Neoarchean Tectonic History of the Teton Range: Record of Accretion

Home , Bivouac Peak, Leigh Canyon, Moran Canyon (Wyoming), Mount Owen (Colorado), Mount Owen (Wyoming), Paintbrush Divide

Research Paper THEMED ISSUE: Active Margins in Transition—Magmatism and Tectonics through Time: An Issue in Honor of Arthur W. Snoke

GEOSPHERE Neoarchean tectonic history of the Teton Range: Record of accretion

GEOSPHERE; v. 14, no. 3 against the present-day western margin of the Wyoming Province doi:10.1130/GES01559.1 B. Ronald Frost1, Susan M. Swapp1, Carol D. Frost1, Davin A. Bagdonas1,2, and Kevin R. Chamberlain1,3 1Department of Geology and Geophysics, University of Wyoming, Laramie, Wyoming 82071, USA 18 figures; 1 set of supplemental files 2Carbon Management Institute, University of Wyoming, Laramie, Wyoming 82071, USA 3Faculty of Geology and Geography, Tomsk State University, Tomsk 634050, Russia

CORRESPONDENCE: [email protected] ABSTRACT INTRODUCTION CITATION: Frost, B.R., Swapp, S.M., Frost, C.D., Bag- donas, D.A., and Chamberlain, K.R., 2018, Neoar- chean tectonic history of the Teton Range: Record Although Archean gneisses of the Teton Range crop out over an The Wyoming Province is an Archean craton that occupies most of of accretion against the present-day western margin area of only 50 × 15 km, they provide an important record of the Ar- Wyoming and portions of Montana, and adjacent states. The Archean rocks of the Wyoming Province: Geosphere, v. 14, no. 3, chean history of the Wyoming Province. The northern and southern are exposed in the cores of basement-involved Laramide uplifts. The early p. 1008–1030, doi:10.1130/GES01559.1. parts of the Teton Range record different Archean histories. The north- mafic crust appears to have been Hadean (Frost et al., 2017), though most of ern Teton Range preserves evidence of 2.69–2.68 Ga high-pressure the exposed area consists of Paleoarchean to Neoarchean quartzofeldspathic Science Editor: Shanaka de Silva Guest Associate Editor: Joshua Schwartz granulite metamorphism (>12 kbar, ~900 °C) followed by tectonic as- orthogneisses that retain an isotopic signature of that ancient crust (Frost, 1993). sembly with isotopically juvenile quartzofeldspathic metasedimen- The Wyoming craton is subdivided into three main provinces (Fig. 1; Mueller Received 16 May 2017 tary rocks under high-pressure amphibolite-facies conditions (~7 kbar, and Frost, 2006). The northwestern province is the Montana metasedimentary Revision received 18 December 2017 675 °C) and intrusion of extensive leucogranites. Together, these events province, which is an area composed of quartzite, pelitic schist, and carbonate Accepted 1 March 2018 record one of the oldest continent-continent collisional orogenies on rock associations that are structurally interleaved with quartzofeldspathic gneiss, Published online 11 April 2018 Earth. Geochemical, thermobarometric, and geochronological data from all of which were accreted at ca. 2.55 Ga. The Beartooth-Bighorn magmatic zone, the gneisses of the southern Teton Range show that this part of the up- which occupies the core of the craton, is dominated by orthogneisses. Most of lift records a geologic history that is distinct from the northern part. the Beartooth-Bighorn magmatic zone contains rocks that were last deformed It contains a variety of quartzofeldspathic gneisses, including a between 2.86 Ga and 2.71 Ga. On the southern and western margins of the 2.80 Ga granodioritic orthogneiss and the 2.69–268 Ga Rendezvous Beartooth-Bighorn magmatic zone, these older gneisses were overprinted by Gabbro. None of these preserves evidence of the granulite metamor- deformation that is as young as 2.63 Ga. The southern margin of the craton con- phism seen in the northern Teton Range. Instead, they have affinities tains the southern accreted terranes, consisting of various fragments of arcs and with rocks elsewhere in the Wyoming Province. The boundary between continents that were accreted to the Wyoming Province at ca. 2.63 Ga. the northern and southern areas is occupied by the Moran deformation The Teton Range, a small range of spectacular mountains in northwestern zone, a broad zone of high strain along which the northern and southern Wyoming, exposes some of the westernmost outcrops of the Archean Wyo- areas were assembled at ca. 2.62 Ga under moderate pressures and tem- ming Province (Fig. 2). The northern portion of the range, described by B. Frost OLD G peratures (T = 540–600 °C and P < 5.0 kbar). The final Archean event of the et al. (2006), Frost et al. (2016a), and Swapp et. al. (2018), contains some of the Teton Range was the emplacement at 2.55 Ga of the Mount Owen batholith, oldest high-pressure granulites in the world. In this paper, we summarize the a peraluminous leucogranite that intrudes the Moran deformation zone. past work on the northern Teton Range, identify a deformation zone that marks The rocks of the northern Teton Range record events that are not the contact between gneisses of the northern and southern Teton Ranges, and OPEN ACCESS present elsewhere in the Wyoming Province. We propose that they discuss how this structural belt relates to the final Neoarchean assembly of the formed at 2.70–2.67 Ga some place distal to the Wyoming Province Wyoming Province. and that they were accreted from the west against the Wyoming Province along the Moran deformation zone at ca. 2.62 Ga. This date GEOLOGIC BACKGROUND is coeval with deformation and metamorphism in the southern accreted terranes and indicates that at this time, accretion was taking place Preliminary geologic mapping of the range was conducted by John C. Reed This paper is published under the terms of the along both the southern margin and western margins of the Wyoming Jr. from 1962 to 1970 (Fig. 2; Reed, 2014). Subsequent studies (Miller et al., CC‑BY-NC license. Province. 1986; B. Frost et al., 2006; Frost et al., 2016a; Swapp et al., 2018) concentrated

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Figure 1. Map of the Wyoming Province, showing the mountain ranges exposing Archean rocks and the location of the three major subprovinces: the Montana metasedimentary province, the Beartooth-Bighorn magmatic zone, and the southern accreted terranes. The study area in the Teton Range is indicated by the rectangle and shown in Figure 2. Colors identify time of final Archean deformation and magmatism; see text for discussion. TTG—tonalite-trondhjemite-granodiorite. Figure is adapted from Frost et al. (2016a).

on the Archean history of the northern part of the range, roughly north of Leigh Canyon, where evidence of high-pressure (high-P) granulite facies is preserved. These studies show that the high-P granulites were metamorphosed at 2695 Ma and were tectonically assembled with layered gneisses at 2685 Ga (Swapp et al., 2018). These rocks were intruded by trondhjemitic leucogranites between 2685 and 2675 Ma (Frost et al., 2016a). The basement rocks of the southern Teton Range include quartzofeldspathic gneisses and hornblende gabbro (Reed, 1973; Love et al., 1992). They lack the 2675–2685 Ma leucogranites and all traces of the 2695 Ma granulite metamorphism. Intruding both domains is the Mount Owen batholith, an undeformed, peraluminous leucogranitic batholith that was emplaced at 2547 ± 3 Figure 2. Geologic map of the Archean rocks of the Teton Range, modified after Love et al. Ma (Zartman and Reed, 1998). The Mount Owen batholith underlies the rugged (1992), showing locations of samples included in this study. high peaks in the central part of the range and forms granitic and pegmatitic dikes throughout the uplift.

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Archean Geology of the Northern Teton Range Basin gneiss and the Layered Gneiss. No older zircon areas were identified,

consistent with mainly positive initial ƐNd values for the Layered Gneiss The Archean rocks of the northern Teton Range consist of three major rock (Swapp et al., 2018). types. Most distinctive of these is the Moose Basin gneiss1, a suite of mafic Reed (1973) mapped large panels of leucogranitic gneisses that intruded and pelitic rocks that were metamorphosed into granulite facies (Swapp et al., the Moose Basin and the northern Layered Gneiss, which he named the Webb 2018). The Layered Gneiss, a sequence of heterogeneous gneisses, consists Canyon Gneiss. Based on the geochemistry of these leucogranitic gneisses, dominantly of psammitic to pelitic paragneiss with minor amounts of quartz- Frost et al. (2016a) recognized two compositionally distinct units, both of ofeldspathic orthogneiss, and mafic and ultramafic rocks. The Layered Gneiss which are calcic and silica-rich. The dominant unit is the Webb Canyon Gneiss, was mapped by Reed (1973) as a single unit that extended the whole length which is ferroan and, based on the classification of Barker (1979), is low-Al. of the range. Frost et al. (2016a) and Swapp et al. (2018) separated the Moose The less-voluminous unit, the Bitch Creek gneiss, forms small dikes and plu- Basin gneiss from the Layered Gneiss. In the northern Teton Range, the Lay- tons within the Moose Basin and northern Layered gneisses. It is distinguished ered Gneiss and the Moose Basin gneiss are intruded by two compositionally from the Webb Canyon Gneiss because it is magnesian and shows high-Al distinct leucogranitic gneisses, the Webb Canyon Gneiss and the Bitch Creek values, based on the classification of Barker (1979). Frost et al. (2016a) postu- gneiss (Frost et al., 2016a). lated that these compositional differences are produced by different modes The Moose Basin gneiss extends from Moose Basin at the northern end of of origin: water-excess melting for the Bitch Creek gneiss, and dehydration the Teton Range southwards along the crest of the range for ~10 km (Fig. 2). melting for the Webb Canyon Gneiss. Because the Webb Canyon Gneiss is by It is composed of kyanite-bearing metapelitic rocks with leucogranitic patches far the most voluminous, both types of leucogranitic gneiss are included in the interpreted as leucosomes formed by partial melting (Swapp et al., 2018). Webb Canyon Gneiss unit on the map shown in Figure 2 and in the maps in Mafic rocks are also found within the Moose Basin gneiss and these likewise Frost et al. (2016a) and Swapp et al. (2018). SHRIMP U-Pb zircon ages from the contain leucosomes formed by partial melting. The leucosomes within the Webb Canyon and Bitch Creek gneisses range from 2675 to 2685 Ma (Frost et mafic rocks contain assemblages that record high-pressure granulite-facies al., 2016a), suggesting that they were emplaced as multiple intrusions during metamorphism. The Zr-in-rutile thermometer records temperatures around or within a few million years following the tectonic assembly of the Moose 900 °C. The assemblage garnet-kyanite-rutile-quartz (Grt-Ky-Rt-Qz; abbrevia- Basin gneiss and the northern Layered Gneiss. Their apparent lack of inherited

tions after Whitney and Evans, 2010) allows application of the garnet-rutile-alu- zircon and mainly positive initial ƐNd values are consistent with an origin by minosilicate-ilmenite-quartz (GRAIL) geobarometer (Bohlen et al., 1984). At partial melting of juvenile sources with little to no involvement of significantly 900 °C, the core garnet compositions yield a pressure of 12.0 kbar; rim garnet older crust (Frost et al., 2016a). compositions yield a pressure of 12.7 kbar. Since ilmenite does not occur as Tabular amphibolite bodies occur within the leucogranites of the Moose Ba- inclusions in the garnets, these pressures are necessarily minimum estimates sin gneiss, within the Layered Gneiss, and within the Webb Canyon and Bitch (Swapp et al., 2018). U-Pb sensitive high-resolution ion microprobe (SHRIMP) Creek gneisses. Locally, some amphibolites contain garnet, but most do not.

