Professional Paper 1661-D

U.S. Department of the Interior

U.S. Geological Survey

Crustal Structure of the Cascadia Fore 

Arc of Washington 

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Professional Paper 1661-D Velocity (km/s)

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Crustal Structure of the Cascadia Fore Arc of Washington

By Tom Parsons, Richard J. Blakely, Thomas M. Brocher, Nikolas I. Christensen, Michael A. Fisher, 

Ernst Flueh, Fiona Kilbride, James H. Luetgert, Kate Miller, Uri S. ten Brink, Anne M. Trehu, and Ray E. Wells

Professional Paper 1661-D

U.S. Department of the Interior

U.S. Geological Survey

Earthquake Hazards of the Pacific Northwest Costal and Marine Regions

Robert Kayen, Editor

Crustal structure and tectonic models of the subduction margin of Washington are derived 

from onshore and offshore seismic data and gravity observations.
U.S. Department of the Interior

Gale A. Norton, Secretary

U.S. Geological Survey

P. Patrick Leahy, Acting Director

U.S. Geological Survey, Reston, Virginia: 2005

For sale by U.S. Geological Survey Information Services

Box 25286, Denver Federal Center

Denver, CO 80225

This report and any updates to it are available online at:

http://pubs.usgs.gov/pp/pp1661-D

For additional information write to:

U.S. Geological Survey

Box 25046, Mail Stop 421, Denver Federal Center

Denver, CO 80225-0046

Additional USGS publications can be found at:

http://geology.usgs.gov/products.html

For more information about the USGS and its products:

Telephone: 1–888–ASK–USGS (1–888–275–8747)

World Wide Web: http://www.usgs.gov/

Any use of trade, product, or firm names in this publication is for descriptive purposes only and does not imply 

endorsement of the U.S. Government.

Cataloging-in-publication data are on file with the Library of Congress (URL http://www.loc.gov/).

Produced in the Western Region, Menlo Park, California

Manuscript approved for publication, November 21, 2005

Text edited by Peter Stauffer

Layout and design by Judy Weathers 

FRONT COVER

Cutaway view into the three-dimensional seismic velocity structure of the Washington margin calculated with traveltime observations from earthquakes, explosions, and airguns. Velocity was found by minimizing residual times between 

calculated and observed. Lower velocities (warm colors) characterize the accretionary prism rocks offshore, whereas 

higher velocities (cool colors) are seen in the Siletz River volcanic rocks, partly outlined by the white dashed lines. 

Surface geology and coastlines are draped onto the topography.
 iii

Contents

Abstract ................................................................................................... 1

Introduction ................................................................................................ 1

Tectonic setting .................................................................................... 2

Active-source seismic experiments in western Washington ................................. 4

Cross-section image of the Cascadia subduction zone and arc .................................... 6

The Juan de Fuca slab .............................................................................. 7

The accretionary complex and Siletz terrane ................................................... 7

The Cascadia volcanic arc and back-arc crust ................................................ 7

Implications of the two-dimensional model for earthquake hazards .............................. 9

Traveltime inversion for three-dimensional upper crustal velocity structure of western 

Washington ....................................................................................... 9

Traveltime data .................................................................................... 10

Velocity modeling methods ........................................................................ 11

Resolution .......................................................................................... 11

The three-dimensional velocity structure of western Washington: observation ............ 22

2.5-km depth ................................................................................. 22

7.5-km depth ................................................................................. 22

12.5-km depth ................................................................................. 22

17.5-km depth ................................................................................. 23

22.5-km depth ................................................................................. 23

The three-dimensional velocity structure of western Washington: interpretation ......... 23

7.5-km depth ................................................................................. 23

12.5-km depth ................................................................................. 24

17.5-km depth ................................................................................. 25

22.5-km depth ................................................................................. 25

Discussion of the three-dimensional seismic-velocity model of coastal Washington ............ 25

Three-dimensional crustal structure of the Puget Lowland from seismic traveltimes and 

Bouguer gravity ................................................................................. 28

Method ............................................................................................. 28

Resolution and test models ........................................................................ 30

Application and evaluation ........................................................................ 31

Recovery of the gravity anomaly ............................................................ 32

Comparison of velocity models with known geologic features ........................ 32

Evaluation from borehole velocity data ................................................... 37

Comparison of velocity models in cross-section view .................................... 37

Conclusions from seismic and gravity modeling of the Puget Lowland crustal structure ...... 37

Acknowledgments ....................................................................................... 37

References ................................................................................................ 37
iv 

Figures

1. Map showing major tectonic elements of the Cascadia subduction zone .................. 2

2. Map showing location of two-dimensional seismic traverse in relation to major geologic 

units, faults, and volcanoes in southwest Washington ....................................... 3

3. Map showing locations of the controlled and earthquake sources and recording stations

used to construct the three-dimensional velocity model ....................................... 4

4. Example of a seismic profile from explosive-source data collected on land in southwest

Washington .......................................................................................... 5

5. Example of a seismic profile from airgun data collected on land .............................. 5

6. Map showing location of Seismic Hazards Investigation in Puget Sound (SHIPS) seismographs and airgun tracks in northwest Washington .......................................... 6

7. Velocity-structure model across Cascadia subduction zone and arc ........................ 8

8. Example of local earthquake data collected on the vertical component ..................... 10

9. Horizontal slice of three-dimensional velocity model at 2.5-km depth ........................ 12

10. Horizontal slice of three-dimensional velocity model at 7.5-km depth ........................ 14

11. Horizontal slice of three-dimensional velocity model at 12.5-km depth ..................... 16

12. Horizontal slice of three-dimensional velocity model at 17.5-km depth ..................... 18

13. Horizontal slice of three-dimensional velocity model at 22.5-km depth ..................... 20

14. Average velocity-depth curves for major classes of sedimentary, metamorphic, and 

igneous rocks ....................................................................................... 24

15. Spaced vertical profiles though the three-dimensional velocity model for coastal 

Washington, with topography and geology draped above .................................... 26

16. Diagrams showing the tectonic model of Siletz-terrane deformation ........................ 29

17. Plot of observed velocity vs. density from Puget Lowland rock samples compared with 

Gardner’s rule ....................................................................................... 31

18. Graph showing velocity-depth profiles derived using traveltimes only and using both 

traveltimes and gravity in relation to seismic coverage with depth ........................... 31

19. Map-view images showing the checkerboard resolution tests at 5-km depth ............... 32

20. Maps showing distribution of gravity observations, SHIPS seismic sources and receivers,

and seismic-ray coverage in the Puget Lowland ................................................ 33

21. Map-view images comparing the gravity anomaly calculated from velocity models with 

the observed gravity anomaly in the Puget Lowland .......................................... 34

22. Horizontal slices from the three-dimensional model at 5-km depth derived from traveltimeonly and combined gravity-traveltime inversions ................................................ 34

23. Graphs showing comparison of tomography results with sonic-log data and lithologies for

six boreholes in the Puget Lowland ............................................................... 35

24. Comparative east-west cross section from the three-dimensional velocity models derived

from combined gravity and traveltime analysis and from traveltime modeling only ......... 36
Crustal Structure of the Cascadia Fore Arc of Washington

By Tom Parsons1, Richard J. Blakely1, Thomas M. Brocher1, Nikolas I. Christensen2, Michael A. Fisher1,Ernst Flueh3, 

Fiona Kilbride4, James H. Luetgert1, Kate Miller5, Uri S. ten Brink6, Anne M. Trehu7, and Ray E. Wells1

Abstract

This paper discusses three crustal-structure models 

developed because of the significant seismic hazard posed by 

Cascadia subduction margin to the highly populated Pacific 

Northwest region of the United States. The first is a 530-kmlong wide-angle onshore-offshore seismic transect across 

the subduction zone and volcanic arc recorded to study the 

major structures that contribute to seismogenic deformation. 

Observed were (1) an increase in the dip of the Juan de Fuca 

slab from 2 ̊-7 ̊ to 12 ̊ where it encounters a 20-km-thick block 

of the Siletz terrane or other accreted oceanic crust, (2) a 

distinct transition from Siletz crust into Cascade arc crust that 

coincides with the Mount St. Helens seismic zone, supporting 

the idea that the mafic Siletz block focuses seismic deformation at its edges, and (3) a crustal root (35-45 km deep) 

beneath the Cascade Range, with thinner crust (30-35 km) east 

of the volcanic arc beneath the Columbia Plateau flood basalt 

province. From the measured crustal structure and subduction geometry, two zones were identified that may concentrate 

future seismic activity: (1) a broad (because of the shallow 

dip), possibly locked part of the interplate contact that extends 

westward from ~25 km depth beneath the coastline perhaps as 

far as the deformation front ~120 km offshore and (2) a crustal 

zone at the eastern boundary between the Siletz terrane and 

the Cascade Range.

In addition to the two-dimensional cross-section model, 

two three-dimensional crustal models are discussed. The first 

is a regional model developed to study the Eocene mafic crust 

with high seismic velocities that underlies much of the Oregon 

and Washington fore arc and acts as a backstop for accretion 

of marine sedimentary rocks from the obliquely subducting 

Juan de Fuca slab. Arc-parallel migration of relatively strong 

blocks of this terrane, known as Siletzia, focuses upper crustal 

deformation and seismicity along block boundaries, which are 

potential sources of earthquakes. In a three-dimensional velocity model of coastal Washington, surface geology, well data, 

and traveltimes from earthquakes and controlled-source seismic experiments were combined to resolve the major boundaries of the Siletz terrane with the adjacent accreted sedimentary 

prism and volcanic arc. In southern Washington and northern 

Oregon, the Siletz terrane appears to be a thick block (~20 

km) that extends west of the coastline and makes a high-angle 

contact with the offshore accreted sedimentary prism. On its 

east flank, the high-velocity Siletz terrane boundary coincides 

with an en echelon zone of seismicity in the arc. In northern Washington, the western edge of Siletzia makes a lower 

angled, fault-bounded contact with the accretionary prism. 

In addition, alternating, east-west-trending uplifts and downwarps of the Siletz terrane centered on the antiformal Olympic 

Mountains may reflect focusing of north-south compression 

in the northern part of the Siletz terrane. This compressional 

strain may result from northward transport and clockwise rotation of Siletz terrane into the relatively fixed restraining bend 

of the Canadian Coast Mountains. 

The final crustal model discussed is one of upper crustal 

structure in the highly populated Puget Lowlands of Washington determined from data gathered during the 1998 Seismic 

Hazards Investigation in Puget Sound (SHIPS) experiment. A 

new method was developed to sequentially minimize seismic 

traveltime and observed gravity residuals in an iterative threedimensional inversion. Resolution tests were performed with 

synthetic traveltime and gravity observations from a checkerboard velocity model using the SHIPS experiment geometry, 

and it was shown that the addition of gravity significantly 

enhances resolution. The new velocity model for the region 

shows better correlation between surface geology and modeled 

subsurface velocity structure, and basin structures in particular 

were much better resolved. 

Introduction

Between 1991 and 1998, a series of controlled-source 

seismic experiments were conducted in Oregon and Washington to determine the velocity structure of the seismically active 

Cascadia convergent margin (Tréhu and others, 1994, Miller 

1

U.S. Geological Survey, Menlo park, California 94025

2

Department of Geology and Geophysics, University of Wisconsin, 

Madison, Wisconsin 53706

3

GEOMAR Research Center for Marine Geosciences, Kiel, Germany

4

Anadarko Petroluem Corporation, P.O. Box 1330, Houston, Texas 77251

5

Department of Geological Sciences, University of Texas at El Paso, El 

Paso, Texas 79968

6

U.S. Geological Survey, Woods Hole, Massachusetts 02543

7

College of Oceanic and Atmospheric Sciences, Oregon State University, 

Corvallis, Oregon 97331
2 Crustal Structure of the Cascadia Fore Arc of Washington

and others, 1997; Flueh and others, 1997; Parsons and others, 1998; Fisher and others, 1999). Crustal structure models 

derived from this work are here presented at three different 

scales: (1) a two-dimensional cross-section model through the 

subduction zone in southern Washington, (2) a three-dimensional seismic tomographic analysis of controlled-source and 

earthquake traveltime data aimed at resolving the large-scale 

geometry of the thick Eocene mafic basement of the Washington fore arc and its relationship to the accretionary prism and 

volcanic arc, and (3) a three-dimensional image of the upper 

crust of the Puget Lowland based on seismic and gravity data. 

