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Click Here GEOPHYSICAL RESEARCH LETTERS, VOL. 35, L10303, doi:10.1029/2008GL033391, 2008 for Full Article Mantle structure beneath the western edge of the Colorado Plateau C. R. Sine,1 D. Wilson,2 W. Gao,3 S. P. Grand,4 R. Aster,5 J. Ni,6 and W. S. Baldridge7 Received 24 January 2008; revised 28 February 2008; accepted 10 March 2008; published 20 May 2008. [1] Teleseismic traveltime data are inverted for mantle Vp and Vs variations beneath a 1400 km long line of broadband seismometers extending from eastern New Mexico to western Utah. The model spans 600 km beneath the moho with resolution of 50 km. Inversions show a sharp, largemagnitude velocity contrast across the Colorado PlateauGreat Basin transition extending 200 km below the crust. Also imaged is a fast anomaly 300 to 600 km beneath the NW portion of the array. Very slow velocities beneath the Great Basin imply partial melting and/or anomalously wet mantle. We propose that the sharp contrast in mantle velocities across the western edge of the Plateau corresponds to differential lithospheric modification, during and following Farallon subduction, across a boundary defining the western extent of unmodified Proterozoic mantle lithosphere. The deep fast anomaly corresponds to thickened Farallon plate or detached continental lithosphere at transition zone depths. Citation: Sine, C. R., D. Wilson, W. Gao, S. P. Grand, R. Aster, J. Ni, and W. S. Baldridge (2008), Mantle structure beneath the western edge of the Colorado Plateau, Geophys. Res. Lett., 35, L10303, doi:10.1029/2008GL033391. 1. Introduction [2] The tectonic history of the southwestern United States is long and varied. Accretion of exotic arcs from 1.8 to 1.1 Ga formed the initial craton but subsequent rifting of the Rodinia supercontinent (0.9 – 0.7 Ga) left a nearly N-S margin along western North America near 115° W and extended lithosphere west of the present-day Colorado Plateau (CP) [Burchfiel et al., 1992]. Later accretion to the margin from the west, beginning in the mid-Paleozoic, established a lithospheric contrast along the rifted Rodinian margin persisting to the present-day with major consequences for continental evolution. Subduction of the Farallon plate beneath the western margin of North America was established by the late Mesozoic. Progressive flattening of the slab caused the Laramide Orogeny (75 – 45 Ma) asso1 Occidental Petroleum Corporation of Elk Hills, Tupman, California, USA. 2 USGS Hawaiian Volcano Observatory, Hawaii National Park, Hawaii, USA. 3 TGS NOPEC Geophysical Company, Houston, Texas, USA. 4 Jackson School of Geosciences, University of Texas at Austin, Austin, Texas, USA. 5 Department of Earth and Environmental Science, New Mexico Institute of Mining and Technology, Socorro, New Mexico, USA. 6 Department of Physics, New Mexico State University, Las Cruces, New Mexico, USA. 7 Earth and Environmental Sciences Division, Los Alamos National Laboratory, Los Alamos, New Mexico, USA. Copyright 2008 by the American Geophysical Union. 0094-8276/08/2008GL033391$05.00 ciated with thick-skinned Rocky Mountain uplifts and thrusting. It has been proposed that the flattened Farallon plate foundered around 40– 20 Ma. [Coney and Reynolds, 1977; Humphreys et al., 2003]. Widespread ignimbrite volcanism followed as hot asthenosphere came in contact with lithosphere previously cooled and hydrated during Laramide flat-slab subduction. Heat from asthenospheric upwelling and the gravitational potential of the Sevier crustal welt led to extension and the formation of the Great Basin (GB) with present-day extension of approximately 100% or more at the latitude of this study [Wernicke et al., 1988; Humphreys et al., 2003]. To the east the Rio Grande Rift (RGR) opened a northern extension of the southern Basin and Range with considerably less Cenozoic extension than the GB [Chapin and Cather, 1994]. During the Cenozoic tectonic events discussed above, the CP remained relatively intact but was uplifted about 2 km [Morgan and Swanberg, 1985]. [3] The Rio Grande Rift Seismic Transect (La RISTRA 1.0), shown in Figure 1, consisted of a deployment of broadband seismometers extending NW along a line from west Texas, across the RGR, to the middle of the CP. The instruments, deployed from July 1999 to May 2001 at 18 km intervals, were used to investigate the crust and mantle beneath the eastern CP, RGR and western Great Plains (WGP) [Gao et al., 2004; West et al., 2004; Wilson et al., 2005]. From June 2004 