GEOLOGY | Volume 46 | Number 3 | www.gsapubs.org 1

Moho temperature and mobility of lower crust in the western United StatesDerek L. Schutt1, Anthony R. Lowry2, and Janine S. Buehler3

1Department of Geosciences, Colorado State University, Fort Collins, Colorado 80523, USA2Department of Geology, Utah State University, Logan, Utah 84322, USA3Scripps Institution of Oceanography, University of California–San Diego, La Jolla, California 92037, USA

ABSTRACTWe use measurements of mantle P-wave velocity from the Moho refracted phase, Pn, to

estimate temperature within the uppermost few km of the western U.S. mantle. Relative to other approaches to modeling the deep geotherm, using Pn velocities requires few assump-tions and provides a less uncertain temperature at a tightly constrained depth. Assuming a homogeneous mantle composition, Moho temperatures are lowest in an arc that extends from the High Lava Plains through western Montana and the high-plains region of Wyoming and western Kansas/Nebraska. Highest temperatures are observed under recent (<10 Ma) volcanic provinces and are consistent with melting. Estimates of lower crustal viscosity suggest that the western U.S. west of the Laramide deformation front likely has regions of mobile lower crust that decouple upper crustal and upper mantle tractions.

INTRODUCTIONA fundamental factor in determining litho-

spheric strength, thickness, buoyancy, and mass flux is temperature. Temperature varia-tions modulate stress and deformation of the lithosphere (Harris et al., 2004; Hyndman et al., 2009; Lowry and Pérez-Gussinyé, 2011), and temperature of the lower crust, which flows if sufficiently warm (e.g., Bird, 1991), is of par-ticular interest. A weak lower crust can decouple upper crustal and mantle strain fields, affecting the style of deformation. The prevalence and expression of lower crustal flow remains hotly debated (e.g., Bürgmann and Dresen, 2008), but constraint of Moho temperatures would be a first step toward greater understanding.

The importance of temperature to lithosphere dynamics has motivated several approaches to estimating deep geotherms, including extrapo-lating surface heat flow measurements to depth (e.g., Pollack and Chapman, 1977; Artemieva and Mooney, 2001); using elevation and/or isostasy (e.g., Hasterok and Chapman, 2007); xenolith geothermometry and geobarometry (e.g., Artemieva, 2009); and mapping of seis-mic shear velocities (VS) to temperature (Goes and van der Lee, 2002). These methods all have their limitations: xenolith sampling is sparse, surface heat flow measurements are perturbed by surface hydrology and sensitive uncertain distribution of radiogenic heat production, and shear wave velocity is sensitive to melts and anelasticity in addition to temperature. Combin-ing complementary observables is an increas-ingly common practice; however, in doing so, sources of uncertainty and the effects of mul-tiple assumptions can be obscured.

Temperature from Pn (a P-wave refracted along the crust-mantle boundary) velocity mod-els is a useful independent complement to exist-ing methods that provides tight constraints on lower crustal temperatures. Mantle composi-tional variations are muted relative to those in the crust, and variations in peridotite mineral assemblages cause much smaller P-wave veloc-ity (VP) perturbations than do plausible Moho temperature variations (e.g., Afonso and Schutt, 2012). Moreover, because Pn traverses the cold-est part of the mantle, the mapping of velocity to temperature is less affected by uncertainties in anelastic effects, which begin to grow large above 900 °C (e.g., Jackson and Faul, 2010).

METHODBuehler and Shearer (2012, 2014, 2017) pro-

duced isotropic and azimuthally anisotropic Pn velocity models using ~206,000 first-arrival Pn travel times collected by EarthScope’s USArray (Trabant et al., 2012). We use their most recent isotropic velocity and crustal thickness models for this analysis (Fig. 1).

To convert Pn velocities to temperature, we use as a null hypothesis a uniform spinel lher-zolite mantle composition (MM3; Baker and Stolper, 1994; Lesher and Baker, 1997). While temperature-dependent metamorphic changes in mineral mode will occur, these have a minimal effect on VP (Schutt and Lesher, 2006). More relevant are changes in the whole rock chemistry. Variations in peridotite composition resulting from melt removal have little effect on VP (e.g., Afonso and Schutt, 2012; Schutt and Lesher, 2006). However, orthopyroxene enrichment may reduce velocities by 1%–3% (Schutt and Lesher,

2010; Wagner et al., 2008). Eclogitic crust attached to stranded subducted slabs (Schmandt and Humphreys, 2011) may also modify veloci-ties (Levander et al., 2006; Worthington et al., 2013). If the mantle is hydrated, minerals such as pargasite or chlorite (e.g., Comodi et al., 1991; Fumagalli and Poli, 2005), or water in olivine (Jacobsen et al., 2008; Jacobsen et al., 2009), will reduce velocities. These possibilities are considered in the discussion.