dating of zircon within the garnet-bearing leucosomes records an age of 2695 Initial ƐNd values (t0 = 2685 Ma) of the Moose Basin gneisses from the

± 7 Ma, which Swapp et al. (2018) interpreted as the age of the granulite meta- northern Teton Range are mainly negative, whereas the initial ƐNd values of morphism. Sparse zircon grains yielding older ages of 3.1–2.8 Ga, negative ini- the Layered Gneiss, Webb Canyon Gneiss, and Bitch Creek gneiss are mainly

tial ƐNd values of metapelitic gneiss samples, and Nd model ages between 3.4 positive. The positive initial values of most of the leucogranitic rocks suggest and 3.0 Ga suggest that the protoliths to the Moose Basin gneiss significantly that they could be derived by partial melting of the Layered Gneiss or other predate granulite-facies metamorphism (Swapp et al., 2018). similarly juvenile sources (Frost et al., 2016a). Contributions from the Moose In the northern Teton Range, the outcrops of the Layered Gneiss that we Basin gneiss, with its Nd isotopic signature indicating a more ancient continen- studied consist mostly of quartz-biotite-plagioclase paragneiss interlayered tal provenance, appear to have been limited. with minor amounts of quartzofeldspathic orthogneiss and low-Ca amphibolite, amphibolite, and metaperidotite. The gneiss has been metamorphosed in upper amphibolite facies and locally is migmatitic. Zircon from the leu- Archean Geology of the Southern Teton Range cosome in the migmatite yields a SHRIMP U-Pb date of 2685 ± 5 Ma, which Swapp et al. (2018) interpreted as the age of tectonic assembly of the Moose The Archean rocks of the southern Teton Range consist of three units, the Layered Gneiss, the Rendezvous Gabbro, and the Mount Owen batholith (Fig. 2). In the southern Teton Range, the Layered Gneiss extends from Snowshoe Canyon south to Rendezvous Mountain. It is dominantly a 1We capitalize unit names established by the U.S. Geological Survey, such as Layered Gneiss and Webb Canyon Gneiss. Those units named informally by the authors are not capitalized, such quartzofeldspathic gneiss that contains minor amounts of amphibolite. Both as Moose Basin gneiss and Bitch Creek gneiss. paragneiss and orthogneiss have been identified, but much of it is a

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quartzofeldspathic gneiss of uncertain parentage. The Layered Gneiss is vari- Initial ƐNd values (t0 = 2685 Ma) of basement gneisses from the southern

ably foliated and folded. Grain size is variable on the outcrop scale (Fig. 3A). domain are variable. Initial ƐNd from the Layered Gneiss vary from -0.2 to +4.0,

Biotite selvages on some leucocratic layers suggest that in some areas, the indicating relatively juvenile protoliths (B. Frost et al., 2006). The initial ƐNd val- Layered Gneiss has undergone partial melting (Fig. 3B). Because both the ues of porphyritic gneiss samples also are negative, and Nd model ages are northern and southern Layered Gneisses are highly heterogeneous, we have 3.1–3.2 Ga, suggesting that the porphyritic gneiss magma sources included

not been able to distinguish them in the field; the distinction is made based some older crustal materials (B. Frost et al., 2006). Initial ƐNd values for the on their occurrence relative to the Moran deformation zone. Rendezvous Gabbro (2685 Ma) are intermediate, at +0.0 and +0.9 (B. Frost et The “Augen” Gneiss (Reed, 1973) extends from Paintbrush Canyon north to al., 2006). Bivouac Peak. We put the word “augen” in quotation marks because by modern terminology, this is not an augen gneiss (i.e., a porphyroclastic rock with feldspar porphyroclasts). This rock is a gray, granodioritic orthogneiss with Mount Owen Batholith conspicuous, large (0.5–4 cm long) phenocrysts of white plagioclase and alkali feldspar. It contains biotite with accessory hornblende, and it is moderately The Mount Owen batholith is a peraluminous leucogranite that contains well foliated and lineated, with little to no compositional layering. The foliation muscovite and locally garnet. The grain size of the rock is heterogeneous; is concordant with the adjacent biotite quartzofeldspathic gneiss. in many outcrops, it grades from pegmatitic to aplitic, suggesting that Another distinctive orthogneiss within the southern Layered Gneiss is the water activity was highly variable during emplacement. Pegmatitic and “bright-eyed gneiss” (Bradley, 1956) found in Death Canyon. This biotite-horn- granitic dikes from the Mount Owen batholith are present throughout the blende orthogneiss is distinguished by magnetite crystals surrounded by Teton Range. In some areas, the pegmatitic dikes contain centimeter-sized white haloes where biotite is absent, which have the appearance of hundreds grains of tourmaline. The 2547 ± 3 Ma Mount Owen batholith (Zartman and of small eyes peering from the rock (Fig. 3C). Reed, 1998) is undeformed, which indicates that deformation had ceased Biotite-rich paragneiss is found throughout the southern Layered Gneiss. by that time. The initial Nd isotopic compositions of the Mount Owen batholith We have examined it in three locations. At the east end of Moran Canyon, the are consistent with an origin by partial melting of rocks with an isotopic paragneiss is interlayered with amphibolite, and leucosomes indicate that it composition similar to the older gneisses of the Teton Range, supporting has been partially melted (Fig. 3D). On the southeast face of the Grand Teton, a crustal origin for the peraluminous leucogranite batholith (B. Frost et biotite paragneiss is preserved as enclaves within Mount Owen granite (Fig al., 2006). 3E). Biotite paragneiss with the assemblage garnet-biotite-cordierite-staurolite The Mount Owen batholith is composed of quartz, alkali feldspar, pla- (Grt-Bt-Crd-St) crops out in Paintbrush Canyon and at Paintbrush Divide (Fig. gioclase, biotite, and muscovite. The alkali feldspar is perthitic. In some 3F). Concordant bodies of amphibolite with well-developed foliations, some rocks, the alkali feldspar is orthoclase, in some, it has been partially in- garnet-bearing and some garnet-free, are found throughout the southern Lay- verted to microcline, and in others, it has been completely inverted. In many ered Gneiss. The southern Layered Gneiss also includes minor occurrences of rocks, muscovite is coarse and tabular and is clearly magmatic. In others, calc-silicate schist. it is a fine, felty intergrowth that is likely to have formed deuterically. In The Rendezvous Gabbro crops out over an area of ~25 km2 in the southern- some rocks, biotite is the dominant mica; in others, it is muscovite. Most most Teton Range. It is a weakly metamorphosed rock that retains an igneous rocks show small amounts of retrogression. Plagioclase in many rocks texture consisting of centimeter-sized hornblende and plagioclase. Several is riddled with sericite, and biotite and garnet are altered to chlorite. textural features indicate that the hornblende has replaced pyroxene: (1) The Biotite locally is altered to epidote; in places, this reaction also produced

hornblende is green and not brown, indicating that it is poorer in TiO2 than titanite. Garnet is present in some samples. In a few samples, quartz contains most igneous hornblendes; (2) many of the hornblende grains contain inclu- tiny needles of sillimanite (<50 mm in maximum dimension), suggesting sions of quartz, which we interpret as having formed by the reaction: augite + that crystallization began in the sillimanite field and then moved into

orthopyroxene + plagioclase + H2O = hornblende + quartz (see Tracy and Frost, the muscovite field as temperature fell or as water fugacity increased. 1991); and (3), the hornblende locally has zones of non-pleochroic tremolite, Zircon is abundant, and allanite and monazite were observed in some which we interpret as replacements of the remnants of the original clinopy- samples. roxene. Plagioclase for the most part preserves the tabular shape of igneous feldspar. Locally, it shows deformation twinning, and, in places, it is bent and contains abundant deformation twins. Some grains are rimmed by plagioclase METHODS neoblasts, indicating that it has undergone some high-temperature deformation involving fast grain-boundary migration. On the whole, the Rendezvous We compiled structural data on the deformed Archean rocks of the Teton Gabbro has not undergone extensive deformation. Range from two sources: measurements by the authors and measurements

GEOSPHERE | Volume 14 | Number 3 Frost et al. | Neoarchean tectonic history of the Teton Range 1011 Research Paper Fig. 3E is about 0.7 m thick. Strongly foliated(F) biotite paragneiss at Paintbrush Divide within the Moran deformation zone. Field of view for A and B is about 1 meter. The horizontal pegmatite dike in Avalanche Canyon. (C) Bright-Eyed gneiss, Death Canyon. (D) Biotite paragneiss, Moran Bay. (E) Biotite paragneiss, Upper Exum route near summit of the Grand Teton. Figure 3. Photographs of the various types of Layered Gneiss in the southern Teton Range. (A) Southern Layered Gneiss, upper Garnet Canyon. (B) Southern Layered Gneiss, B A C D E F

