Many porphyry molybdenum deposits are hosted in multi-phase plutons, but it is unclear in some deposits how these magmas originated and whether the pluton intruded as it fractionated or was intruded by new batches of magma. New mapping has clarified field relationships between units in the White Pine porphyry Mo system hosted in the Little Cottonwood stock, Utah (western United States), including the White Pine intrusion, the Red Pine porphyry, rhyolite dikes, and phreatomagmatic pebble dikes. Geologic relations and geochemistry show the system formed in a continental arc setting during rollback of the subducting Farallon slab rather than during extension related to orogenic collapse. Whole-rock geochemistry shows distinct fractionation trends for each of the major intrusive units in the composite pluton, suggesting they formed separately, which is supported by new U-Pb zircon laser ablation–inductively coupled plasma–mass spectrometry (LA-ICP-MS) ages of ca. 30 Ma for the Little Cottonwood stock, 27 Ma for the White Pine intrusion, and 26 Ma for the previously undated Red Pine porphyry. Mineral textures, cross-cutting relationships, and alteration mineralogy indicate that intrusion of the youngest phase led to a fluid-saturated magmatic system and triggered venting, including emplacement of pebble dikes. In the adjacent east Traverse Mountains, pebble dikes contain clasts that have similar mineral assemblages, textures, and ages as the major igneous units in the White Pine deposit. This indicates that the pebble dikes in east Traverse Mountains and in the pluton are the upper and lower parts of the same magmatic-hydrothermal system, which was decapitated by a mega-landslide that was likely facilitated by alteration in the Oligocene hydrothermal system and by later Basin and Range faulting.

The White Pine porphyry Mo-W deposit is hosted in the Oligocene Little Cottonwood stock of north-central Utah (western United States; Figs. 1 and 2), a product of multi-phase pluton construction. Sharp (1958) and Crittenden (1965) described the White Pine intrusive phase in the northeastern part of the main pluton as a more leucocratic phase, which is itself intruded by the chemically evolved Red Pine porphyry and several rhyolite dikes. The sub economic low-grade deposit is centered on the Red Pine intrusion. There remains some debate as to whether the sequence represents inward differentiation of the Little Cotton-wood granodiorite (Lawton, 1980; Hanson, 1995; Marsh and Smith, 1997; Smyk et al., 2018) or the White Pine and Red Pine intrusions crystallized from separate magmas that intruded the Little Cottonwood stock (Sharp, 1958; Crittenden, 1965; Krahulec, 2015). Other questions about the pluton regard the tectonic setting in which the magma(s) originated, whether the pluton experienced episodes of venting, and the source of the mineralizing fluids. The effect of those hydrothermal fluids on the surrounding rocks is also of interest, particularly regarding subsequent large-scale slope failure and mass movement.

Indeed, the Little Cottonwood stock has evidence bearing on the existence of a megascale landslide adjacent to the pluton. It has recently been proposed that the east Traverse Mountains (Figs. 1 and 2), previously mapped as essentially a horst block oriented perpendicular to the general north-south trend of the Wasatch Range (Biek, 2005), are actually a rootless landslide block (Keith et al., 2017; Chadburn, 2020). This would be among the largest known sub aerial landslides in the world. Some of the evidence for the megalandslide are pebble dikes in the White Pine deposit that resemble those in the east Traverse Mountains. Moreover, alteration related to the dikes may have provided an impetus for the failure.

Addressing questions about the nature of the composite Little Cotton-wood stock and testing the hypothesis of the related east Traverse Mountains landslide required field mapping and detailed comparisons of the petrology, petrography, and radiometric ages of the various intrusive units, including pebble dikes in the mineralized rocks and in the east Traverse Mountains. Based on those results, we outline a sequence of events for the evolution of the Little Cottonwood, White Pine, and Red Pine magmas, the formation of the Mo-W ore deposit and associated hydrothermal alteration, possible volcanic activity, and the production of pebble dikes, all central to the landslide question.

The Little Cottonwood stock (Figs. 1 and 2) is part of the Wasatch igneous belt, an east-west trend of 11 Cenozoic intrusions and a volcanic field in central Utah, all formed during the mid-Cenozoic magmatic flareup that occurred in the western United States. The igneous belt is aligned within the Laramideage Uinta arch along the trend of the much larger Uinta-Cortez axis (Vogel et al., 2001). The deformation associated with the Laramide orogenic uplift of the Uinta arch probably started in the Late Cretaceous (Dickinson et al., 1988; Yonkee and Weil, 2015). Differential uplift continued episodically into the Oligocene. For example, Jensen et al. (2020) concluded that there were two late episodes of relative uplift in the Uinta Mountains—one at ca. 39 Ma to form the clastic sediment of the Duchesne River Formation and another at ca. 34–30 Ma to fold the Duchesne River Formation and form the Bishop Conglomerate. Thus, magmatism in the Wasatch igneous belt, which started ca. 36 Ma and continued until ca. 26 Ma, was partially contemporaneous with this contractional deformation, limiting the possibility for association with regional extension (cf. Vogel et al., 2001). Some workers, including Presnell (1998) and Vogel et al. (2001), proposed that the east-west–trending Uinta-Cortez axis formed along the suture zone between the Archean Wyoming crustal province to the north and Paleoproterozoic accreted crustal terranes on the south (Fig. 1). According to Vogel et al. (2001), the Wasatch igneous belt intruded this structural weakness. Nelson et al. (2011) disputed this, arguing that ages and isotopic compositions of younger granitic and basement rocks of various ages place the suture zone between the two crustal provinces north of Utah. This casts doubt on the hypothesis that the Wasatch igneous belt marks a basement suture zone and places the entire study area in Paleoproterozoic accreted arc terranes.

The rocks of the Wasatch igneous belt generally decrease in age (from 36 Ma to 26 Ma), increase in silica content, and increase in emplacement depth from east to west. The Little Cottonwood stock is the westernmost, largest, youngest, most silicic, and most deeply emplaced of the 11 intrusions (John, 1989; Hanson, 1995). It is compositionally and texturally zoned from a finer-grained less-silicic outer and upper zone to the east to a coarser-grained, more-silicic, structurally lower zone to the west (Marsh and Smith, 1997). It ranges modally from granodiorite to quartz monzonite to granite. Adjacent to the Little Cottonwood stock to the east is the slightly more mafic Alta stock.

The Little Cottonwood stock is a composite pluton. Subsidiary intrusive units include aplite, rhyolite, and lamprophyre dikes and two small intrusions: the White Pine intrusion and the Red Pine porphyry (Fig. 3). The White Pine and Red Pine units host the small White Pine Mo-W deposit, which is surrounded by widespread quartz-sericite-pyrite (QSP) alteration. It also hosts a mineralized breccia pipe and a pebble dike. The deposit has a central stockwork of veins with quartz ± sericite, pyrite, scheelite, and molybdenite as well as veins of ribbon quartz and vug-filling quartz. Toward the southern margin of the White Pine outcrop area, there is a breccia pipe that is Mo mineralized in some locations (Sharp, 1958; John, 1989). It also hosts a pebble dike—a phreatomagmatic feature formed when meteoric and/or magmatic water in low-permeability rocks is superheated and flash-boiled, causing brecciation and the opening of fractures, in some cases all the way to the surface. The hot fluid entrains clasts of brecciated wall rock as it vents upward, and the slurry is emplaced as dikes or pipes in the overlying rock column. Clasts may be rounded via abrasion or thermal spallation (Sillitoe, 1985; Johnson, 2014). Pebble dikes are a common feature in porphyry-type deposits (Bryner, 1961; Mutschler et al., 1981; Sillitoe, 1985).

The plutons of the Wasatch igneous belt intruded cratonic rocks overlain by Neoproterozoic and Paleozoic sedimentary rocks of the allochthonous Charleston-Nebo thrust sheet (John, 1989). These rocks were deposited in a thick miogeoclinal wedge on the rifted western margin of Laurentia (Whitmeyer and Karlstrom, 2007; Yonkee et al., 2014). They were then thrust and stacked eastward in the Sevier orogeny. In the Wasatch Range, these allochthonous sedimentary rocks include thick sections of the Pennsylvanian–Permian Oquirrh Group (Crittenden, 1965; Bryant, 1990). Adjacent to the Little Cottonwood stock are the east Traverse Mountains, a highly brecciated and faulted block made up of Oquirrh Group strata, Paleogene conglomerates, and argillically and propylitically altered volcanic rocks (Biek, 2005; Hoopes et al., 2014). The east Traverse Mountains block also hosts pebble dikes, which contain both intrusive igneous and sedimentary clasts. Separating the Little Cottonwood stock and east Traverse Mountains is the Wasatch fault, a west-dipping regional normal fault system that developed as part of late Cenozoic Basin and Range crustal extension. As much as 11 km of offset along the Wasatch normal fault since 12 Ma, at a slip rate of 0.6–0.7 mm/yr, caused the Little Cottonwood stock to tilt ~20° to the east (John, 1989; Kowallis et al., 1990; Bruhn et al., 2005). This gives rise to incongruous depths of crystallization across the exposed surface of the Little Cottonwood stock: fluid inclusion data give a depth of 6 km on the eastern margin and 11 km on the western margin (John, 1989). Across the valley, 25 km to the west of the Wasatch igneous belt, is the world-class Bingham Canyon porphyry Cu-Au-Mo deposit, which is intermediate in composition and roughly 12 m.y. older than the White Pine deposit.

To better understand the relationship of the White Pine Mo-W deposit, pebble dikes in the east Traverse Mountains, and the Little Cottonwood stock, we conducted new geologic mapping, collected a suite of 80 samples, and examined them using optical microscopy, X-ray fluorescence analysis, and laser ablation U-Pb geochronology of zircon. For analytical details, see text of Item S4 in the Supplemental Material1.

The major geologic units of the composite pluton are described below using field, petrographic, and chemical characteristics, integrating our new findings with those of previous studies.

Little Cottonwood Stock

The Little Cottonwood stock is an Oligocene (Figs. 4 and 5, Table 1) felsic granitoid with an exposed area of 115 km2 (Figs. 1 and 2). It is high-K, magnesian, and metaluminous (Marsh and Smith, 1997). When plotted on a total alkali–silica diagram (Middlemost, 1994), the compositions cluster around the triple point between granite, granodiorite, and quartz monzonite (Fig. 6A). In the International Union of Geological Sciences (IUGS) quartz–alkali feldspar–plagioclase diagram (Le Maitre, 1989), the rocks of the Little Cottonwood stock are primarily granodiorite with less quartz monzonite and granite (Fig. 7). They generally range from coarsely phaneritic-porphyritic with an equigranular groundmass in the inner and structurally lower portion of the stock to the west to having a finer-grained phaneritic to seriate groundmass in the outer and structurally upper portion of the stock to the east (Marsh and Smith, 1997).

In most locations, K-feldspar megacrysts as much as 6 cm long are present, but locally they are as much as 12 cm long. Most of these megacrysts have exsolved albite present as lamellae or as small irregular blebs which are optically continuous in some grains. The matrix crystals are dominantly plagioclase and commonly strained quartz, with lesser amounts of K-feldspar. Mafic minerals are primarily biotite, usually >5% in modal abundance (Fig. 8F; Item S1, footnote 1), with less euhedral amphibole. Accessory phases include ~1.5% titanite (Henze, 2020) as well as allanite, magnetite, apatite, and zircon. Ilmenite occurs only in the cores of some titanite grains where textures show it was partially consumed as titanite grew. There is no significant mineralization, but thin sections of the Little Cottonwood stock show even the freshest samples have minor chloritization of biotite and hydrothermal epidote. The stock is more altered along fractures near the White Pine intrusion.

