AN INVESTIGATION OF LITHOSPHERIC STRUCTURE AND
EVOLUTION IN CONVERGENT OROGENIC SYSTEMS
USING SEISMIC RECEIVER FUNCTIONS AND SURFACE
WAVE ANALYSIS
by
Joshua A. Calkins
________________________
A Dissertation Submitted to the Faculty of the
DEPARTMENT OF GEOSCIENCES
In Partial Fulfillment of the Requirements
For the Degree of
DOCTOR OF PHILOSOPHY
In the Graduate College
THE UNIVERSITY OF ARIZONA
August 2008
2
THE UNIVERSITY OF ARIZONA
GRADUATE COLLEGE
As members of the Dissertation Committee, we certify that we have read the dissertation
prepared by Joshua A. Calkins
entitled “An Investigation of Lithospheric Structure and Evolution in Convergent
Orogenic Systems using Seismic Receiver Functions and Surface Wave Analysis.”
and recommend that it be accepted as fulfilling the dissertation requirement for the
Degree of Doctor of Philosophy.
_______________________________________________________________________
Date: July 31, 2008
George Zandt
_______________________________________________________________________
Date: July 31, 2008
Susan Beck
_______________________________________________________________________
Date: July 31, 2008
Mihai Ducea
_______________________________________________________________________
Date: July 31, 2008
Roy Johnson
_______________________________________________________________________
Date: July 31, 2008
Randy Richardson
Final approval and acceptance of this dissertation is contingent upon the candidate’s
submission of the final copies of the dissertation to the Graduate College.
I hereby certify that I have read this dissertation prepared under my direction and
recommend that it be accepted as fulfilling the dissertation requirement.
________________________________________________ Date: July 31, 2008
Dissertation Director: George Zandt
________________________________________________ Date: July 31, 2008
Committee Co-Chair: Susan Beck
3
STATEMENT BY AUTHOR
This dissertation has been submitted in partial fulfillment of requirements for an
advanced degree at the University of Arizona and is deposited in the University Library
to be made available to borrowers under rules of the Library.
Brief quotations from this dissertation are allowable without special permission, provided
that accurate acknowledgment of source is made. Requests for permission for extended
quotation from or reproduction of this manuscript in whole or in part may be granted by
the author.
SIGNED: Josh A. Calkins
4
TABLE OF CONTENTS
ABSTRACT ............................................................................................................... 7
INTRODUCTION .................................................................................................. 9
PRESENT STUDY............................................................................................... 20
REFERENCES ....................................................................................................... 26
APPENDIX A: CRUSTAL IMAGES FROM SAN JUAN,
ARGENTINA, OBTAINED USING HIGH FREQUENCY
LOCAL EVENT RECEIVER FUNCTIONS ...................................... 30
Permission to reprint from copyright holder ................................................ 30
Abstract .................................................................................................................... 33
Introduction ............................................................................................................ 33
Data and Methods ................................................................................................. 35
Results ...................................................................................................................... 37
Implications and Conclusions ........................................................................... 40
Acknowledgements ............................................................................................... 41
References ............................................................................................................... 41
Figures ..................................................................................................................... 43
APPENDIX B: CHARACTERIZATION OF THE CRUST OF
THE COAST MOUNTAINS BATHOLITH, BRITISH
COLUMBIA, FROM P TO S CONVERTED SEISMIC
WAVES AND PETROLOGIC MODELING ..................................... 46
Abstract .................................................................................................................... 46
Introduction ............................................................................................................ 47
Geologic Setting..................................................................................................... 49
Methods .................................................................................................................... 51
Receiver functions and Vp/Vs estimates ..................................................... 51
Petrologic and Forward Modeling................................................................ 53
Results ...................................................................................................................... 57
South Line .......................................................................................................... 58
5
TABLE OF CONTENTS - CONTINUED
South Line Seismic Results............................................................................. 58
South Line Forward Modeling Results ....................................................... 60
North Line .......................................................................................................... 63
North Line Seismic Results ............................................................................ 63
North Line Forward Modeling Results ....................................................... 65
Discussion ............................................................................................................... 67
Conclusions............................................................................................................. 70
Acknowledgements ............................................................................................... 71
References ............................................................................................................... 72
Figures ..................................................................................................................... 78
Tables ........................................................................................................................ 86
Supplementary Materials .................................................................................... 86
Explanation of petrologic modeling parameters ....................................... 86
Supplementary Figures ................................................................................... 88
APPENDIX C: LITHOSPHERIC SHEAR WAVE
VELOCITY STRUCTURE OF THE CENTRAL COAST
MOUNTAINS BATHOLITH, BRITISH COLUMBIA ............... 93
Abstract .................................................................................................................... 93
Introduction and regional tectonic setting .................................................... 94
Data and Methods ................................................................................................. 97
Results and Discussion ...................................................................................... 100
Conclusions........................................................................................................... 106
Acknowledgements ............................................................................................. 107
References ............................................................................................................. 108
Figures ................................................................................................................... 112
APPENDIX D: LITHOSPHERIC STRUCTURE OF THE
PAMPEAN FLAT SLAB REGION OF CHILE AND
ARGENTINA FROM ANALYSIS OF RAYLEIGH WAVE
PROPAGATION................................................................................................. 122
Abstract .................................................................................................................. 122
Introduction and Regional Geologic Setting .............................................. 123
Data and methodology ....................................................................................... 125
Results and Discussion ...................................................................................... 127
6
TABLE OF CONTENTS - CONTINUED
Conclusions........................................................................................................... 129
Acknowledgements ............................................................................................. 130
References ............................................................................................................. 130
Figures ................................................................................................................... 133
7
ABSTRACT
Whether by accretion, magmatic addition, or refinement of more mafic
lithologies, continental arcs are likely zones for the creation of “average” continental
crust with intermediate silica content. This dissertation contains the results of broadband
seismic studies carried out in two field areas, an active subduction zone and the remnants
of an extinct arc, with the aim of understanding lithospheric evolution at convergent
margins. The analytical techniques of receiver function calculation and surface wave
tomography are applied to data sets collected above the Andean subduction zone in Chile
and western Argentina and in the Coast Mountains Batholith of central British Columbia.
We present the first in-depth comparison of receiver functions calculated using the high
frequencies available in records of intermediate-depth local earthquakes with those
calculated from the lower frequency data in records of larger teleseismic events. The
comparison reveals that the lower crust beneath the Western Sierras Pampeanas contains
a gradational velocity increase over ~20km above a small velocity step at the Moho.
Surface wave tomography confirms the existence of an unusually high velocity anomaly
in the mantle above the slab and yields estimates of slab thickness on the order of 50 km.
To the south of the flat slab region, we see evidence of active mantle wedge convection
above the steep slab, but no evidence of the lithosphere-asthenosphere boundary beneath
the subducting Nazca plate. In the Coast Mountains Batholith (CMB), receiver functions
image a bright, continuous Moho throughout the study region. Combined with petrologic
modeling, the receiver function data point toward convective removal of any ultramafic
8
root that formed beneath the CMB. Low absolute shear wave velocities in the upper
mantle resolved via surface wave analysis strengthen the case for root removal beneath
the eastern section of the CMB. On the far western edge of the CMB, we find evidence
of a partially reformed lithosphere outboard of a major tectonic boundary. These
observations shed light on the distillation of felsic to intermediate continental crust from
more mafic primary magmas in active subduction zones and the eventual return of the
complementary ultramafic residuals to the convecting mantle.
9
INTRODUCTION
Continental arcs, which are characterized by linear, trench-parallel, belts
of volcanoes on the overriding plate at convergent margins, have long been recognized as
probable zones of creation and/or refinement of new continental crust (e.g. Anderson,
1982; Kay and Kay, 1991, Rudnick, 1995; Hollister and Andronicos, 2006). The
processes by which new crust is extracted from the mantle, and pre-existing crust is
chemically and physically altered and redistributed, however, are not entirely understood.
Geologic and geophysical evidence from active and extinct arcs indicates that complex
interactions between the mantle, crust, and associated fluids, can thicken the continental
crust and mantle lithosphere in a variety of ways, including via magmatic additions from
the mantle, pure or simple shear shortening, conversion of mantle materials to rocks with
more crust-like characteristics, and tectonic erosion (Giese et al., 1999). The welldocumented processes of incorporation of pre-existing crustal constituents in arc magmas
(e.g. Patchett and Gehrels, 1998) and the eventual return of fractionated mafic residual
matter to the mantle (e.g. Kay and Kay, 1991; Ducea and Saleeby, 1998a; Zandt, 2004)
complicate the picture of continental growth and felsification at convergent margins.
Seismologic and petrologic studies of the active Andean arc and the extinct Sierran arc
and Coast Mountains Batholith of British Columbia provide insight into the two-way
exchange of materials between the crust and upper mantle throughout the lifespan of a
continental arc.
10
A variety of seismologic techniques are commonly applied to investigate the
elastic properties of the crust and upper mantle, including reflection and refraction studies
to identify the depth to reflecting interfaces and determine velocity structure (e.g. Brown
and McClelland, 2000; ANCORP working group, 2003), inversion of P and S wave
arrival times to locate earthquake hypocenters (e.g. Anderson et al., 2007) and constrain
velocity structure and Vp/Vs ratios (e.g. Graeber and Asch, 1999; Wagner et al., 2005,
2006), and receiver function analysis of teleseismic arrivals to locate impedance contrasts
(e.g. Yuan et al., 2000, Gilbert et al., 2006). These studies alone provide information
only about the elastic properties of the subsurface, however, and no direct information
about lithology, temperature, or hydration. Laboratory and theoretical determinations of
the seismic velocities and Vp/Vs ratios of certain rock types as a function of pressure and
(less frequently) temperature, allow seismic results to be interpreted in terms of density,
mineralogic and water content, and temperature at depth (e.g. Christensen and Mooney,
1995; Christensen, 1996; Hacker et al., 2003). Other geophysical measurements, such as
gravity anomalies, heat flow, and magnetotelluric data, can be combined with
seismologic observations and numerical models (e.g. Sobolev et al., 2006) to further
constrain the possible rock types in the subsurface. Finally, xenoliths of lower crustal
and upper mantle rocks provide the only direct sampling of the lithologies at depth (e.g.
Ducea, 2002) .
Christensen and Mooney (1995; see also Rudnick and Fountain, 1995) used
compiled refraction and laboratory data to calculate global average crustal seismic
velocities, densities, and rock types, and generalized their results based on tectonic
11
localities. Continental crust can be modeled as consisting of three distinct layers,
although transitions between these layers are thought to be largely gradational. Density,
P- and S-wave velocities, and metamorphic grade are all found to increase with depth
regardless of age or location with respect to plate boundaries. Variations in overall
crustal thickness and the thickness of individual layers, however, correlate well with
tectonic setting. Compared to global averages for continental crust, continental arc crust
is found to have slightly elevated P-wave velocities in the upper 20 km and average to
fast velocities in the deeper ranges (see also Rudnick and Fountain, 1995). Christensen
and Mooney conclude that arc-type crust must therefore be altered by the intrusion of
silicic material or inclusion of metasedimentary rocks if it is later converted to average
continental crust. Of particular note are the high seismic velocity (~7.9 km/sec) of
eclogite, a metamorphic rock that can be produced in the pressure and temperature range
of deeply buried lower crust, and the low velocity of serpentinite (~5.3 km/sec), which
can form in the upper mantle via hydration of relatively cold peridotite (e.g. Bostock et
al., 2002). Several workers (e.g. Geise et al., 1999) invoke the high velocity and density
of eclogite to explain the difficulty in imaging the Moho beneath arcs. Furthermore, the
layered, sub-horizontal reflectors often imaged in the lower crust (e.g. Rudnick and
Fountain, 1995) could result from serpentinized mantle rocks interleaved with mafic
lithologies (Geise et al., 1999).
Several seismic and interdisciplinary investigations of the Andean arc in the last
two decades have revealed an extremely complex structure in the crust and upper mantle
(e.g. Myers et al., 1998; Beck and Zandt, 2002; Asch et al., 2006). Much of this
12
complexity is thought to result from the fact that the arc has migrated towards the
continent since the mid-Jurassic, such that the crust that currently lies in the forearc
region has occupied, at earlier times, both the arc and the backarc (Graeber and Asch,
1999). Currently, the active arc resides in the Western Cordillera, which is bordered on
the east by the high (up to 4000m) Puna and Altiplano plateaus. The surface of the
subducting Nazca plate has been located by receiver functions, active source seismology,
and earthquake locations, at approximately 110-130 km beneath the Western Cordillera
(e.g. Yuan et al., 2000). Although seismic characteristics vary along strike in the Andean
crust and images obtained using different seismic methods are not entirely consistent,
several seismic velocity anomalies and the nature of the Moho are sufficiently similar
between studies to be accepted with some degree of confidence.
A 10-20 km thick zone of significantly lowered P- and S-wave velocities, high
Vp/Vs ratio, and high electrical conductivity has been identified at depths of 20-40 km
beneath the Altiplano and Puna Plateaus and has been interpreted as an area of ongoing
metamorphism and partial melting (Geise at al, 1999; Yuan et al., 2000; ANCORP, 2003,
Haberland et al., 2003). Similarly, low velocities and high Vp/Vs are observed beneath
the active arc to depths of at least 50 km and possibly extending into the upper mantle,
indicative of ascending partial melt (Graeber and Asch, 1999). Conversely, the lower
crust to the west of the arc is thought to be rigid, and thus relatively cold, to depths of at
least 40 km based on observations of deep crustal seismicity (Graeber and Asch, 1999).
Directly above the Wadati-Benioff zone at depths of 70-120 km is a zone characterized
by high Vp/Vs, indicating dehydration of the downgoing slab. This is thought to be a
13
result of the blueschist → eclogite phase transition, which releases free water into the
overlying mantle. In the same depth range, velocities on the top of the slab are high,
consistent with the presence of eclogite. Dehydration of the slab also contributes to
upper slab seismicity, via increased pore pressure, and partial melting of the
asthenospheric wedge (Graeber and Asch, 1999).
Seismic detection of the continental Moho beneath the Andes has proven difficult,
particularly beneath the active arc. Seismic reflection studies have been largely
unsuccessful in imaging the Moho, whereas receiver functions show a broad positive
deflection at depths of 60-75 km (Heit et al., 2008; Gilbert et al., 2006; ANCORP, 2003;
Yuan et al., 2000). This distinction reveals that the Moho beneath the arc is likely to be
transitional rather than sharp and may be a zone of underplating of mantle-derived
magmas, hydration of the mantle, or delamination of over-thickened crust (ANCORP,
2003). Tomographic studies have failed to clearly define the Moho, but place it roughly
at 65-70 km below the arc, where Vp goes above 7.6 km/sec (Graeber and Asch, 1999).
Yuan et al. (2000) observed a strong distinction in the depth to the Moho beneath the
Altiplano (75 km) and the Puna plateau (50 km). In the latter case, the discrepancy
between very high topography and relatively thin crust is interpreted by several workers
to be indicative of lithospheric thinning and heating via delamination and inflow of hot
asthenosphere (e.g. Kay et al., 1994; Beck and Zandt, 2002; Schurr et al., 2006).
The above observations reflect a complex history of interaction between the two
lithospheric plates and intervening mantle beneath the western margin of South America
since the Jurassic. Crustal thickening appears to have been accomplished by a variety of
14
mechanisms, as summarized by Geise et al. (1999). True thickening of crustal materials
via tectonic shortening and magmatic input from the mantle can account for
approximately 60% of the Andean crustal depth, but another 15-20% of what appears to
be thickening may in fact be “pseudo-crustal thickening” (Geise et al., 1999). This term
refers to serpentinization of the mantle wedge beneath the arc and forearc via fluid
fluxing such that the uppermost mantle rocks have crust-like seismic velocities and
density. The geophysical Moho, in this case, no longer corresponds with the geochemical
Moho, where upper-mantle peridotites abut lower-crustal mafic granulites. Eclogitization
of the lower crust further complicates this region by giving crustal rocks a distinctly
mantle-like seismic signature. The ANCORP working group (2003) estimates
temperatures at the Moho beneath the Altiplano to be 800-1000°C, well above the
estimated 750°C at which substantial crustal melting occurs (Lewis et al., 1992). Any
partial melting will fractionate the more felsic and heat producing elements into the
magma and leave behind a depleted, ultramafic residue that is likely to be gravitationally
unstable and may founder into the mantle (e.g. Kay and Kay, 1991) Thus, the crust and
mantle beneath the arc are involved in a complex exchange of solids and fluids that,
through several stages of hydration, partial melting, and delamination, results in
distillation of the more felsic components into the crust and return of more mafic
constituencies to the mantle. In this zone of intense interaction between the upper mantle
and crust, the Moho is not as well defined as in stable continental interiors or extensional
regions such as the Basin and Range region of western North America.
15
Studies of relict arcs indicate that their evolution does not stagnate after the
cessation of subduction. The western portion of North America, as the overriding plate in
a convergent margin with the Farallon plate during the later Mesozoic and early
Cenozoic, contains the remnants of many extinct magmatic arcs including the Peninsular
Ranges, the Sierra Nevada, the Idaho Batholith, and the Coast Mountains Batholith. The
locations and ages of these remnant arcs have been used to track the dip angle of the
Farallon slab as well as the termination of subduction along the continental margin
through time. Seismic investigation of these batholiths combined with petrologic studies
of surface exposures and xenoliths indicate that these intrusive bodies are not
homogeneous masses of primary mantle-derived melts, but rather are silica-enriched in
their upper portions and, in some cases, increasingly mafic and layered at depth (e.g.