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by J.C. Reed Jr. and his field assistants made during the period 1962–1970 and HCl and HNO3-HCl ion exchange columns, respectively. Isotope ratios were recorded in field notebooks archived at the U.S. Geological Survey (USGS) in analyzed on the Micromass S54 at the University of Wyoming. Pb was run in Denver, Colorado. Poles to foliations were plotted and interpreted using Stere- static-multicollector mode with 204Pb in the Daly; U was loaded with graphite onet 9 software described by Cardozo and Allmendinger (2013). A compilation and run as a metal in static Faraday mode. Pb blanks for titanite averaged 8 the measurements of the Reed parties as well as the data from the University pg, and U was less than 0.1 pg. Stacey and Kramers (1975) model values were of Wyoming parties is available as Supplemental Tables S1–S52. used for initial Pb isotopic compositions for titanite from 07T23 and 98T12, U-Pb geochronology was undertaken by two methods. U-Pb isotopic data and the Pb isotopic composition of coexisting feldspar was used for samples on zircon from biotite paragneiss and porphyritic gneiss in the southern Teton 07T22 and 04T16. U-Pb isotopic data for the samples analyzed by SHRIMP-RG Range were obtained using the sensitive high-resolution ion microprobe–re- are reported in the supplemental materials Table S6, and U-Pb isotopic data for verse geometry (SHRIMP-RG) at Stanford University and at the Australian Na- samples analyzed by TIMS are reported in Table S7 (see footnote 2). tional University (ANU). Zircon grains were mounted in epoxy along with chips Samples for whole-rock geochemical analysis were prepared by crushing of reference zircon. At Stanford, the standards were Duluth Gabbro (1100 Ma; rock trimmed of weathered surfaces between tungsten carbide plates in a hy- Paces and Miller, 1993) and R33 (Black et al., 2004), and at ANU, the standard draulic press, followed by powdering in a ceramic or tungsten carbide ring was Temora (417 Ma; Black et al., 2003). After polishing, cathodoluminescence mill. Major elements and a limited suite of trace elements were determined (CL) scanning electron microscope (SEM) images were taken for all zircon by X-ray fluorescence (XRF) at the University of Wyoming. Reproducibility de- grains. Isotopic ratios and U, Th, and Pb concentrations were measured using termined from replicate analyses was better than 0.01% for all oxides, except

procedures similar to those given in Williams (1998, and references therein). SiO2, which was better than 0.1%. Trace-element concentrations, including rare The data were reduced using the SQUID Excel macro of Ludwig (2001). Un- earth element (REE) concentrations, were determined by inductively coupled certainties given for individual analyses (ratios and ages) are at the 1σ level, plasma mass spectrometry (ICP-MS) at XRAL Laboratories, Don Mills, Ontario, with correction for common Pb made using the measured 207Pb/206Pb ratios. Canada, or ALS Minerals, Reno, Nevada. Average reproducibility was <5% for A-59 A-60 A-61 A-65 A-62 Concordia plots and linear discordia regression fits were carried out using Iso- all elements, except for Rb, Hf, Eu, Tb, Tm, and Yb, which were <15%. These A-64 A-63 A-55,56 plot (Ludwig, 2003), and uncertainties are reported at the 95% confidence level. data, together with previously published analyses, are presented in Table S8 A-57 D-89 A-58 D-88 D-87 D-70 Zircon grains separated from two samples of the Rendezvous Gabbro and (footnote 2). Sample locations are provided in Table S9 (footnote 2). D-71 D-62 D-86 D-61 D-77 a boudinaged amphibolite dike in the Bitch Creek gneiss were dated by a pro- Table S10 (footnote 2) compiles Sm-Nd isotopic data for the southern Teton D-72 D-60 D-76 D-73 D-69 D-74 D-75 cedure modified from the chemical abrasion–thermal ionization mass spec- Range, most of which were previously published by B. Frost et al. (2006). D-68 D-67 D-66 D-101 D-78 D-79 D-81 D-63 A-71 D-80 D-64 D-65 A-69,70 A-72 A-67,68 trometry (CA-TIMS) method of Mattinson (2005) at the University of Wyoming. Data on sample 07T3 was obtained following methods provided in Frost et D-94 A-66 D-90 A-12 A-13 A-37 D-91 A-23 D-92 A-14 A-21 A-22 D-93 A-20,21 A-39 Representative zircon grains from each morphological subpopulation were al. (2016a). D-95 A-15 A-40 A-38 A-24 A-16 A-36 A-43 A-41,42 A-17,18 A-27 A-31-35 A-25,26, B-10 B-12 annealed at 850 °C for 50 h and then dissolved in two stages in HF and HNO . Mineral analyses were obtained on polished thin sections for garnet, pla- 28-30 3 B-11 B-27 B-19 B-28 B-18 B-17 B-25√ B-24 B-23 B-26 B-29 B-20 The first step was for 12 h at 180 °C, which removed the most metamict gioclase, and hornblende using a JEOL 8900 Superprobe electron microprobe B-21 B-1 B-22 B-2 B-3 B-4,5 B-30 B-6 B-31 205 233 235 B-13 B-7 B-16 B-64 domains. Single grains were then spiked with ET535, a mixed Pb- U- U with a 15 kV accelerating voltage and sample current of 20–30 nA for all min- B-14 B-32 B-9S B-33-35 B-8 B-52 B-53 B-63 B-65 B-66 B-51 B-50 B-54 B-15 B-60 B-59 B-49 B-62 B-56 B-48 B-61 B-57 tracer, completely dissolved at 240 °C for 30 h, converted to chlorides at 180 °C erals. Both natural and synthetic silicates and oxides were used as standards. B-58 B-47 B-46 B-68 B-67 B-36 B-69 B-73 B-72 A-44 B-37 B-70 B-71 A-45,46 B-38 B-95 overnight, and loaded onto Re filaments with H3PO4 and silica gel. Pb and UO2 Analyses were corrected using the ZAF procedure. Errors were less than 1% A-47 B-40 A-48 B-96 A-49 B-97 B-39 A-50 B-98 B-99 A-54 isotopic ratios were measured in single-collector mode using a Daly-photo- of the measured values of major elements and increased greatly for elements multiplier collector on a Micromass S54 TIMS. Pb fractionation was deter- with abundances below 1 wt% oxide. Representative analyses are presented mined by multiple analyses of NIST 981, and U fractionation was determined in Table S11 (footnote 2).

2Supplemental Tables. Table S1: Reed’s field internally. Pb procedural blanks averaged 2 pg, and U was consistently less locations N. Table S2: Reed’s field loca- than 0.1 pg. Isotopic compositions of additional common Pb from each anal- tions S. Table S3: Reed’s field books. Table ysis of the Rendezvous Gabbro were modeled using the model of Stacey and RESULTS S4: USGS field assistant’s field book. Table Kramers (1975); data from the red zircon grains in 04T16 were reduced with S5: University of Wyoming structure data. Table S6: U-Pb SHRIMP data. Table S7: whole-rock Pb values, and data from colorless zircon grains in 04T16 were Structural Geology—Moran Deformation Zone U-Pb TIMS data. Table S8: Geochemi- reduced with the common Pb solution to a three-dimensional (3-D) total Pb cal data. Table S9: Sample location data. isochron. Raw data were reduced using PbMacDat and Isoplot, based on the The major structural feature in the Tetons is a deformation zone that Table S10: Sm-Nd isotopic data. Table S11. Mineral chemistry data. Table S12: U-Pb algorithms of Ludwig (1988, 1991). crops out between Leigh and Snowshoe Canyons (Fig. 2). This deformation data on zircon grains from the Bridger Titanite from three samples of Rendezvous Gabbro and from the boudi- zone, which has not been recognized previously, is over 2 km wide in Moran batholith, Wind River Mountains. Please visit naged amphibolite dike in the Bitch Creek gneiss were selected based on size Canyon and extends south from Moran Canyon through the western por- http://doi.org/10.1130/GES01559.S1 or the full-text article on www.gsapubs.org to view and dissolved in HCl:HF ratio (5:1) at 180 °C in Parrish-style dissolution bombs tion of the Moran massif and over Thor Peak. Because of its exposure in the the Supplemental Tables. within a high-pressure Parr dissolution cell. Pb and U were purified on HBr- Moran Canyon and Moran massif, we call it the Moran deformation zone.

