Gravity and Structures of the Crust and Subcrust in the Northeast Pacific Ocean West of Washington and British Columbia

Home , Explorer Ridge, Jasper Seamount, Lower mantle (Earth), Queen Charlotte Fault

AN ABSTRACT OF THE THESIS OF

RICHARD WILLIAM COUCH for the DOCTOR OF PHILOSOPHY (Name) (Degree)

in GEOPHYSICS presented on J I9 (Major) DateY Title: GRAVITY AND STRUCTURES OF THE CRUST AND SUB-

CRUST IN THE NORTHEAST PACIFIC OCEAN WEST OF

WASHINGTON AN RITISH COLUMBIA Redacted for Privacy Abstract approved:

Gravity measurements made with LaCoste-Romberg surface ship gravity meter S-9, along 20, 000 ku-i of track line, off Washing-. ton an& British Columbia indicate the region is approximately in isostatic equilibrium.Free-air anomalies over the region average near zero except along the base of the continental slope.A -80 to -100 mgalfree-air anomaly occurs west of the Queen Charlotte Islands over the partially sediment-filled Queen Charlotte trough. A free-air anomaly greater than -l50mgal occurs over the Scott Is lands fracture zone off the north end of Vancouver Island.The fracture zone, 170 to 240 km long, exhibits minor bathymetric relief and is interpreted as a graben-like structure filled with approximately 6 km of sediments.Five crustal and subcrustal cross sections consistent with seismic refraction and gravity data indicate that the Mohorovicic (Moho) discontinuity is at depths of approximately 25 to 28 km in the Insular Belt of British Columbia and 19 km in western Washington.In the areas of the Cascadia Abyssal Plain and continental rise off British Columbia, the com-. bination of near-zero gravity anomalies and thin crust indicates that mantle densities are less than under the adjacent continent and Tufts and Alaskan abyssal plains.Still lower densities occur in the upper mantle immediately beneath the Juan de Fuca and Explorer ridges where Moho depths are less than 7 km. Gravity and Structures of the Crust and Subcrust in the Northeast Pacific Ocean West of Washington and British Columbia

by Richard William Couch

A TNESIS submitted to Oregon State University

in partial fulfillment of the requirements for the degree of Doctor of Philosophy June 1969 APPROVED:

Redacted for Privacy Professor of Oceanogphy in charge of major

Redacted for Privacy

Chairm\an of the Departr4ient of Oceanography

Redacted for Privacy

Dean of Graduate School

Date thesis is presented

Typed by Gwendolyn Hansen for Richard William Couch AC KNOW LEDGEMENTS

This research was performed under the direction of Dr. Peter Dehlinger, Professor of Geophysics at Oregon State University.I am grateful to Dr. Dehlinger for providing theopportunity to do this research, for his many suggestions in the analysis of the data and for editing the manuscript. I am grateful to Messrs. E. Robey Banks and Michael Gemperle for their assistance in making gravity measurements at sea and in applying computer operations to gravity problems.Their willingness and ability to help were a constant encouragement. Appreciation is also extended to Dr. John N. Gallagher for assistance in making gravity measurements at sea and for many helpful discussions about the research. Professor Gunnar Bodvarsson and Professor Tjeerd H. van Andel reviewed the thesis and made many helpfulcomments. This thesis reflects the patience and encouragement of Carol, my wife. During my tenure as a graduate student, I was the recipient of a most appreciated National Aeronautics andSpace Administration fellowship. This research was supported by the Office of Naval Research under contracts Nonr 1286(09) and (10).For this assistance I am very grateful. TABLE OF CONTENTS

Page

INTRODUCTION 1

HISTORICAL DEVELOPMENT 3

THEORETICAL CONSIDERATIONS 6

PREVIOUS WORK 9

Physiography 9 Geology 12 Geophysics 14 Seismic Refraction and Reflection Measurements 14 Earthquake Observations 19 Heat Flow 25 Magnetics 26 Gravity 31

NEW GRAVITY DATA 36

Surface-Ship Gravity Measurements 36 Reduction of Gravity Data 41 Distribution of Gravity Measurements 46 The Free-Air Gravity Anomaly Maps 48 INTERPRETATION OF THE FREE-AIR GRAVITY ANOMALY MAPS 51

Gravity Anomaly Trends 51 Isostatic Equilibrium in the Northeast Pacific Ocean 52 Fjords of the British Columbia Coast 54 Gravity Anomaly Trends in Dixon Entrance, Hecate Strait and Queen Charlotte Sound 55 The Continental Shelf Off Washington and Vancouver Island 58 The Randsenken Effect Along the Continental Margin of Washington and British Columbia 60 The Queen Charlotte Trough 63 A Postulated East-West Fracture Zone Near 51° N Latitude 65 Page The Scott Islands Fracture Zone 67 Seamounts 69 The Explorer Trough 72 The Cascadia Abyssal Plain 82

CRUSTAL AND SUBCRUSTAL CROSS SECTIONS ACROSS THE CONTINENTAL MARGINS OF WASHINGTON AND BRITISH COLUMBIA 84 Construction of the Crustal and Subcrustal Cross Sections 84 Sources of Error in the Cross Sections 88 The Central Washington Cross Section 90 The South Vancouver Island Cross Section 96 The Scott Islands Cross Section 102 The Moresby Island Cross Section 109 The Dixon Entrance Cross Section 115 The Continental Margin Structures off Washington and British Columbia 119

CONCLUSIONS 122

REGIONAL TECTONICS 124

BIBLIOGRAPHY 133

APPENDICES 148

Appendix 1.The Reduction of Gravity Measurements 148 Appendix 2.Supplementary Data 161 Appendix 3.Two Dimensional Gravity Computations 171 LIST OF FIGURES

Figure Page 1.Physiographic diagram of the northeast Pacific Ocean off Washington and British Columbia. 11

2.Location map of seismic refraction lines and heat flow stations off the coasts of Washington and British Columbia. 15

3.Earthquake epicenters in the Pacific Northwest. 21

4.Total magnetic anomaly map of the region west of Washington and British Columbia. 28

5.Map of gravity meter track lines off the coasts of Washington and British Columbia. 33

6.Free-air gravity anomaly map of the region off the west coast of Washington and British Columbia from the Columbia River to Cape Scott. 42

7.Free-air gravity anomaly map of the region off the west coast of British Columbia from Cape Cook to Dixon Entrance. 43

8.Bathymetric chart of the Explorer trough and Scott Islands fracture zone. 74

9.Bathymetric, gravity and magnetic profile across Explorer trough along line X-X'. 76

10.Bathymetric, gravity and magnetic profile across Explorer trough along line Y-Y'. 77

11.Bathymetric, gravity and magnetic profile across Explorer trough along line Z-Z'. 78

12.Location map of crustal and subcrustal cross sections. 85

13.Graph of the Nafe-Drake relation between compressional wave velocity and rock density. 87 Figure Page 14.Central Washington crustal and subcrustal cross section BB'. 91

15.South Vancouver Island crustal and subcrustal cross section CC'. 97

16.Scott Islands crustal and subcrustal cross section DD'. 101

17.Free-air gravity anomaly map of the Scott Islands fracture zone. 106

18.Total magnetic intensity map of the Scott Islands fracture zone. 10.7

19.Bathymetric, gravity and magnetic profile across the Scott Islandsfracture zone along line SS'. 108

20.Moresby Island crustal and subcrustal cross section EE'. 110

21.Total magnetic intensity map off Moresby Island. 113

22.Dixon Entrance crustal and subcrustal cross sectithI FF'. 116

23.Diagram illustrating equation parameters for computing gravity over a cone structure. 165

24.Polygon illustrating equation parameters for two- dimensional gravity computations. 171 LIST OF TABLES

Table Page 1.Information on the gravity anomaly size and gravity station density in different physiographic provinces 47

2.Seamount load and compensation parameters .71

3.Seamount locations, dimensions and gravity 170 GRAVITY AND STRUCTURES OF THE CRUST AND SUBCRUST IN THE NORTHEAST PACIFIC OCEAN WEST OF WASHINGTON AND BRITISH COLUMBIA

INTRODUCTION

The northeast Pacific Ocean is an area of great interest to geologists and geophysicists because of its diverse structures, particularly as these structures have global geotectonic implications. Descriptions and studi.es of various parts of this region have been made over a period of years. Bathymetric recordings showed a smooth floor, termed by Hurley (1960) the Cascadia Abyssal Plain, immediately west of the generally narrow continental shelf and steep continental slope off Washington and British Columbia. Between Tufts Abyssal Plain, an area of smooth sea floor approximately 300 miles west of Washington and British Columbia, and the Cascadia Abyssal Plain lies a region of abyssal hills and mountains.This region, first termed the Ridge and Trough Province by Menard (1955b),is now considered a segment of the world-encircling mid-ocean ridge and rise system (Menard, 1960, 1964).In this same region Raff and Mason (1961) made magnetic measurements which showed the now classical long magnetic anomaly lineatilons.Vine and Mathews (1963) suggested that the magnetic anomalies, which are now found paralleling the mid- ocean ridge system in all oceans, represented a magnetic recording of sea floor spreading, a still controversial hypothesis first postulated by Hess (1962) and Dietz (1961).Only recently have sufficient data become available to permit a study of the structures and dynamics of this region. The surface of the solid earth west of Washington and British Columbia is covered by approximately 3. 5 km of water; consequently, direct observations of the physical and structural form of the earth in this region are limited.In this environment geophysical measurements provide valuable indirect observations which contribute greatly to the understanding of the earth beneath the sea. This study is an effort to provide a description of the earth's gravity field in the oceanic region contiguous with Washington and British Columbia and to relate these measurements to possible structures of the earth' s crust and subcrust.The initial part of this study is concerned with the construction of free-air gravity maps of the region which help to delineate and delimit the major structures of the region. The final portion of this study combines the gravity information with other geophysical data, such as seismic refraction measurements, to form hypothetical crust and subcrustal cross sections of the major geologic structures, notably the Juan de Fuca Ridge- -a portion of the mid-ocean rise system- -and the continental margin region where the relatively thin oceanic crust adjoins the much thicker continental crust. 3

HISTORICAL DEVELOPMENT

In the early part of the twentieth century measurements of the earth's gravity field were made with a torsion balance in the search for oil and heavy minerals. The torsion balance, which actually measures the horizontal gravity gradient, was supplementedin the early 1930's by the pendulum, which was usually used to measure relative values of the acceleration of gravity..The gravinieter which measures small differences in gravity has superseded boththe torsion balance and pendulum for obtaining gravity measurements, particularly over limited areas. The early use of gravity measurements in studies of the gross subsurface structures of the earth was concerned with the isostatic balance of the earth's outer layers.Hayford and Bowie (1912) developed a method of determining isostatic gravity anomalies which showed that, regionally, the continental U. S. is in isostatic equilibrium. Measurements of gravity at sea began in 1923 with a three-pendulum apparatus devised by Vening Meinesz (1929) which operated in a submerged submarine. Measurements by Vening Meinesz (1934) combined with extensive land measurements, indicated that at least regionally the earth's outer layers are in isostatic equilibrium.In addition, Vening Meinesz (1934) measured some of the largest negative anomalies known; these occur along the arcuate 4 oceanic trenches.These large negative anomalies, indicating extensive mass deficiencies, were used to support the tectogene theory of the down-warping of the earth's crust due to forces deep within the earth. Additional measurements of gravity at sea were made with pendulums by the Lamont Geological Observatory of Columbia Univer-. sity (Worzel,1965)over oceanic ridges, along trenches and across the transition zone between continental and oceanic crusts.These data were combined with geological and other geophysical data, particularly seismic refraction results, which established the depths of the various layers of hypothetical crustal sections in the areas of study.Talwani etal.(1959b)adapted computer techniques to the calculation of crustal sections.Densities can be obtained from crustal and subcrustal velocities determined by seismic refraction studies by the use of empirical relations such as those proposed by Nafe and Drake (1961) based on data from a large number of core samples.Illustrative of this work is the cross section of the Island of Puerto Rico and the adjacent trench (Talwani etal. ,1959a),which shows large variations in the depth to the Mohorovicic (Moho) discontinuity.

These techniques were applied by Dehlinger etal.(l967a)to the Mendocino Escarpment west of northern California and indicated that in this region the material beneath the crust has, in addition to 5 variations in depth to the Moho, horizontal variations in density. Dehlinger eta]. (1967b) combined their gravity data with seismic measurements reported by Shor eta]. (1968) to analyze structures across Juan de Fuca Ridge and thetransition zone between the continental crust of Oregon and adjacent oceanic crust.This work indicates a; relatively low mantle density beneath Juan de Fuca Ridge and a relatively thin continental crust under western Oregon.The techniques developed in this and other similar research are applied in this study to the interesting and controversial area west of Washington and British Columbia. THEORETICAL CONSIDERATIONS

A force of mutual attraction exists between all bodies. According to Newton's law, the force of attraction between two bodies is directly proportional to their masses and inversely proportional to the square of the distance between them.In equation form m1m2 F = whereFis the force of mutual attraction directed r2 along a line joining the two point massesm1 andm2which are separated by the distancer.. The proportionality constant y is termed the universal gravitational constant and equals 6. 673 Xio8

(cgunit). A force defined at every point in space describes a force field. When a force field is continuous in an open region, and the integral of the force along a path joining two points in the field is independent of the path between the points, there exists a potential function U such thatFvU (e.g. Apostol, 1957).Because work is independent of the path,such a force field is termed conservative and indicates that the force is equal to the gradient of the potential U. The earth's potential is usually assumed positive in geophysical studies. Consequently, the difference in potential U1 between two points is defined as the negative of the difference in potential energy W1 between the two points. 7 The Newtonian gravitational field is a conservative force field and, therefore, all field parameters may be defined in terms of U

y the gravitational potential of the point mass where U r m1 at the location of measurement ofU. Thus, if mass is distributed with density p (x, y, z)throughout a volumeV, the gravitational potential at an exterior pointP,'x,y,z,is given p(x,y,z) U dv. by the equation x ,y ,z = 'Vr(x,y,z) 0 0 0 S v The fundamental problem in the determination of the density distribution in the earth's crust and subcrust is the inverse problem of potential theory, i. e. determining the source from its potential field.The inversion of the potential field equation above may be accomplished by application of Gauss' theorem and vector calculus (e.g. Wylie, 1951).Outside the body the potential satisfies Laplace's

Equation V2U 0. Within the body the potential satisfies Poisson's equation VU -4ir p (x, y, z).In geophysical research actual field data are measurements of forceper unit mass. SinceF U,

F =VZU and Poisson's equation may be written:

p (x, y, z) = F(x, y, z) 4ii This relation indicates that a unique solution to the inverse problem or an exact description of the mass distribution requires knowledge of the gravitation.l potential or its gradient throughout all space. Because in actual field surveys the acceleration of gravity is known only on a plane above the unknown distribution, a unique determination of mass-structural configurations is not possible from gravity data alone.Nevertheless gravity data provide a powerful geophysical constraint on the many possible crustal and subcrustal structure models which might be postulated.The accepted range of rock densities, the geological concepts of horizontal layering with the increase of density downw3rd in the earth's crust and mantle and the addition of other geophysical constraints, such as provided by seismic and magnetic measurements, reduce further the number of possible structural models and allow reasonable hypotheses to be formed concerning subsurface structures and mass distributions within the earth. PREVIOUS WORK

All applicable geological and geophysical data must be considered to determine the structure of the region west ofWashington and British Columbia realistically. A considerable amountof geological and geophysical survey work has been done in thisregion, particularly during the last decade.Detailed bathymetric mapping has outlined the major geologic features, and bottom samplinghas indicated some recent and contemporary processes involvedin their modification.Geophysical measurements, most notably magnetic and heat flow, show the region to have an unusual geologichistory while earthquake studies indicate that tectonic processes arestill continuing. The following sections briefly outline previous geological and geophysical studies of the earth's crust and mantle in the oceanic region west of Washington and British Columbia.

