i
University of Nevada
Reno
Geophysical Constraints on Seismic Hazard and Tectonics in the
Western Basin and Range
A dissertation submitted in partial fulfillment of the requirements for
the degree of Doctor of Philosophy in Geophysics
by
Robert E. Abbott
Dr. John N. Louie/Dissertation Advisor
December 2001
ii
Signature Page
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Abstract
Three studies show the effectiveness of shallow exploration geophysics in solving
problems related to tectonics and seismic hazard. In Chapter 2, 200 new gravity
measurements are used to find depth-to-bedrock in the Reno and Carson City areas.
Depth-to-bedrock and shallow shear-wave velocities have been shown to have an effect
on ground motion from earthquakes, prompting the study. Maximum basin depth in Reno
is 1.2 km and maximum depth in Carson City is 0.53 km. Quaternary fill might only
rarely exceed 200 m. An unexpected and previously unsampled gravity trough in Reno
may increase seismic hazard there relative to other parts of the basins. In Chapter 3, I use
a combination of seismic and gravity techniques to prove that the 1954, Ms 6.8, Dixie
Valley, Nevada, earthquake produced slip on a low-angle (<30°) normal fault. The fault
maintains a low-angle and slightly listric geometry to the depth of constraint (1.75 km
seismic; 2.7 km gravity). As such, the Dixie Valley event may represent the first large,
low-angle normal earthquake on land recorded historically. This result may increase the
awareness of the potential hazard posed by low-angle normal faults. In Chapter 4 I
conclude, from the study of weak ground motion spectral ratios and in situ velocity
characterization, that the continued existence of precarious rocks in their untoppled state
near the San Andreas fault is not due to local site or path effects. Velocity
characterization, using a combination of refraction-microtremor and seismic tomography
methods, shows that the precarious rock sites are slower than the National Earthquake
Hazards Reduction Program BC boundary from 0 to 10 m, and much faster from 10 to 45
m. As a result, synthetic spectral ratios of small earthquakes at the precarious rock sites
iv
relative to the BC boundary show de-amplification below 5 Hz and amplification above 5
Hz. One possible conclusion is that the current seismic hazard relations overestimate the
hazard near large strike-slip earthquakes.
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Acknowledgements
Thanks above all goes to my parents, who have been most patient and supportive over the
many years of schooling. I greatly appreciate the help offered by my committee and coauthors. John Louie, in particular, was extremely supportive and was everything a good
advisor should be. Thanks to the many great friends and fellow graduate students that
helped keep me sane (on average), especially: Aline Concha-Dimas, Mark Engle, Keith
Weaver, Ana Cadena, Nancy Glenn, Mandy Johnson, and Mike Ossofsky.
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Table of Contents
Title page ........................................................................................................ i
Signature Page .............................................................................................. ii
Abstract......................................................................................................... iii
Acknowledgements ........................................................................................v
Table of Contents ......................................................................................... vi
List of Tables ................................................................................................ xi
List of Figures ............................................................................................. xii
Chapter 1: General Introduction .................................................................1
1.0 A NOTE ON ORGANIZATION ....................................................................................... 1
1.1 GENERAL INTRODUCTION .......................................................................................... 1
Chapter 2: Depth to bedrock using gravimetry in the Reno and Carson
City, Nevada, area basins ..............................................................................5
2.0 ABSTRACT.................................................................................................................. 5
2.1 INTRODUCTION........................................................................................................... 7
2.1.1 Geologic Setting ................................................................................................. 8
2.1.2 Seismic Hazard ................................................................................................ 13
vii
2.2 METHODS ................................................................................................................. 13
2.2.1 Discussion on Error ......................................................................................... 17
2.3 RESULTS .................................................................................................................. 18
2.4 DISCUSSION AND CONCLUSIONS .............................................................................. 30
2.4.1 Constraints From a Previous Carson City Study ............................................ 32
2.4.2 Constraints From Well Data ............................................................................ 33
2.4.3 Origin of the Deep Sub-Basin .......................................................................... 35
2.4.4 Seismic Hazard Estimates ................................................................................ 37
2.5 ACKNOWLEDGEMENTS ............................................................................................. 39
Chapter 3: Geophysical confirmation of low-angle normal slip on the
historically active Dixie Valley fault, Nevada ...........................................40
3.0 ABSTRACT................................................................................................................ 40
3.1 INTRODUCTION......................................................................................................... 42
3.1.1 Regional and Tectonic Setting ......................................................................... 43
3.1.2 The 1954 Dixie Valley Earthquake .................................................................. 43
3.1.3 Geologic Evidence for Low-Angle Dip on the Dixie Valley Fault................... 45
3.1.4 Previous Geophysical Work in Dixie Valley .................................................... 46
3.2 METHODS ................................................................................................................. 49
3.2.1 Field Data Acquisition ..................................................................................... 49
3.2.2 Seismic Data Processing.................................................................................. 50
3.2.2.1 Conventional processing and poststack migration .............................................................. 50
3.2.2.2 Velocity analysis ................................................................................................................. 51
viii
3.2.2.3 Prestack migration ............................................................................................................... 52
3.2.2.4 Gravity data processing ....................................................................................................... 54
3.3 RESULTS .................................................................................................................. 54
3.3.1 High-Resolution Line ....................................................................................... 54
3.3.2 Medium-Resolution Line .................................................................................. 56
3.3.2.1 Fault refraction .................................................................................................................... 56
3.3.2.2 Migrations ........................................................................................................................... 59
3.3.3 Gravity Results ................................................................................................. 63
3.4 DISCUSSION ............................................................................................................. 65
3.4.1 Timing of Extension ......................................................................................... 66
3.4.2 Magnitude of Slip ............................................................................................. 68
3.5 CONCLUSIONS .......................................................................................................... 71
3.6 ACKNOWLEDGMENTS ............................................................................................... 72
Chapter 4: Weak ground motion and amplifications predicted from
shear-wave velocities at precarious rocks, near the 1857 rupture of the
San Andreas fault ........................................................................................73
4.0 ABSTRACT................................................................................................................ 73
4.1 INTRODUCTION......................................................................................................... 75
4.2 DATA AND METHODS ............................................................................................... 77
4.2.1 Earthquake Data .............................................................................................. 77
4.2.2 Spectra Calculation ......................................................................................... 81
4.2.3 Spectral Ratios ................................................................................................. 83
4.2.4 Velocity Data ................................................................................................... 83
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4.2.5 Calculation of Synthetics ................................................................................. 84
4.3 RESULTS .................................................................................................................. 85
4.3.1 Local Shear-Velocity Structure ........................................................................ 85
4.3.2 Horizontal and Vertical Spectra ...................................................................... 90
4.3.3 Spectral Ratios ................................................................................................. 92
4.3.4 Synthetic Spectral Ratios with Piute Butte....................................................... 94
4.3.5 Synthetic Spectral Ratios with Two Reference Models .................................... 95
4.4 DISCUSSION AND CONCLUSIONS .............................................................................. 97
4.5 ACKNOWLEDGEMENTS ............................................................................................. 99
Chapter 5: Recommendations for Future Work ...................................100
5.0 ORGANIZATION ...................................................................................................... 100
5.1 RECOMMENDATIONS FOR FUTURE RENO-CARSON CITY BASINS WORK ................ 100
5.1.1 Increased Gravity Coverage .......................................................................... 100
5.1.2 Density Measurements ................................................................................... 101
5.1.3 Finite Difference Modeling ............................................................................ 101
5.1.4 Test Ground Motion Predictions Using New ANSS Stations ......................... 102
5.1.5 Shallow Velocity Measurements .................................................................... 102
5.2 RECOMMENDATIONS FOR CONTINUED DIXIE VALLEY RESEARCH ......................... 103
5.2.1 Static and Dynamic Stress Modeling ............................................................. 103
5.2.1.1 Differentiating among the possible mechanisms for fault weakening. .............................. 104
5.2.2 Electromagnetic Study of the Fault Zone....................................................... 105
5.2.2.1 Dixie Valley Evidence Supporting an Elevated Pore-Fluid Model ................................... 106
5.2.3 Gravity Survey to Study Low-Angle to High-Angle Accommodation Zones.. 109
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5.2.3.1 Survey Plan ....................................................................................................................... 111
5.2.3.1 Modeling and Interpretation .............................................................................................. 112
5.3 RECOMMENDATIONS FOR FUTURE WORK IN THE PRECARIOUS ROCK STUDY ........ 114
5.3.1 Optimize the ReMi equipment ........................................................................ 114
5.3.2 Optimize for 30-m Work................................................................................. 115
5.3.3 More ReMi Data in Varied Environments ..................................................... 117
5.3.4 Finite Differencing ......................................................................................... 118
5.3.5 Measurements at Precarious Rocks in Different Environments .................... 118
References ...................................................................................................119
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List of Tables
Table 2.1 Selected Well Data from the Reno Area Basin .......................11
Table 2.2 Density Contrast vs. Depth for Depth-To-Bedrock
Calculations ..................................................................................................16
Table 4.1 Station Locations........................................................................79
Table 4.2 Event Parameters .......................................................................80
Table 4.3 Average 30-m Shear-Wave Velocities ......................................89
xii
List of Figures
Figure 2.1 Landsat TM over shaded relief map of Reno and Carson
City basins and surrounding area ................................................................9
Figure 2.2 Generalized geologic map of Reno and Carson City basins
and surrounding area ..................................................................................12
Figure 2.3 Reno-Truckee Meadows complete Bouguer anomaly map..19
Figure 2.4 Reno-Truckee Meadows bedrock gravity ..............................20
Figure 2.5 Reno-Truckee Meadows residual basin gravity ....................21
Figure 2.6 Depth-to-bedrock in Reno-Truckee Meadows area..............22
Figure 2.7 Depth-to-bedrock in Reno-Truckee Meadows area and well
data overlay ..................................................................................................24
Figure 2.8 Reno-Truckee Meadows 2.5 –D gravity transect ..................25
Figure 2.9 Carson City-Eagle Valley complete Bouguer anomaly ........26
Figure 2.10 Carson City-Eagle Valley bedrock gravity ..........................27
Figure 2.11 Carson City-Eagle Valley residual basin gravity ................28
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Figure 2.12 Depth-to-bedrock in Carson-City Eagle Valley ..................29
Figure 2.13 Carson City-Eagle Valley 2.5-D gravity transect ................31
Figure 3.1 Map of Dixie Valley and surrounding areas ..........................44
Figure 3.2 Complete Bouguer anomaly map of Dixie Valley .................47
Figure 3.3 Map showing geophysical transects........................................48
Figure 3.4 CDP stack of high-resolution data ..........................................55
Figure 3.5 Raw shot-gather showing simultaneous refraction arrival ..57
Figure 3.6 Optimaztion velocity profile ....................................................58
Figure 3.7 Stolt-migrated stack .................................................................60
Figure 3.8 Kirchoff prestack migration....................................................62
Figure 3.9 2.5-D gravity model of Dixie Valley ........................................64
Figure 3.10 Schematic diagram of fault activity and geomoetry since
late-Oligocene-early Miocene .....................................................................69
Figure 4.1 Basemap showing our site locations .......................................77
Figure 4.2 Sample seismogram ..................................................................82
xiv
Figure 4.3 Shear velocity models ...............................................................86
Figure 4.4 Velocity frequency plots ..........................................................87
Figure 4.5 P-wave velocity profiles ...........................................................88
Figure 4.6 Horizontal and vertical component amplitude spectra .......91
Figure 4.7 Earthquake and synthetic spectral rations relative to Piute
Butte ..............................................................................................................93
Figure 4.8 Synthetic spectral ratios relative to NEHRP BC boundary
and generic rock...........................................................................................96
Figure 5.1 Two possible mecahnisms for creating intra-basin gravity
lows ..............................................................................................................110
Figure 5.2 Bouguer gravity map of Dixie Valley with proposed
traansects ....................................................................................................113
Figure 5.3 Velocity-frequency plot of Lovejoy Buttes with dispersion
curve from tomography velocity model ...................................................116
1
Chapter 1: General Introduction
1.0 A Note on Organization
This dissertation is a collection of three loosely related geophysical investigations.
Chapter 1 is this introduction. Chapters 2, 3, and 4 are reformatted versions of three
publications. Chapter 2 is published in Geophysics [Abbott and Louie, 2000]; Chapter 3
is published in the Journal of Geophysical Research [Abbott et al., 2001a]; and Chapter 4
is, as of this writing, in submission to the Bulletin of the Seismological Society of
America [Abbott et al., 2001b]. Chapters 2 and 3 are presented with formatting changes
only, and the respective journal articles are essentially identical to the chapters herein.
The reader may find it worthwhile to look for the BSSA version of Chapter 4 for more
information, however. Chapter 5 contains recommendations for future work that will only
be found in this dissertation. A supplemental CD-ROM is included that contains digital
copies of all figures in PostScript and PDF formats, including color for Figures 2.1, 2.3-6,
2.9-12, 4.4-5, and 5.3.
1.1 General Introduction
Although the proposals and research plans for the three projects presented in this
dissertation were not written with each other in mind, the projects have much in common.
It is what they have in common that drew my interest and kept my attention over the
years creating this work. Although the three field areas can be loosely grouped in or near
the Basin and Range province of the western United States, it is not their location that is
the uniting thread. Nor is it the precise methodology of investigation used at each site, for
the individual techniques are quite different. Rather, the three projects are united in that
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they show that the cheap and effective tools of exploration geophysics, properly applied,
can answer very important and fundamental questions about the nature of tectonics and
its potential impact on human beings.
Recent studies on how earthquake waves propagate in basins [e.g. Olsen, 1995]
provided strong motivation for the study in Chapter 2. In that Chapter, I present a study in
which I used variations in the earth’s gravitational acceleration in and around Reno and
Carson City, Nevada, to discover the thickness of sediments in the basins. The
interpretation of the complex gravity field was simplified by removing the known gravity
effect of earth tides, latitude and longitude, elevation, etc., and making the very simple
assumption that the remaining anomaly was due to changes in rock density below the
measuring sites. Because basin fill is less dense than basement rock, the areal distribution
of the gravity anomaly as measured at the surface revealed the subsurface basin
geometry.
Advances in computing power in the past two decades has made the study of the
propagation of earthquake waves through strong velocity contrasts feasible for everyone.
Although no simulations of seismic waves through the Reno and Carson City basins have
been made as of yet, my basin mapping may reveal basin geometries that have a
significant effect on the seismic hazard of the region. I predict that basin studies of this
sort will become increasingly common as more and more urban areas feel the need for
increased knowledge of potentially increased earthquake damage within heavily
populated areas.
The results of Chapter 3, although achieved using the methods more commonly
associated with resource exploration, have far-reaching implications for tectonics,
3
seismic hazard, and fault mechanics. I used seismic reflection, seismic refraction,
tomography, and gravity techniques to prove that the Dixie Valley fault is characterized
by subsurface dips as low as 25° at the surface. This result bolsters the argument of many
geologists who have argued for the existence of many low-angle normal faults that have
accommodated large amounts of extension in the upper crust. The results also weaken the
arguments of many fault mechanics researcher who have claimed that low-angle slip is
impossible due to frictional constraints. This study shows that low-angle slip must be
occasionally possible (at least in the upper 2 km) and future research should focus on
mechanisms that make low-angle extension possible.
The results of Chapter 3 also have seismic hazard implications, especially if the
model proposed in Wenicke [1992] is correct. In the 1992 review paper, Wernicke
derived using simple relations, an equation that suggest that the recurrence interval of
normal-mechanism earthquakes as a function of dip θ is proportional to tan(θ)sin(θ). This
relation predicts that low-angle normal earthquakes account for less than 10% of the
number of earthquakes on faults dipping 30° to 60°. However, given the large down-dip
dimension of low-angle faults (i.e. 15 km, the width of a standard seismogenic zone,
divided by the sine of the dip), and the inherit strength of the fault geometry due to
increased normal stress, earthquakes on low-angle faults can pose a serious seismic
hazard [Wernicke, 1992].. This coupled with the fact that low-angle faults have more
surface area near the surface of the earth, could mean that many urban areas have underrecognized faults posing an increased hazard.