Brown and McClelland, 2000).
The most significant difference between the active Andean arc and the preserved
arc batholiths in North America is the disparate crustal thickness. Whereas the subAndean crust is up to 70 km thick, Moho depths beneath preserved batholiths are much
thinner. Zandt et al. (2004) used receiver functions to place the Moho at 35-40 km
beneath the Sierras, and workers have estimated crustal thickness at various places in the
Coast Mountains Batholith to be between 31 and 36 km (Rudnick and Fountain, 1995;
Morozov et al., 1998, Calkins et al., 2008). If the thickness of the Andean arc is taken to
be characteristic of an active arc, some process must occur after the cessation of arc
magmatism that is capable of thinning the arc to roughly half its original dimensions.
Bulk chemical analysis of batholith rocks indicates that their silica content is higher than
16
would be expected of mantle derived melts, which generally have the chemical
composition of basalt regardless of tectonic setting. Furthermore, REE patterns in
batholith rocks suggest an episode of melting within either the mid-crustal plagioclase
stability field (e.g. Gromet and Silver, 1987) or the deep-crustal garnet stability field (e.g.
Ducea and Saleeby, 1998b). When combined, these lines of evidence point towards a
second stage of fractionation within the crust to produce the granitic batholiths in the
upper crust. This second stage of melting must also leave behind a mafic to ultramafic
residue (e.g. Conrad and Kay, 1984; Kay and Kay, 1991; Ducea, 2002), the fate of which
is controversial but key to understanding exchange of materials between the crust and
mantle lithosphere.
Based on simple correlations of crustal thickness and age, Christensen and
Mooney (1995) determine that thick crust that originates in arcs or adjacent orogens
cannot remain thick indefinitely. This argument is based on the observation that all crust
thicker than 55 km is Cenozoic in age, whereas Archean and Proterozoic crust averages
41 km thickness. The authors offer three models for crustal thinning, which are not
necessarily exclusive of one another: delamination; uplift and erosion; and mid-to-lower
crustal flow. In the case of the southern Sierra Nevada arc, there is significant evidence
pointing towards delamination of the ultramafic root in the form of discontinuous Moho
signals (Gilbert et al. 2007; Zandt et al., 2004) and changes in xenolith chemistry
(Saleeby et al., 2003). Seismic data from the northern Coast Mountains Batholith
(Morozov et al., 1998) suggests that the dominant mechanism there has been uplift and
erosion. Rocks exposed at the surface crystallized at approximately 10-15 km depth and
17
the entire crustal column is faster than average crust. These observations have been
interpreted as indicating that the current crust represents only the lower two thirds of a
normal crustal column that has been uplifted and deeply eroded (Hollister and
Andronicos, 2006; Morozov et al., 1998). This interpretation does not preclude the
convective removal of an ultramafic root, however, as isostatic uplift is expected to
follow delamination of dense lower-crustal material. As for mid-to-lower crustal flow,
the Andean Altiplano low velocity zone (e.g. Beck and Zandt, 2002; Haberland et al.,
2003) and similar features observed in Tibet (e.g. Ozacar and Zandt, 2004; Shapiro et al.,
2004) may provide evidence for this type of thinning even as the arc is being built. These
zones of partial melt in the mid crust are thought to decouple upper-crustal brittle
deformation and ductile shearing in the lower crust (Yuan et al., 2000).
One of the most enigmatic features of arcs is the fact that they appear to create
large volumes of batholithic material during so-called flare-up events, which are
separated by long periods of relatively little activity (e.g. Armstrong, 1988). Current
magmatism rates at active continental arcs are orders of magnitude lower than calculated
flare-up rates (Ducea and Barton, 2007). Identification of a distinctly lithospheric or
asthenospheric isotopic signature in melts produced during flare-up magmatism can
differentiate between lithospheric extension or shortening, both of which have been
proposed as triggers for increased arc activity. Indeed, recent isotopic work on intrusive
rocks of the North American coastal batholiths have found strong evidence linking flareups with episodes of crustal thickening (e.g. Ducea and Barton, 2007; Girardi et al.,
2007). Continued petrologic examination of batholith rocks and xenoliths placed in the
18
context of seismic indications of deep crustal composition is key to understanding the
mechanisms involved with arc magmatism.
In this dissertation, I present the results of work aimed at furthering our collective
understanding of the processes of crustal addition and refinement in active and relict arcs.
The work addresses data collected in two field areas whose gross morphology and
tectonics are dominated by long term interaction between oceanic plates of the Pacific
basin and the overriding North and South American continental plates. The first area is
centered on a section of the South American Cordillera where the subducting Nazca plate
assumes an unusual horizontal angle of subduction near 100km depth before descending
into the deeper mantle. In the second field area, focusing on the extinct arc remnants that
make up the Coast Mountains Batholith (CMB), we have no direct evidence of the
current petrologic constituents of the lower crust as tectonic exhumation has exposed
rock only to mid-crustal levels and no deep crustal xenoliths have been identified in the
field. Fortunately, we have some insight into possible composition via outcrops and
xenoliths of lower crustal rocks from the Sierra Nevada, a well-studied batholith
analogous to the CMB in both age and tectonic origins.
Seismic data from both field areas are analyzed using two robust and
complementary research techniques. The first is the well-established technique of
receiver function (RF) calculation, which identifies discontinuities in seismic velocity at
interfaces beneath a recording station. In addition to traditional RF modeling and
interpretation, I have performed a comparison of low and high frequency RFs to improve
resolution at near-Moho depths (Appendix A) and an interdisciplinary study combining
19
RF forward modeling with petrologic modeling constrained by rock samples collected in
the study area (Appendix B). The second analysis technique is that of measuring the
dispersion of surface waves traversing an array to calculate phase velocities and absolute
shear velocities as a function of depth (Appendices C and D). We use a recently
developed technique that models the observed wave field as the sum of two interfering
plane waves (Forsyth and Li, 2005), thereby decreasing the influence of wavefront
multipathing outside the area of interest.
20
PRESENT STUDY
The methods, results, and conclusions of this study are presented in the papers appended
to this dissertation. The following is a summary of the most important findings in this
document.
The topography of Western Argentina near 32°S, where deformation associated
with ongoing subduction of the Nazca plate extends nearly 800 km inboard of the trench,
hints at the unusual horizontal attitude of the plate descending beneath the continent. The
Chile Argentina Geophysical Experiment (CHARGE) targeted the flat portion of the
subducting slab with a deployment of 22 broadband seismometers traversing the Andes in
Chile and Western Argentina. Traditional receiver functions, calculated from recordings
of teleseismic events (occurring 30°-90° from the station), found only faint signals from
the Moho beneath a station located on the flanks of Sierra Pie de Palo, near the city of
San Juan (Gilbert et al., 2006). Motivated by this unusual result and the presence of a
dense cluster of seismicity originating near 100 km depth directly beneath the station, we
undertook a receiver function study comparing results calculated using the intermediate
depth local events and the teleseismic earthquakes (appendix A). We found that the local
events yielded a signal from the Moho that we estimated to originate at 52 km,
approximately 4 km deeper than suggested by a similar calculation made using
teleseisms. Forward modeling revealed that the higher frequency content of the local
events was sensitive to a small, but sharp velocity contrast at 52 km. Conversely, the
21
dominant lower frequency energy contained in the teleseismic recordings was sensitive to
a lower crustal gradient that accounted for the majority of the crust-mantle velocity
increase over a range of ~20 km above the Moho. This result supports the hypothesis of
other workers that eclogite, a dense, seismically fast metamorphic rock, may be a major
component of the lower crust beneath the western Sierras Pampeanas.
Body wave tomography resulting from the CHARGE project also revealed
anomalous structure in the upper mantle above the flat portion of the slab and in the
region where the slab transitions to a more normal dip south of 32° (Wagner et al., 2005,
2006). Due to the highly unusual upper mantle velocities (extremely low Vp/Vs and high
Vs) and the difficulty in interpreting these findings in a petrologic sense, we look to the
measurement of surface wave dispersion (appendix D) as a means of validating these
results and expanding seismic velocity estimates beyond the region that is resolvable
using P and S wave travel times from local earthquakes. We use a recently developed
method of Rayleigh wave tomography that models the observed wave field as the sum of
two interfering planes waves (Forsyth and Li, 2005). Among the many benefits of this
technique are the reduced influence of surface wave multi-pathing or scattering outside
the area of interest, an approach to the inversion that automatically down-weights events
that are not well fit by the two-plane wave assumption, and the ability to resolve shear
wave velocities outside the region covered by the array. Our results provide an important
independent confirmation of the high Vs anomaly above the flattest portion of the
subducting plate. The surface wave tomography also reveals a prominent low velocity
zone approximately 50km below the top of the flat slab (located by patterns of seismicity
22
by Anderson et al., 2007) that we interpret as the lithosphere-asthenosphere boundary at
the base of the subducting Nazca plate. Further south, where the plate subducts at a more
typical 30°, we see low velocities in the uppermost mantle consistent with convection of
asthenospheric mantle in the wedge between the two lithospheric plates.
From these studies and others associated with the CHARGE project, a picture
emerges of a complex, heterogeneous, and dynamic environment in the upper mantle and
crust on the continental side of an active subduction zone. The Moho discontinuity
defines an irregular surface, and the distinction between the crust and the mantle is, in
places, seismically ambiguous. Crustal thickness does not always correspond with
topography, implying dynamic support or anchoring in the upper mantle, but the crust
beneath the arc is generally much thicker than the global continental average of ~41 km
(Christensen and Mooney, 1995). If continental arcs are the modern factories where
continental lithosphere is created, then some post-subduction processes must be at work
to transform arc remnants towards a thinner, seismically slower crust overlying faster
mantle lithosphere with a thickness on the order of 10’s to 100 km. To explore the
possible evolution of a continental arc system towards more typical continental
lithosphere, we look to a relict subduction zone along the North America/Pacific margin.
The Coast Mountains Batholith (CMB) of British Columbia, one of the largest
exposed granitoid batholiths in the world, is an ideal location to study the deep roots of
an extinct continental arc system that was tectonically similar to the modern Andean arc
throughout much of the Mesozoic and early Cenozoic. The volcanic arc has been
removed via uplift and erosion, but remnants of its upper to mid crustal magmatic system
23
are now exposed as plutonic rocks exhumed from depths of 5 to 25 km. The multidisciplinary “Batholiths” Continental Dynamics project focused on resolving processes of
lithospheric evolution at convergent continental margins by investigating magmatic
addition and crustal recycling within the CMB. In particular, the project targeted the
distillation of igneous rocks within the crust from the basaltic compositions extracted
from the mantle to the intermediate compositions that make up “average” continental
crust. A necessary byproduct of this fractionation is a mafic to ultramafic root that may
be dense enough to sink back into the convecting mantle. We search for evidence of an
ultramafic lithospheric root beneath the CMB using receiver functions combined with
petrologic modeling (appendix B) and surface wave tomography (appendix C).
Our receiver function study is novel in combining seismic data with geochemical
analysis of rock samples collected in the same area to construct and test models of crustal
composition. The seismic data contain remarkably consistent crustal reverberations
which allow for the calculation of vertically-averaged Vp/Vs ratios in addition to
resolving the seismic discontinuity structure across the two transects of the Batholiths
array. We find unusually high Vp/Vs values and thin crust in the western portion of the
study region. In the vicinity of an ancient terrane boundary and the Coast Shear Zone, a
major near-vertical ductile shear zone, we image an increase in crustal thickness from
~25 to ~32 km from west to east coincident with a drop in Vp/Vs to values more typical
of continental crust. We construct 1D crustal models by predicting seismic velocities
over a range of pressures and temperatures for a suite of rock types collected at the
surface. Forward modeling with synthetic seismic data indicates that the entire crust of
24
the CMB can be well-matched using only the deeper equivalents of the sampled rock
types, and we find no evidence of the high seismic velocities and diminished Moho signal
that would be indicative of a sub-batholith garnet-pyroxenite root.
Surface wave tomography, again achieved using the two-plane wave
approximation, also suggests that any ultramafic root that formed beneath the CMB
within the project study area has been largely removed. Upper mantle velocities beneath
the majority of both transects are low, an observation that, combined with high heat flow,
thin crust, and high elevations, points towards the presence of warm asthenospheric
mantle in close proximity to the base of the crust. Absolute shear velocities also bolster
the hypothesis that a major tectonic boundary exists in the vicinity of the Coast Shear
Zone. West of this boundary, there are no plutons younger than ~85 Myr (Gehrels et al.,
2008) and fast lithospheric mantle exists down to ~70 km. To the east, where
subduction-related magmatism persisted into the Eocene, the lithosphere/asthenosphere
boundary is very near the base of the crust. Convection around the edge of the Juan de
Fuca/Explorer slab, which continues to subduct to the south of the study area, likely
contributes to quaternary magmatism and the warm upper mantle that is particularly
apparent along the southern transect.
Assuming that the continental arc that created the CMB was analogous to the
modern Andes, the stark differences between the current state of the crust and upper
mantle in the two field areas addressed in this dissertation strongly suggests that
evolution of the continental lithosphere continues after subduction ends. Compared with
the CHARGE study area, the structure beneath the central CMB is much simpler, with a
25
relatively flat and distinct Moho near 30 km and seismically slow and somewhat
homogeneous upper mantle underneath most of the study area. Foundering of the
ultramafic root after magmatism ceases seems an efficient mechanism for not only
creating a sharp velocity contrast at the base of the remaining compositionally
intermediate crust, but also for allowing the influx of hot asthenosphere that appears to
reside beneath the eastern CMB. It is also apparent that while arc magmatism and postsubduction processes may have created typical continental crust within the CMB (as
suggested by Hollister and Andronicos, 2006), the region is not underlain by typical (i.e.
strong, seismically fast) continental lithospheric mantle. Thus, if the CMB is taken to
represent a typical post-subduction continental setting, some processes other than
subduction-related magmatism and subsequent root removal must be responsible for
creation of the lithospheric mantle that forms the base of stable continental interiors.
26
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30
APPENDIX A: CRUSTAL IMAGES FROM SAN JUAN,
ARGENTINA, OBTAINED USING HIGH FREQUENCY
LOCAL EVENT RECEIVER FUNCTIONS
Josh A. Calkins, George Zandt, Hersh J. Gilbert, and Susan L. Beck Dept. of
Geosciences, University of Arizona
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Calkins J. A., G. Zandt, H. J. Gilbert, S. L. Beck (2006), Crustal images from San Juan,
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Abstract
Images of the crust and upper mantle obtained using teleseismic receiver functions
exhibit a weak P-S conversion from the Moho in Western Argentina, where a segment of
the subducting Nazca plate flattens out beneath the overriding South American plate. To
better estimate depth to the Moho and search for mid-crustal impedance contrasts, we
calculate high frequency receiver functions using 37 intermediate-depth local
earthquakes. Radial receiver functions from a station located near San Juan, Argentina,
show a strong signal from the Moho at 7 seconds corresponding to a depth near 50 km,
and conversions from interfaces within the crust at depths of ~ 20 and 35 km. The higher
frequency results obtained using local events have twice the vertical resolution of the
teleseismic results and the combined data sets indicate a gradational increase in S
velocity in the lower crust that cannot be detected in the lower frequency teleseismic data
alone.
Introduction
Teleseismic receiver functions have been used to estimate the depth to the Moho
and other seismic velocity discontinuities with great success in investigations of crust and
upper mantle structure in a variety of geologic locales (e.g. Owens et al., 1984, Gilbert et
al., 2003). Generally, the technique uses differences in arrival time between the direct P-
34
wave and P-to-S converted phases and a known or calculated velocity model to determine
the depths at which incident P-waves encountered significant S velocity contrasts beneath
a seismic station (e.g. Langston, 1977). The same technique, with slight processing
modifications, can be applied to records of local intermediate or deep events in order to
image sub-receiver structure. Whether due to convention or a lack of data from stations
situated directly above clusters of deep seismicity, however, relatively few studies have
been carried out using local event receiver functions. Notable exceptions include an
analysis of the Central Andean crust that combined receiver functions from deep local
events with those from teleseismic events (Beck and Zandt, 2002) and images of the
oceanic Moho in the Nazca slab obtained via P-S conversions from regional intermediatedepth events (Bock et al., 2000). Our goals in this paper are to present the first detailed
comparison of local and teleseismic receiver functions and to add constraints to the
crustal structure in an area distinguished by anomalously thick crust and low elevations.
The city of San Juan, Argentina, lies in the region where the Andean Precordillera
transitions eastward into the Sierras Pampeanas. The accreted terranes in the vicinity of
32° south and 68° west overlie a section of the descending Nazca slab, as outlined by the
Wadati-Benioff seismicity, that flattens at a depth of 100-125 km before resuming a more
normal angle of subduction to the east (Cahill and Isacks, 1992; Figure 1). The Juan
Fernandez Ridge, a zone of over-thickened oceanic crust that has been cited as a possible
cause of this flat subduction, has been traced eastward under the South American
continent in the region of San Juan based on magnetic anomalies (Yanez et al., 2001).
Many workers (e.g. Anderson, 2005; Cahill and Isacks, 1992; Wagner et al., 2005) have
35
observed a relatively dense cluster of small, intermediate depth (80-120 km) earthquakes
in the vicinity of the flat slab that follows the inferred path of the Juan Fernandez ridge
beneath San Juan.
