Detailed three-dimensional shear wave velocity structure of the northwestern United States from Rayleigh wave tomography

Earth and Planetary Science Letters 299 (2010) 273–284

Contents lists available at ScienceDirect

Earth and Planetary Science Letters j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / e p s l

Detailed three-dimensional shear wave velocity structure of the northwestern United States from Rayleigh wave tomography Lara Wagner a,⁎, Donald W. Forsyth b, Matthew J. Fouch c, David E. James d a

Department of Geological Sciences, University of North Carolina-Chapel Hill, Chapel Hill, NC, United States Department of Geological Sciences, Brown University, Providence, RI, United States School of Earth and Space Exploration, Arizona State University, Tempe, AZ, United States d Department of Terrestrial Magnetism, Carnegie Institution of Washington, Washington, DC, United States b c

a r t i c l e

i n f o

Article history: Received 1 June 2010 Received in revised form 24 August 2010 Accepted 3 September 2010 Editor: Y. Ricard Keywords: High Lava Plains Yellowstone Snake River Plains surface waves tomography shear wave velocity Pacific Northwest

a b s t r a c t Since the mid-Miocene, the northwestern United States has experienced extensive flood basalt volcanism, followed by the formation of two time-progressive tracks of silicic volcanism: the Yellowstone/Snake River Plains (YSRP) and the High Lava Plains (HLP). The YSRP track progresses towards the northeast, parallel to North American plate motion, and has therefore often been attributed to a deep mantle plume source. However, the HLP track progresses to the northwest over the same time frame in a direction not consistent with any regional plate motion. The causes of the mid-Miocene flood basalts and the tracks of the YSRP and HLP are a matter of ongoing debate. We present results of Rayleigh wave phase velocity inversions and inversions for 3-D shear wave velocity structure of the northwestern United States using data collected from the High Lava Plains seismic experiment and the EarthScope USArray Transportable Array (TA). The large number of stations used in these inversions allows us to show an unprecedented level of detail in the seismic velocity structures of this tectonically complex area. Our velocity images indicate that low S-wave velocities in the uppermost mantle do not well match the track of HLP volcanism. While at the surface the Newberry caldera appears to anchor the NW end of the HLP hotspot track, the seismic results show that it lies in a separate, north–south trending low velocity band just east of the Cascades that is distinct from the main HLP trace. The ultra-low S-wave velocities beneath the YSRP track extend locally to at least 175 km depth and are by far the most prominent seismic anomalies in the region. Along axis, the YSRP hotspot track is characterized by a discrete low velocity channel in the upper mantle that shallows, narrows and intensifies to the northeast, but then deepens rapidly to the north beneath Yellowstone. The shallowing of the low velocity anomaly to the northeast is consistent with a stationary heat source beneath a moving plate, coupled with Basin and Range extension and heating. © 2010 Elsevier B.V. All rights reserved.

1. Introduction The northwestern United States has been the site of extensive and varied volcanic activity from the mid-Triassic to present as a result of a remarkably broad spectrum of tectonic drivers. The eastern portion of the northwestern United States comprises the 3.2–1.8 Ga Wyoming Craton (Fig. 1) with ages generally younging to the west (Foster et al., 2006; Wooden and Mueller, 1988). The 87Sr/86Sr = 0.706 line that appears to delineate the margin between Proterozoic North America and the younger provinces runs roughly north–south along the Oregon–Idaho border, cutting to the northwest into Washington State (Ernst, 1988; Fleck and Criss, 1985). During the mid-Triassic to early Jurassic, the west dipping Wallowa arc and east dipping Olds Ferry arc ⁎ Corresponding author. Department of Geological Science, CB# 3315, UNC-Chapel Hill, Chapel Hill, NC 27599, United States. Tel.: + 1 919 966 4714; fax: + 1 919 966 4519. E-mail address: [email protected] (L. Wagner). 0012-821X/$ – see front matter © 2010 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2010.09.005

collided with each other upon the closure of an oceanic basin, forming the intervening Baker terrane. These three terranes, collectively known as the Blue Mountain Province (BMP) collided with Nevada during the mid to late Jurassic, with extensive shortening continuing into the early Cretaceous (Dorsey and LaMaskin, 2007; Schwartz et al., 2010). By the mid Cretaceous, dextral slip along the western Nevada shear zone and the western Idaho shear zone translated the BMP to its current location in northwestern Oregon (Wyld and Wright, 2001; Wyld et al., 2006). This was then followed by the intrusion of the Idaho Batholith into both the BMP and the Precambrian crust to the east during the mid to late Cretaceous, with magmatic activity continuing into the early Tertiary (Fleck and Criss, 1985). By mid Eocene (~ 48 Ma), arc volcanism jumped west to approximately its current location with the accretion of the Coast Range Basalt Province (Madsen et al., 2006), also sometimes referred to as the Siletzia terrane (Humphreys, 2009; Wells et al., 1998). Post Laramide Basin and Range extension began ~ 35 Ma in central Nevada and progressed

274

L. Wagner et al. / Earth and Planetary Science Letters 299 (2010) 273–284

Fig. 1. Geologic map of the northwestern United States: the brown shaded area shows the extent of the Columbia River (CRB) and Steens Mountain (SB) flood basalts (Camp and Ross, 2004). The white shaded area represents the Idaho Batholith (IB) (Fleck and Criss, 1985). Also shown are the locations of the Blue Mountain Province (BMP), the Chief Joseph (CJ), Monument (M) and Steens Mountain (SM) dike swarms (shaded ovals), and the Owyhee Plateau (OP). The quarternary basaltic centers of Diamond Craters (DC) and Jordan Craters (JC) are shown with grey and white stars respectively. Red triangles show Holocene volcanism. The black dashed 87Sr/86Sr = 0.706 line is taken from the compilation in Ernst (1988). The boundaries of Basin and Range extension (Wernicke et al., 1988) are shown with a purple dashed line. The black age contours of the HLP track are taken from Jordan (2002). The black age contours of the YSRP track are from Christiansen et al. (2002).

to the WNW. By ~ 22 Ma, volcanism associated with Basin and Range extension had reached southern Oregon (Scarberry et al., 2009), forming the WNW trending Brothers Fault Zone. The most notable volcanism in the area since the mid-Miocene has been associated with the Columbia River Basalts (CRB), the High Lava Plains (HLP) track, and the Yellowstone–Snake River Plains (YSRP) track. Beginning ~ 16.6 Ma, flood basalts began erupting from north– south trending dikes near Steens Mountain in southeastern Oregon (Fig. 1). Later flood basalts erupted to the north and northwest from the Chief Joseph and Monument Dike Swarms respectively (Camp and Ross, 2004). Over the course of 1.5 Ma, over ~ 220,000 km3 of basalt were erupted, covering much of northern and western Oregon, and southern Washington state (Camp and Hanan, 2008). At roughly the same time as the Steens flood basalts erupted, the first silicic volcanic field formed in northern Nevada at the southwestern edge of the Owyhee plateau (Brueseke et al., 2008). By 12 Ma, silicic volcanism had migrated away from the area surrounding the Owhyee Plateau in two distinct directions. The first is the YSRP track whose silicic volcanism progressed to the northeast, roughly parallel to North American plate motion (Smith et al., 2009). The second is the HLP volcanic track whose rhyolitic volcanic centers progressed to the northwest at an oblique angle to the YSRP, with Holocene activity located at Newberry Caldera (Jordan et al., 2004). Unlike the silicic volcanism, the basaltic volcanism on either track does not display an obvious time-progressive pattern with both tracks still displaying active basaltic volcanism along much of their lengths. The relationship between the three post-mid-Miocene volcanic provinces (the CRB, the HLP, and the YSRP) and the tectonic drivers responsible has been the source of considerable controversy. Because the YSRP volcanism trends roughly parallel to North American plate motion, it has often been described as an intra-continental hotspot/ plume track (e.g. Geist and Richards, 1993; Hanan et al., 2008; Morgan, 1972; Pierce and Morgan, 1992, 2009; Pierce et al., 2000; Smith et al.,

