RICE UNIVERSITY

Imaging Lithospheric Structure Beneath the Colorado Plateau and its

Adjacent Regions Using Rayleigh Wave Tomography

by

Kaijian Liu

A THESIS SUBMITTED IN PARTIAL FULFILLMENT OF THE REQUIREMENTS FOR THE DEGREE

Master of Science

APPROVED, THESIS COMMITTEE:

Alan Levander, Carey Croneis Professor, Chair Earth Science

Fenglin Niu, Associate Professor Earth Science

David Alexander; Professor Physics and/Astronomy

HOUSTON, TEXAS

APRIL 2010

UMI Number: 1486042

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ABSTRACT

Imaging Lithospheric Structure Beneath the Colorado Plateau and its Adjacent Regions using Rayleigh Wave Tomography

by Kaijian Liu

This thesis presents a new 3-D shear velocity (Vs) model of the lithospheric structure

beneath the Colorado Plateau and its surrounding regions to understand the complicated

lithospheric modifications caused by major Cenozoic tectonic and magmatic activities.

Prior to this work, lack of an overall velocity model had prevented a comprehensive

understanding of the current encroachment patterns near the margins, and of the

perplexing crust/mantle boundary beneath the plateau. Using the new USArray data, I

have inverted the isotropic Vs model from the Rayleigh wave dispersion curves obtained

by a modified two-plane wave method. The resulting Vs structures not only clearly image

the sublithospheric low-velocity channels, but also identify the high velocity anomalies

underlying the western plateau for the first time in surface wave tomography to support

the recent delamination hypothesis. This work also provides a high-resolution velocity

model for geodynamic modeling on lithosphere-asthenosphere interactions.

11

ACKNOWLEDGEMENTS

I am thankful and indebted to many people who offered me help and encouragement

during my thesis work. Much is owed to my advisor, Dr. Alan Levander, for his great

supervision, valuable suggestion, continuous support, helpful discussion and patient

thesis editing. I have benefited greatly from his profound knowledge of the western

United States and his enthusiasm in Earth science.

I would also like to thank Dr. Fenglin Niu for his invaluable help on the surface wave

tomography code, the data processing and many suggestions. I am grateful to Dr. David

Alexander for his constructive and insightful comments on my thesis.

I would like to thank the talented Rice seismology group for their suggestions and special

thanks go to Yongbo Zhai, Sally Thurner and Jennifer Mackenzie for help on my thesis

writing. Aimin Cao, Maximiliano Bezada, Yang He and Xin Cheng also gave useful

advice on my research. Particularly, I would acknowledge former post-doc Meghan

Miller for her collaboration on the Colorado Plateau project, and express my gratitude to

Dr. Yingjie Yang, Dr. Donald Forsyth and Dr. Aibing Li for their assistance on the

surface wave inversion code. I thank Tobias Hoink and Adam Ringler for discussion on

the results. Dr. Jan Hewitt, Dr. William Symes and Dr. Matthias Heinkenschloss in the

thesis writing class (CAAM 600) also offered plenty of great suggestions and comments.

Finally, I would thank my parents and brothers for their love and unconditional support

all the time. This research was supported by NSF EarthScope grant EAR-0844760.

i i i

TABLE OF CONTENTS

Abstract ii

Acknowledgements iii

Table of Contents iv

Chapter 1 Introduction 1

1.1 Geologic background 1

1.2 Previous seismological investigations 5

1.3 Motivation 6

1.4 Rayleigh wave tomography 9

Chapter 2 Data and Methodology 10

2.1 USArray stations and data processing 10

2.2 Methodology 15

2.2.1 Modified two-plane wave tomography 15

2.2.2 Shear velocity inversion 19

Chapter 3 Parameterization and Inversion 21

3.1 Phase velocity measurement 21

3.1.1 Parameterization 21

3.1.2 Inversion 23

3.1.3 Phase velocity maps 27

3.2 Shear velocity inversion 29

3.2.1 Parameterization and inversion 29

3.2.2 Average shear velocity structures 33

Chapter 4 Results and Discussion 36

iv

4.1 Crustal structure 36

4.1.1 Crustal heterogeneity 36

4.1.2 Interpretation 39

4.2 Upper mantle structure 42

4.2.1 Upper mantle seismic heterogeneity 43

4.2.2 Colorado Plateau mantle lithosphere 46

4.2.3 Basin and Range and CP-BR transition zones 48

4.2.4 Rio Grande Rift and Jemez lineament volcanism 49

4.2.5 Southern Rocky Mountains 50

4.3 Sublithospheric structure 50

4.4 Possible delamination 55

4.4.1 Comparison with receiver function imaging 55

4.4.2 Possible delamination mechanism 59

Chapter 5 Conclusions 63

List of References 65

v

CHAPTER 1

INTRODUCTION

This paper constructs a high-resolution shear velocity model of the lithospheric structure

beneath the Colorado Plateau and its surrounding regions. Such a seismic model is

essential for understanding how the major tectonic and magmatic activities altered the

continental lithosphere during the Cenozoic era, and how the plateau was uplifted to its

current high elevation while suffering little internal deformation. Using the dense

USArray data, I applied a modified Rayleigh wave method to build a better-resolved Vs

model to further examine the Cenozoic lithospheric modification and current complexity

of the crust/mantle boundary (Moho) beneath the plateau and to provide a high-resolution

velocity model for future geodynamic studies.

1.1 Geologic background

The "bowl-shaped" Colorado Plateau (CP) is a high standing and relatively stable block

across four states (Utah, Colorado, New Mexico and Arizona) in the southwestern United

States. As shown in Figure 1.1, this distinct province is bounded by the Uinta Mountains

and the Wasatch front in Utah in the north, the highly deformed Southern Rocky

Mountains (SRM) in the northeast, the Rio Grande Rift (RGR) valley in the east, and the

highly extended Basin and Range Province (BRP) in the south, west and southwest.

Between the Southern Basin and Range (SBR) Province and the plateau, there is a broad

transition zone of about 100 km with transitional geologic and geophysical

characteristics. The Colorado Mineral Belt (COMB) and the Jemez lineament (JL) are

1

two distinct geologic features lying at the transition regions between the CP and the SRM

and the RGR, respectively (Figure 1.1).

-118 -116" -114" -112" -110' -108" -106'' -104*

-118' -116" -114' -112" -110' -108 -106" -104'

Figure 1.1 Topography map of the major physiographic provinces of the western United States in this study, including the Colorado Plateau, the Southern and Northern Basin and Range, the Rio Grande Rift, the Southern Rocky Mountains and the Great Plains. Station distribution inside the study area (white box) from various USArray networks (TA: Transportable Array; CI: Caltech Regional Seismic; US: US National Seismic; IU: IRIS/USGS; UU: University of Utah Regional) are also shown in different symbols (see legend). The topography data is from the ETOPOl model (http://www.ngdc.noaa.gov/mgg/global/global.html; Amante and Eakins, 2009). Locations of the Colorado Mineral Belt (COMB) in Colorado and the NE trending Jemez lineament (JL) at the southeastern margin of the plateau are also indicated.

2

The modern lithospheric structures beneath the Colorado Plateau and its adjacent

provinces in the tectonically active southwestern United States are complicated mainly

due to significant Cenozoic modifications to a Paleozoic passive continental margin

outboard of a Proterozoic and Archean core (e.g. Lipman, 1992; Humphreys et al, 2003).

The northeast-striking continental lithosphere was assembled through progressive

southward accretion of island arc terranes during several Proterozoic tectonism, including

the Yavapai and Mazatzal orogeny (1.75-1.60 Ga), followed by anorogenic intrusion of

granitoids (1.48-1.32 Ga) and mafic intrusion (1.2-1.1 Ga) in the southwestern U.S.

(Karlstrom and Bo wring, 1988; Bowring and Karlstrom, 1990; Karlstrom and

Humphreys, 1998). The major lithospheric modification from the Paleozoic passive

North American continental margin to the cratonic edge (eastern limit of the North

American Cordillera) was associated with the major tectonism, magmatism and

deformation (shortening, extension etc.) occurring within the western U.S. as a

consequence of the Late Cretaceous-Early Cenozoic Farallon plate subduction beneath

the North American plate (e.g. Coney and Reynold, 1977, Humphreys et al, 2003). The

initially rapid subduction of the Farallon slab caused the basement-cored uplifts due to

crustal compression and thickening during the Laramide orogeny (~ 70-45 Ma), and also

initiated an eastward migration of the arc magmatism (Coney and Reynold, 1977;

Lipman et al, 1972; Snyder et al, 1976) in the foreland from Sierra Nevada (~ 74 Ma) to

as far as the NE trending Colorado Mineral Belt (Lipman, 1992; Miller et al, 1992)

several hundred kilometers inland from the trench. Due to the combined effects of the

oceanic plateau subduction and enhanced mantle wedge suction caused by rapid sinking,

the Farallon slab began to flatten and subsequently hydrate the base of the North

3

American mantle lithosphere (Humphreys et ah, 2003; English et ah, 2003; Humphreys,

2009) by introducing water content to the nominally anhydrous minerals and depressing

the solidus for subsequent melting (Lee, 2005; Li et ah, 2008).

Following the slab removal around 30 Ma, the Mid-Tertiary Basin and Range volcanism

("ignimbrite flare-up") and intensive extension were triggered by the combination of

post-subduction orogenic collapse and exposure of the fertile mantle lithosphere to the

hot ascending asthenosphere (Dickinson and Snyder, 1978; Humphreys, 1995). Small-

scale convection at scattered regions (Wilson et ah, 2005a, b; West et ah, 2004;

Hyndman et ah, 2005; Humphreys, 2009) and mantle asthenospheric upwelling from

deep source {e.g. Morgan et ah, 1986; Moucha et ah, 2008, 2009) further modified the

unstable southwestern U.S. lithosphere after the foundering of the slab. Magmatism and

deformation continue presently in the BRP and RGR provinces and are migrating onto

the CP margins at encroachment rates of ~ 3-6 km/Myr (Roy et ah, 2009). During the

past 30 Myr, the RGR at the southeastern edge of the plateau has undergone a two-stage

rifting process characterized by low- and high-angle normal faulting, consecutively, and

is separated by a ~ 10 Myr period of magmatic quiescence (West et ah, 2004; Gao et ah,

2004; Wilson et ah, 2005a, b). It still remains controversial in what manner these tectonic

and magmatic activities (such as plate subduction, slab flattening, post-Laramide

foundering, post-orogenic collapse, rock uplifts and incision) have altered the lithosphere

structure of the North America, especially beneath the CP and the transition zones (TZs)

between the plateau and the surrounding BRP, the RGR and the SRM (Humphreys, 1995;

Humphreys, 2009).

4

1.2 Previous seismological investigations

There have been numerous seismic studies on crustal and upper mantle structures

underlying the CP and the surrounding regions; however, they are limited to either large

continental scales with poorly sampled data, or focus on small regions or profiles. At

large scale, for example, P and S body wave traveltime tomography (e.g. Humphreys and

Dueker, 1994a, b; Burdick et al, 2008) and surface wave tomography (e.g. van der Lee

and No let, 1997) both characterized the western U.S. by pronounced low P and S velocity

anomalies in the upper mantle, from the Mendocino Triple Junction (MTJ) to the Rocky

Mountains front far in the east where the cratonic interior begins. Unfortunately, due to

uneven and sparse station coverage, these results provide little information of localized

heterogeneities on the plateau.