Data and Methods
As part of the Chile Argentina Geophysical Experiment (CHARGE) project, two
arrays of PASSCAL broadband seismometers were deployed in east-west trending lines
across the Andean Cordillera between 2000 and 2002. Station JUAN, one of several
placed between the transects, was situated nearly directly above the cluster of
intermediate-depth seismicity near the city of San Juan (Figure 1). Signals from these
events followed steep ray paths to station JUAN and thus provide a unique opportunity to
compare large numbers of receiver functions from local and teleseismic signals sampling
the same receiver-side structure. While other stations in the array recorded the events
used in this study, only station JUAN was close enough to the earthquake cluster that the
P-wave ray paths were steep enough to result in a sufficient number of high quality
records to produce well-sampled moveout plots and depth stacks.
Roughly one hundred events, with depths near 100 km and epicenters within one
degree of station JUAN, were identified and relocated using the multiple event relocation
code, GMEL, by Anderson (2005). Of these hundred, about half were eliminated from
this study due to very low amplitude or noisy signals or lack of an impulsive P-wave
arrival. Records for the remaining events, recorded at forty samples per second, were cut
ten seconds before and after the direct P-wave arrival in order to retain sufficient data for
36
a robust deconvolution and to preclude interference from the direct S-wave, arriving
roughly eleven to sixteen seconds after the P-wave. The resulting unfiltered traces were
then used to calculate radial and tangential receiver functions using an iterative pulsestripping time domain deconvolution code (Ligorría and Ammon, 1999). This approach
uses a series of Gaussian pulses in deconvolving the input signals, and the width of each
pulse controls the frequency content of the resulting receiver function. In order to take
advantage of the high frequency content of the local events, receiver functions were
calculated using two Gaussian pulse widths corresponding to low pass filters with corner
frequencies of 2.5 and 5 Hz. In terms of wavelength, which controls the vertical
resolution recoverable in the receiver functions, a corner of 2.5 Hz corresponds to crustal
wavelengths larger than approximately 2.7 km and a corner of 5 Hz corresponds to
crustal wavelengths greater than roughly 1.4 km. Receiver functions were also calculated
with higher and lower frequency content, but the results were excessively noisy in the
former case and smoothed to the point of obscuring details in the latter. Of the resulting
traces, thirty-seven were deemed acceptable based on the maximum amplitude occurring
at the beginning of the trace, a minimum of 70% variance reduction in the deconvolution,
and an absence of long-period trends or harmonic oscillations.
In order to minimize the effects of noise and aid in visually enhancing coherent
arrivals, radial receiver functions were stacked and plotted as a function of ray parameter
(Figure 2). The observed moveout of specific phases was then compared with predicted
moveout curves to distinguish between converted arrivals, reverberations, and noise. In
calculating predicted moveout, we use a single-layered crustal velocity model (Vp = 6.4
37
km/sec, Vp/Vs = 1.84), based on waveform modeling of arc and back arc crustal
earthquakes (Alvarado et al., 2005) and analysis of teleseismic receiver functions at
nearby CHARGE stations (Gilbert et al., 2006). Mantle velocities of Vp = 8.15 km/sec
and Vp/Vs = 1.8 are based on analysis of Pn apparent phase velocities from regional
events (Fromm et al., 2004). Gilbert et al. (2006) suggest that the lower crust in the
region may be composed of partially eclogitized material, and hence possess higher
seismic velocities. If a high velocity lower crust is indeed present, the depth to the Moho
calculated herein would be skewed towards shallower depths by our simple velocity
model. However, the established technique of using crustal multiples to constrain the
local velocity structure is not possible in this case because the high amplitude signal from
the direct S-wave drowns out the relatively low amplitude multiples. In the absence of
multiples, we cannot reliably resolve the tradeoff between depth and velocity and
therefore use the single-layered crustal velocity model described above and allow that the
base of the crust may be several km deeper than we calculate. Finally, moveout
corrections and depth conversions based on the local velocity model were applied to
create a depth stack of all thirty-seven receiver functions (Figure 3).
Results
The moveout plots (Figure 2) combined with predicted arrival times reveal
consistent mid-crustal conversions at depths of approximately 20 and 34 km, and, most
importantly, a distinct signal from the Moho at a depth near 52 km. The lower frequency
local event plot (Figure 2a) shows the strongest Moho signal and generally consistent
38
moveout for the mid-crustal arrivals. A strong negative arrival is also apparent just
below the Moho. The higher frequency moveout plot (Figure 2b) suggests complex
layering within the shallow (5-10 km), mid (20 km), and deep (30-40 km) crust. The
moveout plot for teleseismic data, calculated with a corner frequency of 1.25 Hz, (Figure
2c) exhibits very little coherent energy within the crust and a poorly resolved signal from
the Moho.
A comparison of amplitude across the power spectrum shows that, as expected,
the local event data has a significantly higher signal-to-noise ratio (~10 to 100) than the
teleseismic data in the high-frequency range, particularly between 2 and 10 Hz. In the
frequency range of 0.3 to 1 Hz, however, the local events have a signal-to-noise ratio
roughly equal to 1 whereas the teleseismic signal is an order of magnitude higher than the
noise. Depth-converted and normalized stacks of both sets of receiver functions (Figure
3) illustrate the resolution advantage gained by using higher frequency data; many
structures that appear to be single, thick layers in the teleseismic and lower frequency
local stacks (Figure 3c and a) may in fact be composed of multiple thin layers (Figure
3b). The apparent long period signal and associated broad shallow negative in the lower
frequency local stack may be a result of the low signal-to-noise ratio at low frequencies
in the local data. Attempts to filter out this signal had the undesirable effect of producing
a sharp negative side lobe after the direct P-wave arrival.
An obvious difference between the stacks is the 0.6 second delay in the arrival
time of the Moho P-S conversion in the local data relative to the teleseismic data. Both
data sets traverse the subsurface at similar radial distances from station JUAN, and
39
azimuthal analysis of all events reveals no pattern that could explain the timing
difference. Forward modeling using reflectivity synthetics reveals the advantage gained
by using both high and low frequency signals to determine sub-receiver structure. Due to
the complexity of the observed wave forms, our focus was on modeling the timing
differences in the Moho P-S conversions as a function of frequency rather than producing
synthetic receiver functions that matched all of the conversions within the crust. Crustal
velocity models consisting of a gradational increase in velocity in the lower crust (35-55
km) followed by a sharp step in velocity, yield synthetic receiver functions that match the
0.6 second difference in arrival times between the local and teleseismic receiver functions
(Figure 3d). These modeling results suggest that low frequency teleseismic signals are
sensitive to the long wavelength velocity increase centered in the lower crust, while the
high frequency signal can detect the smaller velocity jump at the bottom of the crust.
Figure 3d illustrates several models that can reproduce the observed Moho Ps time
differences, but our data can not differentiate between these possibilities. A wide variety
of models that did not include both the sharp step and the gradual increase in velocity
were also tested, but the resulting P-S conversions showed little or no time separation
across the frequency range covered by our data. Thin layering of high and low velocity
layers in the deep crust might also contribute to the observed time lag, but modeling
indicates that layering alone cannot explain the observed disparity. These findings
support the hypothesis of a partially eclogitized lower crust which Gilbert et. al. (2006)
propose to explain the weak Ps arrival in the teleseismic receiver functions and the
thicker than predicted crust below relatively modest average surface elevation.
40
Implications and Conclusions
The crustal structure revealed by this study is in good agreement with that
observed in the region by other researchers. An extended correlation of seismic
reflection profiles roughly 75 km to the south of Sierra Pie de Palo, near Cerro Salinas,
located strongly reflective horizons near 18 and 32 km (Comínguez and Ramos, 1991).
A similar study in an area north of station JUAN by Zapata (1998) identified complex
layering in the upper 10 km, an imbricated middle crust, suggestions of layering in the
38-45 km range, and a faint Moho signal near 50 km. A consistent strong reflection at
approximately 20 km depth was interpreted to be the signal from the Andean
decollement. Likewise, an investigation of crustal thickness using Pn apparent phase
velocities (Fromm et al., 2004) yielded a Moho depth in good agreement with our 52 km
estimate. While these studies investigated areas tens to hundreds of kilometers from
station JUAN, they represent potential corroboration of our results. Investigations using
S to P converted phases (Regnier et al., 1994) and gravity data (Gimenez et al., 2000)
have also located the Moho near 50 km beneath San Juan province.
In areas of dense seismicity, local earthquakes with hypocentral depths in the 100
km range can be used to generate receiver functions to image the crust and uppermost
mantle. As local event signals are much less attenuated in the high frequencies than
teleseismic signals, receiver functions sensitive to short wavelength features can be
calculated to produce images with significantly improved vertical resolution. These data,
when analyzed in conjunction with teleseismic receiver functions, can illuminate details
of the velocity profile at depth that are obscured when exclusively examining low
41
frequency data. Based on initial investigations of signal-to-noise ratios, the potential
exists to image the crust with data frequencies of up to, and possibly exceeding, 5 Hz.
For station JUAN, in the Western Sierras Pampeanas, local event receiver functions
reveal complex layering in distinct zones throughout the crust, a gradational velocity
increase in the lower crust, and a Moho depth of approximately 52km.
Acknowledgements
We are indebted to two anonymous reviewers, INPRES, the National University of San
Juan (Argentina), and the University of Chile. The National Science Foundation
supported this research under grants EAR9811870 and EAR0510966 and via support of
the PASSCAL facility of the Incorporated Research Institutions for Seismology (IRIS).
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Beck, S.L. and G. Zandt (2002) The nature of orogenic crust in the Central Andes, J.
Geophys. Res.,107(B10), 2230, doi:10.1029/2000JB000124.
Bock, G. B.Schurr, and G. Asch (2000) High-resolution image of the oceanic Moho in
the subducting Nazca plate from P−S converted waves, Geophys. Res. Lett., 27(23),
3929-2932.
Cahill, T.A., and B.L. Isacks (1992), Seismicity and shape of the subducted Nazca Plate,
J. Geophys. Res., 97, 17,503-17,529.
42
Comínguez, A.H. and V.A. Ramos (1991), La estructura profunda entre
Precordillera y Sierras Pampeanas de la Argentina: evidencias de la sísmica de
reflexión profunda. Revista Geológica de Chile, 18,(1), 3-14.
Fromm, R., G. Zandt, and S.L. Beck (2004), Crustal thickness beneath the Andes and
Sierras Pampeanas at 30° S inferred from Pn apparent phase velocities, Geophys. Res.
Lett., 31, L06625.
Gilbert, H.J., A.F. Sheehan, K.G. Dueker, and P. Molnar (2003), Receiver functions in
the western United States, with implications for upper mantle structure and dynamics,
J. Geophys. Res., 108, 1855-1858.
Gilbert, H., S. Beck., and G. Zandt (2006), Lithospheric and upper mantle structure of
Central Chile and Argentina, in press Geophys. J. Int.
Gimenez M. E., M. P. Martinez, and A. Introcaso (2000), A crustal model based mainly
on gravity data in the area between the Bermejo Basin and the Sierras de Valle Fertil,
Argentina, J.S. Am.Earth Sci., 13, 275-286.
Langston, C. A. (1977), The effect of planar dipping structure on source and receiver
response for constant ray parameter, Bull. Seismol. Soc. Am., 67, 1029– 1050.
Ligorría, J. P., and C. J. Ammon (1999), Iterative deconvolution and receiver function
estimation, B. Seismol. Soc. Am., 89, 1395-1400.
Owens, T. J., G. Zandt, and S. R. Taylor (1984), Seismic evidence for an ancient rift
beneath the Cumberland Plateau, Tennessee: A detailed analysis of broadband
teleseismic P waveforms, J. Geophys. Res., 89, 7783–7795.
Regnier, M., J. M. Chiu, R. Smalley, B. L. Isacks, and M. Araujo (1994), Crustal
thickness variation in the Andean foreland, Argentina, from converted waves, B.
Seismol. Soc. Am., 84, 1097-1111.
Wagner, L.S., S. Beck, and G. Zandt (2005), Upper mantle structure in the south central
Chilean subduction zone (30° to 36°S), J. Geophys. Res., 110,
doi:10.1029/2004JB003238.
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interpretation across the southern central Andes (32°–34°S): The role of the Juan
Fernandez Ridge in the late Tertiary evolution of the margin, J. Geophys. Res., 106,
6325–6345.
Zapata, T.R. (1998), Crustal structure of the Andean thrust front at 30° S latitude from
shallow and deep seismic reflection profiles, Argentina, J.S. Am.Earth Sci., 11, 131151.
43
Figures
Figure 1: Regional and local maps showing location of CHARGE seismic stations
(black squares on inset) and piercing points (magenta asterisks) for local events at a depth
of 50 km below station JUAN. Map is after Alvarado et al. (2005) and references
therein; contours to the top of the Nazca slab are after Cahill and Isacks (1992).
44
Figure 2: Radial receiver function moveout plots for station JUAN. Plots (A) and (B) are
composed of local data and calculated with corner frequencies of 2.5 and 5 Hz,
respectively, and plot (C) contains teleseismic traces with a corner frequency of 1.25 Hz.
Numbers on the right side of the plots denote depth to hypothetical interfaces used to
create predicted moveout curves based on the direct P-wave ray parameter (solid lines).
For all moveout calculations, Vp was set to 6.4 km/sec and Vp/Vs to 1.84. Individual
receiver functions are stacked in bins centered every 0.0026 sec/km with sharing between
adjacent bins. All receiver functions within each bin are averaged and normalized to unit
amplitude. Red colors denote positive deflections in the receiver function and blue colors
correspond to negative deflections.
45
Figure 3: Depth converted receiver function stacks for local events low-pass filtered
with corner frequencies of (A) 2.5 and (B) 5 compared with (C) teleseismic stack with
corner frequency of 1.25. Velocities used for depth conversion are the same as in Figure
2. Central blue lines indicate normalized stacks of all traces and outer black lines
indicate one sigma bounds from bootstrap resampling. Phases with positive amplitude
originate at discontinuities where lower velocity material overlies higher velocity
material. Crustal S-velocity profiles in (D) are capable of producing ~0.6 second lags
between Ps arrivals on the low and high frequency synthetic receiver functions. Times
indicated on the legend indicate the delay between Moho Ps arrivals on the teleseismic
(1.25 Hz) and local (5 Hz) synthetic receiver functions. Vp was calculated using a
constant Vp/Vs of 1.84.
46
APPENDIX B: CHARACTERIZATION OF THE CRUST OF
THE COAST MOUNTAINS BATHOLITH, BRITISH
COLUMBIA, FROM P TO S CONVERTED SEISMIC
WAVES AND PETROLOGIC MODELING
Josh A. Calkinsa,, George Zandta, James Girardia, Ken Duekerb, George E. Gehrelsa, and
Mihai N. Duceaa
a
Department of Geosciences, University of Arizona, Gould Simpson Bldg #77, 1040 E.
4th St, Tucson, Arizona, USA.
b
Department of Geology and Geophysics, Dept. 3006, 1000 University Ave., University
of Wyoming, Laramie, WY, 82071, USA
Submitted for publication to Earth and Planetary Science Letters, March 2008.
Abstract
The late Triassic to early Tertiary Coast Mountains Batholith (CMB) of British
Columbia provides an ideal locale to study the processes whereby accreted terranes and
subduction-related melts interact to form stable continental crust of intermediate to felsic
composition and complementary ultramafic residuals. Seismic measurements, combined
with calculated elastic properties of various CMB rock compositions, provide a window
into the deep-crustal lithologies that are key to understanding the process of continental
growth and evolution. We use a combination of seismic observations and petrologic
modeling to construct hypothetical crustal sections at representative locations across the
47
CMB, then test the viability of these sections via forward modeling with synthetic
seismic data. The compositions that make up our petrologic forward models are based on
calculations using the free energy minimization program Perple_X to predict mineral
assemblages at depth for the bulk compositions of exposed plutonic rocks collected in the
study area. Seismic data was collected along two transects in west-central British
Columbia: a southern line that crossed the CMB near the town of Bella Coola (~52° N),
and a northern line centered on the towns of Terrace and Kitimat (~54° N). Along both
transects, seismic receiver functions reveal high Vp/Vs ratios near the
Insular/Intermontane terrane boundary and crustal thickness increasing from ~25 km to
~32 km from west to east across the Coast Shear Zone (CSZ). On the southern line, we
observe an anomalous region of complex receiver functions and diminished Moho signals
beneath the central portion of the CMB. Our petrologic and seismic profiles show that
observed seismic data from much of the CMB can be well-matched in terms of crustal
thickness and structure, average Vp/Vs, and amplitude of the Moho converted phase,
without including ultramafic residual material in the lower crust.
Introduction
The western margin of the Canadian Cordillera has been the locus of significant
growth of the North American continent via complex tectonic interactions at an evolving
ocean/continent margin throughout the Phanerozoic (e.g., Monger, 1993). Along its
leading edge, the accreted terranes of the Insular and Intermontane assemblages hosted
the emplacement of an enormous magmatic arc for the duration of the Mesozoic and into
48
the early Tertiary. Cessation of arc magmatism around 50 Ma and subsequent
exhumation resulted in the exposure of the intermediate to felsic batholithic roots of the
arc, known collectively as the Coast Mountains Batholith (CMB) (Hollister and
Andronicos, 2000). The northern CMB has been presented as a model case for creation
of felsic continental crust via heat and magma input from the mantle combined with
fractionation of more mafic accreted island arc crust (e.g. Hollister and Andronicos,
2006). This process, however, necessarily creates a complementary mafic to ultramafic
residue (Ducea, 2001) that is not commonly observed at the base of typical sections of
continental crust (Christensen and Mooney, 1995). Studies of other Cordilleran arcs have
found strong evidence for convective removal of dense roots (e.g. Ducea and Saleeby,
1998, Zandt et. al, 2004), but the current state of the lithosphere beneath the CMB, North
America’s largest exposed continental batholith (e.g. Barker and Arth, 1990) , remains
largely unknown.