2009). However, the silicic volcanism of the HLP began at the same place and time as the YSRP but progresses in a direction very different from North American plate motion (e.g. Draper, 1991; Jordan et al., 2004; Meigs et al., 2009). The time progression of the HLP track raises questions about the suitability and/or completeness of a plume explanation. Non-plume related theories to explain the region's tectonomagmatism include Basin and Range related extension and rotation (Cross and Pilger, 1982), back-arc extension (Christiansen and McKee, 1978), rollback of the subducting Juan de Fuca plate (Carlson and Hart, 1987), and lithospheric delamination (Hales et al., 2005). In order to be able to deduce the role of past tectonic regimes, it is necessary to first understand the geometry of existing structures, especially in the shallow mantle. Recent seismic studies using surface waves (Pollitz, 2008; Pollitz and Snoke, 2010; Warren et al., 2008; Yang and Ritzwoller, 2008), ambient noise (Moschetti et al., 2007; Yang et al., 2008), and teleseismic body waves (Burdick et al., 2009; Obrebski et al., 2010; Roth et al., 2008; Sigloch et al., 2008; Tian et al., 2009; Xue and Allen, 2010) have taken advantage of the large dataset provided by the EarthScope Transportable Array (TA) in conjunction with other datasets to invert for broad scale maps of shear wave velocities in the western United States. Of these, only Warren et al. (2008) focus specifically on the area including the HLP. In order to better constrain the complex tectonic structures of the HLP/YSRP region, we use vertical component, Rayleigh surface wave data from the Transportable Array and the 100+ stations of the HLP seismic experiment (2006–2009) (Eagar et al., submitted for publication) to create a detailed three-dimensional image of the uppermost mantle. Because fundamental mode Rayleigh waves provide good control only on the top ~200 km of the earth's structure, this study cannot discriminate the depth extent of low velocity zones originating deeper than 200 km. However, the detailed structures observed in the shallowest portion of the mantle do provide constraints on the formation of these volcanic fields.

L. Wagner et al. / Earth and Planetary Science Letters 299 (2010) 273–284

2. Data For this study, we use data collected from two separate but coeval deployments: the 116 station sites of the HLP seismic experiment, and a regional subset of the EarthScope/USArray Transportable Array (TA) (Fig. 2). The High Lava Plains array was an IRIS-PASSCAL supported broadband deployment consisting of over 100 stations installed across southeastern and central Oregon, western Idaho, and northern Nevada (Eagar et al., submitted for publication). The first stations were deployed in early 2006, with the bulk of the stations deployed during summer of 2007. Over the course of this project, some of these stations were relocated, resulting in a total of 118 unique station sites, of which 107 are used in this study. All stations for the HLP array were demobilized by September, 2009. At the same time as the HLP array, the stations of the TA were moving east across our study area. For this study, we use data from rows E–Q and columns 1–17 for a total of 195 TA stations, and 302 stations total. We examined all events with magnitudes Mw N 6.3 occurring between September, 2006 and September, 2009 at distances between 22° and 130°. Of these, 99 events provided data for at least one of the frequencies analyzed in this study. Figure 2 shows the locations of these events with respect to our study area. While not entirely evenly

Fig. 2. Stations and events used in this study: A) locations of the 99 events used in this study relative to the HLP using an azimuthal equidistant projection. Orange lines indicate great circle paths from events to the center of our study area. B) Yellow diamonds are the stations of the Transportable Array (TA) used in this study. Red diamonds are the High Lava Plains seismic experiment stations used. Small black dots are the gridnodes defined for the inversion. Large black diamonds are the corners of the study area (see text for details).

275

distributed, the event distribution used here provides very good azimuthal coverage that helps reduce horizontal streaking of phase velocity anomalies. For this study, we considered 14 periods from 25 to 143 s. Not all events were well recorded at all periods (see Table 1 in the Supplemental Materials for the number of events and traces used at each frequency). To process these data, we follow the analysis procedures first outlined in Forsyth and Li (2005). The instrument responses are first normalized to that of the most frequently used instrument type. For each event, traces are bandpassed using a 7– 10 mHz wide filter centered on the frequency of interest. Bandpassed waveforms are visually inspected for consistency between stations and a minimum signal-to-noise ratio of 10:1, then subsequently windowed around the fundamental mode Rayleigh wave using a variable length window with a 50 s taper on either end to minimize the contributions from noise and other modes. Fourier analysis is then used to determine the phase and amplitude of the Rayleigh wave at the desired period at each station. 3. Methods In order to account for scattering outside of our study area that might result in the Rayleigh wave plane wave arriving off the projected great circle path or in waveform complexities associated with multi-path arrival of energy, we employ the two-plane wave method of Forsyth and Li (2005). This method for determining 2-D phase velocity deviations for Rayleigh waves at a given period proceeds in two steps. First, we model the observed waveforms as the sum of two distinct Rayleigh wave plane waves, each with its own amplitude, phase, and backazimuth. These two plane waves will constructively and destructively interfere across the array, resulting in systematic variations in amplitude. The first step of the Forsyth and Li (2005) method inverts these amplitude variations using a simulated annealing method to determine the best fitting parameters for the two plane waves. The second step of this method then calculates differences in predicted and observed phase and amplitude at each station in order to determine 2-D phase velocities across our study area. In order to account for scattering within our study area and to predict the phase delay and amplitude effects of local structure, we use the finite frequency kernels of Yang and Forsyth (2006), which in turn uses the single scattering Born approximation of Zhou et al. (2004). While previous versions of the two-plane wave method considered only deviations in phase to determine 2-D velocity in this step, these finite frequency kernels consider both phase and amplitude with a resultant significant increase in resolution (Yang and Forsyth, 2006). Observed amplitudes are corrected for geometrical spreading and attenuation assuming Q = 100. We do not explicitly solve for anelasticity in the inversion because the attenuation effect is small and has little effect on the estimated phase velocities (Yang and Forsyth, 2008). The dominant effects on amplitude are multipathing interference and focusing and defocusing due to lateral heterogeneities; the former we represent with the parameters representing the incoming wavefield and the latter we represent with the finite frequency amplitude response functions. Regularization is provided in the form of a priori model covariance, which we set at 0.2 km/s. There always is a trade-off between resolution and variance of the model, or between amplitude of the model corrections and misfit to the data. The choice of the optimum position on these trade-off curves is inevitably somewhat arbitrary. In these inversions for phase velocity, the resolution or regularization is governed by a combination of the smoothing length, Lw, and the a priori model covariance. We choose a smoothing length of 40 km, which seems to provide the maximum resolution achievable within the denser parts of the array. At lengths shorter than this value, both the variance of the model parameters and the amplitude of the short length-scale velocity variations increase rapidly. With this length-