On the other hand, the regional studies in the past two decades investigated small scales,

though with higher lateral or vertical resolution to various degrees. For example, the

seismic reflection/refraction study from the Pacific to Arizona Crustal Experiment

(PACE) (e.g. Wolf and Cipar, 1993; Benz and McCarthy, 1994) and the Colorado

Plateau-Great Basin (CP-GB) PASSCAL experiment (e.g. Sheehan et al, 1995, 1997)

focus on the southern and northern Basin and Range-Colorado Plateau transition zones,

respectively. There are also four major seismic lines, one across the eastern Basin and

Range-western Colorado Plateau transition at 37°N by the University of Arizona (Zandt

et al, 1995), one along the N-S trending Deep Probe using active sources (Henstock et

al, 1998; Snelson et al, 1998; Gorman et al, 2002), one along the Continental Dynamics

of the Rocky Mountains line (CD-ROM) (e.g. Levander et al, 2005, Snelson et al,

5

2005), and the fourth and more recent line across the Rio Grande rifting zone between

western Texas to southeastern Utah by the Colorado Plateau/Rio Grande Rift Seismic

Transect Experiment (La RISTRA) group (West et al, 2004; Gao et al, 2004; Wilson et

al, 2005a, b). Although these active and passive seismic investigations on province

scales reveal parts of the complicated lithospheric architecture, many of which are

discordant with the tectonic features, a comprehensive understanding on how the plateau

interacts with the surrounding provinces requires an overall seismic velocity model that

will be provided in this study on this region.

1.3 Motivation

Such a full-scale crustal and upper mantle model can also be of general importance (1) to

understand the CP uplift mechanism, and (2) to reexamine the existing explanations of

the split and diminished signals of the converted phases in the receiver function studies.

First, a detailed 3D lithospheric model can provide an understanding of the timing and

uplift mechanism(s) of the CP, especially the post-Laramide uplift, which have been the

subject of long-standing debates. The present elevation of approximately 2 km was

mainly formed during the Cenozoic, since the plateau appears to have stayed at or close

to sea level until ~ 65 Ma, as evidenced by the widely distributed Late Cretaceous age

marine strata throughout the plateau {e.g. Oetking et al, 1967; Morgan and Swanberg,

1985). Numerous mechanisms have been suggested to explain the 2-km Cenozoic

rock/surface uplift of the plateau, which is roughly in Airy iso static equilibrium

(Thompson and Zoback, 1979). The buoyancy sources from the crust, lithosphere mantle

6

and asthenosphere that support the present elevation can be attributed roughly to four

major factors: thermal expansion (associated with phase change), mechanical lithospheric

thinning, crustal thickening and dynamic uplift. However, no consensus on the amount of

uplift or the timing contributed by each factor has been achieved yet. For example, based

on an early idea by Christiansen and Lipman (1972), Tompson and Zoback (1979)

suggested that the plateau uplift was caused mainly by gradual thermal warming

associated with eclogite to gabbros phase conversion, after cessation of shallow

subduction. Roy et al. (2009) further developed and quantified this idea based on a

simplified 3-D thermal conduction model, in which they assumed a transient mid-Tertiary

thermal perturbation to the thick heterogeneous and Fe-depleted lithosphere of the

plateau that triggered thermal re-equilibration, expansion and uplift, and the following

magmatism. Crustal thickening, as another significant uplift contributor, might have been

caused by Laramide-orogeny-related compression, magmatic intrusion or underplating

(Morgan and Swanberg, 1985; McKenzie, 1984). Another hypothesis is that intracrustal

transfer from an overthickened hinterland to the north (McQuarrie and Chase, 2000)

raised the CP while maintaining isostatic balance. Besides the thermal and crustal factors,

partial delamination of the mantle lithosphere has been suggested (Beghoul and

Barazangi, 1989; Spencer, 1996; Lastowka et al, 2001), as has the early complete

delamination hypothesis (Bird 1984, 1988). The latter does not appear consistent with

geophysical observations such as relatively high Pn and low heat flow beneath large parts

of the plateau (a thermal signal may not yet have reached the surface) could also produce

much of the plateau uplift while preserving petrologic and seismic signatures in the

uppermost mantle. In addition to all the passive sources to support the high elevation,

7

Moucha et al. (2008, 2009) suggested a regional dynamic topography (~ 500 m)

supported by a strong mantle flow coupled to the sinking Farallon slab. The uplift

contributions from various factors depend on two crucial depths: the Moho and the

lithosphere-asthenosphere boundary (LAB), which can be inferred from the Vs structures.

Second, the complexity and weak signals of converted phases in determination of the

crustal thickness beneath most of the plateau have been noted in many teleseismic

receiver function (RF) studies. From the common conversion point (CCP) stacked RF

profiles using several portable PASSCAL experiments and U.S. National Seismic

Network (USNSN) data, Gilbert and Sheehan (2004) suggested that the weak P-to-Sv

conversion energy in the northern Colorado Plateau might be introduced by gradational

transition from lowermost crust to uppermost mantle with small impedance contrast,

possibly caused by magmatic underplating in the lower crust of the plateau in the vicinity

of the Grand Canyon (Wolf and Cipar, 1993). Most of the converted amplitudes beneath

the northern plateau become separated positive events with weaker amplitude compared

to those of the surrounding tectonic regions. This feature was also identified by other RF

studies at different regions on the plateau using various datasets, e.g. La RISTRA data

(Wilson et al., 2005a), a combination of Source Phenomenology Experiment (SPE) and

La RISTRA array in southern plateau (Gilbert et al., 2007), the CP-GB in northwestern

plateau (Sheehan et al., 1997), the 37°N profile (Zandt et al., 1995), PACE across the

Arizona transition zone (McCarthy and Parsons, 1994; Wolf and Cipar, 1993; Benz and

McCarthy, 1994; Parsons et al., 1996) and the Deep Probe profile (Henstock et al, 1998;

Snelson et al, 1998; Gorman et al, 2002; Levander et al, 2005). The PdS and SdP RF

8

volumes of the western U.S. made by Levander and Miller (2010) from the dense

USArray/Transportable Array (TA) stations reveal a similar splitting of Moho depth

beneath the plateau with improved resolution. Instead, they interpreted this splitting

feature as a detachment of the lower crust and uppermost mantle at the western plateau,

rather than the low impedance contrast as suggested by Gilbert and Sheehan (2004).

Parsons et al. (1996) also claimed that the crust of the southwestern plateau possesses the

same crustal composition and appears as an unextended BRP to its southwest. The

surface wave tomographic result in this thesis will show why the delamination is a

preferred interpretation rather than the gradational Moho transition explanation.

1.4 Rayleigh wave tomography

In this study, to better understand the buoyancy sources for the high-elevated plateau, and

to examine the delamination hypothesis, I built a new high-resolution Vs model based on

a two-step inversion technique. In the next chapter, the details of data processing on the

records by the dense USArray/TA stations (-70 km) and the key ideas of the modified

two-plane wave inversion (Forsyth and Li, 2005; Yang and Forsyth, 2006b) and the

isotropic shear velocity inversion will be summarized. Further details on the

parameterization for both phase and shear velocity inversions and the phase velocity

maps will given in Chapter 3 before going to deep discussions on the Vs structures for the

crustal and upper mantle in Chapter 4. The final inverted Vs model based on the high-

quality data reveals detailed lithospheric structures deep to ~ 250 km, and successfully

images high velocity anomalies where the detached portion possibly occurred, and

therefore supports the delamination hypothesis by Levander and Miller (2010).

9

CHAPTER 2

DATA AND METHODOLOGY

This chapter first gives a brief summary of the selection and processing of the USArray

data to prepare the amplitude and phase data for the final 2-D phase velocity inversion.

Then I will describe the key idea of the two-step inversion method used in this study for

the final shear velocity model construction, especially the modified two-plane-wave

representation to improve the lateral resolution in the 2-D phase velocity inversion.

2.1 USArray stations and data processing

I used the phase and amplitude information of the fundamental mode Rayleigh

wavefields recorded by the EarthScope/USArray Transportable Array (TA) network and

several other seismic networks in the southwestern U.S. within the geographic boundaries

of 32° to 50° north latitude, and -116° to -105° west longitude. About 200 broadband

seismic stations in the TA network with an approximate station spacing of 70 km, and 18

stations from other networks were incorporated for analysis, including ten stations from

the Caltech Regional Seismic Network (CI), five stations from the US National Seismic

Network (US), two stations from the IRIS/USGS Network (IU), and one station from the

University of Utah Regional Network (UU), as shown in Figure 1.1.1 have used a total of

154 teleseismic earthquakes (30° < A s 120°, Figure 2.1a) with shallow focus < 70 km

and magnitude s 5.5 which occurred between June 2007 and September 2009, as the TA

network swept over the CP region. The good azimuthal distribution of events and well-

distributed stations provided extraordinarily dense raypath coverage (e.g. Figure 2.1b) to

10

(a)

m

112u -110"

Figure 2.1 (a) Azimuthal distribution of earthquakes (red dots) used as the Rayleigh wave sources centered on the Colorado Plateau (blue triangle). The neighboring concentric circles are in 30° distance, (b) Raypath (green lines) coverage (50 s) for the Colorado Plateau-Northern Basin and Range sub-region within the white box. Blue triangles represent the USArray stations.

11

better resolve the lateral heterogeneity of the phase velocities propagating across the

plateau region and its adjacent provinces.

I chose and processed the vertical components of the seismograms, which are free of

Love wave contamination in an isotropic medium, to extract the pure Rayleigh wave

signals. After resampling and correcting instrument responses to the same reference

transfer function, I applied 18 Butterworth bandpass filters (narrow bandpass width of 10

mHz, zero-phase shift and 4th order) with central periods at 20, 22, 25, 27, 30, 34, 40, 45,

50, 59, 67, 77, 87, 100, 111, 125, 143 and 167 s for the pre-selected traces (see example

in Figure 2.2). Filtered seismograms with unusual amplitudes or relatively low signal-to-

noise ratios (e.g. < 3) were excluded, and only those clean traces with high coherence

among stations ordered by epicentral distances were retained for further analysis (Figure

2.3). To isolate the fundamental mode from higher order modes and body wave phases, I

windowed the surface waves with variable-length windows with 50s half cosine tapers at

both ends (/. e. windows were several hundred seconds in length and were identical for all

receivers for the same earthquake). Fewer seismograms in short periods (e.g. less than 30

s) can be kept because of severe distortion and incoherent interference patterns among

stations due to strong, probably localized multipathing, focusing, or defocusing effects.

The effects of frequency-dependent anelastic attenuation and geometrical spreading

(Mitchell, 1995; Li et al, 2003) on the Earth's surface were taken into account to provide

a minor correction of the amplitudes after measuring the phases and amplitudes by

Fourier transform. Local site responses (Yang and Forsyth, 2006a) were also considered

to account for individual station corrections of the amplitude amplification and phase

12

20 s

. ' 22 s

A 25 s

f • """"

*l 27 s "I • tt 30s •i

j — —

V / p * l ^ ^ ^

•" • 'Hl i l l ' "

. .* i A i» l» ^ * -

• • j j i ' - ' I •

• ' , « * ;

.«• 34s

-r = • J 40 s tv

t .» 45 s

.» 50 s H '4 ' • f 59 s

67 s

i 77 s »

f 87 s

• - - • • • - - - • _ • • * v ? • - • - • • •

". *

A ' • > " ,

" V • I

—--A •'"',•' • / V / " * 1 , > • * > , — i — • —

> . / \ • ; • '• .' '", / ' " - — | 100 s

j . ,.