Tectonic Setting

The Juan de Fuca Plate subducts beneath North America 

at a rate of ~40 mm/yr on a N68 ̊E azimuth (DeMets and others, 

1990). This oblique subduction has created a complex, geologically diverse, and potentially hazardous region, the Cascadia 

subduction zone and volcanic arc (fig. 1). No great earthquakes 

on the Cascadia subduction zone have been recorded in written 

history, and much of the region is relatively quiet seismically 

(Dewey and others, 1989). However, global comparisons indicate that the Cascadia subduction zone has many characteristics 

in common with those that produce great interplate earthquakes; 

for example, young oceanic lithosphere subducts at shallow dip 

(similar to Central and South America, southwest Japan, and the 

Aleutians) (Heaton and Kanamori, 1984; Heaton and Hartzell, 

1987). Recent studies of the Holocene geologic record have 

shown consistent indications that great subduction-zone and/or 

large upper-plate earthquakes have affected the Washington 

coastal margin. Interpretation of geologic evidence (subsidence, 

tsunami deposits) along the coast has suggested that great earthquakes (M>8) have occurred in the Cascadia subduction zone 

on a recurrence interval of hundreds of years (Atwater, 1996). 

The Cascadia margin has had a long and complex history of deformation and volcanism resulting from Cenozoic 

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Figure 1. Major tectonic elements of the Cascadia subduction zone. There are important structural variations along the margin. The 

accretionary complex broadens progressively along the margin, reaching its widest point at the Olympic Mountains. Volcanic production is greatest in the central arc, where the seismicity rate is lowest. 
Introduction 3

oblique convergence between North America and subducting 

oceanic plates. In the Washington fore arc, Cenozoic marine 

sedimentary rocks overlie Eocene basaltic basement of the 

Crescent Formation, which extend, along with the correlative Siletz River Volcanics of Oregon, from the southern tip 

of Vancouver Island to the Klamath Mountains (Snavely and 

others, 1968; Snavely and Wells, 1996) (figs. 1, 2). Sutured 

to North America at about 50 Ma, these voluminous (5-25 

km thick) submarine and subaerial basalts may represent an 

accreted oceanic island chain (Simpson and Cox, 1977; Duncan, 1982) or a hot spot-generated continental margin rifting 

event (Wells and others, 1984; Babcock and others, 1992). 

Paleomagnetic observations from volcanic rocks of the Eocene 

Crescent Formation in southwest Washington show significant 

(~20 ̊-50 ̊) clockwise rotations, as does the Cascade arc, the 

result of oblique subduction of the Juan de Fuca Plate (Wells 

and Coe, 1985; Beck and Burr, 1979). Marine sedimentary 

rocks accreted against the Eocene volcanic rocks and have 

been exhumed since ~14 Ma in the Olympic Mountains, maintaining a steady-state elevation since that time, with erosion 

balancing tectonic uplift (Brandon and others, 1998).

Some knowledge about the subducted Juan de Fuca slab 

beneath the continent has been acquired through controlledsource experiments, studies of regional earthquake hypocenters, 

and inversion of teleseismic arrival times for velocity structure. 

Taber and Lewis (1986) modeled a 9 ̊ dip on the Juan de Fuca 

slab beneath Grays Harbor. More broadly, the Juan de Fuca Plate 

appears to be arched beneath Washington along an axis oriented 

southwest-northeast that crosses the northern part of the Puget 

Sound (Crosson and Owens, 1987). Interpretation of teleseismic 

traveltime delays suggests a possible tear in the descending Juan 

de Fuca slab, with the dip angle steepening south of the Willapa 

Bay-Columbia River region (Michaelson and Weaver, 1986).

Seismic activity in western Washington is very low along 

the coast and increases inland, with hypocenters most abundant 

beneath the Cascade Range and Puget Sound (fig. 1). Deep earthColumbia River

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Figure 2. Location of two-dimensional seismic traverse (red line) in relation to major geologic units, faults, and volcanoes in 

southwest Washington.
4 Crustal Structure of the Cascadia Fore Arc of Washington

quakes associated with the Juan de Fuca slab seem to mirror 

the distribution of upper plate earthquakes in the fore arc. Two 

distinct north-northwest-trending zones of shallow seismicity 

are associated with Mount St. Helens and Mount Rainier (Stanley and others, 1996) (fig. 1). Heat flow in coastal Washington 

is low (~20-50 mWm-2), increasing to the east in the Cascade 

Range (~50-100 mWm-2) and Columbia Plateau (~50 mWm-2) 

(Blackwell and others, 1990).

Active-Source Seismic Experiments in Western 

Washington

In 1991, the U.S. Geological Survey (USGS), in collaboration with the University of Texas at El Paso, Oregon State 

University, the Geological Survey of Canada, the University of 

British Columbia, and the University of Wyoming, collected a 

series of refraction profiles in Oregon and Washington (Miller 

and others, 1997). Traveltimes from the northernmost refraction profile that trended along the eastern side of Puget Sound 

(fig. 3) were applied in this study. This profile was defined 

by 10 large explosive sources detonated into a 465-channel 

land array (~600 m spacing) that was about 300 km long. Six 

sources from that study are applied here (fig. 3).

In 1995, wide-angle seismic data were collected by the 

same collaborators along a 325-km-long east-west-trending 

profile that crossed southern Washington from Willapa Bay to 

the Columbia Plateau (fig. 2) (Parsons and others, 1998). On 

that profile, about 1,500 instrument deployments were spaced at 

200-m intervals and 17 large explosive sources were recorded. 

An example of these data is shown in figure 4. The explosion 

data show images of the subducting slab (from reflections) and 

provide continuous first arrivals to 230-km offsets.

In 1996, the German research vessel FS Sonne conducted 

an extensive investigation of the offshore Oregon and Washington margins in a joint effort by the Research Center for Marine 

Geosciences (GEOMAR) and the USGS (Flueh and others, 1997). 

A total of more than 14,500 air-gun sources (50-150 m spacing) 

were fired offshore of Washington, with about 6,000 detonated 

along lines instrumented by a cumulative total of 53 ocean-bottom 

recorder deployments (fig. 3). The balance of the air-gun sources 

were fired for marine multichannel profiles. On land, 44 Reftek 

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Figure 3. Locations of the controlled and earthquake sources and recording stations used to construct the three-dimensional velocity 

model. The generalized onshore geology is shown along with the locations of wells used (McFarland, 1979) to define the near-surface 

location of Siletzia, shown by the dashed black line. 
Introduction 5

seismographs distributed (~5 km linear spacing) along three 

profiles (fig. 3) recorded all the air-gun sources continuously. The 

ocean-bottom and on-land profileswere aligned along east-west 

profiles to provide continuous phase coverage across the margin 

and to enable comparison of structure from south to north (fig. 3). 

An example of air-gun data recorded on land is shown in figure 

5. The onshore-offshore data show headwaves traveling down the 

subducting slab as first arrivals in the near offsets and continental 

upper mantle refractions at longer offsets.

In March of 1998, the Seismic Hazards In Puget Sound 

(SHIPS) experiment was conducted in the connected waterways of the Puget Lowland. The SHIPS experiment consisted 

of onshore-offshore wide-angle and multichannel seismic 

(MCS) reflection profiling throughout the Puget Lowland 

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Figure 4. An example of explosive-source data collected on land in southwest Washington, showing the crustal refracted phase, Pg, 

that was used in the traveltime inversion for velocity structure. Reflected phases such as the one shown from the Juan de Fuca slab 

were not applied. Time scale was reduced as time less offset divided by 6.0 km/s (t-x/6.0).

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Figure 5. An example of airgun data collected on land showing the Pg phase and the upper mantle refracted phase, Pn, that were 

used in the traveltime inversion for velocity structure. Time scale was reduced as time less offset divided by 6.0 km/s (t-x/6.0).
6 Crustal Structure of the Cascadia Fore Arc of Washington

using an air-gun array (fig. 6). The total volume of the air-gun 

array varied between 110.3 litres and 79.3 litres, depending 

on whether wide-angle or MCS data were acquired. Wideangle profiling was conducted throughout the study region, 

even in narrow waterways such as the Hood Canal and Lake 

Washington where the multichannel seismic streamer could 

not be towed (Fisher and others, 1999). MCS profiling was 

performed in Puget Sound and the Strait of Juan de Fuca (fig. 

6). Air-gun shotpoint locations and times accurate to a millisecond were determined from global positioning system (GPS) 

navigation and GPS time recorded on the ship. The database 

includes 977,000 traveltime picks from controlled sources. 

The air-gun shots were recorded by a temporary array of 

210 seismographs deployed onshore and on the floor of Puget 

Sound (fig. 6) (Brocher and others, 1999). The SHIPS data 

were acquired with a shot spacing between 50 and 150 m and 

receiver spacing between 5 and 15 km. The quality of the wideangle data obtained during SHIPS is highly variable, although 

most stations provided useful data to source-receiver offsets of 

at least 40-50 km. At bedrock sites remote from urban centers, 

first arrivals can be observed to ranges as far as 200 km. On the 

other hand, few interpretable data were recorded at some of the 

soft soil sites in urban or suburban localities.

Cross-Section Image of the Cascadia 

Subduction Zone and Arc 

The 1995 wide-angle seismic profile was extended offshore in 1996, when the FS Sonne fired about 2,000 offshore 

air-gun sources (~25 m spacing) into 35 instruments located 

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Figure 6. Location of Seismic Hazards Investigation in Puget Sound (SHIPS) seismographs (dots) and airgun tracks 

(dashed and solid lines) in northwest Washington. Black squares show locations of deep wells with sonic velocity 

logs examined in this study.
Cross-Section Image of the Cascadia Subduction Zone and Arc 7

on the sea floor and on land (~5 km spacing). Traveltimes 

from refracted arrivals (crustal and upper mantle; Pg and Pn) 

were merged and inverted for velocity structure, and wideangle reflected arrivals (base of oceanic and continental crust; 

PmP) were forward modeled following the methods outlined 

by Hole and Zelt (1995) and Parsons and others (1996). The 

resulting velocity-structure model is shown in figure 7. The 

root-mean-squared traveltime misfits for refracted arrivals 

were less than 0.1 s, whereas misfits for all reflected arrivals 

were less than 0.15 s. The velocity model was used as input 

for a Bouguer gravity model, and primary model features fit 

well with the observed gravity data. In the following discussion, various features will be described by their location in km 

from the model origin along the profile shown in figure 7.

The Juan de Fuca Slab

The descending Juan de Fuca slab was imaged by refractions and reflections from offshore air-gun blasts and reversing reflections recorded from chemical explosions on land. 

Seismic data coverage extends about 200 km offshore (fig. 7). 

Normal-thickness (6 km) subducting oceanic crust encounters 

the deformation front at about km 60, where its dip increases 

from 0 ̊-2 ̊ to 3 ̊-5 ̊. Coincident multichannel reflection data 

show a series of landward-verging thrusts and fluid diapirs 

beginning at this point, implying low coupling stress along 

the interplate contact (Flueh and others, 1997). About 35 km 

offshore, the Juan de Fuca slab steepens from a 5 ̊-7 ̊dip to a 

12 ̊ dip, as determined from overlapping and reversed wideangle reflections from the base of oceanic crust recorded 

from land and marine sources (fig. 7). The observed 12 ̊ dip, 

slightly steeper than the 9 ̊ dip of Taber and Lewis (1986), 

persists to at least 50 km depth, 75 km east of the coastline 

(fig. 7). Slab earthquakes are located exclusively in the mantle 

part of the slab, counter to the global observations for young, 

warm slabs of Kirby and others (1996). Continuing the slab 

downward along a 12 ̊ dip places it at rather shallow depths 

(60-70 km) beneath the Cascade arc volcanoes. Either the 

slab dip steepens eastward of where it reaches 50 km depth, 

or it is young and warm enough to produce arc volcanism 

at 60-70 km depths. Seismicity appears to coincide with the 

Moho boundary in figure 7B, but this location is a result of the 

choice of events to plot; events were restricted to M 4.0 above 

25 km depth, but all events M 1.0 were allowed beneath 25 km 

depth to enable the best combination of highlighting zones of 

seismicity without obscuring the model with too many events. 

Smaller earthquakes are actually quite widely distributed 

throughout the continental crust (Stanley and others, 1996).