Phys iograpy

Murray (1941) published summaries of the early bathymetric measurements obtained by the U. S. Coast and GeodeticSurvey in the northeast Pacific Ocean. Menard and Dietz (1951) described later soundings of the U. S. Coast and Geodetic Survey.Papers by Menard (1955a and 1955b) and Gibson (1960) discuss the submarinetopography 10 of the Gulf of Alaska, showing the larger seamounts and structural features of the region.Physiographic studies by Hurley (1960) and Menard (1964) delineated and named the major physiographic provinces of the northeast Pacific Ocean.McManus (1964, 1967a, 1967b) has reviewed and extended these studies.Figure 1 is a physiographic map (after Hurley, 1960) showing the larger features of the area of study and their relationship to the North American continent. Dominating this portion of the northeast Pacific Ocean is a section of the world-encircling Oceanic Rise system a chain of high relief features extending from the continental margin off northern California to the continental margin off northern Vancouver Island. McManus (1964) named the segment of the oceanic rise west of Washington and Vancouver Island, Cobb Rise. The Juan de Fuca Ridge, determined primarily from magnetic measurements (Vine and Wilson, 1965), lays along the eastern sector of Cobb Rise.Between Cobb Rise and the continental slope off Washington and Vancouver Island is the depositional topography of Cascadia Abyssal Plain. The Alaskan Abyssal Plain adjoins the Cobb Rise on the northwest A seamount chain beginning west of the Queen Charlotte Islands extends northwesterly across the Alaskan Abyssal Plain to the Aleutian Trench southeast of Kodiak Island, Alaska. West of Cobb Rise is the Tufts Abyssal Plain, separated on the north from the 11

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Figure 1.Physiographic diagram of the northeast Pacific Ocean off Washington and British Columbia (after Hurley, 1960). 12 Alaskan Abyssal Plain by the Whitney, Schoppe and Peters ridges. The Alaskan and Tufts abyssal plains have water depths near 4 km while the Cascadia Abyssal Plain has depths of about 3 km. The high relief of Cobb Rise extends in its most prominent feature;-Cobb Seamount- -to within 40 meters of the surface. More typically, however, Cobb Rise has water depths near 2. 8 km and displayselongate hills and furrows of less than 200 meters of relief which parallel the rise crest. The continental slope, described by Shepard (1938), Hurley (1960), McManus (1964), and McManus and McGary (1967)features continental borderland topography west of Washington and glaciated topography north of the Strait of Juan de Fuca. West of the Queen Charlotte Islands the steepness of the continental slope suggests a fault5carp.Glaciation has produced substantial relief in the Dixon Entrance, Hecate Strait and Queen Charlotte Sound portion of the continental shelf and adjacent fjords of British ColumbIa. The continental shelf west of Vancouver Island and Washington is relatively narrow, glaciated and incised with numerous submarine canyons.

Geology

Very few studies of dredged rock samples are reported for the region west of Washington and British Columbia. The studies are generally restricted to rocks obtained from seamounts of the deep 13

ocean. Nayudu (1962) and Engel and Engel (1963) describe olivine-rich to olivine-poor basalts from CobbSeamount.Palagonite tuff breccia was also obtained off Cobb seamount and from Bowie Seamount off the Queen Charlotte Islands (Nayudu, 1962).Nayudu, Chastam and Woodruff (1968) report studies of prophyritic, vesicular and fine - grained basalts and palagonite tuff breccia dredged from Eickelberg Ridge. Budinger (1967à)pointed out that the basalts of high pyroxene content from Cobb searnount, Mendocino Ridge and the experimental Mohole site (off southern California) differ chemically not only from the nearby continental basalts of the Pacific Northwest (Waters, 1955) but also from other basalts in the Pacific Basin (Engel and Engel,

1963). Budinger and Enbysk (1967b) have reported a crushed whole rock potassium-argon age of 27 ± 6 m. y. for a basalt boulder dredged from 380 meters depth on the northeast slope of Cobb éeamount. Dymond eal (1968) report a whole rock sample potassium-argon age of 1.6 m.y. for a basalt boulder collected from the top of Cobb eamount at a depth of 34 meters. The 27 m. y. age is difficult to reconcile with the 3.5 m.y. age assigned to Cobb seamount by magnetic dating (Vine, 1966),whereas the 1.6 my. age is in reasonable agreement. 14

Menard (1955a) and Hurley (1960) described sediment dispersal in the region.Cobb Rise forms a barrier to western transport of terrigenous sediments which, in the northern Cascadia AbyssalPlain, are derived principally from VancouverIsland and possibly from the Queen Charlotte Sound (Duncan, 1968).The general distribution of the sediments in this region has been discussed by Hurley (1960).Nelson (1968) described sediments onthe Astoria Fan. Royse (1967) and Nelson et al, (1968) described the occurrence of Mazama ash, a key sediment timehorizon and indicator of sediment transport, on the continental slope of Washington andOregon and in the Cascadia Abyssal Plain.

Geophysics

Seismic Refraction and Reflection Measurements

Shor (1962), Milne (1964) and Shoretal. (1968) reported the results of seismic refraction measurements in the oceanic region west of Washington and British Columbia.Tatel and Tuve (1955) and Richards and Walker (1959) presented the results of seismic refraction measurements made in the adjacent continental region. The locations of these measurements are shown in Figure 2.Milne and White (1960) and White and Savage (1965) reported additional seismic measurements about Vancouver Island. 8° 136° 56 134° I° 30° 28° 126° 24° F 122° 20° 118° I 116° I 114° 1120 56° ,f' SEISMIC HEAT REFRACTION PROFILES FLOW VALUES I 54° 410 Mk23 4° I 20 60 1520 UkI8 2°I I 50 500l I 48 8° I 262 S G7 I 46 080 7 58 46°I 6 SF4 515'/ I' 2 63 I 56j 73 42° 40° 138° Figure 2. 136° 34° Location of seismic refraction lines (Shor, 1962;132° 130° 128° 26° 24° 22° 20° Mime, 1964; 118 116° 114° ______44° 112° andHerzen,Shor British etal. 1964; Columbia. and Mesecar, 1967) off the coasts 1968) and heat flow stations (Foster, of Washington1962; Von L71 L1 Reversed seismic refraction lines G 13 through G18 of Shor etal. (1968) across Juan de Fuca Ridge show that theridge is a surface expression of an upbowed crust and mantle. Thecrest and flanks of the ridge exhibit low mantle velocities.Mantle arrivals over the ridge crest are weak or absent suggesting that theMoho is unformed or discontinuous near the crest. The sediments in Cascadia Abyssal Plain are thickest nearthe base of the continental slope.Mantle depths beneath Cascadia Abyssal Plain, however, are relatively shallow 10 to 11km where water depths are 2-i/ kilometers.Refraction work on the continental shelf suggests a structural basement high nearthe outer edge of the slope and large variations in layer and crustalthickness along the Oregon-Washington shelf as well as across the shelf. Milne (1964) shot a reversed refraction profile near thesouth- em edge of the Alaskan abyssal plain which indicated anabnormally thin crust of 11.0 km. The reversed refraction profilesreported by Shor (1962) west of the Dixon Entrance show that althoughthe sedimentary section thickens, toward the continent, theMoho is at depths of 10 to 11 km. Shor (1962) also noted that theMoho becomes shallower near the foot of the continental slope, similar tothe outer ridge of many. trenches, and suggested that a buriedtrench may lie between the ridge and the slope.He also indicated that large. variations in thickness of the sedimentary layer in DixonEntrance implies 17 large-scale faulting.Mime (1964) made similar seismic observations in Dixon Entrance. White and Savage (1965) report an unreversed profile based on the April, 1958,1. 5 kiloton Ripple Rock explosion in Discovery Pas- sage, British Columbia.This profile was interpreted to indicate an average P wave velocity of 7.8km/sec and a corresponding average crustal thickness of 31 kilometers beneath the Coast Range, the Interior Plateau and the Rocky Mountains of British Columbia. No intermediate layer refracted arrivals were noted.Extensive refraction studies (Mime and White, 1960; White and Savage, 1965) in the vicinity of Vancouver Island indicate that considerable variation occurs in both velocity and thickness of upper crustalstructures. Richards and Walker (1959) reported the results of a reversed seismic profile shot in the Alberta Plains of western Canada.This work indicates a depth of 43 km to the mantle, an average Pn velocity of &. Z km/sec, a depth of Z9 km to the intermediate layer, and a P* (intermediate layer) velocity of 7. 2 km/sec. From their refraction studies in Washington, Tatel and Tuve (1955) conclude that the crust is relatively thin, 19 km, beneath Puget Sound and that the Pn velocity is 8. 0 km/sec.East of the sound they conclude that the crust is 30 km thick.The 30 km depth agrees well with the depth given by White and Savage (1965) for the Coast Range about 250 km to the north. Hess (1964) attributed variations in mantle velocities in the Pacific Ocean to velocity anisotropism.He suggests (Hess, 1965) there may be a 5% to 10% azimuthal velocity variation due to anisotropic propagation, such that low velocity paths should parallel the magnetic anomalies and high velocity paths lie perpendicular to the magnetic anomalies (and consequently the active mid-ocean ridge- rise systems).In the region of study the magnetic lineations (Figure 4) trend generally north-south while nearly all the seismic refraction profiles (Figure 2) trend east-northeast.Since the seismic refraction profiles intersect the magnetic lineations at nearly a constant angle in this region it is unlikely that velocity variations along the same line are due to anisotropic propagation; rather, they are more likely to be a function of temperature, petrology or interface irregularities. Hamilton (1967) described results of seismic reflection studies in the Gulf of Alaska, which showed sediment thicknesses of 680 meters in the southeast part of the Alaska Plain and 470 meters in northern Tufts Plain.Turbidite thickness decreased with distance from the continental sources.Reflection profiles (Hamilton and Menard, 1968) across the northern end of Juan de Fuca Ridge to the continental terrace off Washington indicate 100 to 200 meters of turbidites near the eastern edge of Juan de Fuca Ridge, 600 meters near the center of Gas cadia Plain and 1600 meters nearthe base of 19 the continental slope.Hamilton and Menard (1968) concluded that probable uplift of the ridge occurred after turbidite depositionbegan. Hamilton (1967) also noted that the turbidite layers areundistorted, contrary to what might be expected from crustal spreading,and concluded that the whole area has been quiescent since-middleTertiary. Bennett and Grim (1968) and Grim etal. (1968)have described seismic reflection observations on the continental shelf off Washington, which show a broad spectrum of tectonic activity. Broad folds are observed on La Perouse bank; flat-lying bedsexhibit- ing high angle normal faults occur west of the OlympicPeninsula; and gentle folds and piercement folds with exposed diapiric cores are noted south of Quinault Canyon.

Earthquake Observations

The oceanic region west of the coast of North America, between the Columbia River and Dixon Entrance, is part ofthe great circum-Pacific seismic belt.The early compilation of Gutenberg and Richter (1941) indicated that the circum-Pacific seismicbelt branches in the vicinity of southern California. The westerlybranch of the belt, consisting of predominantly shallowearthquakes, passes out to sea in the vicinity of Cape Mendocino, California and re- enters the continental region in the vicinity of the QueenCharlotte Islands and southeast Alaska.Recent compilations by McManus 20 (1965), Mime (1967), Bolt etal. (1968), and Tobin and Sykes (1968) indicate that the earthquake epicenters of this region show preferred trends which appear to relate to known physiographic features. Figure 3 shows the distribution of earthquake epicenters for this region which occurred between 1899 and 1960. In southern California the earthquake activity of the continental margin region of North America is associated with the San Andreas fracture system. Northward, the earthquake activity extends along the coast of northern California, across the Mendocino Escarpment and then northwesterly along the Blanco fracture zone.After an apparent offset, earthquake activity begins anew west of Vancouver Island.It then extends northwesterly along the seaward side of the Queen Charlotte Islands into southeast Alaska where it appears to join the Fairweather-Chugach-St. Elias and Denali fault systems (Benioff, 1954, 1962; St. Amand, 1957; Tobin and Sykes, 1968). Compilations of earthquake activity in Washington by Rasmussen (1967) and in Oregon by Berg and Baker (1963) indicate the relative quiet of southern Washington and Oregon.This area of anomalous quiet between belts of moderate earthquake activity apparently extends westward over Cascadia Plain.The major portion of the earthquake energy released in the northwest North America portion ofthe circum-Pacific seismic belt occurs along the shelf-slope region of southeast Alaska and British Columbia.Lesser amounts of energy 21

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Figure 3.Earthquake epicenters in the Pacific Northwest (after Mime, 1967). 22 are released in California and in Puget Sound, Washington (Mime,

1967). The Juan de Fuca Bidge area is comparatively aseismic except at the extreme northern end.Epicenters of numerous small earthquakes at the north end of the ridge appear to be continuous with the band of earthquake activity associated with the Explorer Ridge-Scott Island fracture zone-Queen Charlotte Trough complex.It is difficult to assign the earthquake activity of this region to any one physlo- graphic feature.One possible exception is the concentration of earthquakes located by Tobin and Sykes (1968) near 500 N lat., 129. 6° W long.Precision Depth Recorder (PDR) records indicate that, at this position, the Explorer Ridge exhibits a'rift valley-like structure. Recent faulting is suggested by the absence of sediment on tilted benches evident along the trench walls.The PDR records also show that approximately 40 to 50 km southeast of the north tip of the trench, ponded sediment is severely tilted.Theabsence or distortion of sediment in this area suggests recent or contemporary tectonism probably related to the observed seismicity. The linaments exhibited by the earthquake epicenters of Figure 3, particularly along the Queen Charlotte Islands, suggest transcurrent faulting.However, in discussing the predominance of transcurrent faulting (determined from fault-plane solutions) in circum- Pacific earthquakes, Hodgson (1957) notes that the zone from 23 southeast Alaska to California is unique in that all types of faulting occur. Six fault-plane solutions have been reported in the region of the continental margin between Cape Mendocino and the Columbia River (Tobin and Sykes, 1968; Hodgson and Milne, 1951; Hodgson and Storey, 1954; A1germissen et a1., 19 65),Epicenters of the April 13, 1949,and April 29, 1965,earthquakes are located in the southern Puget Sound area.The preferred fault plane solutions for these earthquakes indicate thrust faulting for the April 13, l949,. shock south of Tacoma, Washington (Hodgson and Storey, 1954) and normal faulting for the April 29, 1965, shock near Seattle, Washington (Algermissen etal., 1965). The epicenter of the earthquake of June 23, l946,is in the Strait of Georgia (between Vancouver Island and the mainland). The fault plane solutions of this earthquake indicate normal faulting.The preferred solution contains a right lateral strike-slip component with the fault plane oriented N 23° W and dipping 600 east (Hodgson and Mime,

1951). The earthquakes of August 22, 1949, and March 31, 1964, occurred in the Alaskan Abyssal Plain immediately west of Queen Charlotte Sound. The fault-plane solution of the August earthquake indicates normal faulting with the more probable solutions oriented N 56° E and dipping 67° east (Hodgson and Mime, 1951). A right 24 lateral transcurrent component of motion is associated with the preferred solution.Tobin and Sykes (1968) have indicated a transcurrent fault solution for the March earthquake.The fault-plane solution for right lateral motion indicates a strike of N 15° W. The earthquake of August 22, 1949, was located in Dixon Entrance between southeast Alaska and the Queen Charlotte Islands. The corresponding fault- plane solutions indicate transcurrent faulting. The preferred solution has a strike of N 29° W, a dip of 77° east and exhibits right lateral motion.This solution is consistent with observed faulting visible on adjacent land masses (Sutherland-Brown, 1966) and with the motion of the Queen Charlotte fault postulated by Benioff (1954), St. Amand (1957), and Wilson (l965b). Stauder (1967) has noted that the limitations imposed by seismic wave propagation and the geographic distribution of seismic stations favor transcurrent fault-plane solutions, particularly in cases of shallow focus earthquakes.Tobin and Sykes (1968) state that in their earthquake mechanism solutions strikes and dips may be uncertain by as much as 300. The limited number of fault-plane solutions with their expected uncertainty are probably insufficient to describe the tectonics of this region.They do suggest, however, that the continental margin between Cape Mendocino and Dixon Entrance is a region of block faulting between two regions, southeast Alaska and California, which exhibit 25 predominantly right lateral trans current faulting.