through May 2006, 18 broadband seismometers (La RISTRA 1.5) were deployed from the NW end of the La RISTRA 1.0 array, crossed the CP, and extended into the GB at 20 km intervals. The combined La RISTRA 1.0 and 1.5 experiments facilitate imaging of crust and upper mantle seismic structure across a 1400 km transect encompassing four tectonic provinces in the southwestern U.S.: the WGP, the RGR, the CP, and the GB. In this paper we show the results of a tomographic inversion of P and S wave teleseismic traveltimes that combine data from both deployments and present the first detailed (50 km resolution) teleseismic velocity image of the upper mantle that traverses the entire CP. Gao et al. [2004] presented P and S wave tomograms from La RISTRA 1.0 that extended from the WGP to the center of the CP. Here we focus our discussion on newly resolved upper mantle structure beneath the western CP and the GB-CP transition, as imaged using data from La RISTRA 1.5. 2. Methods and Model [4] Instrument responses were deconvolved from La RISTRA 1.5 P and S wave recordings. P waves were bandpassed from 0.2– 1.5 Hz and S waves from 0.03– 0.25 Hz. Traveltime residuals with respect to the IASPEI model [Kennett and Engdahl, 1991] were then measured by a cross-correlation technique for P and S waves indepen- L10303 1 of 5 L10303 SINE ET AL.: MANTLE STRUCTURE BENEATH THE COLORADO PLATEAU L10303 frequency kernel considerations. This would not change our conclusions. [5] The final Vp and Vs models from eastern New Mexico into the GB are shown in cross section in Figure 2d with areas of interest labeled A – E. Notable features are the very slow P and S wave speeds in the upper 200 km beneath the GB (A) and the very sharp transition in seismic wave speed at the western edge of the CP (B) where Vp and Vs change 6% and 12% respectively over a horizontal distance of less than 100 km. The eastern edge of the CP is also marked by a strong lateral gradient from fast velocity beneath the CP to slow velocity beneath the Jemez lineament (D) associated with the RGR [Gao et al., 2004] although velocities beneath the Rift are not as slow as beneath the GB. Interestingly, the core of the CP is underlain by shallow mantle velocities that are slower than its edges. Under the central CP the ‘‘4-corners anomaly’’ zone (in the region of the Navajo volcanic field) of slightly lower than average velocity exists from 200 to 500 km depth (C). At depths below 400 km there is no indication of a strong anomaly beneath the GB or the RGR. Figure 1. Locations of La RISTRA 1.0 and La RISTRA 1.5 stations. Topographic and geologic provinces outlined. dently. Only data with back azimuths within 20° from the trend of the seismic line were used to limit inclusion of offstrike structure. Two stations in the La RISTRA 1.5 array reoccupied La RISTRA 1.0 sites and were used to combine the residuals from both deployments into a single dataset. Residuals were corrected for topography and crustal thickness variations using the models of Wilson et al. [2005] and Wilson et al. [in preparation, 2008] so that final residuals reflect mantle velocity variations. La RISTRA 1.5 produced 767 P wave residuals and 559 S wave residuals. These data were added to the identically processed 5007 P wave and 2164 S wave residuals from La RISTRA 1.0 [Gao et al., 2004]. The mantle beneath the seismic line was parameterized into 1464 blocks 25 km wide and 25 km deep, and the raypath corresponding to each time residual was backprojected through the mantle assuming the IASPEI model. The traveltime residuals were inverted using the LSQR algorithm [Paige and Saunders, 1982] to give a regularized least squares solution for slowness perturbations within the blocks relative to the IASPEI model following the approach of Gao et al. [2004]. To examine model resolution, synthetic traveltime residuals were computed using the raypaths of the measured data for a checkerboard model with 50 by 50 km blocks of alternating velocity (Figure 2a) and were inverted using identical methodology. Figures 2b and 2c show the recovered S and P waves models for this resolution test. The centers and shallower depths of the recovered models are well resolved, but the model edges, lacking crossing raypaths, are significantly less well resolved. Note that most blocks show recovery of the sign of the checkerboard anomalies, but that spreading of information between adjacent blocks generally reduces recovered anomaly amplitude relative to the test model. We would expect the sharpness and the magnitude of the anomalies to increase with 3-D ray tracing or finite Figure 2. Input checkerboard model (a). S and P wave models recovered using La RISTRA raypaths (b) and (c). Final P wave (top) and S wave (bottom) velocity models, topography, regional provinces, longitude, and every fifth station (d). 