276

L. Wagner et al. / Earth and Planetary Science Letters 299 (2010) 273–284

Fig. 3. Results of one dimensional Rayleigh wave phase velocity and shear wave velocity inversions: A) phase velocities for IASP91 are shown in orange (Kennett, 1991). The best fitting average phase velocity at each period studied is shown with black diamonds with horizontal error bars. The red and blue lines show the calculated phase velocities for the red and blue dots in part B of this figure. B) Green lines show the sensitivity kernels to shear wave velocity structure for the periods used in this study. The shortest period used (25 s) has the shallowest peak sensitivity (~40 km). The longest period used (143 s) has a peak sensitivity at ~ 200 km. Orange dots show the shear wave model for IASPEI 91. Red dots show the best fitting 1-D shear wave velocity model to recover the best fitting phase velocities shown with black diamonds in part A. The red line in Figure 3A shows the calculated phase velocities for this best-fit 1-D shear wave velocity model. The blue dots in part B (whose calculated phase velocities are shown with a blue line in part A) show the simplified 1-D shear wave velocity starting model used in the 3-D inversion to depth. The starting model is adjusted at each point to have an appropriate Moho depth for that location.

scale, 0.2 km/s a priori variance for the model parameters corresponds roughly to the amplitude of the phase velocity variations and provides damping for the outer, less well constrained parts of the model space. Grid node locations are established using a pole of rotation 90° from the center of the study area. Grid nodes are then evenly distributed as a function of azimuth and distance from this pole. Grid columns and rows were spaced in 0.25° increments resulting in an average grid node distance of ~27 km (Fig. 2). Outermost grid nodes are left underdamped in order to “absorb” travel-time variations from outside the array that are not adequately modeled with our twoplane-wave approximation. Our model parameters at each grid node are not actual phase velocities, but rather velocity coefficients. The velocity at any point can then be determined by a Gaussian weighted average of the velocity coefficients from surrounding grid nodes, where the weighting falls-off to 1/e times the coefficient at a distance of 40 km from the grid node of interest. We include terms to account for average azimuthal anisotropy across our study area (Forsyth and Li, 2005). Phase velocities at any location are defined as Cðω; θÞ = B0 ðωÞ + B1 ðωÞ cosð2θÞ + B2 ðωÞ sinð2θÞ + B3 ðωÞ cosð4θÞ + B4 ðωÞ sinð4θÞ where ω is the frequency and θ is the azimuth. The 4θ terms are not included in this inversion because they have been shown to be small for Rayleigh waves (i.e. Smith and Dahlen, 1973; Weeraratne et al., 2007). The fast direction of propagation is then defined as 0.5 ⁎ arctan(B2/B1), and the peak-to-peak degree of anisotropy is given by 2 ⁎ (B21 + B22)1/2. Recent shear wave splitting studies show a fairly simple pattern of

anisotropy across much of our study area (i.e. Liu, 2009; Long et al., 2009; Xue and Allen, 2006), so finding the average surface wave azimuthal anisotropy across our study area is sufficient. In total, the number of model parameters in the inversion is equal to six times the number of events (phase, amplitude, and azimuth for two planes for each event) plus the number of grid nodes plus two anisotropy terms. Because the number of events used varies between periods (see Table 1 in the Supplementary Materials), the number of model parameters will vary as well. However, the same grid of 3640 grid nodes is used for inversions at all periods. We first invert for the best fitting average phase velocity for the entire study area at each period. The starting average phase velocities for these inversions are calculated using the P- and S-wave velocities of IASP91 (Kennett, 1991) and the method of Saito (1988). Both the starting IASP91 phase velocities and the resultant best-fit average phase velocities of these inversions are shown in Figure 3A. The bestfit average phase velocities were then used as starting velocities for the inversions for 2-D phase velocity deviations for each period. The resultant 2-D phase velocity maps can be found in Figure 4. We then determine 3-D shear wave velocity structure by inverting for a 1-D shear wave velocity model at each point on a grid of points defined on the 2-D phase velocity maps. We then combine the 1-D shear wave models determined at each point in map view to create a 3-D shear wave velocity model. The 1-D inversions exploit the varying depth sensitivities of Rayleigh wave velocities to shear wave velocity as a function of frequency. At each node location, we determine the phase velocity at each of the 14 periods considered in this study. We then invert these phase velocities for the 1-D shear wave velocity

Fig. 4. Two dimensional phase velocity results and confidence limits: the top two rows show the phase velocity deviations and twice the standard deviations, respectively, for 28 s and 40 s Rayleigh waves. The bottom two rows show the same for 58 s and 91 s Rayleigh waves. Velocity deviation figures (Rows 1 and 3) show deviations from the starting phase velocity in color, and absolute phase velocity values in contours at increments of 0.05 km/s. Contours of 2σ (Rows 2 and 4) are given in km/s. Cooler colors correspond to areas of better lateral resolution. For interpretation of confidence limits, see text.

L. Wagner et al. / Earth and Planetary Science Letters 299 (2010) 273–284

277

278

L. Wagner et al. / Earth and Planetary Science Letters 299 (2010) 273–284

model that would produce the best fitting phase velocities at all periods. Because we use only Rayleigh waves, this shear velocity is roughly equivalent to Sv for a horizontally propagating shear wave, but the kernels we employ assume an isotropic medium. Fig. 3 shows the sensitivity kernels for each of our 14 periods. Our shortest periods have peak sensitivities at depths of ~ 40 km, and our longest periods have broad sensitivity kernels with maxima at depths of ~200 km. As a result, our inversions have little sensitivity at depths shallower than 40 km, and reduced depth resolution at depths greater than 200 km. In order to optimize our starting shear wave velocity model, we first invert our best fitting average phase velocities (black diamonds, Fig. 3) for 1-D shear wave velocity structure (Saito, 1988; Weeraratne et al., 2003). Shear wave velocity models are defined by 128 layers of fixed thickness and uniform P and S-wave velocities extending from the surface to a depth of 660 km. Each layer is assigned an a priori standard deviation of 0.2 km/s to provide regularization. The resultant best-fit average shear wave velocity model (red dots/line, Fig. 3) shows significantly slower Vs at upper mantle depths than IASP91 (Kennett, 1991). Velocity deviations deeper than 400 km are negligible. We then use a simplified version of this best-fit shear wave velocity model (blue dots/line, Fig. 3) as our starting model for the shear wave velocity inversions at each point in map view within our study area so that as little vertical variation as possible will be introduced into the local solutions by the starting model. P-wave velocities for these inversions are defined according to the values of IASP91 (Kennett, 1991); Rayleigh waves in this period range are insensitive to P velocity deeper than about 50 km. Because surface waves are not able to resolve crustal thickness very well, we create a 2-D crustal thickness map using results from the receiver function studies from Kevin Eagar (Eagar et al., submitted for publication) for the greater HLP region, supplemented more broadly by the EARS catalogue (Crotwell and Owens, 2005) (see Supplemental Materials Fig. 1). We then adjust the starting model Moho depth to reflect the crustal thickness at that location. The results of these individual 1-D inversions at each point in map view are then combined to form our 3D shear wave velocity model (Figs. 6–10).