Chapter 4 shows again how small-scale investigations can have large tectonic and
seismic hazard implications. In contrast to Chapter 2, where I investigated the possibility
4
that conditions in the basin amplify seismic waves, I now investigate the possibility that
local de-amplification of seismic waves affects the overall hazard. The scale of the two
investigations bears mentioning. In Chapter 2, large-scale structures, such as the basinbedrock interface, several hundred meters distant, serve to focus waves, increasing the
hazard. In Chapter 4 I discuss the seismic hazard implications of geology less than 100 m
from the measuring station. In this case, the hazard posed by large strike-slip faults was
studied. Motivated by the existence of precarious rocks (which Los Angeles Times
Magazine once termed “The Seismometers of the Gods”) within sight of the San Andreas
fault, I used seismic refraction-microtremor (ReMi) and tomographic techniques to
determine if local effects were responsible for the apparent decreased ground motions at
the sites. Ground motion generated by an earthquake is a convolution of source, path, and
site effects. By using the above techniques and small magnitude earthquake spectral
ratios, I conclude that site and path effects at the precarious rock sites are minimal, and
that current regressions appear to overestimate the hazard inherit to large strike-slip
events. To the millions of people living near the San Andreas, North Anatolian, and other
large strike-slip faults this issue is of paramount importance.
5
Chapter 2: Depth to bedrock using gravimetry in the
Reno and Carson City, Nevada, area basins
Robert E. Abbott and John N. Louie
The University of Nevada, Reno, Seismological Laboratory and Dept. of Geological
Sciences, Reno, Nevada 89557
2.0 Abstract
Sedimentary basins can trap earthquake surface waves and amplify the magnitude
and lengthen the duration of seismic shaking at the surface. Poor existing gravity and
well-data coverage of the basins below the rapidly growing Reno and Carson City urban
areas of western Nevada prompted us to collect 200 new gravity measurements. By
classifying all new and existing gravity locations as on seismic bedrock or in a basin, we
separate the basins' gravity signature from variable background bedrock gravity fields.
We find an unexpected 1.2 km maximum depth trough below the western side of Reno;
basin enhancement of the seismic shaking hazard would be greatest in this area. Depths
throughout most of the rest of the Truckee Meadows basin below Reno are less than 0.5
km. The Eagle Valley basin below Carson City has a 0.53 km maximum depth. Basin
depth estimates in Reno are consistent with depths-to-bedrock in the few available
records of geothermal wells, and one wildcat oil well. Depths in Carson City are
consistent with depths from existing seismic reflection soundings. The well and seismic
correlations allow us to refine our assumed density contrasts. The basin to bedrock
density contrast in Reno and Carson City may be as low as -0.33 g/cm3. The log of the oil
well, on the deepest Reno sub-basin, indicates that Quaternary deposits are not unusually
6
thick there and suggests that the sub-basin formed entirely before the mid-Pliocene.
Thickness of Quaternary fill, also of importance for determining seismic hazard, below
Reno and Carson City may only rarely exceed 200 m.
7
2.1 Introduction
Alluvial basins can amplify the magnitude and lengthen the duration of seismic
waves. In Mexico City, for example, "basin site effects" due to waves trapped in the lowvelocity basin are cited as a primary reason for disastrously high ground motion in the
great 1985 Michoacan earthquake [Campillo et al., 1989; Sanchez-Sesma et al., 1989].
Kawase and Aki [1989] showed that both basin shape (i.e. depth-to-bedrock) and velocity
contrasts within the alluvium were essential parameters needed to model ground motion
in the Mexico City basin. Frankel and Vidale [1992] used water well depth-to-bedrock
data to create a 3-D simulation of seismic waves in the Santa Clara, California, basin.
Efforts to predict ground motions in basins in the Salt Lake City, Utah, and Los Angeles,
California, areas have required knowledge of sediment thickness as well as bedrock
topography [Olsen et al., 1995; Frankel, 1993]. The highest ground motion amplification
in Olsen's simulations of Salt Lake City occurred near the edges of the deepest portions
of the basin (rather than directly over the deepest portion), where the depth gradient was
steepest. In these areas, Olsen states that particle motion can be up to 2.9 times higher
than in bedrock stations. In addition, the duration of the seismic signal is up to 40 times
longer [Olsen et al., 1995]. Although many additional factors must be considered to
estimate seismic hazard, to model wave propagation in western Nevada population
centers, accurate basin models are essential. It is with this in mind that we undertook a
detailed gravity survey of the urban centers of Reno and Carson City, Nevada.
The density contrast between bedrock and unconsolidated or poorly consolidated
sedimentary rocks allows the study of bedrock structures underlying sedimentary basins.
With good gravity data coverage, only changes in rock density affect the shape of any
8
gravity anomaly. Basin shape and depth can be inferred from the spatial distribution of
the anomaly. Examples of this general technique can be found in West [1992, p. 200209]. Schaefer [1983] modeled the Dixie Valley, Nevada, basin using a similar
technique; many researchers have used this method for hydrologic, geothermal, mineral,
and oil exploration. Jachens and Moring [1990] mapped Cenozoic thickness across
Nevada with this principle.
2.1.1 Geologic Setting
Our study area is situated along the western edge of the Basin and Range geologic
province, in the western United States. The cities of Reno and Carson City lie within the
fault-bounded basins of the Truckee Meadows and Eagle Valley, respectively (Figure
2.1). The two basins are bordered on the west by the Carson Range of the Sierra Nevada
Mountains and on the east by parts of the Pah Rah and Virginia Ranges, and the Pine Nut
Mountains. The Carson Range is predominantly Mesozoic granite with older
metamorphic rocks, and the other ranges generally consist of Tertiary volcanic rocks. The
basin fill consists of Quaternary and Tertiary alluvial and lacustrine deposits, and
outwash from the most recent glacial epochs [Bell et al., 1989]. A significant portion of
the Truckee Meadows is underlain by low density diatomaceous sediments.
The subsurface geology in this region is poorly understood. Existing gravity
coverage, as compiled from the 1994 NOAA Gravity CD-ROM [Hittelman et al., 1994],
is too sparse to adequately resolve basin structure. Thompson and Sandberg [1958]
conducted a gravity survey of the Virginia City, Nevada, and Mt. Rose, Nevada,
quadrangles in 1952. However the average station spacing of 1 station per 2 mi2 (5 km2)
was inadequate to characterize the basins. Very few of the existing gravity measurements
9
10
were made over the basins. As a result, basin details are not revealed in the Bouguer
anomaly gravity maps of Plouff [1992] and Saltus and Jachens [1995].
Hess [1996], Garside and Schilling [1979], and associated, recently updated
databases of geothermal and oil wells present some information on 56 boreholes that are
over 150 m deep in the Reno basin. Table 2.1 summarizes data from a few of these that
we use to constrain our basin models. We selected the 26 wells for Table 2.1 because we
could find some minimal location and total depth information for them. In those cases
where a group of wells are on the same property or in very close proximity to one
another, we only list the deepest well and/or the well with the best logs of the group. All
but a few are clumped in the 5 km2 "Moana Hot Springs" area on the southwest side of
the basin. Four deep geothermal wells there logged bedrock at more than 300 m depth.
The bedrock there is the Tertiary Kate Peak formation andesite (part of “Consolidated
Basement Rocks”, Figure 2.2.) Garside and Schilling [1979] report a single wildcat oil
well in the Reno-Truckee Meadows basin. The well was drilled in 1908 on the western
side of the basin, with a log interpreted by Anderson [1910]. The 1890 ft (576 m) hole
encountered only sedimentary rocks, giving a minimum basin thickness in that area. The
other wells outside these limited areas are domestic water wells with only total depths
known. These provide some corroborating minimum basin depth constraints. Seismic
studies of Reno basin velocities are underway, but will not be adequate for describing
bedrock geometry.
Arteaga [1986] mapped depth-to-bedrock in Eagle Valley using a combination of
seismic reflection, seismic refraction, and gravity techniques. Their gravity results
11
12
13
provide independent corroboration of our technique, and their seismic depth soundings
allow for more accurate density calibration.
2.1.2 Seismic Hazard
The seismic hazard of western Nevada is high, with many faults capable of
producing magnitude 7 and greater earthquakes [dePolo et al., 1996]. USGS seismic
hazard maps [Frankel et al., 1996] do not include any evaluation of basin amplification
effects. They show that for both Reno and Carson City there is a 2% probability of
ground motions exceeding 0.6 g in the next 50 years.
As of 1995, approximately 400,000 people live in the Reno, Carson City and
surrounding areas. A hypothetical magnitude 7+ earthquake would represent a
tremendous potential for loss of life and property. Identifying those areas susceptible to
greatest ground movement would be of use to emergency planning personnel.
2.2 Methods
We made approximately 200 gravity measurements with a LaCoste and Romberg
model G gravity meter. The measurements generally follow north-south or east-west
roads in the urban Reno and Carson City, Nevada, areas (Figure 2.2). In Reno, vertical
control was provided by a geodetic-quality GPS. In Carson City, an electronic distancemeasuring theodolite was used for vertical control. The surveys were tied to international
gravity (IGSN 1971) at a gravity base station in Reno (ACIC 0454-1). Local base stations
were re-occupied on a regular basis to monitor tidal variations to gravity as well as
instrument-related drift. Terrain corrections (using 2.67 g/cm3) were estimated by eye in
the field from 1 meter to 54 meters horizontally (Hammer zones B-C) and computed by
algorithm from 54 m to 167 km, using 90-meter digital elevation models. The data were
14
reduced to complete Bouguer anomaly using a reduction density of 2.67 g/cm3. The
curvature correction to the Bouguer slab equation was applied when calculating terrain
corrections beyond 18 km.
Existing gravity coverage [Hittelman et al., 1994] was merged into the dataset to
complete our coverage, since we took fewer measurements outside the basins. The terrain
corrections from 54 m to 167 km were re-computed and re-applied to the existing data,
along with the curvature correction. The total coverage included 600 points.
To differentiate gravity effects due to small-scale basins from broader, regional
anomalies, a "bedrock gravity" value was removed from the data set. Following Jachens
and Moring [1990], all gravity stations are classified as "bedrock" or "basin" stations
using geologic maps [Bonham and Rogers, 1983; Trexler, 1977; Bell and Garside, 1987;
Bonham and Bingler, 1973]. We considered measurements on or near Tertiary Kate Peak
formation andesitic rocks (part of “Consolidated Basement Rocks”, Figure 2.2) to be on
bedrock for our purpose of differentiating low-seismic-velocity sedimentary fill from
relatively high-velocity seismic bedrock. Similarly, points on or near Tertiary Hunter
Creek formation sandstones (Figure 2.2) were considered to be basin fill. Density
measurements by Thompson and Sandberg [1958] indicate that the density of the Kate
Peak formation averages around 2.61 g/cm3. A single density measurement on the Hunter
Creek formation (Truckee formation of Thompson and Sandberg [1958]) indicates a
density of 1.76 g/cm3, although there is evidence from well-log data [Anderson, 1910]
and this study to indicate that the density of this formation varies widely.
Bedrock gravity values were computed by kriging the complete Bouguer anomaly of
those gravity stations known to be in areas where basin fill is minimal, or non-existent.
15
The bedrock gravity is subtracted from the complete Bouguer anomaly of all
measurement points. By removing the perturbations to gravity caused by bedrock density
contrasts, basin structure is emphasized and the gravity effect of deep density variations
below the surrounding mountain ranges is attenuated. The gravity effect of a basin
extends beyond the basin boundaries, however, and these are subtracted as part of our
“bedrock gravity” estimate. Thus, basin depths subsequently estimated will be minima.
Unlike Jachens and Moring [1990], initial basin depth estimates were accomplished
using the infinite slab approximation. We simply scale the basin gravity anomaly value at
each measurement point by a factor that assumes the anomaly results from one or more
slabs of constant density contrast and infinite lateral extent, to find the total sediment
thickness. This estimate is similar to reversing the Bouguer slab calculation, and produces
a smoothed basin-depth profile, with the deepest depths being underestimated.
We initially used the sediment compaction model given in Table 2.2 to find the
alluvium-basement density contrast. The sediment compaction model is the same used by
Blakely et al. [1998] and Jachens and Moring [1990], and represents a regional average
for basins within the Basin and Range province. Depth was calculated by applying the
slab approximation for the shallowest (-0.65 g/cm3, 200 m) slab. If the gravity anomaly
caused by this slab is less than the observed anomaly, deeper layers were taken into
account. It should be noted that the infinite slab approximation works best when the slab
thickness is much less than the lateral extent of the basin. Errors in depth calculations can
occur when nearing the basin edge, where this approximation fails.
We also forward modeled 2.5-dimensional selected linear transects in Reno and
Carson City using the GM-SYS software package, developed by Northwest Geophysical
16
17
Associates. Sediment “blocks” were modeled as extending 3.5 km north and south of the
Truckee River east-west transect in Reno . In Carson City, sediment blocks extend 2 km
north and 6 km south of the 5th Street east-west transect . Care was taken that the
transects followed the trend of the measuring stations as closely as possible. Information
from well data (in Reno) and seismic data (in Carson City) were used to constrain
parameters in basin modeling. In Carson City, due to the complete lack of local density
and lithology information, we made use of average regional density contrasts (Table 2.2).
2.2.1 Discussion on Error
The repeat error of LaCoste and Romberg gravity measurements is estimated to be
0.03 mGal. This is higher than would be expected if the measurements were taken in a
quiet environment under controlled conditions. However, most measurements were taken
along busy urban streets where traffic and other urban vibrations caused measurement
errors. Base stations were carefully chosen to be in quiet, controlled environments. For
those measurements, a repeat error of 0.01 mGal is estimated. GPS and EDM theodolite
locations, accurate to plus or minus one meter, allow us to neglect latitude correction
errors. Vertical position is accurate to within 0.3 meters, as confirmed by GPS or EDM
theodolite re-occupations of sites. Inner-ring terrain corrections, estimated by eye, rarely
approached 0.1 mGal and were 0.01 mGal on average. Still, in areas of high relief, a 20%
error in estimating inner ring terrain effects is possible. In these rare instances, an error of
0.02 mGal could have been introduced. All considered, an error in observed gravity of
plus or minus 0.08 mGal is possible. Measurements of repeated points from different
surveys in the existing data [Hittelman et al., 1994] exhibit a maximum error of plus or
minus 0.5 mGal. This is the limiting factor in the dataset. Given the magnitude of the
18
anomaly in Reno (15-20 mGal) and the coarse contour interval, we view this as an
acceptable amount of error and that the benefits of their inclusion outweigh the problems
caused by decrease in accuracy. The depth error in the infinite slab approximation caused
by a 0.5 mGal error is 36 m using a –0.33 g/cm3 bedrock-alluvium density contrast. In the
forward models, our tolerance fit levels mean that the 0.5 mGal error between
measurement campaigns is essentially invisible.
Our density approximations are the principal source of error in our analysis. The well
logs available in Reno lack density measurements or analyses. This lack of density data
leads to highly speculative density values. With upper and lower limits for basin-bedrock
density contrast set at 0.65 g/cm3 and 0.30 g/cm3, a 50% depth error is conceivable.
2.3 Results
Several products result from our data analysis: (1) Complete Bouguer anomaly maps
derived from all stations; (2) Complete Bouguer anomaly maps derived from bedrock
stations; (3) Basin anomaly gravity maps; (4) Basin depth maps derived from the infinite
slab approximation; and (5) 2.5-D forward models of selected linear transects.