The stock contains ellipsoidal mafic enclaves, which generally range from a few centimeters to nearly a meter across in the northeast part of the pluton but can be as much as several meters across and are commonly associated with mafic schlieren (Hanson, 1995). They have the same mineral assemblage as the host rock but with different modal abundances and more-mafic bulk compositions (Hanson, 1995; Henze, 2020). The enclaves are porphyritic-aphanitic; most enclaves are more than half poikilitic plagioclase, followed by biotite, hornblende, and small amounts of quartz and K-feldspar. Commonly, hornblende has cores of biotite. Accessory minerals include as much as 3% titanite (Item S1, footnote 1), grains of which are large and commonly euhedral, but some grains have small anhedral plagioclase inclusions, and some grains appear to have minor resorption. Other accessory phases include lesser amounts of apatite, Fe-Ti oxides, and zircon. In one sample, sprays of acicular apatite radiate from central points. Alteration phases include chlorite, epidote, and calcite. Henze (2020) suggested an origin of the enclaves from injection of a more-mafic magma followed by incomplete mixing.

White Pine Intrusion

The leucocratic White Pine phase intrudes the northeastern part of the Little Cottonwood stock (Figs. 2 and 3). The contacts between the White Pine phase and the Little Cottonwood granodiorite are difficult to find, both because of the similarity in appearance of the two phases in hand sample and because of widespread cover and pervasive hydrothermal alteration. Sharp (1958) described the contacts as ranging from sharp to gradational. The only location where we observed a sharp contact was a small region on the southern end of the intrusion. Our mapping (Figs. 2 and 3) has extended the range of the White Pine outside that mapped by Sharp (1958) and Critten den (1965) and is intended to encompass the area intruded by possibly several small bodies.

The White Pine phase is texturally and mineralogically similar to the fine-grained facies of the Little Cottonwood stock but is slightly more leucocratic (Item S1, footnote 1), with a medium-to fine-grained phaneritic or seriate groundmass and in some instances poikilitic K-feldspar megacrysts (Figs. 8A8B). The matrix is dominantly plagioclase with lesser amounts of quartz and K-feldspar (Fig. 7). Quartz tends to be larger than plagioclase (as large as 1 cm) and in some samples is square and stands out in prominent relief. The only mafic phase is biotite, and this is usually <5% in modal abundance; unlike in the Little Cottonwood stock, hornblende is lacking. It also differs from the Little Cottonwood stock in having only trace amounts of titanite, which is usually not visible in hand sample. Other igneous accessory mineral phases include apatite, allanite, zircon, and Fe-Ti oxides. Significantly, the White Pine intrusion consistently lacks mafic enclaves, in contrast to the encompassing Little Cottonwood stock.

The White Pine intrusion in most places is moderately to heavily altered with quartz + sericite + pyrite (QSP). In outcrop, it almost universally weathers orange from limonite and jarosite formed by the breakdown of alteration sulfides, which permeate fractures in the rock, especially in the stockwork of mineralized fractures that mark the center of the phase. It also has ribbon quartz in veins and coarse crystals of vug-filling quartz. The fractures are lined with pyrite and sericite—which are also disseminated through the rock—and with quartz, and closer to a mineralized breccia pipe, they also host molybdenite. In thin section, many White Pine rocks have a fine-grained interstitial groundmass of quartz ± K-feldspar, as much as ~15% in modal abundance (Item S1, footnote 1; Fig. 8B). It is not clear whether the groundmass is magmatic or due to post-magmatic silicification or both. Most plagioclase is heavily altered to sericite ± clays, especially the cores, and biotite is commonly partially to completely altered to sericite and/or chlorite with inclusions of rutile. K-feldspar generally has only minor clay alteration, even where biotite and plagioclase are completely replaced, although smaller K-feldspar grains appear more readily altered than megacrystic K-feldspar. Sericite, in addition to replacing plagioclase and biotite, is present as euhedral, tabular blades or fans. In addition to the QSP alteration assemblage, hydrothermal epidote, monazite, barite, rutile, and uranopolycrase [(U,Th)(Ti,Ta)2O6] were identified via electron microprobe. Oxides from late-stage weathering are present.

The White Pine intrusion is Mo mineralized in a few small occurrences, including a breccia pipe and near the west ridge (Fig. 3), as revealed in adits. It is also cut by a distinctive pebble dike. The dike and pipe are described below. In the mineralized rock, molybdenite primarily occurs in a disseminated manner, as well as in small veinlets and rarely in thick quartz veins that cut the stockwork and breccia zones (Sharp, 1958). This mineralized White Pine host rock is distinct in appearance, texture, and mineralogy from the unmineralized White Pine rocks and is yellow in outcrop. The mineral assemblage is simply quartz, very euhedral K-feldspar, and sericite, with trace amounts of plagioclase and hydrothermal pyrite (commonly coarse grained), barite, and monazite. It lacks any mafic minerals or phenocrystic plagioclase. Texturally, it is seriate to phaneritic, though the original magmatic texture has likely been destroyed (Fig. 8C). It also hosts veinlets of alteration clay and oxides as well as quartz veins and veinlets.

Red Pine Porphyry

The Red Pine porphyry, first recognized by Sharp (1958), is found in a few small, scattered outcrops and in float (Fig. 3). A strongly porphyritic variety is most distinctive, but it ranges from aplitic to seriate in texture as indicated by dated samples. Alteration and heavy cover make it difficult to find contacts with the White Pine intrusion, and consequently, the boundaries of the Red Pine porphyry are largely inferred. Widely spaced small outcrops of Red Pine rocks occur on the east ridge of the valley and on the west ridge separating White Pine and Red Pine Canyons and in one outcrop on the northern end of the White Pine deposit (Fig. 3). We concur with Sharp (1958) that the Red Pine porphyry consists of multiple irregular small bodies intruding the White Pine phase.

Although complicated by the overprint of hydrothermal alteration, most Red Pine compositional trends are indistinguishable from those of the White Pine intrusion (Figs. 6 and 9). Thus, textural variations of the Red Pine porphyry may be the best way to distinguish it from other units in the absence of U-Pb ages. The porphyritic variety of the Red Pine is dominantly a quartz-bearing porphyry. It has ~50% fine-grained white to tan altered groundmass with prominent phenocrysts of square and commonly resorbed and/or embayed quartz grains as much as 5 mm across (Fig. 8), plagioclase (some with rounded edges, but otherwise not resorbed), K-feldspar (locally poikilitic and megacrystic), and relict biotite. Alteration is intense; there are no surviving mafic phases because biotite is completely altered to sericite and Fe-Ti oxides, especially rutile, ± chlorite. Phenocrystic plagioclase is almost completely altered to clays and sericite. Pockets of clay minerals are common. Apatite, allanite, zircon, and titanite with a distinctive composition (Martin et al., 2021) are present. Cubes of hydrothermal pyrite, commonly weathered out to leave cubic limonite-rimmed void spaces, are prevalent in altered samples.

The phaneritic variety of the Red Pine porphyry bears little textural resem-blance to its porphyritic variety (Fig. 8E), but our new zircon U-Pb ages show they are part of the same intrusive phase. It is seriate, generally finer grained than the White Pine intrusion or Little Cottonwood stock, and its grains are mostly subhedral to anhedral. It lacks clots of fine-grained groundmass. Phenocrysts include small plagioclase and biotite grains and sparse larger phenocrysts of quartz and K-feldspar. Accessory minerals include apatite, zircon, titanite, allanite, and oxides. It is rather fresh, with minor secondary pyrite, sericite, epidote, and chlorite. Hand samples are light gray to white in color.

Hydrothermal Breccias

The White Pine deposit and the east Traverse Mountains host two different types of hydrothermal breccias. Pebble dikes, which are common in porphyry-type deposits (Bryner, 1961; Mutschler et al., 1981; Sillitoe, 1985), are found in both locations. The White Pine phase also hosts a pipe of brecciated and mineralized granite.

Pebble Dikes in the White Pine Phase

A narrow (~0.5 m) east-west–striking, near-vertical pebble dike cuts the White Pine porphyry near the Mo-mineralized breccia pipe (Fig. 3). The White Pine granite that hosts the pebble dike is more heavily QSP altered than surrounding rock but not Mo mineralized. The matrix of the dike is brownish to grayish green and comprises broken phenoclasts of biotite, quartz, plagioclase, K-feldspar, and finely comminuted material as well as various alteration phases including clays, sericite, finely disseminated pyrite, and calcite—similar to the QSP alteration assemblage in some of the clasts. Rounded to subrounded intrusive igneous clasts (<15 cm in diameter, most commonly ~1–3 cm in diameter, but some as small as a few millimeters) make up 25%–50% of the dike. In some parts of the dike, the clasts are elongated and have a preferred orientation. Based on texture and mineral assemblage, Sharp (1958) identified both White Pine and Little Cottonwood stock clasts. However, using thin sections, we found clasts of White Pine and Red Pine in this pebble dike (Figs. 8 and 10). None of the clasts could be positively identified as Little Cottonwood stock; if present, such clasts are a minor constituent of the clast population. Smaller (5–10 mm) quartz clasts are included as well, with an abundance of ~5%–10% of the rock. These are apparently mostly phenoclasts from brecciated intrusive units or hydrothermal vein quartz. The level of alteration of clasts in this pebble dike varies widely, based on the sericitic alteration of plagioclase and biotite and chloritic alteration of biotite; pyrite is present as larger rounded grains in some clasts and smaller, more numerous grains in the matrix, and some clasts have a thin rind of very finegrained alteration minerals, probably clays. The pebble dike includes both fresh and altered clasts and is itself somewhat altered.

White Pine Breccia Pipe

Other magmatic-hydrothermal features that may occur in porphyry deposits in addition to pebble dikes are breccia pipes (Thomas and Galey, 1982; Taylor et al., 2012). The White Pine intrusion contains one such breccia pipe (Sharp, 1958; Krahulec, 2015; Smyk et al., 2018), which is possibly part of the same phreatomagmatic system as the pebble dike. It is exposed <100 m north of the pebble dike. It forms the focus of Mo mineralization in the deposit and is roughly oval and ~200 m across. Its peripheral part contains angular fragments of the White Pine phase <16 cm in diameter enclosed in a highly sericitized matrix with molybdenite that is locally cemented by pyrite- and molybdenite-bearing vuggy quartz (Smyk et al., 2018). The molybdenite concentration is higher near the margins of the breccia and occurs as aggregates around clast margins (Krahulec, 2015) and as veinlets in the highly altered White Pine clasts. Coarse-grained “bull” quartz locally envelopes the breccia zone, marking its boundary with a zone of weakly mineralized quartz stockwork veins. The breccia and the stockwork zones include QSP alteration along with molybdenite, galena, sphalerite, and minor scheelite. Minor fluorite-sericite zones cross-cut earlier mineralized veins and veinlets (Sharp, 1958).

