmatism has been interpreted as thermal conditions for Late Cretaceous meta- Kelly et al., 2015) and isotopic compositions of the shallow thermal expression of lower-crustal morphism in the Grant Range. We then compare Late Jurassic plutons (Chapman et al., 2015). anataxis, which was initiated as a result of this data to other studied areas of Late Creta- In Nevada, low-magnitude (a few 10s of eastward migration of subduction-related mag- ceous, shallow-crustal metamorphism in eastern km) upper-crustal shortening of likely Creta- matism coupled with conductive relaxation of Nevada, and discuss implications for localized ceous age was accommodated in the Central isotherms within structurally thickened crust thermal weakening of the Sevier hinterland pla- Nevada thrust belt (Taylor et al., 2000; Long, (Miller and Gans, 1989), or alternatively, as a teau crust during the late stages of Cordilleran 2012; Long et al., 2014) and by regional-scale, result of heat influx following delamination of orogenesis. open folding across much of eastern Nevada dense mantle lithosphere (Wells and Hoisch, (Long, 2015). However, thermobarometry of 2008; Wells et al., 2012). CORDILLERAN TECTONIC FRAMEWORK exhumed mid-crustal rocks indicates localized During the Late Cretaceous and Paleocene OF EASTERN NEVADA but significant Cretaceous crustal thickening, (ca. 80–60 Ma), spatially isolated extension including within the Snake Range core complex is recorded in the Nevadaplano, during the fi- From the Neoproterozoic to the Triassic, in east-central Nevada (e.g., Lewis et al., 1999; nal stages of shortening in the Sevier thrust >10 km of shallow-marine sedimentary rocks Cooper et al., 2010), the Ruby–East Humboldt belt (Hodges and Walker, 1992; Camilleri and were deposited on the western Laurentian con- core complex in northeast Nevada (e.g., Hodges Chamberlain, 1997; Wells and Hoisch, 2008; tinental shelf in eastern Nevada (e.g., Stewart, and Walker, 1992; McGrew et al., 2000; Hal- Druschke et al., 2009; Long et al., 2015), and 1980). During the Jurassic, establishment of an lett and Spear, 2014), and in the footwall of the has been interpreted as a consequence of ther- Andean-style subduction system on the western Windermere thrust in northeast Nevada (Camil- mal weakening and isostatic adjustment fol- North American plate margin initiated construc- leri and Chamberlain, 1997). By the Late Cre- lowing lithospheric delamination (Wells and tion of the Cordilleran orogenic belt, which taceous, near the end of crustal thickening, the Hoisch, 2008). In addition, spatially isolated, continued until the Paleogene (e.g., DeCelles, Sevier hinterland is interpreted to have been a post-orogenic, Eocene-Oligocene extension 2004). For the duration of orogenesis, eastern high-elevation orogenic plateau (e.g., Coney has been documented in eastern Nevada (e.g., Nevada was situated within a broad hinterland and Harms, 1984; DeCelles, 2004), named the Gans et al., 2001; Druschke et al., 2009; Long region between the Sierra Nevada magmatic arc “Nevadaplano” after comparison to the Andean and Walker, 2015), and was often associated in and the Sevier fold-thrust belt, the locus of up- Altiplano. The Nevadaplano was underlain by space and time with the sweep of silicic volca- per-crustal shortening (Fig. 1). The timing of the crust that was locally as thick as ~50–60 km nism that accompanied post-Laramide slab roll- initiation of shortening in the Sevier thrust belt (Coney and Harms, 1984; Gans, 1987; DeCelles back (Humphreys, 1995; Dickinson, 2002). The is debated; most agree that the onset of Early and Coogan, 2006; Colgan and Henry, 2009; inception of widespread extension that formed

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White 115˚30’W 115˚22’30”W A Ely B 6 Pine Range

way 39°N

.S. High ange U

alley

V ange Currant

Egan R

ange

ancake R P alley V White River N Grant R km ailroad R area of 01020 116°W Quinn Fig. 2B 115°W Canyon R. 115˚37’30” N Scale (km) 051234

Beatty Canyon 38˚30’N area of Fig. 5A

outer limit A’ of positive areomagnetic A anomaly Heath Canyon

B B’

Valley Grant Canyon

Railroad

Ir Valley win

n Canyo

White River 38˚22’30”N

Quaternary sediment Troy Miocene-Pliocene volcanic Canyon and sedimentary rocks Eocene-Oligocene dikes Eocene-Oligocene volcanic Neogene normal and sedimentary rocks faults Jurassic-Cretaceous granite Paleogene-Neogene detachment faults Mississippian-Pennsylvanian Mesozoic thrust sedimentary rocks faults and folds Silurian-Devonian sedimentary rocks Drill holes that Cambrian-Ordovician intercept granite sedimentary rocks (hashes indicate metamorphism)

Figure 2. (A) Map showing location of the Grant Range and geographic names of surrounding ranges and valleys. (B) Generalized geologic map of the Grant Range, modified from Long and Walker (2015). Compiled source maps include Moores et al. (1968), Kleinhampl and Ziony (1985), Lund et al. (1987; 1988), Fryxell (1988), Camilleri (2013), and Long (2014). Dashed pink line shows outer limit of positive aeromagnetic anomaly interpreted to delineate subsurface extent of Troy pluton (Lund et al., 1987, 1988; Blank, 1993). Locations of Railroad Valley drill holes that intercept granite at depth are from Hess et al. (2004).