The anomaly maps of Reno (Figures 2.3, 2.4, and 2.5) show an extended,
asymmetrical gravity low over the Truckee Meadows. The gravity low represents the
density contrast of bedrock and sediments. The western side of the basin shows the
steepest gravity gradients and the most negative anomaly. The maximum local anomaly
of -16 mGal yields a basin depth of 1160 m using the infinite slab approximation (Figure
2.6) if we assume an alluvium-bedrock density contrast of –0.33 g/cm3. There is evidence
that the residual gravity separation may not have completely succeeded near this subbasin. Figure 2.4, the bedrock gravity grid, also shows a gravity low over this area. Note
19
20
21
22
23
that the infinite slab approximation underestimates basin depth for a given density
contrast. Because we used well-log information to calibrate depth at certain areas, the
density contrast required in the infinite slab approximation was underestimated to
compensate. Sediment density is likely to be less than the 2.34 g/cm3 we used.
The western-most elongation of the basin represents the east-west trending Tertiary
Verdi basin. This basin is underlain by the Miocene-Pliocene Hunter Creek sandstone
formation. The sandstone has a lower average density than the alluvium of the Truckee
Meadows. As such, the depth of the basin may be slightly shallower than indicated on the
depth-to-bedrock maps (Figures 2.6 and 2.7). A sub-basin in the Steamboat Springs area
is represented by another gravity low to the southwest, with -6 mGal local anomaly,
corresponding to a depth of approximately 430 meters.
The east-west cross-section along the Truckee River in Reno (Figure 2.8) yields a
maximum basin depth of 1000 m. This profile shows a striking structural trough in the
western portion of the basin. The maximum basin depth in this model is under West
McCarran Boulevard. A second trough in the eastern portion of the basin is separated
from the western trough by a bedrock ridge that comes within 200 meters of the basin
surface near the Reno/Tahoe International Airport.
The anomaly maps of Carson City (Figures 2.9, 2.10, and 2.11) show an elongate
north-south gravity low over Eagle Valley. The anomaly closely approximates the
anomaly shape of Ateaga [1986], which he mapped using a combination of gravity and
seismic techniques. The magnitude of the local anomaly, -7 mGal, is much smaller than
in the Truckee Meadows, suggesting a shallower basin depth. This corresponds to a 520
m depth with the infinite plate approximation (Figure 2.12, assuming a –0.33 g/cm3
24
25
26
27
28
29
30
density contrast). The northeast-trending contours in the northern part of the basin are
poorly constrained and may be an artifact of the poor data coverage in the area. A subbasin to the northwest is separated from the main basin by the subsurface expression of a
northwest-southeast trending ridge of Triassic metavolcanic rocks. This formation
outcrops at Lone Mountain in northern Carson City [Trexler, 1977].
An east-west cross-section along 5th Street in Carson City (Figure 2.13) yields a
maximum basin depth of 530 m. The 5th Street transect shows Eagle Valley to be a more
symmetrical basin in which the depth increases fairly smoothly to 0.53 km before
returning to bedrock on either side of the basin. The density scheme used is from Table
2.2. The maximum basin depth along this transect is located 1.5 km east of US Highway
395 (Figure 2.13).
2.4 Discussion and Conclusions
Contours indicating negative bedrock depths outside basins allow the estimation of
errors in depth-to-bedrock calculations caused by shallower bedrock density contrasts.
The +2 mGal contour on the western margin of Eagle Valley (Figure 2.10) would
correspond to a -140 meter depth-to-bedrock with the infinite slab calculation (Figure
2.11), where negative depth would mean bedrock above actual elevation. Therefore, our
estimates of bedrock gravity (Figures 2.3 and 2.9) may be in error by 2 mGal, and depth
to bedrock in the basins cannot be considered more accurate than 140 meters. The cause
for this is unclear, but the poor coverage of bedrock gravity measurements, isolated
bedrock density variations such as hidden intrusions, measurement errors, or isostatic
effects near the Sierra Nevada Mts. are possible. Based on our Reno bedrock gravity
estimate (Figure 2.4), our basin gravity difference (Figure 2.5), and basin depth (Figures
31
32
2.6 and 2.7) maps, we estimate a depth uncertainty of 250 meters for the Truckee
Meadows.
To constrain absolute depths, accurate density measurements need to be obtained.
Specific knowledge of how density increases with depth, especially from outside the
geothermal fields, would be particularly useful. We can only use the depth to bedrock
logged in a few of the wells to check our overall density assumptions. In particular, the
thickness of low-density diatomaceous deposits in the Hunter Creek formation varies
widely from location to location. Currently, density uncertainty is the overriding cause of
depth uncertainty. Using our reasonable end-member values for density contrasts, 50%
error in depth calculation is possible, if no other factors, such as seismic depth soundings,
were taken into account. The error in our analysis is probably significantly less than this
maximum value, however.
2.4.1 Constraints From a Previous Carson City Study
Arteaga’s [1986] hydrological study of Eagle Valley included some depth-tobedrock calculations based mostly on seismic depth soundings, supplemented by gravity
measurements. Absolute comparison of gravity values is impossible because the study
did not publish complete Bouguer values. The only published result was the gravity
residual, obtained by subtracting out the regional gradient. The shape of gravity residual
obtained by Arteaga matched extremely well with our basin anomaly residual.
Arteaga’s seismic depth soundings predicted a 620 m maximum basin depth,
compared to our 530 m maximum. The data quality of the seismic depth soundings where
charaterized as only being “fair” in Arteaga’s [1986] study. Overall agreement in depthto-bedrock is generally within approximately 25%. The location of the three
33
measurements closest to the deeper sub-basins is plotted on Figure 2.12, along with the
associated depths in meters.
2.4.2 Constraints From Well Data
Garside and Schilling [1979] and more recent associated databases [e.g. Hess, 1996]
provide some well data that corroborate our basin depths for Reno (Figures 2.6, 2.7, and
2.8; Table 2.1). Neither of these latest databases show any boreholes of record in Carson
City-Eagle Valley. The well casing from the 1908 deep oil prospect [Anderson, 1910]
was located by the Nevada Bureau of Mines and Geology before housing development
overtook the site. The location is thus known to within 30 m and appears near the 800 m
basin thickness contour from the slab calculations on Figure 2.7 (labeled 1). Anderson
[1910] interpreted all but the top few meters of the 576 m total depth drilled as partly
penetrating the middle-Tertiary "Truckee formation" of sands, shales, and diatomites. The
formation is equivalent to the Miocene-Pliocene "sandstone of Hunter Creek" mapped in
the area by Bonham and Rogers [1983]. This deepest boring into basin deposits in Reno
is only 1.5 km south of our lowest basin gravity anomaly. Garside and Schilling [1979]
and the more recent records show an additional 55 wells drilled to depths greater than
150 m in the Reno basin. Table 2.1 lists the 25 of these that provide the best constraints
throughout the basin. Locations of most of these wells are given by partial sections or
permittee addresses, and could easily be 300 m in error. The 30 wells not in Table 2.1
either lacked any reliable depth or location information, or were not as deep or as welllogged as another well on the same property or very nearby. Of the 55 wells, only twelve
are outside the immediate area of the "Moana Hot Springs" geothermal district (the
34
concentration of well locations at the lower center of Figure 2.7). Eleven of these appear
on Figure 2.7; the twelfth lies off the map to the south.
South of the 200 m depth contour, along South McCarran Blvd., Figure 2.7 shows
basin thicknesses of 100 m or less, corroborated by domestic wells such as the Talsma
and the Peterson (labeled 20 and 18 on the Figure), which logged Kate Peak volcanic
bedrock at 49 m and 91 m, respectively. Depths increase rapidly to the north and to the
northwest, in concert with the depth contours derived from gravity. The Pennington
domestic and Warren Estate Geothermal Well No. 3 (labeled 19 and 24 on Figure 2.7)
logged Kate Peak at 283 and 317 m, respectively. The Pennington well is located near the
400 m depth contours on the Figure. The Warren Estate #3 is located by partial sections,
where the first quarter-section may be stated in error in the database; its location on
Figure 2.7 may be 1 km southeast of the true location.
At the north end of the Moana Hot Springs area, the deep Kohlenberg Domestic
Injection Well No. 1, Peppermill Well No. 4, and Salem Plaza Injection Well No. 1
(labeled 16,15, and 21 on Figure 2.7) log Kate Peak depths increasing from 310 m, to 344
m, and to 418 m, respectively, from east to west over a 1 km distance across US 395
Business. The Peppermill No. 4 well is the deepest hole in the region, with logs to 1008
m (n.b.: logged lateral deviation of the hole is less than 100m). These bedrock depths
closely approximate the depth contours on Figure 2.7.
Every logged well in Reno shows Hunter Creek sandstones and diatomites extending
through 20%-90% of the section above the Kate Peak bedrock, averaging about 80%
(Table 2.1). Anderson [1910] summarized observations of similar diatomites throughout
the Great Basin, and proposed that every Miocene basin in the region might host them.
35
The diatomites consist of an open network of silica microfossils having porosities as high
as 70%, and air-filled samples from the surface near the oil well (labeled 1 on Figure 2.7)
will float on water (mentioned by Anderson, 1910]. These diatomites, with pores filled
with water, would have a maximum density of 1.7 g/cm3. It is likely that the diatomites
have less porosity at depth and/or have pore spaces filled with mineralization, thus
increasing the formation’s density.
In Reno outside the Moana Hot Springs district, the 12 deeper wells listed by Garside
and Schilling [1979] have total depths but no logs on file. Figure 2.7 locates them with
white circles, having ‘≥’ marked above their total depths. Given that these are all
domestic water sources, it seems unlikely that they would have been drilled far into any
bedrock formations, although that possibility cannot be ruled out. Assuming that these
wells provide minimum basin thicknesses from their total depths, one of them (labeled 3,
on Figure 2.7) does not match the slab-gravity-derived depth contours. The other water
wells for which there are no logs are all shallower than or close to the infinite-slab depth
contours. The "negative depth" contours on Figure 2.7 surrounding the interchange
suggest bedrock density contrasts that are not constrained by the bedrock data, and have
artificially pushed the depths to more shallow levels. A 198 m well drilled in 1958
(labeled 11 on Figure 2.7) in a sub-basin to the southeast, is very close to bedrock
outcrops and may have been drilled partially into bedrock.
2.4.3 Origin of the Deep Sub-Basin
A novel result of our work is our definition of the 16 mGal gravity low on the west
side of Reno. This anomaly, about 7 km in diameter, had not been sampled by the
previously very sparse gravity coverage of the basin. Within the limits of our data
36
coverage, Figures 2.5, 2.6, and 2.7 show that this low defines a north-south trending
trough about 5 km long and 3 km wide, and up to 1.2 km deep, that we call the West
McCarran sub-basin. It should be noted that the extent of the sub-basin is poorly
constrained to the southwest. The sub-basin is twice as deep as any other sub-basin below
Reno or Carson City, and identifies the location of what could be the largest basin effects
on earthquake ground motion in the western Nevada urban areas.
The 576 m well drilled in 1908 shows that this deep sub-basin was formed entirely in
Miocene and Pliocene time. The well sits nearly atop the deepest part of the basin
(labeled 1 in Figure 2.7), only 1.5 km south of the line of depth section modeled in Figure
2.8. Anderson [1910] mapped the fully exposed Truckee formation (Hunter Creek)
section, inspected the well during drilling, and interpreted the driller's log. He proposed
that the entire borehole had penetrated just the middle diatomite-dominated member of
the Hunter Creek sandstones, perhaps with some sampling of the lower sandy member.
Given the exposed 200 m thickness of each of these members in sections compiled
throughout the region, and the lesser thicknesses in the Moana Hot Springs well logs
[Garside and Schilling, 1979, and recent associated databases], the Hunter Creek
sandstones appear to be thickened in the deep sub-basin by a factor of two or three. This
thickening suggests the sub-basin was actively subsiding during deposition of the
diatomite member near the Miocene-Pliocene boundary (age taken from Bonham and
Rogers, 1983], and probably initiated during deposition of Hunter Creek basal members,
or earlier.
Since the entire 1.2 km-deep sub-basin is filled by early Pliocene and older
sediments, all of the related, basin-forming vertical deformation must have occurred by
37
the early Pliocene. Thus the existence of the West McCarran sub-basin requires no
Quaternary deformation. The 274 m thickness of Quaternary deposits logged by the
driller in the Peppermill Well No. 4 (Table 2.1, number 15) may be overstated. Reinterpretation of this log, and comparison to nearby logs by the authors suggest only 168
m of Quaternary fill above the Hunter Creek sands. This depth thus represents the
maximum observed Quaternary vertical deformation in the Reno basin. All other logged
wells in Table 2.1 show less, between 30 and 165 m. While the total thickness of lowvelocity Miocene through Quaternary sediment varies greatly among the Reno subbasins, the maximum thicknesses of Quaternary deposits in these basins may well be less
than 200 m.
2.4.4 Seismic Hazard Estimates
Our less than 30% error in depth calculation (based on negative depth contours)
should have little effect on seismic modeling, except possibly in one key area. Currently,
modeling of seismic waves in basins is rarely done for frequencies greater than 1 Hz. Any
change in depth equal to or less than one-quarter seismic wavelength may be
undetectable. Work in progress, not presented here, by the University of Nevada, Reno
Seismological Laboratory, Kyoto University, and the Shimizu Corporation of Japan is
employing the microtremor analysis of Horike [1985]. Shear wave velocities for depths
below 100 m at a test site near the Reno/Tahoe Airport (Figure 2.7), are on the order of 2
km/s. At 1 Hz, this would correspond to a 2 km seismic wavelength. Therefore 500 meter
resolution is required for seismic hazard modeling. Only in the deepest section of the
Truckee Meadows would a 30% maximum error in depth even approach this limit. In the
shallower sections of both basins, even 30% depth error will be much less than 500
38
meters, and will have a smaller effect on a smaller seismic hazard. Therefore, the depth
error in the shallow sections may be insignificant. It is the deepest, most poorlycharacterized sub-basin that has the most seismic hazard potential, and this is the area
with the most error.
Depth gradient maps, not presented here, suggest in the manner of Olsen et al. [1995]
that the areas that could most experience basin effects might be near the north and south
corners of West McCarran Blvd. (Figure 2.7). In these areas over the western sub-basin,
the combination of a deep basin and high gradients may produce the ground motions
most amplified by surface waves trapped in the basin. High gradients also exist at the
eastern edge of the Truckee Meadows. However, the basin is not as deep in this area, and
therefore ground motion amplification may be less. Eagle Valley, with a very muted
basin structure as compared to the Truckee Meadows, should show less ground
amplification due to basin site effects altogether. The difference in site effects might be
of significant amplitude; and seismic hazard maps may have to be re-evaluated for these
two areas. The seismic hazard for Reno may increase with respect to the seismic hazard
of Carson City.
39
2.5 Acknowledgements
The authors would like to thank C. Mann, J. Ollerton, and M. Herrick for their
tireless and conscientious fieldwork. Thanks to Kennecott Exploration Company for
generously loaning their geodetic-level GPS, and to J. McKinney and C. Lide of Zonge
Geoscience Inc. for their help in applying terrain corrections to the data. We also thank P.
Cashman and J. Trexler for their thought provoking insights. R. Blakely and P. Milligan
provided careful and thorough reviews; J.G. Anderson, G. Biasi, and A. Cadena helped
refine the manuscript. R. Abbott was supported for this work by a Nevada Seismological
Laboratory Fellowship. Research supported by the U.S. Geological Survey (USGS),
Department of the Interior, under USGS award number 1434-HQ-97-GR-03041. The
views and conclusions contained in this document are those of the authors and should not
be interpreted as necessarily representing the official policies, either express or implied,
of the U.S. Government.