north, positive velocity anomalies are observed at all periods. By 40 s, the full length of the descending Juan de Fuca/Gorda Plate is clearly visible as a positive velocity anomaly at 122°W longitude. The negative velocity anomaly just east of this feature we interpret as being associated with the Cascade arc. The negative velocity deviations that extend from the Oregon/Nevada/Idaho border to Yellowstone along the YSRP track have become more pronounced, especially to the northeast. Areas south and east of the YSRP are characterized by generally positive velocity anomalies, possibly due to the presence of the Wyoming Craton. At 58 s, the negative velocity anomalies previously confined to the areas directly below the Cascades and the HLP/YSRP tracks have now broadened to include much of the northern Basin and Range area in northern Nevada. This contrasts with the positive anomalies further to the east beneath the Wyoming Craton. The velocity anomalies associated with the HLP track have decreased in amplitude, but the anomalies along the YSRP track have increased in amplitude, especially to the northeast. At 91 s, the most prominent features are the YSRP track, the north–south trending negative velocity anomalies just east of the Cascades and northwestern Nevada, and the positive velocity anomalies to the east beneath the Wyoming Craton. The average anisotropy values for azimuth and magnitude at each period are shown in Figure 5. In general, the recovered fast directions are rotated a few degrees counterclockwise from due east–west, consistent with observations from shear wave splitting studies in the area (i.e. Liu, 2009; Long et al., 2009; Xue and Allen, 2006). Magnitudes increase significantly at periods between 91 and 125 s, consistent with a model where the strongest anisotropy originates from depths below 100 km.

4.2. Shear wave velocities Our 3-D shear wave velocity model is shown in Figures 6–10. Figure 6 shows map view slices from 65 to 165 km depth in 20 km depth increments. Using the period range employed in this study, surface wave inversions for shear wave velocities are ill constrained at

4. Results 4.1. Phase velocities The results of the inversions for 2-D phase velocity structure are shown in Figure. 4. Colors show deviations from the starting model, and contours show absolute velocities in increments of 0.05 km/s. Figure. 4 shows the phase velocity maps and confidence limits for 28, 40, 58, and 91 s. Phase velocity maps and approximate 95% confidence limits for all other periods in this study can be found in Supplemental Materials Figures 2–4. In the confidence plots, cooler colors indicate areas that are better resolved, whereas warmer colors indicate areas that are less well resolved. Because these confidence limits include the a priori assignment of standard deviations to the starting models that serve to damp the phase velocity inversions, the proper interpretation of these values is that deviations in phase velocity from the starting average phase velocity that exceed these confidence limits are statistically significant; the values do not represent confidence limits on the absolute velocities. This inversion does not formally account for correlated noise between stations, so the standard deviations are likely somewhat lower than they should be, especially with the inclusion of the densely spaced HLP stations. Longer periods will be more affected by this noise covariance than shorter periods. We are currently developing a formal method for including station covariances, but this is beyond the scope of the current paper. At 28 s, we can see part of the high velocity anomaly that we interpret as the subducting Juan de Fuca plate, along with generally low velocities beneath the Cascade Arc and Yellowstone. Negative velocity deviations are also seen along the HLP and YSRP tracks. To the

Fig. 5. Anisotropy results from phase velocity inversions: For each period studied (xaxis), we show the fast direction of anisotropy (shown in map view with E–W indicated by lines parallel to the x-axis), along with the percent magnitude of the anisotropy (yaxis). Errors in fast direction are shown by the range of azimuths in black. These are very small, and fast directions are generally east west for all periods. Errors in magnitude are shown by vertical error bars. The magnitude of the anisotropy increases significantly between 91 and 125 s, indicating a significant contribution from depths of greater than ~ 100 km.

L. Wagner et al. / Earth and Planetary Science Letters 299 (2010) 273–284

crustal depths and velocities in the vicinity of the Moho are strongly affected by the assumed crustal thickness, so we present our results only for depths greater than 55 km. As with the phase velocities, a number of features are very clear at all depths: the low velocity zones beneath the Cascades, the HLP, Yellowstone, and the YSRP, and the high velocities of the Wyoming Craton. To the north of the HLP–YSRP tracks, we observe fast shear wave velocities, especially north and east of the 87Sr/86Sr = 0.706 line delimiting the boundary between cratonic North America and the accreted terranes. However, high velocities are also observed beneath the northernmost BMP. The high velocities of this area and the Wyoming Craton (~4.4 km/s) are not unusually fast by normal lithospheric mantle standards, but show up as very fast (dark blue) because of the low velocity starting model. In order to test if these velocities are correct or are kept too low due to the slow starting model, we re-ran both the 2-D phase velocity inversions and the inversions for 3-D Vs structure using IASP91 as the starting model for both steps. The results were universally low velocities across the area in comparison to the starting model, with the absolute velocities being very close to our previous results. The robustness of the absolute velocity results gives us confidence in our preferred model presented here. Figure 7 shows the cross-sections A–A′ and B–B′, both in lateral view and in perspective to show their relation to topographic features. The locations of these cross-sections are also indicated in map view in Figure 6. The first cross-section, A–A′, runs from the coast in northern Oregon across the Cascades, along the HLP, through the Owyhee plateau and the Basin and Range, ending in the southern part of the Wyoming Craton. In this cross-section, the descending Juan de Fuca plate is clearly visible, as is the low velocity zone associated with the mantle wedge beneath the Cascade arc. Continuing along this track to the southeast, the low velocity zone seen below the arc diminishes in amplitude and increases in depth slightly, before shallowing again and recovering some of its amplitude beneath Steens Mountain. To the southeast of Steens is the low velocity region associated with the Basin and Range extension. The amplitude of this negative velocity anomaly decreases somewhat before reaching the edge of the Basin and Range, where the amplitudes again increase. The end of this transect beyond the Basin and Range is characterized by the relatively high velocities associated with the southern portion of the Wyoming Craton. The cross-section B–B′ in Figure 7 cuts from the California coast just south of Mendocino up through the northern Basin and Range and Owyhee plateau, and finally up along the YSRP volcanic track, ending just past Yellowstone. Modest low velocities are observed southwest of the Owyhee plateau at depths up to 150 km, similar to those observed in A–A′ beneath the Basin and Range. To the northeast beneath the Owyhee plateau, the negative velocity anomaly associated with the YSRP increases significantly in amplitude. The highest amplitudes for this anomaly exist between 75 and 175 km depth beneath southwestern Idaho, but then shallow to the northeast. Its shallowest point lies just shy of Yellowstone at depths of between 55 km and 120 km. The anomaly directly beneath Yellowstone is the lowest velocity in our entire study area. Beyond Yellowstone, the low velocity anomaly deepens to the north (not visible in Fig. 7; see map view plots in Fig. 6), with very low velocities extending to depths of at least 150 km. At greater depths, the dip and amplitude of this anomaly are difficult to constrain because resolution decays with increasing depth. The form of the anomaly to the east of Yellowstone seen in cross-section B–B′ is less well defined because our station coverage ends (Fig. 2). There is information about the structure outside the array from crossing ray paths and the finite width of the sensitivity kernels, but resolution and confidence decay with increasing depth. 5. Discussion The models from our study clearly show features identified by a number of previous studies such as the downgoing slab, the Wyoming