\ » * 125s

r

* \ - x

I 143 s >

» 167s "i .,

10 20 x tC-2

... i . . .

22 23 24 2S

. • * >

... . ., i . i .*".„ . . V i ,. . . , i y . . . . . . , . . . . . . . . .J . ~ -26 2? 23 2>> 30 31 32 33 34 35 36 37 3&

..I. . 'A 39 «0

Time (s)

Figure 2.2 An example of fundamental-mode Rayleigh waveform at different frequency bands recorded at the M23A station in the USArray/TA network for an earthquake (Magnitude 6.4, depth 35 km, epicentral distance 98.3 degree) on February 17, 2009. The vertical component of the raw data is shown at top and was filtered into 18 frequency bands with central periods ranging from 20 s to 167 s. The delay of arrival time at longer periods (> 50 s) indicates a low velocity zone in upper mantle.

13

02/17/2009 (0.015-0.025 Hz) Depth = 35 km

X1E+01

220

Time (sec) 240 260 280 300 320

oo 00 = _

o

J-..-.-. 220

i 240

...J 260 280 300

L 320

Figure 2.3 Filtered waveform (50 s) at the USArray stations within the study region ordered by epicentral distance for the same individual earthquake used in Figure 2.2.

14

shift that correlates with the localized tectonic features. Earthquake magnitudes were

normalized by scaling the observed amplitudes to the R.M.S. amplitude of all the stations

for each event.

2.2 Methodology

2.2.1 Modified two-plane wave tomography

I have used a two-stage inversion approach developed by Yang and Forsyth (2006b) to

determine the Vs structure beneath the study region from the Rayleigh phase and

amplitude data. In the first step to estimate the 2-D phase velocities, I adopted the

modified two-plane-wave tomography (TPWT) technique (Yang and Forsyth, 2006a, b;

Forsyth and Li, 2005) to effectively take account of interference of energy resulting from

relatively simple multipathing and scattering effects in the incoming Rayleigh wavefields

that will be poorly represented by the traditional one-plane wave approximation. This

TPWT method has been applied successfully to a large amount of datasets to estimate the

phase velocity maps (Forsyth et al, 1998; Forsyth and Li, 2005; Li et al, 2003, 2005;

Yang and Forsyth, 2006a; Yang and Ritzwoller, 2008; Weeraratne et al, 2003; Schutt et

al, 2008; Miller et al, 2009).

In this method, the complexity and distortion in the fundamental mode surface waves in

each individual frequency band are simply considered as the sum of two interfering plane

waves with different amplitudes, initial phases and propagation directions (deviation

angles from the great circle path). The vertical component at the k-th station for the i-th

15

incident event at each frequency co can thus be written in the form of (Forsyth and Li,

2005)

2

]uz{x) = 2 , V<°>>exp(-,"''M0'>) t2-1) ipw=\

with

%PA^ + "(rf " T?) + [r*cos(tf- ,0^) - *f }kiSavgco (ipw = 1, 2)

in which jAipw(co), ,.0/>w, and °0. are the amplitude, derivation angle from the great

circle path, and the phase at the reference station for the primary and secondary plane

waves, x\ and T° are the traveltime of the k-th station and the reference station along the

great circle path with respect to the edge of the study area, respectively, *5 is the

average slowness as a function of the weighted phase velocity coefficients, and r\, y\,

x\ are the local coordinates with respect to the reference station, based on a regional

flattening coordinate system.

In addition, I also followed Yang and Forsyth's (2006b) consideration on the finite-

frequency effect of the surface wave propagation caused by the off-azimuth structures by

introducing amplitude and phase sensitivity kernels during the inversion for the phase

velocity maps at different frequency bands. This approach is effective to further improve

the lateral resolution on the local small-scale structures, especially the heterogeneities at

wavelength scale. Simplified from the 3-D surface wave kernels (Zhou et ah, 2004) based

on the single-scattering approximation (Born approximation), the 2-D phase and

amplitude sensitivity kernels K^(r,co) and KcA(r,a>) for the incoming plane wave

propagation at each single frequency, which are defined as

16

( ** \ rr(KcAr,co)^ dlnA i-Jj \K'A{r,m)j

— dx2, c

can be expressed as (Yang and Forsyth, 2006b):

K;(r,co) = M k2R" • exp[-i(kx" -kAx + JC/4)])

/T(r,a>) = -Re

R^lnkx"

k2R"-exp[-i(kx"-kAx + Jt/4)]h

R-\l2jrkx" J

(2.2)

(2.3)

in which 8c Ic, 6(f) and dIn A are the phase velocity perturbation, phase delay and

amplitude perturbation, respectively; k is the wavenumber of the incident Rayleigh

wave; R" and R are the station polarization vectors for scattered and incoming waves,

respectively; x" denotes the scatter-receiver distance, and Ax denotes the propagation

distance difference between the incoming waves that arrives at the receiver and scatterer.

Figure 2.4 shows the computed sensitivity kernels at 20, 50 and 100 s assuming a single

plane Rayleigh wave incident from left and was recorded in the center of the region. The

broadening and weakening of the sensitivity peaks are well noted at the bottom of Figure

2.4, as the period increases from 20 to 100 s. The introduction of the 2-D finite frequency

kernels were demonstrated to be able to improve the lateral resolution of phase velocity

in regional surface wave tomography by synthetic tests (Yang and Forsyth, 2006b).

There are two main steps involved to invert the phase velocity anomalies as well as the

wave parameters: (1) I inverted the average phase velocity and wave parameters; (2) I

performed a 2-D inversion to obtain the wave parameters and the velocity coefficients

after placing the updated mean phase velocity as the initial velocity at each grid node. In

both steps, I first used the simulated annealing technique to find the best fitting wave

17

-1600

-1000

-500

0

500

1000

1500

-1500

-1000

-500

0

500

1000

1500

-1500

-1000

-500

0

500

1000

1500

Amplitude Kernel (20 s} j

i _ _ 1

I 1

X10"

-1000 0 km

1000

Amplitude Kernel (50 s}

-1000 0 km

1000

Amplitude Kernel (100 s)

-1000 0 km

1000

x 10'

4

2

0

-2

-4

-6

E

-1S00

-1000

-500

0

500

1000

1500

Phase Kerftel (20 s]

210"

1

0 i

-1-1

-2

x10"

-10

-1500-1000 1000 1500 -10 -1500-1000 1000 1500

Figure 2.4 Map views of 2-D amplitude (left) and phase (right) sensitivity kernels on the Earth's surface for left-incident Rayleigh plane waves at 20, 50 and 100 s. The bottom panel shows the cross-section profiles at x=-500 km for these three periods, showing that the sensitivity peaks become broad and weak as the period increases.

18

parameters based on a fixed velocity model for each event individually, and then used a

generalized linearized inversion technique (Tarantola and Valette, 1982) to invert the

wave parameters and phase velocity coefficients, simultaneously. The solution to the

generalized linear inversion technique shows that the change of model Am to the starting

model mo can be given by (Forsyth and Li, 2005; Tarantola and Valette, 1982)

Am = (GTC-;nG + C"l r 1 [GTC-;n Ad - C l (m - m0)] (2.4)

in which m is the current model vector, Ad is the difference vector of the predicted and

observed data, G is the partial derivative matrix, Cmm and Cnn are the a priori

covariance matrices for model and data vectors, respectively. Under this inversion

scheme, the perturbation to the current velocity model and wavefield parameters will be

iteratively updated, until the numerical criterion for the misfit was met or the maximum

number of iterations was reached. The standard deviations of the phase residuals for each

event were estimated after each iteration and were divided to down-weight the noisy

events that were poorly fit by the two-plane wave approximation. In practice, I need no

more than 10 iterations to obtain the final phase velocity variations as no further

significant changes occurred.

2.2.2 Shear velocity inversion

In the second step, I inverted the regionalized shear velocity structures from the localized

dispersion curves (20-167 s) using the DISPER80 package (Saito, 1988) which was

further modified by Forsyth et al.. The forward model shows that the dispersion curves

can be computed through analogy with the mathematically equivalent solutions to the

spheroidal-mode free oscillations at the short wavelength limit, for a 1-D model

consisting of a number of transversely isotropic layers with homogeneous density and

19

elasticity (Takeuchi and Saito, 1972). The DISPER80 package takes into account the

sphericity of the Earth and neglects the self-gravitating effect which is much smaller at

periods > 20 s (Takeuchi and Saito, 1972). A fourth-order Runge-Kutta-Gill (RKG)

scheme was implemented to integrate the motion-stress differential equations upward to

the free surface to satisfy the boundary condition of vanishing stress. Meanwhile, a

compound-matrix method was also introduced to overcome the numerical instability in

searching for the root of the characteristic equation (Saito, 1988; Aki and Richards,

2002). Knowing the forward modeling and starting from a 1-D shear velocity model, I

used a similar iterative and linearized solution given by Equation 2.4 to the nonlinear

least-square regional 1-D shear velocity inversion to construct a pseudo 3-D crustal and

upper mantle shear velocity model beneath the study area. The crustal thickness

constraints are also placed from the PdS receiver function estimates (Levander and

Miller, 2010).

20

CHAPTER 3

PARAMETERIZATION AND INVERSION

Based on the modified two-plane wave tomography technique summarized in the last

chapter, I implemented this method separately to the four subdivisions in my study focus.

For all the eighteen frequency bands, the phase velocities within each subdivision will be

jointed together by averaging overlaps to construct overall phase velocity maps and to

obtain the dispersion curves on each grid node. The following regional 1-D shear velocity

inversions carried on each node lead to the final 3-D Vs model with remarkable vertical

resolution deep to 250 km beneath the CP and the adjacent regions.

3.1 Phase velocity measurement

3.1.1 Parameterization

The phase velocity lateral variations at each frequency can be highly resolved from the

input phase and amplitude data using the modified two-plane-wave approximation. First,

to satisfy the basic assumption of the two-plane wave approximation that the study region

should be small enough so that the recorded surface waveforms can be simply

represented by interference of two plane waves, I divided the entire area by the

perpendicular bisectors at the longitude and latitude axes into four sub-regions (Figure

1.1): (1) Colorado Plateau-Northern Basin and Range Province (CP-NBR); (2) Colorado

Plateau-Southern Rocky Mountains (CP-SRM); (3) Colorado Plateau-Southern Basin and

Range Province (CP-SBR); and (4) Colorado Plateau-Rio Grande Rift (CP-RGR). This

approach also happened to take good advantage of the dynamic deployment status of the

21

USArray/TA station coverage (the majority of the transportable stations lies on the left

side before January 2008, and very few stations remain in the Northern Basin and Range

after December 2008). The phase velocities in the overlapping regions were averaged

after application of the two-plane wave method in each subdivision separately. Different

events were chosen inside each 'box' to ensure good station coverage and to maintain

similar azimuthal distribution to resolve lateral phase velocity variations (Figure 3.1a, b).