40
Chapter 3: Geophysical confirmation of low-angle
normal slip on the historically active Dixie Valley fault,
Nevada
Robert E. Abbott and John N. Louie
Seismological Laboratory and Department of Geological Sciences, University of Nevada,
Reno, Nevada
S. John Caskey
Department of Geosciences, San Francisco State University, San Francisco, California
Satish Pullammanappallil
Optim L.L.C., Seismological Laboratory, University of Nevada, Reno, Nevada
3.0 Abstract
The December 16, 1954, Dixie Valley earthquake (MS=6.8) followed the nearby
Fairview Peak earthquake (MS=7.2) by 4 min, 20 s. Waveforms from the Fairview Peak
event contaminate those from the Dixie Valley event, making accurate fault plane
solutions impossible. A recent geologic study of surface rupture characteristics in
southern Dixie Valley suggests that the Dixie Valley fault is low angle (<30) along a
significant portion of the 1954 rupture. To extend these observations into the subsurface,
we conducted a seismic reflection and gravity experiment. Our results show that a portion
of the Dixie Valley ruptures occurred along a fault dipping 25° to 30°. As such, the Dixie
Valley event may represent the first large, low-angle normal earthquake on land recorded
historically. Our high-resolution seismic reflection profile images the rupture plane from
41
5 to 50 m depth. Medium-resolution reflections, as well as refraction velocities, show a
smoothly dipping fault plane from 50 to 500 m depth. Stratigraphic truncations and
rollovers in the hanging wall show a slightly listric fault to 2 km depth. Gravity profiles
conservatively constrain maximum basin depth and define overall geometry. Extension
along the low-angle section may have occurred in two phases during the Cenozoic.
Current fault motion postdates a 13 to 15 Ma basalt, imaged in the hanging wall, and
inherits from a fault formed during an earlier extensional pulse, concentrated at 24.2 to
24.4 Ma. The earlier extension suggests extraordinary slip rates as high as 18 mm/yr,
resulting in the formation of the low-angle fault break. Sections of the Dixie Valley fault
where there is no evidence for current low-angle slip correlate well with areas where no
pre-15 Ma slip has been documented.
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3.1 Introduction
Despite growing geological and geophysical evidence arguing for the existence of
low-angle normal faults that have accommodated large amounts of extension, the paradox
of the near-complete absence of low-angle normal-mechanism earthquakes in the seismic
record remains. Slip on low-angle normal faults is not predicted in Andersonian theory
[Anderson, 1942] and studies of earthquake focal mechanisms, both regional [Doser and
Smith, 1989] and global [Jackson, 1987], show an extreme scarcity of large (M>5.5)
normal fault mechanisms with dip <30°. However, several researchers [e.g., Abers et al.,
1997; Hatzfeld et al., 2000; Johnson and Loy, 1992] have presented compelling evidence
that substantial extension has occurred along low-angle normal faults in the brittle upper
crust.
Theories to resolve the seismicity paradox fall into two categories: those that do not
require brittle slip at low angles (e.g., “rolling hinge” models [Wernicke and Axen, 1988]
and flexural rotation [Buck, 1988]), and those that argue either for long earthquake
recurrence intervals [Wernicke, 1995], or a current rarity of active low-angle faults
[Burchfiel et al., 1992].
Compelling evidence that brittle slip is possible on at least one low-angle normal
fault would have important ramifications for both fault mechanics theory and seismic
hazard calculations. Here, we present the results of a seismic reflection and gravity
experiment to test whether part of the December 16, 1954, Dixie Valley earthquake
(MS=6.8) produced slip on a low-angle normal fault.
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3.1.1 Regional and Tectonic Setting
Dixie Valley, Nevada, lies in the north central portion of the Basin and Range
province (Figure 3.1). The Basin and Range is a region that has experienced a large
amount of intraplate extension in the Cenozoic. Much of the extension is accomplished
by high-angle (50°-70°) normal faulting, with several large seismic events recorded
historically (e.g., 1915 Pleasant Valley, 1954 Fairview Peak-Dixie Valley, 1983 Borah
Peak). The faulting has created predominately north-south trending mountain ranges and
sedimentary basins. Dixie Valley is one such basin, bounded by the Stillwater Range on
the west and the Clan Alpine Range on the east (Figure 3.1). The Dixie Valley fault, site
of the 1954 Dixie Valley earthquake, is the east dipping range-bounding normal fault
along the eastern front of the Stillwater Range.
The 1954 fault trace lies along the southern portion of the Stillwater Range and
separates Mesozoic and Tertiary footwall rocks from late Tertiary and Quaternary basin
fill. Miocene and Oligocene volcanic rocks and granitic plutons related to the Stillwater
caldera complex [John, 1995] and Mesozoic metasedimentary rocks represent the
“geophysical basement.” The basin fill at the surface consists of alluvial fan and
lacustrine deposits [Wilden and Speed, 1974].
3.1.2 The 1954 Dixie Valley Earthquake
The December 16, 1954, Dixie Valley earthquake was the last of a series of large
earthquakes that took place within a period of 6 months in central Nevada. The preceding
events were the Rainbow Mountain sequence (MS=6.6 and 6.4 on July 6, 1954, MS=6.8
on August 24, 1954) and the Fairview Peak earthquake (MS=7.2 on December 16, 1954).
The Fairview Peak event preceded the MS=6.8 Dixie Valley earthquake by 4 min and 20
44
45
s. Focal mechanisms for the Fairview Peak and Rainbow Mountain events indicate NNW
striking normal-oblique faults with dips ranging from 60 to 78°.
Fault plane solutions for the Dixie Valley event are poorly constrained because the
arrivals are obscured by waveforms from the Fairview Peak event. Doser [1986] used
waveform modeling to determine fault geometry with a strike of N10°W and a dip of
60°E; however, owing to contamination of the Dixie Valley waveforms “large changes in
strike and dip did not significantly change the waveform shape” [Doser, 1986, p.12,583].
Similarly, Okaya and Thompson’s [1985] solution of N11°W, 62°E cannot be relied
upon. Okaya and Thompson [1985, p.116] noted that “Of the four focal parameters
(depth, dip, strike, and slip), only changes in depth are significant; changes in fault plane
strike, dip, or slip have negligible effect.” The Dixie Valley fault plane solution (N8°E,
49°E) of Hodgkinson et al. [1996] using leveling and triangulation benchmarks suffers
from a paucity of data (very few prerupture survey bench marks) in the rupture region,
such that the triangulation network is unable to document slip along most of the rupture.
3.1.3 Geologic Evidence for Low-Angle Dip on the Dixie Valley Fault
Caskey et al. [1996] conducted the most recent and detailed study of the surfacefaulting characteristics of the Fairview Peak and Dixie Valley earthquakes. They noted
substantial geologic evidence for low-angle dip for the Dixie Valley fault along an ~20km-long portion of the rupture zone. Geologic evidence for a low dip angle lies between
The Bend in the north, to just north of Coyote Canyon in the south (Figure 3.1).
Geological evidence for low-angle dip at the surface from Caskey et al. [1996] include
(1) three-point fault plane reconstructions, (2) shallow subsurface modeling of the
46
rupture-trace graben, (3) fault-parallel fracture sets in the footwall, and, (4) geometry of
the Stillwater range front.
3.1.4 Previous Geophysical Work in Dixie Valley
Dixie Valley has been the subject of numerous geophysical investigations. The
studies primarily focused on the northern part of the valley, around 40 km north of our
study area. Okaya and Thompson [1985] combine seismic reflection and gravity data to
model northern Dixie Valley as a half graben with one major normal fault dipping 50°E
to the northwest (the Dixie Valley fault), and three smaller, west dipping normal faults to
the southeast. This model is inconsistent with Smith’s [1967] aeromagnetic study north of
The Bend (Figure 3.1). Smith maps pre-Tertiary basement under Dixie Valley as a graben
within a graben.
Schaefer [1983] collected widespread gravity data throughout Dixie Valley. A
portion of Schaefer’s [1983] Bouguer anomaly map is reproduced in Figure 3.2. As can
be seen in Figure 3.2, two 4-5 mGal local gravity lows are near the latitudes of Coyote
Canyon and Wood Canyon, consistent with a more moderately dipping range front fault
(or uplifted bedrock) between these two latitudes. It should be noted that the shape of
these anomalies is very poorly constrained and the gravity lows may have other
explanations unrelated to a change in range front fault dip. An east-west linear transect of
24 gravity points near the latitude of Little Box Canyon (Figure 3.3) is consistent with a
low-angle Dixie Valley fault, assuming reasonable bedrock-alluvium density contrast.
Meister [1967] and Herring [1967] conducted seismic refraction experiments near
the latitude of IXL Canyon (Figure 3.3). Meister [1967] interpreted southern Dixie Valley
as a composite graben, based on several short refraction lines parallel to the rangefront,
47
48
49
and an east-west cross-valley profile. The east dipping Dixie Valley fault was interpreted
to be a combination of high-angle normal faults and flat terraces, resulting in a “staircaselike” fault geometry. These data allow for an alternate interpretation, however. In our
evaluation of Meister’s [1967] data a single low-angle dip normal fault can replace the
previous stair-step structure. Herring’s [1967] experiment assumed, a priori, high-angle
dip and the assumed geometry was used to test a “sideswipe-refraction” technique.
3.2 Methods
3.2.1 Field Data Acquisition
Figure 3.3 shows the location of our geophysical transects. The Cattle Road profile
consists of two seismic reflection profiles and a gravity profile, while the Scarp profile
consists solely of gravity coverage. The trend of the Cattle Road profile is within 10 of
the gradient of the gravity field near the range front. If the gravity gradient reflects the
direction of true dip of the Dixie Valley fault, then apparent dips measured from the
geophysical profiles will be underestimated by no more than 0.5. The mediumresolution Cattle Road profile extended 3.6 km into the basin and utilized 8-Hz
geophones. It was composed of four stationary setups of 48 receivers with 15.2-m
spacing. The 132 source points were 2-7 kg of high explosive buried 2 m below the
surface. Off-end and longer offset shots were recorded to increase fold and improve deep
velocity information. Maximum source-receiver offset for the medium-resolution line
was 2.8 km, and maximum fold was 24.
The high-resolution Cattle Road profile overlapped the medium-resolution profile
close to the 1954 rupture. It was conducted within 150 m of the range front scarp using
50
100-Hz geophone groups with 2-m spacing. Sixty-seven sledge-hammer source points
were rolled through the array. Six inline geophones per group were used to reduce noise
from ground roll. Maximum source-receiver offset for the high-resolution line was 124
m. Maximum receiver fold was 32.
Gravity data along Cattle Road were acquired with a LaCoste and Romberg Model G
gravity meter. Gravity coverage started at the scarp and extended eastward 12.5 km into
the basin at a 250-m average station spacing. Vertical control was supplied by a geodetic
quality GPS.
Gravity data were also collected parallel to the rangefront scarp (Scarp gravity,
Figure 3.3). The Scarp profile was located along a line where depth-to-bedrock was
assumed to be approximately constant, and therefore any gravity variations would be
largely due to density variations within the geophysical basement. In this way, errors in
depth to bedrock calculations from our 2.5-D forward model along the Cattle Road
profile can be estimated.
3.2.2 Seismic Data Processing
3.2.2.1 Conventional processing and poststack migration
Conventional seismic data-processing techniques were used to remove noisy traces,
mute direct waves, and attenuate other unwanted arrivals. The medium-resolution line
was band-pass-filtered (6-8 Hz, 100-120 Hz trapezoidal filter), and then filtered with a
polygonal, 48 trace, 500 ms f-k domain filter. The f-k filter was designed to eliminate 400
m/s Rayleigh waves. After automatic gain control, constant velocity stacks at 200 m/s
intervals were used to pick stacking velocities. Common depth points were binned at
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15.2-m intervals, with no amplitude variation with fold. The subsequent common depth
point stack was then Stolt-migrated at a constant 2000 m/s velocity.
The high-resolution seismic line was band-pass-filtered (8-12 Hz, 200-202 Hz) and
f-k domain filtered (48 trace, 125 ms). Stacking velocities were chosen using the same
method as with the medium-resolution line. Common depth points were binned at 2-m
intervals, and the resulting CDP stack was Stolt-migrated using the root-mean-square
stacking velocities.
3.2.2.2 Velocity analysis
Owing to complex geometry, strong lateral velocity variation, and steep dips (for
seismic imaging) in the subsurface at Dixie Valley, we chose to compute prestack depth
migrations. For input into our prestack migration, we obtained a detailed velocity image
of the subsurface by performing a nonlinear optimization on first arrivals picked off raw
shot gathers. The optimization technique employs a generalized simulated annealing
algorithm [Pullammanappallil and Louie, 1994] to invert first arrivals for subsurface
velocity structure. We used a commercial package, SeisOpt @2D™ (copyright Optim
LLC, 1998-1999), that implements this method.
The simulated annealing algorithm is a Monte Carlo based estimation process that
has the property of being independent of the starting model and has the ability to find the
global minimum (i.e., solution) for a highly nonlinear problem. These characteristics
make the algorithm a very effective tool for velocity estimation. Travel time inversion is
a highly nonlinear problem because any perturbation in the velocities alters the path of
the ray propagation, changing the travel times recorded at the surface geophones. This
nonlinearity makes linear methods dependent on the starting model; that is, the accuracy
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of the final velocity model is dependent on a good initial guess. The method employed by
SeisOpt @2D “tests” several thousand models before arriving at the optimized velocity
model. The only inputs required by the algorithm are the first arrival picks and survey
geometry (source and receiver coordinates). In addition to the final velocity model,
SeisOpt @2D outputs a ray coverage or “hit count” plot that shows what parts of the
model were sampled by the seismic array. The algorithm outputs only the velocities in
the subsurface that have been sampled by the rays.
A total of 6117 first-arrival picks from 134 shot gathers were used for the
optimization. SeisOpt @2D can handle only two-dimensional array geometry. Hence, in
order to overcome a bend in the medium-resolution seismic line, we project the source
locations to a straight line while maintaining the true offsets of the source-receiver pairs.
As a result of this projection, the optimized velocities might show some lateral smearing
in the vicinity of the bend in the profile. The resulting velocity model was used in a
prestack migration algorithm to image the seismic reflectors directly.
3.2.2.3 Prestack migration
In order to use the optimized velocities to perform a prestack Kirchhoff migration,
we first extended the velocities down to a depth of 2.0 km. Like Pullammanappallil and
Louie [1994] and Chavez-Perez et al. [1998] we extended the optimized velocity models
for migration by finding the maximum constrained velocity value in each column of the
velocity model and substituted that value into the column of the model everywhere below
the depth of the maximum velocity. We then performed a severe lateral smoothing below
the depth of first arrival constraint.
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The resulting velocity model was used in a prestack migration algorithm to image the
seismic reflectors directly. The prestack Kirchhoff summation algorithm was originally
used to image the San Andreas fault zone by Louie et al. [1988] and modified by Louie
and Qin [1991] to account for reflection ray paths that may bend significantly through
strong lateral velocity variations. In Dixie Valley we did not attempt to image nearvertical structure as previous work did, but we needed to account for strong ray bending
through the velocity contrasts at the edge of the basin and through any lateral variability
in the Tertiary basalts within the basin section. So we added to the algorithm a dipdependent obliquity factor.
Another addition is the migration-operator antialiasing criterion of Lumley et al.
[1994], which leads to high-cut filtering of the seismic traces. Completely preventing the
spatial aliasing of the migration operator leads to discontinuous coverage of depth points
for our medium-resolution survey and detracts from the lateral continuity of reflections.
Thus, for the aliasing calculation we used a receiver spacing that is half of the actual
15.2-m spacing. This apparent spacing yields a mild operator antialiasing effect that only
removes the worst aliased frequencies of the most steeply-dipping structure, while
retaining the lateral continuity of the unaliased near-horizontal structure.
We migrated the f-k filtered data also used for stacking, with additional 8-100 Hz
band-pass filtering, locations projected to a straight line, and AGC with a 0.5-s window
for amplitude balancing. Along with migration of the data we estimated noise data in the
manner of Harlan et al. [1984] by resampling to destroy prestack trace-to-trace
coherency. Harlan et al. [1984] provided a Bayesian prestack coherency measure
computed through statistical comparison of the data migration against the noise
54
migration. We screened the data migration through the coherency image to yield an
enhanced structural image. The enhancement emphasizes those structures that produce
the most coherent reflections in the prestack shot gathers.