Heat Flow

Foster (1962), Von Herzen (1964), and Mesecar (1968) have reported heat flow values for the region west of Oregon, Washington, and British Columbia. A composite map of these data is shown in Figure 2. No geophysical interpretation of the heat flow data or correlation of these data with other geophysical measurements has been reported. The average of 25 heat flow values between 44° N lat. and 55° N lat. is 2, 45cal/cm2/sec,approximately double the presently accepted world average. Von Herzen and Uyeda (1963) noted that the highest heat flow values observed over the East Pacific Rise occur in two narrow zones approximately parallel and symmetrically oriented with respect to the rise crest.They suggest these observations can be accounted for by the existence of an upward-moving mantle convection current beneath the crest of the rise and lateral spreading of the material near the surface. The values of the limited number of heat flow stations within a region extending 300 km to either side of Juan de Fuca Ridge show an asymmetry. The average of nine values west of Juan de Fuca Ridge is 1.89cal/cm2/sec.The average of 16 values east of the ridge is 2. 77cal/cm2/sec.Duncan (1968) has indicated average 26 rates of deposition for Cascadia Plain during Pleistocene time of 0. 12 to 0. 18 cm/yr, and rates of 0. 02 to 0. 003 cm/yr during the postglacial period.The high deposition rates for Pleistocene time occurred in the area of high heat flow, suggesting the equilibrium heat flow rates may be 10 to 40% greater than those observed. A higher heat-flow rate, east of Juan de Fuca Ridge, is also suggested by the anomalously high value of 7.27cal/cm2/secnear the north end of the ridge, in an area which has had a lower sediment deposition rate. The differences in the average heat-flow rates east and west of Juan de Fuca Ridge appear significant and suggest major differences in either temperature or conductivity of the crust or subcrust in these areas. The anomalously low heat flow value (0.80cal/cm2/sec)west of Juan de Fuca Ridge may be explained as due to a locally disturbed environment (Von Herzen, 1964).Heat flow values in the proximity of the Queen Charlotte Trough appear normal; however, west of Dixon Entrance, they appear unexplainably high.

Magnetic s

Raff and Mason (1961) and Vacquier etal. (1961) have presented results of an extensive magnetic survey in the Northeast Pacific which extends from 400 N lat, to 520 N lat. and from the continental slope of Oregon, Washington, and British Columbia westward approximately 27 400 km. Emiliaetal. (l968a) presented magnetic maps which extend the magnetic survey between Cape Mendocino and the Strait of Juan de Fuca eastward over the continental slope and shelf. Magnetic measurements were made over the Queen Charlotte Trough and Scott Island fracture in conjunction with Oregon State University gravity surveys.Total field magnetic intensity maps based on these data are shown in Figure 18 and Figure 2 1.Magnetic measurements were also obtained along two tracklines west of Graham Island in 1966 during the YAQUINA long cruise I (YALOC I). These two lines, combined with two lines west of Graham Island reported by Pitman and Hayes (1968), indicate the general magnetic pattern of the area west of Graham Islard.The data from the above surveys, compiled and shaded in Figure 4, show the magnetic anomaly pattern of the region of the Northeast Pacific discussed in this thesis. The magnetic anomalies, delineated by these surveys, are long linear features having a predominantly north-south trend. Two of the magnetic lineations observed over Cascadia Plain extend landward to the shelf break, apparently crossing the continental rise and slope unaltered.Calculations of source depths (Raff and Mason, 1961; Emilia et al.,,1 968a) indicate large variations in depth to anomaly sources. These depths range from the water-crust interface to near the crust-mantle boundary with a tendency for the source depths to increase landward.Emiliaetal. (1968b) explained the nearshore 28

1400 135° 130° 125° 120° 550 55

50° 500

45° 450

140° 135° 130° 125° 120°

Figure 4.Total magnetic anomaly map of the region west of Washington and British Columbia (a compilation from Raff and Mason, 1961; Vacquier etal.1961; Emilia etal. ,1967a; Pitrnan and Hayes, 1968, see text). 29 magnetic anomalies as due to probable seaward extensions of coastal volcanics, Talwanietal. (1965) noted that a magnetic anomaly profile obtained across Juan de Fuca Ridge was similar to magnetic profiles obtained over the East Pacific Rise in the South Pacific.They concluded that Cobb Rise is part of the mid-ocean ridge system with "the important difference that the eastern flank of the rise is under the continent." Vine (1966) interpreted the magnetic anomaly pattern as reflecting sea-floor spreading, contemporaneous with reversals in the earth's paleomagnetic field.In this hypothesis the spreading axis west of Washington and Vancouver Island is represented by Juan de Fuca Ridge.The sea floor is thought to spread at equal rates (approximately 2.9 cm/yr) from the N 300 E trending axis of the ridge toward the continent and oceanward.Large- offsets of the spreading axes are interpreted as due to transform faulting (Wilson, 1965a, 1965b).This hypothesis has been extended by Heirzleretal. (1968), Pitman and Hayes (1968), and Pitmanetal. (1968).These investigators indicated the ages of each of the magnetic anomalies as based on comparative spreading rates of other mid-ocean rises and rifts. Morgan (1968a, 1968b) postulated that the separating of large crustal blocks generates the observed magnetic lineations.He suggests that a Northeast Pacific crustal block is moving northwesterly, away from the North American continent crustal block toward the III] Aleutian and Kurile-Kamchatka trenches.The fracture of separation is indicated by Gorda and Juan de Fuca ridges and the Blanco fracture zone. Numerous oblique faults (or lineations) offset the magnetic patterns (Raff and Mason, 1961).Morgan (1968b) interpreted these lineations as indicating small blocks moving independently of each other,Pavoni (1966) suggested that these lineations are due to wrench faulting, mostly left lateral, which dissected an older crustal unit.An oblique fault is noted offsetting the anomaly which defines the spreading axis of Juan de Fuca Ridge.This suggests that at least some of the oblique faulting has occurred in recent time. Although the interpretation of the magnetic anomalies of Juan de Fuca Ridge suggests a presently active spreading ridge (Vine and Wilson, 1965; Pitman and Hayes, 1968) other geophysical evidence suggests it is probably quiescent (Hamilton, 1967; Tobin and Sykes,

1968).Bostrom (1967) has suggested that the spreading axis between the Blanco and Queen Charlotte transform faults is presently being reactivated beneath western Oregon and Washington. The gradients and intensities of the magnetic anomalies imply a crustal source.Model studies by Mason (1958) suggest an anomaly source located in the crustal layer beneath the sediments.Vine and Wilson (1965) suggest the origin is predominantly due to remanent magnetization existing in the oceanic basalt layer wherein juxtaposed normal and reversely polarized rock generate the observed steep 31 magnetic gradients.Theoretical analysis by Bott (1967) supports the interpretation of a shallow source coupled with normal and reverse magnetization.Ozima et al, (1968) have noted a more than 3 fold increase in saturation magnetization of submarine basalts when heated; they suggest that the magnetic lineations may reflect the effect of localized heating of the basaltic crust.

Gravity

Harrison etal, (1957) and Worzel (1965) have published pendulum measurements west of Washington and British Columbia. No analysis of these data has been reported.It is interesting to note that, although the region west of Washington and British Columbia exhibits large positive and negative gravity anomalies, none of the 43 measurements of Harrison etal, (1957) or Worzel (1965) are located so as to indicate the existence of these anomalies. Woollard and Strange (l96Z) have published regional free-air and Bouguer anomaly maps of the Pacific Basin,Estimates of the depth to mantle, based on these gravity data, were presented. The USC&GS (Erickson, 1967) has made surface ship measurements along a gravity meter check line west of Cape Flattery, Washington. The Dominion Observatory of Canada (Tanner, 1966), in a continuing gravity program, has made bottom gravity meter measurements in the region of Dixon Entrance, Hecate Strait and Queen Charlotte Sound. No analysis of these data are reported and the gravity values are unpublished.Figure 5 shows the station positions of the pendulum measurements of Harrison et al. (1957) and Worzel (1965), the position of the USC&GS surface ship measurements (Erickson, 1967), and the area of bottom meter measurements of the Dominion Observatory. Dehlinger etal. (1966a) have reported gravity measurements along the Inside Passage of Alaska and British Columbia.These measurements show a rapid decrease in free-air anomalies toward the continental interior in traverses along the fjords.Banks (1969) reduced 104 of these rneasurementsto isostatic and Bouguer anomalies.His results indicate a small residual tendency of the gravity anomalies to decrease toward the Coast Range mountains. He also notes a change in the probable depth of compensation north of Juneau, Alaska and south of the San Juan Islands, Washington. Figure 5 shows the locations of the sea gravity measurements described above. Heiskanen (1938) and the Dominion Observatory of Canada (Stacy, 1967) reported land gravity values in the nearshore regions of British Columbia including the Queen Charlotte Islands.Although Heiskanen (1938) presented isostatic anomalies, no other analysis was reported. Walcott (1967) presented a Bouguer anomaly map of southwestern 33

140 J38 136 134 132 130 128 56 26 24 22 156

54 -- .4.

50 0

4611

40 __-_i.ih

Figure 5.Map of surface ship gravity meter track lines off the coasts of Washington and British Columbia. Map also locates underwater gravity meter measurements of the Dominion Observatory of Canada, pendulum stations of Harrison (1957) and Worzel (1965) and computed seamount values (App,2). 34 British Columbia which includes results of previous work in the area by Miller (1929), Garland and Tanner (1957), White and Savage (1965), and Dehlingeretal. (l966b). In his analysis of the gravity anomalies he concluded that the coastal anomalies suggest a structural analogy between Vancouver Island and the Queen Charlotte Islands.He also presented a gravity profile from Vancouver Island across the Strait of Juan de Fuca to the Olympic Peninsula.This profile shows a steep gravity gradient over the Strait of Juan de Fuca which Walcott (1967) suggests indicates large differences in crustal thicknesses on either side of the strait. Woollard and Rose (1963) and Sauers (1966, unpublished) have presented Bouguer anomaly maps of Washington without interpretation. Stuart (1961) constructed a Bouguer anomaly map of the Olympic peninsula and Puget Sound of Washington and presented a two- dimensional crustal model that extends from the Olympic Mountains to the Puget Sound.His model shows a high density block of volcanic rocks extending from the Olympic Mountains where they are exposed to the area between Seattle and Tacoma, where they lie 16 to 20 km deep. Danes etál. (1965) from a detailed gravity survey of the southern Puget Sound area, concludes that the entire Puget Sound area is a major regional gravity low.He presents a north-south cross section to illustrate his analysis, which shows large variations in the depth 35 to Moho. Dehlingeretal. (1968) have combined gravity data west of Oregon, land gravity data in Oregon and Washington (Thiruvathukal, 1968; Sauers, 1966 unpublished) and layer depths and velocities obtained from seismic refraction work (Shor etal., 1968; Hill and Pakiser, 1966) to construct two-dimensional crust- subcrustal cross sections.These cross sections extend from Tufts Abyssal Plain eastward across Juan de Fuca Ridge and Oregon to central Idaho. One cross section passes through southern Washington and the other through north central Oregon.These postulated cross sections indicate the existence of a low-density mantle beneath Juan de Fuca Ridge, a relatively thin crust (23 km) in western Oregon and a crust which thickens to more than 40 km beneath the Cascade Range, eastern Oregon and western Idaho. NEW GRAVITY DATA

Surface-Ship Gravity Measurements

The important sea gravity measurements described in the previous section, although of use in geodetic studies, are insufficient in detail to delineate the major anomalies and do not provide adequate control for structural analysis.The Department of Oceanography at Oregon State University made surface ship gravity measurements from 1964 through 1967 to provide a sufficient density of gravity stations to map the area off Washington and British Columbia. These measurements, combined with adjacent land measurements, provide control for an analysis of the major structures.The initial traverses made in 1964 and 1965 were made perperidicularto the coasts of Washington and British Columbia in anticipation of the structural form of the continental margin. These early measurements, however, revealed new structures, structural trends and areas exhibiting unexpectedly large negative anomalies. To delineate and study these gravity features in detail, additional measurementswere made during 1966 and 1967 along traverses oriented generally perpendicular to the observed anomalies.Figure 5 shows the surface-ship gravity meter track lines. Gravity measurements were made with LaCoste-Romberg 37 gimbal-suspended,surface-ship gravity meter S-9.Gravity meters operating on moving platforms are subjected to horizontal and vertical accelerations in addition to the acceleration of gravity.These accelerations may be as much as several magnitudes greater than the anomaly sought, To achieve an accuracy of a few parts per mu- lion, three corrections are made to the gravity meter measurement: 1) a horizontal or Browne correction for lateral accelerations acting on the meter; Z) a vertical correction for nonlinear responses of the gravity meter under high vertical accelerations and 3) an Ebtvbs correction for the change in the earths centrifugal force due to movement of the ship along the earth's surface.Each of the three corrections is subject to uncertainties.These uncertainties plus the uncertainty in ship position, due to navigation inaccuracies, limit the accuracy of these gravity measurements to about ± 5 mgal (Dehlinger etal., 1966à).Uncertainties in the best measurements are estimated to be ± 3 mgal. The gimbal suspension of meter S-9 permits the gravity meter sensing element to swing freely in response to horizontal ship accelerations.The horizontal accelerations cause the meter to swing or be deflected from the vertical. The sensing element of the gravity meter thus measures the vector sum of gravity plus the negative of the horizontal acceleration,Gravity is obtained from the 2 solution of the relati.on g0 -g whereg = an average value of gravity for the earth,g measured gravity and 8 = angle of deflection of the sensing element from the vertical.This equation is known as the Browne correction or horizontal correction (Browne, 1937) and is applicable for long-period accelerations.LaCoste and Harrison (1961) have presented a comprehensive analysis of surface- ship gravity meter operations involving simultaneous periodic horizontal and vertical accelerations.Their analysis, which included fourth order terms, showed the Browne correction equation is correct for the LaCoste and Romberg gimbal suspended meter, provided 0 remains small and that gravity values are obtained over a length of time that is long compared with ship acceleration periods. The measurement of the horizontal acceleration or Browne correction is accomplished by measuring the angle 0or the deflection of the meter from the vertical.This measurement requires a stabilized reference which, in meter S-9, is provided by a long- period inertia bar (LaCoste, 1967).The free period of the inertia bar in meter S-9 is two minutes.Dehlinger (1964) described the reliability of gimbal suspended gravity meters at sea and the effect of damping of the inertia bar with particular reference to meter S-9. He noted that measurements can be accurate to within ± 3 mgal at Browne corrections up to 300 mgal and that errors became systematically more positive with increasing Browne correction. A statistical study by Dehlinger etal. (l966a) indicated the root mean 39 square (rms) uncertainty in the S-9 Browne correction was less than 2 mgal at Browne corrections between 0 and 400 mgal. No adjustment or modification of the S-9 horizontal accelerometers was made sub- sequent to the above study and consequently it is expected that Browne correction errors in this study are approximately the same as those previously reported. No corrections or adjustments of the gravity measurements were made to compensate or correct errors due to inaccurate horizontal acceleration determinations. The vertical component of measured gravity (g) on a ship at sea is actually a measurement of gravity plus vertical acceleration. Vertical and horizontal accelerations observed on ships are 10,000 to more than 100, 000 times greater than the desired gravity meter accuracy. These accelerations are of relatively short duration, however, and therefore easily removable by filtering.Filtering of the vertical acceleration in the LaCoste and Romberg gravity meter is accomplished electronically.Gravity meter and filter responses must be extremely linear to avoid errors in gravity indications (Harrison and LaCoste, 1968).The LaCoste and Romberg gravity meter S-9 was designed to operate with vertical accelerations of ± 50 gals, on the assumption that vertical and horizontal accelerations experienced at sea were approximately of the same size.It was later determined that vertical accelerations very often exceeded the design limit and consequently exceeded the assumptions of linear response. 40 The causes and results of nonlinear operation have been discussed by LaCoste (1967).Factory linearity adjustments have been made to the LaCoste and Romberg meter S-9; however, the linearity has been noted to change with time, making it necessary to both retest the meter periodically and to apply different correcting equations to the measured gravity.The parameters, method and empirical equations used to correct LaCoste and Romberg meter S-9 measurements were described by Dehlinger etal. (1966a).The particular equations used to correct the data used in this study are listed in Appendix 1. The west to east component of ship velocity increases the centrifugal acceleration experienced by the meter and thus decreases the measured gravitational acceleration.This is termed the EtviSs effect (described in Heiskanen and Vening Meinesz, 1958, p. 158) and can be computed from the relation