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It includes high-strain zones, tens to hundreds of meters wide, that are The central area is characterized by consistent north-northeasterly striking marked by increasing intensity of foliation, grain-size reduction, and the de- foliations that dip moderately to the east (Fig 5B). This fabric is roughly parallel velopment of porphyroclasts. Figure 4 shows one of these high-strain zones to the strike and dip of the shear zones within the Moran deformation zone, located near Triple Glacier on the north face of the Thor Peak–Mount Moran and we conclude that this fabric represents foliations that formed during the massif. The northern end of the Moran deformation zone has not been traced deformation, as well as any earlier foliations on either the footwall or headwall across Snowshoe Canyon, but based upon lithologic contrasts across the de- that were transposed into parallelism with the Moran deformation zone. formation zone, it probably projects to the east of the Archean outcrop north We have done limited field work in the southern Teton Range; the fabric of Snowshoe Canyon. The southern end of the Moran deformation zone is data from this area are mainly those of Reed and his field assistants. Multi- exposed in the slopes south of Leigh Canyon but is obliterated by the Mount ple fold generations exist (for example, see Fig. 3B), but the major foliation in Owen batholith before reaching Paintbrush Divide. We project it to extend ap- the southern Teton Range displays broad open east-trending folds that plunge proximately to Hurricane Pass, west of the outcrops of the southern Layered gently to the east (Fig. 5C). Considering that the trace of the Moran deforma- Gneiss (Fig. 2). tion zone likely lies ~5 km to the west of the westernmost outcrop of the Lay- Foliations in the Moran deformation zone in Moran Canyon give strikes ered Gneiss in this area, we suggest that these folds and foliations represent that average 10° and dips of around 40° to the east. Considering that the sed- the structure of the hanging wall of the Moran deformation zone and that they imentary rocks on the west side of the Teton Range dip ~8° to the west, the were not transposed during the thrusting event. dip of the zone may have been somewhat steeper before uplift on the Teton fault. The foliation in the high-strain zones folds around numerous phacoids of less-deformed gneisses, which account for the variations in strikes and dips Geochronology we obtained from the deformation zone. Lineations generally are downdip, and the sense of shear as determined by asymmetrical porphyroblasts is top- For this study, we obtained U-Pb isotopic analyses of zircon from six sam- to-the-west, which suggest that the deformation zone represents a broad area ples. The first, 07T1, is a sample of the “Augen” Gneiss in the southern Layered of strain marking a crustal-scale thrust fault. Gneiss collected in Paintbrush Canyon (Table S6 [footnote 2]). This sample was analyzed to constrain the intrusive age of the “Augen” Gneiss. Analyses of 13 areas on 10 grains yielded a wide range of U contents, from ~140 to ~2040 Structural Fabrics ppm. The four analyses with the U highest contents (>1450 ppm) were also the most discordant and were not interpreted. Another high-U grain area gave the Apart from the Mount Owen batholith, which is undeformed, the Archean oldest 207Pb/206Pb age of 2870.7 ± 5.4 Ma. The remaining eight analyses yielded rocks of the Teton Range are variably foliated gneisses. Minor folds occur lo- a concordia age of 2803 Ma, but there is some scatter in the 207Pb/206Pb ages, as cally on all scales, as do higher-strain zones. Stereographic projections of fo- indicated by the mean square of weighted deviates [MSWD] of 2.7. If the scatter liations measured across the range suggest that the Archean rocks of Teton is due to slight inheritance in two grain areas, a chord through the remaining Range can be divided into three structural domains (Fig. 5). The northern area six analyses yields a date of 2800 ± 8 Ma (MSWD = 1.4; Fig. 6A). We interpret lies north and west of the Moran deformation zone, the central area extends this date as the best estimate for the magmatic age of the “Augen” Gneiss. from the vicinity of Mount Moran south to the east face of Teewinot, and the We also analyzed detrital zircon from two samples of biotite paragneiss southern area extends from Avalanche Canyon to the southern limits of Pre- from Paintbrush Canyon. One sample, 07T3, was composed mainly of low-U cambrian outcrop. zircon (35–450 ppm), with only a few higher-U, more discordant zircon grains. In the northern area, the Moose Basin gneiss and Layered Gneiss show Neglecting the four analyses that were more than 5% discordant, the remain-

three generations of folds (Swapp et al., 2018). The earliest (F1) folds are seen ing analyses define several age populations. Three grains defined an upper-in-

as isolated isoclinal fold hinges. The second generation (F2) is seen as large iso- tercept age of ca. 2640 ± 21 Ma, six grains yielded a date of 2663 ± 16 Ma, and clinal folds that are up to hundreds of meters wide. These two fabric elements six other grains gave 2711 ± 10 Ma. Two older grains yielded 207Pb/206Pb dates are mostly absent in the Webb Canyon gneiss, which typically shows only a of 2.84 and 2.86 Ga (Fig. 6B). planar foliation. The major foliation in this area is parallel to the fold axes of the Another biotite paragneiss sample, 10T4, contained a few zircon grains,

F2 folds. This foliation has been folded into broad open, north-trending series many of which were relatively high in U and discordant. Of the 13 analyses,

of folds (F3) that become tighter moving from west to east. There is no axial only six were less than 35% discordant. These analyses defined two age

planar foliation related to the F3 folds. Because the fabric on the east limb of the groups of 2630 ± 13 Ma and 2663 ± 9 Ma (Fig. 6C). We note that these two age

F3 folds is parallel fabric to the Moran deformation zone (see Figs. 5A and 5B), groups are also present in sample 07T3.

we interpret the F3 folds to have formed in response to the Moran deformation The age of the Rendezvous Gabbro was constrained by U-Pb dates on zone and that the major foliation formed before the Moran deformation zone. single zircon grains from sample 98T12, a coarse-grained hornblende gabbro

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Figure 4. Field photographs of the Moran deformation zone. (A) Photograph of the north face of the Thor Peak–Mount Moran massif near the westernmost portion of Triple Glacier. Foliated gneisses within the Moran deformation zone are crosscut by several generations of undeformed dikes of Mount Owen granite. Red arrows indicate locations of structural measurements. Triangular peak in the middle of the photograph is about 100 m high. (B) Strongly foliated gneiss from a high-strain zone of the Moran deformation zone located to right of field of view shown in A.

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A Porphyritic “Augen” Gneiss A fold plunges 24° to 21° 07T1 2800

Northern Teton Range 2700 Poles to foliations N = 538 2600

Strike and dip of the Moran deformation zone

Central Teton Range Poles to foliations N = 122 Figure 6 (on this and following three pages). (A) U-Pb concordia diagram for zircon from the “Augen” Gneiss sample 07T1. A regression through six of eight analyses (red ellipses) yielded a concordia age of 2800 ± 8 Ma (mean square of weighted deviates [MSWD] = 1.4). C

euhedral grains produced a range of 207Pb/206Pb dates between 2688 and 2683 Ma, and small euhedral grains gave the youngest weighted mean age of 2678 fold plunges 13° to 102° ± 3 Ma (Fig. 6D). These results suggest that the emplacement and crystallization of the Rendezvous Gabbro involved several episodes of over a period of Southern Teton Range several million years. Poles to foliations The boudinaged mafic dike in the Bitch Creek gneiss (04T16) yielded two N = 83 distinct euhedral zircon populations: red and colorless. Three, single-grain, CA- TIMS analyses of the red population gave a weighted mean 207Pb/206Pb date of 2682.1 ± 1.6 Ma (Fig. 6E; Table S7 [footnote 2]). We interpret these grains to be xenocrysts from the host Bitch Creek gneiss (06T17), which yielded a U-Pb zircon date of 2685.7 ± 4.2 Ma (Frost et al., 2016a). Four, single-grain CA-TIMS Figure 5. Structural data from the gneisses of the Teton Range. (A) Poles to foliation from the northern portion of the Teton Range. (B) Poles to foliation from the central portion of the Teton analyses of the colorless population had significantly lower U concentrations Range. (C) Poles to foliation from the southern portion of the Teton Range. than those of the red zircons and relatively high amounts of common Pb, with blank-corrected 206Pb/204Pb from 1043 to 142 (Table S6 [footnote 2]). A total Pb 3-D isochron solution of these analyses indicated an isotopically evolved excavated from the Jackson Hole Mountain Resort (Table S7 [footnote 2]). Four composition for this common Pb (206Pb/204Pb = 21.4 ± 6.2, 207Pb/204Pb = 16.4 ± single-grain CA-TIMS analyses of red, anhedral grains gave an upper-intercept 4.5). Propagating errors with these values produced relatively large concordia date of 2685.3 ± 2.2 Ma (MSWD = 1.3; Fig. 6D). Sample 07T23, collected from ellipses (Fig. 6E) and a concordia age of 2667.6 ± 4.3 Ma. a garnet hornblende dike that cuts the Rendezvous Gabbro, contained three Three sizes of titanite picks from Rendezvous Gabbro sample 98T12 yielded zircon morphologies, each of which yielded different dates. Three large an- a range of concordant 207Pb/206Pb dates, from 2615.4 ± 1.7 Ma (large ≥300 μm hedral grains yielded an upper-intercept date of 2691.7 ± 2.8 Ma, five medium diameter), to 2614.0 ± 1.7 Ma (medium ~200 μm), to 2611 ± 1.7 Ma (small

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0.55 0.56 Intercepts at B Biotite paragneiss C 71±980 & 2663±16 [±17] Ma 2800 0.54 MSWD = 0.44 0.50 10T4 2600

0.52 Biotite paragneiss 07T3 0.45 0.50 U 2600 U

23 8 Intercepts at / 238 2200 160±1100 & 2711±10 [±11] Ma / 0.40 Pb 0.48 Intercepts at Pb

20 6 MSWD = 0.65 20 6 39±94 & 2630±13 [±14] Ma 0.46 MSWD = 0.034 0.35 Intercepts at 0.44 1098±2600 & 2640±21 [±22] Ma Intercepts at MSWD = 0.60 0.30 0.42 123±67 & 2662.9±9.1 [±11] Ma MSWD = 0.025

0.40 0.25 10 11 12 13 14 15 16 7911 13 15 207 235 Pb/ U 207 235 Pb/ U

Figure 6 (continued ). (C) U-Pb concordia diagram for zircon from the biotite paragneiss sample 10T4. Sparse zircon grains define two detrital age populations at approximately 2630 Ma (blue 0.56 Intercepts at ellipses) and 2660 Ma (red ellipses). 71±980 & 2663±16 [±17] Ma 0.54 MSWD = 0.44 ≤100 μm; Table S7 [footnote 2]). A large single titanite grain from 07T23 0.52 yielded a date of 2622 ± 5 Ma. Different dates for dark and pale titanite were obtained from sample 07T22, a sample of Rendezvous Gabbro cut by dike 0.50 sample 07T23. All the titanite grains from 07T22 were mechanically abraded to U 2600 8

23 Intercepts at try to isolate end-member growths and/or a diffusion gradient. The two most / 160±1100 & 2711±10 [±11] Ma intensely abraded fractions of pale grains produced a weighted mean Pb 0.48 207 206

20 6 MSWD = 0.65 Pb/ Pb date of 2626.7 ± 2.5 Ma (MSWD +0.0005). The four-point weighted mean, moderately abraded pale grains and a single extra-large dark grain gave 0.46 2622.4 ± 0.9 Ma (MSWD = 0.74), and a two-point weighted average of dark Intercepts at titanite picks gave 2616.5 ± 1.1 Ma (MSWD = 0.02). Five titanite fractions from 0.44 1098±2600 & 2640±21 [±22] Ma 04T16, the boudinaged mafic dike intruding Bitch Creek gneiss in the northern MSWD = 0.60 domain, produced an upper-intercept age of 2622.8 ± 2.5 Ma (MSWD = 0.68; 0.42 Fig. 6F; Table S7 [footnote 2]). These ages are consistent with titanite growth at ca. 2626 Ma, ca. 2622 Ma, and possibly as young as 2615 Ma. 0.40 10 11 12 13 14 15 16 207Pb/235U Geochemistry