3.2.2.4 Gravity data processing
We reduced gravity data to simple Bouguer anomaly using standard techniques.
Terrain corrections out to 54 m (Hammer rings A-C) were estimated by eye in the field.
We used a density value of 2.67 g/cm3 for both the Bouguer slab correction and for
terrain correction. Elevations are accurate to within 5 cm, as seen in site reoccupations.
Analysis of loop closures indicates a maximum drift of 0.16 mGal, forming the limiting
error of these data. This amount of error is acceptable, given the magnitude of the basin
anomaly (~35 mGal).
3.3 Results
3.3.1 High-Resolution Line
Figure 3.4 is a no vertical exaggeration common-depth-point stack of our highresolution line near the 1954 scarp. The Dixie Valley fault is the very prominent reflector
dipping to the east from 40 to 160 ms, at 28°. This is interpreted as the bedrock-alluvium
contact. The fault surface is ~6 m below the surface at the range front scarp. This depth to
fault is nearly identical to that predicted by the balanced geological cross section of
Caskey et al. [1996, Figure 12d]. The balanced section is reproduced at the location of the
graben in Figure 3.4. There is no evidence for staircase-like fault geometry at this scale.
Surface-parallel reflections are seen early in the section (above 60 ms) and represent
layering in the alluvial fan above the fault. The coherency of these reflections is disturbed
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in the vicinity of the 1954 scarp (at 25 m, 30 ms), possibly in relation to the formation of
the rupture graben.
3.3.2 Medium-Resolution Line
3.3.2.1 Fault refraction
Fault geometry can be inferred from arrivals seen in several raw shot gathers. The
headwave arriving at the same time at all stations in the gather of Figure 3.5 limits the
range of possible fault dips. If we use our optimization velocity model (Figure 3.6) as a
guide, maximum and minimum bedrock-to-alluvium velocity ratios can be placed at 2.1:1
and 1.4:1. After adjusting for elevation changes along the array, fault dips of 21° to 39°
could result in the vertically propagating headwave seen in Figure 3.5. Using a more
reasonable velocity estimate than the minimum and maximum results in a fault dip of 29°
to 30°. Using the same reasonable velocity estimation, but assuming fault dips of 21° and
39°, results in the synthetic arrivals represented by dashed lines in Figure 3.5.
These results are in general agreement with Meister’s [1967] velocity model of The
Bend area, generated from his Dixie and IXL Canyon refraction lines. His bedrockalluvium velocity contrast of 1.9:1 results in a fault dip of 25°. The coherency of the first
arrival on Figure 3.5 indicates that the fault is smoothly varying and relatively planar
down to 500 m depth. This is in contrast to the segmented first arrivals that would be
generated by refraction and diffraction along a staircase-like fault geometry. Meister’s
[1967] model in which he modeled a staircase-like fault geometry had insufficient
resolution (cross-basin spacing of approximately one geophone per 400 m as compared to
one geophone per 15 m in this study) to resolve downdip fault geometry at a fine scale.
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3.3.2.2 Migrations
A postmigrated stack of our medium-resolution Cattle Road profile is presented in
Figure 3.7. The data were Stolt time migrated at a constant velocity of 2000 m/s. The
interpreted fault plane reflector, dipping eastward at 25° to 30°, can be traced to the Dixie
Valley fault reflector seen in the high-resolution profile (Figure 3.4). Reflections
subparallel to the fault can be seen in the footwall, suggesting the Caskey et al. [1996]
foliation in the footwall bedrock. The fault reflections are obscured below a highly
reflective hanging wall layer at ~400 ms. This first, strong basin reflector has been
interpreted to be a “capping” basalt layer by Hastings [1979] in a reflection profile 25 km
northwest, in the Carson Sink (Figure 3.1). This interpretation is supported by drilling
logs that intersected the profile. Okaya and Thompson’s [1985] reflection profile near the
Dixie Valley Geothermal Field (Figure 3.1) also shows a similar reflection, which they
interpret to be from the same reflector seen in Hastings [1979]. This basalt is the
culmination of a Tertiary volcaniclastic sequence that is seen in all the ranges
surrounding the Carson Sink and Dixie Valley. The sequence appears to be locally
thickened in the southern Stillwater Range, with thicknesses approaching 500 m [Page,
1965]. We interpret the strong series of reflections starting at 500 ms and ending at 900
ms to be originating from this sequence.
Below the capping basalt, hanging wall stratigraphy can be traced to its termination
against the fault at depth. Mapping the terminations allows us to extend our fault plane
interpretation to 1.25 s. Many of the reflections show increased westward dip close to the
fault, forming rollover anticlines. The rollover anticlines form in response to listric fault
geometry.
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Figure 3.8 shows the enhanced pre-stack migration that results from imaging the
medium-resolution data through the extended 8 m velocity model shown in Figure 3.6.
Because of the low-pass filtering inherent in the antialiasing criterion this section does
not have as high a vertical resolution as the poststack migrated section and may not
image more steeply dipping structures such as anticline flanks. As demonstrated in Death
Valley, California, by Chavez-Perez et al. [1998], however, the prestack migration
through an optimized velocity model does assure that this section puts structures at their
true depths. Thus any dips interpreted from Figure 3.8 will be accurate to within a few
degrees.
The top of the Tertiary basalt at 0.5 km depth, and its bend into a west dipping
rollover anticline against the Dixie Valley fault ~1 km east of the 1954 scarp, can be
interpreted in Figure 3.8. The prestack image shows a reflection sequence at 0.8 km
depth, which the stacked data could not image without velocity pull-down effects. These
reflections also show some evidence of rollover and may originate at the bottom of the
Tertiary volcanic sequence at the top of an earlier-Tertiary basin-fill sequence. The image
shows only parts of the Dixie Valley fault plane itself, at 0.5, 1.4, and possibly 1.7 km
depths, dipping east from the 1954 scarp at 28°. The low-frequency response of the
antialiasing obscures the fault plane above 0.5 km depth. The prestack migration, like the
poststack migration, failed to clearly image the fault plane below the strong reflectivity of
the capping basalt. Truncations of the deeper basin stratigraphy do support the linear,
shallowly dipping fault geometry.
The imaging of a flat basin fill sequence at 1.1 km, in on-lap relation to the Dixie
Valley fault, suggests that no appreciable rotation of the fault plane has occurred since its
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formation. It also suggests that significant extension had occurred before the initiation of
volcanism. The flanks of any rollover anticlines at this depth could have been hidden by
the migration antialiasing criterion. Alternatively, a constant average fault dip could have
avoided the formation of rollover anticlines in the deep basin. The two upwarps in the
fault profile suggested by Figure 3.7 between 0.5 and 1 s two-way time may have
resulted in rollovers forming only updip of each. Although the fault reflector itself cannot
be seen below the capping basalt reflector, the Figure 3.8 imaging of a continuation of
on-lap truncations along the projected fault plane to 1.5 km depth supports our
interpretation of an early, now deeply buried, but unrotated episode of basin filling.
3.3.3 Gravity Results
Bouguer gravity along Cattle Road agrees with Schaefer’s [1983] study. The 2.5dimensional gravity modeling (Figure 3.9) is consistent with a shallowly dipping fault. In
the cross section the Dixie Valley fault is modeled as dipping 26°. Density values
generally follow those of Speed [1976] and Thompson [1959] as summarized by Okaya
and Thompson [1985]. A value of 2.5 g/cm3 was chosen for the volcaniclastic units.
Higher in the section, an average value of 2.3 g/cm3 was chosen for volcanic units
intercalated within sedimentary units, as seen in Hastings’ [1979] well log in the Carson
Sink. The gravity effect of topography (terrain correction) was modeled in the plane of
the cross section.
Maximum basin depth is ~2700 m, using our density scheme. The Dixie Valley fault
merges with the steeper Clan Alpine range-bounding fault. It is likely that the Clan
Alpine range-bounding fault shown in Figure 3.9 is actually a series of faults as shown by
Okaya and Thompson [1985], rather than the one large fault modeled. Owing to our lack
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of accurate density information or seismic profiles in this area, an effort was made to
produce the simplest model that fits the anomaly. Although a staircase-like fault
geometry to the east cannot be ruled out solely on the basis of gravity data, the absence of
short-wavelength anomalies suggests no rapid changes in fault dip. Results from the
Scarp profile (Figure 3.2), not shown, indicate that intrabasement density contrast is
negligible along this reach of the fault.
3.4 Discussion
Evidence for low-angle faulting is seen at several scales, from observations at the
surface to over 2.5 km depth. Moving from shallow to deep, the evidence consists of the
following:
1. Geologic evidence from Caskey et al. [1996], in the form of rupture mapping and
geologic cross-sections is valid from 0 to 10 m depth.
2. Our high-resolution profile (Figure 3.4) confirms the geologic observations and
extends the smooth, low-angle fault plane to 75 m depth.
3. The no a priori assumption velocity optimization (Figure 3.6) shows a surface of
increasing velocity dipping shallowly to 480 m depth.
4. Raw shot gathers (Figure 3.5, for example) constrain the fault to be relatively
planar to 500 m and, given reasonable velocity estimations, suggest low-angle dip.
5. The medium-resolution time migration reflection profile (Figure 3.7) shows direct
fault plane reflections from 50 to 750 m. In addition, truncations in hanging wall
stratigraphy seen in the medium-resolution profile allow the interpretation of a slightly
listric low-angle fault plane to ~1.0 km depth.
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6. The same character is seen in the Kirchhoff depth migration of the mediumresolution profile (Figure3. 8), extending observations to 1.75 km depth.
7. Gravity mapping by Schaefer [1983] (Figure 3.2) and this study (Figure 3.9) is
valid from the surface to the maximum depth of the basin, ~2.7 km. The gravity data, too,
are consistent with a low-angle fault geometry.
3.4.1 Timing of Extension
The recognition of an early basin fill sequence below the mid-Tertiary capping basalt
(Figure 3.8) suggests that extension in the southern Stillwater Range started earlier than
extension to the north. Early sediments have been mapped in the Stillwater Range at the
latitude of our transects by John [1995]. Seismic reflection profiles near the Dixie Valley
geothermal field [Okaya and Thompson, 1985] and the northern Carson Sink [Hastings,
1979] record no such sequence. The pervasive capping basalt layer, dated as 8 ± 4
[Hastings, 1979] and 13 to 17 Ma [Nosker, 1981], was interpreted to be preextension. In
southern Dixie Valley a rapid, but spatially limited, pulse of extension is recorded by
tilted fault blocks in the Stillwater caldera complex [John, 1995; Hudson et al., 2000].
The tuffs, flows, and plutons associated with the complex have tilts of 60° to 70°. Similar
deposits in the southern Clan Alpine Mountains and northern Stillwater Range dip < 30°,
suggesting very localized uplift and tilting. Hudson et al. [2000] estimate over 200%
extension in a brief period from a balanced cross-section near IXL Canyon (Figure 3.2).
The extension is well constrained to have started at 24.2 to 24.4 Ma [Hudson et al. 2000].
Parry et al. [1991], on the basis of fluid inclusion and alteration mineral studies, also
estimate extension starting at 20 to 25 Ma. The extension direction was, in general,
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perpendicular to current structural strike, although some vertical axis rotation has
occurred since early Miocene time [Hudson and Geissman, 1987; Hudson et al., 2000].
Two contrasting models of early Miocene extension have been suggested. Parry et al.
[1991] suggest the current Dixie Valley fault is the major fault accommodating extension
and localized uplift and has been active since the early Miocene. Similarly, King and
Ellis [1990] cite the Dixie Valley fault as a potential field example of a fault
accommodating dramatic, localized uplift, starting in the early Miocene. The rotated
crustal blocks mapped in the Stillwater Range represent rotation of footwall blocks of the
Dixie Valley fault in this interpretation. In contrast, Hudson et al. [2000] consider it
unlikely that the current Dixie Valley has been active longer than 13 to 15 Ma (postdating
the capping basalt). As evidence, they point to a sharp accommodation zone between two
dip domains in the Stillwater Range [Hudson et al. 2000]. The sharply defined
accommodation zone separates domains of east and west dip in the rotated caldera
deposits, requiring two large normal faults of opposing dips to form. The current Dixie
Valley fault cuts across both the east and west dip domains. Instead, they favor a model
in which extension along a west dipping (now inactive) detachment rotated the caldera
deposits, which are synthetically and antithetically dipping in the hanging wall of the
low-angle detachment [Hudson et al., 2000]. Current Basin and Range extension
subsequently cut the resulting structure in the middle Miocene [Hudson et al., 2000].
We favor early Miocene initiation of the southern Dixie Valley fault from our
reflection profiles and evidence for low-angle dips. Evidence for early extension comes
from the basin fill and volcanic sequences seen in Figures 3.7 and 3.8 and mapped in the
Stillwater Range [John, 1995]. This flat-lying sequence is directly below a thickened
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Tertiary (early to middle Miocene) volcanic section. The volcanic section is locally
thickened in the southern Stillwater Range [Page, 1965]. The volcanic sequence may be
thickened in our profiles because it filled a preexisting basin at this latitude. The
reflection profiles are adjacent to the west dipping blocks in the Stillwater Range, which
require a large east dipping normal fault in precisely the same location as the current
Dixie Valley fault.
Formation at low-angle may have been in response to rapid extension. Abers et al.
[1997] found a correlation between low-angle faulting and increasing strain rate in the
Woodlark-D’Entrecasteaux rift system, where strain rates of up to 10-8 s-1 are seen.
Hudson et al. [2000] find minimum strain rates in the early Miocene of 10-13 s-1 in the
southern Stillwater Range. That strain rate may be underestimated because the exact
duration of extension is not known. Nevertheless, strain rates at least an order of
magnitude greater than usually encountered may have played a role in forming the lowangle structure. Current Basin and Range extension along the low-angle reach may have
inherited the same fault that accommodated early Miocene extension. North and south of
the low-angle section that formed in the early Miocene, extension had no favorably
oriented structures to inherit and formed a new steeply dipping normal range front fault at
~13 to 15 Ma.
3.4.2 Magnitude of Slip
Estimates for fault slip from our seismic lines can be broken into two phases (Figure
3.10). One phase is constrained by the thickness of the early Miocene basin, and the other
is constrained by the current elevation of the capping 13 to 15 Ma basalt. Our gravity
results indicate a maximum basin depth of 2.7 km. The bottom of the Tertiary volcanic
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sequence and top of the early Miocene basin is interpreted to be at 1.1 km depth from our
Kirchhoff migration (Figure 3.8). Using dates and tilts of caldera deposits, Hudson et al.
[2000] estimate that this phase of extension lasted at most 4 Myr. The majority of the
rotation was most likely accomplished in ~0.2 Myr. The 1.6 km of vertical offset along a
30° fault in ~0.2 Myr corresponds to slip rates of 18 mm/yr (extension rate of 9 mm/yr).
This is an extraordinarily fast slip rate, but it has a current analog in the WoodlarkD’Entrecasteux extensional province, which has extensional rates of over 15 mm/yr and
as high as 40 mm/yr in an area of low-angle normal faulting [Abers et al., 1997].
However, using the most conservative estimate of extension duration from Hudson et al.
[2000] yields a more typical slip rate of 0.9 mm/yr in the early Miocene. Errors in the
gravity model and the duration of extension can make these fault slip rates uncertain,
however.
The second phase of extension postdates the middle Miocene capping basalt (13 to
15 Ma). The basalt is seen at 400 m below the surface (650 m elevation) in Figure 3.8 and
at an elevation of 2500 m in the Stillwater Range. Accommodating the 1850 m of vertical
offset along a 30° fault in 13-15 Myr requires a slip rate of 0.28 to 0.32 mm/yr. These slip
rates compare favorably with Meister’s [1967] estimation of 0.3 mm/yr over 15 Myr for
southern Dixie Valley. Okaya and Thompson [1985] found a fault slip rate in northern
Dixie Valley of 0.47 mm/yr along a 50° fault (using Hasting’s [1979] basalt age of 8
Ma). This would be equivalent to a 13 Myr slip rate of 0.29 mm/yr. Bell and Katzer
[1990] report a Holocene vertical slip rate of 0.5 mm/yr in the Bend area.