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the Basin and Range province, and accompa- white mica porphyroblasts during development thrust faults correlated with the Central Nevada nying lowering of the surface elevation of the of axial-planar cleavage. The second phase in- thrust belt deform Cambrian through Devo- Nevadaplano, was not until the middle Mio- volved growth of randomly oriented white mica, nian rocks (Taylor et al., 2000) (Fig. 2). How- cene (e.g., Dickinson, 2002; Colgan and Henry, phlogopite, biotite, and amphibole porphyrob- ever, in the central and northern Grant Range, 2009; Cassel et al., 2014). lasts, which are interpreted as static metamor- regionally continuous, older-over-younger rela- phic textures. On the basis of mineral assem- tionships indicative of large-scale thrust faults DEFORMATION, METAMORPHISM, AND blage, this stage is interpreted to have yielded have not been documented through the exposed MAGMATISM IN THE GRANT RANGE the peak metamorphic conditions recorded in thickness of Cambrian to Pennsylvanian rocks the Grant Range (Camilleri, 2013). This was (Lund et al. 1993; Camilleri, 2013; Long and An 8-km-thick section of Cambrian-Penn- followed by an additional stage of synkinematic Walker, 2015). sylvanian sedimentary rocks is exposed in the metamorphism that accompanied west-vergent, Paleozoic rocks in the Grant Range are un- Grant Range (Figs. 2, 3), and is dominated by mesoscale folding and thrust faulting, and is conformably overlain by Paleogene sedimen- carbonates (Fryxell, 1988; Lund et al., 1993; characterized by growth of oriented white mica tary and volcanic rocks. 1300 m of Paleogene Camilleri, 2013; Long and Walker, 2015). and chlorite during development of axial-planar rocks are preserved in the study area (Long and Cambrian and Ordovician rocks have experi- cleavage (Camilleri, 2013). Walker, 2015) (Fig. 3), but sections as thick enced greenschist facies metamorphism and In the southern Grant Range, 3–10 km south as ~2000 m are reported in other parts of the low-magnitude, penetrative contractional strain of the study area, Cambrian rocks record a his- range (Moores et al., 1968). In the study area, (Figs. 4B, 4D), both of which die out upsection, tory of metamorphism and large-scale folding, Paleogene rocks overlie Mississippian and and are not observed in Silurian-Pennsylvanian and are intruded by the Troy granite stock (Fryx- Pennsylvanian rock units, with minimal angular rocks (Fryxell, 1988; Lund et al., 1993; Camil- ell, 1988) (Fig. 2). U-Pb zircon geochronology discordance (Long and Walker, 2015). In ad- leri, 2013). Readers are referred to Camilleri indicates two distinct emplacement ages for the dition, Eocene-Oligocene andesite, dacite, and (2013) for detailed descriptions of micro- and stock (Lund et al., 2014). A boudinaged granite granite dikes intruded Cambrian and Ordovician mesoscale ductile structures and metamorphic sill on the western margin of the stock, which stratigraphic levels (Fig. 2) (Fryxell, 1988; Ca- lithologies, textures, and mineral assemblages pre-dates folding, yielded a Jurassic (163.3 ± milleri, 2013; Long and Walker, 2015). In the in the study area in the central Grant Range (Fig. 1.2 Ma) age, and undeformed granite that makes study area, a dacite dike yielded a crystallization 2). Rock unit divisions discussed below and up the bulk of the pluton, and post-dates folding, age of ca. 29 Ma (40Ar/39Ar biotite; Long and shown on Figure 3 are after Long and Walker yielded a Late Cretaceous age (83.7 ± 0.8 Ma) Walker, 2015), and in the southern Grant Range, (2015). Cambrian and Ordovician sandy and (Lund et al., 2014). A stage of static, peak meta- a granite dike is dated at 31.7 ± 0.8 Ma (U-Pb silty limestone protoliths have been metamor- morphism, indicated by randomly oriented por- zircon; Lund et al., 2014). phosed to recrystallized limestone with abun- phyroblasts, is recorded in the southern Grant Paleozoic and Paleogene rocks in the Grant dant mica porphyroblasts, with silty partings Range, and post-dates large-scale folding (Fryx- Range experienced a polyphase history of Ce- metamorphosed to argillite and phyllite (Figs. ell, 1988). The metamorphic grade of Cambrian nozoic extension. The earliest extension was 4A, 4C), and more pure limestone protoliths sedimentary rocks generally decreases upsec- accommodated by low dip-angle, top-to-west have been metamorphosed to crystalline marble tion in the southern Grant Range, but Fryxell detachment faults (Lund et al., 1993; Camil- (Figs. 4B, 4D). Moving downsection through (1988) also documented that metamorphism in leri, 2013; Long and Walker, 2015), which ini- the stratigraphic column (Fig. 3), sericite (de- Cambrian rocks dies out to the south, with dis- tiated in the Oligocene (ca. 29–32 Ma) (Long fined here as fine-grained white mica) and chlo- tance from the granite. This suggests that Late and Walker, 2015). This was followed by one or rite first appear within upper Ordovician rock Cretaceous magmatism may have been a pri- more episodes of extension accommodated by units, tourmaline and white mica porphyroblasts mary source of heat for metamorphism. high-angle normal faults, including Miocene to appear at the top of the Cambrian section, and After these observations in the southern part Quaternary faulting associated with formation phlogopite, amphibole (Fig. 4C), and biotite of the range, the stage of static, peak metamor- of the Railroad Valley structural basin (Fig. 2) porphyroblasts are observed within the lowest phism in the study area in the central Grant Range (Moores et al., 1968; Lund et al., 1993; Camil- Cambrian units in the study area (Camilleri, has also been interpreted to have been, at least leri, 2013). 2013). All metasedimentary rock units observed in part, contemporary with intrusion of the Late in the study area have limestone protoliths, and Cretaceous component of the Troy stock (Lund PEAK TEMPERATURE DATA therefore variation in protolith composition is et al., 1993; Camilleri, 2013). This interpretation not interpreted to have strongly affected the re- is supported by geophysical and drill hole data Structural, Stratigraphic, and Depth sulting mineral assemblages. that suggest that the Troy pluton extends in the Context of Samples In the southern Grant Range, stratigraphi- subsurface under the west-central Grant Range. cally lower Cambrian rocks are exposed, in- A positive aeromagnetic anomaly, interpreted to A suite of 22 samples of Paleozoic sedi- cluding the Pioche shale and Prospect Mountain delineate the subsurface extent of the Troy pluton mentary rocks ranging in age from Cambrian quartzite (Fig. 3). Both of these units exhibit (Lund et al., 1987, 1988; Blank, 1993), extends to Pennsylvanian were analyzed for the peak biotite porphyroblasts, and in one locality ~5 km through the western part of the study area (Fig. temperature that they experienced, using three south of the map area, the Pioche shale exhibits 2). Several drill holes in Railroad Valley that fall different thermometry techniques, RSCM, staurolite porpyhroblasts (Fryxell, 1988). within the area of the aeromagnetic anomaly in- vitrinite reflectance, and Rock-Eval pyroly- Three sequential phases of metamorphism tercept granite at depth (Fig. 2) (Lund et al., 1993; sis. All samples were collected in the central are recorded in Cambrian and Ordovician rocks Hess et al., 2004; Long and Walker, 2015), which Grant Range, within the map area of Long and in the study area (Camilleri, 2013). The first supports this interpretation. Walker (2015) (Figs. 2, 5A). Samples are given phase accompanied mesoscale, east-vergent In the southernmost Grant Range and farther structural context by projecting them to their folding, and is defined by growth of oriented south in the Quinn Canyon Range, east-vergent present-day positions on the deformed cross-

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Peak Temperature (˚C) 0˚ 100˚ 200˚ 300˚ 400˚ 500˚ 600˚ 700˚ 0 Explanation Permian rocks, Lithologies undierentiated Pu 1-1.5 CAI

rmian (Egan Range) (eroded) (n=19) quartzite Pe preserved thickness of shale Paleogene rocks in map area 1,000 limestone Pennsylvanian rocks, dolomite IPu y = 16.798x - 250 undierentiated 2 (Egan Range) (eroded) R = 0.8547 marble highest level of Paleogene unconformity in map area 2,000

ennsylvanian Samples

P Ely Limestone IPe GR09 1-2 CAI GR15

. Chainman Shale Mc GR08B (n=22) iss GR14 Best-t metamorphic eld M Joana Limestone GR04 Mj gradient 3,000 59.5°C / km Guilmette Formation Dg 2-3 CAI Mineral appearance isograds onian (n=10) Simonson Dolomite Doc

