5.1. The Water Content of the Mantle Beneath Big Pine
 The significant water concentrations in primitive BPVF melts (>1.5 wt% H2O) require a source of water in the mantle. The results of the melting model yield a wide range of results, from 250 ppm (at the upper end of MORB-OIB mantle, 100–300 ppm [Dixon et al., 2002]) to 1900 ppm H2O (more typical of back-arc basins [Kelley et al., 2006]). Independent constraints come from a classic inversion of trace elements carried out by Ormerod et al.  following the method in Hofmann and Feigenson . In such an inversion, the variations in a large suite of trace elements are assumed to be related by varying extents of partial melting, and the batch melting equation is co-solved for element pairs to provide an independent estimate of the relative partition coefficient and source concentration for each element. We can use their results to obtain an independent estimate of the H2O content of the source, based on their results for the REE. Taking into account different modal mineralogy (including garnet), Ormerod et al.  predict an average Ce/Y ratio of the mantle source of 1.4. We can arrive at a minimum Ce concentration in the BPVF source of 5.6 ppm by assuming depleted mantle Y concentration of 4.1 ppm [Salters and Stracke, 2004]. Given the restricted range in the H2O/Ce ratio of Big Pine melt inclusions (164 ± 51), and the lack of fractionation between H2O and Ce during crystallization or melting of anhydrous pyroxenes and olivines [Hauri et al., 2006], we calculate a minimum H2O concentration in the Big Pine mantle of 925 ± 280 ppm. This minimum estimation falls near the mid-range of the permissible H2O concentrations to satisfy the wet melting models above (i.e., 1075 ppm). Using the Ce source concentration above, and a D(Ce) of 0.015 (based on values in Kelemen et al. ) and an average Ce concentration in Big Pine primary liquids of 97 ± 18 ppm for >500 ka volcanics and 78 ± 30 ppm for <500 ka volcanics (data from Blondes  and this study, with olivine added until in equilibrium with Fo90), the degree of melting is calculated to be 4.3 + 0.9/−1.3% and 5.8 + 4.5/−2.1%, respectively. This range in melt fraction (3.4–10.3%) not only is within the range of that calculated from the wet melting models (1–10 wt%), but also provides a better-defined mean around 4–6%. Thus, independent methods involving major, volatile and trace element data from lavas and melt inclusions, are consistent with average melting conditions beneath the BPVF involving mantle with ∼1000 ppm H2O melting to ∼5% F. There are likely real differences in F and source H2O with time, but given the uncertainties in the assumptions and models, we consider below only the implications of the average values.
 The above calculation provides a minimum H2O content of the BPVF mantle (925 +/− 280 ppm) that rules out normal upper mantle as a source (100–300 ppm H2O [Dixon et al., 2002]), and even exceeds the water storage capacity for nominally anhydrous minerals in the convecting upper mantle (<500 ppm H2O [Hirschmann et al., 2009]). Thus, an additional source of water is required. There are several possibilities. One is crustal contamination. This seems unlikely at the BPVF, where melt inclusions with as much as 2 wt% H2O occur in near primary Fo89 olivines (e.g., the ones from Jalopy). There is no correlation between the H2O content of melt inclusions, major elements, and Fo content of host olivines, as might be expected from a crustal assimilation process. Also, BPVF magmas bear mantle xenoliths [Ormerod et al., 1991; Beard and Glazner, 1995; Blondes et al., 2007; Kirby et al., 2008], and the lavas with clinopyroxene phenocrysts record mantle equilibration pressures [Mordick and Glazner, 2006]. Therefore, BPVF magmas had minimal residence in the crust. We thus find the crust an unlikely source of the excess H2O. Mantle lithosphere may also be considered a source of water, in the form of hydrous minerals like amphibole and phlogopite, but it is noteworthy that none of the mantle xenoliths from Big Pine contain any hydrous phases [Beard and Glazner, 1995]. Moreover, our new water data for lithospheric mantle peridotite xenoliths demonstrate unusually dry mantle lithosphere beneath Big Pine, with bulk H2O contents (<75 ppm; Figure 7) well below any of the above estimates (250–1900 ppm) for the melt source. Indeed, the mantle lithosphere beneath Big Pine appears to have been cold and dry, which may be why it has survived as an isotopically distinct layer since perhaps the Proterozoic [Beard and Glazner, 1995; Lee et al., 2000].
