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GEOPHYSICAL RESEARCH LETTERS, VOL. 35, L10303, doi:10.1029/2008GL033391, 2008
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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-
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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).
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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
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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
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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. Instruments were provided by
the PASSCAL facility of the Incorporated Research Institutions for Seismology (IRIS) through the PASSCAL Instrument Center at New Mexico
Tech. IRIS facilities are supported by cooperative agreement NSF EAR-
000430 and the Department of Energy National Nuclear Security
Administration.
References
Burchfiel, B. C., D. S. Cowan, and G. A. Davis (1992), Tectonic overview
of the Cordilleran orogen in the western United States, in The Cordillerian Orogen: Conterminous U.S., Geology of North America, vol. G-3,
edited by B. C. Burchfiel et al., pp. 422 – 431, Geol. Soc. Am., Boulder,
Colo.
Cammarano, F., S. Goes, P. Vacher, and D. Giardini (2003), Inferring uppermantle temperatures from seismic velocities, Phys. Earth Planet. Inter.,
138, 197 – 222.
Chapin, C., and S. Cather (1994), Tectonic setting of the axial basins of the
northern and central Rio Grande rift, Spec. Pap. Geol. Soc. Am., 291, 5 – 24.
Coney, P., and S. 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.
Paige, C., and M. Saunders (1982), LSQR, An algorithm for sparse linear
equations and sparse least squares, ACM Trans. Math. Software, 8, 43 – 71.
Poudjom Djomani, Y., S. O’Reilly, W. L. Griffin, and P. Morgan (2001),
The density structure of subcontinental lithosphere therough time, Earth
Planet. Sci. Lett., 184, 605 – 621.
4 of 5
L10303
SINE ET AL.: MANTLE STRUCTURE BENEATH THE COLORADO PLATEAU
Riter, J., and D. Smith (1996), Xenolith constraints on the thermal history
of the mantle below the Colorado Plateau, Geology, 24, 267 – 270.
Schmid, C., S. Goes, S. Van der Lee, and D. Giardini (2002), Fate of the
Cenozoic Farallon slab from a comparison of kinematic thermal modeling
with tomographic images, Earth Planet. Sci. Lett., 204, 17 – 32.
Schutt, D., and C. Lesher (2006), Effects of melt depletion on the density
and seismic velocity of garnet and spinel lherzolite, J. Geophys. Res., 111,
B05401, doi:10.1029/2003JB002950.
Smith, D. (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])
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