2 of 5 L10303 SINE ET AL.: MANTLE STRUCTURE BENEATH THE COLORADO PLATEAU Finally, the dominant deep anomaly is a zone of fast mantle from 300 to 600 km beneath the eastern GB (E). 3. Discussion [6] The sharpness and magnitude of velocity change beneath the GB-CP transition indicate an abrupt change in mantle properties similar to what Zandt et al. [1995] found across the GB-CP boundary south of our line. Assuming that the velocity contrast is due to a sharp lateral temperature gradient and using the temperature derivatives given by Cammarano et al. [2003] (assuming a Qs from 80 to 140 over depths from 150 to 250 km), our model implies an average temperature contrast across the western edge of the CP of 600°C averaged over the upper 200 km. However, temperature estimates from xenoliths [Riter and Smith, 1996] and gravity modeling [Wilson et al., in preparation, 2008] suggest a temperature change of only 200– 400°C. Mantle compositional differences across the western CP boundary are likely insufficient to contribute substantially to the observed anomaly [Schutt and Lesher, 2006]. Hydrous minerals slow seismic velocities significantly but are stable at relatively low upper mantle temperatures (e.g. hornblende is stable up to 1000°C). Hacker et al. [2003] find that typical upper mantle depleted harzburgite contains stable hydrous phases only at temperatures less that 800°C. Smith [2000] found temperatures in the mantle 45 km beneath the GB to be >1000°C, thus ruling out hydrous mineral induced velocity reduction as the primary cause of the deeper slow velocities. [7] Supersolidus conditions in the mantle result in melt between mantle grains which reduces the elastic moduli of the mantle and thereby the seismic velocity. Mavko [1980] and Hammond and Humphreys [2000] model realistic crystal-melt configurations and predict a large range in velocity reduction with a Vp reduction of 1 to 3.6% and a Vs reduction of 2 to 7.9% respectively for a 1% melt fraction indicating that melt in the mantle beneath the GB could contribute to the observed velocity contrast. Hydrogen in nominally anhydrous mantle minerals can also greatly reduce seismic velocities. Starting from a typical mantle (0.01 wt.% hydrogen and Q = 100), the average S wave velocity reduction is 8% for a 0.1 wt.% water content and 19% for a 1 wt.% water content [Karato, 2006]. Hydration of nominally anhydrous mantle minerals localized beneath the GB could thus also partly explain the observed velocity contrast across the GB-CP boundary. Magnetotelluric imaging of this region [Wannamaker et al., 2001] also shows very low mantle resistivity beneath the GB margin at depths below 100 km. [8] We conclude that elevated temperatures combined with partial melting and/or high levels of hydrogen in nominally anhydrous minerals cause the extremely slow mantle velocities beneath the GB. Wyllie [1979] estimates a peridotite solidus drop of 500°C at a depth 100 km for an addition of 0.4% water. Considering the inferred history of flat-slab subduction in the western U.S. hydration of the shallow mantle beneath the GB would not be surprising. After slab rollback and asthenospheric upwelling [Humphreys et al., 2003], water-assisted melting of fertile lithosphere beneath southwestern North America appears L10303 inevitable. We propose the retention of 1 to 2% melt and a temperature increase of 300 to 500°C beneath the GB produces the strong gradient in seismic velocity across the GB-CP boundary. [9] The contrast in seismic velocity across this boundary is not only large but is also very sharp, and our limited resolution inversion results will underestimate the true sharpness. Our smoothed images show the entire change in S wave velocity (12%) occurring over less than 100 km laterally. Some of the contrast could be due to thinning of the lithosphere with subsequent asthenospheric upwelling under the highly extended GB, but the contrast in velocity extends over almost 200 km in depth. Even with strains approaching 100% for the GB, thinning due to extension can not explain the entire seismic anomaly across the boundary. [10] O’Reilly et al. [2001] and Poudjom Djomani et al. [2001] (comparing compositions, ages, and strengths of 16,000 