Dev Dsi Oxyoke Canyon Sandstone GR42 from Camilleri (2013) Dse GR37 Sevy Dolomite GR29 3-5 CAI Laketown Dolomite 4,000 Sl GR38 (n=12) Peak temperature data Oe GR36 projected paleo-surface Silurian Ely Springs Dolomite Oes Eureka Quartzite sericite-in level (15 ± 10 °C) Lehman Formation Kanosh Shale Opl Opk Shingle Limestone Ops GR35 Rock-Eval pyrolysis 5,000 (±0.2 VRo, n=3) Parker Spring Formation Opp GR34 chlorite-in

Ordovician vitrinite re ectance ogonip Group Goodwin Formation Stratigraphic depth (m) COpg P (2σ, n=6) Clm Little Meadows Fm. Raman spectroscopy on Blue Eagle member 6,000 GR50 white mica-in, carbonaceous material Csb GR67 tourmaline-in (2 SE, n=16) Grant Canyon member, Csgu upper unit GR69 subgrain rotation (SGR) quartz recrystallization 7,000 rmatio n (ca. 400-500 °C, n=2) Fo general range from Grant Canyon member, Csgl conodont alteration index lower unit (CAI) of rocks within 75 km GR68 8,000 radius (Craord, 2007) Sidehill Spring GR58 phlogopite-in, metamorphic temperature

Cambrian amphibole-in Willow Spring member Csw GR55 range for biotite-bearing lowest stratigraphic level GR53 units in southern Grant exposed in map area GR56 GR64 biotite-in Range (Fryxell, 1988) Pole Canyon Limestone 9,000 Cpc GR59 GR49 metamorphic temperature Pioche Shale Cp range for staurolite-bearing Cp in southern Grant Range Prospect Mountain Cpm Quartzite (Fryxell, 1988) base not exposed 10,000

Figure 3. Stratigraphic column of Grant Range rock units, using divisions and thicknesses from Long and Walker (2015) (lithologies and estimated eroded thicknesses of Pennsylvanian and Permian rocks in the Egan Range are from Kellogg, 1963, and Brokaw and Heidrick, 1966). Column on right shows peak temperature versus stratigraphic depth below projected paleo-surface level for the 22 analyzed samples.

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A B

cleavage

bedding

C D

white mica West East

amphibole

E F calcite

calcite

SGR

quartz CM

1 mm 0.25 mm

G H I calcite white mica CM CM calcite CM

calcite 100 µm 100 µm 100 µm

Figure 4. Photographs and photomicrographs, illustrating: (A) Thin-bedded, organic-rich limestone with phyllitic partings, characteristic of unit _sgl (refer to Fig. 3 for a guide to all unit abbreviations used in this caption). (B) Marble characteristic of unit _sw, with bedding (subhorizontal) overprinted by spaced cleavage (dipping toward left-hand side). (C) Amphibole (black) and white mica (silver) porphyroblasts within a phyllitic parting in marble of unit _sw. (D) West-vergent, mesoscale fold in phyllitic marble of unit _pc. (E) Black microlithons of carbonaceous material (CM) in marble of unit _pc (sample GR64; plane-polarized light). (F) Equigranular, polygonal quartz subgrain microstructure characteristic of subgrain rotation (SGR) recrystallization (sample GR67; cross-polarized light). (G, H, I) Representative examples of analyzed CM (G: sample GR50, spot 9; H: sample GR36, spot 3; I: sample GR55, spot 6; all taken in plane-polarized light; green spot in each photo is the Raman laser beam).

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A GR04 GR08B GR59 1 km N A’ GR09

A GR64 GR14

GR58 GR42 38°27.5’N GR55 GR56 GR67 GR49 GR29 GR50 GR36 GR53 GR34 GR35 B’ B GR69 GR38 GR68 115°30’W GR37 115°25’W 115°20’W Explanation Pennsylvanian sedimentary rocks Mississippian sedimentary rocks Neogene high-angle normal fault Quaternary sediment Devonian sedimentary rocks unconformity Paleogene-Neogene Silurian sedimentary rocks Neogene valley ll (A-A’ only) low-angle detachment fault unconformity Ordovician metasedimentary rocks bedrock culmination axis Cambrian metasedimentary rocks (units Csgl to Clm) Paleogene volcanic rocks Paleogene dikes Cambrian metasedimentary rocks (units Cpm to Csw) GR38 peak temperature sample unconformity Neoproterozoic metasedimentary rocks (B-B’ only)

W No vertical exaggeration B A A’ W 1 km W 3,000 3,000 115°25’ GR42 ? GR08B GR50 GR14 GR56 GR09 GR64 GR49 GR59 GR55 115°20’ GR58 115°30’ GR04 2,000 2,000 Elevation (m) Elevation (m) 1,000 1,000 0 0 No vertical exaggeration W C B W B’ 1 km W GR29 GR37 GR36 GR38 GR35 GR34 3,000 3,000 GR67 115°25’ ) 115°30’ ) GR53 GR69 115°20’ GR68 2,000 2,000 Elevation (m Elevation (m 1,000 1,000 0 0 02143 5 No vertical exaggeration D Scale (km) A’ 10 GR14

GR04 9 8 GR42

GR09 7 m) GR08B 6 GR50

5 ess (k 4 GR58 3 thic kn GR49 GR59 A GR56 2 1 GR55 GR64 0 02143 5 No vertical exaggeration 10 E Scale (km) B’ 9 8 GR38 GR37 GR36 GR29 7 m) GR35 GR34 6 GR67

5 ess (k GR69 4 GR68

3 thic kn GR53 2 B 1 0

Figure 5. (A) Geologic map of the central Grant Range, generalized from Long and Walker (2015). (B, C) Cross-sections A–A′ and B–B′, generalized from Long and Walker (2015), showing present-day geometry. Translucent areas above modern erosion surface represent eroded rock. Peak temperature samples are projected along-strike to their sampled structural level. (D, E) Cross sections A–A′ and B–B′, restored for motion on Cenozoic normal and detachment faults and for the magnitude of flexural isostatic folding that accompanied extension; generalized from Long and Walker (2015). Translucent areas represent rock that has either been eroded above the modern erosion surface or translated westward off of the map area during extension. The restored positions of peak temperature samples are projected.