 We suggest that the excess water recorded in Big Pine magmas derives from the sub-lithospheric mantle. In fact, at the temperatures inferred for magma formation, mantle that contains significant water is weak and will flow as asthenosphere [Hirth and Kohlstedt, 1996; Karato and Jung, 1998; Karato, 2003]. There are three possibilities for the source of excess water in the mantle beneath Big Pine: hydrous minerals in the asthenosphere, subducting oceanic crust and the transition zone. Phlogopite is a Mg-rich mica and K-richterite is a K-rich amphibole, both of which contain wt% concentrations H2O in their structures and may be stable at pressures >4 GPa along a normal adiabat (Tp = 1350°C) in the upper mantle [Trønnes, 2002]. Phlogopite in the mantle beneath Big Pine may be responsible for the high time-integrated Rb/Sr reflected in the high87Sr/86Sr of the Big Pine magmas [Ormerod et al., 1991], and in the source of potassic volcanism that occurred in the Pliocene of the southern Sierras [Farmer et al., 2002]. On the other hand, both phlogopite and K-richterite have a diagnostic H2O/K2O ratio of 0.38 (from stoichiometry) that can be used to test for their presence in the source of BPVF magmas. In a plot of H2O/K2O vs 1/K2O [after Wallace and Anderson, 1998], BPVF melt inclusions appear to mix to phlogopite-K-richterite for the more K-rich end-member, but they clearly mix to another, higher H2O/K2O component that is not part of the normal N-MORB-E-MORB array (Figure 10a). Arcs and back-arc basin magmas plot to very high H2O/K2O (>7) and may relate to the other end-member. Thus, while phlogopite-K-richterite may exist in the source of BPVF magmas, another water source is still required.
Figure 10. Geochemical constraints on the source of water in BPVF magmas. (a) H2O/K2O versus 1//K2O systematics [after Wallace and Anderson, 1998] suggest mixing between a deep volatile rich phase (Phlogopite/K-Richterite) and an arc source for BPVF melt inclusions. (b) Melt inclusions from active volcanic arcs have uniformly higher Pb/Ce and H2O/Ce than MORB-OIB, generally ascribed to a H2O-Pb rich slab fluid or melt. Melt inclusions from the BPVF have high Pb/Ce like some arcs, but lower H2O/Ce than most MORB-OIB. Some continental arcs (Cascades-Mexico) mix toward a mantle end-member with a high Pb/Ce like the BPVF. Blue circles are least degassed melt inclusions from BPVF (as inFigure 6). Gray circles include H2O calculated from Cl and H2O/Cl = 66 (as in Figure 6). Arc and back-arc data fromGribble et al. ; Sadofsky et al. ; Johnson et al. ; Zimmer et al. ; and Ruscitto et al. .
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 As discussed above, the H2O contents of Big Pine magmas approach those typical of back-arc basin basalts, where water is actively supplied from subduction zones [e.g.,Kelley et al., 2006]. Indeed, BPVF melt inclusions plot in a distinct region on a H2O/Ce-Pb/Ce diagram, with H2O/Ce similar to upper mantle asthenosphere (MORB-OIB), but with clearly higher Pb/Ce, like volcanic arcs. Interestingly, continental arcs like the Cascades, Guatemala, and Mexico plot toward a mantle end-member with Pb/Ce higher than MORB-OIB, in the range of our new data from the BPVF (Figure 10b).
 The trace element signatures of Big Pine magmas also share similarities with arc magmas (i.e., excess Pb, depletions in Nb and Ta; Figure 5). The dehydration of subducted material along slab P-T paths is an obvious supply of water to mantle above subducted slabs, but there is no subduction zone beneath Big Pine at this time. On the other hand, the shear wave tomography inFigures 11a and 11b illuminates the Isabella Anomaly, a prominent fast seismic anomaly located immediately to the west of the Sierra Nevada and BPVF. The Isabella Anomaly has been interpreted as actively foundering Sierran lithospheric mantle or lower crust [e.g., Zandt et al., 2004; Frassetto et al., 2011], or as a fragment of the Farallon Plate that did not detach [Pikser et al., 2012; Y. Wang et al., Fossil slabs attached to unsubducted fragments of the Farallon Plate, submitted to Nature Geoscience, 2012]. This feature could drive mantle upwelling and melting, either as counterflow to the foundering drip, or induced flow as it is dragged northwest with the Pacific Plate. The Isabella Anomaly is one of the largest seismic anomalies in the western USA, comparable in magnitude to the actively subducting Gorda Plate to the north [Schmandt and Humphreys, 2010, 2011]. It is also possible this feature is still dehydrating, although subduction ceased at 20 Ma. In subduction zones, the K2O/H2O and H2O/Ce of slab fluids is related to their last temperature of equilibration at the slab surface [Plank et al., 2009]. Interpreted in this way, the Isabella slab fluids would fall at the high temperature end (>950°C) of fluids that supply active subduction zones, consistent with the heating up of this slab remnant during long residence in the mantle.