= Zcosv mgal where

is the angular velocity of the earth (0.00007Z92/sec), 4 is the

latitude of the point of observation andvis the ships west to east velocity component. The correction so determined is the Etvs cor- r ection. The estimated uncertainty associated with gravity meter measurement, recording, reading, drift, and base station tie is ± 1 mgal. The estimated uncertainty in the vertical and Browne correction is ± 3 mgal when sea conditions produce Browne corrections 41 less than 400 mgal.The estimated uncertainty In Eötvös correction is ± 1 n-igal (± 0.2 knot and ± 1.00 course heading).The total estimated rms uncertainty, exclusive of position error, is (32 +12+12)1/2± 3.3 mgal. Loran A and radar were the principal aids to navigation used

for the gravity surveys west of Washington and British Columbia.In areas where radar ranging is most accurate, such as the fjords of British Columbia, an estimated uncertainty in gravity measurement due to position error, may be as low as ± 0.5 mgal (Dehlinger etal., 1966a).In the fringe areas of radar navigation where a combination radar fixing and dead reckoning are used, gravity uncertainties due to navigation errors may reach ± 3. 5 mgal.The rms uncertainty in gravity measurement due to the positional accuracy attainable with

Loran A varies from ± 3. 0 mgal over the continental shelf to6 mgal over the Tufts and Alaskan abyssal plains. Consequently, the total rms uncertainty in the sea gravity measurements west of Washington and British Columbia ranges from less than± 3.4mgaltonear± 7.0 mgal. For this reason, three levels of uncertainty are assigned the contours of the free-air anomaly maps shown in Figures 6 and 7.

Reduction of Gravity Data

The LaCoste and Romberg sea gravity meter produces a reading 42 51° 135° 134' 133° -: ' 132° 131° 7jrr 130° 129° 28° - 127° - - - 126° - 125° _:- T 124° - --- 123° - 0l ,1 ) _ ' k /\ 0 ° ___ / .'o \ \ i[ ;/J) r (/ \.o/ 8 L 1/) I (/ '° _ r \ ____; /) : 7 _ (( _ I _: _ Ltiip4 Figure 6. ColumbiacoastFree-air of Washingtongravity River toanomaly Cape and Scott. British map of Columbia the region from off the west .. 132° 131° 130° 129° 128° 127° 126° 125° -t 124° 23° l2/ 55. -;a------39° l38 137° 36° 35° - - :134° I33 132° I3I I3O 29° J28 I27 126 55 43 /_- \ - I i- I -:- '\ ]c \ ) \ -I T1' 1 FREE-RGRAVITYANOMALYMAPWEST COAST OF BRITISH COLUMBIA \J , ,) =-/, -- \ ,\ \ 0 t! I1- !II b z OCTOBER, 967

/7 I 52 2 :J\ \ H'\\ /( \ NN 5 __) j \\J ______50 39° 38° 37° 36° r\\\ 35° 34° I______132° Th 131° - 50 1'igure 7. DixonwestFree-air coast Entrance. gravity of British anomaly Columbia from Cape Cook to map of the region off the 33° 30° 29° 28° 27° 126° 44 which when corrected by the Etvs, Browne and vertical corrections and multiplied by a meter (proportionality) constant is directly related to the earth's gravitational acceleration at the point of measurement. The LaCoste and Romberg meter measures difference in gravity rather than absolute gravity; consequently, to determine the earth's gravitational acceleration at a point the meter reading is compared or tied to a base station.Gravity measurements made aboard the USCGC YOCONA in 1964 and 1965 were tied to the USC&GS station US 80 (Duerkson, 1949) in Astoria, Oregon and trans- ferred to the ship using a Worden land gravity meter.Gravity measurements made aboard the R/V YAQUINA in 1965, 1966 and 1967 were tied to the Oregon State University base station at the Marine Science Center (Berg and Thiruvathukal, 1965) in Newport, Oregon. These two base stations are tied to the U. S. Standard Station in Washington, D. C. The output of the LaCoste and Romberg gravity meter S-9 is recorded by six analogue recorders.The recorders log meter reading, frequency and amplitude of vertical motion, displacement of the meter from operating center, Browne correction and two orthogonal components of horizontal acceleration.The meter reading, vertical correction, and Browne correction are determined from these data. The ship's position, velocity, and heading, and an Etvs correction, are determined from data provided by the ship's navigation log.The 45 meter constant, base station gravity value and base station meter reading are logged constants determined prior to a survey. Observed gravity, as a function of space and time is determined

from these data by using the relation:g mbk + mk + E + V

where g= observed gravity,k = meter constant,mbbase meter reading,m = on station meter reading,Ec = Etvbs cor-

rection, and V c= vertical correction.The data obtained aboard the USCGC YOCONA and R/V YAQUINA were digitized by hand and reduced using a computer program written in Fortran IV for the Control Data Corporation (CDC) 3300 digital computer. For mapping and analysis, observed gravity values were reduced to free-air anomalies.Free-air anomalies are defined by the relation:

F.A. = g+

where F. A.is the free-air anomaly, g is observed gravity,

f is the free-air correction and g is theoretical or normal C t gravity.The free-air correction is a correction for elevation above a sea level datum and is zero at sea level.Theoretical or normal gravity was determined from the 1930 International Gravity Formula:

g = 978. 000 (1 + 0.0052884 sin - 0. 0000059 sin22)gals. where 4 is the latitude of the station (e. g.,Heiskanen and Vening Meinesz, 1958).The theoretical or normal gravity was programmed 46 for use on the CDC 3300 computer and incorporated in the general sea gravity reduction computer program.The computer print out indicated the latitude, longitude and free-air anomaly of each gravity station. These data were then plotted and contoured by hand.

Distribution of Gravity Measurements

Figure 5 shows the location of the track lines along which surface- ship measurements of gravity were obtained.The total of the track lengths is approximately 11,000 nautical miles (20, 000 km) and the total number of surface-ship stations is approximately 15, 400. These figures indicate a linear station density of one station approximately every 1.4 nautical miles (2. 5 km). The addition of the pendulum stations reported by Harrison et al. (1957) and Worzel (1965), the underwater gravimeter measurements of the Dominion Observatory of Canada (Stacy, 1967),the surface- ship measurements of the USC & GS and coastline land measurements (Woollard and Rose, 1963; Sauers, 1966; Walcott, 1967; Stacy,

1967) increases the number of gravity stations which control thecon- tours of Figures 6 and 7 to approximately 16, 100.The area mapped is about 635, 000km2.The areal station density is then about one station per 40km2. A station density of one per 40 km2 indicates a relatively thin 47 coverage when compared to a typical land gravity survey.These figures represent an average but in the actual mapping, the track lines and consequently the data points have an increased areal density in the regions of steep gravity gradients.Further the high density of stations along the track lines show the anomaly wavelength to be expected in the region of measurement. If a gravity anomaly can be assumed equidimensional (circular) then the half wavelength of the anomaly observed along the track line which crosses the anomaly approximates the diameter of the anomaly.

In this manner the number of gravity stations which describean anomaly can be estimated.Table 1 shows the average half wavelength for gravity anomalies of different physiographic provinces.

Table 1.Information on the gravity anomaly size and gravity station density in different physiographic provinces

Gravity anomaly Approximate Station half wavelength anomaly area densityper (km) (km2) anomaly

Cascadia Abyssal Plain 7 40 1.0 Washington Continental Shelf 18 250 6.3 Cobb Rise 22 380 9.5 Alaskan Abyssal Plain 45 1600 40.0 Hecate Strait-Queen Charlotte Sound 12 115 2.8 The Cascadia Plain exhibits the shortest wavelength anomalies of any province (and the smallest amplitudes) even though water depths in this region are near 1500 fathoms (3000 meters). Anomaly wavelengths, although of comparative size, are shorter in the Dixon Entrance-Hecate Strait-Queen Charlotte Sound area than on the continental shelf off Washington. The Alaskan Plain has the longest anomaly wavelengths--twice as long as observed over the complex topography of the Cobb Rise.Table 1 also indicates that the contours of the anomalies of the Gas cadia Plain and the Dixon Entrance-Hecate Strait-Queen Charlotte Sound areas are inadequately controlled by available gravity measurements. However, because of the uneven distribution of stations, the Cascadia Plain and the Dixon Entrance- Hecate Strait-Queen Charlotte Sound anomaly contours are better controlled than indicated, while the anomaly contours of the Alaskan

Plain have considerably less control than indicated by Figure5 The actual gravity control for each anomaly can be ascertained by corn- paring Figures 6 and 7 with Figure 5.

The Free-Air Gravity Anomaly Maps

The geographic area covered by the free-air gravity anomaly maps of Figures 6 and 7 extends from the Columbia River of Oregon and Washington north to the Dixon Entrance and Alaska and from the coastline of Washington and British Columbia west to a line 49 approximately 450 kilometers west of and parallel to the edge of the continental shelf. The free-air gravity anomalies are derived primarily from surface ship gravity measurements obtained along the traverse lines shown in Figure 5.The methods used in making the measurements, the reduction of the measurements and methods of determining the uncertainties in the processed data are outlined in the previous section and discussed in Appendices. 1, 2 and 3.Supplementary data and methods of obtaining the final contours are discussed in Appendix 4. The free-air gravity maps have three levels of contouring; solid lines, long broken lines and short dashed lines indicating differences in root mean square uncertainties in measurement. The solid lines have an estimated rms uncertainty of ± 5 mgal.Solid contours occur in areas in which weather conditions, navigation and meter operation have resulted in better than average surface-ship sea gravitymeasure- ments. Within the areas of solid contours are regions in which the rms uncertainty in measurement is considerably better than ± S mgal. For example, Dehlinger etal. (1966b) have estimated anrms uncertainty of ± Z mgal in measurements along the fjords of the Inside Pas- sage of British Columbia and Alaska.Solid contours also imply that the half wavelength of the anomaly described by the contours is equal to or greater than the spacing of the available data. A possible exception to this implication is the flat field of Cascadia Plain.This 50 region is discussed further under the subheading; Cascadia Abyssal Plain. The long broken lines indicate an estimated rms uncertainty of ± 7 mgal. The accuracy of the data in areas described by broken lines has been limited by such factors as marginal meter operation, limited navigation control, and adverse sea conditions.Variations of the smoothed data along the traverse lines suggest that the half wavelength of the observed anomalies is, on the average, at least of the same order as the spacing of the available data, particularly in the areas west of the continental slope. The dashed lines are inferred and are based on isolated gravity values, observed gravity gradients, or known bathymetric trends and gradients.Dashed contours are, therefore, estimations of an unknown gravity field. Undoubtedly the gravity field of this region is more complex than that indicated by the contours of Figures 6 and 7.The limitations of surface ship gravity measurements and navigation facilities necessitates data smoothing which partially removes the high frequency anomalies and limits the useful size of the survey grid.However, the data are sufficiently accurate and numerous to delineate and permit analyses of major structural features. 51

INTERPRETATION OF THE FREE-AIR GRAVITY ANOMALY MAPS

The large density contrast between sea water and the underlying sediments and rock and the varying proximity of the heavier material cause the contours of the free-air anomalies to "follow" the bathymetric features.That is, the overlying water is essentially "trans parentt' to gravity measurements. The density contrast between ocean sediment and basement rock also leads to a partial transparency of the sediment; however, the smaller density contrast between sediment and basement rock necessitates larger volume and/or density changes for a given gravity variation.Therefore, in areas in which the contours of the free-air gravity anomalies do not follow the bathymetric contours large variations in sediment thickness and/or severe changes in the structure or density of the subsurface rock can be expected.

Gravity Anomaly Trends

The free-air gravity anomaly contours in the area between the Columbia River and Cape Cook trend approximately N 500 W, and change in the area between Cape Cook and Dixon Entrance to approximately N 40° W. The trend of the anomalies parallels in general the continental shelf, the coastline and the great structures of the 52 Western Cordillera.This trend is typified by the strong negative anomalies of the Queen Charlotte Trough and Scott Islands fracture zone and by the adjoined positive highs over Eickelberg Ridge. In the physiographic provinces that McManus (1967b)termed the Continental Rise and northern Cobb Rise, the anomaly contours are extensive and irregular.However, in some instances they are somewhat symmetrical and strongly positive, outlining seamounts which typically exhibit the shape of an elliptical cone with the long axis of the basal ellipse along the regional trend. The positive free-air anomaly over Bowie Seamount, the negative anomaly associated with the Explorer Trough, and a small positive anomaly northeast of Cobb Seamount appear as the only large features with a dominant northeast-southwest trend. Numerous anomalies cover Cascadia Plain.These positive and negative anomalies are of small amplitude and area and show no apparent dominant trend.

Isostatic Equilibrium in the Northeast Pacific Ocean

Isostatic equilibrium means that at some depth- -termed the depth of compensation--the overlying crust and subcrustal materials are in hydrostatic equilibrium. More generally the lighter rigid crustal blocks of the earth are "floatingt on a more dense plastic substratum. The observed topographic expressions of crustal blocks 53 such as mountains on a small scale or continents on a large scale are compensated at depth by either lighter 'roots" or lower density materials that extend beneath the Moho. An isostatic gravity anomaly, based on measured gravity and assumptions of the general earth structure, indicates the state of compensation of the crustal elements. Negative anomalies over oceanic regions indicate under compensation and positive anomalies over compensation.Since the free-air reduction is essentially a type of isostatic reduction, corresponding to the case when topographic masses are considered compensated at zero depth (Heiskanen and Vening Meinesz 1958, p. 209) the average regional free-air anomalies over the ocean directly reflect the isostatic status of the region,Therefore, oceanic areas having a negative free-air anomaly are considered to be under compensated or to have a mass deficiency above the isopiestic level,Areas having a positive anomaly have a mass excess and are considered over corn- pensated. The gravity anomalies of the abyssal plains, seamount province and continental shelves, shown in Figures 6 and 7, undulate about the zero level with amplitudes of approximately ± 20 rngal indicating that regionally the area is approximately in isostatic equilibrium.Excep- tions to this are relatively small areas which show large negative or positive anomalies. Such areas are: the Queen Charlotte trough, Scott Island fracture zone, numerous searnounts, the Explorer trough 54 and fjords along the coast of British Columbia. Anomalies associated with these features do not necessarily imply deviations from isostatic

equilibrium but rather special mechanisms of anomaly generationor isostatic compensation. The isostatic equilibrium of these special cases will be discussed separately under the subheading pertaining to each feature. The region approximately enclosed by 50° and 510 N lat. and 133° and 137° W long, has an average free-air anomaly of about -20 mgal. Whether this negative region reflects a deviation from isostatic equilibrium either as unusual crustal or subcrustalmass distributions, or insufficient or incorrect data is not known.This negative tendency is also discussed in association witha postulated fracture zone near 50. 50 N lat.