The Accretionary Complex and Siletz Terrane

At the edge of the continental shelf (km 120), the thickness of accumulated sediment scraped from the descending 

plate reaches 7 km. A 3-km-thick, low-velocity (2.0-3.0 km/s) 

sequence fills a basin just inboard (east) of the continental 

slope edge. A basement high intervenes between this and a 

second, deeper (5 km of sub-3.0-km/s sedimentary rocks) 

basin at about km 170. This part of the shelf and upper slope 

is undergoing extension and collapse as shown by widespread 

normal faulting (McNeill and others, 1997).

Near the coastline lies the surface contact between accretionary rocks and the Siletz terrane, an accreted oceanic block 

of late Paleocene-early Eocene age that forms the basement 

to much of the Cascadia fore arc. Snavely and Wagner (1982) 

map this contact from well and seismic-reflection data as an 

east-dipping thrust that partially cuts a thin Quaternary unit 

about 1 km offshore near Greys Harbor (~20 km north of the 

transect). The Siletz terrane is associated with a strong aeromagnetic signal along the profile where it crops out; this signal 

tapers off abruptly at the coastline (fig. 7A). 

Projecting the Siletz-accretionary boundary to depth on 

the basis of the velocity model is problematic; outcrops of the 

Siletz terrane (Walsh and others, 1987) (fig. 2) correspond with 

the 5.0 km/s contour in the model where it approaches the surface, and independent laboratory velocity measurements from 

the Siletz terrane taken from samples along the seismic profile 

show P-wave velocities between 4.5 and 5.2 km/s for near-surface pressures. However, similar velocities would be expected 

for metamorphosed sedimentary rocks of the accretionary 

complex. High-velocity rocks (6.5-7.7 km/s) are imaged in 

the lower continental crust beneath the surface outcrops of the 

Siletz volcanics in the Coast Ranges (km 190-230) and extend 

to the top of the descending slab, persisting to ~20 km offshore. 

If these high-velocity rocks are interpreted as Siletz terrane, 

then the total thickness offshore is 20 km, and the Siletz-accretionary boundary dips west from the coastline (fig. 7B). If, 

instead, the high-velocity lower crustal rocks offshore are more 

recently accreted oceanic crust, then the Siletz-accretionary 

boundary may dip east (Snavely and Wagner, 1982). There is a 

lateral velocity change at km 240 in the model evidenced by a 

sudden shallowing of turning rays (Pg) (fig. 7C), where highervelocity rocks (> 6.5 km/s) are found at shallower depth (10-11 

km as compared with 17-18 km to the west). This shallowing 

of higher-velocity rocks to the east might result from an eastdipping Siletz boundary (fig. 7). The maximum thickness of the 

Siletz volcanics beneath the Coast Ranges could reach 35 km if 

they extend to the base of the crust (fig. 7B).

The Cascadia Volcanic Arc and Back-Arc Crust

The velocity model shown in figure 7 provides a cross 

section through the Cascade Range and the transition from the 

primarily collisional fore arc to the extensional back arc. A 

significant lateral change in upper-crustal velocity is observed 

across the Mount St. Helens seismic zone (km 340); the 6.5 

km/s contour is deeper on the east side of the seismic zone, 

whereas the 6.0 km/s contour is shallower (fig. 7B). This 

boundary is interpreted as the eastern extent of the Siletz terrane and transition into Cascade arc crust, and the 6.0-6.5 km/s 

zone beneath the volcanic arc is interpreted as silicic intrusive 
8 Crustal Structure of the Cascadia Fore Arc of Washington

rocks on the basis of laboratory velocity measurements made 

on plutonic inclusions taken from the Mount St. Helens lava 

dome (Paine, 1982). This seismically determined position for 

the eastern extent of the Siletz terrane and transition into the 

Cascade arc is in reasonable agreement with the boundary 

modeled from gravity data (Finn, 1990). East of the Cascade 

Range, the combined thickness of Columbia River Basalt 

flows and subbasalt sedimentary units ranges from 6 to 8 km, 

in reasonable agreement with the more detailed studies of 

Saltus (1993) and Jarchow and others (1994).

In the region east of about km 220, turning rays extend 

only to about 10-12 km depth. Thus crustal velocity beneath 

that depth was not directly measured by refracted waves but 

was instead modeled from arrivals reflected off the Moho 

(PmP). East of km 220, turning rays were refracted at about 

10-16 km depth along a horizon that represents rocks having 

6.5 km/s velocities; beneath the 6.5 km/s contour, a velocity 

gradient from 6.5 km/s to 7.0 km/s is shown, and a crustmantle transition from 7.5 to 7.7 km/s (fig. 7B). A number of 

crustal velocity gradients were tested, ranging from a uniform 

6.5 km/s crust to a 6.5-7.5 km/s gradient, but did not introduce 

low-velocity zones. The 6.5-7.0 km/s gradient produced the 

smallest traveltime residuals for continental PmP reflections 

(traveltime misfit to within 0.15 s). The crust-mantle transition zone was required to fit reflection traveltimes from the 

descending slab beneath the crust and is consistent with low 

observed amplitudes of the continental PmP and Pn phases. 

The model shown in figure 7 assumes no lateral lower-crustal 

0 20 40 60 80 100 120 140 160 180 200 220 240 260 280 300 320 340 360 380 400 420 440 460 480 500 520

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Velocity Model

MAGNETIC FIELD STRENGTH, IN GAMMAS DEPTH, IN KILOMETERS DEPTH, IN KILOMETERS

Traverse Line

Coverage: Reflected and Refracted Rays

DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

Figure 7. Velocity-structure model across Cascadia subduction zone and arc. A, Aeromagnetic profile coincident with the twodimensional seismic profile. B, Seismic velocity model across the Cascadia subduction zone and volcanic arc. Model features are 

discussed in text. C, Coverage diagram showing where the two-dimensional model is constrained.
Traveltime Inversion for Three-Dimensional Upper Crustal Velocity Structure of Western Washington 9

velocity contrasts; thus changes in crustal thickness could also 

be modeled by lateral velocity changes. Attempts to model a 

flat Moho by inserting lateral velocity contrasts led to arbitrary 

and in some cases impossible crustal velocities. 

Continental crustal thickness increases from ~20 km 

at the coastline to ~30 km beneath the Coast Ranges and 

Chehalis basin. The crust is thickest (40-45 km) beneath the 

Cascade Range, where there is a broad root. The crust thins to 

about 30-38 km beneath the Columbia Plateau, slightly thinner 

than in the model of Catchings and Mooney (1988). A weak 

but measurable Pn phase was observed from three explosive 

sources and shows upper-mantle velocities of about 7.8 km/s 

at the base of the crust-mantle gradient zone (fig. 7). Upper 

mantle velocities were measured at 7.8-7.9 km/s beneath the 

oceanic crust offshore.

Implications of the Two-Dimensional 

Model for Earthquake Hazards

The image of the subduction process emerging from the 

model is one of a young (~10 m.y. old) and buoyant (Heaton 

and Hartzell, 1987), shallow-dipping (12 ̊) Juan de Fuca Plate 

that appears to bend as a result of overriding contact with rocks 

accreted to the North American Plate. This observed subduction 

geometry concentrates the potentially locked interplate contact 

offshore; the shallow dip broadens the zone out to the deformation front (~120 km offshore), and the steeper part of the 

slab encounters 350 ̊C temperatures at about 25-30 km depth 

(350 ̊C is thought to be the temperature at which stable sliding begins; Hyndman and Wang, 1993) near the coastline (fig. 

7B). A similar seismogenic depth limit was observed beneath 

Mexico, where the comparably young Cocos Plate subducts 

(Tichelaar and Ruff, 1993). The interplate contact above the 

approximate depth of the 350 ̊C contour is devoid of even very 

small earthquakes (fig. 7B). The low coupling stress along the 

interplate contact—as inferred from the presence of landwarddipping thrusts, high pore-fluid pressures (Flueh and others, 

1997), and low frictional heating—implies that on the time scale 

of the seismic cycle, earthquakes are likely to be associated with 

a complete stress drop (Wang and others, 1995).

In the upper plate, the strong, mafic Siletz terrane rocks 

apparently play an important role in delimiting earthquakes 

along their eastern boundary. The correspondence between the 

eastern edge of the Siletz block and the onset of seismicity in 

the Mount St. Helens zone (fig. 7B) may result from clockwise 

rotation (determined from paleomagnetics) of the coherent 

Siletz block that concentrates seismicity at its edges and limits 

internal seismogenic deformation (Wells, 1990; England and 

Wells, 1991; Tréhu and others, 1994; Stanley and others, 

1996). At present, the western edge of the Siletz terrane is not 

a focus of seismic activity. Judging from the apparent influence of this boundary on accretionary-complex deformation 

offshore of Oregon, where the western edge is overlain by 

short-wavelength folds and faults (Tréhu and others, 1995), 

it is possible that an analysis of sedimentary structures in this 

region would reveal similar effects. However, the details of 

this process likely depend strongly on the dip of Siletz terrane and on the nature of the high velocity material at depth 

beneath the shelf (Byrne and others, 1993). 

Traveltime Inversion for ThreeDimensional Upper Crustal Velocity 

Structure of Western Washington

Eocene-age volcanic rocks underlie most of the fore arc 

in Oregon and Washington and are called the Crescent Formation in Washington and the Siletz River Volcanics in Oregon 

(Snavely and others, 1968). Because of its mafic composition, this fore arc terrane, also known as the Siletz terrane, or 

“Siletzia” (Irving, 1979), is thought to be composed of strong 

crustal blocks that play an important role in fore-arc deformation (Magill and others, 1981; Wells and Coe, 1985; Wells and 

Weaver, 1993; Tréhu and others, 1994; Stanley and others, 

1996). Arc-parallel migration of these blocks in response to 

oblique subduction focuses upper crustal deformation and 

seismicity along block boundaries, which may be potential 

sources of earthquakes (Wells and others, 1998).

Along much of the Cascadia subduction margin, accreted 

sedimentary rocks are thrust beneath Siletzia along a major 

terrane-boundary fault presumed to dip eastward beneath the 

Coast Range (Tabor and Cady, 1978; Snavely, 1987). Though 

this boundary is an important locus of strain-accommodation, 

it is presently seismically quiet and its earthquake potential 

is unknown. In the southwest Washington arc, however, the 

northwest-trending Mount St. Helens and west Rainier seismic 

zones are thought to mark the eastern extent of Siletzia (e.g. 

Stanley and others, 1996; Parsons and others, 1998).

The two-dimensional cross-section velocity profile, while 

providing a relatively high-resolution image of the velocity 

structure across the margin, is only a single cross section and 

cannot constrain the dip of the boundary between Siletzia 

and the accretionary complex because of the wide range of 

possible velocities in the metamorphosed accreted rocks at 

shallow depths. Possible interpretations of the velocity structure include landward or seaward dips (fig. 7). Multiple cross 

sections tied to the near-surface geology offer a better chance 

to constrain the dip. Thus, one of the goals for conducting 

the three-dimensional study was to get a greater variety of 

raypaths through the accretionary and Siletz terrane rocks than 

were recorded along the two-dimensional profile.

The Cascadia subduction margin shows many alongstrike tectonic variations, including accretionary-prism 

width, seismicity rate, and volcanic production (fig. 1). 

Because of the important role that the Siletz terrane apparently plays in shaping the margin, defining the along-strike 

variation of its boundaries may provide some of the constraints needed to understand variations in the nature of 
10 Crustal Structure of the Cascadia Fore Arc of Washington

deformation and earthquake potential along major upper 

plate structures in the Washington fore arc.

Traveltime Data

A total of 69,251 P-wave first-arrival traveltimes from 

controlled and earthquake sources were included in the velocity 

modeling. As can be seen in figure 1, coastal Washington is 

seismically quiet; thus very few earthquake-source arrival times 

were available for the coastal regions. The majority (67,898) 

of the traveltimes were from controlled sources, with only 

1,353 arrivals used from earthquakes (fig. 8). There are some 

advantages in applying primarily controlled sources; two of the 

free parameters of a simultaneous inversion, source location 

and origin time, are known, reducing uncertainties. However, 

all the controlled sources were located at the surface, limiting 

the deeper coverage that would result had more earthquake 

sources been available. For the velocity inversion, only first 

arrivals were included, either Pg, the crustal refracted phase, 

or Pn, the upper-mantle refracted phase. All controlled-source 

seismic data were hand picked on a computer screen display; 

the estimated picking errors are 100 ms (one cycle at 10 Hz). 