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3.5 Conclusions
Our results indicate that slip along a section of the December 16, 1954, Dixie Valley
earthquake rupture took place along a fault plane of unusually low dip (<30°). In this
regard, it is the first large historical earthquake on land for which slip on a low-angle
normal fault has been documented. Evidence for the low-angle fault plane is seen at
several different scales, from 0 to 2.7 km depth. A computed velocity model, with no a
priori assumptions, supports a low-angle hypothesis, as does gravity modeling. Our
results suggest extension in the southern Stillwater Range had two distinct phases. The
first period is marked by rapid extension that initiated a low-angle fault. The second
period of extension began at 13-15 Ma and inherited a portion of the previous Dixie
Valley fault along the low-angle section. Fault sections with steep dip formed at 13 to 15
Ma.
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3.6 Acknowledgments
This project was funded by the National Science Foundation under grant EAR-9706255
to J. Louie, S. J. Caskey, and S. Wesnousky. The W. M. Keck Foundation donated
seismic equipment, computers, and modeling software. K. Miller and G. Abers provided
careful reviews and helpful comments. The 1998 Geophysical Applications class at the
University of Nevada, Reno performed all geophysical fieldwork. Class participants were
A. Cadena, T. Rabe, M. Herrick, M. Johnson, A. Rael, T. Blechen, and E. Hobson. C.
Mann, J. Ollerton, and J. Oswald rendered additional field assistance.
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Chapter 4: Weak ground motion and amplifications
predicted from shear-wave velocities at precarious
rocks, near the 1857 rupture of the San Andreas fault
Robert E. Abbott, John N. Louie, James. N. Brune, Abdolrasool Anooshehpoor
The University of Nevada, Reno, Seismological Laboratory and Dept. of Geological
Sciences, Reno, Nevada 89557.
Sathish Pullammanappallil
Optim L.L.C., Seismological Laboratory, University of Nevada, Reno, Nevada
4.0 Abstract
The effects of local geology at fields of precarious rocks are inadequate to explain
the persistence of untoppled rocks near the San Andreas fault, given current attenuation
relationships. As evidence, we present weak ground motion spectral ratios at six locations
from 61 earthquakes, in situ velocity measurements using both refraction-microtremor
and tomographic techniques, and synthetic spectral ratios based on our velocity
measurements. We compare the “known” toppling acceleration of the rocks and the
accelerations predicted by current USGS/CDMG probabilistic seismic hazard maps at the
rock locations. We modify the predicted USGS/CDMG accelerations by accounting for
the measured shallow shear-wave velocity structure at the rock sites and the assumed
velocity structure used for the maps. Comparing the synthetics generated using our
velocity modeling to synthetics generated using the soft rock/soil interface velocity
structure (NEHRP BC boundary), we find ground motions at the precarious rock sites
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0.75 to 1 times the ground motion at an assumed BC boundary location at 3 to 5 Hz.
Above approximately 5 Hz, the ground motion at the precarious rock sites is amplified
above the ground motions at the BC boundary site. Velocity modeling shows that the
precarious rock sites are characterized by velocities significantly slower the than the BC
boundary from 0 to 10 meters, and significantly greater below 10 m. Strong impedance
contrasts at the precarious rock sites in the upper 10 m caused by seismically slow dry
sand over fast fractured granite explains the amplification pattern.
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4.1 Introduction
The geographical distribution of fields of precariously balanced rocks is gaining
increased acceptance as a potential ground motion constraint in seismic hazard models.
“Precarious” rocks (defined as rocks with toppling accelerations of 0.1 to 0.3 g) and
“semiprecarious” rocks (0.3 to 0.5 g) are located as close as 14 km to the Mojave section
of the San Andreas fault [Brune, 1999]. This section of the fault ruptured in the M8,
1857, Fort Tejon earthquake. Sieh [1978], at a trench site approximately 14 km due south
of the precarious rocks, documented 8 large earthquakes in the last 1400 years. At least 6
of those events exhibit ground deformation similar to the 1857 event [Sieh, 1978]. The
continued existence of such rocks in their current configuration for many thousands of
years [Bell et al., 1998] is incompatible with most ground motion prediction models
[Brune, 1996], since the rocks have remained untoppled for many cycles of great
earthquakes.
There are several theories to explain the discrepancy. One line of reasoning is that
current attenuation relationships overstate the hazard of large magnitude earthquakes.
Strong-motion recordings of magnitude 7 and greater events at short distances are rare,
resulting in attenuation curves near great earthquakes that are extrapolated by parameters
found from smaller earthquakes recorded at greater distances. Recent observations of
ground motion from the 1999 Kocaeli, Turkey, MW 7.6, earthquake and the 1999 ChiChi,Taiwan, MW 7.6, earthquakes have dramatically increased the number of strong
motion recordings near large ruptures. Peak accelerations recorded at these stations less
than 100 km from their respective ruptures are significantly less than predicted. Anderson
et al. [2000] summarize these data and conclude that it would be unsatisfactory to
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conclude that the two best-recorded earthquakes are anomalous, and that current
attenuation relationships may overstate the hazard.
Brune [1999] suggested that great earthquakes on the San Andreas fault have
“characteristic” spectra, repeated each earthquake cycle. This alternative hypothesis is
supported by the observation of Wesnousky [1988] that showed a negative correlation in
the “complexity” of a fault (measured by the number of steps or fault segments per km
length) and total fault offset. Longer, more developed faults have fewer steps and other
impedances to rupture, resulting in slower earthquakes with less high frequency ground
motion components. Building upon this, Anderson and Brune [1999a] discuss a thought
experiment created by limiting or eliminating the statistical variation of repeated large
earthquakes and find a better fit to the precarious rock data.
In another line of reasoning, Anderson and Brune [1999b] find that seismic hazard
maps that limit earthquakes to mapped faults (i.e. no “area” sources) are compatible in
most cases for precarious rocks in Nevada. Incomplete knowledge of active fault
distribution, as evidenced by the Landers, Hector Mine, and Northridge earthquakes,
suggests that this particular result of Anderson and Brune [1999b] would not be
appropriate for all situations. It would be most appropriate for situations for areas where
the hazard is dominated by a single, well- characterized, fault, as is probably the case for
the precarious rock sites in this study.
Ground motion is a convolution, of course, of source processes, effects on the
seismic waves along the path from the source to the site, and effects caused by the very
local geology of the site. This project specifically studies the problem not at the source,
but rather at the sites of the rocks themselves. It is currently not known if these precarious
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rocks experience the same ground motions as the surrounding regions they are located in,
or if they define small “islands of stability” in a general ground motion field. Namely, we
aim to test whether or not de-amplification of ground motion at the precarious rock sites
studied herein relative to “representative” sites is sufficient to decrease the level of
shaking to below the rocks’ toppling accelerations, using the currently accepted
attenuation curves. We accomplish this with a combination of earthquake spectral ratios,
in situ velocity measurements, and calculation of synthetic ground motions.
4.2 Data and Methods
4.2.1 Earthquake Data
In October 1999, we deployed matched, calibrated, 3-componant digital stations with
1 Hz L4 and S13 sensors at 6 sites near the San Andreas fault (Figure 4.1; Table 4.1) in
preparation to record blasts from the Los Angeles Regional Seismic Experiment, Phase 2
[Similia et al., 2000]. We calibrated the sites by applying a known voltage in the field and
modeling the resulting waveform (following a damped, harmonic oscillator) with three
parameters: free period, damping, and magnification. Four of the sites (Lovejoy Buttes,
Piute Butte, Alpine Butte, and near Black Butte) are located at or near fields of precarious
or semi-precarious rocks. Precarious rock are less than 200 m away at Lovejoy Buttes,
Piute Butte, and Alpine Butte. A thorough search for precarious rocks has not been made
near the Black Butte site, but the outcrop of rock where the station is located is near an
outcrop where other precarious rocks exist. We also installed a station (Mill Creek
Summit) near the USC strong motion site MCS and one near the town of Llano,
California. Fortuitously, we recorded numerous aftershocks of the October 16, 1999,
Hector Mine earthquake during our 48 hour occupation of the sites (Table 4.2). The poor
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signal-to-noise ratio of the blasts compared to the aftershocks led us to focus on the
earthquake data.
The parameters of the 61 earthquakes in the analysis are presented in Table 4.2.
Table 4.2 also lists which sites recorded each event. The event magnitudes ranged from
1.8 to 5.1 ML. Figure 4.2 shows a sample seismogram from an aftershock. We tried to
include events that have different azimuths to create, as much as possible, a pathindependent amplification scheme. Due to our short occupation time at the sites,
however, the sources are predominately near the Hector Mine rupture area. Each event
was examined and only events with good signal-to-noise ratio on the majority of the
stations were accepted for further analysis.
4.2.2 Spectra Calculation
We computed the spectra of ground-velocity for each component of each record. The
Fourier spectra were calculated for a 5 second time window, starting one-half second
before S-wave arrival. This window was chosen to best contain most of the high
amplitude direct S-wave energy. As pointed out in Bonilla et al. [1997] and Satoh et al.
[2001] using longer times may result in better spectral resolution but also contain energy
from surface waves and other phases.
For each 5-second seismogram, the mean value was removed, and a 5% Hanning
taper was applied. We accounted for attenuation using a frequency-independent Q value
of 1000, based on when the high frequency spectra flattened out. This regional Q is
similar to that found in Adams and Abercrombie [1998]. The ratios are, for the most part,
insensitive to Q, except for ratios at MCS, the station farthest removed from the cluster of
precarious rock stations. After FFT and Q removal, the two horizontal spectra for each
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record were averaged to form one horizontal spectrum. The resulting vertical and
horizontal spectra were smoothed with a three point moving average and then
logarithmically stacked.
4.2.3 Spectral Ratios
One of the most common methods of estimating site response is the spectral ratio
technique [e.g. Borcherdt, 1970]. It is computed by dividing the Fourier spectrum of one
station by the Fourier spectrum of another. The spectral ratio technique assumes that site
effects can be deconvolved from the others if: a) the instrument response is known for
each station; b) the same sources are used; and c) the path from source to station is the
same. By using the same set of earthquakes, with the hypocentral distance being much
greater than the distance between the stations, differences in source and path effects will
most likely be minimal. Local site conditions, therefore, cause the remaining differences
in the spectra.
In general, spectral ratios are computed by dividing the horizontal spectra of a station
for which you wish to know the site response, by the horizontal spectra of a nearby
reference station on rock. The reference station, by virtue of being on rock, is assumed to
have minimal site effects, although there is evidence that this is not necessarily true [e.g
Steidl et al., 1996; Humphrey et al., 1992, Tucker et al., 1984]. We chose Piute Butte as
our reference site because tomography results (discussed below) show it to be the site
with the fastest subsurface velocity.
4.2.4 Velocity Data
We also measured shallow shear-wave velocities at four of the sites for which we
have 3-componant seismograms (LJB, PB, LLA, and MCS). Standard seismic refraction
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equipment was used with 24 8-Hz vertical sensors and linear spread lengths of ~184
meters. The transects were as close as possible to the 3-component stations as the local
topography and access allowed. The resulting analysis yielded the shear-wave velocity
structure at the sites to a depth of 45+ m. The data was acquired using a combination of
refraction tomography and refraction-microtremor (ReMi) techniques. The ReMi
methodology is discussed in detail in Louie [2001], who demonstrated that the resulting
shear-wave velocity depth profiles agree with borehole measurements to within 20% at
other locations.
To supplement these data, we also used a commercial refraction velocity
optimization software, SeisOpt®@2D™ (© Optim LLC, 2001), to derive P-wave velocity
information from 1st arrival picks. This software uses a non-linear optimization technique
called simulated-annealing (Pullammanappallil and Louie, 1994) to map subsurface
velocities from first-arrival picks using no a priori assumptions We picked at least 64 Pwave first arrivals along each array, using hammer source points. After inversion for Pwave velocity, we assume a constant Poisson’s ratio of 0.25, which allowed us to
estimate the shear-wave velocity and compare results from the two fundamentally
different methods.
4.2.5 Calculation of Synthetics
We computed synthetic transfer functions and synthetic spectral ratios for the four
for which we determined velocities. The synthetic ground motions were calculated using
the STK5 program (courtesy of John G. Anderson). STK5 calculates the response of a
stack of sediments on a spectrally flat, vertically incident, SH-wave. The algorithm
calculates reflection and transmission coefficients at each velocity interface for both the
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up-going and down-going waves using the technique described by Luco [1983] and Apsel
[1983]. The solution at the surface includes waves traveling in both directions. We used
our velocity models, as well as the velocity profiles for the NEHERP BC boundary
[Frankel et. al, 1996] and “generic rock” [Boore and Joyner, 1997], to compute the
synthetics. Also, to facilitate comparisons with our earthquake spectral ratios, we ratio
the synthetic spectra at each site with the synthetic spectra from the Piute Butte velocity
profile.
4.3 Results
4.3.1 Local Shear-Velocity Structure
Shear-wave velocities (Figure 4.3) from modeling of refraction-microtremor
dispersion curves (Figure 4.4) and refraction tomography (Figure 4.5) show that all sites
except MCS are in the NEHRP site class B (> 760 m/s) or even A (> 1500 m/s) range. In
the upper 10 m, however all the sites are slower than the BC boundary model used to
generate the seismic hazard maps [Frankel et al., 1996]. The generic rock profile from
Boore and Joyner [1997] matches the shallow depth profiles much more closely. This is
not surprising, given that the generic rock profile is an average of velocity profiles from
57 rock boreholes. Our results show that the precarious rock sites have higher velocities
below 20 m than an “average” rock site.
30-m average velocity from the two methods (Table 4.3) match very well at MCS
but agreement is not good at the other sites. This is because the ReMi method emphasizes
layered geometry (all stations were modeled with 3 or fewer layers), while the
tomography method emphasizes velocity gradients. Also, all the models are not unique,
in that a number of plausible subsurface velocity structures can be postulated that all fit
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90
the data equally well. In general, the two methods agree better when deeper averages (i.e.
averages over 50 or 60 m instead of 30 m) are used. Comparisons of the velocities at the
maximum depth of penetration show that the two methods agree to within 35%, even
with the assumed 0.25 Poisson’s ratio.
4.3.2 Horizontal and Vertical Spectra
Velocity spectra for both the vertical and averaged horizontal components are
presented in Figure 4.6. As can be seen the horizontal spectrum is almost always greater
than or equal to the vertical spectrum. Also notice that for almost all the stations, the
maximum separation of the two components happens around 4-7 Hz. Steidl et al. [1996],
noted that spectral ratios using rock sites were often underestimated as the rock sites had
their own maximum site effect in this approximate frequency range.
Stations LLA and MCS have a slightly different character. The horizontal
component at both sites is elevated for the entire frequency range. A correlation seems to
exist between the relative amount of energy in the vertical and horizontal components
and the subsurface velocity. Sites with lower velocities (LLA and MCS) have more
energy in the horizontal, perhaps because the incoming S-wave is refracted more toward
vertical incidence near the sites. Alternatively, contamination by reflected S and surface
wave energy, due to both regional and local geology may cause the elevated horizontal
energy. Satoh et al. [2001] compare P-wave, P-wave coda, S-wave, S-wave coda, and
microtremor spectral ratios and attribute the differences among the results to the relative
strength of surface waves. The surface waves are caused by conversion of phases at
discontinuities local to the site [Satoh et al., 2001]. A possible cause of the elevated
horizontal spectrum at LLA is its close proximity (< 2 km ) to the San Andreas fault.