Fjords of the British Columbia Coast

Dehlinger etal, (1966b) described gravity measurements along the fjords of the Inside Passage of Alaska and Canada. Themeasure- rnents indicated free-air gravity anomalies increasing negatively toward the continent.The negative increase was tentatively attributed to increasing elevations in terrain from the coastline toward the cordillera.Additional measurements suggest that the free-air gravity contours parallel the sides of the fjords, i.e. follow the topography rather than cross them as suggested by the publishedmap. This is 55 illustrated by the contours in Portland Inlet and Douglas Channel. Although the free-air anomalies rapidly becomemore negative toward the continent, isostatic reductions by Heiskanen (1938) and Banks (1969) indicate that isostatic anomalies show only small negative values along the fjords of British Columbia.This indicates that the

previously glacially loaded fjords--now drowned valleys--arenow near isostatic equilibrium and any residual isostatic anomaly is probably local to the particular fjord. The small negative isostatic anomalies near the heads of the fjords may be associated with the very recent removal of the glacier ice near its most recent ice thickness es. The trend of the negative free-air anomalies originating in fjords can be traced across Hecate Strait and Queen Charlotte Sound to the edge of the continental shelf and suggests eitheran extension of the glacial channels or a structural continuity.

Gravity Anomaly Trends in Dixon Entrance, Hecate Strait and Queen Charlotte Sound

Figure 7 shows a negative free-air gravity anomaly beginning in

Clarence Strait (550 N lat.,13].°50' W long.) which trends southward, turning southeast off the northeast tip of Graham Island.From this poinl it parallels the coast to the vicinity of 530 N lat. 130020I W long., where it turns again nearly straight south passingover the continental 56

shelf at 51° 30' N lat., 130° 30' W long.This feature is reflected in the bathymetry of the area as a troughor bathymetric depression between Clarence Strait and the northeast tip of Graham Island.Else- where the limited bathymetry of the area suggests that it is sediment filled with only small local depressions to hint at its location until it nears the edge of the continental shelf. This feature, originating in a fjord and passing diagonallyacross the continental shelf, suggests that, rather thana synclinal depression as observed along the continental shelves west of Washington and Oregon (described below) it reflects a glacialscour channel or glacial valley now filled with sediment composed largely of glacial till.Den- sities ranging from 1.5 to 2.5 gm/cm3 might be expected for

Pleistocene to Recent glacial till under the shallow water of thecon- tinental shelf.If a density of 2.0 gm/cm3 is assumed for this glacial till and a density of 2. 7 gm/cm3 for the basement rock, thenegative gravity anomalies indicate a fill thickness ranging from 300 to 1000 meters. Higher densities would give a correspondingly thicker fill of sediment and lower densities a thinner fill.Maximum depths of fill might reach 3000 m.

A gravity low can also be observed beginning in thePortland Inlet turning southeastward along the Grenville and Principlechannels and joining with the negative gravity anomaly of the Douglas Channel. The Douglas Channel gravity low thenpasses south-southwestward 57 across Hecate Strait to the continental shelf terminating at the same point as the Clarence Strait low. An additional negative gravity alignment can be noted beginning in the Fisher-Burke channels, crossing Fitzhugh and Queen Charlotte sounds and joining the previously described anomalies.The trend of these joined gravity anomalies passes across the edge of the continental shelf at approximately 510301 N lat. and 130°30' W long.These channels are also partly reflected in the bathymetry and apparently partially concealed with sediment fill of similar thicknesses. If these negative gravity lows are in fact caused by filled relic glacial valleys, they indicate that valley glaciers in this region passed southeast along Dixon Entrance and Hecate Strait and westward across Queen Charlotte Sound, coalescing at a common point on the continental shelf southeast of Moresby Island near the head of the Moresby sea channel. A low ridge suggested by the positive gravity anomaly between the northwest tip of Graham Island and Dali Island, Alaska may have partially blocked access to the sea through the Dixon Entrance during the glacial epoch. These negative gravity anomalies and possibly filled glacial valleys may also be partially structurally controlled.Tectonic structures and known faults of this region parallel these gravity lows over a considerable portion of their length.Further, catalogued earthquake epicenters of this area locate along or near the axis of these gravity lows.The small magnitude of the gravity anomalies suggests the earthquake activity is more likely to be due to continuing tectonism rather than to isostatic adjustment.

The Continental Shelf Off Washington and Vancouver Island

The area of the continental shelf west of Vancouver Island exhibits negative free-air anomalies of approximately -50 mgal southwest of Nootka Island, nearly -70 mgal west of Cape Flattery, and approximately -60 mgal west of Grays Harbor and Willapa Bay. These anomalies are large, generally elongate, and trend parallel to the shoreline.Smaller negative and positive anomalies occur between the three larger anomalies. North of Cape Cook the anomalies of the continental shelf are obscured by the extensive Scott Island fracture zone. South of the Columbia River the character of the gravity anomalies on the shelf changes. A free-air gravity map, prepared by Dehlinger etal. (1968) for the area west of Oregon shows a considerable number of small irregular positive and negative anomalies along the continental shelf.The magnetic anomaly maps of the continental shelf west of Oregon also show many small 'irregular anomalies with considerable magnetic relief (Ernilia, etal. ,1968a). By comparison, the gravity and magnetic anomalies of the continental shelf west of Washington and Vancouver Island are of strikingly different character.They show fewer, more regular anomalies of greater areal extent but of comparable magnitude. The negative gravity anomalies on the shelf off Washington trend toward the Olympic Mountain negative Bouguer anomaly and may be genetically related.The Olympic Mountain anomaly, however, trends northwest-southeast while the associated shelf anomalies turn, near the coastline, to a north-south alignment. The negative anomaly west of Cape Flattery is at least partly due to the bathymetric relief of Astoria Canyon. The negative gravity anomaly along the continental shelf off Vancouver Island appears confined to the shelf region and may represent a sediment-filled synclinal depression. A small gravity high superposed on the regional low west of the Olympic peninsula extends from the- coast of Washington near 43.5° N latitude west then northwestward to the closed gravity high off Vancouver Island near Barkley Sound.Figure 4 indicates a positive magnetic anomaly of similar areal extent which occurs in the same location.The small gravity high joins a similar gravity high on land which appears coincidental with the outcropping of Eocene volcanics about the perimeter of the Olympic Mountains. The gravity and magnetic anomalies suggest that the Eocene volcanics observable on the Olympic peninsula may extend westward beneath the continental shelf as far as the continental slope off northern Washington and southern Vancouver Island. The most striking gravity anomaly along the Washington continental shelf is the - 60 mgal linear minimum west of Grays Harbor and Willapa Bay.This anomaly appears to terminate at Astoria Canyon or suffer a right lateral offset of about 15 km across the canyon. Two possible explanations of this are:(1) the sediment-filled synclinal depression represented by the free-air gravity anomaly is actually offset by a transcurrent fault along the Astoria Canyon- Columbia River lineation, or (2) the gravity low south of Astoria Can- yon has had an apparent westward displacement due tothe upwarping of Tertiary volcanics which, according to Snavely and Wagner (1964), form a northward plunging anticlinorium in northwest Oregon.In this latter explanation the seaward displacement of the anomaly is due to tilting of the continental shelf and the residual sharp east-west anomaly structure over Astoria Canyon is due primarily to the bathymetry of the canyon. Carison (1968) suggested that step-like offsets of the Astoria Canyon indicate that it may have been cut along a zone of structural weakness. Minor faults and anticlines having strikes parallel with the lower Columbia River are observed in northwestern Oregon.

The Randsenken Effect Along the Continental Margin of Washington and British Columbia

Isostasy implies that thick blocks of continental crust sink deeper into the more dense subcrustal material than do thinner blocks 61 of oceanic crust.Therefore, when continental and oceanic crustal blocks are juxtaposed, the edge of the thick continental block pro- jecting deeper into the substratum than the oceanic block produces a step in the crust-subcrust boundary.This Randsenken (sunken-edge) of the continent produces a negative free-air anomaly parallel to and near the base of the continental slope and a positive anomaly parallel to and generally over the landward side of the edge of the continental shelf.The amplitude of the positive anomaly is somewhat reduced by the reverse Randsenken effect of the deep water layer oceanward of the continental slope. The Randsenken along a coastline such as the Pacific coast of North America generally produces a negative free- air gravity anomaly of -60 to -100 mgal. The larger negative anomalies lie along transition zones having the steeper continental slopes.Typical negative free-air gravity anomalies along the west coast of North America due to the Randsenken effect are: Southeast Alaska -80 mgal, Queen Charlotte Island -80 to -100 mgal, Washington -20 mgal, Oregon -60 to -100 mgal, northern California -60 mgal to -70 mgal and Baja California -85 mgal. These data are taken from the free-air anomalies outlined in this thesis and values near the base of the continental slope in the northeast Pacific Ocean given by Worzel (1965), Vening Meinesz (Heiskanen, 1938),

Dehlinger et al.(1 967a), and Dehlinger et al.(1 967b). Figure 7 shows the negative free-air anomaly due to the 62 Randsenken of the continent from the Dixon Entrance to Cape Cook. The continental slope in this region exhibits a very steep gradient and consequently, generates a large Randsenken effect gravity anomaly.The negative free-air anomaly is exaggerated west of the Scott Is lands due to the superposition of a fracture zone mass deficit on top of the expected Randsenken effect. Strikingly, the negative gravity anomaly of the continental Randsenken is exceptionally small west of southern Vancouver Island and Washington.Bathymetric maps of this region, prepared by McManus (1964), show a continental slope having a shallow two step gradient; however, the bathymetric relief is insufficient to account for the near absence of a negative free-air anomaly.The gravity anomaly along the continental slope of Washington and southern Vancouver Island suggests that either the transition from oceanic to continental crust is more gradual than expected, or that the crust of the continental block is thinner than expected. The Randsenken effect or the gravity anomaly observed by Worzel (1965) along the east coast of North America is approximately half as large as observed along the west coast of North America. This difference implies that either the transition from continental to oceanic crust occurs more abruptly along the west than the east coast or that the density contrast between continental and oceanic rocks, including rocks of the upper mantle, is greater along the west coast than the east coast. 63 The Queen Charlotte Trough

Menard and Dietz (1951) outlined, in their early study of the submarine geology of the Gulf of Alaska, an irregular elongated trough along the west side of the Queen Charlotte Islands.St. Amand (1957) states that the continental shelf is narrow in this region, with the 100 fathom contour only about 1 mile offshore.He also postulates on the basis of the narrow shelf and straightness of the western shore that a fault defines the oceanward shoreline of the Queen Charlotte Islands.Hurley (1960), in his study of the geomorphology of the northeast Pacific, includes fathograms of the continental shelf and slope and the abyssal plain west southwest from Dixon Entrance and vest-northwest of Graham Island.These fathograms indicate a steep continental slope terminating in a sharp scarp at the abyssal plain.

Hurley does not show a trough west of the Queen Charlotte Islandson his physiographic map.Gibson (1960) presented a bathymetric traverse west of Moresby Island which shows a complex benchon the continental slope and a trough or depression at the base of the slope. Gibson's map "The Submarine Topography in the Gulf of Alaska" out- lines a "Queen Charlotte Trough" that is approximately 300 km long, with an axis approximately 55 km west of the shoreline of the Queen Charlotte Islands,The U. S. Hydrographic Office map (1952) of this area shows only a small depression west of Moresby Island,The 64 configuration or existence of 'The Queen Charlotte Trought still remains somewhat questionable at this writing.For this reason it is difficult to compare the bathymetry and free-air gravity anomalies in this region. The gravity map indicates a gravity low approximately 150 km in length west and northwest of Graham Island, a more complex gravity low west of Moresby Island, and a small basin shaped low southwest of the south end of the Queen Charlotte Islands.These last two lows are separated by a gravity high which suggests two seamount structures connected by a ridge.Gravity gradients along the steep continental slope west of Moresby Island reach 20 mgal/km. A large part of the negative gravity anomalies west of the extremely narrow continental shelf, in the area about the Queen Charlotte Islands, can be attributed to the Randenken of the continent, indicating that this area is essentially in isostatic equilibrium and suggesting that a very deep sediment filled trench is unlikely.How- ever, the basinal structure suggested by the gravity contours southwest of the south end of the Queen Charlotte Islands has a shallower depth and a less negative gravity anomaly than the gravity minimum to the north suggesting a possible ponding of 1 to 3 kilometers of glacially derived sediment in this basin.The glacial silt streaming down the continental slope near 510301 N lat. and l30°30'W long, was apparently blocked in its northerly flow by the hills and ridge outlined 65 by the gravity high separating the two gravity minima.

A Postulated East-West Fracture Zone Near 510 N Lat.

Menard (1955b) postulated the existence of an eastwest fracture zone about 600 miles north of the Mendocino fracture zone.This was based on the observed regularity in spacing of the then known E-W trending fracture zones which segment the eastern Pacific Ocean. Hurley (1960) discussed this fracture zone further in his study of the geomorphology of the northeast Pacific Ocean, suggesting that Peters Ridge is near the expected position and has the proper trend.He also suggested this trend disappears beneath the Queen Charlotte fan to the east.This east-west lineation is in alignment with the westward set of the north tip of Vancouver Island.This set is visible also in the continental shelf west of the Scott Islands and is emphasized by the gravity anomaly contours of the same area.The gravity anomaly contours between 50° and 51° N lat., however, do not show a pattern which would be expected over a sediment-buried ridge, escarpment or trough, such as those associated with known fracture zones of the eastern Pacific Ocean. The gravity data between 50° N and 51° N lat. and west of 132° W long, have greater uncertainty than the data between 132° W long. and the coast of Vancouver Island and may be of insufficientaccuracy to discern a small fracture. Dehlinger etal. (1967a),in a study of the Mendocino Escarp- ment, have shown that in addition to the variation of gravity due to the bathymetric relief, there is also a part of the anomaly over the fracture zone which is due to density-mass differences in the upper mantle.These differences are reflected in the free-air anomaly map as a predominantly negative region south of the escarpment and zero or slightly positive anomalies north of the escarpment.In the vicinity of the postulated fracture zone between 500 N and 51° N lat. and 1320 to 137° W long, the anomalies are predominantly negative while to the north and west of this region the anomalies are near zero or positive.While the data are inconclusive they do suggest the pos- sibility of variations in distribution of mantle material similar to that hypothesized beneath the Mendocino fracture zone, If the Scott Island fracture zone as outlined by the 100 mgal contour is interpreted as at least partly due to tension, then the expected motion along an east-west fault located just south of 51° N lat. would be right lateral contrary to that which would be expected from the observed hboffsetlt of the continental shelf west of the Scott Islands. The gravity data do not generally substantiate the existence of the postulated fracture but are inconclusive particularly if the fracture zone has its eastern terminus west of 132° W longitude. 67 The Scott Islands Fracture Zone

Benioff (1962) postulated a dextral transcurrent fault west- southwest of the coasts of Alaska and British Columbia, extending from approximately 48° N lat., 127° 30' W long, along a great circle to 60° N lat., 140° 30' W long.Wilson termed this the Queen Charlotte Islands Fault, suggesting that t is a transform fault rather than a transcurrent fault.He also relocated the southern end of this fault near 48° N lat.,128°20' W long.McManus (l967b), from a study of the bathymetry, described the area between Cobb Rise and the northern end of Vancouver Island as consisting of fracture zone topography. The bathymetric map of the region (Figure 8) shows long narrow ridges (50°lO' N lat.,129°10' W long.), trending northwest, with a relatively smooth floor between- -suggestive of sediment-filled troughs. Along a coast with a relatively narrow continental shelf and steep continental slope approximately 80 to 100 mgal of the observed negative free-air gravity anomaly over the ocean may be attributed to the Randsenken of the continent.For this reason the -100 mgal closed contour will be used to define the area of the Scott Islands fracture zone,This fracture zone actually lies partly on the continental shelf and slope between the Scott Islands and Quatsino Sound; and, therefore, the -100 mgal contour line is probably a conservative estimate of the actual areal size of thj.s tectonic feature. This structural feature of the earth's crust defined by a large negative free-air gravity anomaly was first noted on a traverse by the R/V YAQUINA in August, 1965.The traverse passed in a north- easterly direction between the Scott Islands and the northern tip of Vancouver Island nearly over the region of the greatest negative anomaly.Because this feature was first noted off the Scott Islands and is nearly centrally located with respect to the islands, the Scott Islands fracture zone was adopted as a best descriptor. The Queen Charlotte fault as described by Benioff (1962) and Wilson (1965b) actually lies along the extreme southwest edge of this fracture zone.Further, the areal gravity outline is not what would generally be expected for a trans current fault, suggesting that, although this fracture zone is probably associated with the Queen Charlotte fault, it is also an entity in itself. This feature is approximately 35 km in width, at its widest point, and 175 km long.It has the general appearance of a complex graben or a group of down-faulted blocks which, after normal faulting, have been subjected to right lateral shear.If in fact the maximum gravity lows, indicated by small closed hatchured contours, can be assumed to have been aligned originally along a northeast trend, a 30 km right lateral offset can be postulated between the gravity low associated with the north end of the Explorer trough and the central anomalies of the Scott Islands fracture zone. If this anomaly is interpreted as a graben, it is most likely filled with glacial till ubiquitous to this region.Assuming a density contrast of 0. 3gms/cm3between upper crust and loose till (2.5-2.2) and between lower crust and more consolidated till (2. 7-2. 4) maximum fill depths of 4 to 6 km might be expected.