Pebble Dikes in the East Traverse Mountains

Numerous pebble dikes also cut Oligocene volcanic rocks at the crest of the east Traverse Mountains ~17 km away (Fig. 2) (Hoopes et al., 2014). These are quite similar to those in the White Pine intrusion. Although some dikes are visible in natural exposures and recent trenches, evidence of pebble dikes is generally found as float. The few outcrops that are visible range from 30 to 50 cm in width, mostly ~30 cm. The sparse nature of the float and recent housing development of the region makes it difficult to reconstruct dike length and orientations. They generally trend northwest-southeast and can only be confidently traced for a few meters to ~100 m but could have much greater lengths. There is no consistent evidence for wall-rock alteration caused by the dikes. The dikes cut pervasively altered block-and-ash-flow deposits of Eocene age (37–35 Ma), which are part of the “volcanic rocks of the east Traverse Mountains” of Biek (2005). These volcanic rocks are intermediate in composition, ranging from andesite to dacite or trachyte. The alteration is dominated by wide-ranging argillization and silicification with the production of jasperoid and opal.

The matrix of the pebble dikes is invariably very fine grained and ranges from light gray to tan to black. The matrix includes broken phenoclasts of quartz, plagioclase, K-feldspar, and biotite along with grains of pyrite, all in a groundmass of finely comminuted material. Clasts are generally angular but include some well-rounded quartzite pebbles. They vary in size from <0.5 cm to ~10 cm in diameter but are most commonly 0.5–2 cm and can make up 10%– 80% of the rock but generally make up 40%–50% of the rock. The clasts include quartzite, extrusive intermediate-composition rocks, pockets of hydrated clay, possible volcaniclastic sedimentary rocks, and, significantly, felsic clasts of coarse-grained intrusive igneous rock. The origin of the clay pockets is not clear; they could be bleached shales from the underlying Manning Canyon Formation or highly argillically altered igneous clasts. In most dikes, quartzite is the most common clast type, while in other samples, extrusive volcanic clasts are the most common type. Some pebble dikes have intrusive and sedimentary clasts, while others only have sedimentary (Fig. 10). Even within one thin section, the intrusive igneous clasts can be of varying rock type. Among the lithologies present in the intrusive clasts are felsic granitoids composed of quartz, plagioclase, biotite, and K-feldspar; felsic granitoids composed of quartz, plagioclase, biotite, K-feldspar, and amphibole; and porphyritic granitoids composed of plagioclase, quartz, biotite, and fine-grained groundmass. The intrusive clasts vary between nearly fresh to heavily QSP altered and in some cases have alteration selvages around their rims. Based on texture and mineral assemblage, we have correlated the intrusive clasts in these dikes with all three of the intrusive phases mapped 17 km to the east in the White Pine deposit, and U-Pb zircon ages correlate one sample with the Little Cotton wood stock.

Alteration of the pebble dikes in hand sample is evident from the brownishred to yellowish alteration staining on the surfaces of some pebble dikes as well as rimming many of the clasts. This is most likely jarosite and iron oxides after hydrothermal pyrite, and other alteration phases from argillic alteration. Argillic alteration is especially clear in the altered volcanic clasts and as yellowish clays on some faces. Fresh rock fractures reveal unoxidized pyrite. Some of the pebble dikes are intensely silicified, with silica content >90 wt%. They break in a conchoidal manner and host multiple generations of thin silica veins. It is unclear to what extent alteration of the clasts in the dikes occurred post-emplacement; silicification affected many dikes post-emplacement, and some of the dikes have evidence of QSP alteration, while others show little to no indications of post-emplacement alteration or silicification besides possibly unrelated late-stage silica veins. The quartz, sericite, and pyrite in the matrix may have been derived from previously altered intrusive rocks pulverized by pebble-dike emplacement. Some pebble dikes lack QSP alteration and appear fresh in hand sample and in thin section. Accessory and alteration phases in pebble-dike mineral separates identified via scanning electron microscope analysis include molybdenite, barite, uranothorite, arsenopyrite, scheelite, and titanite.

The sedimentary rock clasts are similar to quartzites of the Oquirrh Group that make up the east Traverse Mountains (Biek, 2005), consisting of light brown to tan fine-grained quartzite or sandstone and black chert. Thick sections of these late Paleozoic units are found on the flanks of the Little Cottonwood intrusion and are inferred to underlie the volcanic rock hosts on east Traverse Mountains (Figs. 2 and 3). There appear to be multiple generations of pebble dikes in east Traverse Mountains; in some hand samples, there is a sharp contact between two dikes with different-colored matrix and different-sized clasts.

Rhyolite Dikes

Rhyolite dikes (which Sharp [1958] called “aplite dikes”) are white, off-white, light gray, or tan. They are generally oriented east-west, usually less than a meter wide, and visible mostly as float. In the southern, western, and northeastern margins of the White Pine intrusion, they occur in outcrop, mostly 1–2 m in width, and one rhyolite dike in the northeast part of the deposit is ~10 m in width. The homogeneous aphanitic matrix is made up of quartz, K-feldspar, plagioclase, and fine-grained secondary sericite and clays and comprises 95%–99% of the rock. Square quartz phenocrysts, as much as 0.5 cm in size, appear gray in comparison to the groundmass. These rocks have moderate to strong QSP alteration, as shown by the high silica content, with some having >80% SiO2 (Item S2, footnote 1). A few dikes are cut by discrete quartz veins. In addition to fine-grained quartz and quartz veins, alteration phases usually include small, disseminated pyrite cubes commonly 1 mm in size but as much as 0.5 cm in diameter, mostly weathered away to leave squarish oxide-lined void spaces, and sericite flakes <1 mm in size that permeate the matrix and replace minor relict feldspar phenocrysts. The accessory minerals present in the other units (apatite, titanite, and zircon) have not been found in the rhyolite, even in mineral separates.

The interpretation of laser ablation–inductively coupled plasma–mass spectrometry (LA-ICP-MS) isotopic data for zircons from plutonic rocks is complicated by a number of factors. These include slow cooling of plutonic bodies; their typically composite, multi-stage histories with recycling of zircon from older rocks; the preservation of antecrysts formed in long-lived mush; late hydrothermal alteration; and the relatively low precision of the laser ablation technique (±2%). However, laser ablation analyses of zircon also offer several advantages because small spots (25–50 µm across) can be analyzed and even higher-resolution depth profiles can reveal evidence for events that otherwise might go undetected. Here we leverage the spatial resolution offered by LA-ICP-MS analyses to understand the history of the multi-stage Little Cottonwood pluton.

The U-Pb isotopic data and inferred zircon ages for 12 samples are presented in Figures 4 and 5 and summarized in Table 1. All samples have discordant analyses, a large range of ages, complex age spectra, and high mean standard weighted deviations (MSWDs) indicating that the span of ages exceeds that expected from analytical uncertainty alone and that multiple events are represented by the various zircon populations. Discordant ages were identified as those with a log ratio distance from concordia (before common Pb correction) of >10%; they were rejected from further analyses (even though in the vast majority of cases the discordant 238U/206Pb ages were within the span of the concordant ages). Weighted mean ages and kernel density estimate (KDE) plots were constructed only using the concordant 238U/206Pb ages, which were corrected for common Pb using Stacey and Kramers (1975) lead isotope growth curves and the approximate age. All of the results (concordant and discordant) are reported in Item S3 (footnote 1) and summarized in Table 1.

Below, we describe the details for the samples examined and group them by plutonic event.

Little Cottonwood Stock

The three studied samples of the Little Cottonwood stock (CH-072, CH-073, and CH-074) come from the eastern part of the intrusion in an area intruded by younger “fingers” of the White Pine and Red Pine phases (Fig. 3). Most of the U-Pb zircon ages are concordant (Figs. 4A4F). On this basis, 32 of 43, 33 of 37, and 33 of 37 ages, respectively, were used (Table 1). The weighted mean ages for the concordant 238U/206Pb ages are similar between the three samples—29.3 ± 0.5 Ma (MSWD = 23) for sample CH-072, 29.3 ± 0.3 Ma (MSWD = 17) for CH-073, and 29.6 ± 0.4 Ma (MSWD = 34) for CH-074. These ages are indistinguishable from other published U-Pb ages on zircon from the Little Cottonwood stock (e.g., Stearns et al., 2020; Smyk et al., 2018; Vogel et al., 2001). However, the wide range of concordant U-Pb ages (26.6–33.9 Ma) and the high MSWDs suggest that there is variation in the ages beyond that which can be accounted for by analytical variation alone.

To tease out the cause of that variation, we used the mixing models in the DensityPlotter program (Vermeesch, 2012) to identify separate populations of similar age using KDE plots. These Little Cottonwood granodiorites have some of the simplest age spectra of any samples we analyzed from the pluton, with strong peaks between 28.9 and 30.4 Ma (Fig. 5) demonstrating, as expected, ages similar to the weighted mean ages (Table 1). We interpret the peaks to represent the age of zircon that formed during the crystallization of the Little Cottonwood granodiorite. The KDE curves also have low peaks or shoulders <28 Ma that are probably the result of recrystallization of zircon caused by younger intrusions, in this case the White Pine granite (Fig. 5) or the Red Pine porphyry in the case of sample CH-072 which has three grains that form a peak at ca. 26.6 Ma. Each sample also has zircon ages that are >31 Ma—six of 117 total. Stearns et al. (2020) reported U-Pb titanite ages from the western part of the stock that are as old as 36 Ma. This is also approximately the age of the nearby Alta stock (Figs. 1 and 5) and could imply assimilation of rock from that intrusion.

White Pine Phase

Like the Little Cottonwood stock samples, the five White Pine samples (CH-041, CH-127, CH-160, CH-108, and CH-031) have wide U-Pb age ranges, multiple discordant dates, and complex age spectra (Fig. 5), but all have prominent modes on KDE plots of concordant ages at ca. 27 Ma (26.5–27.3 Ma). Samples CH-031 and CH-041 also have younger zircon ages (Figs. 4 and 5) that probably represent the effects of the subsequent intrusion of the still younger (ca. 26 Ma) mineralizing pluton, the Red Pine porphyry. Ages older than the nominal 27 Ma age of the White Pine intrusion are more obvious and show up as broad tails and minor peaks on the various KDE plots. These most likely represent xenocrysts from the solidified Little Cottonwood pluton or, perhaps, antecrysts derived from a long-lived mushy granodiorite that survived until the intrusion of the geochemically distinct White Pine granite 2 m.y. later. Almost half of the ages were discordant (144 of 247 total) with younger 207Pb-206Pb ages (Figs. 4K4T). Discordance is probably the result of hydrothermal perturbation of U/Pb ratios first established during igneous crystallization.

Red Pine Porphyry

The youngest intrusive phase is the Red Pine porphyry. As with samples from the altered White Pine intrusion, many of the Red Pine zircon analyses produced discordant ages (43 of 86; Figs. 4U4X). The magmatic age inferred from the KDE plots of samples CH-115 and CH-053 is ca. 26 Ma (25.4–25.7 Ma). Xenocrysts (or antecrysts) from the White Pine phase (ca. 27 Ma) are the most common secondary component. Zircons older than 29 Ma, which could be xenocrysts derived from the Little Cottonwood stock, are rare in the two samples (n = 3; Figs. 4U4X and 5). The effects of the Red Pine intrusion on zircon in older rocks are evident in shoulders and minor peaks in the KDE curves (Fig. 5) for the White Pine intrusion (prominent in sample CH-031and subtle CH-108, CH-041, and CH-127), and even in the main Little Cottonwood stock (CH-072 and pebble dike clast CH-036). Although the ages are different, the mineralogy and chemical composition of fresh White Pine and nonporphyritic Red Pine rocks are quite similar. Thus, it is difficult to positively assign a sample to a specific geologic unit unless it has a radiometric age or a strongly porphyritic texture.