Figure 11. Surface wave tomography beneath the Basin and Range and possible dynamic melting models for the BPVF. (a) Shear velocities at 70 km depth from Rayleigh-wave surface tomography [Rau and Forsyth, 2011]. Note the prominent Isabella Anomaly adjacent to BPVF. Melting contours and BPVF melt equilibration depths (red and blue bars) from Figure 9. LAB is the lithosphere-asthenosphere boundary deduced from the shear-velocity structure. Melting beneath the BPVF driven by (b) Extension-melt feedback, (c) small-scale convection related with lithospheric delamination and (d) the Isabella Anomaly as a slab remnant and small-scale convection related with it (see text for discussion). Mantle shear direction fromConrad et al. .
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 Another possible source of water beneath Big Pine is from mantle that may upwell from the transition zone. The storage capacity of water in transition zone minerals is considerably higher (0.5–1 wt% [Ohtani, 2005]) than in the convecting upper mantle (<500 ppm H2O [Hirschmann et al., 2009]) and so mantle that upwells from the transition zone will likely melt [Bercovici and Karato, 2003]. Assuming the maximum water storage, these melts will be water rich (>17 wt% H2O [Hirschmann et al., 2009]), and may ascend and supply water to the primary melting region beneath BPVF. Although there is no obvious dynamic reason why mantle might upwell from the transition zone beneath Big Pine, the high water contents in magmas throughout the Basin and Range [Plank et al., 2009] may require it, and mantle convection models may predict it as part of a large counterflow to Farallon downwelling [Moucha et al., 2008].
5.2. Implications of the Pressure and Temperature Estimates for Rheological Boundaries of Melt Equilibration
 Although mantle-melt thermobarometry provides valuable quantitative constraints on conditions within the upper mantle, the single pressure and temperature recorded in each magma composition is not always straightforward to connect to a melting process. One possibility is that magmas reflect only the final P and T of equilibration in the mantle. Melting prior to this point may have proceeded as a batch (melt remains with solid) or fractional (melt separates from solid) process, but if melt aggregates and equilibrates in one region, it will only reflect the final conditions. The other possibility is that the P-T recorded by magmas reflects a mean of the melts that have equilibrated at different depths. We find this latter scenario unlikely for the <500 ka BPVF magmas, given their very shallow pressures of equilibration (1.1 GPa or 38 ± 7 km, essentially at the Moho) and the lack of garnet signature in the magmas, meaning at least, that the contribution of deep melts to the mixture is minor.
 The pressures recorded in both series of melts may correspond to mechanical boundaries in the melting region. One is the lithosphere-asthenosphere boundary (LAB), which is unusually shallow, ∼55–60 km as deduced from the maximum gradient in the shear velocity profile (Figure 11). Although the relatively high velocity (lithospheric) layer immediately beneath the Moho is too thin to be well resolved from the surface wave tomography alone[Rau and Forsyth, 2011], there is an apparent negative Ps conversion beneath Big Pine observed at ∼55 km from receiver function analysis [Frassetto et al., 2011], also consistent with a shallow LAB. This boundary fundamentally separates the depths of equilibration of the two age groups of magmas. The older magmas (>500 ka) equilibrated at >65 km (70 km on average), possibly stalling at an older, slightly deeper LAB, at that time. The younger magmas (<500 ka) equilibrated uniformly above this boundary, in the inferred lithosphere, and on average at 38 ± 7 km, which is equivalent to crustal thickness here (of ∼36 km) determined from receiver function analysis [Frassetto et al., 2011]. Thus, the older magmas equilibrated within asthenosphere, possibly at the base of an earlier ∼70 km LAB, while the younger magmas have equilibrated above the modern ∼55 km LAB, on average at depths close to the Moho. This is consistent with the fact that both the LAB and Moho are rheological or density boundaries, which may promote melt stalling and equilibration of melts that may have initiated at a much deeper solidus.