global mantle xenocrysts) found lithospheric buoyancy and refractoriness to increase with age. Therefore, under certain conditions Proterozoic lithosphere may resist disintegration while partial melting of fertile Paleozoic lithosphere leads to enhanced heat advection through melt migration and ultimately to pervasive lithospheric weakening. The response of juxtaposed Paleozoic and Proterozoic mantle lithosphere to heat and water input could therefore result in the present-day contrast beneath the GB-CP transition. We conclude that our image reveals the western extent of unadulterated Proterozoic mantle lithosphere which is maintaining the stability of the western edge of the CP and that the factors controlling lithospheric stability in this area were established by preexisting differences between the Proterozoic and Paleozoic mantle lithospheres that meet at our inferred boundary. While Precambrian crust, possibly underlain by a thinning sliver of Paleoproterozoic upper mantle lithosphere, clearly extends west of our boundary to the SR 0.706 line [Burchfiel et al., 1992], we believe the deeper mantle lithospheres separated by our inferred boundary are distinct. The narrow high velocity zone bounding the edge of the CP may also be a manifestation of edge driven convection as proposed by van Wijk et al. [in preparation, 2008]. [11] The fast anomaly imaged from 300 to 600 km beneath the NW portion of the array (E) has Vp and Vs anomalies at 400 km of up to 2% and 4% faster than the surrounding mantle respectively, implying a temperature anomaly of 400°C [Cammarano et al., 2003]. Though checkerboard tests indicate that the model amplitude and precise shape of this anomaly are not reliable, its existence as a strong and large (250 by 250 km) feature is not in doubt. Van der Lee and Nolet [1997] also find transition zone seismic anomalies in this region and interpret them as arising from the trailing edge of the Farallon plate. Schmid et al. [2002] models the present-day thermal signature of the Farallon plate and predicts that fragments remaining in the upper mantle retain a thermal anomaly of 200 – 400°C, consistent with anomaly (E) being the Farallon plate. The volume of the anomaly, however, is too large to be a simple descending oceanic plate of relatively young age. Internal plate deformation due to resistance at the endothermic 3 of 5 L10303 SINE ET AL.: MANTLE STRUCTURE BENEATH THE COLORADO PLATEAU L10303 Figure 3. Inferred mantle state beneath La RISTRA. 660 km phase change [Schmid et al., 2002] could distort the slab to produce the volume of material necessary to give the observed anomaly. [12] Humphreys [1995] observes two migrating areas of magmatism in western North America converging to southern Nevada around 20 Ma and correlates these fronts with the edges of the foundering Farallon plate. In this model the slab buckles in a concave downwards fashion forming a locus of downwelling plate. This area of convergence is spatially associated with the large deep anomaly imaged in this study. Continental mantle lithosphere delaminated from beneath the GB could also contribute to the deep anomaly. Paleozoic mantle lithosphere is gravitationally unstable under normal mantle temperatures and geotherms [O’Reilly et al., 2001; Poudjom Djomani et al., 2001]. Therefore, gravitational instability enhanced by lithospheric cooling during flat-slab subduction might have resulted in delamination of GB mantle lithosphere as lowdensity upwelling asthenosphere replaced the foundering Farallon plate. [13] Figure 3 shows the present-day mantle state interpreted from our tomograms. The sharp velocity contrast across the GB-CP transition corresponds to a boundary between altered GB Paleozoic mantle lithosphere, or possibly juvenile mantle that replaces older delaminated lithosphere, juxtaposed with largely unaltered CP Proterozoic mantle lithosphere. A thinning sliver of Precambrian upper mantle lithosphere may extend west of our inferred boundary. Beneath the eastern GB the Farallon plate or delaminated lithosphere is sinking in the transition zone. Counter-flow associated with the sinking fast anomaly may be coming up beneath the CP and perhaps feeding hot mantle to the GB to the west and the RGR to the east. [14] Acknowledgments. This research was supported by NSF grants EAR-9707188, EAR-9706094, EAR-9707190, and EAR-0207812; by the Los Alamos National Laboratory Institute of Geophysics and Planetary Physics; and by the Geology Foundation of the Jackson School of Geosciences at the University of Texas at Austin. 