Uplift, Rupture, and Rollback of the Farallon Slab Reflected in Volcanic

Home , Abert Rim, Adakite, Bimodal volcanism, Flat slab subduction

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Journal of Geophysical Research: Solid Earth

RESEARCH ARTICLE Uplift, rupture, and rollback of the Farallon slab reﬂected 10.1002/2017JB014517 in volcanic perturbations along the Yellowstone Key Points: adakite hot spot track • Volcanic perturbations in the Cascadia back-arc region are derived from uplift Victor E. Camp1 , Martin E. Ross2, Robert A. Duncan3, and David L. Kimbrough1 and dismemberment of the Farallon slab from ~30 to 20 Ma 1Department of Geological Sciences, San Diego State University, San Diego, California, USA, 2Department of Earth and • Slab uplift and concurrent melting 3 above the Yellowstone plume Environmental Sciences, Northeastern University, Boston, Massachusetts, USA, College of Earth, Ocean, and Atmospheric promoted high-K calc-alkaline Sciences, Oregon State University, Corvallis, Oregon, USA volcanism and adakite generation • Creation of a seismic hole beneath eastern Oregon resulted from thermal Abstract Field, geochemical, and geochronological data show that the southern segment of the ancestral erosion and slab rupture, followed by Cascades arc advanced into the Oregon back-arc region from 30 to 20 Ma. We attribute this event to thermal a period of slab rollback uplift of the Farallon slab by the Yellowstone mantle plume, with heat diffusion, decompression, and the release of volatiles promoting high-K calc-alkaline volcanism throughout the back-arc region. The greatest Supporting Information: • Supporting Information S1 degree of heating is expressed at the surface by a broad ENE-trending zone of adakites and related rocks • Data Set S1 generated by melting of oceanic crust from the Farallon slab. A hiatus in eruptive activity began at ca. • Data Set S2 22–20 Ma but ended abruptly at 16.7 Ma with renewed volcanism from slab rupture occurring in two separate • Data Set S3 regions. The eastern rupture resulted in the extrusion of Steens Basalt during the ascent and melting of a dry

Correspondence to: mantle (plume) source contaminated with depleted mantle. The contemporaneous western rupture resulted V. E. Camp, in renewed subduction, melting of a wet mantle source, and the rejuvenation of high-K calc-alkaline [email protected] volcanism near the Nevada-California border at 16.7 Ma. Here the initiation of slab rollback is evident in the westward migration of arc volcanism at 7.8 km/Ma. Today, the uplifted slab is largely missing beneath the Citation: Oregon back-arc region, replaced instead by a seismic hole that is bound on the south by the adakite hot spot Camp, V. E., M. E. Ross, R. A. Duncan, and track. We attribute slab destruction to thermal uplift and mechanical dislocation that culminated in rapid D. L. Kimbrough (2017), Uplift, rupture, – and rollback of the Farallon slab tearing of the slab from 17 15 Ma and possible foundering and sinking of slab segments from 16 to 10 Ma. reﬂected in volcanic perturbations along the Yellowstone adakite hot spot Plain Language Summary Yellowstone National Park is underlain by a rising plume of hot rock track, J. Geophys. Res. Solid Earth, 122, from the Earth’s deep mantle that has provided the melt source for three supereruptions over the last 2 7009–7041, doi:10.1002/2017JB014517. million years. This hot spot, or mantle plume, appears to be a long-lived feature that resided offshore beneath the oceanic Farallon plate before being overridden by the westward moving North American plate about 42 Received 5 JUN 2017 Accepted 15 AUG 2017 million years ago (Myr). Between 42 and 17 Myr the Yellowstone plume was shielded beneath the subducting Accepted article online 17 AUG 2017 Farallon slab, with little surface expression on the overriding North American plate. After 17 Myr the plume Published online 9 SEP 2017 reemerged to produce a great outpouring of basaltic lava known as the Columbia River ﬂood basalts of eastern Oregon and southeastern Washington. New and compiled chemical and age data suggest that the Farallon slab was uplifted and dislocated by the thermally buoyant Yellowstone mantle plume between 30 and 20 Myr, with oceanic crust of the Farallon slab melting to generate unusual rocks called adakite. Eventual destruction of the Farallon slab beneath eastern Oregon was associated with a long-lived period of thermal erosion and tearing of the slab from 30 to 17 Myr, followed by the foundering and sinking of slab segments from 16 to 10 Myr.

1. Introduction Cenozoic volcanism in the Cascade magmatic arc is derived from subduction of the Farallon plate, with the southern limit of volcanism progressing northward, contemporaneous with northward migration of the Mendocino triple junction [Atwater, 1970; McBirney, 1978; Dickinson and Snyder, 1979]. The modern magmatic arc is composed of active volcanoes of the High Cascades and related volcanic rocks that erupted after ca. 4 Ma. The older portion of the arc erupted from ca. 45 to 4 Ma and includes highly dissected volcanic rocks of the Western Cascades in Oregon, Washington, and northernmost California. These older volcanic products constitute the northern segment of the ancestral Cascades magmatic arc (Figure 1) [du Bray and John, 2011]. A lesser known but concurrent southern segment of the ancestral arc was proposed by Noble [1972] and ﬁ ©2017. American Geophysical Union. Christiansen and Yeats [1992] but de ned more precisely in a series of recent studies as a broader belt of All Rights Reserved. calc-alkaline volcanism straddling the Californian-Nevada border region and active from ca. 30 to 4 Ma

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Figure 1. Location map of geographic and physiographic features in the arc and back-arc regions of southern Oregon, northeastern California, northwestern Nevada, and western Idaho. Distribution of the ancestral Cascades volcanic arc from du Bray and John [2011] and du Bray et al. [2014]. Volcanic rocks from the northern segment of the ancestral Cascades arc range from >45 Ma to 4 Ma, but in Oregon from ca. 35 to 4 Ma [du Bray and John, 2011]. Rocks of the southern segment of the ancestral Cascades arc range in age from ca. 30 to 4 Ma [du Bray et al., 2014]. The proposed tear in the Farallon/Juan de Fuca slab between these two arc segments was suggested by Cousens et al. [2008] and Colgan et al. [2011] to account for the apparent left-lateral offset of their volcanic axes.

(Figure 1) [Putirka and Busby, 2007 Busby et al., 2008a, 2008b; Cousens et al., 2008, 2011; Hagan et al., 2009; Busby and Putirka, 2009; Putirka et al., 2012; John et al., 2012; du Bray et al., 2014]. Volcanism of the ancestral Cascades appears to have occurred above a Farallon slab shown to be intensely dislocated in recent seismic investigations [e.g., Burdick et al., 2008; Schmandt and Humphreys, 2010; Sigloch, 2011; Tian et al., 2011; Liu and Stegman, 2011; Tian and Zhao, 2012]. Other studies attribute this dismemberment to interaction of the Farallon slab with the Yellowstone hot spot, where rupturing of the slab resulted in plume rise and the Miocene eruption of the Columbia River flood basalts [Xue and Allen, 2007, 2010; Obrebski et al., 2010; Humphreys and Schmandt, 2011; Coble and Mahood, 2012; Darold and Humphreys, 2013; Wells et al., 2014; Camp et al., 2015]. Although disruption of the Farallon plate, and/or changes in slab dip, should also result in distinct changes in the style and distribution of arc/back-arc magmatism, such perturbations have not been recognized or thoroughly explored. Here we use field data and a large database of geochemical and geochronological data to define time- dependent variations in the distribution, composition, and eruption style of Oligocene-to-Miocene volcanism in eastern Oregon and adjacent Nevada, California, and Idaho. Our intent is to examine whether or not these variations in time and space are consistent with a dynamic model of plume-slab interaction as conceived in tectonic models [e.g., Coble and Mahood, 2012] and seismic investigations [e.g., Xue and Allen, 2010; Obrebski et al., 2010].

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2. Analytical Data and Methods We present new geochronological data on 20 samples and new geochemical data on 299 samples from volcanic rocks in the Oregon-Nevada-California border region. Major and trace element data were derived from X-ray fluorescence (XRF) at the San Diego State University Geoanalytical Laboratory, with a subordinate number of XRF analyses from the P. R. Hooper Geoanalytical Laboratory at Washington State University. This data set is available at the EarthChem repository [Camp and Ross, 2017]. We use and compare these analyses with a large geochemical database of 2637 analyses compiled from specific site investigations [Norman and Leeman, 1990; Langer, 1991; Mathis, 1993; Camp et al., 2003; Carmichael et al., 2006; Colgan et al., 2006, 2011; Scarberry et al., 2010], together with regional compilations from the ancestral southern Cascades, the ancestral northern Cascades, the Columbia River Flood Basalt Province, and the northern Nevada rift system [du Bray and John, 2011; du Bray et al., 2006, 2014; Hooper, 2000; John et al., 2000; Camp et al., 2013] (North American Volcanic and Intrusive Database (NAVDAT), http://www.navdat.org). Table 1 summarizes geochronological data on 12 mafic lavas and 3 rhyolitic rocks determined by the 40Ar/39Ar incremental heating method at the Noble Gas Mass Spectrometry Laboratory at Oregon State University. The summary age data for five additional rhyolite samples determined by U-Pb zircon ablation geochronology at the University of Arizona Laserchron Center are presented in Table 2. The analytical detail for each data set is available at the EarthChem repository [Duncan et al., 2017; Kimbrough et al., 2017].

3. Northern and Southern Segments of the Ancestral Cascades Plate convergence has been a major contributor to magmatism, terrane accretion, uplift, and deformation of the western U.S. throughout the late Mesozoic and early Cenozoic Eras. Late Jurassic and Cretaceous arc magmatism is evident in an extensive system of granitoid plutons extending from the Peninsula Ranges batholith in the south through the Sierra Nevada and Idaho batholiths in the north. A dramatic change in the style of subduction began in the late Cretaceous and continued into the early Paleocene as arc volcanism migrated eastward while decreasing in volume and intensity, contemporaneous with fold-and-thrust contraction of the plate interior during the Sevier-Laramide orogeny [Humphreys, 2009; Dickinson, 2013]. This event has been widely attributed to a period of flat-slab subduction involving a young and buoyant Farallon plate [Coney and Reynolds, 1977; Engebretson et al., 1985; Burchfiel et al., 1992; DeCelles, 2004]. Others suggest that flat-slab subduction might involve a thicker than normal oceanic crust [Livaccari et al., 1981], perhaps related to the subducted conjugate of the Shatsky Rise [e.g., Saleeby, 2003; Liu et al., 2010; Humphreys et al., 2015]. A transition from flat-slab subduction to a more typical regime of steeper subduction began in the Eocene, but the character and pace of this transition varied significantly with geographic position. In Nevada, the transition was prolonged, manifested in the broad formation and gradual southwest migration of calc-alkaline volcanism from Eocene through Miocene time [Best and Christiansen, 1991]. In the Pacific Northwest, the transition appears to have been more abrupt, resulting in a relatively rapid westward shift of the volcanic axis from Idaho to western Washington and Oregon, where Eocene volcanism was initiated in the northern segment of the ancestral Cascades [e.g., Humphreys, 2009]. In Nevada, the gradual migration of volcanic fronts from 50 to 20 Ma is often referred to as the “mid-Tertiary volcanic sweep.” Two interpretations have been offered for this volcanic propagation. Humphreys [1995] proposed that the shallow-dipping slab began to detach from the overlying lithosphere along an east-west axis in southern Nevada, thus pulling the northern edge of the slab southward. In this scenario, adiabatic decompression of asthenosphere rising into the trailing slab window induced melting and the southward propagation of volcanic fronts from ca. 50 to 20 Ma. The second interpretation, long advocated by many workers, attributes the volcanic sweep to post-Laramide steepening of subduction during slab rollback [e.g., Coney and Reynolds, 1977]. This scenario appears to be more consistent with the southwestward migration of volcanism evident in mapping relationships and palinspastic reconstructions [Cousens et al., 2008; Dickinson, 2013; du Bray et al., 2014]. In the Pacific Northwest, the more abrupt westward shift in volcanism was associated with Eocene accretion of the Siletzia terrane in the coastal region of Oregon and Washington. Several workers describe this terrane as a large oceanic basalt plateau that may have formed above the Yellowstone hot spot before colliding with the continental margin [Duncan, 1982; Schmandt and Humphreys, 2011; McCrory and Wilson, 2013; Wells et al.,