Second, as mentioned in last chapter, I incorporated the 2-D surface wave phase and

amplitude sensitivity kernels, instead of the 'fat ray' approximation with Gaussian-

shaped sensitivity kernels (Forsyth and Li, 2005; Schutt et al, 2008) using same width of

influence zone along the great circle ray path. For all the frequency bands, a Gaussian

averaging function with a smoothing length Lw of 65 km was used in the traveltime

calculation from the edge to the receivers, and it also acted as a 2-D filter with the same

characteristic length to smooth the Born-approximation sensitivity kernels. A checkboard

resolution test for this modified two-plane wave technique has also been done using the

same event distribution and station coverage.

The phase velocity variations were inverted separately at each sub-region. Since the

vertical components of the seismogram are considered as the sum of two independent

interfering plane waves (each with unknown amplitude, initial phase and propagation

angles off the great circle path), there will be six plane wave parameters to represent each

source event. Each 'box' is parameterized on an evenly distributed grid with a spacing of

0.5° in the receiver-covered region, and double-spacing nodes on the edges to absorb the

22

traveltime residuals (by assigning the boundary nodes 10 times of the a priori standard

deviations of the inner nodes) for the waves that are beyond the simple two-plane-wave

representation. Thus, there will be 6M1 + M2 model parameters, where Mi is the total

number of events and M2 is the number of phase velocity coefficients at each grid node.

There are 2N Rayleigh wave observations in which N stands for the total number of

retained traces and the factor 2 comes from the real and imaginary components of the

complex amplitudes in the frequency domain. The starting average phase velocities were

estimated from the synthetic dispersion curve based on the 1-D TNA model (Grand and

Helmberger, 1984). The two main steps, as described in the method section in the last

chapter, were then implemented to invert the 6M1 wave parameters and M2 phase

velocity coefficients simultaneously using both the simulated annealing technique and the

standard linearized iterative inversion technique.

3.1.2 Inversion

The phase velocity variations from 20 to 167 s in each geological 'box' were derived

from the TPWT method described above. The number of earthquakes and raypaths

(Figure 3.1a, b) used in the Rayleigh wave tomography is much smaller at short periods

(< 25 s) due to severe distortion and strong multipathing in the surface waves, and also

reduces slightly due to the decreasing energy of signals at longer periods (> 100 s). The

mean phase RMS misfits (Figure 3. Id) for all of the events stay below 3 s, indicative of a

good fitting except for those at the longest periods {e.g. 143 and 167 s) that have larger

uncertainties of phase measurements. To examine the validity of the two-plane wave

approximation used in the inversion, I checked the distribution of the plane wave

23

150

^00

50

0

6000

6000

4000

3000

2000

1000

0

ICP-NBR

(a) Number of Earthquakes

' .."."..] CP-SRM i i CP-SBR I ICP-RGRI

1 ;• I 20 22 25 2? 30 34 40 46 60 69 67 77 87 100 11<! 125 143 167

(b) Number of Raypaths

I CP-NBR ~~~~~! CP-SRM i ^ CP-SBR I JP-RGR

20 22 25 27 30 34 40 45 50 59 67 77 87 100 11 ' 125 143 167

75

70

65

60

55

60

-

(c) Rank/(number of model parameters) (%)

f CP-NBR CP-SRM CP-SBR CP-RGR| — —|

•j

20 22 2b 27 30 34 40 45 50 59 67 77 87 100 11* 125 143 167

(d) RMS phase misfit (s)

I CP-NBR CP-SRM CP-SBR CP-RGRi i

20 22 25 27 30 34 40 45 50 59 67 77 87 100 114! 125 143 167

Period (s)

Figure 3.1 (a) Number of Rayleigh wave sources, (b) total number of raypaths, (c) the ratio of rank to the total number of model parameters, and (d) the phase RMS misfit in second for the phase velocity inversion at the 18 frequency bands at the four separate sub-regions: CP-NBR, CP-SRM, CP-SBR and CP-RGR.

24

42" 42"

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Phase

i;

0.8

0.8

0.4

0.0

-0.4

-0.8

US

.c 0.

"8 *** o T3 as

CL

Figure 3.2 Comparison of predicted and measured amplitude and phase data at 50 s in the CP-NBR sub-region. The map views of the amplitude (top) and phase (middle) show a good agreement between the observation and the prediction by the modified two-plane wave inversion technique, especially for the phase data. The bottom panel presents the same agreement in another way.

25

(a) Amplitude ratio

50 r

-°50

0.1 0.2 0.3 0.4 0.5 0.6 Amplitude ratio

(b) Primary wave

0.7 0.8 0.9

-15-12-9 - 6 - 3 0 3 6 9 12 15 Deviation angle from great circle path (degree)

(c) Secondary wave

-40 -30 -20 -10 0 10 20 30 40 Deviation angle from great circle path (degree)

Figure 3.3 Statistics of the inversion results of the two-plane wave approximation at 0.02 Hz (50 s). (a) Amplitude ratios of the smaller plane waves to the larger one. (b) Deviation angles of the primary plane waves from the great circle direction, (c) Deviation angles of the secondary plane waves from the great circle path. These angles are much larger than those of the primary waves.

26

amplitude ratio (secondary/primary amplitude), the range of deviation angles from the

great circle, and how well the predicted phase and amplitude fit with the observed data.

One example at 50 s of the inversion for one single earthquake (September 30, 2007) in

Figure 3.2 shows that the predicted phase and amplitude data fit pretty well with

observed data. Additionally, for most of events at the same period the amplitude ratio are

smaller than 0.4 (Figure 3.3a), and the deviation angles of the primary waves mainly fall

into ±10 degrees (Figure 3.3b, c), both of which suggest that the two-plane wave

representation is a reasonable approximation for the incoming wavefields at most of the

relevant frequency bands.

3.1.3 Phase velocity maps

The Rayleigh wave tomographic results of the phase velocity variations from 20 to 167 s

are plotted in Figure 3.4. From the phase velocity maps, several consistent velocity

anomalies associated with relevant geological features are identified: (1) a profound low

velocity area in the SRM at short periods (< 50 s) associated with a relatively low crustal

shear velocity beneath the Colorado Rockies; (2) at short periods less than 30 s, a large

low velocity area covering the COMB to the east of the CP, and relatively fast phase

velocities beneath the SBR which abruptly dropped once entering the CP-SBR transition

zone (TZ); (3) two long NE trending low velocity belts along the CP-NBR and the Jemez

lineament in the CP-RGR transition regions; (4) consistent relatively high phase

velocities in the CP interior compared to the surrounding tectonic margins; (5) when

periods are large than 45 s, the high velocity within the plateau is connected with the high

phase velocity anomaly to the south of the Wyoming Province, and this high velocity

27

34

32 K P L "t] iP9!3 L « i.2 125SJ ^ _1_67sJ 116 **4 112 MO 1CU *06 -116 -114 112 110 128 106 116 M 4 112 MO 108-106

6c/c(%) -6 0 2

Figure 3.4 Rayleigh wave phase velocity maps at 12 of the 18 periods. The phase velocity inversions were performed separately on the four 0.5° x 0.5° grid sub-regions, created by cutting along two geographic lines at 110.5°W and 37°N, and were then combined by averaging the overlaps.

28

anomaly feature exists consistently to longer periods, although it seems to be encroached

from the western, southwestern, southern and southeastern edges from the BRP and the

RGR; (6) the high anomaly beneath the SBR at short periods (e.g. 20, 25 and 30 s in

Figure 3.4) is associated with the relatively shallow Moho depths, since phase velocities

at the same period sense the deeper upper mantle velocity structures in the SBR while in

the other regions they mainly sense the crustal structures.

3.2 Shear velocity inversion

The regionalized isotropic crustal and upper mantle shear velocity structures beneath the

CP and the surrounding tectonic regions were inverted from the local Rayleigh wave

dispersion curves derived above from the two-plane-wave inversion at each grid node

(0.5° spacing). I performed the same standard iterative linearized inversion to solve the

nonlinear least-square problem by estimating the perturbation of the shear velocity

relative to the starting reference model. The Moho topography from teleseismic PdS

receiver functions (Levander and Miller, 2010) also put initial constraints on the Vs

inversion scheme to better resolve the lithospheric structure. A 3-D shear velocity model

was immediately constructed by unifying the grid-based 1-D Vs structures.

3.2.1 Parameterization and inversion

The Rayleigh wave phase velocities are primarily sensitive to the absolute shear velocity,

especially below the Moho, rather than the density and compressional velocity. The

sensitivity kernel profiles for S, P wave velocities and density at various periods (Figure

29

(a) Vs kernel

50

*oc

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JZ 25"

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Vp kernel

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dc/dp

(b) fret- surface

t t : . ") , U

f t . j i . P

/()

IiiU'jirite Upward*

I

\ / /

\

\ ,

Figure 3.5 (a) Sensitivity kernel functions for S-wave, P-wave and density respectively at periods of 20, 40, 67, 100 and 167 s, based on the average shear velocity model, (b) The spherical Earth model with homogeneous Vp, Vs and density in each layer.

30

3.5a) were computed based on the reference inversion model, showing that the strongest

sensitivity comes from the Vs profile, while the P-wave velocity and density have strong

sensitivity mostly to structures within the crust. Thus, I kept density and Vp/Vs ratio fixed

in each layer (Figure 3.5b), and used a combined P and S-wave sensitivity to calculate the

perturbation to the previous velocity model. In practice, no significant change occurred

with fixed shear velocity when slightly varying the ratio of perturbation of Vp with

respect to that of Vs. The longest period (167 s) in the phase velocity measurements

indicates that the upper mantle velocity structure can be well resolved to approximately

250 km in depth, though with much larger vertical uncertainties.

I inverted iteratively for the regionalized shear velocity on a 0.5°x0.5° grid to best fit the

Rayleigh phase velocity observations, and thus built the final 3-D Vs model. The shear

velocity structures within the study area were parameterized as several 10km layers in the

crust and mantle lithosphere, and as layers of increasing thicknesses (20, 30 and 50 km)

in upper mantle. I used a modified TNA model (Grand and Helmberger, 1984), a well-

known shear velocity model for the tectonic North America, as the starting model for

shear velocity inversion. The average Vs structure was first inverted by taking the mean

crustal thickness of 42 km. Subsequently, the local S-velocity structures beneath each

grid point were inverted from the initial TNA model with crustal thickness constraints

from the RF study, in order to maintain typical crustal and uppermost mantle velocity

values above and below the Moho in the starting model. The linearized inversion scheme

used combined criteria of minimum length and smoothing regularization by introducing

nonzero diagonal and off-diagonal terms in the a priori model covariance matrix Cmm.

31

The choice of a damping factor of 0.2 suggested by the obvious turning point in the In­

curve of the velocity model norm and residues of phase velocity (Figure 3.6a), also

agrees well with the choice of Schutt et al. (2008) by bounding the minimum shear

velocity in the upper mantle to a reasonable range.

(a) L curve (b) Rows of resolution matrix

0.035

50

100

150

- * 200

JZ 250 QL T3

300

350

400

450

-0.05 0 0.05 0.1 0.15 0.2 0,25

Residue

500 <— -0.2

15 km 55 km 105 km 155 km 205 km 310 km 425 km

0 0.2 0.4

resolution 0.6

Figure 3.6 (a) Damping parameter estimation from the L-curve of the Vs model norm and residues of the phase velocity prediction. The turning point with damping parameter of 0.2 is chosen for the Vs inversion, (b) Model resolution kernels for layers at depths of 15, 55, 105, 155, 205, 310, 425 km. Rows of the resolution matrix are calculated from the reference model.