Data examples are shown in figures 4 and 5. 

The limited set of earthquake traveltimes came from two 

data sets. The first group of earthquake-source traveltimes was 

recorded by the temporary array deployed during April and 

May of 1996 (figs. 3,8). During this period ~300 earthquakes 

were recorded, distributed across western Washington (green 

stars on fig. 3) and including the M=5.4 Duvall earthquake 

that occurred ~30 km east of Seattle and many of its aftershocks. Only the vertical channel was recorded on the temporary network. A total of 267 high-quality arrivals that could 

be used in the velocity inversion were recorded on the 44 

distributed stations. A high-quality event was defined as one 

recorded by at least five permanent network stations from the 

Pacific Northwest Seismic Network (PNSN), and with a root 

mean square (RMS) error in location less than 2 km. 

The second group of earthquake arrivals came from 

selected PNSN permanent stations that were located near the 

coast (pink triangles on fig. 3). Data from a small group of 

well-located events (1995-1997, M≥2.0, yellow stars on fig. 

15

75

20

25

30

RADIAL OFFSET, IN KILOMETERS

Temporary Network Earthquake P-wave Data: Olympic National Park

150 150

Duvall Aftershocks

TIME, IN SECONDS

Figure 8. An example of local earthquake data collected on the vertical component, showing the Pg phase. Because many 

of the events recorded are aftershocks from the Duvall earthquake, they are not plotted at an equal offset scale; if they were, 

most of the Duvall events would be plotted on top of each other. Instead, the events are plotted at a variable offset scale.
Traveltime Inversion for Three-Dimensional Upper Crustal Velocity Structure of Western Washington 11

3) were recorded at 12 sites. The earthquake source regions 

in western Washington are dominantly the Puget Sound 

and Cascade Range areas; enough events were used to get 

a reasonably uniform sampling of travel paths through the 

three-dimensional model without clustering too many events 

in specific locations. The goal was to solve for the broad-scale 

(25 to 50 km) velocity structure; therefore, adding large numbers of coincident sources would have added to computation 

times without significantly improving resolution. 

Velocity Modeling Methods

The three-dimensional tomographic technique of Hole 

(1992) was modified to simultaneously invert for velocity, 

hypocenters, and origin times (hypocenters and origin times 

only for the earthquake data). This technique applies a finitedifference solution to the eikonal equation (Vidale, 1990; 

updated by Hole and Zelt, 1995) to calculate first arrival times 

through a gridded slowness model. An iterative nonlinear 

inversion is performed as a backprojection along raypaths 

determined from the forward modeling step. 

A larger three-dimensional volume was modeled than that 

containing the targeted coastal region where the Siletz terrane 

meets accreted sedimentary rocks so that important seismic 

source regions could be included, such as the offshore airguns 

and the seismically active Puget Sound region (fig. 3). The 

choice of a large model volume created regions of relatively 

sparse ray coverage. Therefore, as always, the choice of a 

starting model was important in guiding the final model. A 

smoothed version of the detailed two-dimensional cross-section model was extrapolated parallel to the margin to generate 

a three-dimensional starting model. In many cases headwaves 

were observed from the dipping Juan de Fuca slab as first 

arrivals from the offshore airguns; thus, having the approximately correct slab structure in the starting model allowed for 

the proper treatment of these arrivals. Starting models were 

discretized into grids of 2.5-km cells; relatively small grid 

cells were used to ensure accurate calculation of ray paths 

along short source-receiver offsets.

Traveltime picks were compiled for controlled (figs. 4, 5) 

and earthquake (fig. 8) sources for each receiver as a function of their three-dimensional spatial source locations and 

then inverted for three-dimensional velocity structure. Initial 

hypocenter locations and origin times of earthquakes were 

input as determined by the PNSN. A spatial smoothing filter 

was applied to the models between velocity and source-parameter iterations. Early iterations were conducted that applied 

very broad smoothing filters (up to 100 km) and limited 

source-receiver offset ranges to solve the shallowest parts of 

the velocity model first. Subsequent iterations were conducted 

that included greater source-receiver offsets and progressively 

smaller smoothing filters. 

The earthquakes used in this study were initially located 

with a simple vertical (one-dimensional) velocity model; 

thus a degree of coupling between hypocenter location and 

the velocity structure derived from earthquake traveltimes 

is unavoidable and could cause significant errors in the 

resolved velocity models (Thurber, 1993), especially in 

areas where controlled sources were absent. To reduce such 

errors, hypocenters and origin times were relocated while 

controlled-source locations and times were held fixed. The 

events were relocated between velocity iterations (mean 

relocation was 0.28 km). See Hole (1992) for full details on 

the traveltime inversion algorithm. 

Resolution

Resolution in tomography depends on three properties 

of the problem—the signal band width, the source-receiver 

distribution, and the velocity structure itself. For crustal-scale 

experiments the signal wavelengths are typically much smaller 

than the gaps in raypath coverage (as they are in this study, 

2-20 Hz). Three approaches are usually adopted to investigate 

resolution in tomographic problems. The simplest is a hitcount analysis. In this analysis, the number of rays sampling a 

given cell are examined to identify regions of good coverage 

and poor coverage. The second approach to resolution analysis 

is the construction of synthetic tests using the data distribution 

(Humphreys and Clayton, 1988). The synthetic test may be an 

attempt to construct point-spread functions, or it may be an 

attempt to reconstruct the major features of the model simultaneously. The third common method of resolution analysis 

is the use of the resolution matrix of linear inverse theory. 

Typically the diagonals of the resolution matrix are displayed, 

and a certain value is chosen to indicate good resolution. The 

resolution matrix is a construct well suited to the study of linear problems. However, the extension of this tool to nonlinear 

problems is always questionable, particularly when the solution is approached iteratively (Shaw and Orcutt, 1985). 

Each of the above resolution diagnostics depends on 

the velocity structure used to construct the resolution measures. Quantitatively connecting a hit count, synthetic test, or 

resolution matrix with the actual accuracy of the reconstructed 

image is not straightforward. A combination of these resolution indicators can provide some intuition into the resolving 

power of the data. A backprojection was used, thus no formal 

resolution matrix was calculated. Instead, a hit count is shown 

to illustrate the seismic ray coverage and checkerboard tests 

used to estimate the degree of uniqueness of the solution. 

The estimates of spatial and relative velocity resolution based 

on the checkerboard tests vary with depth, and they are thus 

reported individually for slices from the three-dimensional 

velocity model. The checkerboard tests were conducted by calculating synthetic traveltime picks between all the source and 

receiver positions through a model of vertical columns, 50 km 

by 50 km, each with alternating increasing velocity gradients 

that were 0.5 km/s different at all depths (figs. 9-13). Additionally, vertical gradients at various depths were alternated (to 

create checkerboards in cross section) to test the resolution of 

horizontal velocity boundaries. Only the resolution of increasing velocity gradients within column elements were tested.
12 Crustal Structure of the Cascadia Fore Arc of Washington

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DISTANCE, IN KILOMETERS

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A

Figure 9. Velocity model slice at 2.5-km depth. Full discussion of the model slice and resolution test are in the text. See 

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Traveltime Inversion for Three-Dimensional Upper Crustal Velocity Structure of Western Washington 13

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B

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14 Crustal Structure of the Cascadia Fore Arc of Washington

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DISTANCE, IN KILOMETERS

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A

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Traveltime Inversion for Three-Dimensional Upper Crustal Velocity Structure of Western Washington 15

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DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

A

Figure 11. Velocity model slice at 12.5-km depth. Full discussion of the model slice and resolution test are in 

the text. See figure 3 for explanation of symbols. A, Horizontal slice from the three-dimensional velocity model 

volume taken from 12.5 km below sea level. B, Results from a checkerboard resolution test, with the recovered 

model shown above the input model. 

16 Crustal Structure of the Cascadia Fore Arc of Washington
Figure 11.—Continued.

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DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

B

Traveltime Inversion for Three-Dimensional Upper Crustal Velocity Structure of Western Washington 17
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Ray Coverage

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Mountains

DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

A

Figure 12. Velocity model slice at 17.5-km depth. Full discussion of the model slice and resolution test are in the text. See 

figure 3 for explanation of symbols. A, Horizontal slice from the three-dimensional velocity model volume taken from 17.5 km 

below sea level. B, Results from a checkerboard resolution test, with the recovered model shown above the input model. 

18 Crustal Structure of the Cascadia Fore Arc of Washington
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Actual Model

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DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

B

Figure 12.—Continued.

Traveltime Inversion for Three-Dimensional Upper Crustal Velocity Structure of Western Washington 19
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Ray Coverage

Grays Harbor

Willapa Bay

Olympic

Mountains

6.5

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1

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DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

A

20 Crustal Structure of the Cascadia Fore Arc of Washington

Figure 13. Velocity model slice at 22.5-km depth. Full discussion of the model slice and resolution test are in the text. See figure 3 for explanation of symbols. A, Horizontal slice from the three-dimensional velocity model volume taken from 22.5 km below 

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Actual Model

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DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

B

Figure 13.—Continued.

Traveltime Inversion for Three-Dimensional Upper Crustal Velocity Structure of Western Washington    21
An important aspect to tomography is scaling the smoothing dimensions appropriately to the input traveltime data. 

Based on the distribution of seismic sources and receiver stations, resolution of velocity anomalies ≥ 50 km across in the 

lateral dimensions and ≥ 10 km thick in the vertical dimension 

was sought. Thus a 25-km-wide by 5-km-high smoothing filter 

was applied during the final iteration, which yielded an RMS 

traveltime misfit of 0.31 s. A smaller RMS misfit could be 

achieved, but this would require reducing the model smoothness below the appropriate scale for the input data coverage, 

resulting in an artificially detailed velocity model.

The Three-dimensional Velocity 

Structure of Western Washington: Observation

Horizontal slices from the three-dimensional model 

volume at 5-km depth intervals that are windowed around the 

regions of best data coverage, as illustrated by the accompanying hit-count plots, are shown in figures 9-13. An associated 

checkerboard test for each velocity slice is also shown. All 

reported depths are depths below sea level.

2.5-km Depth

The model had limited coverage at this depth (fig. 9A). 

Many of the velocity features shown are actually inherited 

from vertical smoothing of anomalies resolved deeper in the 

model. The 2.5-km model slice does show fairly uniform 

velocities of about 4.5 km/s in the Coast Ranges, where 

there is good coverage. The accompanying checkerboard test 

result (fig. 9B) demonstrates the relatively poor resolution 

in the uppermost layer. Some of the checkerboard pattern 

was recovered, though vertical smoothing of deeper, better 

resolved parts of the test model may have influenced that. 

Because of the limited coverage at shallow depths away from 

the controlled-source transects, the low-velocity Seattle and 

Tacoma Basins beneath Puget Sound (Lees and Crosson, 1990; 

Symons and Crosson, 1997) were not imaged except very near 

to the 1991 land refraction profile (figs. 9A and 3).

7.5-km Depth

The seismic-ray coverage improves with depth into the 

middle crust and is relatively good in the center of the model 

at 7.5 km (fig. 10A). Offshore, fairly low velocity rocks (3.5-

5.5 km/s) are imaged. Onshore, a large body of relatively high 

velocity rocks (>6.0-6.5 km/s) occupies much of the crust, 

extending from near the coastline to the east of Puget Sound 

(fig. 10A). The high-velocity rocks are present at the coastline 

in northern Oregon and southern Washington; farther north 

along the coast, the high-velocity rocks retreat inland at Grays 

Harbor and again at the Olympic Mountains. A pattern of 

alternating east-west-trending bodies of high and low velocity 

is evident along the coastline.

The resolution test at 7.5-km depth shows reasonably 

good recovery of the checkerboard pattern (fig. 10B), with 

edges resolved to within ± 10 km laterally and relative velocities resolved to an average of ± 0.15 km/s. Because of the 

broad lateral smoothing applied to the velocity model, dips 

cannot be resolved very accurately. A ± 10-km lateral shift at 

7.5-km depth implies a ± 55 ̊ change in dip between a surface 

outcrop and the associated velocity anomaly. At this depth, 

therefore, only very high angle vs. very low angle velocity 

boundaries can be discriminated. 