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sections of Long and Walker (2015) (Figs. 5B, The samples span a total stratigraphic thick- on Figure 3 and Tables 1–3 at their stratigraphic 5C). In addition, Long and Walker (2015) re- ness of 6.5 km (Fig. 3). However, assigning pre- depth below 2.5 km of eroded Pennsylvanian stored Cenozoic extension in the Grant Range extensional stratigraphic depths is more difficult, and Permian rocks, after thicknesses reported by retro-deforming offset on all normal and as an unknown thickness of rocks was eroded in the Egan Range (Kellogg, 1963; Brokaw and detachment faults, and presented observations off of the study area prior to the Paleogene (e.g., Heidrick, 1966). Based on the regional obser- indicating that Paleozoic rocks had gentle dip Long, 2012). In the central Grant Range, only a vations summarized above, these stratigraphic magnitudes prior to extension, including docu- few hundred meters of Pennsylvanian rocks are depths are interpreted to represent maximum menting minimal angular discordance and preserved beneath the Paleogene unconformity permissible estimates for burial depths of rocks structural relief across the Paleogene unconfor- (Fig. 3). However, incomplete and complexly in the study area prior to synorogenic erosion. mity, and retro-deformation of the magnitude faulted sections of Pennsylvanian and Permian of flexural isostatic folding that accompanied rocks are preserved along-strike to the north in RSCM Thermometry extension. The samples are projected to their the White Pine Range (Moores et al., 1968), pre-extensional positions on the restored cross- which suggests that stratigraphic levels at least Carbonaceous material (CM), which is de- sections of Long and Walker (2015) (Figs. 5D, as high as Permian were at one time present over rived from solid-state metamorphism of organic 5E), which reveals that they were spread over the study area. To the east in the Egan Range, material (e.g., Buseck and Huang, 1985), is a total east-west distance of 30–40 km prior to Kellogg (1963) and Brokaw and Heidrick common in many metasedimentary rocks. Sev- extension. (1966) documented a total thickness of 2.0–2.5 eral studies have shown that the degree of struc- Based on the reconstruction of Long and km of Pennsylvanian and Permian rocks. Con- tural organization of graphite bonds in CM is Walker (2015), we assume that stratigraphic odont alteration indices from Pennsylvanian and strongly temperature dependent, and therefore depths for the samples are approximately equiva- Permian rocks within a ~75 km radius of the can be used as a quantitative geothermometer lent to pre-extensional structural depths. Samples study area are characterized by values of 1–1.5 (e.g., Beyssac et al., 2002; Rahl et al., 2005; are shown in a stratigraphic column (Fig. 3) using (n = 19) (Crafford, 2007), which corresponds to Aoya et al., 2010). Rahl et al. (2005) calibrated unit thicknesses from Long and Walker (2015), a maximum burial temperature range of 50–80 the RSCM thermometer with an uncertainty of which are based on a combination of complete °C (Königshof, 2003). These data indicate that ± 50 °C (2s) for rocks that achieved peak tem- unit thicknesses measured in their cross sections, Pennsylvanian and Permian rocks in the study peratures between 100 and 700 °C, by measur- complete thicknesses reported in other published area were not deeply buried by a thick Mesozoic ing the height ratio (R1) and area ratio (R2) of studies in the Grant Range (Moores et al., 1968; section that is now eroded away (e.g., Long, four first-order Raman peaks (G, D1, D2, D3) in Fryxell, 1988; Camilleri, 2013), and minimum 2012). On the basis of westward onlap of Tri- the wavenumber offset range between 1200 and tectonic thicknesses for some Cambrian and assic rocks and westward erosional truncation 1800 cm–1. Here, R1, R2, and peak temperature Ordovician rock units. Despite using minimum of Permian rocks below Triassic rocks, several are determined from Equations 1, 2, and 3, re- thicknesses, the cumulative thicknesses of the studies have argued that east-central Nevada (in- spectively, of Rahl et al. (2005). Cambrian and Ordovician sections in the study cluding the study area) was a topographic high CM was analyzed from 16 samples of Cam- area are comparable to nearby estimates in the during much of the Triassic (Burchfiel et al., brian, Ordovician, and Silurian rocks, with li- White Pine and Egan Ranges (Kellogg, 1963; 1974; Collinson et al., 1976; Stewart, 1980), and thologies including marble, phyllite, limestone, Moores et al., 1968), and regionally in eastern therefore did not accumulate a thick section of and dolomite (Table 1; Fig. 3). Most samples Nevada (Stewart, 1980). Triassic rocks. Therefore, samples are reported contained abundant CM, which was either pres-

TABLE 1. PEAK TEMPERATURE DETERMINATIONS FROM RAMAN SPECTROSCOPY ON CARBONACEOUS MATERIAL Sample Map Lithology Stratigraphic Latitude Longitude R1 R2 Peak temperature (°C) n unit depth (m) (dd.ddddd) (dd.ddddd) Mean 1σ Mean 1σ Mean 1σ 2 SE GR29 Sl dolomite 4075 38.44536 115.42425 1.4660.224 0.7210.043 261423613 GR38 Sl dolomite 4090 38.44267 115.42975 1.6570.112 0.7704 0.026225 26 30 14 GR36 Sl dolomite 4210 38.44267 115.42975 1.9840.485 0.7660.071 221805014 GR35 Ops limestone 5095 38.44250 115.43325 0.2450.028 0.3490.045 438443713 GR34 Opp limestone 5235 38.44367 115.43739 0.2610.161 0.3360.100 455674812 GR50 _sb limestone 6150 38.44667 115.45867 0.1900.098 0.3350.111 437956810 GR67 _sgu limestone 6325 38.44711115.48631 0.3090.083 0.3480.057 457433415 GR69 _sgu limestone 6855 38.43297 115.51264 0.4630.131 0.4140.070 425463614 GR68 _sgl limestone 8090 38.42447 115.529110.303 0.0650.330 0.051475 42 36 13 GR58 _sw marble 8470 38.45214 115.48531 0.2850.123 0.3380.087 461624413 GR55 _sw phyllite 8820 38.45019 115.47253 0.1390.055 0.2210.079 544695410 GR53 _sw marble 8910 38.43944 115.47581 0.1170.043 0.1760.054 586464111 GR56 _pc marble 9005 38.45050 115.45886 0.1610.101 0.2240.100 547845115 GR64 _pc marble 9005 38.45531 115.46428 0.2050.045 0.2830.061 498534013 GR59 _pc marble 9050 38.46950 115.47017 0.1760.046 0.2700.043 503353214 GR49 _pc marble 9080 38.44903 115.46467 0.2110.064 0.2570.067 527574114 Note: (1) R1, R2, and peak temperature values calculated using the calibration of Rahl et al. (2005). Internal variability in R1, R2, and peak temperature is indicated by 1σ uncertainty. Temperature is also reported with 2 standard errors (SE), calculated after Cooper et al. (2013), from quadratic addition of 1σ internal error and external error of ±50 °C from the Rahl et al. (2005) calibration, divided by the square root of the number of analyses (n). (2) Stratigraphic depths calculated below estimated paleo-surface level at 2.5 km above base of unit IPe, from thicknesses of Pennsylvanian and Permian rocks in the Egan Range (Kellogg, 1963; Brokaw and Heidrick, 1966).