Figure 6 (continued ). (B) U-Pb concordia diagram for zircon from the biotite paragneiss sample 07T3, identifying detrital zircon ages clustered around 2640, 2660, and 2710 Ma (ellipses with Thirty-two whole-rock geochemical analyses of samples from the Layered green, blue, and yellow centers, respectively). Gneiss, Moose Basin gneiss, Rendezvous Gabbro, and Mount Owen batholith

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Figure 6 (continued ). (D) U-Pb concordia diagram for zircon from Rendezvous gabbro samples. Zircon analyses from sample 98T12 define a concordia age of 2685.3 ± 2.2 Ma (MSWD = 1.3). Three different zircon morphologies from 07T23 were analyzed (see photo inset): Isotopic compositions for large anhedral grains are shown in dark blue, equant medium-sized grains are in green, and small equant grains are in light blue. The large anhedral grains yield an age of 2691.7 ± 2.8 Ma (MSWD = 0.45).

are presented on Table S8 (footnote 2), along with data from these units pre- roan (Fig. 7A). The silica-rich rocks are mainly calcic, but some are calc-alkalic

viously published by Miller et al. (1986), B. Frost et al. (2006), and Wilks (1991). (Fig. 7B). Almost all are peraluminous (Fig. 7C). The Na2O contents of most

We subdivided the Layered Gneiss into northern and southern groups, based on felsic gneiss samples cluster between 2.1% and 4.1%, but K2O varies widely location with respect to the Moran deformation zone. Together with data on the from 0.1% to 4.5% (Figs. 8A and 8B). Almost all Layered Gneiss samples have

Webb Canyon and Bitch Creek leucogneisses published by Frost et al. (2016a), Na2O > K2O. Rb contents are low (45–130 ppm). Sr varies from 65 to 450 ppm, these data represent all the available geochemical analyses on Archean rocks suggesting a role for fractional crystallization or accumulation of plagioclase from the Teton Range. (Fig. 9). Other trace-element characteristics, including Y and Nb contents, are typical of magnesian, calc-alkalic to calcic granitoids generally (Table S8 [footnote 1]; cf. Frost et al., 2016a). There are two important differences between Layered Gneiss the geochemistry of the northern and southern Layered Gneisses. Because the

K2O contents of most northern Layered Gneisses are lower than those of the The northern and southern Layered Gneiss groups are composed of quartz- southern Layered Gneisses, the average northern Layered Gneiss group has

rich rocks and amphibolites, a bimodal assemblage that is reflected in their lower K2O/Na2O than the average southern Layered Gneiss group (Fig 8C). In silica contents (Table S8 [footnote 2]). The quartzofeldspathic rocks range from addition, Zr is high in northern Layered Gneiss (Zr = 436–692 ppm), whereas

68 to 78% SiO2, whereas the mafic rocks contain less than 52% SiO2. The only most southern Layered Gneiss samples have Zr between 100 and 300 ppm. intermediate-composition sample is biotite paragneiss. Most are magnesian, The two biotite paragneiss samples from the southern Layered Gneiss have although one sample of silica-rich northern Layered Gneiss is strongly fer- REE abundances similar to those of average shales, yielding REE patterns with

GEOSPHERE | Volume 14 | Number 3 Frost et al. | Neoarchean tectonic history of the Teton Range 1018 Research Paper Titanite

data-point error ellipses are 2σ 0.510 F

0.506 2640 2630

0.502 2620

U 2610 238 / 0.498

Pb 2600 20 6 2590 0.494

04T16 (Red): 0.490 449±310 & 2622.8±2.5 Ma MSWD = 0.68

0.486 Figure 6 (continued ). (E) U-Pb concordia diagram for zircon from the Moose Basin boudinaged 11.711.9 12.112.3 12.5 mafic dike sample 04T16. Anhedral zircon grains define an age of 2682.1 ± 1.6 Ma, identical within error to the age of the host Bitch Creek gneiss. A second population of lower-U zircons 207Pb/235U yielded a concordia age of 2667.6 ± 4.3 Ma.

Figure 6 (continued ). (F) U-Pb concordia diagram for titanite from the Teton Range. Samples 98T12, 07T22, and 07T23 (data shown by light blue, green, and navy ellipses, respectively) are from light (L) REE enrichment and flat heavy (H) REE patterns (Fig. 10A). One sample the Rendezvous Gabbro in the southern Teton Range, and sample 04T16 (red ellipses) is a boudi- has a slight positive Eu anomaly. The third southern Layered Gneiss sample, naged amphibolite dike intruding Bitch Creek gneiss the northern Teton Range. Five analyses of titanite from 04T16 yielded a concordia age of 2622.8 ± 2.5 Ma; a sixth analysis with a slightly an orthogneiss from Death Canyon, is enriched in LREEs and has a negative older 207Pb/206Pb age was not included in the regression. The 207Pb/206Pb ages of titanite from the Eu anomaly. REE patterns of northern Layered Gneisses are more variable and Rendezvous Gabbro samples range from 2616 to 2627 Ma (Table S7 [text footnote 2]). include a group with high REE abundances and deep negative Eu anomalies (Fig. 10B). These patterns are reminiscent of the “seagull”-shape patterns of the Webb Canyon Gneiss (shaded area on Fig. 10; Frost et al., 2016a). Others have steep LREE/HREE patterns with no Eu anomalies, and one sample has a steep pattern but low REE abundances and a strong positive Eu anomaly (Fig. contents with deep, negative Eu anomalies (Fig, 10D), and share other major- 10B). Clearly, the Layered Gneisses in both domains include rocks with a vari- and trace-element characteristics in common with the Webb Canyon Gneiss ety of composition, origin, and petrogenetic evolution. (Table S8 [footnote 2]; cf. Frost et al., 2016a). We conclude from their geo- Amphibolite bodies and dikes from the northern Teton Range contain 45%– chemical composition that these are samples of the Webb Canyon Gneiss

52% SiO2, are metaluminous, and define an iron-enrichment trend (Fig. 7A). that intruded the Moose Basin gneiss.

K2O, Rb, Nb, Y, and Zr contents are low. REE contents are low, and REE patterns The majority of the Moose Basin gneiss samples have the aluminous are flat. Modest negative or positive Eu anomalies suggest plagioclase frac- nature expected of metapelitic rocks (Fig. 7C) and LREE-enriched REE pat- tionation or accumulation in some samples (Fig. 10C). terns (Fig. 10D). Three metapelitic rocks have silica contents much lower than average Archean shale (for example, average Canadian Archean shale; Cameron and Garrels, 1980; see Figs. 7 and 8) and may represent restites Moose Basin Gneiss from which leucosome has been extracted. Other Moose Basin gneiss sam-

ples have SiO2 contents higher than average Archean shale (Figs. 7 and 8).

The Moose Basin gneiss samples range in silica content from 51% SiO2 The protoliths of these more siliceous Moose Basin paragneiss may have

for biotite-rich metapelitic gneisses to 80% SiO2 for felsic gneiss. The three been more quartzose than shale, or these samples may have incorporated most silica-rich samples are also strongly ferroan (Fig. 7A), have high REE partial melt.

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Figure 7. Geochemical diagrams of (A) Fe-index, (B) modified alkali-lime index (MALI), and (C) alu- Figure 8. (A) Na O, (B) K O, and (C) K O/Na O as a function of silica content for Archean rocks minum saturation index (ASI) for Archean rocks from the Teton Range. Boundaries are from Frost 2 2 2 2 et al. (2001a). Dark- and light-shaded fields represent compositions of the Webb Canyon and Bitch of the Teton Range. Most Layered Gneisses are less sodic than average tonalite-trondhjemite- Creek trondhjemitic leucogranitic gneisses, respectively. Values shown for comparison are aver- granodiorite (TTG). Symbols as the same as in Figure 7. age tonalite-trondhjemite-granodiorite (TTG; Martin et al., 2005) and average Canadian Archean shale (Cameron and Garrels, 1980). MBG—Moose Basin gneiss; NLG—Northern Layered Gneiss.

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Mount Owen Batholith

Analyses of 21 samples of the Mount Owen batholith are shown in Table S8 (footnote 2). Normative quartz–alkali feldspar–plagioclase (Q-A-P) compositions of the Mount Owen batholith calculated from these analyses plot within the stability field of granite, except for one that lies on the granodiorite-tonalite boundary (Fig. 11). To calculate these normative compositions, we subtracted the normative amount of alkali feldspar needed to combine with normative corundum to make muscovite. Because these rocks are dominated by quartz and feldspar, these calculated compositions should come close to true modal compositions. The modal abundance of plagioclase, however, will be over- estimated because small amounts of albite will have dissolved into the alkali feldspar. We observe from Figure 11 that the rocks are true granites and that none of the samples is so poor in silica that it is a quartz monzonite. It is for this reason that we advocate for calling the unit the Mount Owen batholith, rather than the Mount Owen Quartz Monzonite, as named by Reed (1973) and Love et al. (1992). Compositionally, the Mount Owen batholith is similar to other peraluminous leucogranites defined by Frost et al. (2001a) and Frost et al. (2016b) in that it spans the ferroan-magnesian boundary (Fig. 12A). It is both calc-alkalic and alkali-calcic, but the composition range covers a more narrow range of modified alkali-lime index (MALI) values than do peraluminous leucogranites as a group (Fig. 12B). The variation in Fe-index is likely to reflect degree of melting, with small-degree melts being more enriched in Fe than larger ones (Frost et al., 2001a; Frost et al., 2016b). The variation in MALI is ascribed to a different water fugacity during melting. For example, dehydration melting of micas will produce alkalic melts, whereas with increasing water activity, more plagioclase will be involved in the melting reactions, making the melt more calcic. The Mount Owen batholith shows a distinct increase in aluminum Figure 9. Selected trace-element abundances for Archean rocks of the Teton Range. (A) saturation index (ASI) with increasing silica (Fig. 12C). This probably reflects Rb content as a function of SiO . (B) Sr content as a function of SiO . MBG—Moose Basin 2 2 the fractionation of feldspars, which would increase both SiO and ASI in the gneiss; NLG—Northern Layered Gneiss; TTG—tonalite-trondhjemite-granodiorite.Sym- 2 bols are the same as in Figure 7. residual melt. The Zr content of the Mount Owen batholith ranges from 40 to 185 ppm, which is similar to, but on the high side of, other peraluminous leucogranites (Frost et al., 2016b). Zr decreases with increasing silica, indicating zircon crystallization with differentiation; however, there is a broad spread in the pattern, suggesting that some zircon in the Mount Owen batholith may be xenocrystic or cumulate (Fig. 13). The zircon-saturation temperature of the rocks is high, Rendezvous Gabbro ranging from below 750 °C to nearly 900 °C. These results indicate that in the rocks with apparently high zircon-saturation temperatures, some of the zircon Analyses of the Rendezvous Gabbro are limited to four samples: three is probably cumulate, but even the rocks with the lowest Zr abundances record coarse-grained hornblende-plagioclase gabbro samples and one fine- temperatures that are high for melting of felsic crust. grained garnet-bearing gabbro dike (Table S8 [footnote 2]). The coarse gab- The Mount Owen batholith is somewhat richer in REEs than the typical Hi- bros are magnesian, and the gabbro dike is weakly ferroan (Fig. 7A), and malayan leucogranite. The high Zr content of the Mount Owen batholith may all samples are metaluminous (Fig. 7C). Trace-element contents are within reflect the presence of cumulate REE-bearing accessory minerals. The rocks the range defined by the amphibolites within the southern and northern have a distinct negative Eu anomaly, which is probably caused by melting in Layered Gneiss and Moose Basin gneiss (Figs. 8 and 9). the presence of a plagioclase-bearing restite (Fig. 14).