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Table 1. Summary of 40Ar-39Ar Incremental Heating Agesa Age Spectrum Total Fusion Inverse Isochron

Sample IGSN Description Coordinates Age ± 2 s (Ma) 39Ar% K/ca MSWD n/N Age ± 2 s (Ma) Age ± 2 s (Ma) 40Ar/36Ar MSWD

SW-2 gm IEVEC006O Basaltic trachyandesite flow, 40°43040.80″N 11.49 + 0.16 41 0.623 4.51 3/11 11.37 ± 0.09 11.97 ± 0.25 42 ± 74 2.43 Earth Solid Research: Geophysical of Journal Spanish Springs Peak shield 120°12031.52″W volcano, NE California CR-671 gm IEVEC008C Welded ignimbrite, Greaser 42°608.60″N 13.24 ± 0.16 47 14.8 6.62 3/12 13.19 ± 0.08 13.51 ± 0.26 272 ± 21 2.2 Canyon, southern California 119°43023.15″W CR-691 gm IEVEC008B Boulder Mountain shield 41°21049.67″N 11.91 + 0.14 100 0.375 1.52 12/12 11.93 ± 0.13 11.93 ± 0.16 295 ± 5 1.65 volcano, east of Surprise Valley, 119°46037.49″W NW Nevada CR-692 gm IEVEC005X Basalt lava beneath bedded 41°11027.87″N 15.60 ± 0.11 80 0.347 0.99 7/11 15.28 ± 0.12 15.58 ± 0.15 297 ± 6 1.12 pyroclastics, Wall Canyon Road, 119°47034.36″W NW Nevada CR-735 pl IEVEC005H Plagioclase-phyric trachybasalt, 40°54030.47″N 12.99 + 0.71 95 0.027 0.42 6/7 12.77 ± 0.87 12.88 ± 1.95 306 ± 226 0.52 Buckhorn Road NW Nevada 119°51035.33″W CR-713 pl IEVEC004V Basaltic trachyandesite, 40°44023.51″N 14.75 ± 0.39 100 0.048 0.32 9/9 14.82 ± 0.48 14.73 ± 0.45 298 ± 27 0.36 Buffalo Meadows, NW Nevada 119°51015.74″W CR-727 gm IEVEC0059 Basalt lava flow, Buckhorn Road, 40°51037.55″N 14.52 + 0.16 90 0.163 0.04 7/11 14.72 ± 0.19 14.52 ± 0.33 295 ± 131 0.04 Buffalo Hills, NW Nevada 119°53028.64″W CR-726 gm IEVEC0058 Olivine basalt flow, Buckhorn 40°55025.71″N 15.66 + 0.13 91 0.273 0.17 7/11 15.40 ± 0.14 15.65 ± 0.14 299 ± 18 0.17 Road, Buffalo Hills, NW Nevada 119°49010.26″W CR-693 gm IEVEC005Y Basaltic andesite lava flow, 41°10057.86″N 16.81 + 0.13 72 0.718 0.05 6/11 16.54 ± 0.13 16.79 ± 0.23 296 ± 4 0.06 south of High Lake, NW Nevada 119°18051.79″W CR-667 gm IEVEC004L Basalt lava flow, Buffalo Hills, 40°46038.15″N 16.23 + 0.17 88 0.273 1.33 9/12 16.22 ± 0.17 16.19 ± 0.21 297 ± 4 1.41 NW Nevada 119°50045.92″W CR-666 gm IEVEC004K Basalt lava flow, Buffalo Meadows 40°4600.88″N 16.64 + 0.14 54 0.820 0.11 4/12 16.63 ± 0.15 16.74 ± 1.24 288 ± 50 0.11 Road, NW Nevada 119°50034.68″W CR-665 gm IEVEC004J Basalt lava flow, Smoke Creek 40°3207.02″N 15.97 + 0.16 98 0.111 3.1 10/12 15.95 ± 0.15 15.87 ± 0.16 308 ± 11 2.02 Road, NW Nevada 119°54042.77″W CR-673 pl IEVEC004P Rhyolite, Buffalo Hills, Lone 40°49021.92″N 14.67 + 0.98 98 0.008 1.5 8/9 14.10 ± 1.23 15.88 ± 2.48 283 ± 25 1.52 Juniper Canyon road, NW Nevada 119°3204.85″W SC-24 gm IEVEC003N Basaltic andesite lava, second 40°31029.18″N 15.90 + 0.18 49 0.331 1.52 6/12 16.45 ± 0.19 16.05 ± 1.23 286 ± 78 1.88 flow from the top of 600 m thick 119°56018.37″W Smoke Creek section, NW Nevada CR-702 gm IEVEC001L Trachybasalt, Lone Juniper 41°25031.30″N 24.29 + 0.18 96 0.155 1.39 9/10 24.21 ± 0.21 24.37 ± 0.22 294 ± 3 1.28 10.1002/2017JB014517 Springs, NW Nevada 119°38029.39″W aAges calculated using sanidine monitor FCT-3 (28.201 Ma [Kuiper et al., 2008]) and the total decay constant 5.53e 10/yr. N is the total number of heating steps; n the number of steps defining the age plateau/isochron. MSWD is an F-statistic that compares the variance within step ages with variance about the plateau/isochron. PL = plagioclase; gm = groundmass. Journal of Geophysical Research: Solid Earth 10.1002/2017JB014517

Table 2. Summary of Zircon U-Pb Ages Zircon 206*Pb/238 U ±2-Sigma #Grains #Grains Sample IGSN Description Latitude Longitude Age (Ma) Uncert (Ma) MSWDa Analyzed Rejected

CR-771 IEVEC005N Rhyolite lava ﬂow, 40°53020.50″N 119°36019.47″W 15.76 ±0.35 0.33 16 0 Route 447, NW Nevada CR-772 IEVEC005O Bimodal air fall tuff, 40°5304.31″N 119°3608.58″W 15.65 ±0.30 0.43 16 0 underlying CR-771, Route 447, NW Nevada CR-773 IEVEC005P Welded rhyolite tuff, 40°5404.86″N 119°3505.43″W 15.95 ±0.36 0.15 18 0 Squaw Valley Reservoir, NW Nevada CR-776 IEVEC005S Banded rhyolite ash 40°41052.20″N 119°49016.89″W 15.61 ±0.22 0.85 32 1 ﬂow, Buffalo Hills, NW Nevada CR-777 IEVECOO5T Vitrophyric rhyolite, 40°5102.62″N 119°3701.80″W 15.90 ±0.36 0.09 12 1 north of Poodle Mountain, NW Nevada aMSWD, mean square of weighted deviates.

2014]. Geophysical data suggest that the older parts of a greater Siletzia may have inserted and filled the Columbia Embayment of northern Oregon and central Washington at ca. 55–50 Ma, during or very shortly after calc-alkaline volcanism in the Challis-Absaroka arc in Idaho and in the Clarno formation of north-central Oregon [Schmandt and Humphreys, 2011; Darold and Humphreys, 2013]. This docking event necessitated a rapid transition from flat subduction to steep subduction on the western side of Siletzia, which, in turn, initiated early calc-alkaline volcanism in the Cascade arc at ca. 45–40 Ma [Wells et al., 1984; Humphreys, 2009; Schmandt and Humphreys, 2011]. By Oligocene time, two contemporaneous arc segments of the ancestral Cascades were active along continental margin: (1) the northern segment generated by steep subduction beneath coastal Washington and Oregon and (2) the broader southern segment generated by shallower subduction in the Nevada- California border region [du Bray and John, 2011; du Bray et al., 2014]. Humphreys [2009] noted that the boundary separating steep and flatter subduction requires a tear to have existed in the Farallon plate since Eocene time. Cousens et al. [2008] and Colgan et al. [2011] proposed a similar tear near the Oregon- California border to account for the difference in slab dip implied by >100 km of lateral offset between the northern and southern segments of the ancestral Cascades (Figure 1).

4. Yellowstone Hot Spot An important focus of this paper is to examine the potential inﬂuence of the Yellowstone hot spot on Farallon plate subduction and Cenozoic calc-alkaline volcanism associated with the southern Cascades ancestral arc. High-resolution seismic imagery resolves a present-day low-density conduit beneath Yellowstone National Park that has been interpreted by some as upwelling material from the transition zone separating the upper and lower mantle [Sigloch et al., 2008; James et al., 2011]. Others, however, have interpreted the conduit as a mantle plume ascending through the mantle transition zone from the lower mantle [Obrebski et al., 2010, 2011; Humphreys and Schmandt, 2011; Schmandt et al., 2012; Tian and Zhao, 2012; Darold and Humphreys, 2013]. Evidence for a signiﬁcant thermal anomaly beneath the Yellowstone hot spot is likewise controversial. Leeman et al. [2009] use trace elements and seismic constraints to suggest that the source region above the Yellowstone hot spot has a maximum potential temperature of 1450°C, which is cooler than upwellings beneath Hawaii and other oceanic hot spots. In contrast, Schmandt et al. [2012] provide seismic data across the 660 km mantle discontinuity consistent with an anomalously high-temperature for the plume-like upwelling beneath the Yellowstone region. A plume origin is consistent with a number of geological constraints noted by several workers [Camp and Ross, 2004; Hooper et al., 2007; Shervais and Hanan, 2008;

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Pierce and Morgan, 1992, 2009; Coble and Mahood, 2012; Camp, 2013]. These include, for example, the short duration (1.1 Myr) and high eruption rate (0.178 km3/yr) of the main-phase Columbia River Basalts, from 16.7 to 15.6 Ma [e.g., Camp, 2013], and the northeastward propagation of rhyolite volcanism along the Yellowstone-Snake River Plain trend, in the same direction and at the same rate as plate motion once the hot spot track became well established at ca. 12 Ma [Pierce and Morgan, 2009]. Other evidence used both for and against a plume origin comes from the geochemical data. Carlson et al. [1981] and Carlson [1984] studied the trace element and isotopic composition of Columbia River Basalts to suggest a nonplume origin for the flood-basalt eruptions. This interpretation was partly based on their iden- tification of a C2 isotopic source component that they interpreted as sediment-contaminated mantle. This component has since been reinterpreted by several workers as a plume component [e.g., Hooper and Hawkesworth, 1993; Hanan et al., 2008; Wolff et al., 2008; Camp and Hanan, 2008; Wolff and Ramos, 2013]. According to Wolff and Ramos [2013], trace element variations for rocks enriched in the C2 component suggest an enriched mantle II type of oceanic island source. Additional evidence for a deep-mantle source comes from high 3He/4He values (i.e., primordial mantle) for basaltic rocks located on both the Snake River Plain [Graham et al., 2009] and the Columbia Basin [Dodson et al., 1997]. The preponderance of data appears to support a mantle plume origin for flood-basalt volcanism of the Columbia River Basalt Group and for bimodal volcanism along the Yellowstone-Snake River Plain hot spot track. The earliest manifestation of plume arrival was the eruption of Steens Basalt in southeastern Oregon at ca. 16.7 Ma [Camp et al., 2013]. This event was essentially contemporaneous with the initiation of rhyolite volcanism along the Nevada-Oregon border at ca. 16.7–16.5 Ma [e.g., Henry et al., 2006; Brueseke et al., 2008; Coble and Mahood, 2012; Mahood and Benson, 2016], the initiation of bimodal volcanism in the northern Nevada rift at ca. 16.5 Ma [e.g., John et al., 2000], and the initiation of basin and range extension at ca. 16.5 Ma [Colgan and Henry, 2009; Camp et al., 2015]. Two separate models of plume emplacement have been proposed: (1) impingement of a starting plume head that first arrived in the mid-Miocene [e.g., Pierce and Morgan, 1992, 2009; Camp and Ross, 2004; Shervais and Hanan, 2008] and (2) subduction of a long-lived plume tail that was shielded from the overlying plate by the Farallon slab until its reemergence in the mid-Miocene [e.g., Duncan, 1982; Wells et al., 1984, 2014; Geist and Richards, 1993; Pyle et al., 2009; McCrory and Wilson, 2013]. In the latter model, arrival of the Yellowstone hot spot at the Farallon trench resulted in accretion of the Siletzia oceanic plateau at ca. 50–45 Ma, followed by the initiation of the northern Cascades ancestral arc at 45–40 Ma [Schmandt and Humphreys, 2011].