279

Craton, high velocities beneath northern Oregon and Washington, low velocities beneath the Basin and Range, Cascades, High Lava Plains, and the YSRP. However, our ability to provide a higher resolution look at this tectonically complex area allows us to say more about the deep structure of the HLP and YSRP tracks.

5.1. The High Lava Plains Figure 8 shows the −3% surface contour from our shear wave velocity deviation model in 3-D beneath the geologic map from Figure 1. In this view, the discrete low velocity zone beneath the YSRP is the dominant feature, along with a clear north–south trending anomaly just east of the Cascade arc. However, there is no clear linear feature connecting the low velocities in southeastern Oregon to the Newberry Volcano. The lack of a strong continuous low velocity track beneath the HLP akin to the pronounced anomaly beneath the YSRP may suggest that the volcanic progression from east to west along the HLP may not be as continuous as previously assumed, or that the westward migration of volcanism marks a broader thermal anomaly attached to the eastern edge of the Cascades that has migrated to the west with the arc due to extension in eastern Oregon. Previous work based on 40Ar/39Ar dating studies (i.e. Jordan, 2002; Jordan et al., 2004) has suggested that the volcanism of the HLP is a mirror image of the YSRP volcanism. This argument is based largely on the similar timing and location of the start of both tracks, and the apparent age progression of the silicic volcanism to the northwest, with the current end located beneath the Newberry shield volcano just east of the main Cascade range. The age contours for the HLP shown in Figures 1 and 8 are from Jordan (2002). Here, the contours are relatively parallel to one another, progressing to the west until 6 Ma. At that point, there is a distinct “kink” in the contours. However, Jordan et al. (2004) show only smoothly varying contours throughout. We re-evaluate these contours using the dates of rhyolitic volcanism compiled in Meigs et al. (2009) (Fig. 9). Between 16 and 15 Ma, rhyolitic volcanism is widespread across much of the HLP, reaching as far west as 120°W, with samples found from the southern Oregon border up to the northern edge of the HLP. After a period of relative quiescence between 15.0 Ma and 12.0 Ma, rhyolitic volcanism resumes again slowly, this time concentrated along the easternmost part of the HLP track. After that, the rhyolites migrate to the west at a fairly constant rate, with parallel contours trending roughly NE–SW consistent with the contours of both Jordan (2002) and Jordan et al. (2004). The younger rhyolites erupted within the past 5 Ma also show a fairly steady progression northwest towards Newberry Caldera. However, between 7 and 5 Ma, the volcanic pattern deviates from the simple NE–SW trending contours. Instead, volcanism splits into two tracks, one at the northern edge of the HLP trending due east–west between Rattlesnake tuff and Newberry caldera, and the other beginning at ~7 Ma due south of Newberry and trending north from there. In general, most rhyolites older than 7 Ma erupted east of the north–south trending low velocity zone located just east of the modern Cascade arc at ~121°W. We interpret this negative velocity anomaly as being caused by the heat and partial melting associated with the arc. Similarly, all rhyolites younger than 5 Ma occur directly above this negative velocity anomaly. Interestingly, the N7 Ma contour also correlates well with changes in shear wave splitting times as determined by Long et al. (2009) (Fig. 9). They find very large split times (N2 s) beneath the southeastern part of the High Lava Plains, but northwest of our N7 Ma contour, the magnitudes of the splitting times drop off sharply, and do not resume their large magnitudes until well north of Newberry Caldera, beneath the Three Sisters volcanoes. The full implications of the shear wave splitting magnitude patterns are still not well understood, but their alignment with both volcanic patterns and mantle velocities is noteworthy. There also appears to be a similar gap in the low resistivity anomaly

280

L. Wagner et al. / Earth and Planetary Science Letters 299 (2010) 273–284

Fig. 6. Shear wave velocity deviation maps: depths of each map are indicated directly above the map. Colors indicate deviations from the starting velocity model, and contours show absolute velocities in increments of 0.05 km/s. Locations of cross-sections shown in Figure 7 are indicated as well.

between Newberry volcano and areas further to the east (Patro and Egbert (2008) Fig. 2, latitude 45°N cross-section). We suggest that the HLP rhyolitic volcanism does form a track between 12 and 7 Ma in southeastern Oregon that was perhaps

localized to the HLP because the magmas found it easier to erupt through the highly fractured crust of the Brothers Fault zone. After that, the proximity of this track to the arc may have resulted in some complex interaction between arc volcanism and volcanism previously

L. Wagner et al. / Earth and Planetary Science Letters 299 (2010) 273–284

281

Fig. 7. Shear wave velocity deviation cross-sections: cross-sections A–A′ and B–B′ shown in perspective at the top, and from the side at the bottom. Side-view cross-sections show topography (brown line) and Moho depth (blue line) directly above the colored velocity deviations. Contours are for velocity deviations in increments of 1%. Features noted are Newberry Caldera, Steens Mountain, the Juan de Fuca Slab (JdF Slab), the Owyhee Plateau (OP), the Basin and Range province (B&R), the Yellowstone–Snake River Plains track (YSRP) and the Wyoming Craton.

associated with the HLP. The area of the Cascade arc delimited by the N–S trending portion of the 5–7 Ma rhyolitic volcanism corresponds roughly with the “Central” region defined by Schmidt et al. (2008). This region of the Cascades is characterized by its prevalence of low K tholeiites (LKT), which are also common within the rest of the HLP track, suggesting that the thermal anomaly involved in the genesis of N7 Ma silicic volcanism likely still played a role in the genesis of younger volcanism. Unusually hot (seismically slow) mantle, however, is no longer present beneath the central HLP, but is present in both areas of recent basaltic volcanism in eastern Oregon — east of the Cascade front near Newberry and at Diamond and Jordan craters. The segregation of volcanism in eastern Oregon to the margins of the extensionally-thinned crust (Eagar et al., submitted for publication) of the HLP is analogous to the situation in the central Basin and Range where post-5 Ma volcanism is most prevalent along the eastern

margin of the Sierra and west of the Wasatch front marking the west and east boundaries of the extending Great Basin. Thus, both the volcanic history and location of the low velocities in eastern Oregon appear to be more strongly influenced by the history of extension in the HLP than by the migration of a mantle hotspot to the northwest across Oregon over the last 12 Ma. 5.2. The Yellowstone/Snake River Plains track The linear low velocity feature associated with the YSRP track is one of the most prominent features in our models (Figs. 4, 6, 7, 8, and 10). To the southwest, beneath the oldest portion of the track, the anomaly is more diffuse than it is further along towards the northeast (Fig. 6). Also, the top of the low velocity zone shallows and the strength of the anomaly increases to the northeast (Figs. 6 and 7). The