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Figure 10. Normalized rare earth element patterns for Archean rocks from the Teton Range. (A) Southern Layered Gneiss. (B) Northern Layered Gneiss. (C) Mafic rocks, with peridotite shown by dashed line. (D) Moose Basin gneiss. Also shown as shaded areas are the fields for Webb Canyon and Bitch Creek leucogranitic gneiss samples from the northern Teton Range from Frost et al. (2016a).

Sm-Nd Isotopic Data The Nd contents of samples of southern Layered Gneiss vary widely from 26 to 228 ppm (Table S10 [footnote 2]). The analyzed samples are mainly Sm-Nd isotopic data for the leucogranites of the northern domain have been quartzofeldspathic gneiss, but they also include an amphibolite (99T2) and published by Frost et al. (2016a), and Sm-Nd isotopic data for the Moose Basin a biotite paragneiss (07T3). Two samples with high Sm, Nd, and Y contents

gneiss and northern Layered Gneiss were reported by Swapp et al. (2018). Sm-Nd also have the most positive initial ƐNd of 2.7 and 4.0. All other southern Lay-

isotopic data for the southern Layered Gneiss, including the “Augen” Gneiss and ered Gneiss samples have initial ƐNd between 0 and 1.0. Nd isotopic compo- the Rendezvous Gabbro, were reported by B. Frost et al. (2006). Table S10 (foot- sitions of two samples of Rendezvous Gabbro, originally reported in B. Frost

note 2) compiles these data along with one unpublished analysis for the southern et al. (2006), also have ƐNd at 2685 Ma of between 0 and 1.0. These values are

domain, and ƐNd values calculated for 2685 Ma are plotted on Figure 15. within the range of initial ƐNd documented for the northern Layered Gneiss

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silexite granodiorite tonalite 60 60 Figure 11. Normative quartz (Q)–alkali-feldspar (A)–plagioclase (P) diagram showing compositions of Mount alkali feldspar quartz monzodiorite Owen batholith samples. Normative feldspar compo- granite quartz monzogabbro sitions were adjusted for the modal abundance of Co granite (corundum) by combining Co and K-feldspar to make muscovite. quartz quartz diorite alkali feldspar quartz gabbro syenite quartz anorthosite 10 35 65 90 20 20 alkali feldspar quartz quartz monzodiorite syenite syenite monzonite monzogabbro 5 5 diorite syenite monzonite A P gabbro anorthosite

and are more radiogenic than ƐNd for most samples of Moose Basin gneiss Gerald, 2008). Both varieties occur in both the Moose Basin gneiss and in (Fig. 15). the Layered Gneiss. Fitz-Gerald evaluated pressures and temperatures for The two analyses of the “Augen” Gneiss have similar Sm and Nd con- these rocks using the barometer of Kohn and Spear (1990) and the thermom- tents. At the estimated time of crystallization (2800 Ma), these gneisses had eter of Dale et al. (2000). Mineral chemistry appropriate for this barometer

ƐNd of –0.7 to –0.5, values that rule out derivation solely from depleted mantle occurred only in the lower Fe/(Fe + Mg) amphibolites, represented by sam- sources. No older rocks have been identified within the Archean exposures of ple 03T5 and 07T9. the Teton Range; however, the “Augen” Gneiss analyses lie within the range of values for 2.86–2.80 Ga gneisses from the eastern Beartooth Mountains, Washakie block of the Wind River Mountains, and western Owl Creek Moun- Garnet tains (Frost et al., 2006a). Garnet from the garnet amphibolite of the Teton Range is dominantly an almandine-pyrope solid solution with minor amounts of spessartine and gros-

Mineral Chemistry sular end members (Alm70–77Prp10–18Sps3–4Grs8–13; Table S11 [footnote 2]). Some garnet is moderately zoned with margins richer in almandine than the cores. Amphibolites with the assemblage Hbl-Pl-Bt-Ilm-Qz ± Grt ± Ttn ± Ep ± Bt ± Cum occur throughout the Teton Range. They can be divided into two groups. One has high Fe/(Fe + Mg) whole-rock and mineral composi- Plagioclase tions and comparatively sodic plagioclase; garnet is present in virtually all of these samples. The second set has lower Fe/(Fe + Mg) ratios and more Plagioclase from the garnet amphibolite is weakly to moderately normally

calcic plagioclase; only a few samples in this category contain garnet (Fitz- zoned and ranges from An36 to An51.

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1.0 900°C peraluminous 180 0.9 leucogranites 140 0.8 ferroan 850°C magnesian 0.7 100 Fe-index

Mt. Owen Zr (ppm) batholith 800°C 0.6 60 750°C 0.5 60 65 70 75 80 20 70 72 74 76 78 80 SiO2 (wt%) SiO2 (wt %) 12 peraluminous Figure 13. Diagram showing the variation of Zr abundance (in ppm) with weight percent SiO2. leucogranites Contours for zircon-saturation temperature are after Watson and Harrison (1983). 8 c alkali

4 alkali-calcic 103 MALI calc-alkalic Mt. Owen batholith Mt. Owen batholith 0 calcic 100

-4 60 65 70 75 80 10 SiO2 (wt%) Makalu leucogranite

Sample/chondrite 1 1.4 peraluminous leucogranites

1.2 0.1 La Ce Pr Nd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

ASI peraluminous Figure 14. Chondrite-normalized rare earth element abundances of the Mount Owen batholith 1.0 metaluminous compared to Himalayan leucogranites. Field for Makalu leucogranites is from Visona and Lom- Mt. Owen bardo (2002). batholith 0.8

60 65 70 75 80 Hornblende

SiO2 (wt%) IV Hornblende contains 1.49–1.83 atoms of Al pfu and has an XMg = 0.444 Figure 12. Geochemical diagrams comparing the compositions of samples from the Mount –0.496. According to the classification of Leake et al. (1997), the hornblende Owen batholith with other peraluminous leucogranites: (A) Fe-index, (B) modified alkali-lime index (MALI), and (C) aluminum saturation index (ASI). Boundaries on the various diagrams and is magnesiohastingite to edenite, but the edenite lies very close to the fields for peraluminous granites are from Frost et al. (2001a). edenite-magnesiohastingite boundary.

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Porphyritic “Augen” Gneiss Southern Thermobarometry Layered Gneiss Garnet Amphibolites—Moose Basin Gneiss and Layered Gneiss Northern 4 Layered Gneiss Two garnet-bearing samples within the prescribed composition range for Moose Basin the barometer yielded 450–600 °C and 4.0 ± 1 kbar (sample 03T5, Moose 3 gneiss Basin gneiss) and 550–600 °C and 5.0 + 1 kbar (sample 07T9, Layered Gneiss; N Fitz-Gerald, 2008). These results are consistent with the presence of titanite 2 Ma c rocks in the assemblage and the absence of garnet in all but the more iron-rich amphibolites in the region. 1 Sample 03T5 contained the assemblage Grt-Hbl-Pl-Cum-Qz-Ilm-Ttn (Fig. 16). A pseudosection calculated for this sample using Perple-X-6.7.7 of Connolly (2009) and the thermodynamic data set of Holland and Powell (1998, 2011) restricts the appearance of cummingtonite to pressure (P) <4.6 -4 -2 0 2 4 Rendezvous kbar and temperature (T ) in the range 540–600 °C (shaded area in Fig. 17). εNd (2685 Ma) These pressures and temperatures are consistent with the Kohn and Spear

Figure 15. Histogram of initial ƐNd (at 2685 Ma) for the southern Layered Gneiss (including (1990) results above, and we conclude that the garnet amphibolites equil- the “Augen” Gneiss), northern Layered Gneiss, Moose Basin gneiss, mafic rocks, and Ren- ibrated at 540 °C < T < 600 °C and P < 5.0 kbar (Fig. 17). Considering that dezvous Gabbro. Data are from Frost et al. (2016a), Swapp et al. (2018), C. Frost et al. (2006a), and this study. titanite was stable in this assemblage, we conclude that the titanite ages of ca. 2.62 Ga that we obtained from the northern and southern portions of the

range formed during this event, which we call M3.