32

3.2.2 Average shear velocity structures

The average lithospheric shear velocities beneath the entire study region as well as the

distinct tectonic regions were also inverted, as shown in Figure 3.7. I can observe

pronounced low velocity zones (LVZ), compared with the global compilation of the

continental and orogenic averages, with strong or moderate negative gradients underlying

the rigid lithospheric lid beneath the SBR, the NBR, the RGR and the CP. These

characteristics have also been noted in many continent-scale seismic studies (Humphreys

and Dueker, 1994a, b; Burdick et al, 2008; Grand, 1994; van der Lee and Nolet, 1997).

The inversion of Vs structures can be evaluated from the rank of the velocity parameters,

which represents the number of independent pieces of information in the model provided

by the dispersion data and the resolution kernels at different depths. The rank of the Vs

model has the largest value of 4.00 at the Colorado Rockies, the smallest of 3.24 in RGR.

The plateau has an intermediate rank of 3.84 between the SBR (3.62) and the NBR

(3.86). The rank for the rift area is lower than other regions due to larger uncertainties in

the phase velocities. In addition, in order to reveal the depth-dependent layer information

resolved by the observed dispersion relation, the resolution lengths for the reference

model were plotted in Figure 3.6 (b), with the resolution peak displayed around 20 km for

the row of the resolution matrix corresponding to 15 km. The number of adjacent layers

(10 km thick for each above 300 km) needed to recover one entire piece of independent

information of the model increases with the depth, because the resolution kernels become

broader in deeper structure, and more layers are needed to produce a summed resolution

component of 1.0. Thus, phase velocities at longer periods, which have deeper but

33

Dispersion curve Dispersion curve

(a)

•

average

50 100 Period (s)

150

TNA model j Average Vs I

3.4 3.6 3.8 4 4.2 4.4 4.6 4.8 Vs (km/s)

4.4

_ 4.2 45 E * 4 * o •55 3.8

| 3.6 0.

3,4

(c)

..*" ^^^*^ .

/ ' '

-

/ s<

Or | NBR | SBR i RGR i SRM |

E

C

D

0

20

40

60

80

100

120

140

160

160

200

220

240

260

280

300

50 <00 Period (s)

150

(d)

67.5 72.S

82.S 97.S

97.S

TNA modef CP NBR SBR RGR SRM

3.2 3.4 3.6 3.8 4 4.2 4.4 4.6 4.8 Vs (km/s)

Figure 3.7 Dispersion curves fitting and Vs inversion, (a) The predicted dispersion curves for the average (blue) phase velocities, (b) The average Vs (blue line) inverted from the starting modified TNA model (black) to fit the dispersion curve in (a), (c) The dispersion curves for five tectonic regions (CP, NBR, SBR, RGR and SRM). (d) The average Vs models beneath the corresponding geologic regions. The colors are the same as (c), and the black line is the initial TNA model. All error bars stands for the standard deviation of phase velocity measurements. Estimated LAB depths from the Rayleigh wave tomography at the center of the negative velocity gradients.

34

broader peaks of Vs sensitivity (Figure 3.5a), are less able to resolve deeper upper mantle

structures. The combined analysis from the Vs sensitivity and resolution kernels indicate

that the S velocity can be resolved well to a depth of- 250 km, above which vertical and

lateral seismic heterogeneities have been well imaged. The average depths of lithosphere-

asthenosphere boundary (LAB) beneath each tectonic province is inferred at the middle

depth of the low-velocity gradient, and the estimated depths show great consistency (see

Table 1) with the independent PdS and SdP receiver function studies (Levander and

Miller, 2010).

Table 1 Comparison of LAB depths (km) estimated from the PdS, SdP receiver function imaging and surface wave tomography (SWT)

NBR

SBR

RM

CP

RGR

Western U.S.

PRF

71.5±9.1

82.1±7.5

81.7 ±9.3

89.1±11.5

79.8±11.7

SRF

86.4±16.0

89.4±16.4

122.8±18.7

101.7±14.7

99.8±23.8

SWT

72.5

82.5

97.5 (SRM)

97.5

67.5

77.5 (average)

Note: The LAB depth in the last row for the SWT is averaged on the area in this study.

35

CHAPTER 4

RESULTS AND DISCUSSION

The final shear velocity model from the Rayleigh wave tomography reveals complex

lateral heterogeneity of the crustal and upper mantle structures related to the regional

tectonic features. After description and interpretation of the major low and high velocity

anomalies above 250 km, I will briefly discuss the geodynamic interactions between the

lithosphere and underlying asthenosphere based on the well-imaged low-velocity

channels. Last but most importantly, the comparison between this Rayleigh wave

tomographic result with the PdS receiver function volumes shows that the ongoing

delamination beneath the western portion of the CP could be a significant contributor to

the complexity of the crust/mantle boundaries and a cause of the recent CP uplift.

4.1 Crustal structure

4.1.1 Crustal heterogeneity

The crustal structure beneath the CP and the surrounding tectonic regions is characterized

by relatively short-wavelength high and low velocities (Figure 4.2 a, b). Due to lack of

short period surface wave data (< 20 s) and without consideration of the sedimentary

layer information, the upper crustal structure is relatively poorly resolved. However,

several distinct low velocity features can still be identified in the first several layers of the

Vs model with reasonable resolution, including three narrow transition zones (~ 50-100

km in width) along the southwestern, northwestern and southeastern margins of the

plateau, one belt (COMB) beneath the SRM (with the strongest anomalies under the San

36

(a) Surface heat flow -11S: -1*4 : -1*2 •10: -rt»8: -106:

4Z

3S"

32;

150

12b

1C0

E

E J,J

2b

L-J 0

-116: -1'4= -1*2" - '10 : -^08 : -106:

(b) Cenozoic magmattsm i8 : -1CKT

,—i 4500

i

3030

2500 "g"

2000 c

.£ 1500 a

1000 *3

509

0

-500

-1C00 -11G" -1*4" -V.'l -*1

Figure 4.1 (a) Surface heat flow maps (data from www.heatflow.und.edu) after interpolation. Several unusually high heat flow data are excluded, (b) Distribution of the Cenozoic volcanic rocks showing the magmatic pattern in this study area (data from NAVDAT: Western North American Volcanic and Intrusive Rock Database). Different colors indicate different rock ages (red: Tertiary, black: Quaternary), and different symbols indicate different types (circle: felsic; diamond: mafic; inverse triangle: intermediate). Topography from ETOPOl model is also plotted.

37

L JSBH 90kmJ

-'1G'-*'•»'-'" 2'-1'0'-''36'-T Co b / _ SfTOftto 90 kmj

r e _"4'-"!2' -"O'-'CE'-noc-' j^ff _Cru«tal_thlcknesirj ,„

.-•e _--i-_r2-*-0"-'cs-'C5'

33 4Z 50 SO d l n V s (%)

Figure 4.2 Maps of Vs anomalies (a-i) at depths of 15, 30, 60, 80, 100, 125, 150, 200 and 250 km, relative to TNA model, and comparison of Vs maps at 90 km from this study (j) and the S body wave tomography (k; Schmandt and Humphreys, 2010). The crustal thickness (1) is estimated from the PdS receiver functions (Levander and Miller, 2010). Two volcanic fields are marked: the San Juan volcanic field (SJVF) and the Navajo volcanic field (NVF). SMB in fig. a, b stands for the Socorro magma body. The red dot line shows the station distribution of the La RISTRA experiment.

38

Juan volcanic field and Aspen in central Colorado, see Figure 4.2 a, b), and one small

anomaly in the central RGR. The large-scale low crustal Vs anomalies beneath the CP-

BR Utah transition zone and the Colorado Rockies are broader than those in the CP-BR

Arizona transition zone and the NE trending Jemez lineament in the RGR-CP transition

region. The relatively high velocities under the SBR and the NBR are separated by a low

velocity zone located mostly in southern Nevada between 37°N and 40°N. The high NBR

velocities dropped abruptly to the east in north-central Utah near the Wasatch Front.

The plateau interior is relatively complex and has an intermediate average velocity

between the RGR and the stable Great Plains to the east of the RGR and SRM. The NE

trending high velocity anomaly in the southern plateau is bounded by the Jemez

lineament to the southeast and the Yavapai/Southern Yavapai accretionary boundary to

the northwest. It extends from the southeastern Navajo volcanic field (NVF) region to the

western edge of the SRM, and drops abruptly across the SRM-CP boundary with a very

sharp velocity gradient into the COMB. The relatively low velocity area to the west of the

northwestern NVF is consistent with the result of La RISTRA group (West et al, 2004;

Gao era/., 2004).

4.1.2 Interpretation

The seismic complexity within the relatively cold and stable plateau, especially the lower

crustal structure beneath the western plateau, is possibly associated with dynamic crustal

structure reset by a possible delamination beneath this region (see section 4.4). The high

velocities in the eastern part of the Navajo volcanic field could be related now cool mafic

39

intrusions (West et ah, 2004), while the relatively low velocities could result from the

heating from the ascending asthenosphere at the base of the crust under the western

plateau (west of the Navajo Volcanic Field).

The two remarkable low velocity anomalies under the COMB and the Jemez lineament

are consistent with previous tomographic images (e.g. Humphreys and Dueker, 1994b;

Dueker et ah, 2001; Humphreys et ah, 2003) and the refraction/wide-angle reflection

seismic results such as the Deep Probe line (Henstock et ah, 1998; Snelson et ah, 1998;

Gorman et ah, 2002) and the CD-ROM transect (e.g. Levander et ah, 2005, Snelson et

ah, 2005). The low crustal velocity feature under the SRM can be attributed to the results

of one of the most active Laramide-age magmatic activities (Mutschler et ah, 1987) in the

COMB region and the more recent (-37 Ma) large-volume San Juan volcanism. Both

volcanic events have caused compositional modification and produced partial melting of

the crust by heating the base of the COMB crust (Dueker et ah, 2001; Humphreys et ah,

2003). The COMB abutting the SRM-CP boundary is also the site of a large heat flow

contrast. The other crustal low velocity anomaly occurs along the Jemez volcanic

lineament (including the low anomaly under the Mount Taylor with complex sill and dike

structures) and corresponds to the suture between two Proterozoic island arc terranes: the

Yavapai-Mazatzal terrane to the north and the Mazatzal province to the south (Karlstrom

and Humphreys, 1998). The low velocities under the Jemez lineament match well with a

variety of geochemical and magmatic signatures that define the Jemez volcanic trend

(McMillan et ah, 2000). The low crustal velocities coincide with the high heat flow and

recent basaltic magma intrusion into the lower crust (Karlstrom and Humphreys, 1998;

40

Levander et al, 2005), as well as the low Bouguer gravity anomaly (Roy et al, 2005),

suggesting that in places the crust beneath the Jemez lineament is partially molten

(Hammond and Humphreys, 2000; Levander et al., 2005). The lack of a low velocity

signature directly beneath the RGR axis is consistent with recent magmatic quiescence

directly beneath the rift. The lithospheric thinning and rifting in the RGR as well as the

mechanical weakness caused by the plateau rotation could also contribute to reduce the

velocity beneath the Jemez lineament. One of the observed low velocity anomalies in the

central RGR is identified as the Socorro magma body (a mid-crustal sill structure) in New

Mexico (e.g. Balch et al, 1997), which is consistent with the observations by the

RISTRA group (West et al, 2004) and the tomographic result based on combined

ambient noise and surface wave data (Shen et al, 2009).