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Using the group velocities for low frequencies suggested in our ReMi dispersion curves
(Figure 4.4) shows that conversions at the SAF could easily traverse the 2 km and
contaminate our 5-s time window. Even if the conversion of phases are not at the SAF
itself, close proximity to the fault is sure to cause some pervasive subsurface
deformation. Similarly, tectonic deformation the likely cause of the elevated low
frequencies at station MCS, located on pervasively fractured Precambrian rocks in the
San Gabriel Mountains. We ran separate analyses on subsets of the data to see if our goal
of a path-independent amplification scheme was achieved. Spectra from three subsets of
earthquakes were examined: the Hector Mine aftershock area, the Big Bear Lake
aftershock area, and the northern Mojave area (only two events here). The spectra from
all three source areas (not shown) are quite similar, leading us to conclude that path
induced affects are relatively minor. Very fine structure in the spectra generated at the
source might be averaged and smoothed, however.
4.3.3 Spectral Ratios
Traditional spectral ratios at the precarious rock sites, with Piute Butte as the
reference site, show spectral ratios that are generally flat, with amplification factors near
unity (Figure 4.7). Again, the spectral ratios for MCS and LLA (Figure 4.7c-d) are
different, probably due to being on highly fractured rock. The spectra there have a “steplike” geometry with amplification relative to the precarious rock site from 0.4 to 4 Hz,
and de-amplifications relative to the precarious rock site from 4 to 20 Hz. These results
are reminiscent of those in Stirling et al. [2001], who have studied the spectra of PB and
LJB relative to three TRINET stations on NEHRP class B sites. They find negative
residuals below about 4 Hz (meaning de-amplification at the precarious rock sites relative
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to NEHRP B) and positive residuals above 4 Hz (amplification relative to NEHRP B). If
we consider LLA to be a NEHRP class B site, as suggested by the tomography but not
the refraction-microtremor, our results are in agreement. Indeed, we show the same
pattern of de-amplification of precarious rock sites at low frequencies and amplification
at high frequencies relative to our NEHRP class C site, MCS. The flat spectral ratios with
ALP and BB suggest similar site conditions as LJB and PB and add robustness to the
conclusions of Stirling et al. [2001].
4.3.4 Synthetic Spectral Ratios with Piute Butte
Synthetic spectral ratios, based on the velocity models, show some success
predicting overall spectral shape (Figure 4.7). We feel confident that the velocity models
for all stations except MCS include information about the deepest relevant basement
interface (relevant in terms of having a velocity very near true basement velocity with no
significant impedance interfaces at greater depths). The tomography model at MCS
(Figure 4.5d) shows widely varying velocities at the deepest depth of penetration. This
suggests that the average velocity along the 184-m array that both the tomography and
the ReMi methods measure is too low. As such, the synthetic transfer function generated
for MCS is not constrained at the lowest frequencies. Indeed, the synthetic ratios at MCS
show the most disagreement of all stations at low frequencies.
The synthetics from both tomography and ReMi predict a change from amplification
relative to Piute Butte, to de-amplification relative to Piute Butte in the 5 to 20 Hz range,
agreeing with the earthquake spectral ratios. We do not expect the earthquake and
synthetic ratios to be in perfect agreement because of any number of assumptions that
could be violated. Discrepancies between synthetics and data could be due to a number of
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causes, including, but not limited to: surface wave contamination, incorrect velocity
models, and path-induced effects on the input spectra below each site.
4.3.5 Synthetic Spectral Ratios with Two Reference Models
Comparisons to the BC boundary model [Frankel et al., 1996] and generic rock
[Boore and Joyner, 1997] give seismic hazard context to our data (Figure 4.8). Care must
be taken analyzing the synthetics as they still have the same frequency limitations and
caveats as the ratios with Piute Butte. However, the data still show the consistent result of
de-amplification at precarious rock sites at low frequencies and amplification at high
frequencies, seen in this study and in Stirling et al. [2001]. The NEHRP class C site,
MCS, shows the least amplification above 5 Hz, while the faster rock sites show much
more.
The synthetic ratios show amplification factors relative to the BC boundary model
for the precarious rocks of ~0.75 at 3 Hz, and ~1 at 5 Hz. Shake table results (not shown)
demonstrate that the peak ground acceleration threshold at which models of precarious
rocks first topple is linearly proportional to the spectral accelerations at 3 and 5 Hz. That
is, plots of PGA at the minimum toppling threshold of PGA versus 5 Hz acceleration
have a linear trend (with some scatter). The scatter is greater at 3 Hz, with 1 Hz and
below having no relationship to PGA. This relationship implies that we can scale the
probabilistic seismic hazard PGA at the precarious rock sites (0.6 g with a 10%
probability of exceedance in 250 years [Frankel et al., 1996]) by the amplification factors
at 3 and 5 Hz. The rocks at LJB and PB have toppling peak accelerations of ~0.4 g
[Brune, 1999]. Applying the amplification factors at 3 and 5 Hz results in revised PGA
estimates of 0.45 to 0.6 g. Given the mild de-amplification, and rocks that have been
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balanced for around 40 times the exposure time of the PSH estimate, we conclude that the
persistence of the rocks in their current configuration can not be attributable to deamplification due to local shear velocity.
4.4 Discussion and Conclusions
Ground motions at the precarious rock sites in the Mojave are de-amplified below 58 Hz and amplified from 10 to 20 Hz. Therefore the results of this experiment indicate
that site effects due to local geology cannot be used to explain the continued existence of
balanced rocks. Velocity modeling suggests the cause of the amplifications are extremely
strong impedance contrasts in the upper 20 m. The precarious rocks are formed as the
hard granitic rocks are differentially eroded, leaving balanced outcrops. The current arid,
erosional environment at the precarious rock sites leads to the juxtaposition of dry,
unconsolidated granitic sands against hard, fractured rock. Even though the average
velocity to 30-m depth indicates NEHRP class A or B, the upper 10 m is slower than an
average soft-rock/soil interface site. It is the travel time, however, and not average
velocity, that is the critical amplification parameter and even very thin layers can have
large effects on the seismic travel time [Boore and Joyner, 1997]. The Mojave desert
sites, due to the local climate and geology, have slower than average velocities in the
upper 10 m, and therefor greater than average amplification factors above ~5 Hz. This
may not be true for precarious rock sites in different environments, such as those found in
the Sierra Nevada Mountains or near the Nevada Test Site.
We are guilty, however, of using smaller earthquake data at greater distances and
applying it to a large San Andreas fault earthquake. It is widely recognized that deep
basin geometry can be a dominant factor in determining amplification at sites within
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sedimentary basins [e.g., Gao et al. 1996, Hartzell et al. 1997; and Davis et al., 2000].
Subsurface focusing of seismic energy propagating through velocity contrasts at basin
margins is thought to be a mechanism producing the amplifications. If this is true, than
there must be areas of de-focusing and de-amplification. It is possible that the fields of
precarious rocks are located in such an area for a San Andreas event. By necessity, we
essentially examined ground motions from sources near the October, 2000, Hector Mine
earthquake area and the amplification patterns will be different for a San Andreas event.
3-D basin amplification patterns can change dramatically given different earthquakes on
the same fault. Even changing the rupture direction of an earthquake can have dramatic
results in finite difference simulations [Olsen, 2000] and the effects of rupture directivity
have lead some to propose changes to the empirical attenuation relations [Sommerville et
al., 1997]. Given that the rocks have been in their current configuration for as many as 10
great events on the San Andreas fault, some sort of rupture pattern would probably have
resulted in sufficient ground motion to topple the rocks if current attenuation parameters
are correct.
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4.5 Acknowledgements
We thank Glenn Biasi, John G. Anderson, and Mark Stirling for their valuable input
and assistance. This research was supported by the Southern California Earthquake
Center. SCEC is funded by NSF Cooperative Agreement EAR-8920136 and USGS
Cooperative Agreements 14-08-0001-A0899 and 1434-HQ-97AG01718. The SCEC
contribution number for this paper is 601.
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Chapter 5: Recommendations for Future Work
5.0 Organization
In this chapter I will briefly describe my recommendations on how the work and
conclusions presented in the previous three chapters can be strengthened through
additional work. Each project will be treated separately.
5.1 Recommendations for Future Reno-Carson City Basins Work
In Chapter 2 I showed that quick and accurate basin mapping was possible using
gravity measurements in key locations and very simple approximations. Subsequent to
the publication of Abbott and Louie [2000], I became aware of two new wells that have
penetrated bedrock (both in the “Verdi Basin” area. Examination of well logs and
cuttings at Summit Engineering in Reno, NV, revealed that in these two wells, my depthto-bedrock predictions were consistent to within 10%. There are several ways other than
using new borehole results to test the data and to improve seismic hazard prediction.
5.1.1 Increased Gravity Coverage
As noted in Chapter 2, there are several areas of insufficient gravity coverage in key
areas. These areas manifest themselves most clearly when negative depth-to-bedrock
values are predicted. Examples of these areas are I80-US395 interchange in Reno (Figure
2.6) and the area just east of the 5th Street transect in Carson City (Figure 2.12). These are
the most obvious areas, but locations of sparse gravity coverage near sharp gravity
gradients are also areas of concern. The southwest portion of the West McCarran gravity
trough stands out as an example of this.
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Washoe County has recently acquired many new gravity measurement points
between Verdi and Reno that would greatly improve coverage in this area [Mike Widmer,
Washoe County, personal communication]. In the Reno area, I would also suggest
transects along Sutro Street and Clear Acre Lane as being particularly useful.
Carson City coverage suffers from a general lack of measurements in the
surrounding margins of the basin. Additional measurements will be difficult due to access
problems, but just a few measurements in keys areas would be of use in making the
bedrock gravity field more accurate.
5.1.2 Density Measurements
One major weakness in our gravity modeling was our incomplete knowledge of
density contrasts. We tried to make use of the nearby measurements of Thompson and
Sandberg [1958] and the regional model of Blakely et al. [1999], but measurements
within the local outcrops of the Hunter Creek and Kate Peak formations would have been
particularly useful. The Hunter Creek sandstones would probably have to be sampled in
more locations because it seems apparent that the density of the different members of the
formation varies widely. This would allow the bracketing of the density contrasts
between likely minima and maxima for “average” Hunter Creek for the one-dimensional
approximation, or even the introduction two- or three-dimensional models.
5.1.3 Finite Difference Modeling
Although the goal of Chapter 2 was to predict basin depth, it was part of a larger
project whose goal was to estimate seismic hazard due to basin amplification of seismic
waves. Finite difference modeling in the manner of Olsen et al. [1995] has not yet been
undertaken. It is quite possible that the modeling will reveal a significant basin
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amplification contribution to the seismic hazard in the West McCarran area of Reno. The
modeling should include the addition of several model earthquakes, such as magnitude
7+ earthquakes along the Genoa Fault and other possibilities listed in dePolo et al.
[1995]. As is shown in Olsen [2000], the effects of basin geometry on seismic waves can
be very different for different source azimuths, so many models should be run to get a
full picture of the potential hazard.
5.1.4 Test Ground Motion Predictions Using New ANSS Stations
In the coming years, the urban areas of Reno and Carson City will become home to
many new quality digital seismic stations as part of the Advanced National Seismic
System,
including
a
proposed
200
new
stations
for
Reno
alone
[http://quake.utah.edu/anss/Strawman.html]. Sites for the new stations are picked with
such criteria as geologic formation and topography in mind (John Anderson, personal
communication). I would suggest using depth-to-bedrock as a selection criteria as well.
Sampling a number of different gravity gradient and depth environments (e.g., shallow
and flat, deep and flat, deep and steep, etc.) and comparing actual seismic shaking to that
predicted using finite differencing would be extremely rewarding.
5.1.5 Shallow Velocity Measurements
As mentioned briefly in Chapter 2 and more extensively in Chapter 4, prediction of
amplification of seismic waves also requires the knowledge of local shear-wave velocity.
Measurements in a few different geological units (picked using the maps of Bell and
Garside [1987], Bonham and Bingler [1973], and Trexler [1977], could help in creating a
30-m shear wave velocity map for Reno and Carson City as Wills et al. [2000] did for the
state of California. The Wills et al. [2000] map has recently been correlated with some
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success with ground motions observed in southern California [Field, 2000]. This
recommendation would be particularly helpful if all the existing seismic stations and
proposed ANSS sites were measured so that detailed comparisons could be carried out.
5.2 Recommendations for Continued Dixie Valley Research
In Chapter 3, I have proved that a portion of the slip on the 1954 Dixie Valley
earthquake was accommodated along a low-angle reach of the fault. The strongest area of
proof is directly underneath the geophysical transects, with some additional information
coming from both the gravity survey of Schaefer [1982] and the geologic mapping of
Caskey et al. [1998]. Current evidence points to approximately 50% of the fault’s strikelength being characterized by low-angle dip, but this value is uncertain. Neither do we
have a firm handle on the mechanisms that allowed the system to slip at low angle. As the
only truly accessible active low-angle normal fault (as opposed to being on the bottom of
the sea or being a re-activated basement thrust), I believe the Dixie Valley fault warrants
additional study. Below are some financially attractive ideas to study both the kinematics
and the dynamics of the fault.
5.2.1 Static and Dynamic Stress Modeling
With the 1992 M=7.3 Landers, California earthquake came increased recognition
that dynamic strains associated with distal earthquake ruptures could trigger seismicity
[e.g. Hill et al., 1993; Anderson et al., 1994]. Statistically significant increases in
seismicity were noted up to 1250 km (17 fault lengths) from the source, ruling out static
stress changes as a possible mechanism [Hill et al., 1993]. I propose to study the
likelihood of Fairview Peak earthquake waveforms dynamically triggering the Dixie
Valley earthquake.
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The geometry and source-time functions of the Fairview Peak earthquake are known
from the body wave inversions of Doser [1987]. The aftershock studies of Stauder and
Ryall [1967] further constrain the dimensions of the Fairview Peak fault. We also know
the geometry of the Dixie Valley fault from Abbott et al. [2001]. This presents the
opportunity to resolve dynamic strains caused by the Fairview Peak event on the plane of
the ‘known’ Dixie Valley fault. Modeled velocity and reflectivity contrasts along the
fault will partition the incoming disturbance.
As a first step, I propose the re-evaluation of the work of Caskey and Wesnousky
[1997]. They modeled the static stress implications of the 1954 earthquake sequence
using the Coulomb3D program of Ross Stein. Caskey and Wesnousky [1997] did not,
however use the proper dip of the Dixie Valley fault in their model (the low-angle nature
of the fault being merely proposed at the time). The model should be changed to examine
the static stress implications of waves from the Fairview Peak sequence impinging upon a
low-angle Dixie Valley fault.
For the dynamic triggering situation, existing algorithms could be used to first model
the time-varying strain field in 1-dimension, and subsequently in 3 dimensions. For input
into the 3-d inversions, the depth to bedrock data compiled in Blakely et al. [1999] could
act as a starting point in the velocity modeling.
5.2.1.1 Differentiating among the possible mechanisms for fault weakening.
Armed with the dynamic models, it will be possible to conclude which mechanisms
might be responsible for allowing low-angle normal slip. Dynamic mechanisms that have
been proposed are:
High pore flujid pressures
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Transient pore-fluid pressure [e.g. Sleep and Blanpied, 1994; Axen, 2000]
Gouge Liquefaction
Acoustic Fluidization [e.g. Melosh, 1997]
Frictionless Rollers [e.g. Scott, 1996; Anooshehpoor and Brune, 1996]
Opening Modes [Brune, 1997]
Anomalous low frequency pulse [Anderson et al., 1994]
An extensive body of literature has been written on each of these mechanisms.
Strain modeled on the Dixie Valley fault will provide some evidence for and against the
individual mechanisms. For example, acoustic fluidization in the fault zone requires high
amplitude, high frequency strain [Melosh, 1996]. In contrast, the mechanism Anderson et
al. [1994] proposed for Landers-triggered seismicity was a long period (>10 s), high
amplitude pulse. This type of modeling will not uncontrovertibly solve the problem of
dynamic triggering by any means. Rather, it would provide some additional evidence for
this very interesting and difficult problem.