5. Volcanic Stratigraphy of the Oregon-Nevada-California Border Region The southern Oregon border region, between lat. 41° and 43°N, lies near the transition between the southern and northern ancestral Cascades (Figure 1). The volcanic stratigraphy of this region provides essential data on the Cenozoic dynamics of slab subduction and the potential perturbation of subduction associated with mid- Miocene emplacement of the Yellowstone hot spot. Although Eocene volcanic rocks in this region are rare, and volumetrically insigniﬁcant, they are present in the Black Forest, Pine Forest, and Santa Rosa Ranges of northern Nevada and at the base of a thick volcaniclastic sequence of Eocene rocks in the Warner Range of northeast California (Figure 2) [Colgan et al., 2006, 2011; Lerch et al., 2008; Brueseke and Hart, 2009]. The youngest volcanic rocks in the region are Late Miocene to Quaternary in age. These can be subdivided into two major groups (Figure 2): (1) high-alumina olivine basalt and subordinate rhyolite that form the volcanic cover on the Modoc Plateau of northeastern California and much of southern Oregon between long. 122°W and 119°W and (2) the youngest, calc-alkaline volcanic rocks (≤8 Ma) of the ancestral and modern Cascades. Lying between these informal lithostratigraphic units is a sequence of Oligocene to late Miocene volcanic rocks (30–8 Ma) that forms the focus of this paper. These rocks have a shared history with volcanic rocks in the ancestral Cascades and are subdivided here into three major units (Figure 2): (1) Oligocene to early Miocene calc-alkaline to mildly alkaline rocks (ca. 30–20 Ma), (2) mid-Miocene bimodal rocks of tholeiitic basalt and rhyolite (ca. 16.7–15 Ma), and (3) middle-to-late Miocene calc-alkaline rocks (ca. 16.7–8 Ma). The oldest of these three units is separated from the younger two units by a well-established hiatus in calc- alkaline volcanism that began in the southern Oregon border region at ca. 22–20 Ma and ended at 16.7 Ma

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(Figure 2) [McKee et al., 1970; McQuarrie and Oskin, 2010; Coble and Mahood, 2012; this study]. A slightly shorter volcanic gap is present the eastern Santa Rosa-Calico ﬁeld where calc-alkaline volcanic rocks as young as 19 Ma have been reported (Figure 1) [Brueseke and Hart, 2009]. Before discussing more regional interpretations, we ﬁrst describe the overall distribution, chemical character, and eruption style of each of these three volcanic successions.

5.1. Oligocene to Early Miocene Calc-Alkaline Rocks (ca. 30–20 Ma) Paleogene volcanic rocks occur across a broad region south of the Oregon High Lava Plains (Figure 1). Although largely covered by younger rocks, they can be found at the base of footwalls adjacent to large normal faults and as elevated volcanic complexes rising above the younger volcanic cover. These exposure locations are shown in Figure 3, and the general lithology, thickness, and age range of each are denoted in Table 3. Although a few Paleogene outcrops contain Eocene to early Oligocene volcanic rocks, the far greater volume is late Oligocene in age (most between 28 to 22 Ma), with a few rocks from Figure 2. Simplified Cenozoic volcanic stratigraphy of northwestern Nevada, northeastern California, and central Oregon, Rabbit Hills, Oregon extending into the early lying east and southeast of the modern Cascade volcanic axis. Miocene to 20 Ma (Table 3) [Scarberry et al., 2010]. The stratigraphic column includes the main-phase of Columbia The base of these outcrops is rarely exposed, but River Basalt (CRB) volcanism and coeval tholeiitic volcanism along the northern Nevada rift. Not included are post-15 Ma it is not unusual for the volcanic stratigraphy to lavas associated with the younger, more sporadic phase of CRB be several hundred meters thick, with the thick- volcanism in northeastern Oregon (i.e., Wanapum and Saddle est exposures in the Warner Range (up to Mountains Basalt), and coeval lavas associated with Lake 2500 m), at the base of Steens Mountain Owyhee-Powder River volcanic province of eastern Oregon (~1500 m), and in the Pine Forest Range [Ferns and McClaughry, 2013]. (>1000 m) (Figure 3 and Table 3). This sequence along the Oregon-Nevada border is contemporaneous with 30–22 Ma volcanic rocks comprising an 850 m thick composite section of the upper John Day formation in north-central Oregon [Bestland et al., 1997], which overlies volcaniclastic rocks of the 39.7–32 Ma Big Basin Member [Dilhoff et al., 2009]. The upper John Day sequence is dominated by felsic air fall pyroclastic rocks, with a subordinate volume of small alkali olivine basalt flows, trachyandesite flows, and more volumi- nous rhyolitic flows and domes [Walker and Robinson, 1990]. Volcanism appears to have ceased simultaneously in the southern border region and in the John Day basin at 22–20 Ma. It is important to note that the southernmost exposure of the John Day Formation at Hampton Butte is only 92 km from the northernmost exposure in the southern border region at Coleman Hills (Figure 3), and there is no indication that the volcanic rocks of either succession thin into central Oregon. Instead, these coeval exposures appear to represent a contiguous succession that is covered today by a younger sequence of Neogene volcanic rocks of the central Oregon High Lava Plains (Figure 3). The easternmost exposure of the Paleogene succession is present in the Owyhee Mountains of Idaho, where the Salmon Creek Volcanics form a 1225 m thick sequence of calc-alkaline basaltic andesite [Norman and Leeman, 1990]. The age range for these lavas is not well-constrained; however, whole-rock K-Ar ages have been reported between 30.9 and 26 Ma [Norman et al., 1986]. Volcanic rocks of equivalent age are absent in the deep-canyon exposures of the Snake and Imnaha Rivers of northeastern Oregon; however, they are present in the upper drainage basin of the Grande Ronde River, where trachyandesite

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Figure 3. Surface exposure of Oligocene-to-early Miocene volcanic rocks in Oregon, northwest Nevada, and northeast California. Most exposures not only range in age from ca. 30 to 20 Ma but also include minor exposures of Eocene to early Oligocene volcanic rocks (ca. 40–35 Ma) in the Pine Forest and Black Forest Ranges of northern Nevada and in the Big Basin member of the John Day Formation in north-central Oregon. Age ranges for the northern segment of the ancestral Cascades (light gray and dark gray) are from du Bray and John [2011].

lavas with K-Ar ages of 28.8 and 27.7 Ma underlie the mid-Miocene Columbia River flood basalts [Fiebelkorn et al., 1983; Robinson et al., 1990]. The northernmost extent of the John Day Formation is not known, but field evidence shows that it does not extend beneath Columbia River Basalt in central Washington [Campbell, 1989]. The style of eruption appears to be consistent with numerous, relatively small volcanic centers spread over a broad region of southern Oregon border region, and as far east as the Oregon-Idaho border. There is no evidence of large calderas with extensive pyroclastic sheetflows in the southern border region that are similar in age and size to the ca. 30 Ma Crooked River caldera and the ca. 32 Ma Tower Mountain caldera found in the lower John Day Formation to the north [e.g., McClaughry et al., 2009; Seligman et al., 2014] or contemporaneous calderas found in the interior andesite-rhyolite assemblage of central Nevada to the south [e.g., Ludington et al., 1996]. The chemical trends of individual outcrops in the southern Oregon border region are quite variable, from calc-alkaline to mildly alkaline and bimodal to intermediate (Figure 4). The Salmon Creek volcanics are intermediate-composition rocks that are typically calc-alkaline, whereas the Hart Mountain volcanics are largely felsic and strongly alkaline, with common pantellerites and peralkaline rhyolites [Norman and Leeman, 1990; Mathis, 1993]. None of the locations are bimodal in the traditional sense of basalt being the most mafic rock type. Instead, intermediate rocks are the dominate composition in all locations, with basalt being a minor component. In contrast to abundant rhyolitic rocks found in the John Day formation to the north and the andesite-rhyolite assemblage to the south, rhyolite comprises ~13% of collected and compiled samples. The majority of these, however, are from the lower Steens Mountain exposure (Figure 3), where 53% of collected samples are rhyolites concentrated in the lower part of the section [Langer, 1991].

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Table 3. Paleogene Volcanic Rocks of Southeastern Oregon, Northern Nevada, Northeastern California, and Southwestern Idaho Area/Formation Lithology Age Thickness Reference

Adel Mildly alkaline to calc-alkaline tachyandesite lava flows ? >100 m This study, Dooley [2010] South Abert Rim Mildly alkaline to calc-alkaline tachyandesite lava flows ? >150 m This study Coleman Hills and Mildly alkaline trachyandesite to rhyolite complex 23–20 Ma ~100 m Scarberry [2007], Rabbit Hills Scarberry et al. [2010] North Abert Rim Mildly alkaline basaltic trachyandesite to rhyolite volcano 22 Ma ~300 m Scarberry [2007], Scarberry et al. [2010], this study Lower Steens Mountain Calc-alkaline; felsic tuffs/flows; abundand rhyolite at base of ~24–22 Ma ~1500 m Fuller [1931], Hart section with andesite lavas at top of section and Carlson [1985], Minor et al. [1987], Langer [1991] Hays Volcano Calc-alkaline to mildly alkaline basalt to andesite lava flows, 24 Ma >550 m Carmichael et al. [2006] agglutinated spatter and felsic pyroclastic tocks Lone Juniper Springs Calc-alkaline to mildly alkaline basaltic trachyandesite flows 24.3 Ma >100 m This study Hart Mountain Mildly alkaline to strongly peralkaline; trachybasalt to alkali 26 Ma >600 m Mathis [1993] rhyolite lava flows and tuffs Warner Range Calc-alkaline to midly alkaline lava flows, andesite lava 28–24.5 Ma 700–2500 m Colgan et al. [2011] flows and block-and-ash flows, and felsic pyroclastic rocks Salmon Creek Calc-alkaline basalt-andesite 30.9–26 Ma 1225 m McIntyre [1972], Ekren et al. [1981], Volcanics Norman and Leeman [1990] Black Rock Range Calc-alkaline to mildly alkaline; basalt lava flows 35.7 Ma27–22 Ma >1000 m Lerch et al. [2008] (35.7 Ma) and younger bimodal basaltic andesite lava flows and felsic pyroclastic rocks Pine Forest Range Calc-alkaline to mildly alkaline; trachyandesite intrusion 38 Ma30–23.6 Ma ~800 m Colgan et al. [2006] (38 Ma) and younger bimodal basaltic lava flows and felsic pyroclastic rocks Hampton Butte Rhyodadicte to dacite domes 30.4–28.6 Ma >300 m Iademarco [2009] John Day Formation and Bimodal; felsic pyroclastic rocks and local andesite 39.7–22.6 Ma ~1100 m Walker and Robinson [1990], equivalent-age rocks Dillhoff et al. [2009], Iademarco in the northern High [2009], Evans and Brown [1981] Lava Plains Clarno Formation Calc-alkaline, intermediate 54–40 Ma >1800 m Walker and Robinson [1990], Bestland and Retallack [1994]