The Late Tertiary and Quaternary magmatism could have caused partial melt and reduced

the shear velocities along both the Utah and Arizona transition zones between the CP and

the Basin and Range (Wolf and Cipar, 1993; Benz and McCarthy, 1994; Parsons and

McCarthy, 1995; Parsons et al, 1996; Frassetto et al, 2006). Outside the transition

zones, the relatively high crustal velocities in the SBR and the NBR arise because of the

thinner crust in these regions (Figure 4.2 1, Levander and Miller, 2010; Frassetto et al,

2006)): The surface waves are sampling the lower crust and upper mantle compared to

middle crust of the CP interior at the same depth. In particular, in Figure 4.1b, the

strongest high velocity anomaly under the SBR at 30 km is larger than those elsewhere,

because the velocities at that depth represent both lower crust and upper mantle

41

lithologies. The low velocity zone between the SBR and the NBR, with distinct

nonbasaltic magmatism and extension to the north and to the south respectively,

coincides with the "amagmatic zone" close to the Utah-Arizona-Nevada tristate area (e.g.

Axen et al, 1993; Zandt et al, 1995). In a similar interpretation of the velocity difference

between the Basin and Range and the plateau interior, the seismic velocity reduction can

be attributed to thickened crust (-30-40 km, Levander and Miller, 2010; Zandt et al,

1995) in this region that have experienced more recent extension (<15 Ma, 60-100%)

than the Early and Late Tertiary intraplate extension (100-200%) in the BRR

4.2 Upper mantle structure

The Rayleigh wave tomography of the upper mantle is in very good agreement with the

independent P and S body-wave traveltime tomographic images from Schmandt and

Humphreys (2010) using USArray/TA and other data. For example, compare the shear

velocity anomaly map at 90 km shown Figure 4.2(j) with the surface wave tomography at

similar depths (Figure 4.2 k), the lateral heterogeneity patterns are similar, both showing

a large and continuous low-velocity anomaly surrounding the high velocity body in the

CP interior and its north. The slightly different magnitudes of low and high velocity

anomalies are largely due to the radial anisotropy of the S velocity structure, in which the

Rayleigh waves are sensing the vertical component (Sv) while the body-wave traveltime

tomography is inverting the SH component. Such radial anisotropy could reach more than

5% beneath the BRP while it stays low inside the plateau (Moschetti et al, 2010). In

general, the Rayleigh wave tomography model shows a high consistency with the

previous studies. The pattern of the lateral seismic heterogeneity strongly correlates with

42

the recent magmatic activities and the surface heat flow data (Figure 4.1a, b). The shear

velocity variations in the upper mantle are described first, followed by the interpretation

at each tectonic province.

4.2.1 Upper mantle seismic heterogeneity

The strong lateral seismic heterogeneity in the upper mantle beneath the CP and the

adjacent regions is clearly imaged by the surface wave tomography. A large high velocity

anomaly beneath much of the CP (1/2 to 2/3), extending from the Navajo volcanic field

in the south across the NE striking accretionary boundary (Cheyenne Belt) to the

southern edge of the Archean Wyoming Province. This high velocity body is consistently

observed from the uppermost mantle (60 km) to deeper upper mantle (-200 km) (Figure

4.2c-h). It is surrounded by a continuous, large-scale and strong low velocity anomaly

around the plateau rim beneath the southernmost Rocky Mountains, the rifted RGR, and

the extended SBR and NBR regions. At greater depths, the relatively high velocity

anomalies weaken, and a NE-trending narrow low velocity belt (e.g. Figure 4.2 f-h) is

observed to have penetrated into the high velocity body in the northwest of the Four

Corner region, which may indicate that the lithospheric root of the plateau are suffering

convective erosion from the underlying asthenosphere.

The tomographic images (Figure 4.2 c-h) also show one of the largest amplitude low-

velocity anomalies from depths of 60-150 km beneath the NBR to the west of the

Wasatch Front in Utah. The strong and sharp contrast between this low and high velocity

anomaly under the abutting CP interior is in good agreement with the extended traveltime

43

tomographic result from the new La RISTRA 1.5 data (Sine et al, 2008). The strong low

velocity zones, though bounded by the normal faults at the western edge of the plateau,

penetrate the margins of the less deformed plateau and become broader at greater depths.

The penetration distance from the SBR is about 100-150 km, greater than that from the

NBR (50-100 km). The inferred thickness of the mantle lid beneath the NBR from the

Rayleigh wave tomography model is relatively small (-20 km), consistent with the

observations from the inter-station surface wave technique using the CP-GB data

(Sheehan et al, 1997; Lastowka et al, 2001). The thin mantle lid is underlain by a more

steep negative velocity gradient than beneath the CP (Figure 3.7), with minimum Vs as

low as ~ 4.1-4.2 km/s at ~ 100 km (~ 7-8% slower than the global average). In the

southwestern margin of the plateau, a similar distinct LVZ with a thin mantle lid in the

transition region (identified by the PACE group, Benz and McCarthy, 1994; Parsons and

McCarthy, 1995) as well as the recent uplift, are suggested to be a delayed mechanical

response to the lithospheric thinning after the low-angle slab subduction and foundering

(Parsons and McCarthy, 1995). From the southwestern margin to the plateau interior, the

physical state of the upper mantle becomes more complicated, and an ongoing

delamination may provide an explanation for this complexity (see section 4.4).

In the southeastern plateau and the rift, a "lambda-shaped" (k) low-velocity pattern in

New Mexico is observed beneath the southern RGR and the Jemez volcanic fields

(Figure 4.2 c-e). These two low velocity trends merge at the Jemez Mountains close to

the New Mexico-Colorado boarder, and extend northwards into the SRM with slight

increase of Vs. The smallest shear velocities of ~ 8% lower than the global average are

44

centered directly beneath the rift axis and along the Jemez lineament (Figure 4.2 c-e),

extending from shallow uppermost mantle ~ 40 km to ~ 100 km deep, at which depth the

"X" shaped low velocity feature (Figure 4.2 f, g) blends with the large-scale SBR low

velocity anomaly. The velocities of the Colorado Rockies, though not as low as the Jemez

lineament, are still relatively low compared with those under the northern plateau and

north of the Cheyenne Belt. This "X" shaped low-velocity pattern is bounded by high

velocity anomalies beneath the plateau interior to the northwest of the Jemez lineament

and by the western edge of the Great Plains to the east of the rift. A large velocity

contrast, up to -10%, exists between the Jemez lineament and the high velocity anomaly

that separates the Jemez lineament and low velocities beneath the northwestern Navajo

volcanic field. This remarkable lateral velocity gradient from the rifting-related volcanic

fields to the stable plateau interior has also been clearly imaged by the RISTRA group

using both P and S traveltime tomography (Gao et ah, 2004; Sine et ah, 2008) and

surface wave measurements (West et ah, 2004), though their tomographic results are

limited to a single line (as indicated by the red dot lines in Figure 4.2 a-j). The higher

velocities to the east of the rift, determined by numerous seismic investigations {e.g. van

der Lee and Nolet, 1997; Dueker et ah, 2001), stop the rift low velocity zone at its eastern

edge which is associated with the westernmost edge of the cratonic lithosphere of the

North American Plate.

At a given depth above 200 km (Figure 4.2 c-g), the upper mantle shear velocity

variations can reach as high as ~ 12%, generally due to the contrast between high

velocities in the northern plateau and low ones along the CP-BR transition zones. This

45

strong lateral heterogeneity weakens with greater depths (Figure 4.2 h, i) due to relatively

weak sensitivity of the phase to shear velocities and lateral averaging effect at long

periods in the Rayleigh wave tomographic technique. The relatively high velocity relative

to the TNA model starts to dominant below 250 km (Figure 4.2 i), showing signature of

the base of sublithospheric low velocity zones (see Figure 4.3). This velocity increase

could be related to the change of grain size from ~ mm to ~ cm (Faul and Jackson, 2005).

4.2.2 Colorado Plateau mantle lithosphere

The CP interior has low heat flow (< 50 mW/m ) and relatively thick Proterozoic

lithosphere compared to the surrounding BRP and RGR, as inferred from the surface

wave model (see Table 1, Figure 3.7, Figure 4.1a) and previous studies {e.g. Thompson

and Zoback, 1979). Such cold and thick Proterozoic mantle lithosphere beneath the

plateau forms a thick thermal boundary layer, and thus shields the crust from deformation

and magmatism by limiting the heat transfer from the convecting mantle to the crust. This

thermal boundary layer is less dense and more ion depleted mantle lithosphere with

greater strength compared with the BRP, as proposed by Lee et al. (2001) based on the

peridotite xenoliths data. The slightly decreased Vs (~ 4.4 km/s) in the mantle lid of the

plateau compared to the global average (4.5 km/s), is consistent with the measurement of

relatively low Sn velocities (Nolet et al, 1998) in the same region. Even though a large

part of the plateau mantle lithosphere survived from the Laramide flat-slab subduction,

moderate hydration from the de-watering after the sinking of the Farallon slab could

cause minor reduction of the mantle lid shear velocity. This slightly low Vs and the

relatively high measured Pn of 8.0-8.1 km/s {e.g. Beghoul and Barazangi, 1989) yield a

46

high Vp/Vs ratio (-1.80-1.85), which has been observed by receiver function studies (e.g.

Levander and Miller, 2010; Bashir et al, 2009).

The crustal thickness of the CP used in this Rayleigh wave tomography from the PdS RF

estimates, is shallower than previous RF studies (e.g. Zandt et al, 1995; Gilbert and

Sheehan, 2004; Wilson et al., 2005a), and does not provide sufficient isostatic support for

the present 2-km elevation. The lithospheric thickness (-100 km) of the plateau inferred

from the surface wave model (see Table 1, Figure 3.7) rules out the almost complete

removal of the continental lithosphere beneath the plateau following flat-slab subduction,

as suggested by Bird (1988). In fact, the estimated LAB depth is largely consistent with

the partial delamination model using the mantle buoyancy to explain the anomalous

plateau elevations (Beghoul and Barazangi, 1989; Spencer, 1996; Lastowka et al, 2001).

It is also consistent with the fact that the cored mantle lithosphere imaged in the Vs model

preserves the Proterozoic signature including the relatively high seismic velocities and

strength, as observed in the xenoliths, despite small thermal modifications after removing

partial base of the lithosphere at - 30 Ma (Spencer, 1996). The relatively low surface heat

flow indicates that the recent heating at the base of the thick plateau lithosphere has not

yet reached the surface (Thompson and Zoback, 1979). In addition to the isostatic support

of the crust and dynamic uplift (Moucha et al, 2008, 2009), a mantle contribution to the

current uplift of the plateau is necessary in the form of asthenospheric support beneath

the low-density plateau lithosphere. Based on the surface wave tomography model in this

study, I favor the partial delamination model (Beghoul and Barazangi, 1989; Spencer,

1996; Lastowka et al, 2001) combined with dynamic uplift, isostatic response, and

47

asthenospheric buoyancy to support the current elevation of the plateau.