 It is important to consider that the younger magmas, which have equilibrated in the lithosphere near the Moho, clearly did not form there, as their temperatures are too high (1220°C) to sustain at the Moho without massive melting of the crust. Also, their temperatures exceed those recorded by the Oak Creek mantle xenoliths, 1000–1100°C at <1.5 GPa as reported by Lee et al.  (although Ducea and Saleeby  report up to 1200°C). Moreover, the lithosphere is too dry to be the source of these magmas, as recorded in the low water contents (<100 ppm H2O) of the Oak Creek xenoliths (Figure 7), and the lack of hydrous minerals contained within them [Beard and Glazner, 1995]. Thus, the extent of re-equilibration within the lithosphere is partial, reflected in the pressures (which only requires precipitation of olivine), but not the temperatures (which is retained in the transported melts) nor water contents (which would require equilibration with larger volumes of dry mantle). It is also important to keep in mind that the seismic results represent averages over significant areas; very localized thinning or erosion or infiltration of the lithosphere by melt would not be recognized.
5.3. Geochemical Indicators of Lithosphere Versus Asthenosphere Mantle Sources in the Evolution of BPVF
 Our thermobarometry results, in combination with the seismic structure of the upper mantle, point to different melt equilibration scenarios for the older and younger BPVF suites, with the >500 ka magmas equilibrating below the LAB in the asthenosphere, and the <500 ka magmas equilibrating within the lithosphere. This interpretation differs from previous work that considered all Big Pine magmas to have a lithospheric source, based on their Proterozoic Sm-Nd model ages, their high87Sr/86Sr (0.7054–0.7064), and their arc- or continent-like trace element signatures [Ormerod et al., 1988, 1991]. On the other hand, based on our new data here and that in Blondes et al. , we find a systematic variation in diagnostic trace element ratios in both space and time. The hotter and deeper melts (>0.5 Ma) have higher Ce/Pb (14–21) and constant Ba/La ∼ 20 values, closer to an upper mantle asthenosphere end-member (Figures 12a and 12b). Shallower and cooler melts (<0.5 Ma) have lower Ce/Pb (<14), and higher Ba/La (25–30) values, more typical of arc magmas and consistent with the available data from lithospheric mantle xenoliths from BPVF [Lee, 2005], suggesting a lithosphere component for these melts (Figures 12a and 12b). In the data set as a whole, there is a highly significant (R2 = 0.85) correlation between Ce/Pb and the depth of equilibration, from 90 to 35 km. Thus there is a clear contribution from old, chemically distinct mantle that increases at shallower levels, consistent with a chemical boundary layer above the seismically imaged LAB (55–60 km depth, Figures 11b and 12a). On the other hand, the temperatures (>1200°C) and water contents (>1.5 wt% H2O) of all the magmas seem inconsistent with formation in the cold and “dry” lithosphere.
Figure 12. Shear wave velocity profiles compared with depth of melt equilibration and the geochemical evolution of BPVF. (a) Shear wave profile (from Rau and Forsyth , Rayleigh-wave inversion) and depth of equilibration for <500 ka and >500 ka BPVF melts. (b) Strong correlation between melting depths and the Ce/Pb ratio in BPVF melt inclusions and volcanics. The ratio decreases with time, toward values more typical of lithospheric mantle xenoliths [fromLee, 2005] and subduction zone magmas. (c) Melts >500 ka have constant Ba/La ∼ 20, closer to upper asthenosphere values, while the samples of <500 ka melts have higher and variable Ba/La (25–30), also approaching subduction zone magmas. These strong correlations between trace element ratios and melt depths delineate a thin chemical boundary layer in the upper mantle that may coincide with the seismic LAB. With time, melts interact more extensively with this shallow lithosphere. Upper asthenosphere (MORB-OIB) values from Georoc database (http://georoc.mpch-mainz.gwdg.de/georoc/).
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 Although the correlation is clear with trace elements (Ce/Pb, Ba/La), isotopic compositions do not vary systematically with depth, except within the Papoose section [Blondes et al., 2008]. The petrologic modeling together with the trace element correlations collectively suggest that both asthenosphere and lithosphere have played a role, and possibly evolved, over the formation of the BPVF magmas. It is possible that mantle interpreted as asthenosphere today (at 90 to 60 km depth from shear wave tomography in Figure 11b) was evolved from lithospheric mantle during melt infiltration and re-heating, thus explaining the old model ages of all BPVF magma sources [Beard and Glazner, 1995].