Seamounts

Cobb Rise and the adjacent area toward the northwest contain numerous seamounts (Figure 1).These seamounts seldom occur as solitary prominences but rather as contiguous groups with a general northwest trend.The method of determining the free-air gravity anomaly value above the peaks of the seamounts and the average seamount densities is described in Appendix 4.The location and dimensions of 44 seamounts occurring in the area of the free-air gravity anomalies of Figures 6 and 7 are given in Table 6.Figures 6 and 7 show that a negative free-air gravity anomaly occurs near the base around solitary seamounts and along adjoined seamounts. The larger negative free-air anomalies are most prominent along the southwest and northeast sides of seamounts even near solitary, symmetrical seamounts such as Union searnount (49°32. 5' N lat., 132 041. 6' W long.).This indicates that the substructural trend is also predominantly northwest-southeast. 70 The negative anomaly surrounding the central positive peak anomaly suggests that the seamounts in this regionare in or nearly in isostatic equilibrium. The average ratio of thearea of compensation, defined by a circle about the seamount axis approximately enclosing the negative ttmoat" anomaly, and the area of the load, defined by a circle approximately enclosing the positive peak anomaly, is 7.That is the weight of the conical seamount, sustained by the crust, is distributed over an area approximately 7 times thearea of the base of the seamount.Load and compensation areas for 10 seamounts are given in Table 2. A comparison of the area of load, defined by the positive free- air gravity anomaly and the area of load defined by the physical base of the seamount shown on bathymetric charts, shows theyare, within the accuracy of measurement, very nearly identical in the region

south of 50° N lat.However, north of 50° N late the load area defined by gravity is approximately 5 times the loadarea defined by bathymetry.Since the structure of seamounts above the ocean floor is probably similar, this difference in loadareas suggests that differences in structure or constituents of the crustor upper mantle occur between these two areas.In this respect McManus (1965) noted that the seamounts of the Alaskan Abyssal Plainappear to occur on rises.

If it is assumed that the source of magma to forma seamount is Table 2. No. Seamount load and compensation parameters Physical Dimensions v Gravity Dimensions Table 3 53 (km)ra9.9 2.02(km) h (km3) 216 (km)rp8.6 34.2(km)r (r/r) 16.0 2 (m) 43h (mgal) 4.1 251514 6 18.4 8.3 1.673.062.762.73 1082 700977119966 14.319.218.5 7.8 47.830.621.438.5 4.67.44.06.6 169240 9860 15.922.8 5.89.3 40383728 17.511.016.516.2 7.4 1.912.951.832.852.55 1080 937815233 19.335.612.813.526.4 100.5 40.747.734.238.5 4.58.08.13.37.1 151160109 2124 19.415.310.4 2.12.3 = radius of seamount base (km) Average 7 108 10 vhra = volumeheight of of seamount seamount (km) (km3) rnr p = radius of negativepositive F.F. A.A. anomalyanomaly (km)(km) x E. g = (V/p r2) (2. 4/3. 3) yL p t h = . 095 h (mgal) (m) I- 72 a subcrustal layer beneath the area of compensation, the thickness of this layer may be determined from the volume of the seamount obtained from bathymetry and the area of compensation.Since the source material of the effusive has a density near 3. 3gm/cm3and the resulting seamount a density near 2. 4 gms/cm a proportional adjustment of the source layer thickness is made. The range of thickness of such a layer is 20 to 170 meters with an average of 108 meters.In transferring the material of such a layer to the surface as a seamount, a negative free-air anomaly would be formed in the area of compensation.This anomaly would range from 2 to 23 mgal with an average of 10 mgal.This figure is in good agreement with the average anomaly observed about seamounts.It is not possible by this means to say whether the material in the layer beneath the area of compensation has actually erupted on the surface or whether it has been displaced horizontally out of the area of compensation by a surface load derived from a deeper source.It does, however, indicate that isostatic equilibrium is achieved and that the weight of the searnount is not likely sustained by the elastic strength of the crust.

The Explorer Trough

Gibson (1960) depicts the Explorer trough on his map of the submarine topography in the Gulf of Alaska as a 130-mile long, 5-10- 73 mile wide depression extending from approximately 48°20' N lat., 131° W long., to 50°20' N lat., 129°40' W long.This topographic feature is shown as a nearly linear feature trending N 15° E on the longer southern and central portion and bending to N 500 E at the northern end.Gibson presents cross-sections of the Explorer trough suggesting they have the form of a thrust fault with a raised escarpment on the eastern edge and favors a tectonic origin for this topographic feature. Figure 8 shows a portion of a bathymetric map by McManus (1964) which portrays the topography about the Explorer trough.This map shows the location of the Explorer seamount (49°05' N lat. 130°45' W long..) andtheSethinole seamount (49°45' N lat. ,129°50' W long.) with respect to the Explorer trough and indicates that the northern end of the trough (50°05' N lat. ,129° 45' W long. ) terminates abruptly against a long narrow northwest-southeast trending ridge. The free-air gravity contours outlining the Explorer trough show a pattern similar to the bathymetric maps of Gibson (1960) and McManus (1964); however, the anomaly contours appear more sinuous and irregular.The southern end of the trench encircles the Explorer seamount anomaly in the manner of a moat typical of large seamounts. The central portion of the negative free-air gravity anomaly is elongate and narrow with the axis of the -50 mgal anomaly displaced to the northwest of the topographic axis.The magnitude 74

132° 1310 1300 129° I28° 51° z:,i

500 500

490 49°

4R° 480 132° 1310 130° 129° 128°

Figure 8.Bathymetric chart of the Explorer trough and Scott Islands fracture zone (after McManus, 1964) showing location of profiles S-S', X-XT, Y-Y', and Z-Z'. 75 and shape of the gravity anomaly suggest that the central portion of the trough is partly filled with sediment. The free-air gravity anomalies over the northeast end of the Explorer trough show two -50 mgal depressions. The northeast gravity anomaly low is directly correlative with the bathymetry; however, the gravity anomaly approximately 20 km to the south has no apparent topographic expression implying a total filling by approximately 1.5 km of sediment of a pre-existing basinal or graben structure. Two possible explanations of the plan view character of the Explorer trough reflected by the sinuous free-air gravity anomalies are:1) the trough is a series of circumstantially interconnected moats about the Explorer seamount, Seminole seamount and the end of the Dellwood seamount range; or 2) the trough originallywas a more linear feature, formed by tectonic activity, which has been reshaped in the south and central areas by volcanic flows from adjacent seamounts and in the north end by transcurrent faulting normal to the trough axis.In the first case the trough postdates the seamounts and in the second case the trough pre-dates the seamounts. Figures 9, 10 and 11 show three selected gravity, magnetic and bathymetric profiles across and approximately orthogonal to the Explorer trough.Figure 8 shows the location of the three profiles: Figure 9 along line X-X, Figure 10 along line Y-Y' and Figure 11 along Z-Z'. 7£) ALONG LINE XX'

568

60 566 a50 40 564 30 20 562 10 > 0 560 -10 -20 558 -30 -40 556 -50 -J Ii- -60 554

552 GRAVITY

1.5

2.0

2.5

I- 3.0

3.5

0 5 10 15 20 25 30 35 40 45 50 DISTANCE (km)

Figure 9.Bathymetric, gravity and magnetic profile across Explorer trough along X-X'. 77

ALONG LINE V-V

560

20 559

I0 558

557 -

556

-20 555

554

553 ;

-50 552

GRAVITY

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0 5 10 15 20 25 30

0/STANCE (km)

Figure 10.Bathymetric, gravity and magnetic profile across Explorer trough along Y-Y'. 78

ALONG LINE Z - Z' 560

\ E / 558 20 1 0 / 556 I0 /'Y 1 0 // 1'N/-'\\\j ( 552 () 550 ::k 548 -30 MAGNETICS OBSERVED GRAVITY 546 2 COMPUTED GRAVITY -40

2.5

p 03

S I \ p2.I 3.5 / - E \

\ / / / I. \/

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/ 50 I I I I I I I I 0 0 20 30 40 50 60 70 80 90 100 110 120

DISTANCE (k,n)

Figure 11.Bathymetric, gravity and magnetic profileacross Explorer trough along Z-Z'. 79 The bathymetric profile over the northeast end of the Explorer trough along line X-X' indicates a 3. 5 km deep trough or trench with steep scarp-like walls.Benches occur on the northwest wall.Scarps of the northwest wall appear to have slopes of 600 or greater while the southeast wall slope is near 30°.The angle of the walls or scarps suggests gravity faulting along the northwest wall and low angle normal faulting along the southeast wall possibly indica- tive of tensional forces normal to the trough axis.Several benches occur on the northwest wall which exhibit little or no sediment cover. Projection of the wall slopes downward to intersection suggests that sediment fill in the trough axis is 0. 1 km or less.No significant amount of sediment cover or fill is detectable on the ridge structures which flank the northern end of the Explorer trough. PDR records from approximately 40 to 50 km southeast of the trench axis show ponded sediment which is severely tilted. The bathymetry along profile Y-Y' near the mid-length of the Explorer trough shows a relatively small trough approximately 7 km wide and 300 m deep.The sediment fill in this portion of the Explorer trough, estimated from gravity measurements, is 0. 5 to 1 km. The topography along line Z-Z', indicated on precision depth records, shows a basinal feature approximately 60 km wide and 300 to 400 m deep in the center.Free-air gravity anomalies show a sharp 20 mgal negative depression over the axis of the topographic depression suggestive of a sediment-filled trough. To estimate the thickness of the sediment fill (or the actual trough depth) two subsurface 2 dimensional structures were assumed and the gravity anomaly calculated for each. A density of 2. 1gm/cm3was used for the sediment and a density of 2. 9gm/cm3for the underlying basement rock. The structure, indicated by the open triangles connected by dashed lines, produces the gravity anomaly shown by the open triangles.Although the calculated gravity anomaly, shown by the open traingles, does not coincide with the details of the observed gravity anomaly (shown by the solid 1ine)it does exhibit the same amplitude.The model structure shows an estimated minimum anticipated sediment thickness of 0. 6 km in the filled trough. The structure indicated by the solid circles connected by a solid line produces the gravity anomaly shown by the solid circles. The gravity anomaly produced by this structure coincides well with the observed gravity anomaly and the model structure shows an estimated maximum sediment thickness or trough depth of 1.7 km. The effect of two-dimensional models and the high density contrast (2.9gm/cm3- 2. 1gm/cm3= 0.8gm/cm3)cause these calculations to produce conservative estimates of sediment fill thicknesses.The sediment thickness and consequently the actual EiJ relief of the Explorer trough south of the Explorer seamount is therefore probably near 1.5 km. Magnetic measurements were made concurrently with the gravity measurements along profiles X-X', Y-Y' and Z-Z'. The total magnetic intensity exhibits a relief of 1 600 gammas over the northeastern end of the trough, 800 gammas in the central section and a relief of approximately 500 gammas over the southwestern end of the trough near the base of the Explorer seamount. The gradient of the regional magnetic field is approximately perpendicular to the magnetic profiles.Therefore, removal of the regional field to produce a magnetic anomaly causes little change in the magnetic profile except to change the reference level.Figures 9, 10 and 11 show the magnetic anomalies are strongly positive over the troughand the axis of the magnetic anomaly is displaced towards the southeast with respect to the trough axis.The magnetic anomalies are very nearly centered on the southeast scarp of the trough.Each magnetic profile shows a dip or decrease of 100 to 200 gammas near the magnetic axis and also a greater amplitude of the southeast edge of the positive anomaly than the northwest edge. The amplitudes of the gravity and magnetic anomalies increase along the strike of the Explorer trough from southwest to northeast whereas the sediment thickness in the trough decreases from approximately 1.5 kilometers to 0.1 kilometers.Maps of the free-air 82 gravity anomalies (Figure 6), magnetic anomalies (Figure 4) and the bathymetry (Figure 8) suggest that the changes in anomaly values and sediment thickness may not be gradual but rather that the Explorer trough may be segmented. The absence of significant amounts of sediment in the northeast end of the trough suggests recent deforma- tionand earthquake epicenters about the immediate trough area indicate continuing tectonism whereas the trough south and east of the Explorer searnount appears vestigial.

CascadiaAbyssal Plain

The contours of the free-air gravity anomalies in the area designated by Hurley (1960) as the Gas cadia Abyssal Plain are small in amplitude and areal size and exhibit relatively smooth closed out- lines.The half wave length of the free-air gravity anomalies along the survey ships track line indicate that a number of these anomalies have smaller dimensions than the survey grid.For this reason additional small anomalies can be anticipated in the regions about

48° N lat., 128° W long, and 46° N 120° W long.A1so it is quite possible that with additional measurements the anomalies

indicated in this region may subdivide into many smallerones. These small anomalies suggest the presence of sediment-covered hills of low relief.The magnitude of the gravity anomalies, as sum- ing a density contrast between sediments and abyssal hills of 0. 6 to 0.7gms/cm3,indicates that these abyssal hills may exhibit an undulat- ing relief of 300 to 400 meters which is completely covered by sediment.Examination of PDR records obtained across the Cascadia Abyssal Plain shows the tops of a number of hills protruding through the sediment (Kulm, 1967), particularly near the western flank of the Gas cadia Abyssal Plain where the sediment thins. A seismic reflection profile across the north end of the Gas cadia

Plain (Ewing etal., 1968) indicates that the negative free-air anomaly near 48.8° N lat., 128.8° W long, is due to a graben.The graben trends about N 30° E, generally parallel to Explorer ridge and trough, and is partially filled with sediment 0.5 to 1. 2 km thick. Earthquake epicenters (Tobin and Sykes, 1965) in the vicinity of the grab en indicate continuing tectonism.