4.2.3 Basin and Range and CP-BR transition zones

These extremely low upper mantle velocities beneath the Great Basin cannot be merely

caused by the ~ 200-400°C temperature (Riter and Smith, 1996; Sine et al, 2008) or

compositional variations, but they also require the presence of partial melt and water

content in the nominally anhydrous minerals in the upper mantle lithosphere (Hammond

and Humphreys, 2000; Karato, 2006; Sine et al, 2008). The hydration-induced melting

is a consequence of the slab rollback (Humphreys et al., 2003) and the intensive Basin

and Range extension in Early and Late Tertiary. The differences between the magnitudes

of the low Vs in the northern and southern Basin and Range are probably related to

different intensity and history of the Cenozoic extension and magmatism separated by the

"magmatic gap" zone in the central Basin and Range, such as the chemical and thermal

variations in the upper mantle in the NBR and SBR (Armstrong and Ward, 1991; Axen et

al, 1993; Humphreys and Deuker, 1994b; Zandt etal, 1995).

The various penetrating distances agree well with the different magmatic encroachment

rates of 6.3 km Myr"1 in the SBR and only 4.0 km Myr"1 in the NBR (Roy et al, 2009).

These dominant low velocity features could be associated with the small-scale convection

(e.g. van Wijk et al, 2008) around the flanks of the plateau lithosphere with driving

forces from the warm BRP (characterized by 100-200% extension since Mid-Tertiary)

asthenosphere at relatively shallow depths (50-100 km), and is accompanied by increased

seismicity and a high surface heat flow regime (Figure 4.1a). This small-scale convection

48

could also provide dynamic uplift support from the northwestern and southwestern edges

of the plateau, as explanation for the excess marginal elevation. Additionally, the recent

Late Tertiary and Quaternary magmatism and the significantly metasomatic alternation at

the margins of the plateau (Roy et al, 2005, 2009; Smith and Griffin, 2005) could have

produced certain percentages of melt when the ascending asthenospheric materials heated

the base of the thin lithosphere between the extension regime and the relatively thick

Proterozoic lithosphere under the plateau.

4.2.4 Rio Grande Rift and Jemez lineament volcanism

This "X" shaped anomaly structures could have resulted from the Proterozoic structural

control on the mantle lithosphere similar to the crust (Karlstrom and Humphreys, 1998),

enhanced by the intrusion of hot asthenospheric materials to the base of the lithosphere

after removal of the Farallon slab. The temperature variations across the rift and the

mechanical thinning during the rifting could both contribute to such large velocity

gradients across the CP-RGR-GP zone. More recent magmatic and tectonic activities (in

the last 10 Myr) after the post-Laramide slab removal and Basin and Range extension is

likely caused by small-scale convection at the edge of the lithospheric root roughly

beneath the rift axis, which introduces a lateral temperature difference. Recent volcanic

fields along the Jemez lineament, rather than the RGR, coincide with a regional extension

induced by the small clockwise rotation of- 3° (Aldrich et al, 1986; Hamilton, 1988) of

the plateau that weakened the lithosphere abutting the rift and allowed magma ascent.

49

4.2.5 Southern Rocky Mountains

The upper mantle velocity structures beneath the SRM bear certain resemblance to the

Deep Probe, the CD-ROM and the RMF (Rocky Mountain Front Broadband Seismic

Experiment) projects (Snelson et ah, 1998; Henstock et ah, 1998; Gorman et ah, 2002;

Levander et ah, 2005; Dueker et ah, 2001; Li et ah, 2005), and exhibit more complicated

heterogeneous structures in 3-D perspective views of the shear velocity model (Figure 4.2

& 4.3). The less pronounced low velocity zones beneath the SRM could be caused by

about 1% partial melt in the upper mantle along the Colorado Mineral Belt (Dueker et ah,

2001; Humphreys et ah, 2003). The southern edge of the Rockies along the Jemez

lineament also correlates well with the Yavapai-Mazatzal boundary, further supporting

the structural control on tectonism by the ancient accretionary boundaries that weakened

the lithospheric zone (Karlstrom and Humphreys, 1998). The mantle velocity

perturbations beneath the SRM are not as strong as those in the upper and middle crustal

structures (Figure 4.2 c-g). However, the moderately low velocities compared to the

continental velocity still need a small amount of partial melt in the upper mantle

combined with a high temperature.

4.3 Sublithospheric structure

The surface wave tomography model has clearly imaged the sublithospheric channels of

low seismic velocity beneath the CP and the surrounding areas. Low velocity channels

are typically associated with low-viscosity zones, frequently attributed to the combined

weakening effects of water and partial melt. It stands to reason that I am imaging a sub­

continental low-viscosity asthenosphere. The asthenosphere plays a key role in

50

-116 -114 -112 -110 -108 -106

E

I

-116 -114 -112 -110 -108 -106

Vs (km/s) 4.0 4.1 4.2 4.3 4.4 4.5 4.6

Figure 4.3 Vs cross-sections along 33.5°, 34.5°, 35.5°, 36.5°, 37.5°, 40.5°N showing location of low-velocity channels, with tectonic boundaries in red ticks and suggested delamination location in white arrows. The low-velocity channel beneath the RGR is relatively shallow compared with the Basin and Range Province.

51

understanding fundamental mantle convection and plate tectonics questions. Recent

numeric and analytic work revealed that flow in the asthenosphere is a Poiseuille-Couette

flow associated with the wavelength of convection and the relative strength of lithosphere

and asthenosphere (Hoink and Lenardic, 2008, 2009). The accurate depths of the

sublithospheric low-viscosity channels are difficult to be identified with Rayleigh waves

because these are relatively insensitive to the seismic boundaries of velocity changes. The

integrating nature of the measurements cannot distinguish sharp from gradual velocity

gradients beneath the lithosphere (Eaton et al, 2009). However, the period-dependent

sensitivity peaks to the absolute shear velocity over a certain depth range still provide

information on the relevant depths and thicknesses of the low-velocity channels in the

sublithospheric structures.

Six cross-sections along 33.5°, 34.5°, 35.5°, 36.5°, 37.5° and 40.5°N, as shown in Figure

4.3, clearly present the low-velocity channels beneath the southwestern US. The base of

the asthenosphere, which can hardly be inferred from the PdS or SdP receiver function

analysis, can be well estimated here from the surface wave tomographic images to study

the geodynamic interaction between the lithosphere and underlying asthenosphere at

various regional tectonic regions. The dominant strong low-velocity channels beneath

both southern and northern Basin and Range provinces are shallow (60-120 km) but

strong, and the magnitudes decrease to some extent when the region is far from the

plateau margins (Figure 4.3). The short-wavelength low-velocity asthenosphere beneath

the RGR is even shallower (to ~50 km) and a bit stronger compared to the Basin and

Range region, and this low-velocity channel is thinner (see the 34.5° cross-section in

52

Figure 4.3). The variety of the depths and thicknesses of the low-velocity channels

beneath the Basin and Range provinces, the RGR valley and the CP is essential to

understanding the current geodynamics after a complex superposition of multiple tectonic

and magmatic effects during the Cenozoic era.

The relative shallow but strong low-velocity zones were likely developed during Basin

and Range extension, as the lithosphere was thinned by a factor of up to 2. Basin and

Range asthenosphere might exert strong convective erosion to the relatively undeformed

plateau with increased depths. It correlates well with the migration of magmatism into the

margins around the Colorado Plateau from south, southwest and northwest. The

development of a strong low velocity zone was probably enhanced by hydration from the

Farallon slab, as well as the strongly coupled mantle convections rising to the shallow

base of the Basin and Range lithosphere. The slightly decreasing magnitudes of the low-

velocity (low-viscosity) channels away from the front of the magmatic encroachment

may indicate a cooling effect after the magmatic front moved away.

The consensus on classification of the Rio Grande Rift as a passive or active rift has yet

to be achieved. The observed emplacement of the hot asthenosphere (roughly to 150 km

at the base) at the replaced lithosphere below the rift in the Rayleigh wave model is

consistent with both surface and body wave traveltime tomographic results from the La

RISTRA group (Gao et al, 2004; West et al, 2004; Wilson et al, 2005b) and the

dynamic test from the numerical modeling (van Wijk et al, 2008), indicating that the Rio

Grande Rift may be a recent passive rift caused by the small-scale convection in response

53

to far-field-driven forcing of lithosphere. The consistency argues that it is relatively

unlikely that the RGR experienced an active rifting process controlled by the mantle

upwelling from deep source suggested by Moucha et al. (2008) based on the joint

seismic-geodynamic simulation.

The shear velocity model documents large lateral heterogeneity in depths thickness, and

minimum velocity in the low-velocity channels and maximum velocity in the overlying 1

lithosphere. The variations in the asthenosphere could be caused by the lateral

temperature gradients, which can generate strong pressure-driven flow in the

asthenosphere. In addition, a small amount could also be ascribed to the compositional

variations, and water content or melt. The asthenospheric channel beneath the Colorado

Plateau is moderately weak and shows less shear velocity reduction compared to the

surrounding BRP and RGR valley. The step in lithospheric thickness between these

distinct tectonic regions could cause strong lateral interaction between the ascending

asthenosphere and the lithosphere besides the vertical interaction (van Wijk et al, 2008).

The transitional regions of the lithospheric thicknesses strongly correlate with strongest

low-velocity zones that are at the margins between different tectonic regions. The well-

imaged low-velocity channels beneath the study area can help to understand the regional

lithosphere-asthenosphere interactions beneath the extended, rifted and undeformed

regions from various evolutions of mantle structures since Late Cretaceous.

54

4.4 Possible delamination

4.4.1 Comparison with receiver function imaging

The Rayleigh wave tomography model shows strong correlation with the PdS teleseismic

receiver function (RF) image volumes at the same area, despite the differences in

frequency bands (0.1-0.75 Hz in RF; 0.006-0.05 Hz in the Rayleigh wave tomography).

The receiver function imaging, which represents the converted teleseismic wave

responses beneath the stations, was obtained by applying a time-domain iterative

deconvolution approach to 106 teleseismic events recorded by the USArray/TA stations

(Levander and Miller, 2010). After converting the individual receiver functions from the

time domain to the depth domain, common conversion point (CCP) stacking was used to

migrate them laterally to the conversion point based on a 3-D reference velocity model.

The model was constructed from the Crust 2.0 model (Bassin et al, 2000), the University

of California at San Deigo upper mantle model (Buehler and Shearer, 2009), and the

University of Oregon P and S body wave tomography model (Schmandt and Humphreys,

2010). CCP stacking enhances the image by stacking redundant signals from the same

conversion point. The final stacked RF images reveal unusually complicated Moho

topography beneath the Colorado Plateau, especially beneath the northwestern portion of

the plateau, in which the Moho broadens and separates into two distinct positive-

amplitude events in the interior of the plateau. The top positive event is relatively shallow

compared with previous estimates of crustal thickness in the CP, while the deeper one

extends to depths (> 50 km) greater than the typical estimates and can hardly be

interpreted as the thickened Moho.

55

Figure 4.4 Map view of LAB depths picked from the PdS receiver function profiles with surface wave tomographic result as a guide.