Pebble Dikes

To understand the sequence of rocks that the east Traverse Mountains pebble dikes might have explosively sampled on their routes to the surface, we analyzed 30 zircon grains from the matrix of one dike and 42 sites on 38 zircon grains from a granodioritic clast included in a different pebble dike.

As expected, the dike matrix (sample CH-039) contained many zircons with ages much older than the intrusions. U-Pb zircon ages range from 28.5 Ma to as old as 1613 Ma. Despite the wide range of ages, most are concordant, including those of Precambrian age (Figs. 4I4J). Thirteen (13) out of 30 analyzed zircon grains from this dike had ages >800 Ma (Fig. 4I). There is a cluster of ages ranging from 994 Ma to 1162 Ma. These ages are similar to those of detrital zircons in the Little Willow and Big Cottonwood Formations, which have a prominent age peak from 900 to 1900 Ma (Spencer et al., 2012; Dehler et al., 2010). The meta-sedimentary rocks have a detrital age peak at ca. 2700 Ma that we did not find in the pebble dike, probably because of the small number of grains analyzed. These Neoproterozoic strata underlie Paleozoic sedimentary rocks in the Wasatch Range and form a domical carapace around the Little Cottonwood complex (Fig. 2).

Seventeen (17) of the zircon ages from the pebble dike matrix are Cenozoic with a concordia age of 30.5 ± 0.6 Ma and an MSWD of 5.2 (Fig. 4I). The weighted mean age (30.5 ± 0.5 Ma; Fig. 4J) is similar to the concordia age. The KDE plot of the Cenozoic ages also has a strong mode at ca. 30.5 Ma—essentially the age of the Little Cottonwood stock through which the pebble dike was emplaced (Fig. 5). It also has a small peak at 32.5 Ma formed by three grains. Stearns et al. (2020) analyzed zircon from the granodioritic stock as old as 35 Ma and titanite as old as 36 Ma. Perhaps these older zircons were derived from such rocks or from the 32–35 Ma Alta stock (Vogel et al., 2001; Smyk et al., 2018; Stearns et al., 2020).

Zircons extracted from a hornblende-bearing granodiorite clast (sample CH-036) taken from a pebble dike on east Traverse Mountains provided a relatively simple age spectrum with concordant ages ranging from 30.5 to 26 Ma (Figs. 4G and 4H). Only one of the 42 ages was rejected because of discordance. The weighted mean age of the 41 concordant grains was 28.8 ± 0.3 Ma (Fig. 4H), but the MSWD is high (62) and the KDE spectrum is not unimodal (Fig. 5). The mixing model identified two peaks—a prominent one at 29.3 Ma (accounting for 66% of the data) and a smaller one formed by spots with ages from 27.2 to 26.0 Ma. Based on the presence of hornblende and the U-Pb ages, we conclude that the clast was a xenolith from the solidified Little Cottonwood granodiorite (which has the same age as the older zircons) and the small younger peak is the result of recrystallization of igneous zircon or of forming hydrothermal zircon when the White Pine magma intruded it ca. 27 Ma. Similar variation was found in age spectra of the Little Cottonwood stock.

The preponderance of Little Cottonwood and White Pine ages in the pebble dikes, the presence of quenched fragments of Red Pine magma, and hydrothermal alteration of all three phases suggest that the Red Pine phase was the source of the dikes and associated hydrothermal alteration.

Geochronologic Summary

Taken together, the zircon age spectra reveal three broad peaks at ca. 29.5, 27, and 26 Ma, which we interpret to be crystallization ages for the Little Cotton-wood stock, the White Pine phase, and the Red Pine porphyry, respectively (Fig. 5). The first two peaks are consistent with published ages for the Little Cottonwood stock and for the White Pine phase (Constenius, 1998; Vogel et al., 2001; Smyk et al., 2018; Fig. 5). However, the KDE plots reveal other peaks and shoulders that show each intrusive phase has inherited zircons from an older phase. Moreover, most of the samples also have a younger peak that corresponds with the age for a younger intrusion. For example, deconvolution of the age spectra for a mineralized sample of the White Pine phase (sample CH-031; Fig. 5; Table 1) produces a prominent peak at ca. 26.5 Ma (40%) that we interpret to be the age of zircons crystallized from the White Pine magma; it has another smaller inferred peak at ca. 30.2 Ma (4%) that is probably formed by xenocrysts from the Little Cottonwood stock. A third population of zircons in sample CH-031 peaks at ca. 24.9 Ma (19%), which we interpret to be zircon that either reequilibrated at the time of the younger Red Pine intrusion or newly crystallized from hydrothermal fluids associated with the Red Pine phase in an already solid portion of the White Pine phase.

The older zircons in the pebble dike matrix are detrital zircon from Proterozoic clastic metasedimentary rocks, which would have been cut by pebble dikes emanating from the Red Pine porphyry. The younger clast in the pebble dike apparently included partially recrystallized zircon from the Oligocene Little Cottonwood stock that would have been completely solid when the 3–4-m.y.-younger Red Pine magma generated these hydrothermal explosion dikes.

The Little Cottonwood stock is compositionally similar to other volcanic and plutonic rocks of middle Cenozoic age from the adjacent Great Basin (e.g., Christiansen and McCurry, 2008). Major element–SiO2 variation diagrams (Fig. 9) show relatively straightforward trends in the freshest rocks, with TiO2, Al2O3, Fe2O3, MnO, MgO, CaO, and P2O5 varying inversely with SiO2 across all units. Na2O slightly decreases and K2O slightly increases with increasing SiO2, but both show some scatter. Marsh and Smith (1997) concluded that the stock crystallized at a relatively high fO2 of 10−12 to 10−15 bar at 750–850 °C, which is consistent with its magnesian character (Fig. 6B). The Little Cottonwood stock plots as granodiorite to granite in a total alkali–silica diagram (Fig. 6A) and is part of a high-K2O series (Fig. 9G). It falls in the volcanic arc granite field in the Pearce et al. (1984) tectonic discrimination diagram (Fig. 6C). Zr concentrations are quite low for granodiorite (100–170 ppm; Fig. 6E). Consequently, the average zircon-saturation temperature (Watson and Harrison, 1983) for fresh rocks is also rather low at 768 °C for the Little Cottonwood stock (n = 15, σ = 15 °C), near the hornblende-plagioclase temperatures calculated by Hanson (1995) but distinctly higher than the Zr-in-titanite temperature (717 °C) calculated by Henze (2020).

The chemical composition and fractionation trends of the White Pine rocks are similar to those of the Little Cottonwood stock (Figs. 6 and 9). The White Pine intrusion is magnesian and high-K and plots as granite in a total alkali–silica diagram (Figs. 6A6B and 9G). It is generally more silicic than the Little Cottonwood stock, although there is much overlap of the two intrusive units (Figs. 6 and 9). The White Pine intrusion is compositionally different from the Little Cottonwood stock in some ways. For example, it has higher FeOT/(FeOT + MgO) for a given SiO2 concentration and typically lower concentrations of compatible elements like Ti, Mg, P, Sc, Sr, Cr, and Ni (Figs. 6B and 9). However, the White Pine phase has higher concentrations of Zr for a given TiO2 concentration (Fig. 6E), implying that it crystallized at higher temperatures despite its higher silica content and lower abundance of mafic minerals. The average zirconsaturation temperature for fresh rocks is 798 °C for the White Pine intrusion (n = 17, σ = 10 °C), ~30 °C hotter than the Little Cottonwood stock. In addition, White Pine titanites are chemically distinct from those in the Little Cottonwood stock, with different Fe/Al ratios and F contents (Martin et al., 2021). Like the other rocks of the composite pluton, the White Pine falls in the volcanic arc granite field in the Pearce et al. (1984) tectonic discrimination diagram (Fig. 6C).

The Red Pine porphyry is geochemically similar to the White Pine intrusion in most geochemical diagrams, though with only three samples, two of which are heavily altered, solid geochemical trends cannot be reliably established. The average zircon-saturation temperature is 813 °C (n = 3, σ = 16 °C), also higher than that of the granodiorites of the Little Cottonwood stock.

Chemically, the rhyolite dikes are much more differentiated than the other units and lie at the extreme end of the chemical trends. They have the highest silica and incompatible element concentrations (K2O, Rb, Nb, Y, Th, and U) and lowest compatible element concentrations (TiO2, FeOT, Al2O3, MgO, CaO, P2O5, Sr, and Zr; Figs. 6 and 11; Item S2, footnote 1). They also have no surviving ferromagnesian silicates, but it is unclear whether that absence is due to alteration or a primary magmatic feature. In the other units, Y declines to quite low levels as SiO2 increases. However, Y is strongly enriched in the rhyolite dikes (Fig. 11). The average zircon-saturation temperature is 729 °C for the rhyolites (n = 6, σ = 17 °C), lower than for the more mafic units, but these temperatures may be compromised by high K/Na ratios, indicative of alteration (Fig. 12). Negative Nb anomalies are inferred from high Rb/Nb, Th/Nb, and K/Nb ratios but not La/Nb ratios, which are low (Fig. 11). Low concentrations of light rare earth element (LREEs) are probably the result of fractionation of LREE-enriched allanite ± monazite. The rhyolites have low Zr (40–60 ppm; Fig. 6E) and positive Pb and negative TiO2 anomalies (Fig. 11). All of the above suggests that these are highly differentiated arc-type magmas, rather than A-type rhyolites like the topaz rhyolites erupted during the late Cenozoic in the adjacent Basin and Range province (Christiansen et al., 1986). However, they have high Nb + Y contents, moderately high Ga concentrations (21–30 ppm), and Ga/Al ratios (Figs. 6C and 6D) all of which are taken to be characteristic of within-plate or A-type granites (Pearce et al., 1984; Whalen et al., 1987).

As noted, geochemical characteristics of some rocks in the Little Cotton-wood suite are obscured by effects of hydrothermal alteration, especially the concentrations of more mobile elements. For example, Rb and K2O have a strong linear positive correlation and K2O is enriched beyond magmatic values in a few samples, while Na2O and CaO both have negative correlations with K2O and extend to low concentrations in the most altered samples of all units (Fig. 12). CaO concentrations <0.5 wt% are uncommon in unaltered igneous rocks, even high-silica rhyolites, but the most altered White Pine and Red Pine rocks have concentrations below this threshold (Fig. 12). Because of this, we used less-mobile trace elements, especially Zr and TiO2, to study the magmatic evolution of the system. A Zr-TiO2 variation diagram (Fig. 6E) indicates two separate and parallel positive trends for these two immobile elements, one for the Little Cottonwood stock and another for the White Pine and Red Pine units. The rhyolites are far more depleted than the other rocks in both of these compatible elements.

Mantle-normalized (McDonough and Sun, 1995) trace element patterns (Fig. 11) have strong similarities for the Little Cottonwood stock, White Pine intrusion, and Red Pine porphyry. These three units all have large positive Pb and large-ion lithophile element anomalies and negative Nb and Ti anomalies. The rhyolites have positive Pb and U anomalies, strongly negative Ba, P, and Ti anomalies, and small negative Sr anomalies. Moreover, the rhyolites have low concentrations of LREEs but are enriched in Y.

The chemical variations defined by each unit indicate that fractional crystallization of the observed minerals from different parent magmas controlled the geochemical trends. In this system, SiO2, K2O, Rb, and U behave incompatibly and TiO2, Al2O3, FeOT, MgO, CaO, P2O5, Sr, Y, Zr, and Ba behave compatibly. Most of the differentiation trends are typical of an I-type calc-alkaline system (Marsh and Smith, 1997). We concur with Marsh and Smith (1997) who proposed that the negative trends of FeOT, MgO, TiO2, Zr, and CaO versus SiO2 are driven by fractionation of biotite, hornblende, titanite, magnetite, and zircon, while depletion of Sr is driven by plagioclase fractionation, and depletion of Ba by biotite and K-feldspar removal. P2O5 depletion is likely due to apatite fractionation.