DISCUSSION

Cum Archean Assembly of the Teton Range

The Archean rocks exposed in the northern Teton Range record geologic events that did not affect the southern part, nor are these events preserved elsewhere in the Wyoming Province. The ca. 2.7 Ga, >12 kbar high-pressure granulite metamorphism is the most striking of these features, particularly because granulite-facies metamorphism is rare throughout the Wyoming Province. Small exposures of granulite-facies rocks of approximately this Ttn Hbl age occur locally in the northwestern Wind River Range (Frost et al., 2000; Zrn Chamberlain and Frost, 2005), but they record lower pressures and temperatures (6–8 kbar, 800–900 °C) than the granulites in the Teton Range. The 2685 Ma Webb Canyon and Bitch Creek leucogranitic orthogneisses in the northern domain are other notable features that also have no counterparts, either in the southern domain or elsewhere in the Wyoming Province. These trondhjemitic orthogneisses are interpreted to have formed by water-excess melting and by dehydration melting of amphibolite (Frost et al., 2016a). Both melting mechanisms are expected in collisional environments, a process that appears not to have operated in the southern domain. Finally, the broad,

100 µm open folding event (F3) recorded in the northern Teton Range (Fig. 4A) is not present to the south.

Figure 16. Photomicrograph of hornblende (Hbl), cummingtonite (Cum), and titanite (Ttn) in The southern Teton Range exposes rock units that are not present in the quartz and plagioclase from 03T5. Garnet is present elsewhere in the slide. Zrn—zircon. northern area but that have affinities to rocks elsewhere in the Wyoming

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5.0 Detrital zircon ages from biotite paragneiss of the southern Layered Gneiss Chl - Amp - Pl - Grt - Bt - Ttn - Ilm - Qz also suggest that the northern and southern parts of the Teton Range had distinct crustal histories for most of the Archean. The dominant age populations Bio - Ilm - Qz of ca. 2.63 Ga and ca. 2.66 Ga and sparse grains with older ages in the range

Grt - 2.86–2.84 Ga do not match the ages of any known geologic events recorded - Pl in the northern area. However, these are the ages of voluminous felsic mag-

M3 mp - matism elsewhere in the Wyoming Province, including the 2.63 Ga Louis Lake A

- 4.0 and 2.64–2.71 Ga Bridger batholiths (Frost et al., 1998; Aleinikoff et al., 1989;

ilm - Qz Table S12 [footnote 2]) and the 2.86–2.84 Ga Bighorn batholith (Frost and Fan- -

Opx - Chl

- Bt ning, 2006). The fact that the biotite paragneiss incorporates zircon that could

- Bt - ilm - Qz Grt have been derived from these sources, but none from units in the northern - Pl

Pl area (such as the zircon-bearing Webb Canyon and Bitch Creek leucogranitic

Amp - gneisses), suggests either that the rocks of the northern Teton Range were not P (kbar ) Amp - - exposed to erosion, or that the biotite paragneiss depocenter was not located Chl - Minimum P from Kohn Opx and Spear (1990) where it could receive detritus from the northern domain. 3.0 - Grt - Bt - Ilm - Qz The northern and southern parts of the Teton Range both preserve ca. Pl 2.62 Ga titanite associated with amphibolite-facies metamorphism. As noted

Amp - above, we conclude that these grains grew during M3 (540 °C < T < 600 °C and P < 5.0 kbar), which is well below the closure temperature of titanite (Frost et al., 2001b). The boundary between the northern and southern domains of the

Chl - Cum - Teton Range is marked by the Moran deformation zone. The biotite paragneiss of Paintbrush Canyon lies within the Moran deformation zone and exhibits a 2.0 strong NNE-striking foliation that is characteristic of the deformation zone. The 500550 600650 700 T (oC) older age limit of deformation is well constrained by detrital zircon in both samples 07T3 and 10T4 to be no older than the ages of the youngest detrital Figure 17. Metamorphic conditions for garnet amphibolites from the Teton Range. Light-shaded zircon grains, which are ca. 2.63 Ga. The available ages constrain the juxtaposi- field shows the conditions reported by Fitz-Gerald (2008) using the garnet-hornblende-plagioclase-quartz barometer of Kohn and Spear (1990). Pseudosection is for a garnet amphibolite tion of northern and southern domains to a short interval of ~10 m.y. between from the Moose Basin gneiss (sample 03T5) containing the assemblage Hbl-Grt-Cum-Pl-Qz-Ilm- deposition of the biotite paragneiss and amphibolite metamorphism, which Ttn, where abbreviations are after Whitney and Evans (2010). The heavily shaded triangular we interpret as coincident with collision. region shows the stability field for this assemblage. This plot was calculated assumingP H2O = The pressure, temperature, and age constraints for M3 allow us to put fur- Ptotal. These relations suggest that metamorphism (M3) occurred at temperature (T ) and pressure (P) of 540 °C < T < 600 °C and P < 5.0 kbar. ther constraints on the pressure-temperature-time (P-T-t ) path followed by the

Teton Range (Fig. 18). The M1 and M2 events occurred during the collisional orogeny in the northern Teton Range between 2695 and 2685 Ma. They are Province. These include the 2.69 Ga Rendezvous Gabbro and the 2.80 Ga connected by the dashed arrow because they represent part of a single oro-

“Augen” Gneiss. The Rendezvous Gabbro is younger than the Stillwater Com- genic event (Swapp et al., 2018). The M3 event was associated with the accre-

plex (2709–2710 Ma; Wall and Scoates, 2016; Wall et al., 2016), and it may be tion of the rocks of the northern Teton Range with those of the south. The M3

more closely related to mafic magmatism that is manifest across the Wyoming event is separate from the M1 and M2 events. It occurred as much as 60 million Province as a series of mafic dikes that are the same age within uncertainty: years later. 2689 ± 5 Ma dikes in the Bighorn Mountains (Chamberlain et al., 2010), 2688 ± 6 Ma dikes in the Granite Mountains (Grace et al., 2006), 2683 ± 8 Ma dikes in the western Owl Creek Mountains, and 2681 ± 2 Ma dikes in the Wind River Implications for the Accretion of the Wyoming Province Range (B. Frost et al., 2006). The 2.80 Ga “Augen” Gneiss within the southern Layered Gneiss is much Several crustal fragments of oceanic and continental affinity were accreted older than any rock dated from the northern Teton Range. However, felsic along the southern margin of the Wyoming craton (present-day coordinates) rocks of this age dominate the Archean rocks of the Beartooth uplift, where the between 2.65 and 2.62 Ga, and comprise the southern accreted terranes ~5000 km2 Long Lake magmatic complex was emplaced between 2.83 and 2.79 (Chamberlain et al., 2003, C. Frost et al., 2006b). This accretion occurred prior Ga (Mueller et al., 2010). to the intrusion of undeformed 2.63–2.62 Ga granitic batholiths (C. Frost et al.,

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14 Our data show that the Moran deformation zone occurred at roughly the same time as the later phase of accretion on the southern margin. During this deformation, the rocks in the northern Teton Range, which have a geologic rutile in garnet, record distinct from the rest of the Wyoming Province, were accreted to the GRAIL Wyoming Province. This exotic terrane was accreted to the craton from the 12 M1 2695 Ma west at approximately the same time as accretion occurred along the Wyo- ming province’s southern boundary. The discovery of the 2.62 Ga Moran deformation zone is a major addition 10 to our understanding of the Neoarchean structural development of the Wyo- ming Province. From the precise dates and kinematics of individual accretion events that we have been developing from across the province, it is evident that the ancient core of the Wyoming Province was affected by multiple colli- sions during the span from 2.65 to 2.62 Ga. 8 GRAIL P (kbar) Significance of the Mount Owen Batholith M2 2685 Ma 6 The 2.55 Ga Mount Owen batholith is a typical peraluminous leucogranite: It is a muscovite-bearing, silica-rich granite that is not associated with garnet amphibolites more mafic rocks. The undeformed batholith intrudes the Moran deformation zone and, as such, may be considered a stitching pluton. Peralumi- nous granites are commonly interpreted to have formed by melting of a 4 Ky M3 ~ 2.62 Ga. sedimentary source (Chappell and White, 1974), but they also may form by melting of biotite-bearing metaluminous felsic rocks (Miller, 1985; Frost and Frost, 2011). The Nd isotopic composition of the Mount Owen batholith is And Sil similar to that of the Layered Gneiss. From this, we conclude that the Mount 2 Owen batholith was melted from a pelitic or intermediate to felsic crustal 400 600 800 1000 source with Nd isotopic compositions of the Layered Gneiss. T (°C) The Mount Owen batholith is the largest 2.55 Ga granitic intrusion in the Wyoming Province. Imprecise 2.55 Ga dates reported for granite plutons in Figure 18. Pressure-temperature-time path for gneisses from the Teton Range. M and M are 1 2 the Wind River Range (Stuckless et al., 1985) and Granite Mountains (Lud- recorded only in the northern Tetons. The M1 high-pressure granulite event was dated at 2695 Ma and determined by garnet-rutile-aluminosilicate-ilmenite-quartz (GRAIL) barometry and Zr- wig and Stuckless, 1978) have not been confirmed by more recent U-Pb

in-rutile thermometry to have occurred at >12 kbar and a minimum of 900 °C. The M2 event geochronology, which suggests older ages of 2.62 Ga (Wall, 2004; Bagdo- is marked by partial melting of Layered Gneiss in the presence of staurolite at 2685 Ma. M is 3 nas et al., 2016). However, small volumes of 2.55 Ga granitic rocks have recorded in garnet amphibolites from the northern portion of the range. It is interpreted as an amphibolite-facies metamorphic event accompanying accretion of the northern and southern been identified in the northern Wyoming Province. Mogk et al. (1988) estab- domains of the Teton Range along the Moran deformation zone at ca. 2.62 Ga. Ky—kyanite, lished an age of 2564 ± 13 Ma for an augen gneiss sill in the North Snowy Sil—sillimanite, And—andesite. block of the Beartooth Mountains, and Mueller et al. (2016) identified 2.55 Ga leucogranites emplaced along ductile shear zones in southwest Mon- tana. Mueller et al. (1996) dated a 2.55 Ga hydrothermal zircon from a 3.5 Ga 2006b). Timing of the earliest period of Neoarchean deformation along this trondhjemitic mylonite from the boundary of the Beartooth-Bighorn mag- southern margin was established by syndeformational titanite and metamor- matic zone and the Montana metasedimentary province. phic zircon growths at ca. 2.65 Ga (C. Frost et al., 2006b; Grace et al., 2006). The Foster et al. (2006) also documented 2.55 Ga orthogneisses in the most precise timing of the later phase accretion along the southern Wyoming Grouse Creek block, an area underlain by Archean crust located in south- Province was established by McLaughlin (2016), who interpreted a strongly ern Idaho and northern Utah and Nevada that may have a shared his- foliated 2635.8 ± 1.6 Ma anatectic granite in the western Granite Mountains to tory with the Wyoming Province (Gaschnig et al., 2013). Latest Archean have formed in response to deformation accompanying northwesterly conver- granitic gneiss is exposed in the Albion Mountains, where Strickland et al. gence of the southern accreted terranes against the Wyoming craton. (2011) established a 2532 ± 33 Ma age for the Green Creek orthogneiss, a