282

L. Wagner et al. / Earth and Planetary Science Letters 299 (2010) 273–284

Fig. 8. Geological map from Figure 1 with − 3% low velocity surface contour shown in 3-D below the map. This 3-D surface contour can also be seen in Figures 9 and 10. Also indicated are the locations of the cross-sections shown in Figure 7. The yellow and blue contours in the HLP track show our b 5 Ma and N 7 Ma contours for rhyolitic volcanism in the area (see Fig. 9 and the text for details).

broader anomaly at the beginning of the track may be in part due to thermal diffusion and in part due to spreading of the hot, low-density mantle over time. The deepening of the top of the anomaly may be in part due to cooling from the earth's surface, which is likely to occur once a moving heat source has passed. The variation in strength of the anomaly is consistent with the heat source currently coming from

below Yellowstone. These observations are consistent with a heat source moving relative to the North American plate that started near the Owyhee Plateau and moved towards the northeast from there. The variation in strength of the anomaly is consistent with the heat source currently coming from below Yellowstone. The northward trending deepening of the anomaly directly beneath Yellowstone (Fig. 10) is

Fig. 9. High Lava Plains rhyolites, shear wave splitting, and low velocity anomaly: This shows the same − 3% surface contour plotted in Figure 8, overlain by the shear wave splitting results of Long et al. (2009) and the rhyolitic volcanism reported in Meigs et al. (2009). The rhyolite locations (triangles) are color-coded according to age, with blues and purples being older than 7 Ma, and warmer colors younger than 5 Ma. All volcanism younger than 5 Ma (with one exception) falls to the north and west of the yellow contour. All volcanism older than 7 Ma falls east of the blue contour. In between are the 5–6 Ma rhyolites in white, and the 6–7 Ma rhyolites in black. Black outlines define the E–W and N–S trends of the rhyolitic volcanism during this time. Shear wave splitting directions are shown by the black lines. Circle size and greyscale colors indicate the magnitude of the shear wave split in seconds with larger circles and darker colors indicating larger split times. The quarternary basaltic centers of Diamond Craters (DC) and Jordan Craters (JC) are shown with grey and white stars respectively.

L. Wagner et al. / Earth and Planetary Science Letters 299 (2010) 273–284

283

Fig. 10. Yellowstone/Snake River Plains low velocity zone: shown is the − 3% surface contour from Figure 8 in perspective, looking from the north down towards the southeast. Cross-sections from Figure 7 are shown, as is the location of Yellowstone Caldera (blue triangle).

similar to that observed by Schutt and Humphreys (2004), Yuan and Dueker (2005), and Waite et al. (2006). We note, however, that our anomaly does not trend to the northwest, but rather migrates due north from the center of the anomaly with increasing depth. This may reflect the differing raypaths and sensitivities of surface waves versus teleseismic body waves, with the latter being more vulnerable to vertical streaking along a dominant raypath. Other shear wave velocity models developed using surface waves (e.g. Pollitz and Snoke, 2010; Schutt et al., 2008; Stachnik et al., 2008) do not show particularly pronounced trends to the northwest at these depths. The one feature of the YSRP anomaly that is not consistent with a simple model of a relatively stationary heat source coming from below and with the overriding plate moving to the southwest is that the bottom margin of the YSRP LVZ also appears to shallow to the northeast. This shallowing is subtle, and may simply be an artifact of worsening resolution and the tendency of surface wave inversions to oscillate somewhat with depth. If real, this particular observation does not obviously favor any one particular model of YSRP genesis, and may simply reflect hot asthenosphere along the southern portion of the track associated with Basin and Range extension.

from below, though the depth of this heat source cannot be constrained below ~ 200 km in this study. Supplementary materials related to this article can be found online at doi:10.1016/j.epsl.2010.09.005.

Acknowledgements This project benefited greatly from the combination of two very large seismic datasets: the Transportable Array and the High Lava Plains seismic experiment. We are very grateful to the many people involved in making both of these a reality. Many thanks also to the PASSCAL Instrument Center and the IRIS Data Management Center for their ongoing support. This paper was greatly improved by generous input from Rick Carlson and Julie Donnelly-Nolan. The High Lava Plains project was funded through NSF award EAR-0507248 (MJF) and EAR-0506914 (DEJ). LW's participation was supported by NSF award EAR-0809192 and DWF was supported by NSF award EAR0745972. References

6. Conclusion Our high-resolution surface wave tomographic images show a number of tectonic features in the northwestern United States including the downgoing Juan de Fuca slab, the Wyoming Craton, the Basin and Range, and the Yellowstone–Snake River Plains track. The one feature that is not clearly observed is a continuous low velocity track associated with the High Lava Plains volcanic trend. This suggests that there may be a significant difference between the tectonic processes behind the formation of the YSRP and HLP tracks. We suggest that our results are consistent with the formation of the HLP volcanism due to the interaction of Basin and Range extension together with subduction zone/arc processes. Beneath the YSRP, we observe a discrete low velocity anomaly that shallows towards the northeast, but then deepens directly beneath Yellowstone. This anomaly is broadly consistent with a moving heat source coming up

Brueseke, M.E., Hart, W.K., Heizler, M.T., 2008. Diverse mid-Miocene silicic volcanism associated with the Yellowstone-Newberry thermal anomaly. Bull. Volcanol. 70, 343–360. Burdick, S., van der Hilst, R.D., Vernon, F.L., Martynov, V., Cox, T., Eakins, J., Mulder, T., Astiz, L., Pavlis, G.L., 2009. Model update December 2008: upper mantle heterogeneity beneath North America from P-wave travel time tomography with global and USArray Transportable Array data. Seismol. Res. Lett. 80, 638–645. Camp, V.E., Hanan, B.B., 2008. A plume-triggered delamination origin for the Columbia River Basalt Group. Geosphere 4, 480–495. Camp, V.E., Ross, M.E., 2004. Mantle dynamics and genesis of mafic magmatism in the intermontane Pacific Northwest. J. Geophys. Res. 109, B08204. doi:10.1029/ 2003JB002838 Carlson, R.L., Hart, W.K., 1987. Crustal genesis on the Oregon plateau. J. Geophys. Res. 92, 6191–6207. Christiansen, R.L., McKee, E.H., 1978, Late Cenozoic volcanic and tectonic evolution of the Great Basin and Columbia inter-montane regions, Geol. Soc. Am. Memoir 152, 283–311. Christiansen, R.L., Foulger, G.R., ad Evans, J.R., 2002. Upper-mantle origin of the Yellowstone hotspot. Geological Society of America Bulletin 114, 1245–1256. Cross, T.A., Pilger, R.H., 1982. Controls of subduction geometry, location of magmatic arcs, and tectonics of arc and back-arc regions. Geol. Soc. Am. Bull. 93, 545–562.