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ent within organic-rich microlithons (Figs. 4E, from Cambrian units _pc and _sw (see Figure GR29, spot 10 R1=1.653 4I), or as isolated patches, typically ≤50 µm in 3 for a guide to all unit abbreviations used in the T=246°C R2=0.751 diameter (Fig. 4G, 4H). CM was analyzed in situ text), range between ~500 and 580 °C. on polished petrographic thin sections that were GR38, spot 14 R1=1.407 cut normal to bedding. Measurements were Vitrinite Reflectance Thermometry T=224°C R2=0.755 made at the LeRoy Eyring Center for Solid State Science at Arizona State University, using a Vitrinite reflectance, a thermal maturation pa- GR36, spot 7 R1=1.678 custom-built Raman spectrometer. The 532 nm rameter widely applied in the petroleum industry, T=234°C R2=0.764 laser was operated at a power of 3 mW, and was can be used to estimate the peak temperature of focused using a 50× ultra-long working distance sedimentary rocks over the range of ~50–300 °C Mitutoyo objective. The probed area of CM for (e.g., Barker and Pawlewicz, 1994; Mukhopad- GR35, spot 1 R1=0.229 T=443°C R2=0.341 each measurement was ~1 µm in diameter (Figs. hyay, 1994). Random vitrinite reflectance, the 4G, 4I). Instrument parameters, settings, and proportion of normal incident light reflected by a procedures follow those outlined in Cooper et polished surface of vitrinite, increases systemati- GR34, spot 19 R1=0.180 al. (2013). Where possible, the laser was focused cally with peak temperature, as the result of a se- T=456°C R2=0.315 on CM situated beneath a transparent grain (typ- ries of chemical transformations that accompany GR50, spot 7 R1=0.385 ically calcite), after procedures outlined in Beys- hydrocarbon maturation (Mukhopadhyay, 1994). T=355°C R2=0.463 sac et al. (2003). CM was analyzed for 120 sec- Six samples of limestone, dolomite, and –1 onds over a spectral window of 1100–2000 cm , shale from Devonian, Mississippian, and Penn- GR67, spot 7 R1=0.388 and typically 15 separate spots were analyzed in sylvanian rocks were analyzed for vitrinite re- T=453°C R2=0.372 each sample, to allow evaluation of in-sample flectance thermometry (Table 2). Analyses were variation. Peak positions, heights, widths, and performed by Weatherford Laboratories, Inc., GR69, spot 17 R1=0.491 areas of the Raman spectra (see supplementary using procedures outlined in ASTM (2014). No T=424°C R2=0.423 information for supporting data from individual primary vitrinite fragments were identified in analyses1) were determined using a custom peak the samples; instead, measurements of random GR68, spot 10 R1=0.285 fitting program written in Matlab by E. Soignard. reflectance were made on grains of solid bitu- T=473°C R2=0.327 The program allowed peak shapes to be fit by a men, a vitrinite-like maceral (e.g., Landis and combination of gaussian and lorentzian peaks, Castano, 1995). As few as two and as many as and any background slope to be removed by us- 30 measurements of solid bitumen reflectance GR58, spot 12 R1=0.212 T=486°C R2=0.296 ing a first-order polynomial between 1100 and (RSB) were made from individual samples (sup- 2000 cm–1. Examples of representative Raman porting data in supplementary information; see GR55, spot 13 R1=0.119 spectra for each sample are shown in Figure 6, footnote 1). RSB values were converted into T=544°C R2=0.216 and summary data for peak temperature deter- equivalent vitrinite reflectance (RVE) values minations are shown in Table 1. Mean peak tem- using the equation of Jacob (1989) (Table 2). peratures of multiple measurements are reported Peak temperatures were then obtained from GR53, spot 7 R1=0.148 T=589°C R2=0.182 on Table 1, and internal uncertainty within each the mean RVE value, using the calibration equa- sample is represented by the reported 1s error in tion for burial heating from Barker and Pawle- R1, R2, and peak temperature. However, after wicz (1994), as samples were not collected in GR56, spot 8 R1=0.177 T=537°C R2=0.239 Cooper et al. (2013), peak temperatures are re- stratigraphic continuity to allow proper thermal ported with 2 standard errors (SE), which takes modeling. Error is reported at the 2s level, into account the external error of ± 50 °C from which corresponds to a typical error range of GR64, spot 4 R1=0.222 T=502°C R2=0.284 the Rahl et al. (2005) calibration (see footnote ±10–25 °C. Peak temperatures determined of Table 1). At 2 SE, typical error ranges are ± from the six samples range between ~130 and 30–50 °C (Table 1). 150 °C (Table 2; Fig. 3), and temperatures from GR59, spot 14 R1=0.136 Resulting RSCM peak temperatures exhibit all samples overlap within estimated error. T=506°C R2=0.256 a total range between ~260 and 580 °C (Table GR49, spot 10 G R1=0.198 1; Fig. 3), and temperatures generally increase Rock-Eval Pyrolysis Thermometry T=539°C R2=0.242 with stratigraphic depth. The three Silurian sam- D1 D2

ples (GR29, GR38, GR36) yielded temperatures Rock-Eval pyrolysis, another technique 1200 1300 1400 1500 1600 1700 1800 between ~220 and 260 °C. Seven samples from widely applied in the petroleum industry to es- Raman Shift (cm-1) the Ordovician section and the upper and middle timate the thermal maturity of organic matter, Figure 6. Examples of representative Raman parts of the Cambrian section (GR35, GR34, can be used to estimate the peak temperature of spectra from each of the 16 samples analyzed GR50, GR67, GR69, GR68, GR58) range be- sedimentary rocks over the range of ~50–200 °C for Raman spectroscopy on carbonaceous mate- tween ~420 and 480 °C. The lowest six samples (e.g., Clementz, 1979; Peters, 1986). Pyrolysis rial. The positions of the graphite band (G) and (GR55, GR53, GR56, GR64, GR59, GR49), involves heating of a pulverized rock sample in defect bands (D1, D2) are shown on the bottom the absence of oxygen, in order to release and spectrum. Samples are stacked in stratigraphic measure organic compounds (Peters, 1986). order. Peak temperatures (T) and R1 and R2 1 parameters are calculated after Rahl et al. (2005). GSA Data Repository Item 2016051, Tables DR1– T , the oven temperature at which the maxi- DR3, Figure DR1 and related supporting information, max Peak center position, height, amplitude, and is available at www.geosociety.org/pubs/ft2016.htm, mum amount of non-volatile hydrocarbons are area are listed for individual analyses in the sup- or on request from [email protected]. released, can be converted into a calculated vi- plementary information (see footnote 1).