In conjunction with the multi-disciplinary Batholiths project, we collected an
average of 14 months of 3-component broadband seismic data along two roughly strikeperpendicular transects (Fig. 1) between the summer of 2005 and fall 2006. Each of the
43 sites, deployed with an inter-station spacing of ~10km, consisted of either a Guralp
CMG3T or a Streckheisen STS-2 seismometer linked to a Quanterra 330 data acquisition
system. The southern line ran from Burke Channel in the west, through the town of Bella
Coola, and onto the Chilcotin Plateau near the town of Anahim Lake. The northern
transect traversed Douglass Channel from Hartley Bay to Kitimat and then followed the
Skeena River northeast from Terrace to New Hazelton. Data quality was high across the
49
study area with the exception of the stations on the far southwestern end of the northern
line which, due to siting issues and ocean storms, had higher noise levels and were out of
service for a portion of the deployment.
Herein, we describe our use of receiver functions to image the Moho and other
crustal discontinuities and constrain crustal Vp/Vs along the two transects. Combining
the seismic data with deep crustal compositions inferred from thermodynamic modeling,
we test various structural and compositional models for the crust at a series of
representative stations across the seismic array.
Geologic Setting
The Batholiths project study area (Fig.1) is located in central western British
Columbia where subduction-related magmatism of the CMB intrudes upon and obscures
the boundary between the outboard Insular superterrane and the inboard Intermontane
superterrane (tectonostratigraphic nomenclature follows Monger, 1993). The
Intermontane group is an amalgamation of peri-cratonic arc-related units that formed
along the Laurentian margin during the Paleozoic and early Mesozoic. The oceanic and
continental slivers of the Insular group (primarily Alexander and Wrangellia terranes)
appear to have an origin in the arctic realm (Colpron et al., 2007). Although the timing
and geometry of accretion of the Intermontane and Insular superterranes to the North
American margin is a matter of vigorous debate (e.g. Cowan et al., 1997; Butler et al.,
2001), there is considerable geologic evidence supporting the consolidation of the two
superterranes by mid-Jurassic (Colpron et al., 2007), and post accretion dextral transport
50
of the combined terranes of at least 850 km along the North American margin (Gabrielse
et al., 2006, Butler et al., 2001).
The Coast Shear Zone (CSZ), a steeply east-dipping zone of intense ductile
deformation, splits the CMB along strike on a >1200 km line roughly paralleling the
suture between the two superterranes (Rusmore et al., 2001). Geologic studies indicate a
complex history of strike slip, reverse, and normal offset along the CSZ, with the last
documented motion occurring around 51 Ma (Rusmore et al., 2001, Andronicos et al.,
1999, Klepeis and Crawford, 1999). In the northern portion of the study area, the CSZ
marks the western boundary of the Central Gneiss Complex (CGC), a package of
amphibolite to granulite grade mid-crustal rocks exhumed from depths of up to 25 km by
~50 Ma (Rusmore et al., 2005, Hollister and Andronicos, 2000). The CGC is bounded on
the east by the Eastern Boundary Detachment (hereafter the EBD; also known as the
Shames mylonite zone), a more shallowly NE dipping ductile shear zone active between
67 and 52 Ma (Rusmore et al., 2005). Neither the CGC nor the EBD have been mapped
in the southern portion of the study area, where rocks now exposed at the surface have
been exhumed from depths of only 5-10 km.
Jurassic to Tertiary magmatism across the CMB consists largely of diorite to
granitic plutons with an average composition of tonalite. Age of emplacement generally
decreases from west to east across the Insular superterrane and the heart of the batholith,
but plutons within Stikine terrane appear to follow no particular age progression and
range from late Triassic to Eocene (Gehrels et al., 2007). Large pulses of magmatism
occurred at ~150 Ma, ~120-80 Ma, and ~52 Ma. Massive tonalite plutons were emplaced
51
along the CSZ from 70 to 57 Ma and appear to be concurrent with motion on the shear
zone. Magmas related to the Eocene flare-up appear only east of the CSZ and are
significantly more voluminous in the northern part of the study region than in the south.
Recent age-dating resulting from the Batholiths project seems to indicate that two
independent pluton-generating arcs were active between the Late Jurassic and the Early
Cretaceous. The western arc is separated from the eastern arc by the Gravina belt, a band
of marine sediments that was being deposited slightly west of the present location of the
CSZ until ~100 Ma. (Gehrels et al., 2007).
Methods
Receiver functions and Vp/Vs estimates
In order to image the Moho discontinuity and intra-crustal structure, we calculated
P-wave receiver functions using the iterative time domain deconvolution method of
Ligorría and Ammon (1999). Included in our data set were over 100 teleseismic P and
PP phases from earthquakes with broad azimuthal coverage (Fig. 2), magnitudes > 5.5,
and good signal to noise ratio. After cutting and bandpass filtering the signals between .3
and 30 seconds, we calculated radial receiver functions, discarding any resulting traces
with a variance reduction of less than 70%, long-period trends, or high amplitude
harmonic oscillations.
The receiver function method is limited by the inherent trade-off between depth to
an interface where a converted phase (e.g. Pds, the conversion of a P wave to an S wave
at the Moho) originates and the average velocity, principally the Vp/Vs value, of the
52
structure above the interface. By measuring the differential arrival times of the converted
phase and its crustal multiples, however, we can estimate the average Vp/Vs ratio in the
crust using the relationship:


2
V p Vs =  1 − p 2VP

(
1

)2 t t
2
 
 + 1 + p 2VP 2 


 
2
− tP
  PpPms − t Pds
Pds
(1)
(Zandt et al., 1995) where Vp and Vs are the average P and S wave velocities in the crust,
p is the ray parameter of the incoming P wave, tPds is the arrival time of the P-s converted
phase from the Moho, tP is the arrival time of the direct P wave, and tPpPms is the arrival
time of first crustal multiple of the Moho conversion (hereafter referred to as the 2P1S
phase). The Vp/Vs ratio is a particularly useful indicator of crustal composition (via its
relation to Poisson’s ratio) and constrains bulk lithology better than either compressional
or shear wave velocity alone (e.g. Behn and Kelemen, 2006; Zandt and Ammon, 2005).
The Batholiths data set displays remarkably clear Moho multiples (Fig. 3) allowing us to
make robust Vp/Vs estimates over much of the study area.
Once we have an estimate for Vp/Vs, depth to the Moho is calculated using:
−2
−2
h = t Pds − t p  VS − p 2 − V p − p 2 


(2)
(Zandt et al., 1995). Equations 1 and 2 have a small dependence on absolute Vp, which
we have assumed to be constant at 6.2 km/sec throughout the study area based on the
ACCRETE project refraction line, approximately 150 km to the north of our study area
(Morozov et al., 2003). However, Zandt et al. (1995) show that the effect of varying Vp
53
within the range of reasonable values changes Vp/Vs by no more than 0.03, well below
the standard deviation in our data set (see below).
Rather than picking arrival times of the converted phases and their multiples at
individual stations, we follow the method of Kind et al. (2002) and create common
conversion point (CCP) stacks of depth-converted receiver functions along both transects
(Fig. 3), then employ an automated picking routine to select the maximum amplitude of
the Pds and 2P1S phases at the center point of each stacking bin. Stacking has the benefit
of emphasizing coherent arrivals observed at adjacent stations and decreasing the
influence of noise and small scale variations that might obscure more significant
structures. Sharp interfaces where velocity increases with depth, such as the Moho,
produce distinct red (positive) arrivals in the stacks; blue (negative) arrivals, conversely,
indicate layers where velocity decreases with depth. We use bootstrap resampling (e.g.
DiCiccio and Efron, 1996) of the entire data set to estimate the variance in the imagedepth of the converted phases and multiples as well as the Vp/Vs values calculated from
these arrivals (see section 4.1 and Fig. 4).
Petrologic and Forward Modeling
After correcting the CCP stacks for the calculated Vp/Vs variations (Figs. 5 and
6), six representative stations across the array were selected as candidates for forward
modeling: BS01, BS07, BS22, BN07, BN19, and BN22. We chose stations so as to
elucidate the major changes in observed structure and Vp/Vs within the study area while
avoiding excessively complex receiver functions and areas where the data are poor or
observations are few (e.g. the far western end of the northern transect). One-dimensional
54
petrologic columns were generated via forward modeling to produce synthetic receiver
functions that match our observations at the six “target” stations. To construct the
columns, we used primarily data gleaned from rock samples collected within the CMB
(see Table 1) but supplemented our modeling suite, where necessary, with estimates of
lower-crustal compositions taken from the geologic literature.
In order to predict the seismic properties of probable compositions throughout the
modern crust, we use the free energy minimization program Perple_X (Connolly, 1990,
2005) and the thermodynamic database of Holland and Powell (2001) to calculate modal
abundances for representative rock types at a range of temperatures and pressures
appropriate for the active arc environment. Once we have determined the appropriate
lithologies, the method of Hacker and Abers (2004) allows us to calculate the P and S
wave velocities and densities of these predicted compositions at pressures and
temperatures appropriate for the modern CMB.
To validate the accuracy of our calculations, we compare the assemblages
predicted by Perple_X to the observed modes (determined by thin section observation) of
exhumed rocks for which the peak pressure and temperature are known (via the methods
of Hollister et al., 1987, and Blundy and Holland, 1990, respectively). In all cases, we
obtain a match between predicted and observed modes of >90%. The good match for
rocks exhumed from .1 - .5 GPa is an important constraint for the accuracy of the
thermodynamic calculations which are used to predict the modal abundances over a range
of pressures, temperatures, and fO2 for both plutonic and metamorphic lithologies. (For
55
additional details regarding the petrologic forward modeling parameterization, see the
supplementary materials available online.)
Predicted mineral assemblages at depth rely largely on the amount of exhumation
along our seismic/sample collection transects of the CMB. Since a significant fraction of
the upper crust has been removed since the late Mesozoic via erosional and/or tectonic
processes (Hollister and Andronicos, 2006), plutons emplaced in the lower crust
equilibrated at depths significantly deeper than the modern 32 km Moho. Relict kyanite
in the CGC suggests that ~25 km of uplift has occurred along our northern transect since
about 90 Ma, with as much as 15 km occurring in the last 60 Ma. West of the CSZ,
preserved kyanite suggests 25 km of uplift since 90 Ma and 10 km since 70 Ma (Rusmore
et al., 2005). Along the southern transect, the depth of exhumation has been estimated at
5-10 km east of the CSZ (Rusmore et al., 2000, 2001) and ~12-18 km in the Western
Metamorphic Belt west of the CSZ (Rusmore et al., 2006).
The implication of a deeply exhumed arc for our modeling is that to predict the
composition of the lower crust using Perple_X, we must calculate mineral modes using
pressures and temperatures corresponding to (or approximating) their position in the crust
before the widespread period of normal faulting and exhumation which occurred during
increasingly oblique convergence and gravitational collapse of the CMB from 58-48 Ma
(Hollister and Andronicos, 2006). Pre-exhumation pressures and temperatures used for
calculation of mineral modes should, in theory, correspond to points along a hot
geotherm, with Moho temperatures of ~800 °C – 1000 °C, intended to simulate
temperatures in the arc lower crust during pluton emplacement (Barton, 1990). For
56
plutonic lithologies, however, this association is simplified since the temperature at
which the widespread and voluminous granitoid plutons intruded the crust (~800 °C)
typically far exceeds a conductive geotherm. This means that from isothermal
calculations over a range of pressures, we can use Perple_X to predict the modal
abundance of a rock for which we would otherwise only have a surface observation.
To predict current velocities and densities of the assemblages, however, we make
corrections for post-emplacement exhumation and the change in the thermal regime since
phase equilibration. To accomplish the first correction, we shift the predicted mineral
assemblages upward in the crust by an amount equivalent to the depth of solidification
for each rock sample (as determined by Aluminum in hornblende geobarometry using the
method of Hollister et al., 1987). We then use the Excel macro of Hacker and Abers
(2004) to calculate Vp, Vs, and density for each rock sample at a range of pressures and
temperatures appropriate for current crustal conditions. Modern crustal temperatures are
estimated using a cooler conductive geotherm with a gradient of 11 °C /km based on
Hollister’s (1982) estimate for the CGC, and calculated using the same parameterization
as Tassara (2006).
With one exception, the rock units that make up our models are based on hand
samples collected in the vicinity of the seismic array. These include a high silica tonalite
(62% SiO2), an intermediate silica diorite (57.5 % SiO2), a low silica diorite (51% SiO2),
a gabbro, and an orthogneiss. We also include a serpentinite which is based on Hacker et
al.’s (2003) phase diagram for serpentinized harzburgite at modern deep crustal pressures
and temperatures. Table 1 lists the compositions and peak pressures and temperatures,
57
where known, for the suite of model rock types; sample collection locations are labeled in
Figure 1. Supplementary Figure 2 contains in-situ velocity and density profiles as
predicted by the method of Hacker and Abers (2004) for the mineral assemblages
predicted at depth for each sampled rock. To create layers with gradational seismic
properties or intermediate velocities not predicted by our six basic rock types, we use a
simple mixing equation where the proportion of one rock type decreases linearly while
the proportion of a second type increases. This scheme is not necessarily meant to imply
chemical mixing of rock types, but rather small-scale compositional variations such as
layering which would be indistinguishable seismically at the frequencies contained in our
data. For all models, we use an upper mantle Vp of 8.0 km/sec, Vs of 4.43 km/sec, and
density of 3.3 g/cm3. As targets for forward modeling, we selected sites BS01, BS07,
BS22, BN07, BN19, and BN22 (Fig. 7). These sites span the major mapped geologic
units within the array and capture the most significant structures in the crust as revealed
by the CCP stacks.
Results
Figure 3 shows our constant Vp/Vs CCP stacks along both transects. The direct
Pds converted phase from the Moho and the first two crustal multiples (2P1S and 2S1P)
are remarkably well recorded in the CCP stacks, allowing for good Vp/Vs estimates
across the array (Fig 4). Two-sigma error estimates for Vp/Vs are generally less than +/0.1 with a few exceptions as noted in sections 4.1 and 4.2. Error estimates for crustal
thickness are in the range of +/- 1-3 km. As the surface geology and structure revealed in
58
the CCP stacks differ significantly between the two transects, we discuss the crustal
thickness, structure, and Vp/Vs estimates for the southern and northern lines separately.
In this study we have focused on crustal structure, however we calculated receiver
functions of a sufficient length to image discontinuities as deep as ~100 km. In the upper
mantle, somewhat scattered but widespread negative (blue) arrivals may indicate a
diffuse lithosphere-asthenosphere boundary. Receiver functions constructed using lower
frequency pulses focus this energy between 50 and 70 km beneath both lines
(supplementary Figure 1). These observations are in general agreement with the findings
of Peterson and Russell (2007), who estimate the base of the lithosphere to be near 60 km
based on geothermometry of upper mantle xenoliths collected just east of the Batholiths
study area. Heat flow data from the northern Canadian Cordillera (Hyndman et al., 2005)
and northern Cascadia (Hyndman and Lewis, 1999), which bracket our study area,
provide further corroboration for our observation of a shallow lithosphere-asthenosphere
boundary.
South Line
South Line Seismic Results
Vp/Vs estimates for the southern transect decrease nearly monotonically from
west to east. Our estimates range as high as 1.9 near station BS01, where the array
touches the eastern edge of the Insular superterrane, suggesting the presence of
serpentinite in the lower crust (e.g. Singh and McKenzie, 1993). Within the central CMB
and the Intermontane assemblage, the calculated Vp/Vs drops to more typical crustal
59
values of 1.75 to 1.78 (Fig 4a). The largest decrease takes place in the vicinity of the
CSZ, and the ratio stays nearly constant across the CMB and onto the Chilcotin Plateau
except for a slight decrease to 1.75 in the vicinity of stations BS12 through BS14.
East of station BS06, where the line crosses the CSZ, the signal from the Moho
brightens and deepens, from approximately 28 km to 32 km. The increased brightness is
particularly apparent in the amplitude of the multiples in the CCP stacks (Fig. 3). A
positive mid-crustal interface is imaged near 15 km depth between stations BS06 and
BS09. This interface appears to be truncated or offset by the CSZ in the CCP stacks, but
stacks of receiver functions from individual stations suggest that it may actually rise
toward the surface near the CSZ and that the much lower amplitude mid-crustal arrival to
the west derives from a different structure. This interface disappears beneath the central
portion of the CMB east of station BS10. The aforementioned zone of decreased Vp/Vs
near stations BS12-BS14 corresponds with a slightly diminished and more diffuse signal
from the Moho and some faint conversions originating in the mid crust. The complexity
of the observed signal in this region is the target of an ongoing comparative analysis of
the resolving power of different deconvolution methodologies and will be the subject of a
future publication (Hansen et al., manuscript in prep.) Further east, the crust appears to
thicken slightly beneath station BS21, but a strong negative arrival in the mid crust
indicates the top of a low velocity zone. Significantly diminished P wave velocities
would tend to artificially increase the imaged depth to the Moho, so the apparent crustal
thickening may therefore be an artifact of our assumption of constant Vp. The low
velocity zone could be associated with excess heat or partial melt related to the nearby
60
Anahim volcanics, or simply reflect layering of higher velocity quaternary basalt over
lower velocity sedimentary or plutonic rocks. We do not observe the high Vp/Vs ratio
that is normally associated with partial melt (e.g. Watanabe, 1993), however, and
therefore favor the latter interpretation.