.

284

L. Wagner et al. / Earth and Planetary Science Letters 299 (2010) 273–284

Crotwell, H.P., Owens, T.J., 2005. Automated Receiver Function Processing, Seismo. Res. Lett: Electronic Seismologist November/December, 2005. Dorsey, R.J., LaMaskin, T.A., 2007. Stratigraphic record of Triassic–Jurassic collisional tectonics in the Blue Mountains Province, northeastern Oregon. Am. J. Sci. 307, 1167–1193. Draper, D., 1991. Late Cenozoic bimodal magmatism in the northern Basin and Range province of southeastern Oregon. J. Volcanol. Geotherm. Res. 47, 299–328. Eagar, K.C., Fouch, M.J., James, D.E., Carlson, R.L., Group, H.L.P.S.W., submitted for publication, Crustal structure beneath the High Lava Plains of eastern Oregon and surrounding regions from receiver function analysis, J. Geophys. Res. Ernst, W.G., 1988. Metamorphic terranes, isotopic provinces, and implications for crustal growth of the western United States. J. Geophys. Res. 93, 7634–7642. Fleck, R.J., Criss, R.E., 1985. Strontium and oxygen isotopic variations in Mesozoic and Tertiary plutons of central Idaho. Contrib. Mineral. Petrol. 90, 291–308. Forsyth, D.W., Li, A., 2005. Array analysis of two-dimensional variations in surface-wave phase velocity and azimuthal anisotropy in the presence of multipathing interference. In: Levander, A., Nolet, G. (Eds.), Seismic Earth: Array Analysis of Broadband Seismograms: AGU Geophysical Monograph Series 157, Washington, DC, pp. 81–98. Foster, D.A., Mueller, P.A., Mogk, D.W., Wooden, J.L., Vogl, J.J., 2006. Proterozoic evolution of the western margin of the Wyoming craton: implications for the tectonic and magmatic evolution of the northern Rocky Mountains. Can. J. Earth Sci. 43, 1601–1619. Geist, D., Richards, M., 1993. Origin of the Columbia Plateau and Snake River plain: Deflection of the Yellowstone plume. Geology 21, 789–792. Hales, T.C., Abt, D., Humphreys, E., Roering, J.J., 2005. A lithospheric instability origin for Columbia River flood basalts and Wallowa Mountains uplift in northeast Oregon. Nature 438, 842–845. Hanan, B.B., Shervais, J.W., Vetter, S.K., 2008. Yellowstone plume-continental lithosphere interaction beneath the Snake River Plain. Geology 86, 51–54. Humphreys, E., 2009. Relation of flat subduction to magmatism and deformation in the western United States. In: Kay, S., Ramos, V.A., Dickinson, W. (Eds.), Backbone of the Americas: Shallow Subduction, Plateau Uplift, and Ridge and Terrane Collision 204: Geological Society of America Memoir, pp. 85–98. Jordan, B.T., 2002. Basaltic Volcanism and Tectonics of the High Lava Plains. Oregon State University, southeastern Oregon. Jordan, B.T., Grunder, A.L., Duncan, R.A., Deino, A.L., 2004. Geochronology of ageprogressive volcanism of the Oregon High Lava Plains: implications for the plume interpretation of Yellowstone. J. Geophys. Res. 109, B10202. doi:10.1029/ 2003JB002776. Kennett, B.L.N., 1991. IASPEI 1991 Seismological Tables. Bibliotech, Canberra, Australia. Liu, K.H., 2009. NA-SWS-1.1: a uniform database of teleseismic shear-wave splitting measurements for North America. Geochem. Geophys. Geodyn. G3 10, Q05011. doi:10.1029/2009GC002440. Long, M.D., Gao, H.G., Klaus, A., Wagner, L.S., Fouch, M.J., James, D.E., Humphreys, E., 2009. Shear wave splitting and the pattern of mantle flow beneath eastern Oregon. Earth Planet. Sci. Lett. 288, 359–369. Madsen, J.K., Thorkelson, D.J., Friedman, R.M., Marshall, D.D., 2006. Cenozoic to Recent plate configurations in the Pacific Basin: ridge subduction and slab window magmatism in western North America. Geosphere 2, 11–34. Meigs, A., Scarberry, K., Grunder, A.L., Carlson, R.L., Ford, M.T., Fouch, M.J., Grove, T., Hart, W.K., Iademarco, M., Jordan, B.T., Milliard, J., 2009. Geological and geophysical perspectives on the magmatic and tectonic development, High Lava Plains and northwest Basin and Range. In: O'Connor, J.E., Dorsey, R.J., Madin, I.P. (Eds.), Volcanoes to Vineyards: Geologic Field Trips through the Dynamic Landscape of the Pacific Northwest: Geological Society of America Field Guide, 15, pp. 435–470. Morgan, L.A., 1972, Plate motions and deep mantle convection, Geol. Soc. Am. Memoir 132, 7–22. Moschetti, M., Ritzwoller, M., Shapiro, N., 2007. Surface wave tomography of the western United States from ambient seismic noise: Rayleigh wave group velocity maps. Geochem. Geophys. Geodyn. G3 8, Q08010. doi:10.1029/2007GC001655. Obrebski, M., Allen, R.M., Xue, M., Hung, S., 2010. Slab-plume interaction beneath the Pacific Northwest. Geophys. Res. Lett 37, L14305. doi:10.1029/2010GL043489. Patro, P.K., Egbert, G.D., 2008. Regional conductivity structure of Cascadia: preliminary results from 3D inversion of USArray transportable array magnetotelluric data. Geophysical Research Letters 35. doi:10.1029/2008GL035326. Pierce, K.L., Morgan, L.A., 1992. The track of the Yellowstone hotspot: Volcanism, faulting, and uplift, in: P.K. Link, Kuntz, M.A., Platt, L.B., (Eds), Regional geology of eastern Idaho and western Wyoming, Geololgical Society of America Memoir 179. Pierce, K.L., Morgan, L.A., 2009. Is the track of the Yellowstone hotspot driven by a deep mantle plume? — Review of volcanism, faulting, and uplift in light of new data. J. Volcanol. Geotherm. Res. 188, 1–25. Pierce, K.L., Morgan, L.A., Saltus, R.W., 2000. Yellowstone Plume Head: Postulated Tectonic Relations to the Vancouver Slab, Continental Boundaries, and Climate, USGS Open-File Report 00-498. Pollitz, F.F., 2008. Observations and interpretations of fundamental mode Rayleigh wavefields recorded by the Transportable Array (USArray). J. Geophys. Res. 113, B10311. doi:10.1029/2007JB005556. Pollitz, F.F., Snoke, J.A., 2010. Rayleigh-wave phase velocity maps and three dimensional shear velocity structure of the western US from local non-plane surface wave tomography. Geophys. J. Int. 180, 1153–1169. Roth, J.B., Fouch, M.J., James, D.E., Carlson, R.L., 2008. Three-dimensional seismic velocity structure of the northwestern United States. Geophys. Res. Lett. 35, L15304. doi:10.1029/2008GL034669.