The observation of an alternating pattern of high and low 

velocities along the coast looks at first glance to be a result of 

a coverage bias, in that the three east-west onshore-offshore 

profiles (where coverage tends to be best) are associated with 

lower velocities (figure 10A). However, the two southern lines 

were located in basins, and the northenmost line was located 

on the accretionary rocks of the Olympic Mountains, where 

velocities are expected to be lower. South of the Olympic 

Mountains, the high velocities were modeled from traveltimes to PNSN stations. Immediately south of Grays Harbor, 

the high velocities may not extend all the way to the coast as 

shown, given that there is a hole in the coverage there (fig. 

10B). However, the high velocities modeled farther south near 

the coastline around the mouth of the Columbia River are better constrained (fig. 10B). The generally high velocities modeled north and south of Grays Harbor are not artifacts from 

the starting model; instead, they result from the addition of the 

PNSN data. If the PNSN data are excluded from the inversion, the observed velocity pattern is one of lower velocities 

smeared across the coastal areas. The addition of the PNSN 

stations reduced the overall RMS misfits; therefore, the higher 

velocity areas are not a result of any systematic timing errors 

associated with the earthquake traveltimes.

It is possible that some vertical smearing of very low 

velocities near the surface could be influencing the velocity 

image where low-velocity anomalies are observed because of 

poor coverage in the upper 2 km. However, the low velocities associated with these regions persist through the 12.5-km 

depth range (see subsequent discussion), exceeding the vertical smoothing dimension (5 km), and passing through the best 

covered parts of the model, and they are sampled by rays from 

sources and receivers away from the areas at the surface that 

correspond with the lower-velocity regions.

12.5-km Depth

At 12.5-km depth, an overall northeast trend of the west 

edge of higher velocity rocks (>6.5 km/s) from Willapa Bay in the 

south to Puget Sound in the north becomes clear. An embayment 

of lower velocity rocks (<6.0 km/s) at Grays Harbor is still evident 

at 12.5-km depth (fig. 11A). To the east, high velocities drop off 

dramatically along a north-trending boundary that corresponds 

roughly to the west Rainier seismic zone.

Lateral spatial resolution at 12.5-km depth is comparable to 

that at 7.5-km depth (± 10 km) (fig. 11B), though the recovered 

22 Crustal Structure of the Cascadia Fore Arc of Washington
relative velocity resolution is a little worse at about ± 0.2. km/s. 

The checkerboard pattern recovery continues to be better along the 

western part of the model window than along its east side.

17.5-km Depth

The pattern of higher (>6.5 km/s) and lower (<6.5 km/s) 

velocity rocks at 17.5-km depth resembles that at 12.5-km 

depth (fig. 12A). The western edge of the higher velocity block 

continues to have a northeast trend. Higher velocity rocks are 

observed in the offshore region beginning at this depth. Spatial 

(± 20 km) and velocity resolution (± 0.3. km/s) are reduced 

relative to shallower slices (fig. 12B).

22.5-km Depth

The relatively consistent observation of a higher velocity 

body of rocks onshore as compared with the offshore region 

between 2.5-km and 17.5-km depth is lacking at this depth (fig. 

13A). Instead, fairly uniform 6.5-7.0 km/s rocks are continuously distributed. Vertical and horizontal resolution tests were 

conducted on the onshore high-velocity anomaly at 22.5-km 

depth. The checkerboard pattern was shifted from that in layers above to test the lateral resolution while eliminating the 

possibility that the recovered pattern might incorporate some 

smoothing from layers above where there was better seismicray coverage. It was found that the checkerboard pattern could 

be recovered to some extent, with a lateral resolution of ± 

15-20 km, and ± 0.2 km/s resolution in velocity. Thus, if there 

was a lateral velocity contrast such as was observed at all other 

depths in the model, it should be evident at 22.5 km depth.

The Three-dimensional Velocity Structure of 

Western Washington: Interpretation

The primary goal in generating the three-dimensional 

velocity model was to identify the boundary between Siletzia and the sedimentary rocks accreted against it. Siletzia in 

western Washington is made up of the Crescent Formation and 

comprises massive and pillowed basalt flows cut by diabase 

dikes (Tabor and Cady, 1978). In contrast, where exposed in 

the Olympic Mountains, the sedimentary rocks accreted against 

Siletzia are primarily sandstones and mudstones of increasing 

metamorphic grade from east to west (Tabor and Cady, 1978; 

Brandon and Calderwood, 1990). Thus it is expected that Siletzia rocks are much higher velocity than the accreted sedimentary complex. To interpret the actual boundary between these 

rocks, their velocity-depth functions are needed; although some 

data exist for the Siletz terrane (N.I. Christensen, written communication, 1999; Parsons and others, 1998), little is known 

about the accretionary complex. However, some insight can 

be gained from a global compilation of velocity measurements 

that shows a fairly narrow transition from metamorphosed 

sedimentary rocks into metamorphosed mafic igneous rocks 

(Christensen and Mooney, 1995). Given the coarse smoothing 

applied to the three-dimensional model and resulting velocity 

resolution (~± 0.2 km/s), such an approach provides reasonable 

guidelines for interpretation. 

In figure 14, the results of Christensen and Mooney 

(1995) are summarized graphically; the metasedimentary rocks 

at depth fall below a velocity range of ~6.1-6.4 km/s, while 

the mafic intrusive rocks tend to fall above that curve, with 

the exception of unmetamorphosed basalt. Where exposed in 

Washington, Siletzia rocks are zeolite to prehnite-pumpellyite 

facies (Tabor and Cady, 1978). Thus the velocity range between 

6.0 and 6.5 km/s is interpreted as containing the Siletz-accretionary boundary, which takes into account the average ± 0.2 

km/s lateral velocity resolution (fig. 14). In Washington no 

ultramafic rocks are exposed at the surface in outcrops of Siletzia, but on Vancouver Island the sequence is floored by gabbros 

(Massey, 1986). Therefore the highest velocities (~7.0 km/s) 

at depth are interpreted as a gabbro layer near the base of the 

Siletz terrane. Below, geologic interpretations are made of the 

velocity anomalies identified on slices from the three-dimensional model in the previous sections.

7.5-km Depth

High-velocity rocks at depth match fairly closely with 

the surface outcrops of Siletz rocks and with well data 

(McFarland, 1979) (figs. 3 and 10A). Thus the large body of 

high-velocity rock that occupies most of the crust onshore is 

interpreted as the Siletz terrane (fig. 10A) (specifically, the 

Crescent Formation of the Siletz terrane). The lower velocity 

anomalies west of the Siletz terrane and offshore are interpreted as accreted sedimentary rocks. At depth, high velocities are seen at varying offsets to the east of the near-surface 

contact between Siletzia and accreted sedimentary rocks (fig. 

10A). The near-surface and 7.5-km-deep contacts are nearly 

coincident in northern Oregon and southern Washington, while 

a significant eastward shift north of Willapa Bay is observed 

(fig. 10A). This appears as a high-angle contact to the south 

and a lower angle contact to the north; Snavely and Wagner 

(1982) also concluded that the Siletz-accretionary contact is 

low angle (25 ̊) north of Willapa Bay. North of Grays Harbor, 

the higher velocities of the Siletz terrane are evident wrapping around the Olympic Mountains uplift. Correspondence 

between the surface outcrops of the Siletz terrane and the 

high-velocity anomalies at depth suggests a high-angle Siletzaccretionary prism contact in the uppermost crust.

The alternating pattern of high and low velocities 

observed along the coast (fig. 10A) may be the result of 

regional north-south directed compression (McCrory, 1996) 

that causes folding, variable thinning, and imbrication of the 

Siletz terrane. Regional north-south compression may have 

domed the Siletz terrane south of Grays Harbor and in the 

Olympic Mountains along roughly east-west axes and created 

an intervening synform. The relatively low velocities associated with the southern Grays Harbor region correspond with 

Traveltime Inversion for Three-Dimensional Upper Crustal Velocity Structure of Western Washington 23
22.5

17.5

12.5

7.5

5 5.5 6 6.5 7 7.5 Serpentinite Gabbro

Amphibolite Greenschist Basalt

Diabase

Quartz Schist Zeolite Basalt

Phyllite Slate

Quartzite

Metagraywacke

Basalt

VELOCITY, IN METERS PER SECOND

DEPTH, IN KILOMETERS

the Chehalis Basin and imply that the basin is deep and has 

only a thin veneer of Crescent Formation near the surface; 

north of Grays Harbor the low velocities may indicate that the 

Crescent Formation is downwarped. Relatively low velocities 

also coincide with the accreted sedimentary core of the anticlinal Olympic Mountains uplift (Tabor and Cady, 1978).

The eastern edge of the Siletz terrane is more poorly constrained by the model; high-velocity rocks appear to be absent 

beneath the 1991 refraction profile at 7.5-km depth where 5.0 

km/s velocities are observed. Miller and others (1997) analyzed 

the complete 1991 line and show velocities consistent with Siletz 

terrane at greater depth. Moran (1997) conducted a three-dimensional inversion using local earthquake sources centered east of 

Puget Sound and found a prominent north-south trending velocity boundary 50 km east of Puget Sound beginning at ~10km 

depth that was interpreted as the east edge of the Siletz terrane.

12.5-km Depth

The boundary between accreted sedimentary rocks and the 

Siletz terrane makes a northeast trending boundary that occurs 

somewhat farther inland than it did at 7.5-km depth (fig. 11A). 

Missing at this depth are the alternating high- and low-velocity patterns evident at 7.5-km depth. In the southern Olympic 

Mountains and at Willapa Bay, significant northeast shifts in the 

terrane boundary with increasing depth can be seen by comparing its location at the surface and at depth (figs. 11A and 

15). These shifts may be evidence for a thin west edge of the 

Siletz terrane or for shallow crustal faults that offset the surface 

outcrops of Siletz terrane from their deeper roots in the depth 

range between 7.5 and 12.5 km. In an inversion centered in the 

Puget Sound region, Lees and Crosson (1990) measured an 18 ̊ 

to 28 ̊ dip on this contact near the eastern Olympic Mountains, 

comparable to that imaged by Clowes and others (1987) and 

Calvert (1996) at southern Vancouver Island. An interpretation 

of magnetotelluric data shows that the accreted sedimentary 

rocks do not underthrust the Siletz block at depths greater than 

~10 km beneath Puget Sound (Aprea and others, 1998). 

The low-velocity anomaly at Grays Harbor persists to 

12.5-km depth (fig. 11A); the uniform velocities from the 

accretionary prism to the onshore parts of this area may imply 

that accretionary rocks were pushed onshore in similar fashion 

Figure 14. Velocity-depth curves from global averages of hundreds of measured samples of the major 

classes of metamorphic, and igneous rocks (Christensen and Mooney, 1995) of likely composition and grade 

to be found in the upper 25 km of the Cascadia subduction zone (Best, 1982). Given the 2 km/s lateral velocity 

resolution of the three-dimensional velocity model, the Siletz-accretionary complex boundary is contained in 

the velocity interval between 6.0 and 6.5 km/s.

24 Crustal Structure of the Cascadia Fore Arc of Washington
as to the north beneath the Olympic Mountains. To the east, 

high velocities drop off dramatically along a north-trending boundary that corresponds roughly to the west Rainier 

seismic zone, consistent with previous interpretations that the 

deformation is occurring along the Siletzia boundary based on 

gravity data (Finn, 1990; Stanley and others, 1996).

17.5-km Depth

The primary change between 12.5-km and 17.5-km depth 

is the observation of high-velocity rocks in the offshore region. 

These higher velocities are interpreted as oceanic crust of the 

downgoing Juan de Fuca slab. Velocities appropriate for accretionary rocks make a nearly continuous linear boundary along 

the western edge of the Siletz terrane. A possible explanation 

for this may be that the compressional deformation (folding or 

faulting) of Siletzia is more pronounced at shallower depths. 

The north-trending eastern boundary also persists at 17.5-km 

depth, consistent with a high-angle boundary.

22.5-km Depth

At this depth the northeast-trending lateral velocity 

contrast as observed from 2.5-km to 17.5-km depth is not 

resolved. This change may indicate that the base of the Siletz 

terrane occurs near this depth (to within ± 5 km). The modeled 18-28-km thickness is in reasonable agreement with the 

estimates made from two-dimensional profiles (Tréhu and 

others, 1994; Parsons and others, 1998). This depth coincides 

with a horizontal band of seismicity beneath Puget Sound 

(Stanley and others, 1996) and might imply that there is some 

deformation along the base of the Siletz terrane.