Only in the vicinity of station BS06 do we observe anomalously high data
variance, an effect we attribute to the offset in the Moho at or near the CSZ. Arrivals
coming from different back azimuths therefore encounter the Moho at different depths
which causes significant scatter in aggregate receiver functions created by stacking
different bootstrap sample sets.
South Line Forward Modeling Results
On the southern line, we identified stations BS01, BS07, and BS22 as target sites
for forward modeling (Fig. 7a). BS01 was installed near the boundary of the Insular and
the Intermontane (Fig 1) and displays a unique pattern of velocity decreasing with depth
in the lower crust accompanied by high average Vp/Vs (Fig. 7a). This combination of
seismic properties is suggestive of serpentinite, a low-density hydrated ultramafic rock
that forms in the mantle wedge above subduction zones (e.g. Bostock et al., 2002;
Brocher et al., 2003) and is inferred to be a common component in the lower crust (e.g.
Singh and McKenzie, 1993). Metamorphosed accretionary wedge type sediments have
also been observed to have high Vp/Vs ratios, approaching 1.95, but typically have svelocities more similar to granitoid rocks (e.g. Rudnick and Fountain, 1995; Christensen,
1996). Forward models containing metasediments, therefore, do not produce the negative
arrivals in the mid to lower crust we observe in the data. Serpentinite has been found
61
near the Insular/Intermontane suture outside of the study area (e.g. Bier et al., 2006) and
is often associated with transpressive ocean/continental margins, such as the San
Andreas transform fault in California (e.g. Catchings et al., 2002). To our knowledge,
however, there are no known field occurrences of serpentinite within the Batholiths
project study area. Our best matching forward model consists of 5 km of tonalite
overlying a 5 km layer in which gabbro gradually replaces tonalite and a lower crust
composed of gabbro mixed with an increasing percentage of serpentinite with depth (Fig.
8a). We prefer a model with a maximum of 55% serpentinite at the base of the crust,
although we note that this results in a slightly over-predicted Pds amplitude. Decreasing
the amount of serpentinite, however, has the undesirable effect of lowering the Vp/Vs
and decreasing the amplitude of the mid to lower crustal negative arrivals that are so
striking in the data.
BS07 is located western Stikine terrane (Fig. 1) and displayed a strong midcrustal positive arrival at 15 km flanked by lower amplitude negative arrivals (Fig. 7a)
The Moho signal is strong, ~25% of the direct P amplitude, and originates at ~32 km.
Creating a 1-D profile to match the data proved difficult, as velocity contrasts in the
shallow crust produce numerous multiples that complicate the portion of the signal where
the Moho arrival is observed. We found that a model with a single mid-crustal velocity
increase produces a positive arrival at 15 km, but the maximum amplitude we can model
(by placing the slow tonalite over fast low silica diorite) is only ~50% of the observed
amplitude. Furthermore, this model does not reproduce the negative arrivals observed
before and after the positive arrival in the data. We note that the observed receiver
62
functions for stations BS07 and BN19 are similar, and we have therefore presented a
simple tonalite-over-low-silica diorite forward model for BN19 (Fig. 8b). We find a
better match for BS07, however, in a model that contains a thin, shallow, low velocity
layer (Fig. 8a). The top of this low velocity layer causes the first negative arrival
observed in the data, and the second, larger, negative pulse is a reverberation of this
phase. The bottom of the low velocity layer at 15 km causes the large positive arrival in
the mid-crust. We note, however, that the high amplitude of the positive phase in the
data (~20% of the direct P wave amplitude) is impossible to reproduce with the rock
samples in our forward modeling suite. A thin sill of the (slow) tonalite inserted between
layers of (fast) low silica diorite produces a positive arrival that is only ~10% as large as
the direct P wave amplitude, or roughly half that of the observed mid-crustal arrival.
Station BS22 was the furthest east of our southern stations and was located on the
Chilcotin plateau where Neogene and younger volcanics cover the dominantly Jurassic
intrusives associated with the CMB. The receiver function stack reveals relatively simple
structure, with a sharp Moho at ~33 km, no major mid-crustal interfaces, and an average
Vp/Vs of 1.79. The relatively high Vp/Vs and moderate Moho amplitude (~18% of
direct Pds) suggest a slightly more mafic average crustal composition than that beneath
late Cretaceous to Eocence intrusions in the central CMB. This observation is in good
agreement with field samples which indicate that the Jurassic plutons are generally lower
in bulk SiO2 than the younger plutons to the west. Our best-matching model is composed
of 15km of 80% intermediate silica diorite/20% low silica diorite grading linearly to
100% low silica diorite between 15 and 32 km (Fig. 8a).
63
North Line
North Line Seismic Results
Significant Vp/Vs variations are largely confined to the southwestern end of the
northern transect, primarily between the surface exposures of the 59 Ma tonalite
Quottoon pluton and the CGC (see bottom of Fig. 4b for names of geologic units). We
note that the data in this region are generally sparse and noisier than that recorded
elsewhere in the array as suggested by the large error bars in Fig. 4b. However, the
general trends, if not the absolute Vp/Vs and crustal thickness values, are roughly
consistent with those observed on the south line. Southwest of the CSZ, rocks of the
Intermontane Yukon-Tanana member abut a sliver of the Gravina belt that marks the
boundary with the Insular-affiliated Alexander Terrane (Fig. 1). The suture and
surrounding units are heavily intruded by mid-cretaceous granodiorite plutons. Our
Vp/Vs estimates are extremely high, approaching 1.9, and again suggest the presence of
serpentinite in the lower crust (see section 4.1). Traversing northeastward across the
Scotia-Quaal metamorphic complex (Gareau, 1991) and the CSZ into the Quottoon
pluton, the velocity ratio drops abruptly to 1.65 before stabilizing around 1.8 in the
volcanic and sedimentary rocks of the Intermontane.
Due to the large fluctuations in Vp/Vs across short lateral distances on the
southwest end of the north line, the uncorrected CCP stack (Fig. 3b) displays apparently
jagged, short wavelength topography on the Moho beneath Douglass Channel.
Corrections to the stacking algorithm to account for the variations in Vp/Vs, however,
64
restore the imaged Moho to a more reasonable geometry (Fig. 6). The gross crustal
structure is similar to that observed on the southern transect: the Moho is bright and
shallow at ~28 km to the southwest of the CSZ and deepens slightly, to just over 30 km,
to the northeast. There is some complication and a possible bifurcation in the Moho
conversion between stations BN09 and BN12 in the Kitimat River valley west of Terrace.
Beneath stations BN15 through BN21, where the array crosses the southern reaches of
the Jurassic to Cretaceous Bowser Basin, a prominent positive arrival can be traced from
near surface levels to the lower crust moving away from the mapped trace of the Eocene
EBD. Although this structure appears to be dipping shallowly in Fig. 6, a correction for
the angle between the strike of the detachment and the line of the array results in a
steeper dip angle more consistent with the detachment structure mapped by Rusmore et
al. (2005). This anomaly tapers out north of station BN21, where the line of
seismometers jumps 30km ENE, approximately parallel to the dip of the EBD.
Errors in Vp/Vs and crustal thickness are highest along the Douglass channel sites
(BN01-NB07), ranging up to +/-0.15 and +/-2.5 km, respectively. These estimates may
be lower than the realistic uncertainty due to the aforementioned scant and noisy data in
this region. East of the CGC, data are more plentiful and errors are consistently lower, in
the range of approximately +/-.07 (Vp/Vs) and +/- 1.5 km (Moho depth).
Due to the noisy, sparse data recorded on the southwest end of the northern line,
we have not attempted to model any single station from this region. The general structure
observed beneath stations BN01 and BN02, however, is largely similar to that imaged on
the far western end of the south line, where station BS01 sits very near the boundary
65
between the Insular and the Intermontane. The extremely high Vp/Vs ratio indicates the
likely presence of serpentinite in the lower crust, as does the large amplitude Moho
converted phase. The low average Vp/Vs observed in the vicinity of station BN04 likely
results from the highly felsic and voluminous Paleocene Quottoon pluton emplaced along
the CSZ.
North Line Forward Modeling Results
On the north line, we use stations BN07, BN19, and BN22 to provide seismic
constraints for our forward models (Fig. 7b). BN07 was situated squarely in the surface
exposure of the CGC where intermediate Eocene plutons intrude deeply exhumed
metamorphic rocks. The composite receiver function suggests a layered crust, with fairly
sharp velocity increases near 10 and 18 km. The average Vp/Vs is 1.8 above a slightly
diffuse Moho at 28km. We obtain a good match with a 3-layer model consisting of 10
km of orthogneiss over a mid crust composed of 40% orthogneiss and 60% low silica
diorite and a gabbroic lower crust (Fig. 8b). While low silica diorite in the lower crust
would produce a receiver function with a similar form, we prefer gabbro, with its higher
Vp/Vs, in order to bring the average modeled velocity ratio closer to the observed value.
Station BN19 was chosen as a forward modeling target based on its location atop
the Jurassic/Cretaceous sediments of the Bowser Basin and as a representative of the
group of stations on the northern line that display a prominent positive arrival in the mid
crust. The forward model presented for station BN19 in Fig. 8b serves as a counterpoint
to that for station BS07 (Fig. 8a), as the observed data from both stations are very similar.
The column is simple: 17km of tonalite in the upper crust above a lower crust of 100%
66
low silica diorite. This model has several shortcomings, however, including the absence
of negative arrivals at 8 and 22 km, a too-weak mid crustal positive, and an average
Vp/Vs lower than the observed value of 1.80. A model similar to that presented for BS07
is therefore preferable to the simple 2-layered model for BN19. As noted in section 4.1,
the amplitude of the mid-crustal arrival is less than half that in the data, implying that
some physical characteristic of the rock other than bulk composition must contribute to
the impedance contrast at the base of the low velocity zone. We hypothesize that the
source of these low velocities may be a zone of brecciated and/or fluid-rich rock, possibly
caused by Eocene displacement on the EBD and subsequent intrusion of surface waters
along the fractured rock in the fault zone. Higher resolution images of the shallow
structure that will result from the active seismic component of the project, currently
scheduled for the summer of 2008, will allow for further testing of this hypothesis.
Station BN22 was deployed on the far northeast end of the Northern line near the
mapped contact between the Bowser basin and plutonic rocks of Stikine terrane.
Noticeably absent from the stacked receiver function for BN22 is the positive arrival near
15 km so striking at stations BN19 through BN21. The overall form of the receiver
function is very similar to that recorded at station BS22; the crust is largely transparent
and a sharp Moho is imaged near 30 km. The Vp/Vs is slightly lower than that estimated
for station BS22 – 1.77 vs. 1.79 – and this is reflected in our choice of plutonic rocks in
the preferred model. The model presented in Fig. 8b is comprised of a 16 km thick upper
crust containing 50% intermediate silica diorite and 50% tonalite with the lower crust
grading from intermediate to low silica diorite between 17 and 30 km. At both BS22 and
67
BN22, our models predict a slightly narrower Pds phase than that observed in the data,
suggesting that the transition from crust to mantle velocities is likely somewhat less
abrupt than the step-like increase we have modeled.
Discussion
Our results indicate that the majority of the crust beneath the study area can be
modeled using only simple combinations of the plutonic rocks sampled at the surface and
their bulk compositional equivalents at depth. With two notable exceptions, we are able
to obtain very good matches for the observed receiver functions, in terms of amplitude of
the Moho converted phase and average Vp/Vs, without including any ultramafic restites
or other exotic assemblages not found in outcrops. The first exception is the far western
section of the study area, where the high average Vp/Vs and low absolute velocities at
depth strongly suggest the presence of serpentinite in the lower crust. We hypothesize
that the low density serpentinite formed as a result of mantle wedge hydration during
Mesozoic subduction and subsequently migrated into the crust, possibly exploiting
structural pathways opened along margin-parallel transform faults (e.g. Bier et al., 2006;
Catchings et al., 2002).
The second location in the study area that proves difficult to model using only
intermediate plutonic rocks is the central portion of the southern line, roughly between
stations BS12 and BS15. We chose not to attempt a forward model in this region because
stacked receiver functions from individual stations are complex and highly variable over
short distances, thus making the choice of a “representative station” arbitrary if not
68
impossible. Arrivals from the Moho at stations BS12-BS15, where evident at all, are low
amplitude and inconsistent across different back-azimuth groups. Weak Pds arrivals
indicate a small or gradational velocity contrast across the Moho, suggesting that the
lower crust is extremely fast, that the upper mantle is considerably slower than average,
or that the transition from low to high velocities is smeared over several kilometers rather
than abrupt. Body wave tomography performed by Peck et al. (2006) indicates that the
lower crust is perhaps 1-2% faster than average here, which alone is insufficient to cause
the diminished Pds amplitudes. The tomographic study shows no velocity anomaly in the
upper mantle beneath the central portion of the southern line, and initial analysis of Pn
waves that travel in the uppermost mantle (Jasbinsek, 2007) find near normal Vp values
of ~8.1 km/sec. The nature of the lower crust and upper mantle in this area thus remain
somewhat enigmatic, but the ongoing work of Hansen et al. and the active seismic study
are expected to yield higher resolution images of the central portion of the southern
transect.
To further test the possibility that garnet pyroxenites are present in large fractions
in the lower crust, we constructed a forward model with a hypothetical lower crustal
“root” containing a 50/50 mix of the low silica diorite from our modeling suite and the
“average Sierra pyroxenite” as described by Ducea (2002). The Vp/Vs of the eclogitefacies garnet pyroxenite is 1.79 at 1 GPa compared to ~1.80 for our lower crustal rocks
(gabbro and low silica diorite), and thus does not serve as a clear differentiator. The Vs
of the residual pyroxenite, however, should be in the range of 4.5-4.6 km/sec - greater
than the mantle velocities used in our forward models (4.43 km/sec). Our hypothetical
69
50/50 root model would give the lower crust a Vs of ~4.2 km/sec, Vp of ~7.5 km/sec,
Vp/Vs of 1.79, and a density very similar to that of the upper mantle, ~3.3 g/cm3.
Synthetic receiver functions predict the amplitude of the converted phase from an
interface between the root and upper mantle peridotite should be on the order of ~6% of
the direct P amplitude. With the exception of the stations near the center of the southern
line, the observed Pds amplitudes, averaging ~20% of the direct P amplitudes, are much
larger than would be expected from a Moho separating upper mantle lithologies from
compositions similar to our hypothetical root.
The obvious differences between the eclogite/pyroxenite forward model and the
observations from the majority of the study area indicate that the arrival we have
interpreted as Pds is not consistent with the expected converted phase between a mafic
restite and the mantle. We acknowledge the possibility that what we are imaging is the
interface between intermediate intrusives and ultramafic residues within the crust, which
have a seismic signature very near that of the mantle, but consider this scenario to be
unlikely. Combined petrologic and seismic studies of the southern Sierra Nevada
batholith in California (Saleeby et al., 2003) suggest that the transition from liquidfractionated intermediate rocks to mafic residual assemblages should be gradual,
occurring over tens of kilometers, and therefore incapable of producing a sharp seismic
converted phase. Recently acquired seismic data from a portion of the northwestern
Sierra Nevada batholith believed to be underlain by a residual root provide additional
evidence that the Moho in such a setting is transitional and produces a diffuse, long
wavelength Pds arrival (Zandt et al., 2007).
70
Conclusions
Analysis of broadband seismic records collected as part of the Batholiths project
yields the first quantitative estimates of the structure and composition of the crust beneath
the central Coast Mountains of British Columbia. On the southern transect, we find thin
crust (27 +/- 4km) and high Vp/Vs (1.86 +/- 0.12) in the outboard terranes. East of the
CSZ, we image the Moho at an average depth of 32 +/- 1.5 km and Vp/Vs decreases to an
average value of 1.77 (Fig. 4a). Along the northern transect, we observe high Vp/Vs
(1.91 +/- 0.12) in the outboard Insular superterrane, a sharp decrease to 1.71 +/- 0.08 in
the western Yukon Tanana block, and values fluctuating around 1.8 across the central
gneiss complex and western Stikine Terrane. The crust is thinnest (~26 +/- 3km) beneath
the outboard terranes and averages ~30 km east of the EBD (Fig. 4b). The bright midcrustal positive flanked by negative arrivals beneath western Stikinia (Fig. 7b) requires
the presence of a ~5 km thick layer in the upper to mid crust with seismic velocities
lower than the most felsic plutonic rocks in the area (Fig. 8a). Based on these low
velocities and the correlation with the mapped exposure of the EBD, we hypothesize that
this feature may be related to motion along the EBD at depth. A diffuse negative arrival
observed in lower frequency receiver functions between 50 and 70 km beneath both lines
may mark the lithosphere-asthenosphere boundary (supplementary Fig. 1).