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TABLE 2. PEAK TEMPERATURE DETERMINATIONS FROM VITRINITE REFLECTANCE

Sample Map Lithology Stratigraphic Latitude Longitude Mean RSB RSB Mean RVE RVE nTpeak ± 2σ error unit depth (dd.ddddd) (dd.ddddd) (%) (1σ) (%) (1σ) (°C) (°C) (m) GR09 IPe limestone 2435 38.46167 115.37467 1.14 0.2531.110.146 8144 +19/–25 GR08BMc siltstone 2695 38.46536 115.38319 1.17 0.3681.120.161 2145 +20/–27 GR14 Mc shale 2845 38.45800 115.40414 0.95 0.1500.990.090 16 135+13/–16 GR04 Mc limestone 2860 38.47006 115.51256 1.26 0.1361.180.083 30 149+11/–12 GR42 Dsi dolomite 3875 38.45653 115.43808 0.91 0.1170.960.071 23 132+11/–13 GR37 Dse dolomite 3975 38.42458 115.42819 0.91 0.1960.96 0.112 7132 +21/–17

Note: (1) RSB was converted into RVE using the equation of Jacob (1989): RVE = 0.618*RSB + 0.40. (2) Tpeak calculated from RVE values, using equation for burial heating from Barker and Pawlewicz (1994): Tpeak = (ln(RVE) + 1.68)/0.0124. (3) Stratigraphic depths calculated below estimated paleo- surface level at 2.5 km above base of unit IPe, from thicknesses of Pennsylvanian and Permian rocks in the Egan Range (Kellogg, 1963; Brokaw and

Heidrick, 1966). Abbreviations: RSB = solid bitumen reﬂectance, RVE = equivalent vitrinite reﬂectance (from Jacob, 1989), Tpeak = peak temperature.

trinite reflectance value (cal. VR ) (Jarvie et al., Additional Peak Temperature Constraints the upper 6 km of the Paleozoic section (Fig. 3). 2001), and therefore can be used to estimate Pennsylvanian-Permian rocks are characterized peak temperature. Three additional data sets, including quartz by CAI values of 1–1.5 (<50–80 °C; all CAI tem- Three samples of limestone, siltstone, and recrystallization microstructure, published con- perature ranges listed here are from Königshof, shale from Mississippian and Pennsylvanian odont alteration indices, and published meta- 2003), Mississippian rocks 1–2 (60–140 °C), De- rock units were analyzed for Rock-Eval py- morphic temperature ranges based on mineral vonian rocks 2–3 (110–200 °C), and Ordovician- rolysis (Table 3). Analyses were performed by assemblage, supply semiquantitative peak tem- Silurian rocks 3–5 (190–480 °C). The CAI tem- Weatherford Laboratories, Inc., using instrument perature constraints for Paleozoic rocks in the perature ranges are generally in agreement with settings and procedures outlined in Clementz Grant Range. the vitrinite reflectance, Rock-Eval pyrolysis, and (1979) and Peters (1986). Additional supporting Characterization of the morphology of dy- RSCM peak temperature determinations (Fig. 3). data are included in the supplementary informa- namic recrystallization of quartz in thin section In the southern Grant Range, Fryxell (1988)

tion (see footnote 1). Tmax was converted into cal. allows bracketing of deformation temperature documented that lower Cambrian rock units, in-

RV using the equation of Jarvie et al. (2001), and range (e.g., Stipp et al., 2002). In the study area, cluding units _pc, _p, and _pm, exhibit biotite

peak temperature was calculated from cal. RV two samples of sandy limestone from map units porphyroblasts, and in one locality phyllite from using the calibration of Barker and Pawlewicz _sgu (GR67) and _sb (GR50) display evidence unit _p exhibits staurolite porphyroblasts. Us- (1994). Peak temperatures for the three samples for subgrain rotation recrystallization (Fig. 4F), ing mineral stability fields presented in Fryxell range between ~100 and 125 °C (Table 3). In which is characterized by ~0.1 mm, equigranu- (1988), over a pressure range of 2–3 kbar, a peak the absence of a statistically meaningful proce- lar, polygonal quartz subgrains, and indicates de- temperature range of ~480–580 °C is estimated dure for calculating uncertainty, errors of ±0.2 formation temperatures of ~400–500 °C (Stipp for biotite-bearing rocks, and a range of ~580–

cal. RV were assigned, which results in a typical et al., 2002) (Fig. 3). This temperature range is in 630 °C is estimated for the staurolite-bearing temperature error range of ±15–30 °C, which is agreement with RSCM peak temperatures from sample. The ~480–580 °C range is in agreement comparable to the ± 2s error range for the vitrin- these samples (457 ± 34 °C, 437 ± 68 °C). with the ~500–550 °C range of RSCM tempera- ite reflectance analyses. Peak temperatures for Conodont alteration indices (CAI) provide tures from the four biotite-bearing samples in the all three pyrolysis samples overlap within this semiquantitative estimates of the peak tempera- study area (Fig. 3). estimated error range (Table 3). ture that sedimentary rocks have experienced Peak temperature estimates for these three during diagenetic burial or metamorphism (e.g., DISCUSSION samples were obtained from both vitrinite re- Epstein et al., 1977; Königshof, 2003). Compila- flectance and pyrolysis. Pyrolysis temperature tion of characteristic CAI values for Ordovician Calculation and Interpretation of estimates were consistently lower, by a margin through Pennsylvanian sedimentary rocks within Metamorphic Field Gradient of ~20–45 °C; however, temperatures from both a 75 km radius of the study area, using the da- techniques overlap within estimated error for all tabase of Crafford (2007), gives a general range The peak temperature samples span a 6.5 three samples. of peak temperature versus stratigraphic level for km stratigraphic thickness, and therefore can be

TABLE 3. PEAK TEMPERATURE DETERMINATIONS FROM ROCK-EVAL PYROLYSIS

Sample Map unit Lithology Stratigraphic Latitude Longitude Tmax cal. RV Tpeak ±0.2 cal. RV depth (dd.ddddd) (dd.ddddd) (°C) (%) (°C) error (m) (°C) GR09 IPe limestone 2435 38.46167 115.37467 433 0.63 99 +22/–31 GR08B Mc siltstone 2695 38.46536 115.38319 447 0.89 126 +16/–21 GR14 Mc shale 2845 38.45800 115.40414 440 0.76 113+19/–25

Note: (1) Tmax = oven temperature of maximum release of non-volatile hydrocarbons. (2) cal. RV = calculated vitrinite reﬂec- tance percentage (from Jarvie et al., 2001, conversion). (3) Tpeak = peak temperature rock has experienced (from Barker and Pawlewicz, 1994, calibration). (4) Stratigraphic depths calculated below estimated paleo-surface level at 2.5 km above base of unit IPe, from thicknesses of Pennsylvanian and Permian rocks in the Egan Range (Kellogg, 1963; Brokaw and Heidrick, 1966).