Discussion of the Three-Dimensional 

Seismic-Velocity Model of Coastal 

Washington

The large-scale boundaries between the Siletz terrane and 

the adjacent accretionary prism and volcanic arc in western 

Washington were imaged in three dimensions. Internal deformation in Siletzia that manifests as alternating thinning and 

thickening along the Washington coast (fig. 10A) may result 

from folding and imbrication of the Siletz block on low-angle 

faults in the shallow crust (fig. 15). The thick Siletz terrane of 

Oregon (Tréhu and others, 1994) persists near the coast as far 

north as southern Willapa Bay. North of Willapa Bay, at the 

latitude of Grays Harbor, the west edge of the Siletz terrane 

is thinned and has a lower angle contact with the accretionary prism (fig. 15). This is consistent with deformation of the 

Crescent Formation basalt along the coast, where it is exposed 

in seaward-vergent overturned anticlines in the hanging walls 

of landward-dipping thrust faults (Snavely and Wagner, 1982; 

Wells and Coe, 1985). Seismic models from east of the Olympic Mountains and on Vancouver Island also show shallow 

easterly dips of the Siletz terrane in the upper 10 km, with 

accretionary rocks beneath it (Symons and Crosson, 1997; 

Lees and Crosson, 1990; Clowes and others, 1987; Calvert, 

1996), and Pratt and others (1997) suggest a low-angle detachment beneath Puget Sound between 14-km and 20-km depth. 

Therefore, either the western edge of Siletzia was originally 

thinner at Willapa Bay and in the Olympic Mountains and 

is thus more susceptible to deformation, or the more intense 

deformation in the northern Coast Ranges has involved Siletzia in low-angle faulting.

There is a high-amplitude antiformal uplift centered in 

the Olympic Mountains, and it appears that similar (but lower 

amplitude) deformation extends along much of the Washington margin, with uplift between Willapa Bay and Grays 

Harbor and synforms along the Columbia River and between 

Grays Harbor and the Olympic Mountains. Outside of the 

study area, north of the Olympic Mountains, this pattern persists with a synformal structure that coincides with the Strait 

of Juan de Fuca (Snavely, 1987) (fig. 16). The Bouguer gravity 

anomaly associated with the Siletz terrane also reflects this 

pattern, with a coherent gravity high along the Oregon coast 

that grades into a more complex alternating pattern of highs 

and lows in Washington (Finn and others, 1991).

The shape of the western edge of the Siletz terrane in 

Washington may have implications for the geologic evolution of the margin. In Oregon, there is clear paleomagnetic 

evidence of clockwise rotation of the Siletz terrane (Simpson and Cox, 1977). However, on the Olympic Peninsula 

to the north, the Crescent Formation shows no rotation or 

translation (Warnock and others, 1993). Thus Siletzia must 

have accommodated this variable motion by internal deformation. Northwest-directed translation of central Oregon 

and southern Washington (Pezzopane and Weldon, 1993) 

must also have been accommodated. The internal strain 

of the Siletz terrane may be focused along the coast of the 

Olympic Peninsula. The large-scale arching of the mafic 

basement along east-west trends transverse to the margin 

can be explained by margin-parallel northward transport 

and rotation of Siletzia into the relatively fixed restraining 

bend of the Canadian Coast Mountains at Vancouver Island 

(Wells and others, 1998). The Olympic Peninsula acted as a 

soft hinge-point, undergoing significant north-south directed 

deformation. Alternatively, the size of the accretionary complex reflects proximity to the major proto-Columbia River 

sediment source (Brandon and Vance, 1992) The deformation of Siletzia in Washington then represents the impact of 

major sediment accretion and possible tectonic erosion of 

the Crescent Formation along the thrust that locally forms its 

base in the Olympic Mountains (Tabor and Cady, 1978)

The Siletz terrane acts as a backstop against which 

accreted rocks are thrust, and this contact may have seismogenic potential. A three-dimensional map of the contact zone 

at depth is provided that shows a fairly steep angle in northern 

Oregon and southernmost Washington. To the north, it appears 

that the Siletz backstop deforms along with the rocks accreted 

Discussion of the Three-Dimensional Seismic-Velocity Model of Coastal Washington 25
26 Crustal Structure of the Cascadia Fore Arc of Washington

2 km

440m

Siletz Terrane

45 ̊

30 ̊

20 ̊

10 ̊ 25 ̊

Siletz Terrane Juan de Fuca P.

Juan de Fuca P.

Juan de Fuca P.

45 ̊

30 ̊

20 ̊

10 ̊

?

A

B

O l y m p i c M t s

a c c r e t i o n a r y

c o m p l e x

Siletz

Terrane

C o l u m b i a

P l a t e a u

Quaternary

Volcanics

N o r t h C a s c a d e

p r e - C e n o z o i c

t e r r a n e s

C o a s t R a n g e

Te r r a n e

P u g e t

L o w l a n d

Te r t i a r y 

C a s c a d e 

A r c

O l y m p i c M t s

a c c r e t i o n a r y

c o m p l e x

Siletz

Terrane

C o l u m b i a

P l a t e a u

Quaternary

Volcanics

N o r t h C a s c a d e

p r e - C e n o z o i c

t e r r a n e s

C o a s t R a n g e

Te r r a n e

P u g e t

L o w l a n d

Te r t i a r y 

C a s c a d e 

A r c

Velocity (km/s) Velocity (km/s)

Figure 15. Spaced vertical profiles through the three-dimensional velocity model for coastal Washington, 

shown with the topography and generalized geology from figure 3 draped above. Vertical and horizontal 

scales are distance, in kilometers. A, Cross section showing the western Siletz-accretionary terrane boundary 

at the coast of northern Oregon; basalt is present in the well indicated offshore, and thus a high-angle terrane 

contact is interpreted. The eastern edge of Siletzia is poorly constrained in this cross section. B, Basalt is 

observed in the well marked at the coast in this cross section; however, the main high-velocity anomaly is 

shifted to the east as compared with the northern Oregon section (A). A 25° dipping boundary is shown, following Snavely and Wagner (1982), who derived it from well data, shallow seismic-reflection data, and 
Discussion of the Three-Dimensional Seismic-Velocity Model of Coastal Washington 27

Siletz Terrane

45 ̊

30 ̊

20 ̊

10 ̊

?

?

Siletz Terrane

45 ̊

30 ̊

20 ̊

10 ̊

~25 ̊?

Juan de Fuca P.

Juan de Fuca P.

Juan de Fuca P.

Juan de Fuca P.

C

D

O l y m p i c M t s

a c c r e t i o n a r y

c o m p l e x

Siletz

Terrane

C o l u m b i a

P l a t e a u

Quaternary

Volcanics

N o r t h C a s c a d e

p r e - C e n o z o i c

t e r r a n e s

C o a s t R a n g e

Te r r a n e

P u g e t

L o w l a n d

Te r t i a r y 

C a s c a d e 

A r c

O l y m p i c M t s

a c c r e t i o n a r y

c o m p l e x

Siletz

Terrane

C o l u m b i a

P l a t e a u

Quaternary

Volcanics

N o r t h C a s c a d e

p r e - C e n o z o i c

t e r r a n e s

C o a s t R a n g e

Te r r a n e

P u g e t

L o w l a n d

Te r t i a r y 

C a s c a d e 

A r c

Velocity (km/s) Velocity (km/s)

magnetic modeling. Farther inland, a well finds basalt at 2-km depth west of the Doty Fault, which 

thrusts Siletzia rocks back to the surface (Snavely and Wagner, 1982; Wells and Coe, 1985). A highangle eastern boundary between Siletzia and the Cascade arc is interpreted, coincident with the 

west Rainier seismic zone (Stanley and others, 1996; Moran, 1997). C, The Siletz terrane can be seen 

beneath the southern rim of the Olympic Mountains; a 25 ° dipping boundary between Siletzia and the 

underthrust accretionary complex is approximated. D, Nearer to the axis of the Olympic Mountains 

uplift, the Siletz terrane is bent sharply upwards west of Puget Sound, where most of its thickness is 

exposed at the surface (Tabor and Cady, 1987).
28 Crustal Structure of the Cascadia Fore Arc of Washington

to it and may have active internal low-angle faults. North of 

Willapa Bay and in the western Olympic Mountains, the seaward edge of Siletzia appears to be a gently landward-dipping 

thrust flake overlying imbricated and underthrust sediments of 

the Cascadia accretionary prism. 

Three-Dimensional Crustal Structure 

of the Puget Lowland from Seismic 

Traveltimes and Bouguer Gravity 

The Puget Lowland of Washington is seismically active 

and is crossed by many faults that offset Quaternary deposits, 

or that have a history of late Holocene rupture (Gower and 

others, 1985; Atwater and Moore, 1992; Bucknam and others, 

1992; Johnson and others, 1994, 1996, 1999). Major crustal 

fault zones bound the Tacoma, Seattle, and Everett Basins (fig. 

6), large geological features that may prolong and amplify the 

strong ground motions. The Tacoma and Seattle Basins introduce a significant difference in the delay times of first arrivals 

produced along shot lines in Puget Sound versus Hood Canal. 

In Puget Sound, traveltime delays produced by the Seattle 

Basin exceed 1 s (Brocher and others, 2001).

Method

In this study, observed gravity is incorporated as a 

sequential step in the inversion process. Thus the method is 

not a joint inversion in a rigorous sense, where residuals from 

traveltime and gravity are minimized simultaneously through a 

linear expression that relates the two. Instead, the gravitational 

field predicted by a velocity model obtained from traveltime 

inversion is compared with observed gravity. Velocities are 

modified to reduce residuals, and then used as the starting 

model for the next traveltime iteration. A fixed velocity-density relation enables all output models to be expressed as both 

velocity and density; the technique could be readily modified 

to incorporate spatially variable velocity-density relations if 

desired.

As before, the three-dimensional tomographic technique 

of Hole (1992) was used for seismic traveltime inversion. The 

updated velocity model from traveltime inversions was passed 

to the gravity step as a starting model. The gravity algorithm 

was applied sequentially in three dimensions, but it can easily 

be adapted for use with any gridded or node-based traveltime 

modeling method in two dimensions or three.

For the example presented here, the gravity anomaly 

was calculated on the first iteration of the velocity model by 

converting velocity to density using Gardner’s nonlinear rule 

of � ����� ��� � for velocities (in km/s) below 6 km/s (Gardner and others, 1974) and q = 2920 kg/m3

 for velocities greater 

than 6 km/s, which correspond to the Crescent formation 

(Brocher and others, 2001) (fig. 17). These choices may be 

modified for other regions as necessary. The gravity anomaly 

g(x,y) of a three-dimensional density distribution r(x,y,z) is 

the sum of the effecdividual layer of the model. Let r(x,y,z) be 

discretized by layers, so that r(x,y,z) = rk(x,y), direct the z axis 

downward, and denote the two-dimensional Fourier transform 

by 

where kx and ky are wavenumbers for the x and y directions, 

respectively. The gravity anomaly due to the entire density 

distribution is given by

where

The functionF ���is the gravitational earth filter (Blakely, 

1995) given by 

where g is the gravitational constant, z1 and z2 are depths to 

the top and bottom of layer k, respectively, and � � ��

�

� � �

� . Thus, the gravity anomaly due to each layer was calculated by 

(1) Fourier transforming the two-dimensional density of each 

layer, (2) multiplying by the appropriate earth filter, and (3) 

inverse Fourier transforming the product. Layers with uniform 

density (inferred from uniform velocity) were ignored, as they 

would add only a constant to the overall anomaly.

The calculated anomaly is subtracted from observed gravity, which is gridded at the same scale as the top surface of the 

three-dimensional velocity model (squares 1 km by 1 km in 

this study), yielding a set of residual gravity values. The velocity model, represented by 1-km cubed cells of constant velocity and density, is then changed to minimize the residuals. A 

challenge is posed in fitting gravity observations because the 

observed anomaly at the surface is relatively independent of 

the depth of the source, yielding many possible models that 

satisfy the observations. Some assumptions were made (as 

described below), and constraints from seismic traveltime 

modeling were applied to make targeted changes in the velocity-depth relation in each column in the velocity model. These 

changes resulted in a model that is consistent with observed 

gravity and seismic traveltimes. 