5.4. The Cause of Mantle Melting at Big Pine
 Together, the petrological and seismological data provide a consistent view of the melting region beneath Big Pine. The base of the low velocity zone is around 225 km beneath Big Pine (Figure 11b), which could correspond to the carbonatite-carbonated silicate melt transition [afterHirschmann, 2010]. This would occur at ∼8 GPa (250 km) for mantle with 1000 ppm H2O and 250 ppm CO2 km and F = 0.1% (note that primary CO2 is generally not possible to constrain even from melt inclusions, which may already be saturated in a CO2-rich vapor when trapped; 250 ppm CO2 in the mantle is twice that needed to explain maximum CO2 contents in BPVF magmas, and so permissible and not unreasonable). The region of the lowest seismic velocities beneath Big Pine (<4.2 km/s) begins at around 120 km, which could correspond to the onset of significant silicate melting (F = 1% in mantle with 1000 ppm H2O at 4 GPa, 140 km). Melting then continues up to the base of the 500 ka LAB (∼70 km), producing ∼5% melt. Although the melting boundaries are tens of km deeper than the seismic boundaries, these differences are not likely significant given the uncertainties in both inversions. The >500 ka magmas then erupted rapidly above this depth, possibly by diking in cold lithosphere, carrying mantle xenoliths to the surface, with little residence in the crust. With time, the mantle lithosphere warms, magmas stall, react and equilibrate there. This leads to an erosion of the lithosphere, and possibly shallowing of the LAB from 70 to 55 km in ∼500 ka. Crustal storage regions are likely near the brittle-ductile transition (∼20 km), where magmas stall due to a rheological contrast, melt inclusions start to be trapped by crystal cooling (at ≤5 kb;Figure 6), magmas fractionate, and no longer bear mantle xenoliths. Volcanic vents form very near active faults in Owens Valley [Kirby et al., 2008], which enhanced ascent pathways.
 Despite this self-consistent view of the melting process, many questions still remain as to the ultimate cause of volcanism here. Why do mafic magmas erupt at Big Pine? What is special about the crust and mantle that leads to melting and eruption? There are many special tectonic features and events that appear to characterize the Big Pine region. (1) Active oblique strike-slip faulting and extension in Owens Valley [Phillips and Majkowski, 2011]; (2) proposed Pliocene foundering of adjacent Sierran lithosphere [Manley et al., 2000; Jones et al., 2004; Zandt et al., 2004]; (3) and the prominent Isabella anomaly to the west and thicker LAB to the east (Figure 11b). Which of these are the critical drivers for melting and eruption of mafic magmas here? And how can they be consistent with the constraints we have provided on the pressures, temperatures, water contents and evolution of the mantle melting region? We provide two views on melting, one top-down (1. Extension-Melt Feedback) and the other bottom-up (2. Small-Scale Convection).
 1. Extension-melt feedback. In the first view, lithospheric extension is the initial driver for melting and eruption. Owens Valley, where Big Pine magmas erupt, is a graben that has developed largely since 3.5 Ma, with active normal and oblique strike slip faults [Phillips and Majkowski, 2011]. Most of the modern motion across the valley (determined geodetically) drives strike-slip on N-S faults, although as much as a third of the total motion is driving extension at a rate of ∼1.5 mm/yr [Phillips and Majkowski, 2011]. Applied over the past 3.5 Ma, this extension could account for ∼5 km of opening in Owens Valley. Although active and significant, this extension (1.5 km/Ma) is more than an order of magnitude too small to drive either the above proposed LAB shallowing (30 km/Ma) or the apparent regional-scale thinning of the LAB from 90 km to the east of Big Pine to 55 km beneath it (Figure 11b). On the other hand, it is possible that feedbacks between thinning and melting lead to progressive conversion of lithosphere to asthenosphere, in a thermal and chemical corrosion process similar to that outlined by Holtzman and Kendall . Initial extension leads to some enhanced melt production below the LAB, followed by melt infiltration, reaction and diking within the lithosphere that cause it to weaken and thin further. Further melting leads to additional feedbacks as lithosphere topography drives stress-driven melt segregation [Holtzman and Kendall, 2010] or shear-driven upwelling [Conrad et al., 2010, 2011], which leads to further corrosion, thinning, and eventually upwelling and decompression melting. It is possible this process commenced at 3.5 Ma, coinciding with the onset of the current strain regime in Owens Valley, and that ∼2 Ma were necessary for lithosphere-to-asthenosphere conversion to progress to a critical stage for melting, segregation and eruption of the first Big Pine magmas at 1.3 Ma. Today, asthenospheric melts have been generated up to ∼60 km depth, above which the lithosphere is actively being infiltrated and warmed. Melts are now equilibrating there, and no longer diking through this region.