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used to estimate a metamorphic field gradient. field gradient will be the most representative es- magmatism, and partial melting, at peak pres- However, interpretation of the significance of timate available for the peak thermal conditions sure and temperature conditions of 6–10 kbar this field gradient involves several assumptions attained in the upper crust at this locality. and 600–800 °C (e.g., Hodges and Walker, and caveats: A best-fit line constrained by the surface tem- 1992; McGrew et al., 2000; Hallett and Spear, 1. Three independent lines of evidence in- perature datapoint and the three peak tempera- 2014). To the east, Paleozoic metasedimentary dicate that Paleozoic rocks in the study area ture data sets yields a metamorphic field gradi- rocks in the Wood Hills and Pequop Mountains were not deeply buried beyond observed re- ent of ~60 °C/km (Fig. 3). Using this gradient, (Fig. 7) are interpreted to have experienced regional stratigraphic depths. These include: (a) approximate temperatures for the appearance of gional peak metamorphism at ca. 84 Ma, as a regional preserved sections of Pennsylvanian metamorphic minerals can be estimated, with result of structural burial to ~4–6 kbar under the and Permian rocks up to 2.5 km thick (Kellogg, sericite (fine-grained white mica) appearing at Windermere thrust sheet (Camilleri and Cham- 1963; Brokaw and Heidrick, 1966), which con- ~275 °C, chlorite at ~300 °C, white mica and berlain, 1997). Amphibolite facies metasedi- sistently yield low CAI values (Crafford, 2007), tourmaline porphyroblasts at ~375 °C, phlogo- mentary rocks exhumed in the footwall of the and multiple studies that have argued that this pite and amphibole porphyroblasts at ~500 °C, Northern Snake Range décollement (Fig. 7) region of Nevada did not accumulate a thick and biotite porphyroblasts at ~550 °C. However, were structurally buried to ~6–8 kbar, and ex- section of Triassic rocks (Burchfiel et al., 1974; the data are not perfectly linear, and between perienced peak temperatures of ~500–650 °C, Collinson et al., 1976; Stewart, 1980); (b) a lack depths of 3 and 5 km, several datapoints lie be- prior to exhumation (Lewis et al., 1999; Cooper of evidence for large-scale thrust faults and folds low (GR42, GR37 from Devonian rocks) and et al., 2010). in the central and northern Grant Range (Lund above (GR35, GR34 from Ordovician rocks) In contrast, in the ranges surrounding the et al. 1993; Long and Walker, 2015), which the best-fit line. This could be a consequence of Northern Snake Range, Late Cretaceous peak precludes significant structural burial; and (c) heterogeneous heat distribution either vertically metamorphism is also recorded within metased- retro-deformation of Cenozoic extension in the or laterally through the crust. On the basis of the imentary rocks that, prior to Cenozoic exten- study area, which indicates that Paleozoic rocks spatial proximity of metamorphic rocks to the sion, restore to upper-crustal levels (Miller et al., had low pre-extensional dip angles (Long and exposed and inferred subsurface extent of the 1988; Miller and Gans, 1989). The Deep Creek Walker, 2015). Therefore, pre-extensional strati- Troy stock (Fig. 2) (Fryxell, 1988; Lund et al., Range, Kern Mountains, Schell Creek Range, graphic depths are interpreted to approximate 1987, 1988, 1993), and the interpretation that and Egan Range (Fig. 7) exhibit conformable structural depths during peak metamorphism, peak metamorphism was contemporary with stratigraphic successions that span the Neopro- indicating that the samples spanned crustal intrusion of the Late Cretaceous component terozoic to the Triassic, encompassing the upper depths of 2.5–9 km. The surface level during of the stock (Fryxell, 1988; Lund et al., 1993; 12–14 km of the crust (Stewart, 1980; Miller peak metamorphism is interpreted at the top of Camilleri, 2013), the elevated temperatures are and Gans, 1989). The deeper stratigraphic levels the Permian section, and is assigned a tempera- interpreted as the result of addition of magmatic of these ranges, which consist of Neoprotero- ture range of 15 ± 10 °C (Fig. 3). heat. Therefore, proximity to subsurface intru- zoic to Cambrian rocks that restore to pre-ex- 2. All samples are assumed to have achieved sive bodies is one factor that may have yielded tensional depths of 8–14 km, exhibit greenschist peak temperature conditions at approximately nonlinearity in temperature versus depth. (chlorite- and biotite-bearing) to amphibolite fa- the same time, during the pulse of static, peak cies (staurolite-, garnet-, and locally andalusite- metamorphism recorded in Cambrian and The Grant Range in the Context of Late and sillimanite-bearing) mineral assemblages, Ordovician rocks in the central and southern Cretaceous Metamorphism in the Sevier which encompass an approximate total peak Grant Range (Fryxell, 1988; Camilleri, 2013). Hinterland temperature range of 300–650 °C (Miller et al., Fryxell (1988) and Camilleri (2013) both in- 1988). In general, metamorphic grade increases terpreted that peak metamorphism was, at least Late Cretaceous (ca. 70–90 Ma) granitic with stratigraphic depth, and peak metamorphic in part, synchronous with intrusion of the Late magmatism, metamorphism, and east-vergent, field gradients, estimated from mineral stability Cretaceous (ca. 84 Ma) component of the Troy ductile contractional deformation have been fields and CAI data, are in many places as high granite stock. This interpretation is supported documented in several areas in eastern Nevada, as ~50 °C/km (Miller and Gans, 1989). Upper by peak Late Cretaceous (ca. 70–90 Ma) meta- including the Snake Range metamorphic core crustal metamorphic rocks in these ranges are morphism documented across much of eastern complex and surrounding ranges (e.g., Miller observed over a modern (post-extensional) map Nevada (Miller and Gans, 1989). and Gans, 1989; Lewis et al., 1999; Cooper et area of ~135 km (N-S) by ~100 km (E-W) (Fig. 3. When restored for Cenozoic extension, al., 2010), and within and east of the Ruby–East 7), which indicates a broad spatial distribution the positions of the peak temperature samples Humboldt metamorphic core complex (e.g., of elevated upper-crustal temperatures. How- are spread over a total east-west distance of 30– Hodges and Walker, 1992; Camilleri and Cham- ever, abrupt increases in metamorphic grade are 40 km (Figs. 5D, 5E), and therefore do not rep- berlain, 1997; McGrew et al., 2000) (Fig. 7). also observed in proximity to Late Cretaceous resent a vertical crustal column. Therefore, the While significant attention has been paid to the plutons, defining locally high vertical and lat- temperature gradient that these data define will metamorphic conditions recorded in mid-crustal eral temperature gradients, which indicates that not represent a true estimate of the peak geo- (~4–10 kbar) rocks exhumed in the footwalls of the rise of granites was a primary mechanism thermal gradient, but instead will be a compos- these core complexes, we emphasize here that for heat transfer to the upper crust (Barton et ite estimate of peak thermal conditions obtained there is also evidence for metamorphism and el- al., 1988; Miller et al., 1988). In addition, Cam- both laterally and vertically through the upper evated temperatures at upper-crustal depths (≤4 brian and Neoproterozoic rocks exhibit syn- crust of the study area. In the absence of quan- kbar) preserved in several localities. metamorphic, east-vergent shear fabrics, which titative peak temperature data from adjacent In the Ruby–East Humboldt core complex, increase in intensity with depth, and are inter- ranges, or the ability to collect data through a exhumed mid-crustal rocks record a history of preted to have accommodated layer-parallel vertical column in the Grant Range due to over- Late Cretaceous, upper amphibolite- to granu- shear related to deformation in the Sevier thrust printing extension, the resulting metamorphic lite-facies metamorphism, ductile thickening, belt to the east (Miller and Gans, 1989). Peak