Forward models based on rock samples from the field area and their predicted
equivalents at deep crustal pressures and temperatures match the observed seismic data
well (Fig. 8), with two notable exceptions. We find it necessary to include a significant
amount of serpentinite in the lower crust to model the low velocities and high Vp/Vs
71
ratios observed near the boundary of the Insular and Intermontane terranes, and we are
unable to model the complex and highly variable receiver functions observed in the
central portion of the southern transect. While our results do not preclude the existence
of a mafic/ultramafic root beneath some portion of the CMB, particularly in the vicinity
of stations BS12-BS15, the evidence presented herein suggests that much of the modern
CMB is not in possession of a high velocity residual root. Furthermore, the abrupt
transition from low to high velocities indicated by the sharp, high amplitude converted
phases suggests that we are imaging a tectonic Moho rather than a gradual transition from
granitoid arc rocks to their complementary ultramafic residuals.
Acknowledgements
The Batholiths project was funded under the U.S. National Science Foundation’s
Continental Dynamics program, grant EAR-0309885. The instruments used in the field
program were provided by the PASSCAL facility of the Incorporated Research
Institutions for Seismology (IRIS) through the PASSCAL Instrument Center at New
Mexico Tech. Data collected during this experiment will be available through the IRIS
Data Management Center. The facilities of the IRIS Consortium are supported by the
National Science Foundation under Cooperative Agreement EAR-0004370 and by the
Department of Energy National Nuclear Security Administration. Compositions of the
rock samples used for forward modeling were determined by x-ray diffraction by Jeff
Thole at Macalaster College, St. Paul, MN, 55105.
72
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Sobolev, S.V., Babeyko, A.Y., 1994. Modeling of mineralogical composition, density
and elastic wave velocities in anhydrous magmatic rocks. Surv. Geophys. 15, 515-544.
Tassara, A., 2006. Factors controlling the crustal density structure underneath active
continental margins with implications for their evolution. Geochem. Geophy.
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Watanabe, T., 1993. Effects of water and melt on seismic velocities and their application
to characterization of seismic reflectors. Geophys. Res. Lett., 20(24), 2933-2936.
Wei, C. J., Powell, R., 2003. Phase relations in high-pressure metapelities in the system
KFMASH (K2O-FeO-MgO-Al2O3-SiO2-H2O) with application to natural rocks.
Contrib. Mineral. Petr., 145, 301–315.
White, R. W., Powell, R., Holland. T. J. B., 2001. Calculation of partial melting
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(NCKFMASH). J. Metamorph. Geol., 19, 139– 153.
White, R. W., Powell, R., Phillips, G. N., 2003. A mineral equilibria study of the
hydrothermal alteration in mafic greenschist facies rocks at Kalgoorlie, Western
Australia. J. Metamorph. Geol., 21, 455– 468.
Zandt, G., Ammon C.J., 1995. Continental crust composition constrained by
measurements of crustal Poisson’s ratio. Nature 374, 152-154.
Zandt, G., Gilbert, H., Frassetto, A., Owens, T., Jones, C., 2007. Insights on lithospheric
foundering from the Sierra Nevada Earthscope Project (SNEP). Eos Trans. AGU,
88(52), Fall Meet. Suppl., Abstract T31E-01.
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Zandt, G., Gilbert, H., Owens, T. J., Ducea, M., Saleeby, J., Jones, C.H., 2004. Active
foundering of a continental arc root beneath the southern Sierra Nevada in California.
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Zandt, G., Myers, S.C., Wallace, T.C., 1995. Crust and mantle structure across the Basin
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78
Figures
Figure 1: Batholiths project seismic stations overlain on the GSC Canadian Geologic
map (1996). The light blue rectangle on the inset Google Earth map shows the regional
location of the study area. Seismic stations are marked with green stars and rock sample
locations with yellow circles and GJP-XX labels. See Table1 for sample compositions.
W=Wrangellia, A=Alexander, Y=Yukon Tanana, CSZ=Coast Shear Zone, S=Stikine,
B=Bowser Basin EBD=Eastern Boundary Detachment.
79
Figure 2: Earthquakes used to construct receiver functions in this study (black circles).
Center of the array is marked by the star.
80
Figure 3: Constant Vp/Vs cross sections for the south (A) and north (B) transects. Red
arrivals indicate upward transitions from higher velocities to lower velocities and blue
arrivals the opposite. From top to bottom, cross sections consist of receiver functions
stacked on the direct (red) Pds arrival, the (red) 2P1S multiple, and the (blue) 2S1P
multiple. The multiples are stacked using lower frequency receiver functions than the Pds
phase. Note the striking prominence of the 2P1S and 2S1P arrivals.
81
82
Figure 4 (previous page): Summary of receiver function results correlated with local
geology for the south line (A) and the north line (B). From top to bottom, the panes
represent: surface elevation averaged over 50 km around each station; Moho image depth
for the Pds and 2P1S phases; Vp/Vs calculated based on the method of Kind et al.
(2002); calculated crustal thickness; and major geologic features along strike. All
calculated values are plotted +/- one standard deviation based on bootstrap resampling.
Figure 5: South line CCP stacks corrected for Vp/Vs combined with a geologic strip
map. Green stars mark station locations and colored lines mark the boundaries between
major tectonostratigraphic units. A=Alexander, Y=Yukon Tanana, CSZ=Coast Shear
Zone, S=Stikine.
83
Figure 6: North line CCP stacks corrected for Vp/Vs combined with a geologic strip
map. Green stars mark station locations and colored lines mark the boundaries between
major tectonostratigraphic units. A=Alexander, Y=Yukon Tanana, CSZ=Coast Shear
Zone, S=Stikine, B=Bowser Basin.
Figure 7: Single station stacks for receiver functions selected as targets for forward
modeling. Sub-horizontal dashed lines denote Moho arrivals and jagged vertical lines
signify major structural boundaries. From left to right, the boundaries are the
Insular/Intermontane terrane boundary (solid line), the CSZ (dashed line) and the EBD
(dash-dot line). See figures 5 and 6 for station locations.
84
85
Figure 8 (previous page): Best fitting forward modeled receiver functions for each target
station compared with observed data for the south (A) and north (B) lines. From lef to
rith in each panel are the stacked observed receiver functions, the synthetic receiver
function from the preferred forward model, a schematic diagram of the petrologic column
used to create the preferred forward model, and the predicted Vs for the forward model
based on Hacker and Abers (2004). TON= tonalite, GABB=gabbro, SERP=serpentinite,
LSD=low silica diorite, ISD=intermediate silica diorite, ORTH=orthogneiss.
86
Tables
Rock Type
Tonalite
Sample Name
SiO2
TiO2
Al2O3
Fe2O3*
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Total
peak P (kbar)
peak T (°C)
GJP-79
62.10
0.84
16.54
5.50
0.08
2.81
5.49
4.05
1.64
0.25
0.60
99.90
3.20
783.00
Int. silica
Diorite
GJP-1
57.48
0.81
16.77
7.87
0.15
4.12
6.37
3.52
1.83
0.18
1.93
100.04
1.90
769.00
Low silica
Diorite
GJP-43
50.99
1.14
19.15
9.74
0.17
5.01
9.34
3.76
0.34
0.29
0.39
100.30
4.10
?
Gabbro
Orthogneiss
GJP-50
45.15
0.79
20.45
10.36
0.12
7.33
14.70
1.05
0.26
0.04
0.39
100.64
4.10
?
GJP-76
63.05
0.56
16.69
5.56
0.15
1.67
6.12
3.18
2.20
0.28
0.64
100.09
?
?
Table 1: Bulk compositions, peak pressure, and peak temperature (where known) of
rock samples used in forward modeling. Bulk compositions determined via x-ray
diffraction by Jeff Thole at Macalaster College. Pressures are from Hollister (1982) and
temperatures are from Blundy and Holland (1990). All detected iron was assumed to be
Fe2O3. See Fig. 1 for sample collection locations.
Supplementary Materials
Explanation of petrologic modeling parameters
The method of sub-solidus phase equilibrium calculations using the free energy
minimization algorithm Perple_X (Connolly, 1990, 2005) and thermodynamic database
87
of Holland and Powell (1998) enables the accurate probing of deep crustal compositions
(e.g. Behn and Kelemen, 2006; Tassara, 2006) which may otherwise be only roughly
constrained by seismic observations of rock physical properties. In this study,
equilibrium mineral assemblages were calculated for initial bulk compositions (Table 1)
that include the major element oxides: SiO2, Al2O3, FeO*, MgO, CaO, Na2O, K2O, and
H2O (estimated from the lost on ignition (LOI) value measured during XRF analysis).
The solution models used in these calculations are shown below in Supplementary Table
1.
Following the precedent of similar studies (e.g. Behn and Keleman, 2006), we
have excluded the oxides TiO2, MnO, and P2O5 from our equilibrium calculations as their
inclusion would needlessly complicate the calculations without significantly impacting
the resulting modes. The oxide K2O, which also requires additional poorly characterized
solid solution models, was not omitted because it is a major component in some of the
intermediate composition samples generated in the continental arc setting (unlike the
more mafic lithologies generated in the island arc setting that was the subject of the study
of Behn and Kelemen, 2006),. Furthermore, the intermediate (on average tonalitic) bulk
composition samples for which we are calculating phase equilibria (Table 1) contain
modal biotite (potassium iron magnesium aluminum silicate, a common igneous rock
forming mineral). Excluding the K2O oxide would eliminate our ability to predict its
thermodynamic stability by removing all possible end-members for solution models.
Pressure and temperature-dependent fO2 buffers ranging from fayalite magnetite-quartz
(FMQ) to Nickel-nickel oxide (NiNiO) were used to control oxygen fugacity.
88
Supplementary Figures
89
Supplementary Figure 1 (previous page): CCP stacks created using a smaller Gaussian
width factor to emphasize low frequency energy. The broad negative (blue) arrivals in
the 50-70 km range may originate at a diffuse lithosphere-asthenosphere boundary.
90
91
Supplementary Figure 2 (previous page): In-situ velocity and density profiles for the
predicted mineral assemblages for each rock type used in the forward models. Sharp
velocity increases in the 20-30 km range arise from the increased stability of mafic
minerals such as pyroxenes, garnets, and alumino-silicates over feldspars (e.g. Sobolev
and Babeyko, 1994). Mineral asseblages are based on Perple_X calculations for the bulk
compositions in Table 1. Velocities and densities were calculated with the Hacker and
Abers (2004) Excel macro. ISD=Intermediate silica diorite, LSD=low silica diorite,
orth=orthogneiss.
92
Source b
HP98
HP96
HP96
HP98
WP03,
Amphibole
WPH03
c
Biotite
PH99, HP98
Plagioclase
NCK80
HP01,
Melt
Na-Mg-Al-Si-K-Ca-Fe hydrous silicate melt
WPH01
a
Compositional variables (w,x,y, and z) are calculated as a function of temperature and
pressure using the free energy minimization algorithm Perple_X'06 (Connolly, 1990,
2005) and vary between 0 and 1. See Perple_X documentation for more details.
Sources: HP96 (Holland and Powell, 1996); HP98 (Holland and Powell, 1998); HP01
(Holland and Powell, 2001); WP03 (Wei and Powell, 2003); WPH01 (White et al.,
2001); WPH03 (White et al., 2003); PH1999 (Powell and Holland, 1999); NCK80
(Newton et al., 1980).
b
Holland and Powell solutions models used wherever possible to maintain internal
consistency.
c
Mn component in all calculations is zero since MnO is excluded from initial bulk
compostions.
Phase
Olivinec
Orthopyroxene
Clinopyroxene
Garnetc
Formula
Mg2xFe2yMn2(1–x–y)SiO4, x+y≤1
[MgxFe1-x]2-yAl2ySi2-yO6
Na1-yCayMgxyFe(1-x)yAlySi2O6
Fe3xCa3yMg3zMn3(1-x-y-z)Al2Si3O12, x+y+z≤1
Ca2-2wNaz+2w[MgxFe1-x]3+2y+zAl3-3y-wSi7+w+yO22(OH)2,
w+y+z≤1
K[MgxFeyMn1-x-y]3-wAl1+2wSi3-wO10(OH)2, x+y≤1
NaxCa1-xAl2-xSi2+xO8
Supplementary Table 1: Solid Solution Modelsa (Modified from Behn and Kelemen,
2006, table 2)
93
APPENDIX C: LITHOSPHERIC SHEAR WAVE
VELOCITY STRUCTURE OF THE CENTRAL COAST
MOUNTAINS BATHOLITH, BRITISH COLUMBIA
Josh Calkins, George Zandt, Ken Dueker
Abstract
The western margin of central British Columbia was the site of voluminous
granitoid batholith emplacement beneath a continental arc throughout much of the
Mesozoic and into the earliest Cenozoic. The exposed roots of this system, collectively
known as the Coast Mountains Batholith (CMB), form the core of the mountain range
that runs the length of the western Canadian margin. Isotopic evidence and comparison
with the Sierra Nevada Mountains of California strongly suggest the formation of a
garnet rich root beneath the CMB that would likely have been denser than the underlying
mantle and thus prone to foundering. We analyze phase and amplitude measurements of
Rayleigh waves from teleseismic events recorded in the central CMB to gain insight into
the modern day structure of the mantle lithosphere and search for evidence of an
ultramafic root. Our results show slow upper mantle velocities that deviate from global
averages by 4-9% and two major localized anomalies. We interpret a slow anomaly near
the town of Bella Coola to be related to flow around the northern edge of the subducting
Juan de Fuca slab. A fast anomaly imaged on the west side of the Insular/Intermontane
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terrane boundary likely results from thicker mantle lithosphere beneath an older portion
of the arc.
Introduction and regional tectonic setting
The Coast Mountains Batholith (CMB) of British Columbia is the remnant of an
arc magmatic system fueled by subduction of oceanic plates of the Pacific basin beneath
the western margin of North America from mid-Paleozoic through early Tertiary time
(Monger et al., 1993). The exhumed plutonic and metamorphic rocks of the CMB form a
mountain range with elevations up to 4000 meters spanning from northern Washington to
Southern Alaska, and stand as one of the largest exposed granitic masses on earth. The
Batholiths project study area covers the region where rocks of the CMB intrude upon the
boundary between the Insular and Intermontane superterranes (Monger, 1993), which
accreted to the North American continent by late Jurassic time (Colpron et al., 2007).
The toe of the wedge of Proterozoic strata thought to underlie the eastern margin of the
accreted terranes has been placed several hundred kilometers east of our easternmost
station based on seismic reflection data (Snyder et al., 2002). The Coast Shear Zone
(CSZ), a >1200km long steeply east-dipping zone of ductile deformation with a complex
history of strike slip and orogen-normal deformation (Rusmore et al., 2001), splits the
CMB along strike near the western portion of the study region (Fig. 1). Emplacement
depths of plutons now exposed at the surface range from 25km in the northern portion of
the study region to as little as 5-10 km in the south (Rusmore et al., 2005, Hollister and
Andronicos, 2000).
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While simple mass and chemical balance calculations necessitate the creation of a
complementary mafic to ultramafic root at least as massive as the granitic plutons in
continental arc magmatic systems (e.g. Conrad and Kay, 1984, Kay and Kay, 1985,
Ducea, 2002), the composition and long term fate of the root beneath western British
Columbia are unknown. The Batholiths continental dynamics project was conceived with
the aim of probing the CMB from the surface to the upper mantle in search of evidence of
the ultramafic root and to investigate the dynamics of root formation and removal
processes.
The lithologic makeup of a sub-arc root is largely dependent on the depth at
which melting and equilibration occur (e.g. Kay and Kay, 1991, Ducea, 2002). At the
pressures expected below depths of approximately 40 km, garnet becomes a major
component and plagioclase completely absent in the residue at the base of the melt
column, dramatically increasing both density and seismic velocity (Ducea, 2002). Garnet
pyroxenite xenoliths from the Sierra Nevada that formed at depths between 35 and
100km are inferred to be the residua from the extraction of granitoid melts from a
basaltic-andesite source (Ducea, 2002). Calculations performed using the method of
Hacker and Abers (2004) predict densities of ~3.5 g/cm3 and P wave velocities in excess
of 8.0 km/sec for these eclogites, rendering them both prone to foundering into the
underlying mantle and difficult to differentiate from upper mantle peridotites based on
seismic velocities.
Unfortunately, no deep crustal xenoliths have been found in the central CMB.
Geochemical evidence in the form of steep rare earth element patterns and the absence of
96
Europium anomalies, however, indicates that dense, garnet-rich rocks remained in the
melt-source region (Girardi et al., 2007.). These patterns are particularly pronounced
during high flux magmatic events centered at ~100 and 52 Ma (e.g. Gehrels et al., 2008),
possibly as a result of overthickening of the continental crust, introducing more fertile
material into the melt zone (Ducea and Barton, 2007). By comparison with the Sierra
Nevada, a tectonically analogous setting known to be underlain, at least in part, by an
ultramafic root (e.g. Ducea and Saleeby, 1998; Zandt et. al, 2004), the minor element
chemistry of plutonic rocks of the CMB strongly suggest the formation of a garnet
pyroxenite root.