An unusual characteristic of most of the igneous rocks in this area is their very low Y concentrations, mostly <10 ppm (Fig. 13A) in all the intrusive igneous units except for the rhyolite dikes. When plotted on a Sr/Y–Y diagram (Fig. 13B), most of these samples fall into the adakitic range (Defant and Drummond, 1993; Richards and Kerrich, 2007), with Sr/Y ≥20 and Y ≤18.

Origin and Evolution of the Magmatic Units

White Pine and Red Pine Units versus Little Cottonwood Stock

We interpret the White Pine intrusion and Red Pine porphyry as forming from two magmas derived separately from the Little Cottonwood magma, not differentiated from it. At a macroscopic level, the White Pine and Red Pine phases can be distinguished from the Little Cottonwood granodiorites based on the lack of hornblende and of mafic enclaves. Although the White Pine and Red Pine units typically have lower concentrations of compatible elements and lie on linear trends that include the Little Cottonwood stock, some variation diagrams, especially TiO2-Zr (Fig. 6E), show a differentiation trend that is parallel to but separate from that of the Little Cottonwood stock. The typically more evolved White Pine and Red Pine magmas have higher Zr concentrations for a given TiO2 value. Thus, the White Pine and Red Pine rocks, which extend to higher silica concentrations, have higher average zircon-saturation temperatures than the Little Cottonwood stock (798 °C and 813 °C versus 768 °C). Also, the White Pine and Red Pine units have higher FeOtotal/(FeOtotal + MgO) ratios for a given SiO2 concentration than the Little Cottonwood stock (Fig. 6B), which we interpret to indicate that they had a lower fO2 than the granodiorites of the main phase. Finally, both units are roughly 3–4 m.y. younger than the Little Cottonwood stock based on their U-Pb zircon ages (Fig. 4; Table 1).

The KDE plots of the zircon ages for the White Pine and Red Pine magmas show that they likely incorporated some zircon grains from intrusions that pre-dated each of them. Furthermore, shoulders or small peaks of younger age in the Little Cottonwood and White Pine samples show possible influence on the zircon populations of these units by subsequent intrusions. The Mo-mineralized White Pine sample CH-031 shows a prominent example of this, with older inherited zircons (ca. 30.5 Ma) from the Little Cottonwood stock, zircons that crystallized from the White Pine magma (ca. 27.7 Ma), and an irregular young peak for secondary zircon (ca. 26 Ma) that formed when the Red Pine magma intruded. This hydrothermal influence appears to indicate that the Red Pine magma was the source of the Mo mineralization in the White Pine intrusion.

Rhyolite Dikes

The rhyolites are enigmatic because their composition is so different from that of the other units in the composite pluton: they have higher silica and incompatible element concentrations and lower concentrations of compatible elements except for anomalous enrichment of Y (Fig. 13A). The ages of the rhyolite dikes are not clear from cross-cutting relations, and an attempt to separate zircons for U-Pb geochronology was unsuccessful. However, the restriction of the rhyolite dikes to the White Pine intrusion area, their roughly east-west orientations, and nature of the secondary QSP alteration all suggest that the rhyolite dikes were part of the same White Pine–Red Pine magma system that was subsequently altered under conditions that stabilized the QSP assemblage.

Indeed, most trace element contents in the rhyolites follow the differentiation trends defined by the White Pine and Red Pine phases except for Nb, Y, and Sc, which are enriched in the rhyolites (Figs. 6F, 11, and 13A). This may be evidence that the rhyolites evolved from a separate magma. However, these anomalous trends could be explained by the lack of titanite and mafic phases in the final stages of differentiation. Therefore, during near-eutectic fractional crystallization, partition coefficients for Nb, Y, and Sc would be low. These elements could have been concentrated in the residual, highly differentiated melt, which then intruded upward as a series of dikes. It is unlikely that these anomalous trace element enrichments are simply another manifestation of alteration; such high-field-strength elements are generally immobile during hydrothermal alteration. Consequently, we interpret the rhyolites as being highly differentiated residual liquids from the White Pine or Red Pine magma. These enrichment trends are not unique to the rhyolites in this region; anomalous enrichment in Sc also occurs in the Pine Grove rhyolites (Keith, 1982) and Bishop Tuff rhyolites (Hildreth, 1979).

Strontium-Yttrium Ratios and the “Adakitic” Signature

The high Sr/Y ratios (almost all >20 and as high as 200 in fresh samples of the White Pine intrusion) and low Y (<10 ppm) concentrations in the Little Cottonwood stock and related units (Fig. 13) are puzzling, given that these are among the identifying characteristics of adakitic rocks. This is significant because a true adakitic slab-melt source for the pluton would require reinterpreting the tectonic setting of the region. It is also significant because some workers have suggested a link between adakites and porphyry-type mineralization (Defant et al., 2002; Lee and Tang, 2020), although others have disputed the relationship (Richards and Kerrich, 2007; Richards, 2011). The issue is fraught with debate, including what constitutes a truly adakitic signature (Richards and Kerrich, 2007; Castillo, 2012). For example, Smyk et al. (2018) suggested the Little Cotton wood stock is adakitic because of high La/Yb and La/Y ratios, which they interpreted to be caused by partial melting of granulite facies metamorphic rocks in the lower crust. However, rather than being a product of slab melting or melting of a garnet-rich plagioclaseabsent eclogite (as implied by Smyk et al. [2018]), the low Y in the Little Cotton wood stock could be due to fractionation of certain mineral phases (Moyen, 2009; Richards, 2011; Castillo, 2012). Amphibole, apatite, allanite, and titanite, all found in the Little Cottonwood stock, preferentially incorporate Y compared to melt. For example, Ackerson (2011) found the partition coefficient for Y in titanite in various dacitic to rhyolite ignimbrites is very high and varies from 490 to 1920 (with one major outlier), with a geometric mean of 1247. In the Fish Canyon Tuff dacite (Colorado, USA), Y has partition coefficients of 633 (for titanite), 13.5 (for hornblende), and 181 (for zircon) (Bachmann et al., 2005). The partition coefficient for Y is 162 for apatite in peraluminous granite (Bea et al., 1994), 28.8 in allanite in low-silica rhyolite, and 11.3 in amphibole in dacite (Ewart and Griffin, 1994).

There is a general decrease in Y with increasing SiO2 in the Little Cotton-wood stock, the White Pine intrusion, and the Red Pine porphyry (Fig. 13A), with partially overlapping fields. The fractionation of amphibole, apatite, allanite, zircon, and especially titanite, would cause the residual liquids to be progressively depleted in Y and enriched in silica. In at least one Little Cottonwood stock sample, CH-073, these Y-rich minerals combined make up >2.5% of the rock. Assuming titanite is the only phase important for Y variation, a drop from 17 ppm to 6 ppm Y would require <20% fractionation (using our average modal abundance of 0.7% titanite and a Y partition coefficient of 900). Similar calculations on the White Pine and Red Pine phases show that a much larger degree of crystallization, as much as ~75%, is required to reproduce the measured depletion of Y from the observed range of 13 ppm to 1.5 ppm. This is admittedly a large amount of fractionation, but the large range in Sr concentration in the White Pine intrusion, decreasing from 650 ppm to <200 ppm (Fig. 6G; Item S2, footnote 1), supports extensive fractionation in the evolution of the younger granitic magmas.

Indeed, the low concentrations of Sr argue against the deep crustal melting model, which requires the absence of feldspar in the source and would consequently enrich Sr in the parent melts of Y-poor magmas. Thus, melting of garnet-bearing and feldspar-absent rocks (i.e., eclogites) is necessary to acquire the high Sr/Y ratios of true adakites but appears to play no role in these lower-Sr rocks. Moreover, the notion that the Little Cottonwood stock and associated younger intrusions are “fertile” for ore mineralization is not supported by our data (cf. Smyk et al., 2018), given that the low Y concentrations are not an indication of a relationship with adakite.

Strontium-yttrium (Sr/Y) ratios have also been used to infer paleo–crustal thickness in magmatic arcs (e.g., Profeta et al., 2015) based on the notion that plagioclase is unstable and replaced by garnet in thick crust where the pressure is >~12 kbar (~45 km thick). Using the filtering criteria and empirical relationship defined by Profeta et al. (2015), the median Sr/Y ratio for the Little Cottonwood stock of ~61 implies a crustal thickness of 68 km during the Oligocene. Of course, the high Sr/Y ratio that implies this extreme crustal thickness could be exaggerated by titanite fractionation, which would have lowered the Y concentrations and raised the ratio.

Construction of the Pluton

The new field, geochronological, and geochemical data allow us to refine the sequence and time span over which the composite intrusion formed, intruded, chemically evolved, and became Mo mineralized. As Stearns et al. (2020) pointed out, the Little Cottonwood stock magmatic-hydrothermal system evolved in an episodic manner. The most common U-Pb zircon ages of the Little Cottonwood stock are 30–29 Ma. The intrusion is large—with an average exposed diameter of ~12 km and a thickness of at least 6 km based on depths of emplacement from John (1989), which give an approximate minimum volume of 680 km3—and therefore must have cooled slowly. Hot, mafic magma, now represented by the mafic enclaves, injected into the Little Cottonwood magma while it was still partially molten. This may have provided heat to keep the magma reservoir partially molten for an extended period of time. About 2–3 m.y. later, the more silicic White Pine magma formed separately and intruded as a series of small bodies in the northeastern section of the stock. Inherited zircons and gradational contacts suggest that the Little Cottonwood magma may have still been partially molten at the time of intrusion. After the White Pine intrusions had mostly crystallized, the Red Pine magma intruded it ca. 26 Ma. Vegetation, alteration envelopes, and glacial deposits make it difficult to find contacts within the pluton, but it is clear that there are several unconnected dike-like bodies of porphyritic and phaneritic Red Pine rock of varying textures. This supports the idea that repeated intrusions of small batches of water-rich Red Pine magma formed the Mo-W deposit and its alteration envelope.

Such complex multi-stage, mutually intruding small intrusions are common in other porphyry Mo systems, including the Henderson (Carten et al., 1988), Climax (Audétat, 2015), and Mount Emmons (Thomas and Galey, 1982) deposits, all in Colorado, as well as the Pine Grove (Utah; Keith et al., 1993), Questa (New Mexico, USA; Gaynor et al., 2019), and Endako (British Columbia, Canada; Villeneuve et al., 2001) deposits. At Endako, for example, the deposit is hosted within granitic rocks with zircon ages ranging from 149 to 145 Ma. The oldest, outermost biotite-hornblende granodiorite is intruded by the youngest, ore-bearing, altered, biotite monzogranite. This mirrors the pluton construction of the White Pine deposit, with an outer, fresh biotite-hornblende granodiorite being intruded by biotite granites that provided the mineralizing fluid and molybdenum over a period of ~5 m.y. Their common center of emplacement may be the result of spatiotemporal magmatic “focusing” along an already weakened channel through the crust, as outlined for nested Sierra Nevada plutons (California, USA) by Ardill et al. (2018).