Saito, M., 1988. DISPER80: a subroutine package for the calculation of seismic normal mode solutions. In: Doornbos, D.J. (Ed.), Seismological Algorithms: Computational Methods and Computer Programs. Elsevier, New York, pp. 293–319. Scarberry, K., Meigs, A., Grunder, A.L., 2009. Faulting in a propagating continental rift: Insight from the late Miocene structural development of the Abert Rim fault, southern Oregon, USA. Tectonophysics 488 (1–4), 71–86. doi:10.1016/j.tecto.2009.09.025. Schmidt, M.E., Grunder, A.L., Rowe, M.C., 2008. Segmentation of the Cascade Arc as indicated by Sr and Nd isotopic variation among diverse primitive basalts. Earth Planet. Sci. Lett. 266, 166–181. Schutt, D.L., Humphreys, E., 2004. P and S wave velocity and Vp/Vs in the wake of the Yellowstone hot spot. J. Geophys. Res. 109, B01305. doi:10.1029/2003JB002442. Schutt, D.L., Dueker, K., Yuan, H., 2008. Crust and upper mantle velocity structure of Yellowstone hot spot and surroundings. J. Geophys. Res. 113, B03310. doi:10.1029/ 2007JB005109. Schwartz, J.J., Snoke, J.A., Frost, C.D., Barnes, C.G., Gromet, L.P., Johnson, K., 2010. Analysis of the Wallowa–Baker terrane boundary: implications for tectonic accretion in the Blue Mountains province, northeastern Oregon. Geol. Soc. Am. Bull. 122, 517–536. Sigloch, K., McQuarrie, N., Nolet, G., 2008. Two-stage subduction history under North America inferred from multiple-frequency tomography. Nat. Geosci. 1, 458–462. Smith, M.L., Dahlen, F.A., 1973. The azimuthal dependence of Love and Rayleigh wave propagation in a slightly anisotropic medium. J. Geophys. Res. 78, 3321–3333. Smith, R.B., Jordan, M., Steinberger, B., Puskas, C.M., Farrell, J., Waite, G.P., Husen, S., Chang, W., O'Connell, R., 2009. Geodynamics of the Yellowstone hotspot and mantle plume: Seismic and GPS imaging, kinematics, and mantle flow. J. Volcanol. Geotherm. Res. 188, 26–56. Stachnik, J.C., Dueker, K., Schutt, D.L., Yuan, H., 2008. Imaging Yellowstone plumelithosphere interactions from inversion of ballistic and diffusive Rayleigh wave dispersion and crustal thickness data. Geochem. Geophys. Geodyn. G3 9 (6), Q06004. doi:10.1029/2008GC001992. Tian, Y., Sigloch, K., Nolet, G., 2009. Multiple-frequency SH-wave tomography of the western US upper mantle. Geophys. J. Int. 178, 1384–1402. Waite, G.P., Smith, R.B., Allen, R.M., 2006. Vp and Vs structure of the Yellowstone hot spot from teleseismic tomography: evidence for an upper mantle plume. J. Geophys. Res. 111, B04303. doi:10.1029/2005JB003867. Warren, L.M., Snoke, J.A., James, D.E., 2008. S-wave velocity structure beneath the High Lava Plains, Oregon, from Rayleigh-wave dispersion inversion. Earth Planet. Sci. Lett. 274, 121–131. Weeraratne, D.S., Forsyth, D.W., Fischer, K.M., 2003. Evidence for an upper mantle plume beneath the Tanzanian craton from Rayleigh wave tomography. J. Geophys. Res. 108 (B9), 2427–2446. doi:10.1029/2002JB002273. Weeraratne, D.S., Forsyth, D.W., Yang, Y., Webb, S.C., 2007. Rayleigh wave tomography beneath intraplate volcanic ridges in the South Pacific. J. Geophys. Res. 112, B06303. doi:10.1029/2006JB004403. Wells, R.E., Weaver, C.S., Blakely, R.J., 1998. Fore-arc migration in Cascadia and its neotectonic significance. Geology 26, 759–762. Wernicke, B., Axen, G.J., Snow, J.K., 1988. Basin and Range extensional tectonics at the latitude of Las Vegas, Nevada. Bull. Seismol. Soc. Am. 100, 1738–1757. Wooden, J.L., Mueller, P.A., 1988. Pb, Sr, and Nd isotopic compositions of a suite of Late Archean, igneous rocks, eastern Beartooth Mountains: implications for crust– mantle evolution. Earth Planet. Sci. Lett. 87, 59–72. Wyld, S.J., Wright, J.E., 2001. New evidence for cretaceous strike slip faulting in the United States cordillera and implications for terrane-displacement, deformation patterns, and plutonism. Am. J. Sci. 301, 150–181. Wyld, S.J., Umhoefer, P.J., Wright, J.E., 2006. Reconstructing northern Cordilleran terranes along known Cretaceous and Cenozoic strike-slip faults: Implications for the Baja British Columbia hypothesis and other models. In: Haggart, J.W., Enkin, R.J., Monger, J.W.H. (Eds.), Paleogeography of the North American Cordillera: Evidence For and Against Large-Scale Displacements: Geological Association of Canada, Special Paper, 46, pp. 277–298. Xue, M., Allen, R.M., 2006. Origin of the Newberry Hotspot Track: evidence from shearwave splitting. Earth Planet. Sci. Lett. 244, 315–322. Xue, M., Allen, R.M., 2010. Mantle structure beneath the western U.S. and its implications for convection processes. J. Geophys. Res. 115, B07303. doi:10.1029/ 2008JB006079. Yang, Y., Forsyth, D.W., 2006. Regional tomographic inversion of the amplitude and phase of Rayleigh waves with 2-D sensitivity kernels. Geophys. J. Int. 166, 1148–1160. Yang, Y., Forsyth, D.W., 2008. Attenuation in the upper mantle beneath Southern California: physical state of the lithosphere and asthenosphere. J. Geophys. Res. 113, B03308. doi:10.1029/2007JB005118. Yang, Y., Ritzwoller, M., 2008. Teleseismic surface wave tomography in the western U.S. using the Transportable Array component of USArray. Geophys. Res. Lett. 35, L04308. doi:10.1029/2007GL032278 Yang, Y., Ritzwoller, M., Lin, F., Moschetti, M., Shapiro, N., 2008. Structure of the crust and uppermost mantle beneath the western United States revealed by ambient noise and earthquake tomography. J. Geophys. Res. 113, B12310. doi:10.1029/ 2008JB005833. Yuan, H., Dueker, K., 2005. Teleseismic P-wave tomogram of the Yellowstone plume. Geophys. Res. Lett. 32, L07304. doi:10.1029/2004GL022056. Zhou, Y., Dahlen, F.A., Nolet, G., 2004. Three-dimensional sensitivity kernels for surface wave observables. Geophys. J. Int. 158, 142–168.

.