Geophysical studies that would illuminate details of the current lithospheric
structure in the central CMB are limited, however. Regional partitioned waveform
(Frederiksen et al., 2001) and body wave (Mercier et al., manuscript in prep)
tomographic models have found upper mantle velocities roughly 5-8% lower than global
averages beneath the Coast Mountains, but do not reveal detailed structure in the
Batholiths project study area. High heat flow observations from the northern Canadian
Cordillera (Hyndman et al., 2005) suggest the source of the depressed upper mantle
velocities may be the presence of shallow asthenosphere. A study combining seismic
receiver functions and petrologic modeling (Calkins et al., 2008) indicates that crustal
rocks immediately above a shallow seismic Moho (~32km) are characterized by
relatively low velocities and densities, and are thus inconsistent with the mafic to
ultramafic lithologies inferred to have formed during batholith formation.
97
Herein, we use measurements of the relative phase and amplitudes of surface
waves traversing the study area to extend our observations into the upper mantle.
Rayleigh waves are a particularly useful tool for investigating lithospheric structure due
to their dispersive nature; high frequency, short wavelength energy is sensitive to the
seismic velocities of rocks in the crust and longer period energy is sensitive to deeper
structure. As the higher frequency components of the teleseismic data used in this study
are generally attenuated and noisy, we restrict our observations and interpretations to the
better imaged upper mantle. A future publication combining Rayleigh wave and ambient
noise tomography (J. Stachnik et al., manuscript in prep.) will provide higher quality
images of crustal structure and velocity contrasts across the Moho.
Data and Methods
Our data set consists of 15 months worth of broadband data collected at 42 sites
between June 2005 and September 2006. Each site was equipped with either a
Streckeisen STS-2 or a Guralp CMG3T sensor linked to a Quanterra Q330 digitizer. Due
to the rugged topography and scant road access to much of the central CMB, the array
was deployed in two transects paralleling the only major roads that cross the orogen
between 52° N and 55° N. To image the west side of the CSZ and the
Intermontane/Insular superterrane boundary, both lines were extended offshore along the
deep fjords that cut the western CMB. Inter-station spacing along each transect averaged
roughly 12 km (Fig. 1).
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From the set of all earthquakes recorded during the deployment with magnitudes
(Ms) greater than 5.5 and epicentral distances between 30 and 120 degrees, 54 events
with clear Rayleigh wave energy and high signal to noise ratios were selected for
processing (Fig. 2). We avoided records with more than 2 dominant pulses in the
fundamental Rayleigh wave and those lacking clear separation between the fundamental
and higher-order modes. Subsets of events vary based on the quality of the signal at
various frequency ranges, but a minimum of 34 events were used at each bandpass range.
Crossing rays originating from the north and east of the array are somewhat limited due
to the paucity of sources from these azimuths and the complex source-receiver travel
paths (Figs. 2 and 3), but coverage along both transects and in the western portion of the
region separating the two lines is sufficiently dense to resolve major features.
To solve for phase velocities, we used the method of Forsyth and Li (2005),
which models the observed Rayleigh wavefield as an interference pattern between two
plane waves. This approach has the benefit of decreasing the influence of multipathing
and scattering which occur between the source and the array, and has been successfully
employed in studies of both continental (e.g. Li et al., 2003, Weeraratne et al., 2003,
Yang and Forsyth, 2006) and oceanic (e.g. Forsyth et al., 1998) lithosphere. The
inversion uses an iterative, two-staged approach. A simulated annealing inversion is first
performed (with the velocity model fixed) to solve for the real and imaginary components
of the best-fitting two plane wave solution. The plane wave parameters are then used as
inputs to a linearized least squares inversion to solve simultaneously for phase velocities
and corrections to the plane wave parameters. Events that are poorly represented by the
99
two plane wave model, as determined by the standard deviation of the residuals, are
automatically down-weighted in successive iterations and thus have minimal impact on
the final solution. Using the technique described by Yang and Forsyth (2006), phase
velocities were smoothed between adjacent grid points with a Gaussian function with a
characteristic wavelength of 50km. This value was chosen to maximize the resolution
afforded by close station spacing within the array while limiting the degradation of model
variance that results from using a shorter smoothing length. For a complete description
of the inversion and associated error estimates, see Forsyth and Li (2005).
For each event, fundamental mode Rayleigh waves were isolated on the vertical
component for 14 bandpass ranges centered on periods between 20 and 100 seconds.
Due to the long distance separating the two halves of our array, we assumed that seismic
energy reaching the northern transect may have traveled a significantly different path
from the event than that reaching the southern transect. We therefore treated each
earthquake as two independent observations, and solved for different plane wave
parameters on each transect. Errors are calculated as twice the standard deviation of the
phase velocity at each grid point and are presented herein as the percent perturbation to
average velocity which can be considered significant at the 95% confidence limit. On the
shear velocity maps in Figure 5, the 1.5% and 2.5% contours for the 30 second data set
are plotted and used to mask the poorly resolved area (see also supplementary figure 1).
To invert for shear velocity as a function of depth from the gridded phase velocity
data, we use the linearized least-squares inversion program ‘surf’ (Hermann, 1987), in an
iterative approach whereby damping is decreased after each iteration (Snoke and James,
100
1997; Larson et al., 2006; Warren et al., 2008). To calculate an appropriate starting
model for the 2-D phase velocity inversion, we ran an inversion using the average
dispersion curve for the entire study area (Fig. 4a). Upper mantle starting velocities for
this preliminary inversion were based on the tomographic studies of Frederiksen et al.
(2001) and Mercier et al. (2008). Starting crustal thickness and velocities are consistent
with both seismic refraction data from just north of our study area (Morozov et al., 2003,
Hollister and Andronicos, 2006) and with combined receiver function/petrologic
modeling of data collected as part of the Batholiths project (Calkins et al., 2008).
Throughout the inversions, Vp, density, and attenuation parameters were held constant.
To test the dependence of the average 1-D Vs model on the starting model, we also ran a
series of inversions beginning with the global model AK135 (Kennett et al., 1995) as a
starting point and repeatedly updating the Vs model with the solution from the previous
run. After several iterations, the results converged towards our preferred local model
(Fig 4b).
Results and Discussion
Results of the phase and shear velocity inversions are summarized in figures 4
and 5. In general, we find extremely slow upper mantle velocities throughout the study
area and therefore plot deviations from the average Vs at each depth slice (Fig. 5) rather
than deviations from a reference model. Compared with the global model AK135
(Kennett et al., 1995), our results show velocities slower by 4-5% in the uppermost
mantle (Moho to 75km) and by 6-8% between 75 and 200km. Our 1D Vs profile is
101
slower even than the Tectonic North America (TNA) model of Grand and Helmberger
(1984), which found sharply diminished shear velocities in the upper mantle beneath the
tectonically active portions of the western United States and southernmost Canada
relative to the continental shield. Our model averages ~1-5 % slower than TNA in the
upper mantle.
We see maximum deviation from the global models near the very well-resolved
center of the southern transect, where the mantle Vs at 110km is 4.1 km/sec, or 9% (+/~1%) slower than AK135. The anomaly is less pronounced near Moho depths (Fig. 5a),
but trends towards the most recently active portion of the late Cenozoic Anahim volcanic
belt (see Fig. 1 for location). The depressed velocities near 100 km are in remarkably
good agreement with the observations of Frederiksen et al. (2001), who found the lowest
upper mantle velocities in all of Canada beneath the southern Coast Mountains. Receiver
functions from the same region (Calkins et al., 2008) show very faint converted phases
from the Moho, indicative of a small velocity contrast between the base of the crust and
the upper mantle. A -9% anomaly is difficult to explain based on the effects of
composition alone and is likely to indicate the presence of fluids or partial melt (e.g. van
der Lee and Nolet, 1997). Although the relationship between temperature, partial melt,
and Vs is complicated and not entirely understood (e.g. Priestley and McKenzie, 2006),
as little as 1% partial melt has been shown to be capable of producing a ~8% decrease in
mantle Vs (Hammond and Humphreys, 2000).
The geometry of the slow anomaly, visible in the 155 km depth slice (Fig 5e) near
the town of Bella Coola and trending to the northeast at shallower depths, suggests a
102
possible linkage to the descending Juan de Fuca slab. Mercier et al. (2008) have imaged
the edge of the slab just south of our study area, and Frassetto et al. (2008) interpreted
shear wave splitting results from the Batholiths data set as owing to possible toroidal flow
around the slab’s northern edge. A similar model has been proposed to explain shear
wave splitting observations near the southern edge of the subducting Gorda-Juan de Fuca
plate (Zandt and Humphreys, 2008) and the effect of upwelling mantle may increase
magmatism in the northern Basin and Range in a manner similar to that proposed here.
Isotopic evidence (Bevier, 1989) is indicative of an oceanic mantle source for the Anahim
volcanics, similar to that responsible for much of the Basin and Range magmatism.
The second major anomaly apparent in the Vs depth maps is the fast upper mantle
that occurs west of the CSZ down to depths of 70km. Particularly on the north line, the
upper mantle depth slices reveal a sharp velocity contrast juxtaposing higher velocities
west of the CSZ with velocities 4-5% lower to the east. Determining whether this
velocity contrast occurs at the CSZ or at the boundary between the Insular and
Intermontane superterranes, approximately 30 km to the west, is unfortunately beyond
the limits of our resolution. Sub-Moho velocities near station BN01 are approximately
4.49 +/- .08 km/sec, within the expected range of both a Sierra Nevada type garnet
pyroxenite and an average upper mantle peridotite. As such, we see two reasonable
interpretations for the high velocity anomaly to the west of the CSZ: it is either a
Jurassic/Cretaceous ultramafic root petrologically complementary to the overlying
batholith, or a layer of more typical mantle lithosphere attached to the Insular
103
superterrane that has been tectonically juxtaposed alongside the western flank of the
Intermontane superterrane.
To assess the first possible interpretation, we look for comparison to the southern
Sierra Nevadas, where a strong case has been made for the ongoing delamination of a
dense lithospheric root (Gilbert et al., 2007). Beneath the portion of the Sierra Nevada
batholith believed still to be in possession of its root, the thick crust is gravitationally
anchored, causing locally subdued surface elevations, and the Moho is nearly transparent
to receiver functions. Beneath station BN01, conversely, the crust is thin (~25 km), and
strong P-S conversions from the Moho indicate a sharp velocity contrast (Calkins et al.,
2008). While surface elevations are low beneath the western portion of our northern
transect, the thin crust implies, at least to first order, that the region is isostatically
balanced. Furthermore, a batholithic root would have to be extremely garnet rich and
therefore very dense, ~3.5 g/cm3, to achieve velocities approaching 4.5 km/sec (Jull and
Kelemen, 2001; Ducea, 2002). Residence times for such a dense root in a warm upper
mantle are estimated to be on the order of 10 Myr (Jull and Kelemen, 2001). Due to the
fact that the majority of the plutons west of the CSZ were emplaced prior to 85 Ma
(Gehrels et al., 2008), a gravitationally unstable ultramafic root would likely have
foundered into the deeper mantle long ago.
We therefore prefer a tectonic juxtaposition model. One such model is that
proposed by Monger et al. (1994) and supported by age dating results from the Batholiths
project (Gehrels et al., 2008). According to this model, the northwestern portion of the
Jurassic/early Cretaceous-age batholith was translated to the south along a sinistral strike
104
slip fault around 100Ma, effectively doubling portions of the same arc on a line
perpendicular to the strike of the trench. As subduction continued throughout the late
Cretaceous and early Tertiary, a back-arc basin separating the Insular and Intermontane
superterranes collapsed and the arc migrated eastward. Magmatism on the eastern flank
of the CMB, conversely, continued until ~48 Ma and overprinted the eastern portion of
the Jurassic/early Cretaceous arc (see Gehrels et al., 2008, fig 11). The Paleocene
tonalite sills and Eocene granodiorite plutons immediately east of the CSZ (Gehrels et al.,
2008; Rusmore et al., 2001, 2005) are particularly voluminous and have rare earth
element patterns that strongly suggest deep melting and dense root formation (Girardi et
al., 2007).
We add to this model the hypothesis that the root that formed beneath the
Jurassic/early Cretaceous arc foundered into the mantle sometime before ~85 Ma.
Continuing compression throughout the latest Mesozoic and Cenozoic in the forearc
region re-thickened the lithosphere to a depth of ~70 km. Studies of modern subduction
zones have imaged a “cold nose” of lithospheric mantle beneath the forearc that is largely
isolated from convection in the mantle wedge (Abers et al., 2006). The geometry of the
preserved lithosphere that we image agrees well with observed cold noses, both in terms
of thickness and distance from the trench. Beneath the eastern CMB, which would have
been exposed to actively convecting asthenosphere of the mantle wedge, several cycles of
root formation and foundering may have elapsed between accretion of the Insular
superterrane and the cessation of voluminous arc magmatism in the early Eocene (Girardi
et al., 2007). While the timing of the last cycle of root removal is unknown, any
105
lithosphere present beneath the eastern portion of the CMB is significantly younger and
thinner than that beneath the western portion.
We also image two minor velocity anomalies in the northern portion of the study
area. The first is a somewhat slower (~2-3%) than average region in the uppermost
mantle beneath the Bowser basin. The second anomaly is an isolated region of slightly
elevated velocities, up to 2.5% faster than average, between 70 and 110 km beneath our
northeastern-most stations. Due to their limited spatial extent and minor percent velocity
perturbation, we refrain from speculating about the origin of these anomalies.
Comparison of shear velocity variations in cross-section with the discontinuity
structure imaged by receiver function analysis (Calkins et al., 2008) reinforces the pattern
of eastward crustal thickening across the CSZ (figure 6). Positive (red) deflections in the
receiver functions, indicative of increasing velocity with depth, place the Moho
discontinuity very near the same depth as the increase in shear velocities from ~3.5
km/sec to ~4.2 km/sec. Some smearing in the vertical location of the velocity increase is
apparent in the cross-sectional shear velocity map and is due to the fixed depth of
velocity steps and damping of the solution relative to the starting model in our inversion.
Despite this smearing, the pattern of crustal thickness across the CMB correlates well
between the two independent techniques. Receiver functions from stations BN01 and
BS01, both located west of the CSZ, also show high amplitude negative arrivals near 100
and 60 km depth, respectively, which we interpret to originate at the
lithosphere/asthenosphere boundary. As station BS01 is only slightly closer to the
surface expression of the CSZ than is station BN01 (~35 km vs. ~41 km), the receiver
106
functions suggest a possible southward thinning of the western arc lithosphere. This
pattern is also apparent in the shear velocity maps in figure 5, and may be attributable to
the proximity of the active subduction zone to the south. Receiver function arrivals
originating in the upper mantle beneath stations east of the CSZ are of significantly lower
amplitude and may be attributable to small scale heterogeneities or noise in the data.
Conclusions
The Coast Mountains Batholith of British Columbia provides an ideal locale to study
the post-magmatic evolution of the lithosphere in a continental arc setting. Isotopic
evidence strongly suggests the formation of a dense ultramafic root beneath the granitoid
Jurassic to Eocence plutons within the Batholiths project study area (Girardi et al., 2007).
We have used phase and amplitude measurements of Rayleigh waves to probe the crust
and upper mantle in search of a currently resident ultramafic root or evidence of
detachment and foundering of the same. The major findings of our analysis are as
follows.
(1) The upper mantle beneath the central CMB is very slow relative to global models.
On avegage, the sub-CMB mantle is ~8% slower than AK135 at 200 km and 46.5% slower between 70 and 110km.
(2) A low Vs anomaly that appears to originate at a depth of at least 155 km roughly
beneath the town of Bella Coola trends to the northeast at shallower depths. We
posit that this low velocity zone may funnel fluids or partial melt flowing around
the edge of the subducting Juan de Fuca slab to the Anahim volcanic belt.
107
(3) Upper mantle shear wave velocities jump roughly 5% from east to west across the
Coast Shear Zone. West of the CSZ at 36-46 km, estimated upper mantle
velocities approach 4.5 km/sec. If this anomaly corresponds to an ultramafic root,
the high seismic velocity implies density greater than 3.5 g/cm3. We deem it
unlikely that such a dense residue would remain perched above the slow, warm
upper mantle imaged beneath the central CMB for ~100Myr since the cessation of
magmatism in the western arc. Our preferred interpretation for this anomaly is
therefore that it stems from ~70 km-thick forearc lithosphere that has been
tectonically abutted against the slower upper mantle to the east of the CSZ. The
fairly typical lithospheric velocities on the west contrast sharply with the low
upper mantle velocities east of the CSZ that are a signature of delamination of the
Paleocene/Eocene root and resulting infill by asthenosphere. Comparison of the
shear velocities with receiver stacked functions suggests this lithospheric layer
may thin to the south.
Acknowledgements
The authors would like to thank Don Forsyth, Yingjie Yang, and Aibing Li for sharing
their computer programs and experience related to surface wave processing. The
Batholiths project was funded under the U.S. National Science Foundation’s Continental
Dynamics program, grant EAR-0309885. The instruments used in the field program were
provided by the PASSCAL facility of the Incorporated Research Institutions for
Seismology (IRIS) through the PASSCAL Instrument Center at New Mexico Tech. Data
108
collected during this experiment will be available through the IRIS Data Management
Center. The facilities of the IRIS Consortium are supported by the National Science
Foundation under Cooperative Agreement EAR-0004370 and by the Department of
Energy National Nuclear Security Administration.
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Figures
Figure 1: Map of the study area highlighting seismic station locations (green stars) and
major tectonostratigraphic boundaries (after Gehrels et al., 2008). Underlying geology is
from the GSC Canadian Geologic map. The Anahim volcanic track is the E-W trending
series of roughly circular features located north of Anahim Lake and outlined in the
dashed black box. The light blue box on the inset Google Earth image shows the regional
location of the array. A=Alexander terrane, B=Bowser basin, CSZ=coast shear zone,
EBD=eastern boundary detachment fault, G=Gravina belt, S=Stikine terrane,
W=Wrangellia terrane, Y=Yukon Tanana terrane.