Tectonic Setting

It is important to understand the tectonic setting of the Wasatch igneous belt in order to formulate a reasonable model for the origin and evolution of the Little Cottonwood magmas. However, there is little consensus on the processes involved. For example, Marsh and Smith (1997) concluded that the Little Cottonwood magma formed in response to back-arc extension and partial melting of high-K calc-alkaline tonalite in the crust but that magma from an underlying subduction zone provided the heat for crustal melting. Following Farmer and DePaolo (1983), Marsh and Smith (1997) saw no role for a mantle component in the felsic magma. In contrast, Vogel et al. (2001) concluded that the magmas of the Wasatch igneous belt were not subduction related. They proposed a complex structural model in which a pull-apart structure developed in a strike-slip zone related to middle Cenozoic extension, which generated magma by decompression melting of the mantle. This was supposed to have been localized on the inferred suture between Archean and Proterozoic crust, but the suture may actually lie much farther north (Nelson et al., 2011). According to this model, hot mantle-derived magma caused recently metasomatized lower crust to partially melt and then rise into the shallow extensional structures. More recently, Smyk et al. (2018) revisited the origin of the Wasatch igneous belt and concluded that the magma systems formed in response to melting in the asthenosphere caused by extension, followed by magma intrusion into and assimilation of the “delaminated lower crust” to create geochemical signatures that resemble those of magmas generated in a subduction zone and imparted an adakitic signature.

All three of these models invoke a role for extension and decompression melting—either of the lithospheric mantle or of the asthenosphere. Moreover, all invoke melting of the lower crust to generate the magmas of the Wasatch igneous belt without an important contribution of mantle-derived magma to the plutons; such magmas acted only as sources of heat in these hypotheses. We question these conclusions about the role of extension and the role played by mafic magmas as outlined below.

The extension model of Vogel et al. (2001) was largely based on an interpretation that the southern margin of the Little Cottonwood stock (Fig. 2) is an extensionally reactivated lateral ramp of a Sevier-age thrust system. In addition, a mylonitic shear zone between Mississippian sedimentary rocks and the Little Cottonwood stock is described as the roof fault of an extensional duplex. In contrast, we interpret both features to be part of a much younger shear zone (Fig. 2) related to the slip of a large landslide at 6 Ma that decapitated the Little Cottonwood pluton and overlying volcanic field, as described below (Keith et al., 2017; Kindred et al., 2018; Jordan et al., 2018; Chadburn, 2020). In addition, Kowallis et al. (1995) inferred that the northeast strike of middle Cenozoic dikes and microcracks in plutonic quartz is probably related to northeast-directed compression rather than northwest-southeast extension. Another structural argument against extension at this time lies in the observations of Jensen et al. (2020) that contractional deformation in the Uinta arch continued until ca. 30 Ma when strata of the Duchesne River Formation were folded and the erosional detritus formed the Bishop Conglomerate. Thus, magmatism in the Wasatch igneous belt, which started ca. 36 Ma and continued until ca. 26 Ma, was partially contemporaneous with this contractional deformation. Geochemically, felsic igneous rocks in this region that are older than ca. 24 Ma lack geochemical indicators of extension (A-type or within-plate granites) and are typical continental volcanic arc magmas (Christiansen et al., 2007).

Nonetheless, the age of onset of post-orogenic extension in the western North American Cordillera is debated (e.g., Best and Christiansen, 1991; Best et al., 2009; Canada et al., 2019; Zuza et al., 2021; Lund Snee and Miller, 2022). Although there may have been local extension during the emplacement of granitoids ca. 32–26 Ma, most Great Basin extension began in earnest much later at ca. 18 Ma, and extension in the study region possibly later than that. Kowallis et al. (1990) suggested displacement along the Wasatch normal fault began at 11 Ma based on apatite and zircon fission-track ages, while Armstrong et al. (2003) proposed 10–12 Ma using the same techniques along with apatite (U-Th)/He ages. Thus, magma origin as a result of extension is unlikely.

Instead, most regional reconstructions place the shallow oceanic lithosphere of the Farallon slab beneath this region at the time of pluton construction (e.g., Best et al., 2009; Yonkee and Weil, 2015; Smith et al., 2017). Thus, the Wasatch igneous belt and the Little Cottonwood stock are expected to be genetically related to subduction. Initially, the slab dipped at a shallow angle to create the Sevier and Laramide orogenies but eventually rolled back and steepened (e.g., Best et al., 2016; Copeland et al., 2017). In our view, the slab continued to dehydrate as it rolled back and fluid came into contact with hot asthenosphere, causing the overlying wedge of metasomatized mantle to melt. These subduction-related parent magmas rose into the crust, stalled because of density constraints, and assimilated large fractions of the continental crust, as indicated by their relatively high 87Sr/86Sr and especially by low εNd values (Fig. 14). A subduction-related origin is supported by the trace element compositions of the Little Cottonwood stock and White Pine intrusions, which fall in the volcanic arc granite field in the Pearce et al. (1984) tectonic discrimination diagram (Fig. 6C). Some of the more altered Red Pine and White Pine samples do not fall into the volcanic arc granite field, but we interpret that as being due to hydrothermal alteration enriching these rocks in mobile Rb. Enrichments in large-ion lithophile elements over high-field-strength elements, such as high Ba concentrations, positive Pb anomalies, and negative Nb and Ti anomalies (Fig. 11), are indicators of a subduction-zone setting as well (Rustioni et al., 2019). Finally, the oxidized character of the magmatic suite supports this idea, given that subduction-zone magmas typically have high fO2 in contrast to many extension-related magmas (Wood et al., 1990). Marsh and Smith (1997) reported that the granodiorite crystallized at an fO2 of 10−12 to 10−15 bars at 750°–850 °C, about three log units above the quartz-fayalitemagnetite (QFM) buffer. Low FeOtotal/(FeOtotal + MgO) ratios (Fig. 6B), indicative of relatively high fO2 which stabilized magnetite in the Little Cotton-wood, White Pine, and Red Pine magmas, are consistent with this conclusion. The slightly higher FeOtotal/(FeOtotal + MgO) ratios and TiO2 concentrations and presence of minor ilmenite in the younger White Pine and Red Pine magmas are probably due to a slightly lower fO2 for these somewhat younger and hotter magmas (798 °C and 813 °C in White Pine and Red Pine, respectively, versus 768 °C in the Little Cottonwood stock, based on zirconsaturation temperatures). Mg-rich biotite and amphibole compositions along with Cl-rich apatite (Hanson, 1995) are typical of oxidized subduction-related magma systems with high fH2O and fHCl as a result of slab dehydration.

Role of Mafic Magma

The contribution of mantle-derived magma to the mass of the Wasatch igneous belt is also controversial, with most of the previous investigations calling for a minimal role. However, its importance is evidenced by the abundant mafic enclaves in the Little Cottonwood stock, which Henze (2020) interpreted to be the result of mixing with mafic magma. The enclaves have SiO2 contents as low as 51 wt% and MgO as high as 6 wt% (Henze, 2020). A preferred orientation of the rock fabric elements, inclusions of K-feldspar megacrysts derived from the granite, titanite with distinctive reaction textures in both components, the association with mafic schlieren with wispy margins, and sprays of elongate apatite indicate that the enclaves were still molten when intruded into the felsic magma. The enclaves have elongate flattened shapes with complex, cuspate margins that may have formed by shearing as hotter mafic magma injected and then quenched in the cooler Little Cottonwood stock magma, or by new injections that ballooned the magma and compressed the enclaves outwards (Dorais et al., 1990). Nonetheless, mineral compositions in the enclaves reached equilibrium with the surrounding granite (Hanson, 1995). Smyk et al. (2018) proposed that mafic lamprophyre dikes in the Little Cottonwood stock are evidence that mantle-derived magma may have added mass to the evolving magma system.

Moreover, the major element compositions of the plutons themselves support the argument for a mantle component in that partial melts of the continental crust are unlikely to be as mafic as some of the rocks in the Wasatch igneous belt. For example, rocks from the Clayton Peak and Indian Hollow stocks in the belt are as mafic as gabbro, with silica as low as 48 wt% (Hanson, 1995; Vogel et al., 2001; Smyk et al., 2018); the nearby, slightly older Bingham Canyon complex has shoshonite as low as 54 wt% SiO2, and melanephelinite as low as 44 wt% SiO2 (Maughan et al., 2002). Production of such low-silica magmas by partial melting of normal continental crust that has an average silica content of ~61 wt% would be unlikely (Sisson et al., 2005; Rudnick and Gao, 2014). Even melting of rocks comparable in composition to the lower crust (53.4 wt% SiO2) leads to more silicic magmas for reasonable melt fractions (<35%). For example, experimental partial melting of an oxidized gabbro with 54 wt% SiO2 produced granitic liquids with SiO2 >72 wt% for all melt fractions (<30%) examined by Sisson et al. (2005).

Commonly, Nd and Sr isotopic data can be used to assess the role of mantle-derived mafic magma in the origin of plutonic rocks. Here, isotopic data are not conclusive, but they are at least consistent with a mantle contribution to the magma systems if accompanied by significant crustal assimilation (Fig. 14). For the Little Cottonwood stock, initial 87Sr/86Sr ratios range from 0.7074 to 0.7084 and εNd values range from −18.5 to −17.3 (Farmer and DePaolo, 1983; Vogel et al., 2001; Smyk et al., 2018). Other igneous rocks along the Wasatch igneous belt also have very low εNd but ranging from about −18 to −13. Given the very low εNd values, it is clear that a significant fraction of the Nd in these rocks originated in the continental crust, which regionally is ca. 2–2.6 Ga (Farmer and DePaolo, 1983; Yonkee et al., 2014). No primitive or nearly primitive magma with εNd values this low has been identified nearby (Kempton et al., 1991; Tingey et al., 1991; Nelson and Tingey, 1997; Allen, 2012). Even for alkaline magmas generated from extensively metasomatized subcontinental lithospheric mantle, εNd is no lower than −14 (Fig. 14). Nonetheless, the relatively low initial 87Sr/86Sr ratios (~0.708) of the Little Cottonwood stock contrast with those of many crustally derived magmas. For example, Late Cretaceous peraluminous granites—thought to be derived exclusively from Proterozoic basement rocks in eastern Nevada—have much higher initial 87Sr/86Sr values, ranging from ~0.710 to 0.733 (Wright and Wooden, 1991). Farmer and DePaolo (1983) interpreted the low 87Sr/86Sr ratios to mean that the Little Cottonwood magma was derived from depleted lower continental crust with low Rb/Sr ratios but normal Sm/Nd ratios. On the other hand, assimilation and fractional crystallization models involving a mantle-derived magma assimilating continental crust can reproduce the isotopic composition of the Little Cottonwood stock if the crust has an 87Sr/86Sr of <~0.710—but the ratios of assimilation to crystallization rates need to be high (>0.6). Generation of such low Sr isotopic compositions by assimilation of upper continental crust (with 87Sr/86Sr of 0.720) would be unlikely unless the mafic magma had a high Sr/Nd ratio (>10). Another indication of the involvement of mafic mantle-derived magma comes from the range of εNd values (−18.5 to −12.8) in the Wasatch igneous belt as a whole. Variable degrees of interaction between mafic magma with high εNd and crust with low εNd could explain the observed range of values.

The strong crustal overprint precludes knowing much about the presumed mantle component. It also makes it difficult to ascertain whether the ultimate origin of the mafic magmas was in the asthenosphere above a subducting slab or entirely within the lithosphere. Nonetheless, an important role for mafic magma, however generated, is indicated for the Little Cottonwood stock and the Wasatch igneous belt in general.