5.2.2 Electromagnetic Study of the Fault Zone
There exists, of course, other ways to study the consequences of differing fault
triggering mechanisms. In the following section, I discuss the motivation for attempting a
magnetotelluric investigation of the Dixie Valley region. The proposed study is to deploy
a 41 km long line, consisting of 15 MT stations, perpendicular to the fault. The line
would terminate at the Dixie Valley fault in the west, and extend over the Clan Alpine
Range to the east. The MT method, it is hoped, would be able to detect any briny, and
therefore conductive, fluids trapped in the fault zone.
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In a model proposed in Axen (1992), following Rice (1992), super-hydrostatic
pressured fluids from the ductile regime migrate up a permeable low-angle fault. Pressure
is maintained as minerals are deposited by released, decompressing high-temperature
fluids. The minerals continually seal the fractures releasing fault-zone overpressure. In
addition, low permeability chloritic breccia can help seal the fault. The high pore-fluid
pressure acts to change the principal directions of the stress tensor local to the
detachment, allowing low-angle slip. Another consequence of applying the overpressure
hypothesis to dipping, rather than strike-slip faults, is that seals on both sides of the fault
are not necessary. If an impermeable seal restricts fluid flow from the footwall to the
hanging wall, normal stress is reduced as the hanging wall “floats” with respect to the
footwall. An intriguing possibility is that high pore-fluid pressure, in conjunction with
static stress changes brought upon by the Fairview Peak event, led the Dixie Valley fault
to failure. There are many lines of evidence in Dixie Valley that leads one to believe that
elevated pore fluid pressures are contained within the fault zone.
5.2.2.1 Dixie Valley Evidence Supporting an Elevated Pore-Fluid Model
Northern Dixie Valley is a geothermal area. The hot (250 degree C) Dixie Valley
Geothermal system produces 65 Mw of energy [Hulen et al., 1999]. The geothermal
fluids are concentrated along the Dixie Valley fault [Hulen et al., 1999]. Greater than
hydrostatic pore-fluid pressure is sometimes seen in geothermal areas where downhole
temperatures are above 350 C [Grawinkel and Stockhert, 1997].
Forster et al. [1997] integrated outcrop studies and subsurface data from the
geothermal field. In outcrops along the fault they noted that: “Sample and outcrop
observations suggest that the breccia zones may represent localized sites of tectonically-
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driven hydrothermal implosion brecciation [Sibson, 1981] that originally produced zones
of high permeability that are now silicified.” [Forster, et al., 1997].
Barton et al. [1997] conducted borehole experiments at the geothermal plant and
noted a systematic relationship in the orientation of hydraulically conductive and nonhydraulically conductive fractures. The hydraulically conductive fractures had distinct
orientations and were preferentially oriented parallel to the fault zone [Barton et al.,
1997]. In addition, in situ stress measurement showed these permeable fractures to be in a
state of critical stress, and therefore close to brittle failure [Barton et al. 1997].
Mineralization along the Dixie Valley fault at the Dixie Comstock mine (approx. 13
km north of the 1954 ruptures) has led to economic concentrations of gold [Vikre, 1994].
The ore host consists of “quartz that cements quartz breccia, in quartz stockwork, veins
with quartz, pyrite, chalcopyrite, and montmorillonite in breccia clasts, and in veins of
incipiently crushed gabbro” [Vikre, 1994]. Mineralization of this sort is often seen in the
hanging walls above detachment faults, perhaps forming the impermeable cap that allows
pore fluid pressures to exceed hydrostatic levels [Axen, 1992]. Altered colluvium near
the Dixie Valley fault at the latitude of the Dixie Comstock mine shows that
hydrothermal fluids have been present at shallow levels in the Quaternary [Vikre, 1994].
These observations, although mostly restricted to the much more heavily studied
Dixie Valley Geothermal field, leads one to believe that the presence of elevated pore
fluid pressures may have played a role in the low-angle faulting. It is also possible that
the Fairview Peak earthquake caused a transient migration of pore fluids into the Dixie
Valley fault system. The permeable Dixie Valley fault may have in effect “tapped” into a
new reservoir of pore fluids at depth.
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Early, ca. 1975, magnetotelluric geothermal studies by Morrison et al. [1979] 50 to
100 km to the north of the Dixie Valley study area, first identified an unexpectedly
shallow (15 km +/- 4 km) low resistance (1 to 10 ohm-meters) zone over a series of
widely-spaced MT stations. This conductive horizon is the best candidate for a source of
over pressurizing fluids to facilitate low angle-normal faulting. If the horizon contains
brines in permeable rock, as has been widely proposed [Wannamaker et al., 1997] to be
consistent with observations, this fault geometry positions gently upward dipping lower
crustal units into the low-angle fault plane. It is conceivable this could provide the
necessary large reservoir for repeatedly injecting over pressured fluids into what is likely
an episodically depressurized fault system.
With regard to detecting the fluid filled fault zone, it must be said that this is a
possibility, but not a certainty. It is a useful point of reference that a ten meter thick fault
zone with a 0.5 ohm-meter resistivity has a conductivity thickness product (20 S) equal to
4 km of 200 ohm-m material and hence should be easily detected. However, a thinner
fault zone or smaller resistivity contrast with surrounding units would cause the zone to
be lost in the background. Forward modeling using a range of fault zone conductances
should be undertaken to define resolvability limits within the resistivity structure
recovered from the field survey.
Even if the fault system were unable to contain elevated pore fluid pressures over
geologic time, the fault may not have been able to release the new excess pressure fast
enough. The new fluids need not have migrated all the way to the surface, which would
be extremely unlikely given the time scale (approx. 4 minutes). The focal depth (approx.
12-15 km, [Doser, 1987; Okaya and Thompson, 1985]) indicates that rupture started at
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depth. Increased stress at the “crack tip” of the rupture, coupled with a critically stressed
fault, may have been enough to rupture the fault all the way to the surface.
5.2.3 Gravity Survey to Study Low-Angle to High-Angle
Accommodation Zones
As mentioned above, geologic evidence for low-angle rupture essentially terminates
to the north of Wood Canyon and to the south of Coyote Canyon (Figure 3.1). Gravity
lows in Dixie Valley mapped by Schaefer [1982] are laterally coincident with the lowangle rupture terminations mapped by Caskey et al. [1996] (Figure 3.2). This coincidence
argues for the existence of an accommodation zone as the range-front Dixie Valley fault
changes from high angle, to low angle, back to high angle, along this reach. This is not,
however, the only possible conclusion. Several mechanisms can create the observed
anomaly, but the present inadequate sampling of the anomalies makes differentiating
among the hypothetical mechanisms impossible. With increased data coverage one would
be able to constrain models in which the gravity lows are caused by:
Intrabasinal, “scissors-type” faults (Figure 5.1a). If it assumed that the range front
fault is marked by a change in dip at the mapped geologic terminations, east-west striking
normal faulting may account for the change in basin depth evidenced by the gravity lows.
Assuming that slip is constant along both high- and low-angle sections of the range front
fault, fault offset on the intrabasinal, east-west faults will increase to the east. Maximum
offset will be at the deepest portion of Dixie Valley; minimum offset would be along the
Stillwater range front. This is analogous to the increasing separation of blades away from
the hinge in open scissors. Maximum gravity gradients would be observed in north-south
transects.
110
111
Intrabasinal, north-south normal faults (Figure 5.1b). If it is assumed that the entire
rupture length of the Dixie Valley fault is low-angle, the gravity lows could be explained
by north-south striking normal faults in the basin. The gravity signature in this scenario
should show maximum gravity gradients in east-west transects over the anomaly.
Low density sediments in the hanging wall. The gravity lows may be caused by
density differences in hanging wall sediments unrelated to the normal fault system. The
gravity signature in this case would seem unpredictable (i.e. have any shape), but the
depth of the mass causing the anomaly would be shallow. Using a three-dimesnional
gravity inversion program, such as GRAV3D, one would be able to determine the
approximate depth to the anomalous mass, as discussed later. Therefore one could
determine if the anomaly is caused by basement deflections (as with faulting) or lowdensity deposits.
Sub-basins that formed during an earlier depositional phase.
difficult anomaly to characterize.
This is the most
The shape of the anomaly can not be predicted
beforehand and may exhibit the characteristics of any of the above three classes.
The spatial coincidence of the gravity lows and low-angle surface rupture
terminations suggest that case 1) is the most likely scenario. The use of 2.5-D and 3-D
modeling and inversion programs will aid in interpretation. Simple horizontal directional
derivatives of the gravity field will help delineate the strikes and locations of intrabasinal
faults.
5.2.3.1 Survey Plan
Approximately 1000 gravity measurements are proposed. The measurements will be
made on generally north-south and east-west transects, centered on the gravity lows in the
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transition zone (Figure 5.2). Transects will jog around small blocks of land owned by the
U.S. Navy, if necessary. Existing gravity coverage from Schaefer [1982] should be
merged into the dataset to complete the coverage. The additional coverage will allow the
calculation of the regional gravity gradient and avoid edge effects while modeling. To
differentiate gravity effects due to small-scale basins from broader, regional anomalies, a
"bedrock gravity" could be removed from the data set in the method used by Chapter 2.
5.2.3.1 Modeling and Interpretation
The reduced data will be modeled with 3-D inverse and 2.5-D forward modeling
programs. The 3-D modeling will be done with GRAV3D under free academic license
from the University of British Columbia Geophysical Inverse Facility (UBC-GIF)
Outreach Program. Details of the inversion can be found in Li and Oldenburg [1998]. The
inversion program has been successfully used on both synthetic and field data with good
results. A crucial element in gravity inversion is the ability to determine the depth to the
anomalous mass. GRAV3D allows the use of a depth weighting function to separate
anomalies caused by density contrasts at varying depths. This will be particularly useful
in discriminating between case 3) and cases 1) and 2) above. A case history relating to a
gravity inversion of the Voisey’s Bay deposit using this program may be found at
http://www.geop.ubc.ca/ubcgif/casehist/voisey/intro.html. Selected transects will be 2.5D forward-modeled using Northwest Geophysical Associates’ GM-SYS software
package. Initial depth estimates will draw from Abbott et al. [2001], and Meister’s [1967]
seismic refraction study.
113
114
5.3 Recommendations for Future Work in the Precarious Rock
Study
In many ways, the methodology used in Chapter 4 is in its infancy. In particular, the
acquisition and modeling of the ReMi data could be improved. In addition, more detailed
knowledge of how surface waves behave at this scale and the consequences of 3-D
geometry need to be better understood. Below are some suggestions.
5.3.1 Optimize the ReMi equipment
Cluster and sensitivity tests show that the 21-bit floating point digitizer in the Bison
seismic recorder used for the precarious rock study has sufficient dynamic range to
record data well below the corner frequency of our 8 Hz vertical sensors. Tests at the
Reno/Tahoe airport and UNR campus (not shown) using 4 Hz vertical sensors (RefTek
“Texans”) show interpretable energy down to 2 Hz (as opposed to 4 Hz in the best of
circumstances using 8 Hz sensors). My experience forward modeling the data has led me
to conclude that sampling the half-space velocity is very important to achieving a
satisfactory model. Proper half-space velocity is also needed if one is to calculate
synthetics accurately (witness the results of station MCS (Figure 4.7, Section 4.3.4)). At
MCS, energy to 2 Hz, rather than 4 Hz, might well have penetrated to the “basement”.
Another benefit to using the Texans was an increase in field productivity. The
current UNR setup consists of heavy refraction cables on spools. The takeout distance
(i.e. the distance between geophone connecters) is 15.24 m. Our optimal geophone
spacing of 8 m means that we are essentially carrying around 190% of the weight
necessary to complete the job. This is a critical factor in deciding where it is feasible to
115
make measurements. Potential sites in Chapter 4 were ruled out solely on the merits of
access problems.
The Texans ameliorate that problem significantly. Each sensor acts independently
without the need for a bulky cable. Each field assistant could easily carry 12 or more and
deploy them all in less than an hour over varied terrain. For a method that emphasizes
cost effectiveness and rapid results [see Louie, 2001], this is even more attractive.
5.3.2 Optimize for 30-m Work
The ReMi method, on virtue of its ease and cost, has a potential to be used
pervasively for 30-m average shear wave velocity modeling. The modeling results in
Chapter 4 demonstrate that additional work needs to be done in order to make the method
more accurate in the upper 30 m. In order to test whether the results of the tomography
and ReMi methods were complementary, I generated a dispersion curve given the
tomographic velocity model for Lovejoy Buttes. I then converted the dispersion curve
from velocity-period to slowness-frequency and displayed the result on the LJB p-f plot
to facilitate the comparison. Even assuming a 0.25 Poisson’s ratio, the dispersion curve
from the tomography follows the energy in the p-f plot remarkably well (Figure 5.3).
Clearly, the disagreement in the two methods was introduced during the forward
modeling. The obsolescent modeling program used in Chapter 4 might be partially to
blame, and alternate sources for modeling programs are being explored. Alternatively,
inversion of dispersion picks, rather than forward modeling might prove more accurate,
although I think that experienced and well-trained geophysicists can probably do a better
job given the right information. This brings me to my next point.
116
117
5.3.3 More ReMi Data in Varied Environments
Nothing teaches so well as experience, or, failing that, the experience of others. The
ease and accuracy of forward modeling would be improved if there was some knowledge
of how p-f plots look given different environments. Ideally, the ReMi measurements
would be made where independent measurements (such as borehole studies) have been
made. The “look and feel” of p-f plots and how they were successfully or unsuccessfully
modeled would aid the interpreter immensely. This step is currently underway with
measurements on soil, hard rock, and softer rock being compiled.
The quality of the ReMi data at Lovejoy Buttes and Piute Butte were especially poor.
We speculate that this is due to the heterogeneous sub-surface (e.g. alluvial fans between
rock outcrops and large boulders surrounded by sand). The poor data quality was
especially apparent when studying sources external to the Buttes (microtremor, as
opposed to the nearby hammer source ). An alternative hypothesis is that the region
surrounding the Buttes is too quiet and the urban noise that can dominate microtremor is
absent. I propose a test of both hypotheses in which the Buttes are instrumented with as
many as 200 RefTek Texan sensors. The sensors would be placed in an array surrounding
the Buttes with an array in the Buttes themselves. Simultaneous measurement of
microtremor at all stations will produce a “map” of ground motion on this small scale.
There would be the possibility of recording a distant train (10-15 km away?) that runs
parallel to the San Andreas Fault. This would facilitate measurement of waves
propagating with a known azimuth before, during, and after passage through the Butte
area.
118
5.3.4 Finite Differencing
The combination of the ReMi and tomography methods allows for an exciting
possibility. The velocity profile produced by the tomography can be simplified and used
as input to finite difference investigations of shallow surface wave propagation. Resulting
synthetic data can then be subjected to the same processing flows as the real data for a
site and comparisons can be made. A current working hypothesis, given my experience,
is that surface waves are not as well-developed in areas with highly variable velocity.
Alternatively, it may be that higher mode surface waves are obscuring the fundamental
mode by placing energy at different p-f coordinates. Subsequent spectral ratios would
decrease the signal-to-noise ratio (with noise consisting in this case of higher mode
surface waves) along the main dispersion curve. Either way, synthetic data would greatly
improve the understanding of both surface wave propagation and the ReMi method.
5.3.5 Measurements at Precarious Rocks in Different Environments
In Chapter 4, I mentioned that the conclusions on the site effects at the Mojave are
not necessarily applicable to precarious rocks at other locations, such as the Nevada Test
Site or the Sierra Nevada Mountains. This assertion needs to be verified. At first glance,
it would seem that the three environmental regimes would be very similar. The precarious
rocks in the Sierra share a similar lithology (fractured granite) but wetter climactic
conditions. The precarious rocks at the Nevada Test Site share a similar arid
environment, but different lithology (more volcanic rocks). Speculating, the basement
velocity in the Sierras might be virtually the same, with higher velocities near the surface
due to the wet soil. Velocities at the Test Site, on the other hand, might be very similar in
the upper 10 m, with different basement velocity.
119
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