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are thanked for hosting our analytical sessions on the sensitive high-resolution ion microprobe– two-mica granite with alkali-feldspar megacrysts. Circa 2.55 Ga zircon reverse geometry (SHRIMP-RG) instruments at Stanford and Australian National University. Lee grains are also present in xenoliths in Snake River Plain basalts (Wolf et Finley-Blasi prepared whole-rock sample powders, and Norbert Swoboda-Colberg performed al., 2005), and as inherited zircon in the Southern Atlanta lobe of the Idaho the X-ray fluorescence analyses. This manuscript was completed while C.D. Frost was serving at the National Science Foundation. K.R. Chamberlain was partially supported from Mega-Grant batholith (Gaschnig et al., 2013). 14.Y26.31.0012 of the government of the Russian Federation. This is publication 66 of the Large It is evident that ca. 2.55 Ga deformation and magmatism are widespread Igneous Provinces–Supercontinent Reconstruction–Resource Exploration Project (CAMIRO Proj- in the western United States, but the lack of outcrop and the overprinting ect 08E) 3, and Natural Sciences and Engineering Research Council of Canada (NSERC) Collab- orative Research and Development publication CRDPJ 419503–11. See www.supercontinent.org by younger deformation and metamorphism make it nearly impossible to and www.camiro.org/exploration/ongoing-projects for more information. We would like to thank characterize the tectonic environment of this event. The best indication for Charles Nye for assistance with the figures. the tectonic origin of this event is the fact that the Mount Owen batholith is undeformed and that we have no obvious indication of metamorphism at 2.55 Ga. It is reasonable to assume that the Mount Owen batholith formed REFERENCES CITED from magmas generated from intermediate to felsic crustal rocks that were Aleinikoff, J.N., Williams, I.S., Compston, W., Stuckless, J.S., and Worl, R.G., 1989, Evidence for an early Archean component in the middle to late Archean gneisses of the Wind River Range, thrust beneath the western Wyoming Province at 2.55 Ga. west-central Wyoming: Conventional and ion microprobe U-Pb data: Contributions to Miner- alogy and Petrology, v. 101, p. 198–206, https://doi.org/10.1007/BF00375306. Bagdonas, D.A., Frost, C.D., and Fanning, C.M., 2016, The origin of extensive Neoarchean high- silica batholiths and the nature of intrusive complements to silicic ignimbrites: Insights from CONCLUSIONS the Wyoming batholith, U.S.A.: The American Mineralogist, v. 101, p. 1332–1347, https://doi .org/10.2138/am-2016-5512. (1) The Teton Range exposes two areas with contrasting Archean histo- Barker, F., 1979, Trondhjemite: Definition, environment, and hypotheses of origin,in Barker, F., ed., Trondhjemites, Dacites, and Related Rocks: Amsterdam, Netherlands, Elsevier, p. 1–12, ries. The northern portion of the range preserves a high-pressure granulite https://doi.org/10.1016/B978-0-444-41765-7.50006-X. event at ca. 2.70 Ga, followed by amphibolite metamorphism and intrusion of Black, L.P., Kamo, S.L., Allen, C.M., Aleinikoff, J.N., Davis, D.W., Korsch, R.J., and Foudoulis, C., leucogranitic gneisses at 2.68 Ga. These features, which have not been docu- 2003, TEMORA 1: A new zircon standard for Phanerozoic U-Pb geochronology: Chemical Geology, v. 200, p. 155–170, https://doi.org/10.1016/S0009-2541(03)00165-7. mented elsewhere in the Wyoming Province, have been interpreted as record- Black, L.P., Kamo, S.L., Allen, C.M., Davis, D.W., Aleinikoff, J.N., Valley, J.W., Mundil, R., Camp- ing a collisional orogeny (Frost et al., 2016a; Swapp et al., 2018). We interpret bell, I.H., Korsch, R.J., Williams, I.S., and Foudoulis, C., 2004, Improved 206Pb/238U micro- these rocks to represent a terrane exotic to the rest of the Wyoming Province. probe geochronology by the monitoring of a trace-element–related matrix effect; SHRIMP, ID-TIMS, ELA-ICP-MS and oxygen isotope documentation for a series of zircon standards: (2) The southern part of the Teton Range contains rock units that differ from Chemical Geology, v. 205, p. 115–140, https://doi.org/10.1016/j.chemgeo.2004.01.003. those in the northern area but that share affinities with mafic and felsic rocks Bohlen, S.R., Wall, V.J., and Boettcher, A.L., 1984, Experimental investigations and geological elsewhere in the Beartooth-Bighorn magmatic zone, suggesting that it should applications of equilibria in the system FeO-TiO2-Al2O3-SiO2-H2O: The American Mineralo- gist, v. 68, p. 1049–1058. be considered a part of the Wyoming Province. Bradley, C.C., 1956, The Pre-Cambrian complex of Grand Teton National Park, Wyoming, in (3) Both northern and southern domains experienced amphibolite-facies Berg, R.R., ed., Jackson Hole: Eleventh Annual Field Conference Guidebook, 1956: Casper, metamorphism at ca. 2.62 Ga. We interpret this to have occurred when the two Wyoming, Wyoming Geological Association, p. 34–42. Cameron, E.M., and Garrels, R.M., 1980, Geochemical compositions of some Precambrian shales domains were juxtaposed along the Moran deformation zone. This Archean from the Canadian Shield: Chemical Geology, v. 28, p. 181–197, https://doi.org/10.1016/0009- assembly of the Teton Range was approximately coeval with the last phase of 2541(80)90046-7. accretion of the southern accreted terranes along the southern margin of the Cardozo, N.C., and Allmendinger, R.W., 2013, Spherical projections with OSXStereonet: Comput- ers & Geosciences, v. 51, p. 193–205, https://doi.org/10.1016/j.cageo.2012.07.021. Wyoming Province. The Moran deformation zone is the only exposed Neoar- Chamberlain, K.R., and Frost, B.R., 2005, Granulite metamorphism in the northeastern Wind chean terrane boundary in the Wyoming Province. River Range at 2705 Ma: Evidence for crustal underplating during the Stillwater thermal (4) The intrusion of the 2.55 Ga Mount Owen batholith was the latest igne- event and plume-style crustal growth: Geological Society of America Abstracts with Pro- grams, v. 37, no. 7, p. 505. ous event in the Teton Range and the last major Archean event in the Wyoming Chamberlain, K.R., Frost, C.D., and Frost, B.R., 2003, Early Archean to Mesoproterozoic evolution Province. The next youngest event, the intrusion of the voluminous peralumi- of the Wyoming Province: Archean origins to modern lithospheric architecture: Canadian nous leucogranites of the Wyoming batholith on central Wyoming, occurred Journal of Earth Sciences, v. 40, p. 1357–1374, https://doi.org/10.1139/e03-054. Chamberlain, K.R., Schmitt, A.K., Swapp, S.M., Harrison, T.M., Swoboda-Colberg, N., Bleeker, 65 m.y. earlier. The tectonic setting for the crustal melting that made the Mount W., Peterson, T.D., Jefferson, C.W., and Khudoley, A.K., 2010, In situ U/Pb SIMS (IN-SIMS) Owen batholith is cryptic. micro-baddeleyite dating of mafic rocks: Method with examples: Precambrian Research, v. 183, p. 379–387, https://doi.org/10.1016/j.precamres.2010.05.004. Chappell, B.W., and White, A.J.R., 1974, Two contrasting granite types: Pacific Geology, v. 8, p. 173–174. Connolly, J.A.D., 2009, The geodynamic equation of state: What and how: Geochemistry Geophys- ACKNOWLEDGMENTS ics Geosystems, v. 10, Q10014, https://doi.org/10.1029/2009GC002540. We acknowledge financial support provided by National Science Foundation (NSF) grant EAR- Dale, J., Holland, T., and Powell, R., 2000, Hornblende-garnet-plagioclase thermobarometry: 0537670 to B.R. Frost, C.D. Frost, and S.M. Swapp and by University of Wyoming–National Park A natural assemblage calibration of the thermodynamics of hornblende: Contributions to Service Research Center grant 49182 to B.R. Frost and C.D. Frost. Joe Wooden and Mark Fanning Mineralogy and Petrology, v. 140, p. 353–362, https://doi.org/10.1007/s004100000187.

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Fitz-Gerald, D.B., 2008, Evidence for an Archean Himalayan-Style Orogenic Event in the Northern Association Commission on new minerals and mineral names: Mineralogical Magazine, v. 61, Teton Range, Wyoming [M.Sc. thesis]: Laramie, Wyoming, University of Wyoming, 159 p. p. 295–321, https://doi.org/10.1180/minmag.1997.061.405.13. Foster, D.A., Mueller, P.A., Mogk, D.W., Wooden, J.L., and Vogl, J.J., 2006, Proterozoic evolution of Love, J.D., Reed, J.C., Jr., and Christiansen, A.C., 1992, Geologic Map of Grand Teton National Park, the western margin of the Wyoming craton: Implications for the tectonic and magmatic evolu- Teton County: Reston, Virginia, U.S. Geological Survey, scale 1:62,500. tion of the northern Rocky Mountains: Canadian Journal of Earth Sciences, v. 43, p. 1601–1619, Ludwig, K.R., 1991, ISOPLOT for MS-DOS, a Plotting and Regression Program for Radiogenic- https://doi.org/10.1139/e06-052. 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