In short, we conclude that the composition and age of the magmas that generated the Little Cottonwood complex are consistent with formation in a continental magmatic arc, generated by subduction and rollback of the Farallon plate in the mid-Cenozoic, and that they include significant lower-crustal and likely mantle components. Although Vogel et al. (2001) proposed that an inherited weakness in the lithosphere along a Proterozoic suture zone contributed to the alignment of the plutons along the Wasatch igneous belt, an alternative control on the alignment could have been a tear in the subducting plate formed during rollback. This model would allow both flux melting (dehydration upon heating of the slab as it rolled back) and decompression melting (rise of asthenosphere and decompression through the slab tear) along a linear trend. Such a scenario could produce the parental magmas of the Wasatch igneous belt, either as a heat source to induce melting of the crust or as a source of primitive magma followed by assimilation of continental crust (e.g., Best et al., 2016).

Evidence for Venting

Some porphyry Mo deposits are known to have had extrusive episodes in addition to intrusive components (Mutschler et al., 1981; Keith et al., 1993). Good examples include mineralized magmatic systems at Pine Grove (Utah; Keith et al., 1986), Questa (New Mexico; Lipman, 1983; Gaynor et al., 2019), and three in Colorado: Mount Emmons (Thomas and Galey, 1982; Sillitoe, 1985), Climax (Audétat, 2015), and Henderson (Carten et al., 1988; Mercer et al., 2015).

No volcanic rocks directly linked to the Little Cottonwood composite pluton have yet been positively identified; the Keetley Volcanics in the eastern Wasatch igneous belt are too mafic and too old at 34.6–35.5 Ma (Smyk et al., 2018) to have vented from the Little Cottonwood stock magma, and the volcanic rocks on east Traverse Mountains, at 35.25 ± 0.13 Ma to 35.7 ± 0.6 Ma (Biek, 2005), are likewise too old. However, Stearns et al. (2020) proposed that the adjacent Alta stock acted as a conduit for magma from the Little Cottonwood stock to reach the surface and erupt, and that the Keetley Volcanics to the east and the volcanic rock of east Traverse Mountains are erupted equivalents of a combined Little Cottonwood stock–Alta stock system. They proposed that the White Pine and Red Pine intrusions may have erupted as well. The lack of extant volcanic rocks of the same age as the Little Cottonwood to Red Pine intrusions may be a result of erosion following offset on the Wasatch normal fault and isostatic uplift of the footwall.

Regardless of whether or not known volcanic units can be tied to the Little Cottonwood stock or associated units, there is both field and textural evidence that the system vented. Phreatomagmatic pebble dikes and the hydrothermal breccia pipe are the most obvious and give strong evidence of volatile exsolution from a magma chamber, whether or not they breached the surface. Another line of evidence of venting is the presence of quench textures and features in the White Pine intrusion and especially the Red Pine porphyry (Figs. 8A, 8B, and 8D). Elongate apatite and zircon grains and the overall porphyritic-aphanitic texture of the Red Pine porphyry, and to a lesser extent the White Pine intrusion, point to venting, which causes quench growth of elongate grains and rapid drops in pressure along with the loss of volatile fluid. The White Pine and Red Pine intrusions are also the only units that have embayed and/or resorbed quartz, which can form by decompression (Nekvasil, 1991).

Another example of this process is the Johnson Granite Porphyry of the Tuolumne Intrusive Suite in the Sierra Nevada batholith, suggested by Bateman and Chappell (1979) to have textures consistent with the degassing, depressurization, and quenching that occur during eruption. Among these textures are embayed quartz and a fine-grained groundmass. Like the Red Pine porphyry, the Johnson Granite Porphyry is porphyritic-aphanitic with resorbed quartz phenocrysts and K-feldspar megacrysts.

Based on the above evidence, we suggest that the highly fractionated Red Pine magma partially crystallized, concentrating volatiles in the remaining melt phase until it became fluid saturated. This induced fracturing and venting to fuel one or more eruptions, as suggested by Stock et al. (2016). Decompression melting occurred as Red Pine magma ascended in the volcanic vent, which destabilized quartz and caused the embayed texture. Then, degassing of water from the magma during eruption raised the solidus temperature, causing the remaining interstitial melt fraction to crystallize (Bateman and Chappell, 1979).

When this happened in relation to the pebble-dike emplacement is not completely certain, but there must have been venting and degassing sufficient to quench some Red Pine magma prior to the emplacement of some of the pebble dikes because multiple dikes (e.g., samples CH-061A and CH-066; Item S2, footnote 1) have entrained clasts of porphyritic Red Pine rock as well as clasts of Little Cottonwood stock and the White Pine phase (Fig. 10). Mineralization pre-dated some of the pebble dikes, as indicated by certain clasts in dike samples from the unmineralized east Traverse Mountains (CH-027 in Item S2, CH-036, and CH-039) having QSP alteration and containing molybdenite. It is not unusual for an ore district to have several generations of pebble dikes (e.g., Johnson, 2014). Cross-cutting dikes show that is the case here as well, so magmatism, fluid saturation, venting, and mineralization were likely partially contemporaneous.

Molybdenum-Tungsten Mineralization

The architecture of the subeconomic Mo-W deposit (Figs. 3 and 15) shows that multiple small bodies of White Pine magma intruded the Little Cottonwood stock at its apex. Small plugs or fingers of Red Pine porphyry subsequently intruded the White Pine phase, possibly providing the source of the ore mineralization, which is manifest in the fracture-controlled alteration of the stockwork zone (Sharp, 1958; Krahulec, 2015). A few late-stage pebble dikes and a breccia pipe, also part of this magmatic-hydrothermal system, emanate upward from the White Pine and/or Red Pine units and are now mostly either eroded away or covered within the intrusion. Many late-stage rhyolite dikes cut through a wide area in the White Pine intrusion, Little Cottonwood stock, and possibly Red Pine porphyry as well, oriented northeast-southwest, which is in accordance with regional stress at the time (Sharp, 1958; Kowallis et al., 1995). QSP alteration overprints the White Pine intrusions, the Red Pine bodies, and the surrounding parts of the Little Cottonwood stock, while potassic alteration is centered in small zones of rock that are very similar around the breccia pipe and on both sides of the ridge more than half a kilometer from the pipe. This mineralized and altered rock may also be the same rock previously mined in adits 2 km to the north of the breccia pipe (Sharp, 1958; Crittenden, 1965), which are now filled with waste rock and inaccessible. Scheelite occurs in an area of at least 6 km2 in an elevation range of ~1 km, mostly in the distal fringe zone, although there is some scheelite in the stockwork zone. It was deposited on some of the QSP-altered fracture surfaces in the White Pine unit and in the surrounding Little Cottonwood stock (Sharp, 1958).

The presence of tungsten in the form of scheelite (CaWO4) is unusual for a granitic host; W in Mo deposits is typically present as wolframite [(Fe,Mn)WO4]. While scheelite is recovered from Mo skarns in high-CaO country rocks associated with arc-related porphyry Mo deposits (Taylor et al., 2012), the scheelite here is found in veinlets in the host intrusive igneous rock and disseminated in rhyolite dikes. Scheelite forms in preference to wolframite if the fO2 is lower, as established by Hsu (1976, 1977). We suggest that the lower oxidation state may be the result of magmatic assimilation of organic-rich shales found in the country rock around the intrusion (e.g., Manning Canyon Shale) or of the passage of hydrothermal fluids through such rocks. Sphalerite was also present in the potassically altered breccia pipe. While it is not highly unusual to have sphalerite in the heart of an ore-related intrusion, it usually forms along with galena and quartz in more distal or peripheral zones (Ludington and Plumlee, 2009; Taylor et al., 2012).

Sharp (1958) noted some similarities between the White Pine and the Climax (Colorado) porphyry Mo deposits. Subsequent workers have identified two discrete types of porphyry Mo deposits, and the White Pine deposit seems less like Climax types and more closely associated with arc-related porphyry Mo deposits such as Endako (British Columbia), Quartz Hill (Alaska, USA), and Thompson Creek (Idaho, USA), as classified by Taylor et al. (2012). Among the parameters in which it fails to meet the Climax-type characterization are: (1) the deposit has a lower ore grade than typical for Climax-type deposits; (2) the deposit lacks widespread gangue fluorite; (3) the trace element compositions of the ore-related intrusions are unlike Climax-type intrusions for several trace elements (F, Zr, Sr, Rb, etc.); (4) the host granitoid straddles the boundary of I-type and A-type granites, and is not clearly A-type like those associated with Climax deposits, according to the Whalen et al. (1987) classification (Fig. 6D), and it was emplaced in a subduction-related magmatic arc, not a post-orogenic extension environment.

Climax-type deposits are typically higher grade than arc-related porphyry Mo deposits, at ~0.3%–0.5% molybdenite (0.2%–0.3% Mo) (Guilbert and Park, 1986), with typically 100–1000 Mt of ore (Ludington and Plumlee, 2009), while arc-related deposits have 0.03%–0.22% Mo and typically have >50 Mt of ore (Taylor et al., 2012). The White Pine deposit is estimated to hold between 14.5 and 16 Mt of Mo at 0.1% Mo grade, or 0.16% molybdenite, and 0.02% WO3 (Krahulec, 2015; Smyk et al., 2018). The grade and tonnage of ore in the White Pine deposit is thus more characteristic of arc-related porphyry Mo deposits.

The mineralogy and geochemistry of the deposit are also closer to those of arc-related porphyry deposits than Climax-type deposits, although in some ways the deposit is transitional. In Climax-type deposits, fluorite is a significant gangue mineral (Ludington and Plumlee, 2009), but in arc-related deposits, quartz is the dominant gangue mineral and fluorite is only minor (Guilbert and Park, 1986; Taylor et al., 2012). Sharp (1958) found fluorite only in a geographically limited area in and near the Red Pine porphyry on the west ridge, whereas quartz is abundant in stockwork veins, ribbon veins, vugs, bull quartz, and silicified granite throughout the White Pine deposit.

The source plutons of Climax-type deposits generally have >250 ppm Rb, >20 ppm Nb, >2000 ppm F, <100 ppm Sr, and <120 ppm Zr (Ludington and Plumlee, 2009). Few whole-rock analyses from the unaltered plutonic rocks of the White Pine deposit fit any of these criteria, and none fit all of them (Fig. 6; Item S2, footnote 1). Of the moderately to heavily altered and mineralized plutonic rocks, a few fit some of these criteria. Three of 44 samples have >250 ppm Rb, four have >2000 ppm F, and only one has <100 ppm Sr. The rhyolite dikes do have most of these characteristics but are unlikely to be the source of the mineralization because they are all themselves barren of Mo mineralization. Igneous rocks in arc-related Mo deposits, however, usually have <300 ppm Rb, <30 ppm Nb, and >100 ppm Sr, which describes all samples except the late rhyolite dikes and the three highly altered White Pine samples with high Rb probably introduced during alteration (Fig. 12).