Initially a grid of residual scaling factors was generated, 

each a function of the magnitude and sign of the corresponding residual gravity value, namely

F � � � � � ���� ������ ����� ����� � ���� ��

�

� ��

�

� ������� (1)

���� �� � � � ��� ���

���

�

� � � � ���

F �� � � �F �� � � � � � �∙F ���� ���

F ��� � ���

� ��� ��� � �� � ��� ��� � � � ��, z1>0, z2>z1 , (4)

���� �� � ����� �� � ��

�

��� �
Three-Dimensional Crustal Structure of the Puget Lowland from Seismic Traveltimes and Bouguer Gravity 29

N68 ̊E

North-Directed Compression

Alternating

SynformalAntiformal

Deformation

A "fixed" block

Coast

Northern Oregon Cross Section

Siletz Terrane

B

Coast

Olympic Mountains Cross Section

 Siletz Terrane

More Deformed

Less Deformed

Oregon

Washington

Figure 16. Tectonic model of Siletz-terrane deformation. A, North-directed compressional stress and southern Siletz terrane rotation 

are accommodated at a hinge point along the coast of the Olympic Peninsula. The greatest amplitude of uplift occurs at the Olympic 

Mountains, but the pattern of folding persists north and south of the Olympics. This may occur as a result of the margin-parallel northward transport and rotation of Siletzia into the relatively fixed restraining bend of the Canadian Coast Mountains at Vancouver Island. B, 

Comparative cross sections through northern Oregon and the Olympic Mountains indicate the relationship between backstop deformation and sedimentary accretion, although whether this relationship is causative remains an open question.
30 Crustal Structure of the Cascadia Fore Arc of Washington

where r(x,y) is the residual at a grid point, �

� is the mean of all 

gravity residuals, and  is an input constant used to scale the 

magnitude of velocity change. Larger values of  cause smaller 

incremental changes in the velocity model with each iteration, 

allowing the effects of traveltime and gravity steps to gradually evolve a model consistent with both data sets. Appropriate  values were determined by experimenting with different 

numbers, using lowered RMS misfit in traveltime and gravity 

as the criterion. 

The second step of the algorithm is calculation of the 

velocity-depth gradient (dV/dz) at each model cell. The 

gradients are found by finite differencing the gridded velocity 

model. Gradient values are used to scale the velocity perturbations by maximizing changes where the gradients are highest, based on the assumption that much of the variation in the 

gravity anomaly results from sedimentary basin structures 

near the surface. Slight increases or decreases in the model 

depth to basement can quickly satisfy the residuals. In many 

cases, crustal tomography studies are undertaken to resolve 

basin structures (models for assessing likely earthquake strong 

ground motion for example). Preferentially changing velocity 

where gradients are highest has an effect akin to subtly moving the depth to the basement-sediment interface. Subsequent 

traveltime iterations allow such changes only within the allowable range that still satisfy the traveltime data.

The last scale applied on velocity perturbation resulting 

from gravity residuals is a seismic coverage factor. Perturbations were scaled relative to the degree of calculated seismic 

ray coverage in a given model cell by

where h(x,y,z) is a hit count (number of times a seismic ray 

is calculated to encounter a model cell), α is the maximum 

hit count in the model, and ω is a weighting factor (can range 

from 0 to 1; a value of 0.1 was used in this study). Use of this 

scaling factor enables the model to be preserved where seismic 

coverage is high and allows more variation where seismic coverage is low or nonexistent. Thus the gavity field can be used 

to supplement traveltime coverage. The weighting value w is 

an input parameter and allows control on how much weight 

to give the traveltime data versus the gravity data. In the test 

cases, most weight was given to seismic traveltimes because 

they are more sensitive to depth than residual gravity values, 

which are affected primarily by variations in basin structure.

Model velocity is changed by a simple product of the three 

scaling factors as

If any of the scaling factors approach zero at a given cell, 

then velocity is unchanged. Thus when the residuals grow 

very small, or when the velocity gradient is zero, the velocity perturbation approaches zero. The coverage scale C(x,y,z) 

is equal to 1 when the seismic hit count in a cell is zero (see 

equation 6), meaning that the perturbed velocity model is a 

function solely of the starting model and the gravity residuals 

in that case. Maximum and minimum allowable velocity-depth 

ranges are set that prevent unrealistic model velocities from 

developing. In this study, these ranges were set to the minima 

and maxima obtained from laboratory analysis at confining 

pressure of regional rock samples (Brocher and others, 2001). 

Lastly, a spatial smoothing filter is applied that minimizes 

vertical streaking, since perturbations are made column by 

column. 

An example of the iterative process is shown in figure 18. 

The largest changes in the velocity-depth curve occur at depths 

with steepest gradients and relatively little seismic coverage. 

The velocity-depth curve is slightly smoothed as a result of 

introducing the gravity constraint, especially where seismic 

coverage is minimal (fig. 18). This example is a representative column of the test model and traveltime database (more 

fully discussed in a later section) and demonstrates the fairly 

minimal velocity perturbations necessary to satisfy both the 

traveltimes and the gravity anomalies.

Resolution and Test Models

Following the methods used in the three-dimensional 

velocity modeling of coastal Washington, checkerboard tests 

were conducted by calculating synthetic traveltime picks 

between all the SHIPS source and receiver positions. In addition, a synthetic gravity field was calculated with a model 

of vertical columns 10 km by 10 km, each with alternating 

increasing velocity gradients that were 0.5 km/s different at 

all depths. The synthetic traveltime picks and synthetic gravity grid were then used with a 1-D starting velocity model 

to recover the checkerboard pattern by sequential inversion, 

and the result was compared with the result of using only the 

traveltime picks.

Results from resolution tests limited to seismic traveltimes show recovery of a smoothed version of the checkerboard pattern only in regions where seismic-ray coverage is 

high (figs. 19, 20) (see also Brocher and others, 2001). Significant improvement is made by sequentially incorporating the 

gravity constraints, particularly in areas with minimal seismic 

coverage. The improvement is primarily recovery of the shape 

and scale of relative velocity variation rather than recovery 

of absolute velocities because a fixed velocity-density relation was used. The estimated confidence in locating the actual 

boundaries improves with the addition of gravity; Brocher and 

others (2001) estimated that anomalies greater than 15 km 

across were resolvable using only traveltimes. Here 10-km 

anomalies are well resolved. 

The checkerboard resolution tests indicate some utility 

of the gravity inversion in extending the model outside of 

the seismic coverage because the checkerboard pattern is 

recovered in those regions (fig. 19). Spatial smoothing of the 

model extends the influence of seismically determined velocities by the half-width of the smoother. Beyond this distance 

�������� ����

��������

�

�

���� ����� ��� �

����� ���� �� ���� ���� � ��� ��� ���

��� ����

��

�������� �

��

����

�

��

���� • • �� ���
Three-Dimensional Crustal Structure of the Puget Lowland from Seismic Traveltimes and Bouguer Gravity 31

1.5

1.8

2.1

2.4

2.7

3.0

1 2 3 4 5 6 7 8

Gardner's Rule � = 1740 v0.25

DENSITY, IN GRAMS PER CUBIC CENTIMETER

Puget Lowland wells

Olympic core complex

Crescent basement

COMPRESSIONAL WAVE VELOCITY, IN KILOMETERS PER SECOND

Figure 17. Plot 

of observed velocity vs. density from 

Puget Lowland rock 

samples, compared 

with the Gardner’s 

rule (Gardner and 

others, 1974) relation 

(heavy black line) 

used in sequential 

seismic traveltime 

and gravity inversion 

for crustal structure.

3.0 4.0 5.0 6.0 7.0

0.0

5.0

10.0

15.0

20.0

1,000 3,000 5,000

Seismic coverage (hit count)

Velocity-seismic only

Velocity-fit to gravity

VELOCITY, IN KILOMETERS PER SECOND

DEPTH, IN KILOMETERS

Figure 18. Comparison of velocity-depth profiles taken from threedimensional models derived from seismic traveltimes only (black 

line) with those derived from combined analysis of traveltimes and 

gravity (gray line). Adjacent bar graph shows depth variation of 

seismic hit counts (see text for explanation) associated with the 

velocity-depth curves. Perturbation introduced by incorporating 

gravity is greatest where the velocity gradient is steepest (depths 

between 1 and 4 km) and where seismic coverage is sparse. 

the gravity inversion is highly sensitive to the starting-model 

velocity gradient and will tend to suggest sedimentary basins 

where the observed anomaly is low and basement highs corresponding with gravity highs. Without some guidance from 

the starting velocity model, features such as low-density 

granitic plutons or high-density carbonates or mafic intrusions would probably be mishandled. Low-velocity zones are 

also unlikely to be properly resolved in areas lacking seismic 

coverage unless the starting model includes reverse gradients 

with depth.

Application and Evaluation

Resolution tests using synthetic data sets based on a 

real geometry show improvements to a velocity model by 

incorporating sequential gravity iterations. In this section the 

crustal structure of the Puget Lowlands was calculated using 

traveltime data from the SHIPS experiment and the observed 

gravity anomaly for the Puget Lowland study area of Brocher and others (2001) (figs. 6, 21). The gravity observations 

were discretized onto a surface grid of cells 1 km by 1 km 

that corresponds to the 1-km cubes in the three-dimensional 

velocity model (fig. 21). The best results in terms of reducing traveltime and gravity residuals were achieved by running 

five iterations using traveltime only, and then introducing the 

gravity constraints sequentially with traveltime iterations for 

another five iterations. Because of the distribution of seismic 

sources and receiver stations, resolution of velocity anomalies ~5-10 km across in the lateral dimensions was sought. A 

5-km halfwidth smoothing filter was applied during the final 

iteration, which yielded an RMS traveltime misfit of 0.09 s, 

and a 5 mGal RMS misfit to the gravity. Introduction of the 

gravity constraint necessitated some minor station corrections 
32 Crustal Structure of the Cascadia Fore Arc of Washington

to achieve the 0.09-s RMS traveltime misfit. At some stations, 

traveltime misfits were consistently shifted either positively or 

negatively, without dependence on source azimuth and range; 

in these cases the mean traveltime residual was subtracted 

from the RMS residual. Introduction of the gravity constraint 

caused this problem because seismic coverage is very sparse 

in the upper 1-2 km of the model, and lateral smoothing of 

nearby localized velocity anomalies affected velocity immediately beneath some stations. 

Recovery of the Gravity Anomaly

In figure 21, a comparison is made of predicted gravity 

calculated from velocity models developed with and without 

sequential gravity steps. As expected, the velocity model calculated with gravity and seismic traveltimes reproduces more 

features of the observed gravity anomaly than does the traveltime-only model. However, in the Tacoma Basin the calculated 

anomaly does not match well with the observed; the anomaly 

associated with the Tacoma Basin is reproduced only slightly 

better by the combined model. It appears that density is lower 

for a given velocity in the Tacoma Basin as compared with the 

Seattle Basin, where the calculated anomaly agrees well with 

the observed (fig. 21). This circumstance indicates that, in some 

cases, a spatially varying velocity-density relationship might be 

desired to best resolve structure. 

Comparison of Velocity Models with Known 

Geologic Features

The velocity model that resulted from sequential inversion of traveltimes and gravity is shown at 5-km depth in figure 

22, where a comparison is made between a model generated 

from the traveltime data only and one that was fit to traveltimes 

and the gravity anomaly. Evaluation of results from applying 

the sequential gravity steps to the SHIPS model is subjective, 

because the actual subsurface structure is unknown. There are, 

however, some features in the combined model that appear better correlated with surface geology and inferred, related deeper 

structure. For example, the accreted sedimentary core of the 

anticlinal Olympic Mountains uplift is expected to have lower 

seismic velocity than their backstop, the Eocene volcanic rocks 

of the Crescent Formation that form a rim around the Olympic 

Mountains (Tabor and Cady, 1978; Parsons and others, 1999). 

The rim of high-velocity Crescent rocks is well resolved by 

the traveltime model, but there is only a limited image of lower 

velocity Olympic-core rocks inside the rim because of poor 

seismic coverage (figs. 20,22). With the incorporation of gravity, lower velocity core rocks are more evident, and the rim of 

Crescent Formation rocks is resolved farther west than before. 