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Abbreviations 114°W Explanation Wells DC: Deep Creek Range Jurassic: E: Egan Range P plutons G: Grant Range WH K: Kern Mountains amphibolite facies

T °N P: Pilot Range PQ greenschist facies PQ: Pequop Mountains 41 Elko Cretaceous: NS: Northern Snake Range WFW NSRD: Northern Snake plutons Range decollement Wendover amphibolite facies REH: Ruby-East Humboldt REH greenschist facies core complex SC: Schell Creek Range ada Cambrian rocks Utah SS: Southern Snake Range Nev with no reported T: Toano Range metamorphism WFW: exhumed footwall (pre-extensional of Windermere thrust WH: Wood Hills stratigraphic depth WW: Wah Wah Mountains range of ~5-10 km) 116°W Roberts Mts. thrust

K E

Eureka

SC NS

NSRD Ely

SS N

9°

G WW

N old-thrust belt

0 km 20 Pioche Sevier f

Figure 7. Map showing locations of Jurassic and Cretaceous plutons in the Sevier hinterland, and spatial distributions and grades of associated metamorphic rocks (compiled from: Stewart and Carlson, 1978; Miller et al., 1988; Miller and Gans, 1989; Miller and Hoisch, 1992, 1995; Camilleri and Chamberlain, 1997; Hintze et al., 2000; Crafford, 2007). Range and valley polygons from McQuarrie and Wernicke (2005).

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metamorphism and accompanying deforma- do not preserve evidence for anomalously high low-crustal metamorphism in eastern Nevada tion is interpreted as a protracted event that may geothermal gradients, which illustrates the het- (Miller and Gans, 1989; Camilleri, 2013), and have spanned ca. 70–90 Ma, the age range of erogeneous spatial distribution of upper-crustal have been interpreted as representing the top of the majority of magmatism (Miller and Gans, metamorphism. a diffuse zone of layer-parallel shear related to 1989; Barton, 1990). Late Cretaceous K-Ar and Late Cretaceous magmatism and associated east-vergent shortening in the Sevier thrust belt 40Ar/39Ar ages obtained from several areas sug- metamorphism in the Sevier hinterland has been (Miller and Gans, 1989). At the latitude of the gest that the thermal regime cooled shortly after interpreted as the shallow thermal expression large area of shallow-crustal metamorphism in peak metamorphism (Miller et al., 1988). of an episode of lower-crustal anataxis (Bar- east-central Nevada, ca. 88 Ma marks a slowing Metamorphism and deformation in the ton et al., 1988; Miller and Gans, 1989; Wells in crustal shortening rates in the Sevier thrust Grant Range was similar in style and timing, and Hoisch, 2008). Causative mechanisms for belt by a factor of ~3 (DeCelles and Coogan, and is interpreted here as genetically related to lower-crustal melting are actively debated (e.g., 2006). Slowing of shortening rates could have the Late Cretaceous, shallow-crustal metamor- Wells et al., 2012; Miller et al., 2012), with com- been the result of the interplay of several dy- phism documented in the Deep Creek Range, peting hypotheses including eastward migration namic factors, including, but not limited to, Kern Mountains, Schell Creek Range, and Egan of subduction-related magmatism in concert changes in interplate coupling stress, changes Range (Fig. 7). Metamorphism in the Grant with conductive relaxation of isotherms within in décollement or wedge strength, or changing Range is spatially associated with the ca. 84 Ma thickened crust (Miller and Gans, 1989), or the the geometry of the sedimentary basin that the Troy granite stock, and was accompanied by influx of magma, heat, and fluids following de- thrust front migrated into. We suggest that ther- low-magnitude, ductile contractional deforma- lamination of dense mantle lithosphere (Wells mal weakening of the hinterland crust may have tion (Camilleri, 2013), similar to observations in and Hoisch, 2008; Wells et al., 2012). Regard- also contributed to slowing shortening rates, by east-central Nevada. However, the Grant Range less of their origin, the rise of anatectic melts, locally initiating ductile thickening at middle is exceptional in that it records metamorphism to their eventual emplacement as granitic plu- and lower-crustal levels. This could, at least in at shallower stratigraphic levels, up to the top tons at middle and upper crustal levels, provided part, help explain the apparently disparate ob- of the Ordovician section, which corresponds an efficient mechanism for heat transportation, servations of minimal upper-crustal shorten- to pre-extensional depths as shallow as 5 km. yielding metamorphism and conditions condu- ing (a few 10s of km) recorded across much of Similar to interpretations in east-central Nevada cive for ductile deformation at shallow depths eastern and central Nevada (Taylor et al., 2000; (Miller and Gans, 1989), the rise and emplace- (Barton et al., 1988; Miller and Gans, 1989). Long, 2015) and crustal thickness estimates of ment of the Troy pluton is interpreted to have The higher pluton density observed in east- ~50–60 km interpreted to have been attained been the primary mechanism for heat transfer to central Nevada produced upper-crustal heating across this region by the end of orogenesis (e.g., the upper crust. over a broad area (Miller et al., 1988; Barton, Coney and Harms, 1984; Gans, 1987; Chapman The Grant Range is separated from the 1990), while the Grant Range represents a local- et al., 2015). larger area of upper-crustal metamorphism in ized thermal anomaly associated with a single east-central Nevada by ~100 km (Fig. 7). One intrusion, which is surrounded by a large region ACKNOWLEDGMENTS proxy for analysis of the spatial distribution of without elevated upper-crustal temperatures. This work was supported by a grant from Makoil, Inc. (Gregg Kozlowski, CEO). We are indebted to Jerome Walker for ad- upper-crustal heating is to look at the map dis- Late Cretaceous magmatism, though spa- vice on the collection and interpretation of vitrinite reflec- tribution of exposures of Cambrian rocks, which tially localized, has implications for significant tance and pyrolysis data. Conversations with Karen Lund, Sue Beard, Joan Fryxell, and Phyllis Camilleri also contrib- correspond to a general, pre-extensional strati- thermal weakening of areas of the thickened uted to several ideas presented here. Thoughtful reviews graphic depth range of ~5–10 km (e.g., Stewart, Sevier hinterland crust. Metamorphic field gra- from Adolph Yonkee, Karen Lund, and an anonymous re- 1980), with no reported metamorphism (Fig. dients of 50–60 °C/km indicate that tempera- viewer greatly improved an early version of this manuscript. 7). Quantitative peak temperature data are not tures above the ~300 °C quartz crystal-plastic available from the majority of these localities, transition (e.g., Stipp et al., 2002) were locally REFERENCES CITED American Society for Testing and Materials (ASTM), 2014, but some published constraints do exist. Near achieved at depths as shallow as 5–6 km, and Standard test method for microscopical determination Eureka, Long et al. (2015) obtained Paleozoic indicate the potential for locally attaining the of the reflectance of vitrinite dispersed in sedimentary rocks, in Annual Book of ASTM Standards, West zircon fission-track and zircon (U-Th)/He ages ~700 °C minimum temperatures necessary for Conshohocken, Pennsylvania, v. 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