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Figure 2: Locations of earthquakes used in this study (red dots). The center of the
Batholiths array is marked by the black star; inner and outer concentric rings mark 30°
and 90°, respectively, from the array center.
114
Figure 3: Ray paths crossing the study area. Seismic station locations are marked with
black triangles outlined in grey.
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Figure 4: Average dispersion curve (a) and 1D shear velocity profile (b) for the central
Coast Mountains Batholith. For comparison, the global model AK135 (Kennett et al.,
1995) and the tectonic North America (TNA) model (Grand and Helmberger, 1984) are
also plotted on the Vs diagram.
116
117
Figure 5 (previous page): Vs at selected depth slices from the inversion. Panels a-e show
percent perturbation from the average Vs for the entire gridded area at each depth. The
dashed lines around the two transects outline the area contained by the 1.5 % error
contour, and the masked (white) region lies outside the 2.5 % error contour. The heavy
blue line marks the location of the Coast Shear Zone. Note the change in color palette
range between 70 and 110 km. Panel f is a color map of two times the standard error,
expressed as a percentage of phase velocity, for the 30 second Rayleigh wave data. The
1%, 2%, and 3% contours are plotted. Dashed white line A-A’ in panel f shows the
location of the cross section in figure 6.
118
Figure 6 Combined surface wave and receiver function results. Color map in the
background reflects absolute shear wave velocities along the line marked A-A’ (53.2° N)
in figure 5f. Stacked receiver functions from selected stations are overlain at their
approximate distance from the surface expression of the Coast Shear Zone (CSZ),
marked by the black triangle. Receiver functions are normalized such that the direct P
arrival (truncated) has unit amplitude. Positive (red) arrivals in the receiver functions
denote interfaces where seismic velocity increases with depth and negative (blue) arrivals
the opposite. See Calkins et al. (2008) for details of the receiver function analysis.
119
120
Supplementary Figure 1 (Previous page): Phase velocity maps for selected period
ranges
121
Supplementary Figure 2: Vertical Vs profiles at stations (a) BS07 and (b) BN01
compared with (c) the average 1D Vs profile for the entire study area. For each plot,
solutions from variably damped phase to Vs inversions are presented. Blue lines result
from the most lightly damped inversions, black lines from intermediate damping, and red
from high damping. See figure 1 for locations of seismic stations.
122
APPENDIX D: LITHOSPHERIC STRUCTURE OF THE
PAMPEAN FLAT SLAB REGION OF CHILE AND
ARGENTINA FROM ANALYSIS OF RAYLEIGH WAVE
PROPAGATION
Josh Calkins, Susan Beck, Aibing Li
Abstract
We image the crust and upper mantle of the Pampean flat slab region of Chile and
Western Argentina using the two plane wave approximation for surface wave
tomography. Analysis of Rayleigh wave data recorded by an array of broadband
instruments that spanned the horizontal section of the slab and a more normally dipping
(30°) portion of the subduction zone to the south reveals dramatic differences in upper
mantle shear velocities. Our results confirm the previously discovered zone of
anomalously high S-wave velocities above the flat slab and indicate that a similar
structure is absent above the steeper slab section to the south. The mantle wedge above
this southern section, near 36°S, is instead characterized by very low Vs (~4.2km/sec),
indicative of active mantle wedge convection. We image the base of the flattened section
of the slab near 150 km, suggesting total plate thickness of the order of 50km, but find no
evidence of the lithosphere-asthenosphere boundary beneath the steeply dipping portion
of the slab.
123
Introduction and Regional Geologic Setting
Between 29° and 32° south, the oceanic Nazca plate flattens near 100 km depth
beneath Western Argentina before steepening to a more typical subduction angle of ~2530° as it progresses eastward (Cahill and Isacks, 1992; Ramos et al., 2002; Anderson et
al., 2007; see Fig 1). The effects of this unusual pattern of subduction on the accreted
terranes that make up the overriding South American plate are dramatic, both in the
patterns of volcanism and topographic structure. Geologic studies find that magmatism
initially progressed eastward in response to the rising slab around 16 Ma, then ceased
altogether as the mantle wedge was cut off from asthenospheric flow ~5 Ma (Ramos et
al., 2002, Kay et al., 1988). Uplift of basement cored crustal blocks on thrust faults
lagged behind the advancing magmatic front by approximately 2 Myr (Ramos et al.,
2002). Deformation is ongoing in the Sierras Pampeanas, up to 750 km inboard of the
trench (Ramos et al., 2002). Due to their tectonic and morphologic similarities to the
Rocky Mountain foreland region of the Western United States, the Sierras Pampeanas
have been proposed as a modern analogue for the processes responsible for the basement
cored uplifts of the Laramide orogeny (Jordan and Allmendinger, 1986). While the root
cause of slab-flattening is not entirely understood (e.g. Gutscher et al., 2000) the spatial
correlation of the Pampean flat slab with the aseismic Juan Fernandez ridge and the
associated increased buoyancy of the downgoing plate are thought to be at least partially
responsible for the process (e.g. Ramos et al., 2002; Yanez et al., 2001; Anderson et al.,
2007).
124
Recent investigations of the lithospheric structure of the Pampean flat slab region
via broadband seismic observations have revealed a highly heterogeneous and atypical
upper mantle. The Chile Argentina Geophysical Experiment (CHARGE) targeted the
area between 30° and 36° S with a dense array of broadband seismic instruments in order
to image the morphology of the Nazca plate at depth and gain insight into the effects of
flat slab subduction on the overriding plate. Receiver functions show that the Moho
obtains a maximum depth of 60-70 km beneath the high topography of the Principal
Cordillera, but that the crust stays thick well to the east beneath the lower elevations of
the Sierras Pampeanas (Heit et al., 2008, Gilbert et al., 2006; Calkins et al., 2006).
Seismic body wave tomography from CHARGE data by Wagner et al. (2005, 2006)
revealed a high Vs, low Vp/Vs anomaly between the top of the flat slab and the base of
the crust. This anomaly was interpreted as a chunk of extremely depleted, cold
lithosphere, and forms the basis of a model whereby the asthenospheric mantle wedge
was completely squeezed out in advance of the shallowing subducting plate.
In this paper, we present our findings from inversion of Rayleigh wave data
recorded by the CHARGE array, a technique that allows us to extend seismic velocity
estimates beyond the region resolved by body wave tomography. Our results corroborate
the major features imaged by Wagner et al. (2005), correspond well with the crustal
thickness estimates obtained via Pn modeling (Fromm et al., 2004) and receiver function
analysis (Heit et al., 2008, Gilbert et al., 2006), and add constraints regarding the
thickness of the subducting plate and extent of the convecting mantle wedge.
125
Data and methodology
To investigate the lithospheric structure of the Pampean flat slab region, we
measure the phase and amplitude of dispersive Rayleigh waves and then invert for phase
and shear velocity. Our data set consists of earthquakes recorded primarily by the
CHARGE array supplemented by records from stations of the GEOSCOPE and Global
Seismograph Networks (Fig. 1). The CHARGE array consisted of 22 broadband
instruments deployed in two roughly east-wet transects across the Andes, with a few
stations filling the gap between lines, for 18 months beginning in late 2000. For
processing, we selected 27 large earthquakes with epicenters between 30° and 120° from
the center of the array and clearly visible fundamental mode Rayleigh wave energy on the
vertical component (Fig 2). Due to the position of the network in relation to patterns of
global seismicity, azimuthal coverage from the east side of the array is sparse. Large
events occurring on the Central American trench generally had high signal-to-noise
ratios, but excessively complicated waveforms owing to scattering along the continental
margin between source and receiver. In spite of these geographically-imposed
constraints, however, raypath coverage within the network is sufficiently dense to resolve
the major phase velocity variations (Fig. 3).
To invert for phase velocities (Fig 4), we employ the two plane wave (2PW)
approximation for Rayleigh wave propagation (Forsyth et al., 1998, Forsyth and Li,
2005). The 2PW method accounts for the effects of multi-pathing and scattering of
surface waves outside the area of interest by modeling the observed wave field as the sum
of two interfering plane waves. The two stage iterative inversion technique also has the
126
benefit of down-weighting those earthquakes that are not well modeled by the 2PW
approximation (Forsyth and Li, 2005; Yang and Forsyth, 2006). For the final step of
estimating shear wave velocity as a function of depth, we use the linearized, least squares
inverse method of Hermann et al. (1987) in the iterative approach described by Larson et
al. (2006). For a more complete discussion of our technique, see Calkins et al. (2008;
appendix 3 of this document) and references therein.
The shear velocity inversion is dependent on having a 1D Vs starting model that
is relatively close to the velocity structure within the study region, and a damping factor
in the inversion controls how far the solution is permitted to deviate from the starting
model. We calculated a starting model by first constructing a “best-guess” model based
on the crustal velocities resolved by Alvarado et al. (2007) and mantle velocities from the
global model AK135 (Kennett et al., 1995). We then performed a series of inversions
using the average dispersion curve for the study area and the “best guess” model,
updating the velocity model after each iteration with the output from the previous model.
After several iterations, each of which was run with the damping set to a lower value than
in the previous iteration, the output stabilized at a solution that we then used as the
starting model for each point in our grid. In anticipation of significant Vs heterogeneity
in the region (e.g. Baumont et al., 2002, Wagner et al., 2005), we opted for a lightly
damped inversion which allows our solution at each grid point to deviate significantly
from the starting model. Figure 5B shows a comparison of the moderately and heavily
damped average 1D Vs solutions.
127
Results and Discussion
As anticipated, phase velocities within the study region show significant lateral
variations that reflect the large scale morpho-tectonic features of the region. Figure 4
contains absolute phase velocity maps for five of the 14 period ranges at which we
measured phase and amplitude. At shorter periods (Fig. 4a), the low velocities (red) in
the thick crust of the Principal and Frontal Cordilleras contrast sharply with the higher
velocities (blue) of the surrounding upper mantle to the east and west. At longer periods
(Figures 4c-e), the sub-horizontal portion of the Nazca slab is evident as the broad high
velocity anomaly centered near 32°S, 68°W. The average dispersion curve and resulting
1D Vs model reveal a crust and uppermost mantle that are, on the whole, slightly slower
than global averages.
We restrict our shear velocity maps to the region that is well resolved in our
inversions (Figs. 6 and 7). At the margins of this region, which is based on the inversion
of the 50 second data set and bounded by the heavy black line in figure 4f, changes in Vs
greater than 1.75% are significant at the 95% confidence interval. Towards the center of
both transects, Vs variations as small as 1% are within the 95% confidence limit (see
Forsyth and Li, 2005, for a description of error estimation). We chose the 1.75% limit as
the minimum contour that allows us to gain some insight into the structures that span the
transition zone between the horizontal and steeply dipping sections of the Nazca plate.
In the mid to upper crust (Fig 6a), we see evidence of lower velocity material in
the south and east of the study region. To the south, low velocities are likely attributable
to Quaternary volcanism in the active arc above the normally dipping portion of the slab.
128
The low velocities near 67°W, 32°S are more difficult to explain, and do not correlate
well with velocities estimated by Alvarado et al. (2007) using regional body waveform
modeling. As the highest frequencies in our data set were often noisy and highly
attenuated, we have more confidence in deep crust and upper mantle Vs estimates which
are reflected in the longer period data. The low relative velocities in the 35, 45, and 49
km maps (Fig 6 c and d) in the region of the high Cordillera reflect the influence of
adjacent regions with thin crust. As our maps plot deviations from average velocities at
each depth slice, the low velocities beneath the high Cordillera stem from crustal material
occupying the depth range corresponding to the upper most mantle further east. Gilbert
et al. (2006) estimate crustal thickness of at least 60 km beneath the high Andes and in
the range of 40 km beneath the adjacent Sierras Pampeanas.
The most prominent feature in the upper mantle is the high velocity anomaly
above the flat slab in the 80-100km depth range (Fig. 7a, b). Shear velocities here are 46% higher than the surrounding mantle, an effect attributed by Wagner et al. (2008) to
high levels of orthopyroxene enrichment. While the velocities reported by Wagner et al.
(2006, 2008) are slightly higher than those in our solution, the general shape and location
of this high Vs anomaly are remarkably consistent. The difference in absolute Vs is
likely attributable to details of the inversion techniques used to analyze body waves
(Wagner) and surface waves (this study). Low velocities in the southern portion of the
study area at 100km (Fig 7b, 8b) depth reflect the presence of convecting mantle
asthenosphere above the steeply dipping portion of the slab.
129
Directly beneath the high Vs anomaly at depths greater than ~140km (Fig 7d-f),
shear velocities are lower than the surrounding region. We interpret this region of low Vs
to reflect the lithosphere/asthenosphere boundary (LAB) at the base of the Nazca plate.
This hypothesis is buttressed by the S-wave receiver functions analyzed by Heit et al.
(2008), which showed evidence of a velocity decrease at the same depth beneath the
northwestern CHARGE stations. We note, however, that the cross section through our
southern transect (Fig 8b) shows no consistent signature of the LAB and indeed no direct
evidence of the presence of the subducting plate. Our inability to image the slab may be
due to lateral smoothing of the phase velocity inversion using a Gaussian function with a
characteristic wavelength of 100km. This smoothing length is the smallest reasonable
value afforded by station spacing within the array, and likely results in horizontal
smearing of the higher velocities in the surrounding mantle across the expected narrow
low velocity signature of the slab’s interior. The plate’s steeper subduction angle in the
south is implied, however, by the very low velocity material near 80 km on the east end
of line B-B’, a feature which we interpret to be convecting asthenospheric mantle wedge.
The depth profile of this feature is very similar to a pervasive low velocity zone mapped
by Gilbert et al. (2006, Fig 5).
Conclusions
On the basis of the surface wave analysis presented herein, we conclude the
following:
130
(1) The shear velocity structure of the uppermost mantle beneath Chile and
Western Argentina between 30° and 36°S is highly heterogeneous, likely due
to the influence of variations in morphology of the subducting Nazca plate.
(2) Above the horizontal portion of the slab, we agree with previous findings of
unusually high shear wave velocities, in excess of 5% faster than the
surrounding upper mantle. This anomaly is noticeably absent above the
steeply dipping portion of the slab to the south.
(3) A region of low Vs approximately 40-50 km directly beneath the high
velocity anomaly likely indicates the lithosphere-asthenosphere boundary on
the downgoing Nazca plate.
(4) At 35.6° S, we image a zone of very low shear wave velocities (~4.2km/sec)
near 70-90 km which we interpret as convecting asthenospheric mantle.
Acknowledgements
Our thanks go to Don Forsyth and Yingjie Yang for sharing their computer code
and experience in analyzing the surface wave data. The National Science Foundation
supported this research under grants EAR9811870 and EAR0510966 and via support of
the PASSCAL facility of the Incorporated Research Institutions for Seismology (IRIS)
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Figures
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Figure 1 (previous page): Regional topographic map. Locations of seismic stations used
in this study are in white diamonds and quaternary volcanoes are denoted by black
triangles. Grey line marked “JFR” shows the projected down-dip extension of the Juan
Fernandez Ridge. Thin black contours map the depth to the top of the Nazca plate as
determined by Anderson et al (2007). Thin blue line shows the approximate location of
the Cuyania/Pampia terrane border (Ramos et al., 2002). Lines A-A’ and B-B’ mark the
locations of cross sections in Fig. 8.
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Figure2: Locations of CHARGE array (white star) and earthquakes used in this study
(red circles).
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Figure 3: Example ray coverage plot showing the source-receiver great circle paths for
the 50 second data set. Seismic station locations are marked by grey triangles
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Figure 4 (previous page): Phase velocity maps and error plot. Panes a-e show absolute
phase velocity at selected period ranges. Pane f shows 2 x standard error, expressed as a
percent of phase velocity, for the 50 sec. period range. The heavy black contour outlines
the area inside which velocity variations greater than 1.75% are significant at the 95%
confidence interval.
139
Figure 5: Average (a) dispersion curve and (b) 1D Vs profile for the study area.
140
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Figure 6 (previous page): Shear velocity maps at selected depths within the crust and
uppermost mantle. The color palette in each pane expresses Vs in terms of % deviation
from the average at that depth within the study area, and labeled thin black contours show
absolute Vs in km/sec. Note the change in color palette range between the 49km (e) and
69km (f) maps. Velocities are masked at the contour defined in figure 4.
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Figure 7 (previous page): Shear velocity maps at selected depths within the mantle. The
color palette in each pane expresses Vs in terms of % deviation from the average at that
depth within the study area, and labeled thin black contours show absolute Vs in km/sec.
Note the change in color palette range between the 80 km (a) and 100km (b) maps.
Velocities are masked at the contour defined in figure 4.
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Figure 8: Cross sections at (a) 30.4 °S and (b) 35.6°S. See figure 1 for location of cross
sections. White dashed line shows approximate location of the Moho based on the
receiver function analysis of Gilbert et al. (2006). Heavy black line shows the
approximate location of the top of the Nazca slab based on earthquake relocations of
Anderson et al. (2007). Dashed black line marks 40 km below the slab surface. The color
palette in each cross section represents % deviation from the average Vs at each depth
within the study area, and labeled thin black contours show absolute Vs in km/sec.