Another example is the southern Whidbey Island Fault, 

discussed by Johnson and others (1996), which is resolved 

by the traveltime model as bounding the Seattle Basin on 

Figure 19. Checkerboard resolution tests of three-dimensional velocity model at 5-km depth. Synthetic traveltime and gravity observations were calculated with the input model using the SHIPS experiment geometry. The two recovered models used traveltimes only and 

traveltimes in combination with observed gravity. Use of gravity in combination with seismic traveltimes resolves the checkerboard pattern 

much better than traveltimes alone, and over a larger area.
Three-Dimensional Crustal Structure of the Puget Lowland from Seismic Traveltimes and Bouguer Gravity 33

Figure 20. Comparison (left) between distribution of gravity observations (small yellow dots) and SHIPS seismic sources (dashed lines) and 

receivers (blue dots) in the Puget Lowland. Gravity coverage is more uniform and the number of observation points is greater than the coverage 

by wide-angle recorders in the SHIPS seismic experiment. Also shown (right) is seismic-ray coverage at 5-km depth below sea level expressed 

as a hit count, the number of times a 1-km cubic cell in the model is crossed by seismic raypaths between sources and receivers. The darkest 

areas have the highest hit counts.

the southeastern end of the fault. The combined traveltimegravity model shows additional influence of this fault on 

Whidbey Island, where a distinct high-velocity anomaly is 

bounded by the northwest extent of the southern Whidbey 

Island Fault (fig. 22) but is not evident from the traveltimeonly model. Much of this high-velocity anomaly occurs offshore, where there are many airgun sources but few receivers, 

causing unreversed seismic arrivals that limit resolution. 

A final example is the Everett Basin, which is located north 

of the Seattle Basin and was poorly resolved by the traveltime 

model because it lies at the edge of seismic coverage (Brocher 

and others, 2001; figs. 20,21). Incorporation of the gravity 

observations introduces this basin into the velocity model where 

it was previously absent (fig. 22). Even relatively small gravity 

anomalies can have important effects on the velocity model in 

the upper crust where seismic coverage is sparse.

It appears that structures around the edges of seismic 

coverage are better resolved as a result of fitting the gravity 

anomaly (fig. 22). Subtle changes are also evident within the 

region of high coverage (figs. 20, 22) where the Tacoma Basin 

and, to a lesser extent, the Seattle Basin have slightly different shapes. The resolution tests showed better recovery of the 

checkerboard pattern in the region of highest ray coverage 

(figs. 19, 20); the modifications to the Tacoma and Seattle 

Basins may thus indicate refinement of these basins in the 

model, though this is difficult to verify. 

Evaluation from Borehole Velocity Data

Six deep industry boreholes in the study area provide 

sonic velocity logs (fig. 6). Comparison of the tomography 

model to these sonic velocity logs suggests that the shallow 

(upper 3 km) velocities are generally well recovered by both 

velocity models (fig. 23). This comparison, as originally 

made by Brocher and others (2001), showed that the velocity structure in the upper 3 km of the traveltime-only model 
34 Crustal Structure of the Cascadia Fore Arc of Washington

0 50 100 150

200

150

100

50

0

0 50 100 150

5 6 7

Sesimic velocity from traveltimes only Sesimic velocity from Bouguer gravity and traveltimes

Olympic 

Mountains S. Whidbey Island Fault

Everett 

Basin

Seattle

Basin

Tacoma

Basin

Cross-sections

DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

VELOCITY, IN KILOMETERS PER SECOND

Figure 22. Horizontal (map-view) slices from the three-dimensional model volumes at 5-km 

depth below sea level are shown from the traveltime-only inversion (left) and the combined 

gravity-traveltime inversion (right). White dashed lines are crustal faults. Incorporation of gravity 

highlights features not previously evident, such as the Everett Basin and offset of a high-velocity 

body along the northwest end of the South Whidbey Island Fault. The image of low-velocity rocks 

of the Olympic Mountains ringed by higher-velocity, mafic Crescent Formation rocks is improved. 

East-west line marked “cross sections” gives the location of the vertical slices shown in figure 24. 

Figure 21. Calculation of the gravity anomaly from velocity models developed from (A) traveltimes only and (B) combined traveltime 

and gravity analysis in the Puget Lowland. As expected, more features of the observed gravity anomaly (C) are apparent from the combined analysis. This exercise highlights areas like the Tacoma Basin, where the traveltime residuals were satisfied but the full gravity 

anomaly was not, implying that relatively lower density rocks for a given velocity may fill the basin. 

0 50 100 150

0

50

100

150

200

0 50 100 150

0

50

100

150

200

-80 -60 -40 -20 0 20 40 60 80

0 50 100 150

0

50

100

150

200

Gravity calculated from traveltime-only model Gravity calculated from combined model Observed Bouguer gravity

Everett 

Basin

Seattle

Basin

Tacoma

Basin

A B C

DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS DISTANCE, IN KILOMETERS

GRAVITY, IN MILLIGALS
Dungeness Spit #1, Standard-Union 

1000

5200

0 500 1,000 1,500 2,000 2,500 3,000 3,500

Kingston #1, Mobil Oil Corporation 

1000

1600

2200

2800

3400

4000

4600

5200

0 500 1,000, 1500 2,000 2,500 3,000 3,500

Pope & Talbot #3-1, Union Oil Co. of Calif.

1000

1600

2200

2800

3400

4000

4600

5200

0 500 1,000 1,500 2,000 2,500 3,000 3,500

Socal-Schroeder #1, Standard Oil Co. of Calif.

1000

1600

2200

2800

3400

4000

4600

5200

0 500 1,000 1,500 2.000 2.500 3.000 3.500

1000

1600

2200

2800

3400

4000

4600

5200

0 500 1,000 1,500 2000 2,500 3,000 3,500

Socal Whidbey #1, Standard Oil Co. of Calif.

1000

1600

2200

2800

3400

4000

4600

5200

0 500 1,000 1,500 2,000 2,500 3,000 3,500

Pope & Talbot #18-1, Union Oil Co. of Calif.

1000

1600

2200

2800

3400

4000

4600

5200

Crescent Fm.

Pleistocene 

deposits

DEPTH, IN METERS DEPTH, IN METERS

DEPTH, IN METERS DEPTH, IN METERS

DEPTH, IN METERS DEPTH, IN METERS

SONIC VELOCITY, IN METERS PER SECOND SONIC VELOCITY, IN METERS PER SECOND SONIC VELOCITY, IN METERS PER SECOND

SONIC VELOCITY, IN METERS PER SECOND SONIC VELOCITY, IN METERS PER SECOND SONIC VELOCITY, IN METERS PER SECOND

Crescent Fm.?

Crescent Fm.

Pleistocene deposits Pleistocene deposits

Renton Fm.

Mt. Perris 

volc. rocks

Puget

Group

Pleistocene deposits

Blakeley Fm.

Blakeley Fm.

Pleistocene deposits

Blakeley Fm.

Blakeley Fm.

Scow Bay Sandstone

and unnamed strata

Pleistocene deposits

Blakeley Fm.

Blakeley Fm.

Puget Group

correlative

Scow Bay

Sandstone

Scow Bay

Sandstone

Twin 

River

Fm.? 

Velocity-depth traveltime-only

Velocity-depth gravity-traveltime

Figure 23. Comparison of tomography results with sonic-log data and lithologies for six boreholes (Brocher and Ruebel, 1998; Brocher 

and others, 2001). Borehole locations are shown in figure 3. The velocity profile from the model calculated from traveltimes only is 

shown with a heavy dashed line, and the velocity profile from combined traveltime and gravity inversion is shown with a solid line.

Three-Dimensional Crustal Structure of the Puget Lowland from Seismic Traveltimes and Bouguer Gravity 35
36 Crustal Structure of the Cascadia Fore Arc of Washington

6.0

5.0

4.0

7.0

0 50 100 150

4.0

5.0

6.0

7.0

0

10

20

30

0

10

20

30

0

10

20

30

0 50 100 150

0 50 100 150

Velocity model from traveltimes only

Velocity model from Bouguer gravity and traveltimes

Seismic coverage

2.0 5.0 8.0

0 250 500 750 1,000

Hit count

Olympic Mountains Crescent Fm. Seattle Basin

W E

DEPTH, IN KILOMETERS DEPTH, IN KILOMETERS DEPTH, IN KILOMETERS

DISTANCE, IN KILOMETERS

DISTANCE, IN KILOMETERS

VELOCITY, IN KILOMETERS 

PER SECOND

DISTANCE, IN KILOMETERS

Figure 24. Comparative east-west cross sections from the three-dimensional velocity models derived from 

combined gravity and traveltime analysis and from traveltime modeling only (see fig. 6 for location). The region 

constrained by seismic raypaths is shown by the hit-count map below. The velocity structure is subtly changed 

by introduction of the gravity constraint: The low-velocity rocks of the Olympic Mountains are shifted deeper, and 

the eastern edge of the Seattle Basin is more completely resolved. Slight variation in velocity structure beneath 

the depth of seismic coverage (~15 km) results mostly from smoothing conducted during the gravity inversion.
References 37

matched the data from four of the six boreholes to within 

0.2 to 0.3 km/s, particularly for the sedimentary rocks in the 

Oligocene Blakeley Formation filling the Seattle Basin. Thin 

(<300 m thick) layers of high-velocity Crescent Formation 

in the Socal Whidbey #1 and Pope and Talbot #18-1 boreholes were too thin to be resolved by the tomography method 

(fig. 23). The greatest misfit occured at the Pope and Talbot 

#3-1 and Dungeness Spit #1 wells at the northern end of the 

tomography model in regions having sparse receivers (figs. 6, 

23). Incorporation of gravity improved the fit to wells where 

seismic coverage was sparse, but degraded the fit slightly 

where seismic coverage was best, such as at the Mobil Kingston #1 well. This occurrence illustrates the tradeoffs in seismic 

resolution that can result from the gravity constraint. 

Comparison of Velocity Models in Cross-Section 

View

 A cross section (location shown in fig. 22) from the 

velocity model shows how structure was modified by the 

incorporation of gravity in the inversion. The differences are 

relatively minor, but potentially important (fig. 24). A deeper 

zone of low velocity beneath the Olympic Mountains was 

modeled, with a ~3-km downward shift of isovelocity contours resulting from the observed gravity low in the Olympic 

Mountains (fig. 21). A higher velocity at the surface is modeled where Crescent Formation volcanic rocks are exposed 

and where seismic coverage is absent (fig. 24). The modeled 

shape of the Seattle Basin is changed with the incorporation of 

gravity; the basin appears more symmetric because the eastern 

edge is modified from the starting model and is better resolved 

where seismic coverage fades out. By design, most of the 

variation occurs where seismic coverage is low or nonexistent. 

Resolution of basins near their edges may enable more accurate simulation of earthquake strong ground motion. 

Conclusions from Seismic and Gravity 

Modeling of the Puget Lowland Crustal 

Structure

Improvement to synthetic and real three-dimensional 

seismic-velocity models was demonstrated by incorporating 

gravity modeling as a sequential step in traveltime inversion. 

The technique is simple and can be easily added to any iterative, grid-based traveltime modeling scheme. Inversions of 

seismic traveltimes and gravity observations have resolving 

power in different parts of the crust. Additionally, there are 

usually more gravity stations than seismic recorders in a given 

region, and the two data sets often have different spatial distribution. Thus, adding a gravity constraint as a sequential step 

in the seismic-velocity inversion can significantly improve 

resolution of shallow crustal structure. 

Acknowledgments

Parts of this work were supported by the German Ministry of Education, Research, Science, and Technology, the U.S. 

National Earthquake Hazards Reduction Program, the USGS 

Continental Surveys Program, and the USGS Coastal and 

Marine Geology Program. Bob Crosson, Ruth Ludwin, and 

Steve Malone at the University of Washington are acknowledged for helping with arrival-time data from the Pacific 

Northwest Seismic Network. Reftek seismographs were 

provided by the IRIS-PASSCAL instrument center at Stanford 

University; Marcos Alvarez and Anthony Wei supported the 

controlled-source experiments. Guy Cochrane, Michael Ellis, 

Gary Ernst, Michael Fisher, Harvey Kelsey, Simon Klemperer, 

Vicki Langenheim, Kate Miller, Walter Mooney, and Anne 

Tréhu reviewed various parts of this compilation.

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40 Crustal Structure of the Cascadia Fore Arc of Washingtons