The great majority of Oligocene to early Miocene trendlines fall within the high-K calc-alkaline series of Peccerillo and Taylor [1976] (Figure 4c), with Fe-depletion trends typical in all locations (Figures 4e and 4f). In Figure 4d we plot compositional fields for volcanic rocks from the northern and southern segments of the ancestral Cascades; in Figures 4c and 4d we compare these fields with coeval rocks from the southern Oregon border region. The field boundary delineated in Figure 4d was formulated by Putirka et al. [2012] to separate 90% of rocks from the Walker Lane segment of the ancestral southern Cascades (high K) from rocks of the modern Cascades arc in northern California (lower K). Here we demonstrate that this boundary can be used in a wider context to separate the majority of rocks from the southern and northern segments of the ancestral Cascades (Figure 4d). Although some analyses from the Salmon River volcanics lie below the field boundary, all remaining analyses from the southern Oregon border region lie above it. These Oligocene to early Miocene sequences are distinctly more K-rich than volcanic rocks from northern segment of the ancestral Cascades, but with few exceptions identical to the high-K character of the coeval southern segment of the ancestral arc (Figures 4c and 4d). The overall high-K, calc-alkaline character of these rocks is also consistent with trace element variations. Individual eruptive centers and well-sampled outcrop areas in the southern Oregon border region show dis-

tinct enrichment in H2O-soluble elements Rb, Ba, K, and Pb, and depletion in Nb, as reflected in normalized trace element patterns for mafic rocks that are similar to those from the southern segment of the ancestral Cascades (Figure 5). The Hart Mountain pattern is the sole exception that appears to lack a significant negative Nb anomaly (Figure 5). We conclude that the overall major and trace element data for Oligocene to early Miocene rocks from the southern Oregon border region closely align with the chemical character of the ancestral southern Cascades. As a group, however, these rocks also display local variations between individual eruptive centers, most evident in Sr, Ti, and Zr concentrations, as well as Zr/Nb ratios (Figures 5 and 6). Although the Zr/Nb

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Figure 4. Analyses of Oligocene-to-early Miocene (ca. 30–20 Ma) volcanic rocks plotted on (a and b) the total alkali versus silica diagram of Le Bas et al. [1986], (c and d) the K2O versus silica diagram of Peccerillo and Taylor [1976] (modified in 4d to include the field boundary of Putirka et al. [2012]; see text for discussion), and (e and f) the total alkali-FeO*-MgO diagram of Irvine and Baragar [1971]. Separate compositional fields include 162 analyses from the ancestral southern Cascades with ages between 35 and 18 Ma (from du Bray et al. [2014]), 45 analyses from the Warner Range and Hays Mountain field of northeastern California (from Carmichael et al. [2006] and Colgan et al. [2011]), and 427 analyses from the northern ancestral Cascades (aka western Cascades; from du Bray et al. [2006]). The remaining analyses are shown as separate data points from Adel (18 analyses) [this study; Dooley, 2010], Lone Juniper Springs [this study], Abert Rim (6 analyses) [this study], Coleman and Rabbit Hills (24 [this study]; 15 analyses [Scarberry, 2007]), the Pine Forest Range [this study; Colgan et al., 2006], lower Steens Mountain [Langer, 1991], Hart Mountain [Mathis, 1993], Salmon Creek Volcanics [Norman, 1987], and miscellaneous locations in northeastern California, southern Oregon, and northern Nevada [this study].

ratios for the Warner Range and Lone Juniper Springs typify those found throughout the southern ancestral Cascades, the remaining outcrop regions have higher Zr and Nb concentrations and typically lower Zr/Nb ratios than those sampled elsewhere in the southern segment of the ancestral arc. Disparities in Zr/Nb cannot be readily attributed to variations in partial melting or fractional crystallization but are more likely related to compositional variations in the source rock beneath each locale (Figure 6).

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Figure 5. Incompatible trace element diagrams normalized to the primitive mantle of Sun and McDonough [1989] for Oligocene to early Miocene (ca. 30–20 Ma) basalts and basaltic andesites with SiO2 contents <57%. Only those analyses with Pb and Nb data were selected. Sources are from this study (Abert Rim, Adel, Coleman Hills, and Lone Juniper Springs; lower Steens Mountain [Langer, 1991]; Pine Forest Range [Colgan et al., 2006]; Pine Forest Range; Hart Mountain [Mathis, 1993]; and Salmon Creek Volcanics [Norman, 1987]). The gray pattern is the range in normalized trace element data for Oligocene to early Miocene basalts and basaltic andesites in the Warner Range of the ancestral southern Cascades volcanic arc (from Colgan et al. [2011] and du Bray et al. [2014]).

5.2. Mid-Miocene Bimodal Rocks of Tholeiitic Basalt and Rhyolite (ca. 16.7–15 Ma) The period of volcanic quiescence that began at ca. 22–20 Ma was terminated abruptly at 16.7 Ma with the outbreak of tholeiitic ﬂood-basalt volcanism of the Columbia River Basalt Group. The main phase of ﬂood- basalt eruption was from 16.7 to 15.6 Ma when 93% of the volume erupted in ~1.1 million years. Flood-basalt eruption was contemporaneous with early emplacement of massive rhyolite calderas at the western end of the Yellowstone-Snake River hot spot track from 16.5 to 15.5 Ma and with bimodal eruptions along the northern Nevada rift system from 16.5 to 15 Ma (Figure 7) [Zoback et al., 1994; John et al., 2000; Brueseke et al., 2008, 2014; Coble and Mahood, 2012].

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From oldest to youngest, the main phase of the Columbia River flood basalts includes Steens, Imnaha, and Grande Ronde Basalt, all of which erupted from prevalent dike swarms in eastern Oregon and eastern Washington. Less extensive dikes that fed basaltic lavas in the northern Nevada rift system appear to be at least partly coincident with aeromagnetic anomalies delineating large, mafic intrusions in the middle to shallow crust [Zoback et al., 1994; John et al., 2000]. Together, these dike systems, calderas, and aero- Figure 6. Zr versus Nb plot for 30 to 20 Ma volcanic rocks from outcrop loca- magnetic anomalies define the tions shown in Figure 3. Data from this study (Abert Rim, Adel, Colman Hills, Lone Juniper Springs, and Pine Forest Range; lower Steens Mountain [Langer, Nevada-Columbia Basin magmatic 1991]; Pine Forest Range and Warner Range [Colgan et al., 2006, 2011]; Hart belt, a relatively narrow (~300 km), Mountain [Mathis, 1993]; Salmon Creek Volcanics [Norman, 1987]; and short-lived (16.7–15 Ma) belt of bimo- Coleman and Rabbit Hills [Scarberry, 2007]). dal magmatism that extends for

Figure 7. (a) Surface exposure of Miocene bimodal volcanic rocks of tholeiitic basalt (black) and rhyolite (gray). Basalt from ca. 16.7 to 15 Ma includes the main phase eruptions of the Columbia River Basalt Group (Steens, Imnaha, Picture Gorge, Prineville, and Grande Ronde Basalt), as well as basalt outcrops in the northern Nevada rift. Rhyolites vary from middle-to- late Miocene, but most are 16.5 to 15 Ma in age. (b) Compositional fields for 688 analyses of the Columbia River Basalt main- phase lavas, from Hooper [2000] and Camp et al. [2013] plotted on the total alkali versus silica diagram of Le Bas et al. [1986], the K2O versus silica diagram of Peccerillo and Taylor [1976], and the total alkali-FeO*-MgO diagram of Irvine and Baragar [1971]. Compositional fields exclude a few analyses that are clear outliers to the bulk data set. (c) Compositional fields for the main phase of the Columbia River Basalt Group, as shown in Figure 7b, together with analyses of mafic lavas and dikes from in the northern Nevada rift and contemporaneous mafic rocks in the Humboldt Formation and Ruby Mountains-East Humboldt Range, from this study (38 analyses), John et al. [2000], and the North American Volcanic and Intrusive Rock Database (NAVDAT, http://www.navdat.org).

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~1000 km, from southern Nevada into eastern Washington (Figure 8) [Camp et al., 2015]. This magmatic belt marks a major change in the style of eruption following the 20–16.7 Ma volcanic hiatus, as the Oligocene eruption of intermediate to felsic lavas from widespread volcanic centers was followed by more focused fissure eruptions of flood basalt and localized bimodal volcanism along this linear trend. All main-phase units of the Columbia River Basalt Group show tholeiitic Fe- enrichment trends on the AFM diagram, with the majority of lavas lying along differentiation trends from basalt to andesite on the total alkali versus silica diagram (Figure 7). It is noteworthy, however, that the liquid line of descent in the upper part of the Steens stratigraphy differentiates toward more alkali- rich compositions of trachybasalt and basaltic trachyandesite (Figure 7). Figure 8. The short-lived (16.7–15 Ma) Nevada-Columbia Basin mag- Contemporaneous mafic rocks in Nevada matic belt extends along a north-south trend for over 1000 km lie along similar Fe-enrichment trends [Camp et al., 2015]. This belt marks an abrupt period of basaltic to fi bimodal volcanism coincident with the main eruptive phase of the but fall on both sides of the AFM eld Columbia River Basalt Group. It is defined by the main-phase dikes of boundary separating tholeiitic and the Chief Joseph, Monument, and Steens Mountain dike swarms from calc-alkaline rock types. These include 16.7 to 15.6 Ma, contemporaneous dikes and bimodal volcanic rocks of coeval lavas and dikes in the northern the northern Nevada rift from 16.5 to 15 Ma, rhyolite centers along the Nevada rift, dikes in the Ruby and East western end of the Snake River Plain hot spot track from 16.5 to 15.5 Ma, and large, midcrustal mafic intrusions delineated by an Humboldt Ranges, and basaltic lava extensive system of arcuate aeromagnetic anomalies [Blakely and flows interbedded in the mid-Miocene Jachens, 1991; Glen and Ponce, 2002]. These deep-seated intrusions Humboldt formation of northeast > vary in length from 100 to 400 km long with average widths of about Nevada (Figure 7). 5km[Ponce and Glen, 2002, 2008].

5.3. Middle to Late Miocene Calc-Alkaline Rocks (ca. 16.7–8 Ma) A dramatic change is evident in the distribution of calc-alkaline volcanism following the 20–16.7 Ma volcanic hiatus. Calc-alkaline volcanism in eastern Oregon extended as far north as the Blue Mountains near the Oregon-Washington border from 30 to 20 Ma; however, when it reemerged in the mid-Miocene, it was largely restricted to the Nevada-California border region, extending northward only ~75 km into southern Oregon (Figure 9). The ca. 16.7 to 8.0 Ma calc-alkaline rocks described here are considered to be part of the southern segment of the ancestral Cascades magmatic arc [Colgan et al., 2011; du Bray et al., 2014], which was active during eruption of contemporaneous 17–10 Ma volcanic rocks of the northern segment of the ancestral Cascades (Figure 9) [du Bray and John, 2011]. In western Nevada, these middle to late Miocene volcanic rocks form part of the western andesite assemblage of Ludington et al. [1996], and they correlate directly with the ca. 15.5 to 13 Ma Pyramid Sequence mapped on the east and west side of Pyramid Lake (Figure 1) [Bonham and Papke, 1969; Ressel, 1996; Henry et al., 2004, 2009a, 2009ab; Hinz, 2004]. In northeastern California, this ancestral southern Cascade sequence includes a group of ca. 16 to 14 Ma maﬁc lavas exposed in the Warner Range and a younger group of ca. 14 to 8 Ma lavas of intermediate compositions located south and west of the Warner Range [Colgan et al., 2011; du Bray et al., 2014]. Where exposed, the arc assemblage lies unconformably above the sequence of Oligocene to early Miocene volcanic rocks described earlier [Henry et al., 2009a; Colgan et al., 2011].

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Figure 9. (a) Surface exposure of middle-to-late Miocene (ca. 16.7–8 Ma) calc-alkaline rocks of the southern and northern segments of the ancestral Cascades (from du Bray and John [2011], Colgan et al. [2011], du Bray et al. [2014], and this study). The dashed line is western limit of ancestral arc volcanism. Outcrops of the southern ancestral Cascades closely correspond with a series of young (<5 Ma) transtensional faults of the northern Walker Lane belt [e.g., Busby, 2013]. The southern arc succession is dominated by basaltic to trachyandesitic lava ﬂows, with subordinate interbedded rhyolitic to dacitic tuff and tuffaceous sediment; however, felsic tuffaceous rocks become more common north and northwest of Lakeview, California into southern Oregon. The greater volume of these rocks erupted between ca. 16.7 and 13 Ma, with a few analyses in northeastern California as young as 8 Ma. (b) Analyses of this younger phase (ca. 16.7–8 Ma) of the ancestral southern Cascades rocks plotted on the total alkali versus silica diagram of Le Bas et al. [1986], the K2O versus silica diagram of Peccerillo and Taylor [1976], and the total alkali-FeO*-MgO diagram of Irvine and Baragar [1971]. The gray ﬁeld for shield volcanoes of northeastern California incorporates 73 analyses from this study and 12 analyses from Colgan et al. [2011]. The remaining analyses from northwestern Nevada and the Warner Range include 143 from this study, 13 from Carmichael et al. [2006], and 14 from Colgan et al. [2011].