The lithosphere-asthenosphere boundary (LAB) depths can also be better estimated from

the RF volumes with additional constraints from the surface wave tomography model in

this study. The LAB depth estimates of- 60-120 km from the PdS RF images, as shown

in white dash lines corresponding to the negative amplitude events, are well correlated

with the top of the low-velocity channel in the Vs images. This strong correlation

suggests that the Vs images can assist with determining LAB depths, particularly when

there are multiple negative events due to contamination from noise, reverberation, or

56

subsurface discontinuities within the upper or middle crustal structure, even though there

are larger uncertainties in the estimate of LAB depths directly from the Vs images than

from the RF due to the limitation on vertical resolution in the Rayleigh wave study.

Figure 4.4 shows the final estimated LAB depth map from the combined analysis of these

two studies, showing that the Basin and Range Province has relatively shallow LAB at

approximately 50-80 km, the CP has a deeper LAB of- 80-120 km, and the SRM LAB is

even deeper to -150 km.

Most importantly, the combined interpretation of these two studies suggests that a portion

of the lower crust and uppermost mantle are delaminating along the crust/mantle

boundary, and a new Moho is being formed at relatively shallow depths. Figure 4.5

shows a comparison of two profiles along 37°N and 111°W in these two independent

studies (except that crustal thickness constraint in the Rayleigh wave study comes from

the PdS receiver function estimates), with selected receiver function profiles (light red

means positive amplitude, while light green means negative) at every degree are also

displayed in the Vs images for direct comparison. As shown in Figure 4.5, a high velocity

anomaly is clearly imaged between 113°W and 111°W as hanging on to the western edge

of the plateau along the 37°N line, though more tilted than the second positive event as

imaged in the RF cross section. The key features in the RF images and the Vs images are

consistent, showing that a relatively high velocity anomaly appears to detach from the

plateau interior towards the western margins of the CP. Along the N-S profile at 111° W, I

also observed a high velocity anomaly with horizontal scale of ~ 1 ° in the Vs image, the

top of which we correlate with the lower part of the detachment as imaged in the RF

57

(a) 37 N Latitude

BRP CP

W

*

E W

\ '••

, It ;.

CP SRM .

'..: ^ * T .,»^;"-

KAJ aU^.^J MLu m i d rt , Ay , *j*W

•1C -*.J .'.2 .',.3 -V- -lb 4.6

1 3

(b)

SRM • H -

,'tf » 1' v**i

111W Longitude CP SBR SRM CP

S N

I » tl.

^ V ^ ^/ ̂ H- /*4

SBR

Amplitude

Figure 4.5 Comparison of cross-sections of shear velocity structure (left) and 1.25 Hz PdS receiver function images (right) along 37°N (a) and 111°W (b). Top of each image shows the elevation along the lines with arrows pointing at the Monument uplift (MU) where the delamination may start. I set the crustal velocity to 3.9 km/s for display, and plotted receiver function profiles (positive in light red and negative in light green) on the Vs cross-sections. Black solid lines and white dash ones indicate the Moho and LAB depths, respectively.

58

cross-section. This part of the high Vs anomaly is connected to a vertical high anomaly in

the northern part of the plateau, but cannot be imaged in the PdS images because the

receiver functions are not able to resolve vertical interfaces. The Vs model from the

Rayleigh wave tomography in this study supports the delamination hypothesis suggested

by the PdS RF study (Levander and Miller, 2010) to explain the two split positive events

in RF profiles beneath much of the plateau.

4.4.2 Possible delamination mechanism

One of the generally accepted descriptions of the delamination process is that the mantle

lithosphere peels off the overlying crust along the crust/mantle boundary and the

detached lithosphere sinks into the underlying asthenosphere (Morency and Doin, 2004).

The suggested ongoing delamination beneath the western and northwestern portions of

the Colorado Plateau develops from the observations of the high velocity anomaly in the

surface wave tomography and the second Moho-related positive event in the PdS receiver

function images. Although the mechanism of this regional delamination is not clearly

understood yet, the combined effects of regional surface uplift, magmatic underplating in

the lower part of the crust, and the high seismic attenuation might initiate the detachment

of the lower crust and uppermost mantle beneath the western plateau, probably starting

from the west of the Navajo volcanic field.

The newly re-dated Monument Uplift could be associated with the initiation of the

delamination in the western plateau. Along the 37°N profile in Figure 4.5a, the regional

59

elevation of the southern edge of the Monument Uplift (west of the Navajo volcanic

field) between 111°W and 110°W is relatively high and strongly correlates with the

delamination location where the deeper warm asthenospheric materials might rise up to

the base of the plateau crust, as indicated by the low velocity anomaly in the Vs cross-

section, or the negative event between the two separated positive events in the receiver

function image. Although the dominant rock/surface basement-cored uplifts and

monoclines in the CP are associated with the Sevier-Laramide orogenic deformation

(140-45 Ma) from Late Cretaceous to Early Tertiary, the rise of several small-scale

regions in the western and northwestern plateau such as the Circle Cliffs, the San Rafael,

the Monument, and part of the Kaibab uplifts were estimated to be formed more recently

(approximately 33 to 11 Ma after Laramide orogeny) by using the (U-Th)/He

thermochronology technique (Stockli et ah, 2002; Flowers et ah, 2008; Davis and Bump,

2009). This more recent regional uplift could respond to the initiation of delamination

that recently occurred.

The thickened crust caused by the regional crustal compression and magmatic

underplating could probably initiate the delamination, occurring along the localized and

weakened lower crust due to hydration and heating after the rollback of the Farallon slab.

The nearly west-east compression in the western part of the CP near 37°N (Saleeby,

2003) could cause a small part of the crustal thickening, and furthermore, the magmatic

intrusion or underplating at the base of the lower crust was suggested to contribute most

of the crustal thickening beneath the western plateau (Morgan and Swanberg, 1985; Wolf

and Cipar, 1993). The crustal thickening could stretch the geotherm under the southern

60

Monument Uplift and the magmatic underplating could also raise the Moho temperature,

which further thermally weakened the bottom of the lower crust. Additionally, due to the

suggested strong small-scale convection from the BRP around the flanks of the mantle

lithosphere beneath the plateau interior, the lithosphere of the western plateau was further

modified convectively and heated at shallower depths by the asthenosphere than the

preserved Proterozoic lithosphere beneath the plateau interior. The dehydration of

serpentinites in the cold Farallon subduction plate brought aqueous fluids to the

overriding continental lithosphere during subduction, further hydrated the plateau

lithosphere (e.g. Lee, 2005; Li et ah, 2008) and reduced the viscosity at the basal crust

(Humphreys et ah, 2003; Dixon et ah, 2004; Hyndman et ah, 2005). Furthermore, the

delamination started coincidently at the place with strong seismic attenuation of P and S

velocity (Ringler et ah, pers. comm., 2010). The thickening of the crust and the

weakening at the base of the crust could trigger the delamination that the lower part of the

crust and the upper mantle detached from the low-viscosity lower part of the crust and

fell into the warm and mobile asthenosphere driven by negative buoyancy (Meissner and

Mooney, 1998), and such delamination may be the cause of recent regional uplift of the

western plateau. This delamination mechanism described here is similar to the one

suggested by Morency and Doin (2004) based on 2-D numerical simulations using a

brittle-ductile rheological model. The possible ongoing delamination will reset the Moho

above its former depth and is therefore a dynamic crust/mantle boundary.

The initiation location of the delamination, such as under the southern Monument uplift

along the 37°N profile, is limited to the west of the Navajo volcanic field where no

61

significant magmatic underplating was suggested based on mantle xenolith evidence

(Mattie et al, 1997; Condie and Selverstone, 1999; Selverstone et al, 1999). Though the

garnetite and the associated eclogite xenoliths provide some evidence of mantle hydration

and metasomatism related to the Farallon subduction (Smith et al, 2004; Smith and

Griffin, 2005), the eastern side of the delaminating region has a relatively thick

lithosphere that could prevent the heating and weakening across the crust/mantle

boundary from the underlying asthenosphere (Figure 4.3). However, no significant

volcanic activity or thermal signals have been observed to be associated with the

suggested delamination. The moderately low velocity anomaly is indicative of the

replacement of the lower crust and upper mantle by the upward migration of the

asthenosphere though the velocity is not as low as under the BRP. An estimate of the

conductive time constant of about 13 Myr (using thermal diffusivity of 10"2 cm2/sec and

crustal thickness of ~ 40 km) may indicate that even if the base of crust was heated by the

ascending asthenosphere, the heat has not reached the surface yet, providing the

delamination occurred within 13 Myr. More geochemical and geophysical evidence is to

be investigated to support this delamination hypothesis.

62

CHAPTER 5

CONCLUSIONS

This study has developed a new high-resolution 3-D shear velocity model (to the depth of

~ 250 km) from the Rayleigh wave tomography technique using the latest USArray data

to reveal in detail strong seismic heterogeneity in the crust and upper mantle beneath the

CP and the adjacent tectonic provinces in the southwestern U.S.. Relatively high

velocities are observed in the Proterozoic mantle lithosphere beneath the CP interior, and

extend to the Archean province in the north. Continuous large-magnitude upper mantle

low velocity anomalies surround and underlie the edges of the plateau in the northwest,

south and southeast, including the CP-SBR and CP-NBR transition zones, the COMB,

the Jemez lineament and the RGR. These may be indicative of small-scale convection

involving shallow asthenosphere. In addition, the sublithospheric low-velocity (low-

viscosity) channel is observed to vary in thickness, depth, and minimum velocity. The

strong lateral heterogeneity of Vs can be attributed to the combined effects of temperature

variations, lithospheric thickness differences, compositional change, partial melt and

presence of water content in the mantle minerals. The complexity of the modern crustal,

upper mantle structures are strongly correlated with the Cenozoic lithospheric

modification related to the Farallon slab subduction, foundering, and the major the

extension, rifting and the associated magmatism.

The final shear velocity model agrees well with numerous previous regional or

continental-scale seismological investigations, such as the PACE, the CP-GB, the Deep

Probe, the CD-ROM, the RMF, as well as the more recent La RISTRA projects. It also

63

shows a good agreement with the new body-wave traveltime tomography models

(Schmandt and Humphreys, 2010) and the new PdS and SdP receiver function results

(Levander and Miller, 2010) using different frequency bands of the US Array data or

different imaging methods. The LAB depths inferred as the centers of the negative

velocity gradients from this Rayleigh wave tomography model have strong correlation

with the negative events in the PdS receiver function volumes.

The surface wave tomography model also successfully images, for the first time, high

velocity anomalies in the same location where lithospheric delamination has been

suggested based on receiver function imaging (Levander and Miller, 2010) to explain the

complexity of the crust/mantle boundary. This delamination event may be related to the

recent uplift in the Monument, where could be triggered by the magmatic underplating

and strong localized seismic attenuation to the west of the Navajo volcanic field.

The development of this high-resolution shear velocity model in this thesis can be used to

numerically study the geodynamic interactions between the lithosphere and

asthenosphere in regional scale. The successful implementation of the modified two-

plane wave technique makes it possible to apply the combined the surface waves and

ambient noise tomography to the new available Flexible Array Mendocino Experiment

(FAME) dataset to reveal the complicated upper mantle in the Mendocino Triple Junction

(MTJ) region. The dispersion curves in the Rayleigh wave tomography also provide the

surface data for the joint inversion with the receiver functions to examine the upper

mantle of the United States covered by the US Array stations.

64

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