Finally, the inferred tectonic setting at the time of emplacement and the type of granite do not support a Climax-type origin for the Mo-W deposits. As described above, the timing is controversial, but regional contraction probably persisted until after 30 Ma and extension started significantly later in the Miocene. This implies that the 27–26 Ma White Pine deposit did not originate in a truly extensional, post-subduction tectonic setting as is called for in the model for Climax-type deposits (Ludington and Plumlee, 2009). In addition, Climax-type deposits are typically related to A-type granites, but the Little Cottonwood stock, White Pine intrusion, and Red Pine porphyry are generally I-type intrusions based on the chemical criteria of the Whalen et al. (1987) Ga/Al versus Zr discrimination diagram and plot as volcanic arc granites on the Pearce et al. (1984) Rb versus Y + Nb discrimination diagram (Fig. 6C). The ore-related intrusions also have low FeOtotal/(FeOtotal + MgO) ratios (Fig. 6B), have hornblende and biotite with low Fe/Mg ratios (Hanson, 1995), and have magnetite and titanite—all indicators of a fairly high oxidation state. John (1989) also concluded the stock is in an oxidized calc-alkaline I-type granitoid, based on the mineralogy and chemical composition. These mineralogical and chemical characteristics are more consistent with the type of setting that hosts arc-related deposits—a subduction zone with oxidized calc-alkaline granitoids. Taylor et al. (2012), however, did not preclude the possibility of hybrid deposit types between Climax type and arc-related porphyry Mo type, and that may be the most fitting characterization of the White Pine deposit, given that it bears some similarities to Climax-type deposits as well.

Connection to the East Traverse Mountains Mega-Landslide

An outgrowth of our investigations is the possibility that the mineralized composite pluton was decapitated by a large landslide, as first proposed by Keith et al. (2017). Among the various lines of evidence in support of the mega-landslide hypothesis are the pebble dikes in the east Traverse Mountains that are similar to the pebble dikes in the White Pine deposit. The pebble dikes provide not only support for this hypothesis but also clues to a possible mechanism for how the landslide formed.

Pebble-Dike Evidence Supporting the Landslide Hypothesis

Intrusive igneous clasts in the east Traverse Mountains dikes have the same textures and mineral assemblages as the White Pine, porphyritic Red Pine, and Little Cottonwood stock units (Figs. 8 and 10). Both sets of dikes have similar east-west to northeast-southwest orientations and are spatially associated with argillic and QSP hydrothermal alteration. Uranothorite, molybdenite, and scheelite in the east Traverse Mountain dikes, indicator minerals of molybdenum deposits, are also found in the White Pine Mo deposit.

The two pebble-dike samples from the east Traverse Mountains have U-Pb zircon age populations that include White Pine and Red Pine ages as well as Little Cottonwood ages (Fig. 5; Table 1). Sample CH-039, the matrix of a pebble dike, has a strong age peak at 30.4 ± 0.6 Ma (an uncertainty of 2% is assumed here), within the range of U-Pb zircon ages for the Little Cottonwood stock (Fig. 5), but has many other zircon xenocrysts with ages extending back to 1.6 Ga. An intrusive clast, sample CH-036, has a peak at 29.3 ± 0.6 Ma, also within the range of ages for the Little Cottonwood stock. Both samples include a small number of zircons with younger ages that lie within the age ranges for the White Pine and Red Pine phases (Fig. 5). We interpret the intrusive clast to be derived from part of the Little Cottonwood stock that was altered and silicified when the younger phases intruded it, creating young zircon rims.

We propose that the Red Pine porphyry, and perhaps also the earlier White Pine intrusion, generated the pebble dikes. These pebble dikes sampled already-solidified or mostly solidified portions of the Little Cottonwood, White Pine, and Red Pine intrusions at depth before penetrating the sedimentary and volcanic rocks now exposed in the east Traverse Mountains but which rested above the Little Cottonwood stock at that time. Thus, the White Pine pebble dikes represent the lower part of the magmatic-hydrothermal system while the east Traverse Mountains pebble dikes represent the upper portion of the same system.

We propose that about 20 m.y. later (Jordan et al., 2018), following a radical change in the tectonics of the region accompanying the onset of extension and uplift of the Wasatch Range, the east Traverse Mountains block broke loose in a catastrophic landslide and slid rapidly to its current location.

Areas of cataclastic deformation and pseudotachylyte in the region (Fig. 2) are interpreted to be damage zones caused by sliding (Perfili et al., 2017; Kindred et al., 2018; Chadburn, 2020). This event cut the pebble dike swarm, with the upper part entrained in the landslide block and the lower portion rooted in the Little Cottonwood stock–White Pine intrusion host.

The presence of clasts of the three lithologies found in the White Pine deposit, the presence of molybdenite, scheelite, and uranothorite, and the presence of QSP alteration in the pebble dikes make it unlikely that there is any other possible source for the east Traverse Mountains pebble dikes.

Possible Alteration Contribution to the East Traverse Mountains Landslide

The intrusion of the Little Cottonwood stock below the paleo-location of the east Traverse Mountains may have contributed to the slide. The middle Cenozoic volcanic rocks of the east Traverse Mountains are argillically altered, and pockets of clays in pebble dikes and clays in samples of the Red Pine and White Pine units are evidence of a convecting hydrothermal fluid system, which circulated from the cooling intrusion into the overlying east Traverse Mountains block. This argillic alteration may have produced or enhanced a zone of weakness, which has since eroded away. Stearns et al. (2020) concluded that the Little Cottonwood–Alta hydrothermal system eventually focused on the margin of the Alta stock to the east, which may have been the where the incipient landslide first broke away on its eastern margin (Fig. 2).

Alteration could have occurred along a fault, or along the contact between the Little Cottonwood stock and the overlying Paleozoic sedimentary strata, particularly if there was an impermeable layer such as the thick Manning Canyon Shale or Doughnut Formation (Biek, 2005), along which circulating hydrothermal fluids could have locally concentrated. Evidence for this alteration is found in the overlying slide block in the brecciated quartzite of the Geneva Rock quarry on the west end of east Traverse Mountains. One quartzite sample (CH-081; Fig. 2) has ~5.3% Fe2O3 and ~2.4% Al2O3 as a result of alteration of more pure quartzite in the area (e.g., sample CH-083), which has about 96% SiO2, 1.2% Al2O3, and 0.5% Fe2O3. A possibility for the location of the slide surface is a weakness in the altered pluton itself, especially given the coincidence between stockwork fractures and alteration in the White Pine intrusion. The slide surface is no longer visible in most locations, having largely been eroded away from the surface of the autochthonous block; only small zones of pseudotachylyte and cataclasite have been preserved, and the base of the east Traverse Mountains is buried by Quaternary valley-fill sediment.

Investigation of the timeline of events in the construction of the White Pine ore deposit shows a complex magmatic and hydrothermal history over >5 m.y. (Fig. 15). About 30–29 Ma, the Little Cottonwood magma appears to have formed in a continental-arc tectonic setting, rather than in a post-orogenic extensional environment as proposed by others. The magma intruded folded and faulted Proterozoic to Paleozoic sedimentary rocks (Fig. 15A). Along with two feldspars, biotite, and quartz, it crystallized titanite, allanite, hornblende, zircon, and apatite, which depleted heavy rare earth elements and Y in the residual melt as it differentiated and gave the resultant pluton an apparent adakitic geochemical signature. U-Pb ages and whole-rock compositions show that ca. 27–25 Ma, the separate White Pine (Fig. 15B) and then Red Pine magmas (Fig. 15C) intruded the Little Cottonwood stock as distinct batches. Later rhyo lite dikes also intruded, although their absolute age is uncertain. Regional geology, mafic enclaves, lamprophyre dikes, elemental compositions, and Sr and Nd isotopic values suggest some mantle component was incorporated in the magmas, although they extensively assimilated lower continental crust.

Hydrothermal alteration accompanied each intrusion of the long-lived magmatic system (Stearns et al., 2020). Eventually, the Red Pine porphyry formed a hydrothermal system that led to the arc-related porphyry Mo-W deposit hosted in the White Pine intrusion. QSP alteration pervasively affected the entire region of the deposit. Venting may have occurred as early as just after intrusion of the Little Cottonwood stock, more probably after intrusion of the White Pine magma and certainly following intrusion of the Red Pine magma (Fig. 15D). This is manifested in the form of pebble dikes, if not volcanic activity as well. The venting was voluminous enough to cause the quenching, via release of volatiles, of the apical Red Pine magma, creating its quench textures and resorbed minerals. Pebble-dike emplacement, fueled by venting of the volatile-saturated zone or circulation of hydrothermal fluids (Fig. 15E), injected clasts of Little Cottonwood stock, White Pine intrusion, and quenched Red Pine porphyry into the older overlying sedimentary and volcanic section now exposed on the east Traverse Mountains. Alteration pre-dated and post-dated emplacement of the pebble dikes, as seen in some samples with juxtaposed heavily altered and nearly fresh intrusive igneous clasts of rock types that are usually only altered. Pebble-dike injection and QSP alteration processes partially overlapped, allowing altered and unaltered clasts to be included in the dikes. Subsequent alteration of some pebble dikes after emplacement at a shallow level occurred during multiple generations of dike emplacement.

Overprinting of lower-temperature alteration followed injection of most pebble dikes and solidification of the system. This includes propylitic alteration in the form of chlorite and epidote seen in many intrusive samples from the region. Argillic alteration is manifested by the pervasively altered volcanic rocks in the central and eastern portions of the Traverse Mountains as well as pockets of clay in the pebble dikes and crandallite-jarosite alteration rimming some of the clasts (Fig. 15E). Local silicification occurred as well. Alteration in the thrust sheet of sedimentary rocks overlying the intrusion-related hydrothermal system and at the contact between the composite pluton and the overlying Paleozoic sedimentary rocks assisted the slide's eventual catastrophic failure (Fig. 15F).

Inception of the Wasatch fault along the western margin of the Little Cotton wood stock by 12 Ma resulted in several kilometers of displacement. This caused a 20° tilt to the east in the pluton and provided the steep topographic gradient necessary for the east Traverse Mountains mega-landslide. The slide is thought to have formed ca. 6.5 Ma (Keith et al., 2017; Chadburn, 2020); sub sequent displacement and erosion stripped away the source region and exposed the intrusion. Key to this interpretation are pebble dikes in the east Traverse Mountains that bear clasts resembling the Little Cottonwood stock, White Pine intrusion, and Red Pine porphyry in mineralogy, texture, age, and alteration mineral assemblage. These features suggest that the pebble dikes are the decapitated upper part of the same pebble-dike system found in the White Pine deposit 17 km away, supporting other evidence that the east Traverse Mountains once rested atop the White Pine porphyry Mo system. Continued tilting and subsequent Pleistocene glacial erosion (Fig. 15G) obscured the original geometry of the intrusion.

Funding for this study came from the Brigham Young University (BYU) College of Physical and Mathematical Sciences. Ron Harris, Bob Biek, Mike Dorais, Ken Krahulec, and Tim Thompson provided valuable insights on the project. Scott Thayne provided access to the Geneva Rock quarry on the east Traverse Mountains. Several people helped with analytical lab work at BYU and at the University of Utah, including Mike Stearns, Diego Fernandez, Grant Rea Downing, Mike Standing, David Tomlinson, and Tober Dyorich. Thanks for many hours of field and lab work go to the Traverse Mountain students, including Cameron Harrison, Ryan Chadburn, Bret Young, Spencer Burr, Chris Perfili, Lars Jordan, Thane Kindred, Riley Dunkley, Tia Misuraca, Alec Hoopes, Sam Martin, Alec Martin, Joseph Tolworthy, and Steven Hood. Reviewers Gary S. Michelfelder and N. Seymour provided important feedback to shape the paper.

1Supplemental Material. Sample locations, modal mineralogy, whole-rock geochemistry, U-Pb isotopic data, instrumental methods, and cathodoluminescence images of zircon grains. Please visit https://doi.org/10.1130/GEOS.S.19401212 to access the supplemental material, and contact editing@geosociety.org with any questions.
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Associate Editor: Graham D.M. Andrews
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