Previous 40Ar/39Ar ages place the base of this mid-Miocene succession at ca. 16 Ma [e.g., Colgan et al., 2011]. Here we present new data from the Smoke Creek/Buffalo Hills region north of Pyramid Lake (Figure 1) where our three oldest 40Ar/39Ar samples yield ages of 16.81 ± 0.13, 16.64 ± 0.14, and 16.23 ± 0.17 (Table 1). The base

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of the succession is thus slightly older than previous estimates. Notably, it is equivalent in age to the short- lived eruption of Steens Basalt, the oldest unit of the Columbia River Basalt Group [e.g., Camp et al., 2013]. This suggests that the reinitiation of calc-alkaline volcanism was contemporaneous with the initiation of flood-basalt volcanism to the east. At least some of these arc-like lavas and interbedded ash flow tuffs overlie Steens Basalt in the Warner Valley, east of Adel, Oregon (Figure 1) [Dooley, 2010], but there is no evidence that they underlie the Steens succession elsewhere. We therefore suggest that this arc sequence is no older than 16.72 Ma, which is the age of the Steens geomagnetic reversal determined by Jarboe et al. [2008, 2010]. In northwestern Nevada, this middle to late Miocene succession is primarily composed of mafic to intermediate lava flows, 2–10 m thick, with massive flow interiors and scoriaceous flow tops, interbedded in places with volcaniclastic sedimentary rocks and/or felsic air fall pyroclastic rocks that vary from distal to proximal facies. The majority of interbeds found in the lower part of the section are composed of juvenile and/or reworked felsic clasts most likely derived from the ca. 16.5 to 15 Ma High Rock caldera complex that erupted northeast of the Buffalo Hills [Coble and Mahood, 2012, 2016]. Younger, smaller, and more localized felsic rocks are also evident in the upper part of the stratigraphy in the form of small rhyolitic domes, rare ash flow tuffs and thinner, more discontinuous volcaniclastic interbeds (Tables 1 and 2). Although basaltic to andesitic lava flows dominate the sequence in the Nevada-California border region, they become less prominent north of Lakeview, California and into southern Oregon (Figures 1 and 9), where the Miocene succession is instead dominated by younger (ca. 13–8 Ma) tuffaceous sedimentary rocks, air fall pyroclastic rocks, and interstratified welded and nonwelded ash flow tuffs, with only minor basaltic to andesitic lava flows. This succession is well-exposed in Greaser Canyon east of Adel, Oregon (Figure 1), where an ash flow tuff yielded a 40Ar/39Ar age of 13.23 ± 0.16 Ma (Table 1). The earliest lavas (16.7–16.0) were derived from fissure eruptions associated with an extensive swarm of north-south dikes in the Smoke Creek/Buffalo Hills region of northwestern Nevada (Figures 1 and 10). Progressively younger lavas (ca. 16.0–8.0 Ma) were associated with a broad westward shift of eruptive activity into northeastern California, with a contemporaneous change in eruption style to central-vent eruptions and the production of numerous, well-preserved shield volcanoes that range in age from ca. 14 to 8.0 Ma (Figures 10 and 11). Still farther west the style of eruption changed once more after ca. 8 Ma with small-volume lava flows generated from much smaller shield volcanoes and an increasingly larger number of scoria cones (Figure 10). Chemically, the 16.7 to 8.0 Ma volcanic rocks resemble the ca. 30 to 20 Ma volcanic rocks in being mildly alkaline to calc-alkaline, with the majority of intermediate rocks falling in the high-K calc-alkaline field of Peccerillo and Taylor [1976] (Figure 9). They differ, however, in being less variable in overall composition and more mafic in average composition. Whereas basaltic rocks are a very minor component in the ca. 30–20 Ma succession, they comprise 38% of the analyses from the younger sequence, with basaltic andesites forming an additional 38%. The younger 14 to 8.0 Ma lavas derived from shield volcanoes of north-

eastern California are even more uniform in composition, with a limited range SiO2 between 49% and 60% (Figure 9).

6. Volcanic Perturbations Resulting From Cenozoic Dislocation of the Farallon Slab The Cascadia back-arc region of eastern Oregon contains a long record of Cenozoic volcanism marked by noteworthy perturbations in the style and composition of the erupted products. There is little field evidence that back-arc volcanism involved significant crustal extension during the 30 to 8 Ma time interval considered here [Henry, 2008; Colgan et al., 2004, 2006; Meigs et al., 2009; Scarberry et al., 2010; Camp et al., 2015]. The 30 to 20 Ma and 16.7 to 8 Ma periods of high-K calc-alkaline volcanism with elevated LILE/HFSE values suggest the involvement of a hydrated and metasomatized mantle source. In contrast, the 16.7–15.6 Ma period of main-phase tholeiitic flood-basalt volcanism suggests the involvement of a relatively dry mantle source. The 20–16.7 Ma hiatus in back-arc volcanism marks a significant transition separating the earlier period of widespread high-K calc-alkaline volcanism from the simultaneous initiation of tholeiitic volcanism in eastern Oregon and the reinitiation of high-K calc- alkaline volcanism that becomes more localized to the California-Nevada border region.

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Figure 10. (a) Simplified bedrock geology map of lavas and vents of mid-Miocene to recent age in NE California (CA) and NW Nevada (NV), between lat. 40°150 to 41°0N and long. 119°300 to 121°W. White background = mid-Miocene to Holocene volcanic rocks and vents; dark gray background = pre-Miocene rock types; light gray = shallow lakes and playas. Age ranges for lavas and vents are subdivided into three zones separated by gray dashed lines that become younger to the west. Lava flows west of the Black Rock desert in Nevada (NV) were largely derived from fissure eruptions along the Smoke Creek dike swarm from 16.7 to 14 Ma. In adjacent California (CA) to the west, the dominant mode of eruption was from a system of younger shield volcanoes (rounded polygons with or without circular craters) that were active from 14 to 8 Ma. Still younger vents west of Susansville and Eagle Lake, California (small filled circles; typically <8 Ma) are more variable in volcano type with scoria cones becoming more prevalent. Modified from Grose et al. [1990, 2013] and Muffler and Clynne [2015], with additional data from this study. (b) Geochronological data from area outlined in Figure 10a (i.e., between lat. 40°150 to 41°0N and long. 119°300 to 121°W) plotted against longitude. Data derived from Grose et al. [1992, 2013], Sloan et al. [2014], and this study (Tables 1 and 2).

These surface changes in the style, distribution, and composition of volcanism cannot be readily attributed to abrupt changes in tectonic environment. Instead, we believe that they reflect sublithospheric changes in mantle source that we attribute to slab disturbance, dislocation, and dismemberment. We describe here the evolutionary causes and effects of slab disruption consistent with field evidence, seismic data, and a large database of geochemical and geochronological analyses. 6.1. Contemporaneous Calc-Alkaline Volcanism in the Arc and Back-Arc Regions of the Ancestral Cascades From ca. 30 to 20 Ma Late Cretaceous to Paleogene flat-slab subduction in the northern Great Basin was followed by slab rollback that began at ca. 50 Ma, progressing to the southwest across Idaho, southern Oregon, and northern Nevada during Eocene, Oligocene, and Miocene time to produce the “mid-Tertiary volcanic sweep” [Coney and Reynolds, 1977; Christiansen and Yeats, 1992; Dickinson, 2006; Best et al., 2009; Colgan et al., 2011]. Although a general southwestward progression of volcanism seems evident from compiled ages on volcanic rocks (e.g., NAVDAT data) [Walker et al., 2004], the wide and disparate distribution of Eocene to early Oligocene

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Figure 11. Field images of mid-Miocene dikes and vents in northwestern Nevada and adjacent California. (a) Observation Peak shield volcano viewed from the southeast. This volcano is typical in size and shape to several well-preserved shield volcanoes lying south of the Warner Range in northeastern California (see Figure 10). (b) Observation Peak viewed from the north. (c) Small basalt dikes (~1 m wide) exposed at the northern end of the Smoke Creek dike swarm in northeastern Nevada. (d) Oblique view from Google Earth Pro 2016 of three large basaltic dikes, each about 5 m wide and ~1 km long, extending beyond the image to the north. Smoke Creek road in lower left of the image.

volcanic rocks does not delineate a well-defined volcanic arc [du Bray et al., 2014]. In the area considered here, Eocene to early Oligocene volcanic rocks are rare and volumetrically insignificant (Tables 1 and 3). As noted by du Bray et al. [2014], it was not until steeper subduction was reestablished by the rollback process at ca. 30 Ma that volumetrically significant volcanism defined the southern segment of the ancestral arc. The eastern boundary to the southern arc segment is confined by du Bray et al. [2014] to westernmost Nevada. This location is based on a broad change that takes place in the style of volcanism, from mafic, intermediate, and felsic volcanism associated with composite arc volcanoes, to partly contemporaneous (>20 Ma) bimodal, caldera-related volcanism (the ignimbrite flare-up) of the interior andesite-rhyolite assemblage of central Nevada [Ludington et al., 1996]. To the north, the ancestral arc is delimited by an apparent left-lateral offset of the southern and northern segments near the California-Oregon border region (Figure 1) [Cousens et al., 2008; Colgan et al., 2011; du Bray et al., 2014]. Termination of the ancestral southern Cascades arc near the border region of southern Oregon (Figure 1) [Colgan et al., 2011; du Bray et al., 2014] appears reasonable in that it prevents the arc from extending into the Oregon back-arc region of the coeval northern arc segment. The geologic data described here, however, are consistent with the extension of arc-like volcanism into the back-arc region from ca. 30 to 20 Ma. The style of volcanism for the Oligocene to early Miocene (ca. 30–20 Ma) high-K rocks of eastern Oregon differs from the bimodal, caldera-related eruptions of the interior andesite-rhyolite assemblage of central Nevada. Instead, volcanism of mafic, intermediate, and felsic rocks appears to be associated with composite volcanoes [e.g., Mathis, 1993; Scarberry et al., 2010; this study], consistent with the style of volcanism in the southern ancestral Cascades arc [du Bray et al., 2014]. The high-K calc-alkaline character and trace element patterns for the majority of ca. 30–20 Ma volcanic rocks of eastern Oregon closely match those for the ca. 35–18 Ma volcanic rocks of the ancestral southern Cascades (Figures 4 and 5). We are unaware of any chemical discri- minant that can separate these two volcanic regions.

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Figure 12. (a) Distribution of Oligocene to early Miocene volcanism in the ancestral Cascade arc and back-arc regions. The eastward shift in the northern segment of the ancestral Cascades from 35–26 Ma to 25–18 Ma is from du Bray and John [2011]. Location of the southern segment of the ancestral southern Cascades arc in the California-Nevada border region at ca. 30–23 Ma is based on the age constraints of Cousens et al. [2008] and du Bray et al. [2014]. The eastern limit of the ancestral southern Cascades arc in the Oligocene is juxtaposed against Cenozoic calderas of the interior andesite-rhyolite assemblage, separated into the western and central Nevada volcanic fields (WNVF and CNVF, respectively). These calderas are mostly Oligocene in age but vary from 36 to 18 Ma [Best et al., 2013, 2016]. Largely contemporaneous calderas in northern Oregon include the Crooked River (CR; ca. 30 Ma), Wildcat Mountain (WM; ca. 40 Ma), and Tower Mountain (TM; ca 32 Ma) eruptive centers [McClaughry et al., 2009; Seligman et al., 2014]. These northern calderas lie adjacent to a suture zone (connected brown dots) separating early Cenozoic rocks of a greater Siletzia terrane inserted into the Columbia Basin to the north from older oceanic terranes to the south [Schmandt and Humphreys, 2011; Seligman et al., 2014]. Y1 and Y2 are potential migration paths for a long-lived Yellowstone plume [Wells et al., 2014], based on the hot spot reference frames of O’Neill et al. [2005] and Müller et al. [1993], respectively. The adakite hot spot track lies along a geographic regression line derived from the location of Oligocene adakites and rocks of adakite affinity, in contrast to Oligocene locations that lack rocks with the chemical signature of adakite. The curved arrow delineates clockwise rotation of the Siletzia terrane, ~250 km from its original location during terrane accretion at ca. 50 Ma. (b) Schematic cross section proposed for the Farallon plate beneath southern Oregon from ca. 30 to 22 Ma, modified from Coble and Mahood [2012]. The ancestral northern Cascades arc is generated by moderately steep subduction, but calc-alkaline magmatism is also active in the back-arc region due to the release of volatiles during uplift of the Farallon plate by a thermally buoyant Yellowstone plume.