The thickness of sediment reported (Shor etal., 1968; Hamilton, 1967; Duncan, 1968) in the Cascadia Abyssal Plain is sufficient to produce a free-air anomaly of 10to60 mgal when superimposed on an isostatically-balanced crust.An observed zero average anomaly over the Cascadia Abyssal Plain strongly suggests the region has subsided to accommodate the additional load.This subsidence would be proportional to the sediment thickness and of the order of 380 meters near the center of the Gas cadia Abyssal Plain. CRUSTAL AND SUBCRUSTALCROSS SECTIONS ACROSS THE CONTINENTAL MARGINS OFWASHINGTON AND BRITISH COLUMBIA

Construction of the Crustal andSubcrustal Cross Sections

Figure 12 shows the locationof five 2-dimensional crustaland

subcrustal sectionsacross the continental margins of Washingtonand British Columbia whichwere constructed to study the structural transition from ocean to continent.Each of the sections is approxi-

mately 800 to 1000 km in length.They begin in the Tuftsand Alaskan abyssal plains and extendeastward to the continentalinterior east of the coastranges. A 50 km vertical dimensionwas arbitrarily selected for thecross sections.On the basis of seismicrefraction investigations (e.g. James andSteinhart, 1966), isostasy(e.g. Heiskanen and Vening Meinesz,1958), previous studies(e.g. Worzel, 1965) of continentalmargins, and thewave lengths of the observed gravity anomalies, majorchanges in structure in thezone of transition were expected tooccur within 50 km of the earth'ssurface. Land elevations on thecontinents were obtainedfrom topographic maps of the US Geological Surveyand the DominionObservatory of Canada.Free-air gravity anomaliesat sea are basedon measurements made with gravitymeter S-9 along track lineswhich very closely approximate thesurface trace of thecross sections. Water 140 138 136 rfl ,.. ,,. ,',n I1f 56 54, ¶I 54 52' ______- 52

Uico depths were determined from PDR measurements obtained simultaneously with the gravity measurements. Depths to layers in the cross sections were obtained from seismic refraction measurements.The location of the seismic surveys are shown in Figure 2. Seismic refraction results were extrapolated laterally along isobaths when they did not coincide exactly with the section profiles.The topographic, bathymetric and seismic refraction results were scaled onto working drawings and then connected by straight lines so that the scaled cross section consisted of joined polygons.The cross sections were effectively extended 5000 km in each direction to prevent end effects in the final computation. The observed seismic velocities were converted to densities using the relation shown in Figure 13.This empirical relation, established by Nafe and Drake (1961) is based on:1) measurements on sediments and sedimentary rocks (V< 6 km/sec), 2) Birch's measurements at 10 Kbar for various igneous rocks, and 3) the Bullen A model of the mantle (Nafe, 1967),The Nafe-Drake relation between compressional wave velocity and rock density is generally considered more reliable in sedimentary and upper crustal rocks than in materials of the mantle.The densities obtained from the seismic wave velocities using the Nafe-Drake curves were assigned to the appropriate polygons of cross section. 87

_ff4.J rai LSI -II.i±IIIi

7 -- -H-

i;= 5

---H-i-----

2 EEEEEEEEEEEE

:1.] 2 3 4 Figure 13.Graph of the Nafe-Drake relation between cornpres- sional wave velocity and density (Nafe and Drake, 1961). The vertical component of gravitational acceleration at points along the earth's surface due to the sw-n of the polygons of a cross section was computed using the line integral method as applied by Talwani et al.(1959b) (see Appendix 3).The computed gravity was compared with observed gravity, and iterative adjustments of the polygons were made until observed and computed gravity values coincided.

A common crustal and subcrustal section 50 km in depth,was adopted for the ocean end of the five cross sections.The common section is a computational reference which permits the five cross sections to be compared in three dimensions. The section, which is

not shown on the cross sections, consists of five layers:1) a 3. 7 km water layer of density 1.03 gm/cm3,2) a 0.5 km sedimentary layer of density 2.15 gm/cm3,3) aO.8 km upper crustal layer of density 2.74 gm/cm3,4) a 5.0 km lower crustal layer of density 3.00 gm/cm3,and 5) a 40.0 km mantle layer of density 3.30 gm/cm3 A free-air gravity anomaly of zero is assumedover the section.

Sources of Error in the Cross Sections

Bouguer gravity anomalies were used for structural control over land.The Bouguer correction assumes a density of 2. 67 gm/cm3which may differ considerably from the true density insome places.Walcott (1967), in constructing a Bouguer anomalymap of Vancouver Island, estimates the anomaly error may beas large as 6 mgal. Sea gravity measurements have an rms uncertainty of approximately SmgaJ.. Changes in gravity of approximately S nigal in the Bouguer and free-air anomalies would cause less than 0. 5 km vertical change in a Moho depth of the cross section, Shor (1962) indicated the accuracy of seismic refraction velocities is difficult to assess and attributes an "error of determination' of 0. 05 km/sec to good stations.An error of 0. 05 km/sec in the velocity determination will result in an error of 0. 2 km in layer depth of a deep-sea profile and about 0. 5 km in a shelf profile (Shor, 1962).Errors in refraction studies may also occur due to uncertainties in ship position particularly in areas of large structural gradients.Shor etal, (1968) estimates an uncertainty of ± 5 km in position in refraction studies off Oregon. The velocity-density relation of Nafe and Drake (1961) is an average fit to many rock types.The composition of mantle rocks is unknown; and rock types postulated for the mantle,which have been measured in the laboratory under conditions believed to exist in the upper mantle, indicate that for a velocity of 8 km/sec densities may range from 3. 0gm/cm3to 3. 7gm/cm3(e. g. W o olla r d,

1 9 6 2).The empirical relation between crustal thickness indicated by seismic measurements and changes in surface elevation suggest mantle densities are between approximately 3. 28gm/cm3 and 3.46gm/cm3(Woollard, 1962).Clearly the thicknesses and densities of upper mantle materials can be varied to conform with the observed gravity anomalies.However, seismic refraction control and a small range of reasonable densities in the crust restrict acceptable layering for the models. A uniform mantle is assumed below 50 km.If, however, differences in the upper mantle beneath oceans and continents extend greater than 50 km as has been suggested as possible by MacDonald (1963), lateral density contrasts in the upper mantle would be reduced.

The Central Washington Cross Section

The central Washington crustal and subcrustal section BB! is shown in Figure 14.The western end of the central Washington section is located in Tufts Abyssal Plain, near 45. 3° N lat., 134.2° W long. The cross section strikes N 74° E cutting obliquely across Juan de Fuca ridge and the Cascadia Abyssal Plain. The cross section is normal to the trend of the continental slope and shelf and crosses the coast north of Gray's Harbor, Washington. It traverses Puget Sound near Seattle and has its eastern terminus in the foothills of the Cascade Range east of Everett, Washington near 48.0° N lat., 120. 7° W long. The short wave length gravity anomalies over Juan de Fuca Ridge are due primarily to the effects of bottom topography. The 100 50 q(.0-J -50 I-. -150-100 0 (0 5 20 15 330 / 327 /i '\ \\\ 327 502 295 20 ci. 25 25 0I- Lu S Lu 3530 CRUST A CENTRAL WASHINGTON SUBCRUSTAL CROSS SECTION BB 332 3035 40 i/ / \\\ FREE-AIR AND BOUGUER GRAVITY ANOMALY PROFILE COWIJTEDOBSERVED GRAVITY SRAVITY0000 - 40 45 SEISMIC REFRACT1OI LINES (AFTER SHOR H at, 19681 OREGONDENSITES STATE OF LAYERS UNIVERSITY IN GMiCM3 MARCH 1968 050100HORIZONTAL SCALE KILOMETERS 45 50 Figure 14. Central Washington crustal and subcrustal cross section BB'. 50 '.0 92 average free-air anomalies across the ridge and across the Cascadia Abyssal Plain, shown in Figure 14, are near zero, indicating that these two provinces are essentially in isostatic equilibrium.Adjacent to the ridge in Tufts Abyssal Plain, a small number of gravity measurements indicate a negative free-air anomaly.Beginning at approximately 127° W long, near the base of the continental slope gravity anomalies become very irregular and show a gradual decrease toward Puget Sound. Gravity data on land are Bouguer anomalies (Woollard and Rose, l963 uncorrected for topography; Sauers, 1966, unpublished).Gravity anomalies in the Puget Sound region are not consistent with values expected for the observed elevations.Danes étal. (1965) noted that Puget Sound region exhibits a regional gravity low in addition to individual large negative anomalies.Isostatic reductions by Heiskanen (1938) and Banks (1969) indicate a -100 mgal negative isostatic anomaly in the vicinity of Seattle.It may be concluded from these observations that the Puget Sound region is not, at least locally, in isostatic equilibrium.The free-air and Bouguer anomalies suggest that the region of isostatic disequilibrium may extend westward as far as the continental slope.

Refraction lines (Shor etal.,1968) across the ridge show the mantle to be as shallow as 7 km under the crest and to dipaway from the ridge to depths of about 10 km under the adjacent Cascadia and 93 Tufts plains.Velocities in the mantle are low on the ridge (7. 7 km/ sec or less) and slightly low (7.9 to 8.0 km/sec) under the adjacent plains. Crustal layering and mantle velocities in western Washington are not defined by seismic refraction data; however, they are in general agreement with the results of Tatel and Tuve (1955) who indicate a depth to the mantle of 19 km in Puget Sound and 30 km east of PugetSound. Crustal velocities beneath the Puget Sound region show large variations (Neuman, 1959).Mantle velocities are near 8.0 km/sec (Tatel and Tuve, 1955). Seismic refraction measurements indicated by heavy lines on the figure define the crustal layers of cross section BB' beneath Juan de Fuca Ridge and Gas cadia Abyssal Plain.The seismic ties along the continental shelf and slope were extrapolated a short distance along isobaths from measurements off the Oregon-Washington coast. The layers indicated by dashed lines are less reliable than those indicated by solid lines.The lines of parallel dashes outline the sides of polygons used in the computations and are used to suggest either a gradual density transition or an indefinite structural dimension. The upper curve in Figure 14 shows the gravity anomalies. The solid line represents observed free-air anomalies at sea and Bouguer anomalies on land, and the circles represent calculated anomalies. 94 The sections assume a constant density beneath an arbitrary depth of 50 km. For this condition a mantle density of 3. 32gm/cm3 beneath Washington and 3. 30gm/cm3beneath Tufts Abyssal Plain

i sconsistent with the observed gravity anomalies. A lower mantle density of approximately 3. 27gm/cm3appears to exist under Juan de Fuca Ridge and Cascadia Plain down to a depth of 50 km; a lower density is also possible in which case the mantle layer would be less thick. The gravity data require the presence of a still lower density material in the upper mantle beneath Juan de Fuca Ridge.The density and configuration of the low density mantle material beneath the ridge can be varied provided the gravity anomalies are satisfied. Dehlinger etal. (1968) also postulated a low density upper mantle beneath Juan de Fuca Ridge and similarly beneath the Gorda Ridge (Dehlinger etal., l967a). At the continental margin refraction results indicate the

Moho discontinuity dips steeply beneath the continent.Few data on crustal layers are available for Washington.Based on a mantle density of 3. 32 gm/cm the section indicates a depth to the Moho of 19 km in western Washington, 36 km beneath Puget Sound and approximately 34 km beneath the western flank of the Cascade Range. The 19 km depth west of Puget Sound and 34 km east of Puget Sound agree well with the seismic refraction results of Tatel 95 and Tuve (1955).Danes etal. (1965) attribute the large negative gravity anomaly in the vicinity of Seattle to a 10 km deep sedimentary basin.The cross section agrees in general with their interpretation and suggests displacements of the crust which form the sedimentary basin may extend to mantle depths. A large distortion in the sedimentary structures near 125° W long, is required by the gravity profile.The inferred structure may be interpreted as severe folding of the sedimentary layers or normal faulting of the layers with the west side down relative to the east side. Gravity data cannot distinguish between these two interpretations. Von Hueneetal. (1969) concluded, on the basis of seismic profiler records, that very little if any horizontal crustal compres- sion has occurred along continental margins where there are no oceanic trenches.Very small graben-like features, filled with sediment of density 1.90 gm/cm3, are observable on the cross section near 126° W long.These structures have also been observed on seismic reflection records obtained by the University of Washing- ton, and appear to be filled with unconsolidated sediments (McManus, l967c).These structures suggest crustal tension. A seismic reflection profile, reported by Ewing etal. (1968), across the continental slope and Cascadia Plain approximately 200 km north of the Washington cross section, shows a sediment-filled trough at the base of the continental slope.The trough appears to have been formed by faulting of the basement rock near the base of the slope with the oceanic side down relative to the continental side.Inter- preted as a normal fault, this structure also suggests tenson along the continental slope off Washington. The occurrence of tension along the continental slope would suggest that the structure inferred from the gravity measurement is a large normal fault offsetting the upper sedimentary layers and extending into the lower crustal layer.The trough along the scarp formed by the fault is apparently sediment filled.

The South Vancouver Island Cross Section

The south Vancouver Island cross section CC' is shown in Figure 15.The western end of the south Vancouver section is located in Tufts Abyssal Plain, near 44.7° N lat.,132. 7°W long.The cross section strikes N 53° E cutting obliquely across Juan de Fuca idge and the Cascadia Abyssal Plain.The section is normal to the trend of the shelf and crosses the coast of Vancouver Island near the south side of Barkley Sound.It traverses the Coast Mountains and has its eastern terminus on the east flank of the Rocky Mountains, near 52.3° N lat., 117.l°.W long. The average free-air gravity anomalies over Juan de Fuca Ridge and Cascadia Abyssal Plain are near zero, indicating that these 3 0 -Sc0C IC(A 0('SC 0 0 I'-3- : at, : 15to 25 508 2520 3C FREE-AIR *2(1 BOUSUERCRUST GRAVITY *2(1 AP4ALYSJBCRUSTAL PROFLE CROSS SECTION CC' SOUTH VANCOUVER ISLAND 0 908IZQVThL 50 sraE 68 40 SEISMIC REFRAC11ON UNES (AFTER 514CR NI. 960- SENSTESCJTEDORSERVED AN088ALY NALY LAYERS GOACM' 00000 66 DIETERS 35 5045 OREGON STATE INVERSITY APRL 968 5045 Figure 15. Southsection Vancouver CC'. Is land crustal and subcrustal cross J 0 two provinces are essentially in isostatic equilibrium. Thefree-air gravity anomalies become negative over the continental slope and dip sharply negative over the continental shelf.The negative anomaly over the shelf may be interpreted as reflecting asedimentfilled synclinal depression. Bouguer anomalies on land are based on gravity studies reported by Miller and Hughson (1936), Garland and Tanner (1957), and Walcott (1967).Isostatic gravity anomalies reported by Miller and Hughson (1936), Garland and Tanner (1957), Walcott (1967), and Banks (1969) indicate that regionally Vancouver Island, Georgia

Strait and the Coast Ranges of British Columbia areinisostatic equilibrium. Seismic refraction measurements reported by Shor etal. (1968) provide data on crustal layers and velocities in the mantle across Juan de Fuca Ridge and Cascadia Abyssal Plain.Refraction results reported by Richards and Walker (1959), which indicate a Moho depth of 43 km beneath the Alberta Plains were used for control along the east end of the south Vancouver section.The crustal layers and mantle velocities in western British Columbia are not defined by seismic refraction data.However a Moho depth of 34 km under the Interior Plateau between the Coast and Rocky mountains of British Columbia is near the 31 km depth determined by White and Savage (1965) from an unreversed refraction profile. Beneath the continent, layer thicknesses and depths depend on the assumed densities; consequently, the structures are indicated by dashed lines.Garland and Tanner (1957) attribute the smaller Bouguer anomalies of interior British Columbia to anomalous densities in the upper crust and the larger anomalies to warping of the base of the crust.The crustal structure in Figure 15 shows a similar interpretation. White and Savage (1965) concluded that the crust of Vancouver Island is 45 km thick.Ellis etal. (1968) have reanalyzed the seismic explosion data obtained by the Dominion Observatory in the Vancouver Island region from 1953 to 1963 and conclude that the Moho is deeper than 56 km. These estimates of Moho depth are based on the absence of identifiable Pn arrivals at distances greater than 400 km. The depth to Moho beneath Vancouver Island as indicated on the South Vancouver Island section is approximately 27 km. A crustal thickness of 56 km would require an intermediate crustal layer of density 3. 22gm/cm3approximately 42 km thick to contain the same overburden mass as indicated for the Cascadia Abyssal Plain.The high density and thickness of the intermediate layer required by a deep Moho seem excessive for an acceptable crustal section. Walcott (1967) presented a gravity profile in a line of section across Juan de Fuca Strait which extended north over the Vancouver 100

Island gravity high and south over the Olympic Mountains low,He postulated that the anomaly is produced by the edge affect of two juxtaposed crustal slabs of different density and thickness. He assumed that the thickness of the crust below Vancouver Is land is 51 km based on the results of White and Savage (1965) and showed that a crustal block beneath the Olympic Mountains 33 kmthick and 0. 2 gm/cm3less dense than the Vancouver Island crustal block would satisfy the gravity requirements.If, however, the crust is assumed to be 19 km thick beneath western Washington in accord with the Washington section,the same density contrasts (Walcott, 1967) lead to .a 28 km thickcrustbeneath Vancouver Island in good agreement with the South Vancouver Island section. The Washington and Vancouver Island sections suggest that approximately half of the large gravity anomaly observed over the Strait of Juan de Fuca is due to an 8 to 10 km vertical step in the Moho and half due to an approximately 0. 2gm/cm3average density contrast in the upper crustal rocks. Crustal deformation at the base of the slope, apparent on the Washington (Figure 14) and Scott Islands (Figure 16) sections, is not present on the South Vancouver Island section.Figure 6 shows a saddle in the gravity anomalies where the section crosses the junc- ture between abyssal plain and continental slope.The discontinuity in the crustal deformation along the base of the slope at this location 101 may be due to an unexplained structural change or simply due to a change in strike of the edge of the continental block.