Density and elastic moduli variations are calculated using the method and parameters of Schutt and Lesher (2006, 2010). Anelastic effects are based on an empirical study of anelasticity in olivine (Jackson and Faul, 2010). This study infers that grain size significantly affects anelas-ticity. Grain sizes in lithospheric xenoliths follow a power-law distribution from roughly 0.2 to 25 mm (Armienti and Tarquini, 2002; Tabor et al., 2010). However, because few studies have sys-tematically investigated lithospheric grain size,

GEOLOGY, March 2018; v. 46; no. 3; p. 1–4 | GSA Data Repository item 2018046 | https://doi.org/10.1130/G39507.1 | Published online XX Month 2018

© 2018 Geological Society of America. For permission to copy, contact editing@geosociety.org.

Figure 1. Pn (a P-wave refracted along the crust-mantle boundary) velocities used (Buehler and Shearer, 2017). States mentioned in the text are in white: Colorado (CO), Kansas (KS), Montana (MT), Nebraska (NE), Oregon (OR), Washington (WA), and Wyoming (WY). Tectonic and geographic locations: Bighorn Mountains (BgH), Black Hills (BH), Cascade Volcanoes (CV), Colorado Plateau (CP), Colo-rado Rocky Mountains (CRM), Great Plains (GP), High Lava Plains (HLP), Rio Grande Rift (RGR), Sierra Nevada (SN), Wyoming Craton (WC), and Yellowstone hotspot track (YT). Dashed lines show commonly used tectonic provinces.

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and none consider the effect of a distribution of grain sizes on anelastic effects, the contribution of grain size to anelasticity is poorly constrained. Uncertainties associated with anelastic effects are calculated by conservatively assuming grain size varies in the range of 0.2–25 mm.

RESULTSNinety-five percent (95%) of Moho tem-

perature estimates are between 444 and 892 °C, with extrema at 200 and 1160 °C (Fig. 2). Tem-peratures correlate well with upper mantle body wave tomography (e.g., Schmandt and Hum-phreys, 2011), and temperatures >800 °C cor-relate strongly with occurrences of <10 Ma magmatism (EarthChem database, http://www .earthchem.org /portal).

Temperatures >800 °C are found along the Yellowstone hotspot track, in the Colorado Rocky Mountains, on the southwest margin of the Colorado Plateau, just east of the Rio Grande Rift, and along the Sierra Nevada Mountains and Cascade Volcanoes (see Fig. 1 for geographic locations). Temperatures <500 °C occur along a trend from under the High Lava Plains through the Wyoming Craton and into eastern Kansas and central Texas. Median and mean tempera-tures are 648 and 659 °C, respectively, with a standard deviation of 118 °C. Estimated one-sigma uncertainties in absolute temperature vary laterally, and have a mean of ~77 °C (see the GSA Data Repository1 for details).

The Pn velocities used for this analysis are perturbations to a best fitting uniform starting model, calculated as a least-squares inversion (Buehler and Shearer, 2012). As with most itera-tive nonlinear inversions, the final results are

1 GSA Data Repository item 2018046, supplemen-tary information and Figures DR1–DR8, is available online at http://www.geosociety.org/datarepository /2018/ or on request from editing@geosociety.org.

non-unique and depend upon the starting model and inversion regularization (e.g., Aster et al., 2013). We have run a series of inversions using different starting models to explore the model null space (Fig. 3). Here, we calculate the Pn velocity for an assumed starting temperature, and use this as a starting model for an inver-sion of Pn arrival times following the method of Buehler and Shearer (2013).

Inversions with a 600 °C uniform start-ing temperature tend to perturb toward slower velocity; i.e., in many places the starting velocity model is too fast to fit the observations. If the crustal thickness model is correct, and if com-position does not vary, then the data imply that the western U.S. Moho temperature is >600 °C everywhere southwest of a band that extends

roughly from the High Lava Plains, through the Bighorn Mountains in Wyoming, and into Kan-sas. A uniform 700 °C starting model results in structures similar to the final model shown in Figure 1. Nearly the entire western U.S. Cordil-lera appears to be >700 °C at the Moho, except for the aforementioned high-plains regions. A uniform 800 °C starting model results in only a few areas of perturbation toward lower veloci-ties, corresponding to <10 Ma volcanic prov-inces and the Black Hills. Most regions need higher velocities, implying colder temperatures, to fit the data. These tests suggest that the bulk of the western U.S. uppermost mantle is bounded by 600 °C and 800 °C, with the volcanic areas being either hotter or having partial melt, con-sistent with Figure 2.

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Figure 3. Effect of starting model on Pn (a P-wave refracted along the crust-mantle boundary) velocity inversion results. Starting velocities (left panels) are calculated using the pressure at the base of the crust for a constant temperature, and the velocity change from the start-ing model (right panels) is shown for starting models of 600 °C (top), 700 °C (middle), and 800 °C (bottom).