This evidence leads us to conclude that the southern arc segment transitions into the back-arc region of the northern arc segment, while maintaining its high-K calc-alkaline character (Figure 12). Compositional distinc- tions between the northern and southern arc segments have been attributed to variations in the thickness and composition of the crust or to mantle-source enrichment. du Bray et al. [2014], for example, suggested that the high-K rocks of the southern ancestral Cascades were generated beneath thicker, more evolved crust. Putirka et al. [2012] and Putirka and Platt [2012] attributed high-K volcanism in the Walker Lane segment of the southern ancestral Cascades to the presence of enriched subcontinental mantle lithosphere (SCML) associated with relatively small degrees of partial melting. These scenarios, however, have greater difﬁculty explaining high-K volcanism in the back-arc region of eastern Oregon. This region is not underlain by enriched SCML, but rather by thin oceanic lithosphere composed of Paleozoic to Mesozoic accreted terranes. The crustal composition of these terranes varies from low-K tholeiite to primitive calc-alkaline rocks that show little indication of being derived from an enriched or fertile mantle source [e.g., Vallier, 1995]. Instead, we attribute the high-K character (Figure 4) and high LILE/HFSE ratios (Figure 5) of these rocks to an Oligocene-to-early Miocene period of mantle-source enrichment by subduction-related ﬂuids derived from the Farallon slab.

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6.2. Thermal Uplift and Source Melting From ca. 30 to 20 Ma Initial volcanism in the northernsegment of the ancestral Cascades wasconcentrated insouthwest Washington State at 45–40 Ma [Smith, 1993], but it did not become well-established in Oregon until sometime after 35 Ma [SherrodandSmith,2000;duBrayandJohn,2011].Thiseventwascloselyfollowedbyhigh-Kcalc-alkalinevolcan- ism in the Oregon back-arc region from 30 to 20 Ma (Figures 3 and 4). In the absence of significant lithospheric extension,calc-alkalinevolcanismintheback-arcareawouldseemtorequireasignificantchangeinthemelting region, most likely reflecting an increase in temperature and/or fertility of the mantle source. We attribute such a disturbance to regional uplift of the Farallon slab as it interacted with a thermally buoyant Yellowstone plume (Figure 12), as first suggested by Coble and Mahood [2012]. Thermal uplift of the slab provides an explanation for an eastward shift in the volcanic axis of the northern segment of the ancestral Cascades arc in Oregon that began at ca. 25 Ma (Figure 12), coincident with a significant increase in the eruptive volume [du Bray and John, 2011]. Although mantle plumes have low viscosities and little mechanical strength [Stegman et al., 2006; Betts et al., 2012], their high thermal energy and high buoyancy provides them with the capability to uplift slabs while simultaneously decreasing slab strength by thermal diffusion [e.g., Macera et al., 2008; Betts et al., 2012]. We suggest that slab uplift and decompression combined with plume-induced heating had two significant effects in the back-arc region: (1) it resulted in the liberation of

H2O from the slab into the mantle wedge, thus promoting partial melting and calc-alkaline volcanism over a broad region of eastern Oregon from 30 to 20 Ma, and (2) it decreased slab strength by thermal erosion that began at ~30 Ma, leading to slab dislocation and dismemberment that began at ca. 20–17 Ma. The recent field and plate motions studies of McCrory and Wilson [2013] and Wells et al. [2014] support the original contention of Duncan [1982] for a long-lived plume tail that provided the source for hot spot volcanism and generation the Siletzia oceanic plateau. Wells et al. [2014] concluded that accretion of Siletzia in Oregon was completed between 51 and 49 Ma. This was followed by overriding of the Yellowstone hot spot (plume tail) by the North American plate at 42 Ma, with probable hot spot volcanism extending into the Cascade fore-arc region from 42 to 35 Ma. The flux of a feeding plume tail should result in significant mass accumulation and flow displacement of plume material beneath the subducting slab between the time it was overridden at ca. 42 Ma until plate rupture and the initiation of flood-basalt volcanism at ca. 17 Ma. Location of the hot spot track during this time is difficult to determine from the surface geology but can be estimated from plate-reconstruction models. Wells et al. [2014] used the reconstruction model of Seton et al. [2012], but they also noted that the track location will vary depending on which hot spot reference frame is used. In Figure 12, we reproduce the estimated hot spot tracks Y1 and Y2 from Wells et al. [2014], created from the reference frames of Müller et al. [1993] and O’Neill et al. [2005], respectively. Seligman et al. [2014] suggested that the earliest eruptions of plume-related volcanism in the back-arc region correspond with the alignment of the Wildcat Mountain (40 Ma), Crooked River (32–28 Ma), and Tower Mountain (32 Ma) eruptive centers lying north of Y1 in Figure 12. These eruptions sites fall along an orientation adjacent to the suture zone separating the greater Siletzia terrane inserted into the Columbia Embayment to the north from older oceanic terranes to the south [Schmandt and Humphreys, 2011]. Seligman et al. [2014] argue that the eruptive products along this alignment may have been derived from interactions between the Yellowstone plume and delamination of the overlying crust, as suggested by whole-rock trace element 18 data from basalts and rhyolites, and δ O and εHfi data from zircons in the rhyolitic rocks. We do not deny the possibility of plume material escaping through a rupture in the Farallon plate to trigger early Oligocene eruptions adjacent to the greater Siletzia suture zone, as advocated by Seligman et al. [2014]. Farther south, however, the evidence is more consistent with the back-arc eruption of Oligocene to early Miocene calc-alkaline rocks (~30–20 Ma) derived from a wet-mantle source with no direct evidence of a genetic relationship with plume-derived melts until ca. 16.7 Ma. 6.2.1. Adakite Generation Plume impingement beneath an oceanic slab appears to be capable of generating a broad-based melt zone of hydrated mantle above the site of slab uplift and an equally broad area of surface calc-alkaline volcanism. In addition to this extensive zone of mantle melting, continued heating and uplift could also result in direct melting of oceanic crust in the subducting slab. The lower solidus temperature of a mafic source (compared with ultramafic sources at similar mantle depths) makes oceanic crust particularly susceptible to melting in an environment of plume advection and elevated mantle temperatures [e.g., Yaxley, 2000].

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Building on the earlier work of Kay [1978], Defant and Drummond [1990] ﬁrst introduced the term adakite for andesite-dacite-rhyolite rock types (ADRs) produced as direct melts of the basaltic portion of a subducted oceanic plate. In the typical melt zone for adakite, the basaltic oceanic crust has been metamorphosed to eclogite or amphi- bolite, a source composition that gives adakite a distinct set of compositional characteristics. Pure adakites appear to

be characterized by SiO2 typically ≥56%, Al2O3 ≥ 15%, and MgO usually <3% (rarely above 6%); relative to island – Figure 13. Y versus Sr/Y plot of ca. 30 20 Ma volcanic rocks with normal- arc ADRs, they have lower HREEs con- ized SiO2 contents >55%. Fields for “normal” calc-alkaline rocks and adakites are from Defant and Drummond [1993] and Castillo et al. [1999]. centrations, lower Y (18 ppm), and Defant and Drummond [1993] define adakitic compositions as andesitic higher Sr (rarely <400 ppm) and there- rocks with Sr/Y ratios >20, which suggests that a significant number of fore higher Sr/Y ratios [Defant and the late Oligocene to early Miocene rocks might contain an adakitic Drummond, 1990; Castillo, 2006]. component derived from slab melting of oceanic crust. The Sr/Y ratio is a widely used discrimi- nant to separate adakites from normal arc ADRs. During partial melting of a garnet-bearing and/or amphibole-bearing mafic source, at depths below the stability field of plagioclase, Sr will behave incompatibly and concentrate into the fluid phase, but Y will be retained in the bulk solid because of its compatibility in garnet and amphibole. The resulting high Sr/Y ratios in adakite melts are opposite to those found of peridotite partial melts that differentiate toward higher Y and lower Sr/Y ratios typical of island arc ADRs. We use the Sr/Y versus Y diagram to differentiate the ca. 30–20 Ma volcanic rocks in the southern Oregon border region into those with adakite compositions and those with normal calc-alkaline ADR compositions

(Figure 13). We eliminate silica-poor compositions by restricting analyses to those with normalized SiO2 values ≥55%. A large but fairly isolated group of analyses with Sr/Y values between 25 and 50 fall either within the adakite field or in a transitional region between the adakite field and the field of normal island arc rocks of Defant and Drummond [1993]. We suggest that these rocks could be derived from various mixing proportions of pure adakite melts with differentiated melts from a basaltic source derived from the hydrated mantle. These rocks with moderately high Sr/Y include all samples at Adel and Lone Juniper Springs and the majority of andesitic samples from lower Steens Mountain and the Warner Range (Figure 3). Rocks with the strongest adakite signatures (Sr/Y ≥ 60 and Y ≤ 18) are confined to two geographic regions, the Salmon River volcanic field of western Idaho and the Pine Forest Range of northernmost Nevada (Figures 3 and 13). Adakite-like rocks may well form from a variety of processes. Recent compilations of chemical data from magmatic arcs, for example, suggest a correlation between Sr/Y and crustal thickness attributable to variations in partial melting of lower crust and/or crystal fractionation of mafic magmas in the lower crust, where plagioclase is unstable and garnet or amphibole is a stable phase [Chiaradia, 2015; Chapman et al., 2015; Putirka, 2017]. An example of this correlation comes from the Nevadaplano orogenic plateau of Nevada and western Utah. Here crustal thickening in the Mesozoic culminated at 85–90 Ma to a maximum thickness of 55–65 km with coeval adakite-like rocks having median Sr/Y values of ~45; thickening was followed by basin and range extension, mostly in the Miocene, where crustal thickness was reduced to 30–40 km and median Sr/Y values were correspondingly lowered to ~25 [Chapman et al., 2015]. The southern Oregon border region lies north of the Nevadaplano where the accretion of thin oceanic terranes was associated with only a moderately thickened crust. In this area of minor basin and range extension (~20%), the median crustal thickness is ~30 km [Lerch et al., 2007; Gashawbeza et al., 2008; Eagar et al., 2011], with no evidence of unusually thick crust (>40 km) prior to basin and range extension Lerch et al. [2007]. Here the median Sr/Y value is 35 for the combined lava sequences at Adel, Lone Juniper Springs, lower Steens

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Mountain, and the Warner Range and 75 for the combined lava sequences in the Pine Forest Range and the Salmon River volcanics. Unlike the Nevadaplano, lower crustal processes of partial melting and crystal fractionation cannot readily account for these high median values in significantly thinner pre-Miocene crust. Assuming an environment of high mantle temperatures and high mantle flux during plume-slab interaction, we believe that a more realistic model of adakite generation in this area is from the partial melting of oceanic crust from the Farallon slab. 6.2.2. Yellowstone Adakite Hot Spot Track The northwesternmost ca. 30–20 Ma outcrop areas at Abert Rim, Hart Mountain, Coleman Hills, and Rabbit Hills lack rocks of adakite affinity; instead, they are composed entirely of rocks with significantly lower Sr and higher Y values typical of the normal range of island arc ADRs. There appears to be a distinct east-northeast-trending line of demarcation separating the outcrops of normal ADRs from those that contain adakites or rocks of adakite affinity. The adakite exposures are not only spaced over a broad area but also lie along a geographic regression line of east-northeast trend (Figure 12). We speculate that they were produced by the melting of Farallon oceanic crust above the Yellowstone hot spot between ca. 30 and 20 Ma. Such an interpretation is consistent with a greater degree of heating created by the thermal impact of the mantle plume tail [McKenzie and Bickle, 1988; Campbell, 2005]. We refer to this belt as the Yellowstone adakite hot spot track. Although it is not manifested as a surface trend of age-progressive volcanism, its orientation parallels North American plate motion along a trend similar to that of the younger Yellowstone-Snake River Plain hot spot track to the east but offsets very slightly to the north. Its axial trend lies midway between and parallel to the two hypothetical Yellowstone hot spot tracks calculated by Wells et al. [2014], Y1 and Y2 in Figure 12. The strongest adakite signatures in the Pine Forest Range and in the Salmon River volcanics are located in the central region and eastern margin of the ca. 16.7–15 Ma Nevada-Columbia Basin magmatic belt, respectively. About 80 km north of the Pine Forest Range, the axial trend of the adakites bisects the site of the initial flood-basalt eruptions at Steens Mountain. This evidence suggests that significant melting of oceanic crust prior to ca. 20 Ma occurred at the eventual site of slab rupture, followed by plume ascent, plume melting, and the initiation of tholeiitic flood-basalt volcanism at ca. 16.7 Ma.