 This view of the origin of the BPVF fundamentally relates melting to active lithospheric deformation. Deformation initiates upwelling that drives melting and starts a feedback process that also creates melt pathways to the surface. The broader implication here is that magmas in the Basin and Range will tend to form and erupt where deformation is active. This is true in a broad sense (most of modern Basin and Range deformation and volcanism is occurring at its margins; along the eastern side of the Sierras (as in Big Pine) and along the western side of the Colorado Plateau [Bennett et al., 1998; Hammond and Thatcher, 2004]. Moreover, seismic images are providing abundant evidence for the presence of melt in the mantle over much of the Western USA, but in most places, it does not manifest at the surface (e.g., in the Amagmatic Zone of southern Nevada [Rau and Forsyth, 2011]), possibly because active deformation does not provide the melt enhancement, segregation and ascent pathways required for eruption. Further work linking melting in the mantle to surface deformation and volcanism across the Basin and Range will test these ideas. On the other hand, this view of deformation and melting does not provide a ready source of the excess water that is required in the source of Big Pine magmas. Normal Farallon subduction did occur in this region as recently as 18–20 Ma, and likely hydrated the mantle, but as we have argued above, the water concentration reflected in the source of Big Pine exceeds the normal storage capacity of anhydrous mantle at TP = 1350°C (500 ppm), and phlogopite, which can be stable at these temperatures, cannot supply all the water (based on K2O/H2O, Figure 10).
 2. Small-scale convection. The other driver of melting could be related to lithospheric drips and the Isabella anomaly to the west of the BPVF (Figures 11c and 11d). The Isabella anomaly is a fast seismic anomaly that extends to at least 200 km depth, and is located beneath the western Sierra foothills and Great Valley. Its origin is debated. Some ascribe it to Sierran lithosphere (garnet clinopyroxene lower crust and peridotitic mantle) that foundered in the Pliocene [Zandt, 2003; Ducea and Saleeby, 1998; Jones et al., 2004] while others link it to the fossil Monterey microplate, a remnant of the Farallon Plate that was left when subduction ceased ∼20 Ma (Wang et al., submitted manuscript, 2012). In the latter view, the Isabella anomaly is a slab fragment, still attached to the Pacific lithosphere, and possibly being dragged northwest along with it. In either case, a sinking or dragging structure (lithospheric drips or remnant subducting oceanic crust) may induce flow in surrounding mantle, either in an upward counterflow or in a small-scale convective circulation, in combination with the high mantle shear in the western Basin and Range [Zandt et al., 2004; Elkins-Tanton et al., 2001; Conrad et al., 2010]. It is also possible that the Isabella anomaly could supply excess water to the mantle melting beneath Big Pine, through heating-induced dehydration of hydrous Sierran lower crust or of hydrous Farallon slab. In this view, the thin LAB beneath Big Pine was not generated by extension, but by lithospheric foundering in the Pliocene, and melting occurs today due to an influx of water and convective upwelling. In this way, the BPVF is situated above a long-lived mantle wedge, like an arc, a southern extension of the ancestral Cascades (sensu [Cousens et al., 2008]). Although it is unclear whether Mesozoic arc crust or a slab subducted 20 my ago could retain enough water to supply the BPVF, the Isabella anomaly does appear spatially related to volcanism along its eastern margin, from Long Valley to Coso. On the other hand, this volcanism does not include the high-Mg# andesite compositions of bajaites erupted in the south, where similar Guadalupe and Magdalena Farallon remnants have been proposed (Wang et al., submitted manuscript, 2012), nor the much lower87Sr/86Sr of the ancestral or modern Cascades erupted in the north and related to Gorda and Juan de Fuca subduction [Cousens et al., 2008]. The isotopic composition of Big Pine magmas overlaps that of Sierra Nevada granites. The role of the Isabella anomaly as a water source remains to be tested by study of the water contents across the Basin and Range, to see how widespread or locally “wet” the mantle is.