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Reynolds (1977), Cordilleran benioff zones, Nature, 270, 403 – 406. Gao, W., S. Grand, W. S. Baldridge, D. Wilson, M. West, J. Ni, and R. Aster (2004), Upper mantle convection beneath the central Rio Grande rift imaged by P and S wave tomography, J. Geophys. Res., 109, B03305, doi:10.1029/2003JB002743. Hacker, B., S. Peacock, G. Abers, and S. Holloway (2003), Subduction factory 2. Are intermediate-depth earthquakes in subducting slabs linked to metamorphic dehydration reactions?, J. Geophys. Res., 108(B1), 2030, doi:10.1029/2001JB001129. Hammond, W., and E. Humphreys (2000), Upper mantle seismic wave velocity: Effects of realistic partial melt geometries, J. Geophys. Res., 105, 10,975 – 10,986. Humphreys, E. (1995), Post-laramide removal of the Farallon slab, western United States, Geology, 23, 987 – 990. Humphreys, E., E. Hessler, K. Dueker, G. Farmer, E. Erslev, and T. Atwater (2003), How laramide-age hydration of North American lithosphere by the Farallon slab controlled subsequent activity in the western United States, Int. Geol. Rev., 45, 575 – 595. Karato, S. (2006), Remote sensing of hydrogen in earth’s mantle, in Water in Nominally Anhydrous Minerals, Reviews in Mineralogy and Geochemistry, vol. 62, edited by H. Kepler et al., pp. 343 – 375, Mineral. Soc. Am., Chantilly, Va. Kennett, B., and E. Engdahl (1991), Traveltimes for global earthquake location and phase identification, Geophys. J. Int., 105, 429 – 465. Mavko, G. M. (1980), Velocity and attenuation in partially molten rocks, J. Geophys. Res., 85, 5173 – 5189. Morgan, P., and C. Swanberg (1985), On the Cenozoic uplift and tectonic stability of the Colorado Plateau, J. Geodyn., 3, 39 – 63. O’Reilly, S., W. Griffin, Y. Poudjom Djomani, and P. Morgan (2001), Are lithospheres forever? Tracking changes in subcontinental lithospheric mantle through time, GSA Today, 11, 4 – 10. 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(2000), Insights into the evolution of the uppermost continental mantle from xenolith localities on and near the Colorado Plateau and regional comparisons, J. Geophys. Res., 105, 16,769 – 18,920. Van der Lee, S., and G. Nolet (1997), Seismic image of the subducted trailing fragments of the Farallon plate, Nature, 386, 266 – 269. Wannamaker, P., et al. (2001), Great Basin-Colorado Plateau transition in central Utah: An interface between active extension and stable interior, in The Geologic Transition, High Plateaus to Great Basin: A symposium and field guide: The Mackin Volume, Utah Geol. Assoc. Publ., vol. 30, edited by M. C. Erskine, pp. 1 – 38, Utah Geol. Assoc., Salt Lake City, Utah. Wernicke, B. P., G. J. Axen, and J. Snow (1988), Basin and Range extensional tectonics at the latitude of Las Vegas, Nevada, Geol. Soc. Am. Bull., 100, 1738, 1738 – 1757. West, M., J. Ni, W. Baldridge, D. Wilson, R. Aster, W. Gao, and S. Grand (2004), Crust and upper mantle shear-wave structure of the southwest United States: Implications for rifting and support for high elevation, J. Geophys. Res., 109, B03309, doi:10.1029/2003JB002575. Wilson, D., R. Aster, M. West, J. Ni, S. Grand, W. Gao, W. Baldridge, S. Semken, and P. Patel (2005), Lithospheric structure of the Rio Grande rift, Nature, 433, 851 – 855. L10303 Wyllie, P. (1979), Magmas and volatile components, Am. Mineral., 64, 469 – 500. Zandt, G., S. Myers, and T. Wallace (1995), Crust and mantle structure across the basin and range - Colorado Plateau boundary at 37°N latitude and implications for Cenozoic extensional mechanism, J. Geophys. Res., 100, 10,529 – 10,548. R. Aster, Department of Earth and Environmental Science, New Mexico Institute of Mining and Technology, Socorro, NM 87801, USA. ([email protected]) W. S. Baldridge, Earth and Environmental Sciences Division, Los Alamos National Laboratory, Los Alamos, NM 87545, USA. (sbaldridge@ lanl.gov) W. Gao, TGS NOPEC Geophysical Company, 2500 Citywest Blvd. Suite 2000, Houston, TX 77042, USA. ([email protected]) S. Grand, Jackson School of Geosciences, University of Texas, University Station C1100, Austin, TX 78712, USA. (steveg@maestro. geo.utexas.edu) J. Ni, Department of Physics, New Mexico State University, Las Cruces, NM 88003, USA. ([email protected]) C. R. Sine, Occidental Petroleum Corporation of Elk Hills, P.O. Box 1001, Tupman, CA 93276, USA. ([email protected]) D. Wilson, U.S. Geological Survey Hawaiian Volcano Observatory, Box 51, Hawaii National Park, HI 96718, USA. ([email protected]) 5 of 5