6.3. Slab Rupture, the Initiation of Flood-Basalt Volcanism, and the Rejuvenation of Calc-Alkaline Volcanism at ca. 16.7 Ma Calc-alkaline volcanism in the back-arc region of eastern Oregon and northernmost Nevada waned dramati- cally at 22 Ma and ceased at 20 Ma. We concur with Coble and Mahood [2012] who attribute the ensuing 5 Myr hiatus in back-arc volcanism to the cessation of corner flow resulting from slab uplift. This was followed by eventual rupture of the Farallon slab, thus allowing rapid rise and decompressional melting of hot plume mantle with attendant flood-basalt volcanism and crustal intrusion of magmatic rocks along the entire length of the Nevada-Columbia Basin magmatic belt from ca. 16.7 to 15 Ma (Figure 13) [Liu and Stegman, 2012; Camp et al., 2015]. Although the earliest products of plume ascent at 16.7–16.5 Ma indicate that slab rupture was focused beneath the accreted terranes of southeastern Oregon (Figure 14b), plate reconstruction models suggest that the hot spot center may have arrived farther east beneath the cratonic lithosphere of western Idaho at ca. 16.5 Ma [Jordan, 2005; Pierce and Morgan, 2009; Wells et al., 2014]. This apparent discrepancy has been explained by one or a combination of factors: (1) westward diversion of the accumulated bulk of plume material against the inclined slab [Pierce and Morgan, 2009], (2) westward deflection of the plume tail by upper mantle flow [Smith et al., 2009], (3) bending of the plume tail due to plate interaction [Wells et al., 2014], and/or (4) westward deflection of buoyant plume material across the cratonic boundary as it ascended preferentially to shallower depths beneath thinner oceanic lithosphere [Jordan, 2005]. In a nonplume scenario, Liu and Stegman [2012] attribute slab rupture to the breaking of a weakened hinge in the plate beneath southeast Oregon, with model simulations that show a rapid propagation of tearing to the north and south, coincident with the Nevada-Columbia Basin magmatic belt. In the plume-uplift model, an abrupt downward bend is likely to develop on the eastern side of the uplifted slab (Figure 14b), similar to that envisioned by Coble and Mahood [2012]. Camp et al. [2015] note that such a bend is analogous to the weakened hinge of Liu and Stegman [2012] but weakened further by plume heating, thus generating a site of rupture that would propagate in much the same way as modeled by Liu and Stegman [2012]. This scenario is also consistent with the modeling results of Macera et al. [2008], who show that slab strength will decrease by

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Figure 14. (a) Distribution of middle to late Miocene 17–8 Ma volcanic rocks in the ancestral Cascades arc and back-arc regions. The distribution of 17–10 Ma rocks from the northern segment of the ancestral Cascades is from du Bray and John [2011]. Locations for 17–8 Ma radiomentric ages for the southern segment of the ancestral Cascades are from du Bray et al. [2014], this study (Tables 1 and 2), and the U.S. National Geochronological Database (http://mrdata.usgs.gov/geochronology/geochron.html). Tholeiitic magmatism in the back-arc area from 16.7 to ca. 15.5 Ma is denoted by mafic dike swarms of the main-phase Columbia River flood basalts, related mafic dikes of the northern Nevada Rift, and large midcrustal mafic dikes delineated by prominent aeromagnetic anomalies [i.e., Glen and Ponce, 2002]. Together, these mafic intrusions define the central portion of the Nevada-Columbia Basin magmatic belt [Camp et al., 2015]. The cratonic boundary is based largely on location of the initial 87Sr/86Sr isopleth of Pierce and Morgan [2009]. (b) Schematic cross-section A-A0 at ca. 16.5 Ma, modified from Coble and Mahood [2012]. The uplifted portion of the Farallon slab is marked by tears to the west and to the east, consistent seismic imagery [see for example Liu and Stegman, 2012]. See text for discussion.

heating above a plume head by 40% to 80% in 10 million years, which they conclude is a reasonable interval to reach the conditions of slab rupture and detachment. In Figure 15, the Farallon slab is resolved as a high-velocity feature beneath northern Nevada and Utah at 300– 800 km depth, bounded by a slab gap to the west [Schmandt and Humphreys, 2010; Obrebski et al., 2010]. It seems probable that this gap was derived from subduction-related tearing of the uplifted slab, thus allowing the rejuvenation of arc volcanism to begin at 16.7 Ma (Figures 14b and 15). The termination of slab subduction occurs at a shallow depth of ~350 km beneath northern Nevada, at depths that vary from 350–400 km beneath southern Oregon and at ~160 km beneath northern Oregon [Burdick et al., 2008; Xue and Allen, 2007; Sigloch et al., 2008; Obrebski et al., 2010; Schmandt and Humphreys, 2010; Tian et al., 2011]. Based solely on plate motion estimates from Gripp and Gordon [1990, 2002], these depth ranges account for subduction since 17–15 Ma and 6–8 Ma, respectively [Xue and Allen, 2010; Schmandt and Humphreys, 2010]. A deeper slab segment, similar to that beneath Nevada and Utah, appears to be missing north of the Nevada-Oregon border. The shallow termination of the subducting slab, and lack of its deeper continuation, provides evidence of a seismic hole beneath eastern Oregon that others have attributed to breakup of the Farallon plate by plume arrival [Xue and Allen, 2007, 2010; Obrebski et al., 2010; Schmandt and Humphreys, 2010; Darold and Humphreys, 2013]. The eastward extent of the seismic hole is partly ambiguous, but best delineated in southeastern Oregon as an east-northeast-trending, 200 km wide slab window that extends into western Idaho (Figure 15) [Sigloch et al., 2008]. The northern boundary of this feature lies adjacent to a remnant of the Farallon slab that appears to have been delaminated during arrival of the Yellowstone plume [Darold and Humphreys, 2013]. The southern boundary of this hole lies adjacent to the adakite hot spot track, and its axis lies parallel to plate motion (SG line, Figure 15), midway between the Y1 and Y2 hypothetical hot spot tracks of Wells et al. [2014].

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Figure 15. Mantle profiles A-A0 and B-B0 from the Vp model of Schmandt and Humphreys [2010]. Profile locations are dis- played on a map slice at 350 km depth. The color represents velocity perturbations of ±3% relative to the mean velocity, with blue representing the higher velocities. The boundary of the slab window centered in eastern Oregon is open to interpretation of the geophysical data. The outline shown here is derived largely from Sigloch et al. [2008], but with modifications from Schmandt and Humphreys [2010] and Tian and Zhao [2012]. SG, tear in the Farallon plate from Sigloch et al. [2008].

According to Sigloch et al. [2008], the slab window reproduced in Figure 15 is part of a more extensive tear or slab gap extending through Montana into southern Saskatchewan (SG line, Figure 15). They suggest that this early Cenozoic feature has operated as a self-perpetuating tear close to the trench and parallel to plate motion, but without any obvious surface continuation on the Juan de Fuca seaﬂoor. We suggest instead that (1) this older segment of the SG line beneath central Idaho and Montana is a linear zone of early Cenozoic slab disruption created by the active Yellowstone plume before it was overridden by the North American plate at 42 Ma, and (2) the younger segment in southeastern Oregon is largely the result of plume-triggered uplift and rupture of the subducted slab from ca. 30 to 17 Ma. Slab breakup and the rejuvenation of more normal subduction in the Nevada-California-Oregon border region at 16.7 Ma were immediately followed by a period of slab rollback. In northwestern Nevada and adjacent California, rollback is expressed by the westward migration of arc volcanism from the Smoke Creek dike system to its current position near Lassen Peak, California. We use new and published isotopic ages to estimate a migration rate of 7.8 km/Ma across the arc axis at lat. 40°300N (Figure 10), which is nearly identical to the 8 km/Ma estimate of Henry et al. [2009b] for the entire arc axis since the mid-Miocene.

7. Subduction Component in Steens Basalt Evidence described herein suggests that the eruption of Steens Basalt in the mid-Miocene was preceded by an Oligocene to early Miocene period of mantle-source enrichment. It is therefore conceivable that these early flood basalts could contain a melt component from this enriched mantle source. Steens Basalt differs from all other formations in the Columbia River Basalt Group in showing a progression of differentiation from early tholeiitic flows to later, more evolved flows of mildly alkali or transitional calc-alkaline compositions [Johnson et al., 1998]. Camp et al. [2013] noted that such a progression could derive from protracted fractional

crystallization combined with multiple cycles of maﬁc recharge. They also noted that the high K2O compositions of the most evolved upper Steens lavas of trachybasalt and basaltic trachyandesite compositions

require addition of K2O, perhaps derived from the liberation of LILEs from a crustal assimilant.

Here we suggest the alternative explanation that these K2O-rich lavas were derived from partial melting of the hydrated mantle enriched in the LILEs during the earlier period of slab uplift (see section 6.2). It seems likely that the ascent of hot plume material and plume melts through the region of slab rupture at ca. 20–17 Ma (Figure 14) would have encountered this water-rich, readily fusible metasomatized mantle, thus providing a subduction-like source component for Steens Basalt. In such a case, a crustal component may

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not be necessary to explain the high-K character of trachybasalts and basaltic trachyandesites found in upper Steens Basalt. Taken as a whole, Steens Basalt shows only minor chemical evidence for a subduction component in the mantle and/or contamination by accreted terrane crust [Wolff and Ramos, 2013]. On the other hand, normalized trace element profiles for the most evolved lavas of upper Steens Basalt show patterns consistent with a subduction component. This is evident in the remarkably similar patterns for trachy- Figure 16. Incompatible trace element diagrams normalized to the primi- basalts and basaltic trachyandesites tive mantle of Sun and McDonough [1989] for upper Steens trachybasalts from upper Steens Basalt compared and basaltic trachyandesites compared with the pattern in gray derived with those from the southern ancestral fl – from basaltic trachyandesite ows of the middle-to-late Miocene (16.7 Cascades (Figure 16). These two lava 8 Ma) ancestral southern Cascades in northeastern California. Selected groups do not represent primary mag- samples include only basaltic trachyandesites with SiO2 compositions between 53 and 60% and a complete set of ICPMS trace elements. These mas, but rather magmas that may have include seven samples of the ancestral Cascades (gray pattern: 08-C-12, undergone some degree of fractional JC08-WR413, WR08-AE08, JC08-WR414, WR07-AE08, WR07-AE16, and crystallization ± (mafic) recharge. WR07-AE43) from Colgan et al. [2011] and du Bray et al. [2014] and four Nevertheless, the high LILE/HFSE ratios, upper Steens samples (JS-32, JS-57, JS-65, and JS-72) from Wolff et al. [2008] and Johnson et al. [1998]. Pb peaks, and Nb-Ta troughs of both groups are diagnostic of a metasomatized mantle source (Figure 16). The high-K concentrations, common to both, are more likely the result of source enrichment rather than crustal contamination. The Steens lavas differ from those of the southern ancestral Cascades only in being somewhat more enriched in the HFS elements Zr, Ti, and Y, and in the middle-to-heavy rare earth elements Sm, Eu, Dy, and Yb (Figure 16). These variations probably result from source differences, and/or variations in the degrees of partial melting/fractional crystallization, and/or variations in the composition of fractionating phases. The contemporaneous initiation of Steens Basalt volcanism and the rejuvenation of subduction-related volcanism at ca. 16.7 Ma (Table 1) suggest that these lava sequences might be interstratified where they occur in close proximity. Although some arc-like lavas overlie Steens Basalt near the arc axis [e.g., Dooley, 2010], most of the basaltic trachyandesites interbedded in upper Steens Basalt lie too far east of the arc axis to be directly associated with subduction. For that reason, it seems more likely that these lavas were derived instead from plume-triggered melting of the hydrated mantle that was previously enriched during the earlier period of slab uplift.