Figure 2. Moho temperature at depths varying from 20 to 50 km (see also Fig. DR5 [see foot-note 1]). Black dots are locations of <10 Ma volcanism (EarthChem database, http://www .earthchem.org/portal).

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DISCUSSIONThe uppermost mantle under the western U.S.

Cordillera has long been considered anomalously warm (e.g., Lowry et al., 2000), but if MM3 closely approximates the composition of litho-spheric mantle, most of the western U.S. Cor-dillera is not just warm but hot. The regions of >800 °C temperatures are near or above the soli-dus of most presumed lower crustal compositions, and would suggest that pervasive low electrical resistivity of the lower crust in these regions (e.g., Meqbel et al., 2014) indicates partial melt. The presence of melt is not considered in the velocity-to-temperature mapping, so high temperatures will be overestimated where mantle partial melt exists. Hydrous minerals also are present in the upper mantle under parts of the Cordillera (e.g., Dixon et al., 2004), and these would lower veloc-ity, resulting in overestimated temperature.

Whether or not melt and/or hydrous minerals are present, the lower crust can be presumed to be weak and mobile over significant portions of the U.S. west of the Cordilleran deformation front. Figure 4 shows example calculations of effective viscosity at the estimated Moho tem-peratures, given flow law parameters compiled by Bürgmann and Dresen (2008) for various minerals at end-member hydration states. Dis-location creep flow strength, σ, is calculated at pressures derived from crustal thickness (Fig. DR5 in the Data Repository), assuming a mean crustal density of 2960 kg m–3 and a strain rate of 10−14 s–1. Effective viscosity is ηeff = σ / 2 (see the Data Repository for more details). Flow strength will reflect the weakest abundant (>10%–15%) mineral, which in the lower crust is likely to be either quartz or feld-spar. A dry feldspar composition implies lower crust >~800 °C is fluid with viscosity of order 1021 Pa·s (Fig. 4C). Wet and/or quartz-rich lower crust would be even more mobile.

High Moho temperatures east of the Cordil-lera might at first seem counterintuitive. How-ever, this primarily reflects greater crustal thick-ness there (Fig. DR5): geotherms extrapolated to a constant depth exhibit cooler temperatures for stable North American lithospheric mantle, consistent with tomographic imaging suggestive of thicker lithosphere (Schmandt and Lin, 2014). An important corollary is that lithospheric stabil-ity is not defined solely by temperature-depen-dent viscosity of the lower crust: Archean-aged cratonic mantle lithosphere with high flexural rigidity (e.g., Lowry and Pérez-Gussinyé, 2011) and strong resistance to tectonism, such as in the northern Great Plains region, has a Moho temperature similar to that of actively deforming regions of the Cordillera (Fig. 2), while the very low rigidity northern Rocky Mountains have cold lower crust. Flexural rigidity variations in North America indicate that stable versus mobile lithosphere reflects a stark difference in upper mantle strength (e.g., Lowry et al., 2000),

so geotherms carefully constructed using Pn hold great promise as a foundation for future explorations of compositional controls on litho-spheric strength and stability.

The Colorado Plateau exhibits anomalously high temperatures along its margins, consis-tent with thermal advection by melts encroach-ing laterally toward the interior (e.g., Roy et al., 2016). The northwestern Siletzia province (western Oregon and Washington) may be an ophiolitic block still attached to an abandoned slab on its eastern edge (Schmandt and Hum-phreys, 2011). Hence cold temperatures and also perhaps eclogite may be present, with VP

a few percent higher than peridotite at identical pressure-temperature conditions (Worthington et al., 2013). High VP/VS and null anisotropy in the central High Lava Plains also support an eclogite interpretation (Buehler and Shearer, 2014). The coldest temperatures in the Wyoming Craton seem inconsistent with steady-state geo-therms extrapolated from surface heat flow mea-surements (Blackwell et al., 2011) and may be a compositional anomaly, perhaps due to rem-nant eclogite from Archean cratonic assembly or garnet-rich pyroxenite from past melt migration. Alternatively, this may reflect transient thermal behavior. Under the Sierra Nevada, Pn velocity

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may also be modified by a compositional com-ponent (e.g., Zandt et al., 2004). Metasomati-cally altered mantle in central Montana (Downes et al., 2004) could provide a compositional com-ponent to lower velocities there.

ACKNOWLEDGMENTSWe thank Derrick Hasterok, an anonymous reviewer, and editors Dennis Brown and Bob Holdsworth for constructive and insightful reviews. This work was supported by National Science Foundation grants EAR-1358664 (Schutt), EAR-0955909, EAR-1358622 (Lowry), and EAR-1358510 (Buehler).

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Manuscript received 9 July 2017 Revised manuscript received 8 November 2017 Manuscript accepted 1 December 2017

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