8. Calc-Alkaline to Mildly Alkaline Volcanism in Easternmost Oregon The easternmost volcanic region of the back-arc area contains several Miocene exposures of calc-alkaline, mildly alkaline to transitional lavas of intermediate composition (Figure 17). Many of these lava flows were originally described as “Steens-Type” Basalt by Hart and Carlson [1985]. In some places they sit directly above flows of upper Steens Basalt, with 40Ar/39Ar ages that range from 16.5 to 15.2 Ma [Brueseke et al., 2007; Benson et al., 2017]. These lavas are present at the top of the lava successions in the Steens Mountain-Sheepshead Mountain area of southeastern Oregon (SSM, Figure 17) [Johnson et al., 1998; Brueseke et al., 2007] and in the border region near the Whitehorse Caldera in Oregon to the Bilk Creek Mountains in Nevada (WCBC, Figure 17) [Mankinen et al., 1987; Benson et al., 2017]. These younger lavas are similar in age and composition to lavas found farther east in the Santa Rosa-Calico volcanic field (16.5–14 Ma) [Brueseke and Hart, 2009] of northern Nevada and farther north in the Strawberry Volcanics of eastern Oregon (>16–12 Ma) [Steiner and Streck, 2013]. Still younger intermediate

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Figure 17. Distribution of late Miocene calc-alkaline to mildly alkaline volcanic rocks of eastern Oregon unrelated to subduction. The genesis of these rocks may be associated with a foundering Farallon slab and/or crustal extension and decompressional melting of lithospheric mantle that was previously enriched by Farallon slab-derived fluids on the eastern side of slab uplift from ca. 30 to 20 Ma. The distribution of 17–10 Ma rocks from the northern segment of the ancestral Cascades is from du Bray and John [2011]. Locations for 17–8 Ma radiomentric ages for the southern segment of the ancestral Cascades are denoted by filled and open circles, from the same sources as in Figure 14. KS, Keeney Sequence; OM, Owhyee Mountains; PRV, Powder River Volcanics; SMV, Strawberry Mountain Volcanics; SSM, Steens Mountain and Sheepshead Mountain; SRC, Santa Rosa-Calico volcanic field; WCBC, area from Whitehorse caldera to the Bilk Creek Mountains; USB, area of evolved trachybasalt and trachybasaltic andesite lava flows of Upper Steens Basalt.

calc-alkaline to mildly alkaline lavas, mostly between 14 and 10 Ma, include those in the Keeney Sequence exposed in the Oregon-Idaho graben [Hooper et al., 2002] and in the south fork drainage basin of the Malheur River [Camp et al., 2003], latites and related lavas in the Owyhee Mountains in western Idaho [Ekren et al., 1981], and the Powder River Volcanics in the vicinity of the Baker and La Grande grabens of northeastern Oregon [Bailey, 1990]. All of these younger volcanic fields occur along a north-south zone of crustal extension and dike intrusion defined as the La Grande-Owyhee eruptive axis of Ferns and McClaughry [2013]. Many of these lavas that postdate the bulk of the Steens Basalt succession are high-K calc-alkaline to transitional lavas similar to trachybasalts and basaltic trachyandesites found interbedded in upper Steens Basalt, while others are more typical of low-K calc-alkaline andesites. Brueseke and Hart [2009] and Steiner and Streck [2013] show that lavas of intermediate composition in the Santa Rosa-Calico and Strawberry volcanic fields could have been derived from mixing of tholeiitic, Steens-like magmas with more felsic components. However, the youngest of these arc-like lavas are spread throughout most of the La Grande-Owyhee volcanic axis where they erupted several millions of years after the main-phase flood basalt eruptions. Hooper et al. [1995] suggested a more regional model where their calc-alkaline signature was derived from crustal extension and decompressional melting of lithospheric mantle enriched by a previous event. We suggest that this enrichment event occurred during the earlier period of Farallon slab uplift and that calc-alkaline volcanism in eastern Oregon, mostly from 14–10 Ma, was therefore derived from the same mantle source as the more widespread volcanism above the uplifted slab from ca. 30 to 20 Ma. We speculate that this period of volcanism may have been associated with the foundering and sinking of slab vestiges marking final demise of the uplifted slab against the cratonic margin (Figure 17).

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Figure 18. Chronology of volcanism in arc and back-arc regions of the ancestral Cascades under changing conditions of plume-plate interaction.

9. Conclusions The Yellowstone hot spot has had a profound effect on the volcanic and tectonic evolution of the Pacific Northwest. Several studies have linked hot spot arrival to (1) the generation and accretion of the Siletzia oceanic plateau [Duncan, 1982; McCrory and Wilson, 2013; Wells et al., 2014], (2) flood-basalt volcanism of the Columbia River Basalt Group [Brandon and Goles, 1988; Hooper and Hawkesworth, 1993; Camp and Ross, 2004; Hooper et al., 2007; Camp and Hanan, 2008; Wolff et al., 2008; Camp, 2013; Wolff and Ramos, 2013], (3) bimodal volcanism along the Snake River Plain-Yellowstone hot spot track [Pierce and Morgan, 1992, 2009; Hanan et al., 2008; Shervais and Hanan, 2008; Smith et al., 2009; Jean et al., 2014], (4) midcrustal and near-surface dike intrusion associated with the northern Nevada rift [Zoback and Thompson, 1978; Zoback et al., 1994; Glen and Ponce, 2002; Camp and Ross, 2004], and (5) large-scale crustal stretching in the northern and central basin and range [Parsons et al., 1994; Pierce and Morgan, 2009; Camp et al., 2015]. The geologic evidence for these genetic connections is substantial and well-documented. In contrast, there has been little data from the surface geology that documents the effect of plume arrival on the subducting Farallon slab. This presumed interaction instead has been based largely on inverse interpretations from seismic profiles of contemporary mantle [e.g., Xue and Allen, 2007, 2010; Obrebski et al., 2010; Humphreys and Schmandt, 2011]. Here we supplement these seismic investigations with geochemical, geochronological, and field data used to define volcanic perturbations in the arc and back-arc region of the ancestral Cascades (Figure 18). These perturbations form the basis of our conclusions, which we summarize chronologically in the following sequence of events.

9.1. Uplift of the Farallon Slab Recent studies add credence to the idea that the Yellowstone hot spot is a long-lived feature that resided offshore in Paleocene and Eocene time, providing a mantle source for basaltic volcanism on the Farallon plate that produced the Siletzia oceanic plateau from ca. 56 to 49 Ma. Accretion of Siletzia into the Columbia embayment of Washington and Oregon at ca. 50 Ma was followed by overriding of the hot spot that began

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at ca. 42 Ma [McCrory and Wilson, 2013; Wells et al., 2014]. This event shifted subduction westward, thus initi- ating calc-alkaline volcanism in the northern segment of the ancestral Cascades arc at ca. 45–40 Ma in Washington [Smith, 1993] and volumetrically significant volcanism in Oregon at ca. 35 Ma [Sherrod and Smith, 2000; du Bray and John, 2011]. The low-K calc-alkaline composition of rocks from the northern ancestral Cascades differs markedly from the high-K character of rocks from the southern segment of the ancestral Cascades. From ca. 35 to 18 Ma, the axis of the southern arc segment was aligned with the California-Nevada border region, striking northward into the broader back-arc region of the northern arc segment where a coeval period of high-K volcanism occurred in eastern Oregon from ca. 30 to 20 Ma. This volcanic perturbation into the back-arc region is attributed to uplift of the Farallon slab as it interacted with the thermally buoyant Yellowstone plume. Uplift-induced decompression and plume-induced heating resulted in the liberation volatiles from the slab into the mantle wedge, thus promoting partial melting of a hydrated mantle and calc-alkaline volcanism over a broad region of eastern Oregon. The high-K character and relatively high LILE/HFSE ratios of the back-arc lavas are attributed to source enrichment by these slab-derived fluids. We argue that the center of plume impingement on the Farallon slab is expressed at the surface by 30–20 Ma volcanic successions of adakites and rocks of adakite affinity in southeastern Oregon. These adakite sequences lie along a broadly defined trend parallel to model calculations of the Yellowstone hot spot track (Figure 12). Petrochemical and crustal thickness considerations suggest that these adakites are unlikely to be produced by fractional crystallization in the lower crust or by partial melting of the lower crust but instead by melting of oceanic crust from the Farallon slab. Location of the adakite hot spot track is attributed in part to high thermal flux above the Yellowstone plume tail.

9.2. Rupture of the Farallon Slab The waning and cessation of back-arc volcanism from 22 to 20 Ma is attributed to the termination of corner ﬂow due to continued slab uplift [Coble and Mahood, 2012]. The greatest amount of uplift was near the thermal center of plume impingement in southeastern Oregon, which might well have resulted in juxtaposition of the slab against lithosphere of the overlying plate. Rupturing of the slab between ca. 20 and 17 Ma occurred in two separate, north-trending zones. The eastern rupture allowed for the adiabatic ascent and melting of a dry, plume-contaminated mantle source underlying the uplifted slab, thus generating the eruption of Steens Basalt at 16.7 Ma. Continued motion of the Farallon slab west of the uplifted segment resulted in a contemporaneous north-trending rupture near California-Nevada border, thus resulting in the simultaneous rejuvenation of subduction-related volcanism at 16.7 Ma along the Smoke Creek dike swarm of northwestern Nevada. This north-south rupture continues into Oregon where the subducting slab terminates at 350 to 160 km depth. We concur with Xue and Allen [2010] who concluded that the large, deep-seated remnant of the Farallon slab beneath northern Nevada (Figure 15) is a broken relict that sank and continues to sink due to its negative buoyancy. This previously uplifted and broken remnant is not present in eastern Oregon, its place occupied instead by a seismic hole that others have attributed to interaction with the Yellowstone plume [Xue and Allen, 2007, 2010; Obrebski et al., 2010; Schmandt and Humphreys, 2010; Darold and Humphreys, 2013]. This large gap was created through a young, thin and warm Farallon slab that was subject to additional thermal input from constant ﬂux of the underlying plume tail. The result was a prolonged period of uplift and heating that began at ca. 30 Ma. Mechanical dislocation culminated in rapid tearing of the slab from ca. 17 to 15 Ma and possibly by the foundering and sinking of slab segments adjacent to the cratonic margin of eastern Oregon from ca. 16 to 10 Ma.

9.3. Rollback of the Farallon Slab Although calc-alkaline volcanism was widespread in the back-arc region of eastern Oregon from 30 to 20 Ma, its reemergence at 16.7 Ma was restricted to the much smaller area of northwest Nevada. This event marks the rejuvenation of calc-alkaline volcanism in this portion of the southern segment of the ancestral Cascades after a3–5 million-year hiatus, and it was also contemporaneous with the initiation of tholeiitic ﬂood-basalt eruptions in southeastern Oregon. Rollback of the Farallon slab since 16.7 Ma is evident in the new and compiled isotopic age data for a broad westward shift in eruptive activity at a rate of 7.8 km/Ma (Figure 10), from the Smoke Creek dike swarm of northwestern Nevada to its present position near Lassen Peak, California.

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