The Scott Islands Cross Section

The Scott Is lands crustal and subcrustal section DD' is illustrated in Figure16. The western end of the Scott Islands section is located in Tufts Abyssal Plain, near 47.30N lat., 135.2° W long.The section strikes N 55° E and intersects the continent between the north end of Vancouver Island and the Scott Islands. Its eastern terminus is in the Interior Plateau of British Columbia near 52.40 N lat. ,12 5.10W 1 ongThe average free-air anomalies over Cobb Rise and Explorer Ridge are near zero, indicating these two regions are approximately in isostatic equilibrium. A large negative anomaly associated with the Scott Islands fracture zone occurs over the continental slope.The free-air anomaly is positive landward of the shelf break between the north end of Vancouver Island and the Scott Is lands and near zero over Queen Charlotte Strait.Bouguer anomalies along the Inside Passage of British Columbia, reported by Banks(1969),decrease rapidly toward the Coast Mountains. Average free-air anomalies over Queen Charlotte Strait and isostatic anomalies along connecting fjords (Banks,1969)are near 50 0 so -50 -500 -4 - -100 I50 -too L 5 5 0 0 274 103 N55E. 132 130 2.40 0 5 2.00 2S0 I 2.74 5 C 3 ------, '--__ \1 10 2O / IIIj 3. I. 3.00 20 3025 ;JIJliI11IIrI 3025 4 35 FREE-AIR AND BOLJGUERCRUST GRAVITY AND ANALY SUBCRUSTAL PROFILE CROSS SECTION DD OBSERVED GRAVITY SCOTT ISLANDS 332 35 40 SEISMC REFACTION LUES (AFTER MLPL 964)- DNS1TIESCOtI'UTED GR/ffY 0000 LAYERS Il GMICM3 0 NOR14TAL SCALE 50 100 40 45:50 Figure 16. OREGON STATEScott INVERSITY Islands crustal and subcrustal MAY 968 cross section DD'. KILOTERS 5045 0 103 zero indicating that the area about the continental end of the Scott Is lands section is in isostatic equilibrium. Seismic refraction measurements reported by Milne (1964) show a mantle velocity of 8. 4 km/sec and a relatively thin, 10 km thick crust in the southern part of the Alaskan Abyssal Plain.These measurements provide data on crustal layers and velocities on the west end of the section.Beneath Cobb Rise, Explorer Ridge, and the continent layer, thicknesses and depths depend on assumed densities; consequently, the structures are indicated by dashed lines. Gravity values of Miller and Hughson (1936) and Garland and Tanner (1957) and the seismic refraction results of Richards and Walker (1959) were extrapolated over the Alberta Plains to provide computational control for the east end of the section.The section shown, however, is terminated in the Interior Plateau of British Columbia. If near normal crustal densities are assumed for Cobb Rise and Explorer Ridge, light material must be postulated in the upper mantle to contain the same over-burden mass as indicated in the Alaskan Abyssal Plain.Figure 16 shows an upper mantle structure similar to the structure shown in Figure 14 beneath Juan de Fuca Ridge. A material of low density, here assumed to be 3. 18 gm/cm3, overlies a less low density material of 3. 27 gm/cm3 and extends from Eickelberg Ridge near 132.8° W long, to the continental shelf. 104 The Moho off the north end of Vancouver Is land is shown as 27 km deep; if lighter average crustal densities are assumed the Moho will be correspondingly shallower. The Bouguer anomaly over the coast mountains is attributed to relatively light rocks in the upper crust and a crust extending to depths of approximately 35 km. No attempt was made to fit exactly the short wave length gravity anomalies caused primarily by the very irregular topography of Cobb

Rise and Explorer Ridge. The negative anomaly near 131.10W long. suggests a sediment-filled 2 km deep moat along the west side of Explorer seamount, similar to the trough along the south side.The sharp negative anomaly near 129.70W long, is primarily due to the median trough (Explorer trough) of Explorer Ridge.The higher elevation of Explorer Ridge is shown compensated by a thicker section of the 3. 18gm/cm3mantle material. The large negative anomaly of the Scott Islands fracture zone near 129° W long, is interpreted on the section as a graben. A crustal block of 2. 8 gm/cm is shown downthrown 6 km and the resulting valley filled with two 3 km thick sediment sections of average densities 2, 0gm/cm3and 2. 4gm/cm3. Ewing etal. (1968) reported a seismic reflection profile which passes near the north end of the Scott Islands anomaly. The profile shows normal faulting and narrow-filled trenches near the base of the slope but no evidence of folding.The interpretation of the negative 105 anomaly of the Scott Islands fracture zone as crustal downfaulting is supported by the seismic reflection observations. The large negative anomaly associated with the Scott Islands fracture zone is shown in Figure 17,The -100 mgal contour used as an outline indicates that the structure is not a simple graben. The irregular anomaly contours suggest that either it is composed of a large number of small blocks or that a large graben has been subjected to shear and consequently deformed. Figure 18 shows the total magnetic intensity observed over the Scott Islands fracture zone.This zone exhibits very little magnetic relief.The magnetic highs near 50.7° N lat. ,129.1° W long, and 50. 3° N lat., 128. 3° W long, coincide with changes in the contours of the free-air anomalies and are probably due to near-surface basement rocks along the outer edge of the continental shelf.The large magnetic highs near 50.0° N lat,129. 7° W long, and 49.9° N lat., 129. 6° W long, are the north end of a long magnetic lineation and in this area coincide with the negative gravity anomalies associated with the north end of the Explorer trough. Figure 19 shows a composite profile along line S-S' across the Scott Is lands fracture zone where the observed gravity anomaly is a minimum. The location of profile S-S' is shown in Figure 8.The gravity and magnetic anomalies are negative from the continental shelf to the abyssal depths of the sea floor,The minimum (-160 mgal) 106

1300------129°- 128°

510 510

500 500

1290 1280

Figure 17.Free-air gravity anomaly map of the Scott Islands fracture zone. 107

130° p - - - - qo- _ - - - - - SCOTT ISLANDS - TOTAL MAGNETIC INTENSITY J 1 CONTOUR INTERVAL \\'\. 510 51°

010

: '1' -

50° 50°

129 128

Figure 18.Total magnetic intensity map of the Scott Islands fracture zone. SCOTT ISLAND FRACTURE ZONE a IaE GRAVITY, MAGNETIC, ALONG LINE S-S -c BATHYMETRIC PROFILE zC,IdI-C-, +250 4 0 Ea IdI&Ia-e-. 0 -100 4000 -250 -200 800 16001200 -500 -300 24002000 0 Figure 19. 20 theBathymetric, Scott Islands gravity fracture and zone magnetic along line profile S-S'. across 40DISTANCE (kilometers) 60 I.1.' 109 in the free-air anomaly 60 km from the southwest end of the profile may indicate a fault at the base of the slope.The bathymetric highs at 8 km and 26 km are named Paul Revere Ridge and Winona Ridge respectively on a bathymetric map of the region compiled by McManus

(1964).

The Moresby Island Cross Section

The Moresby Island crustal and subcrustal section EF' is shown in Figure 20.The western end of the Moresby Island section is located in the Alaskan Abyssal Plain near 49.9° N lat. ,138.2° W long.The section strikes N 42° E across the Alaskan Plain and intersects the continent near the center of Moresby Is land.The section crosses Hecate Strait and intercepts the Coast Mountains in the proximity of Douglas Channel.It5 eastern terminus is in the Interior Plateau of British Columbia near 55° N lat., 126° W long. The average free-air anomalies over the Alaskan Abyssal Plain are near zero indicating that the plain is in approximate isostatic equilibrium. A large negative free-air anomaly occurs over the continental slope off Moresby Island.Bouguer anomalies are positive over Moresby Island (Stacy, 1967).Free- air anomalies inHecate Strait average near zero, and Bouguer anomalies (Banks, 1969) decrease rapidly along the Douglas channel toward the Coast Mountains. 100 50 100 -.4 0 O50 K -(00-50 -100-50 -ISO H -I5O K' 5- 138 136 ZL 134 130 267 I28 0-5 - 200 I 03 210 .1 \I 273 L'., I / 276 10 _-_ 297 'A - -5 10 IS -j_ ------IS 2025 297 - 20 EL 30 - FREE-AIR AND BOUGUER GRAVITYCRUST ANOMALY AND SUBCRUSTAL PROFILE CROSS SECTION EE MORES BY ISLAND 40 35- - SEISMIC REFRACTION LINES (AFTER SHOR,I962) - DENSTIESCOMPUTEDOBSERVED OFGRAVITY GRAVITY LAYERS IN000000 GM/CM3 332 40 45 - OREGON STATE UNIVERSITY MAY 1968 0 HORIZONTAL SCALE KILOMETERS 50 00 50 Figure 20. Moreeby Island crustal and subcrustal cross section EE'. 50 C 111

Average free-air anomalies over Hecate Strait andisostatic anomalies along Douglas Channel (Banks, 1969) are near zero, indicating that the area about the continental end ofthe Moresby Island section is in approximate isostatic equilibrium. Shor (1962) reported the results of seismic refraction measurements along a profile southwest from Dixon Entrancein the Alaskan Abyssal Plain.The Moho is about 1.5 km shoaler between Hodgkins Ridge and the Queen Charlotte trough than west ofHodgkins Ridge. Mantle velocities are approximately 8. 3km/sec west of Hodgkins ridge and 8. 1 km/sec east of the ridge. A long compoundrefraction profile from the northern tip of Graham Is land east toCelestial Reef suggests that Moho depths are approximately 25 km nearDixon Entrance. The Alaskan Plain shows no marked changes intopography between the seismic refraction profile of Shor(1962) and the similar refraction measurements near the southern edge of theAlaskan Plain reported by Milne (1964).The seismic and bathymetric observations suggest the southeast sector of the plain has a uniformstructure. The refraction results of Shor (1962), therefore, wereextrapolated southeast along isobaths to provide a guide to crustal andsubcrustal layer depths in the Alaskan Plain and the depth of theMoho beneath Hecate Strait.Structures of the upper crust beneath Moresby Island and east and west of Moresby Is land are not controlled byseismic 112 measurements and, therefore, are shown as dashed lines. The Moresby Island section indicates the Moho shoals to approximately 9 km immediately west of the Queen Charlotte troug.h and then descends to a depth near 25 km beneath Moresby Island and Hecate Strait.The densities assumed for the crust at the east end of the section result in a Moho depth of 40 km beneath the Coast Mountains.The combined gravity and seismic observations require a decrease in density of the upper mantle beneath the continental rise. Short wave length gravity anomalies near 134° W long, suggest faults or intrusions in the upper crust beneath the surface sediment layer.The wave lengths of the gravity anomalies are approximately the same as the wave lengths of the magnetic anomalies in the region. However, magnetic measurements were not made along the same track line as gravity measurements; consequently, the results of a correlation of gravity and extrapolated magnetic data were inconclusive, Figure 21 shows total magnetic intensities southwest of Moresby Island.The magnetic lineations observed over the Alaskan Plain trend N 20 E off Moresby Island and appear to terminate at the base of the N 43° W trending continental slope.The gravity and magnetic anomalies show little correlation; the structure producing the well- defined gravity high at 52.1° N lat.,131.6° W long, causes only a small perturbation in the magnetic field.The magnetic high over 113

I 5 N 530

MORESBY ISLAND TOTALMAGNETIC INTENSITY CONTOURINTERVAL 00 gammas 50 gammas

/ / -- / Ii ', I, I I Ii I I 7 IllI It

: 1

I I I )

I' - e4

52 -

64 1063

(a/lu'

1310

Figure 21.Total magnetic intensity map of the area off Moresby Island. 114 Moresby Island may be associated with the syntectonic plutons, described by Sutherland-Brown (1966), which are concentrated along the western flank of the Queen Charlotte Islands. The interpretation of the crustal structure of Moresby Is land is in good agreement with the structure proposed by Sutherland-Brown (1966) based on geologic observations except that the large fault along the west side of Moresby Island is shown dipping away from the island rather than under the island.The actual dip of the predominantly right lateral strike-slip fault is unknown. Sutherland-Brown (1966) described a graben-like trough of early Cretaceous age in the central area of the Queen Charlotte Islands and also large northwesterly faults along the eastern flank of the islands which, in addition to right lateral strike-slip movement, exhibit normal, east block down displacement.This suggests east-west tensional stresses over the region contemporary with transcurrent motion along the Denali-Queen Charlotte Islands fault system off the west coast of the Queen Charlotte Islands.In a region of tension normal or transcurrent faulting rather than thrust faulting is expected; and, therefore, the large trans current faults along the west side of the Queen Charlotte Islands may dip seaward and have a normal fault component with the ocean down relative to the islands. Sutherland-Brown (1966) indicated that folding is less important 115 than faulting in the Insular Belt of British Columbia; consequently, the short wave length gravity anomalies in Hecate Strait are interpreted as caused primarily by abrupt changes in water depths and faulting and/or glacial gouging of the upper crustal layers.

The Dixon Entrance Cross Section

The Dixon Entrance crustal and subcrustal section FF' is shown in Figure 22.The western end of the Dixon Entrance section is located in the Alaskan Abyssal Plain near 52.90 N lat.138.90 W long.The section strikes N 69° E and intersects the continent in Dixon Entrance off the north end of Graham Island.Its eastern terminus is in the Interior Plateau of British Columbia near 55. 6° N lat. ,1 27. 20 W long. The average free-air anomalies over the Alaskan Plain, except for the local anomaly over Hodgkins Ridge, are near zero indicating the region is essentially in isostatic equilibrium. A large negative anomaly occurs over the continental slope.The average free-air anomaly is slightly positive landward of the edge of the continental shelf and decreases to near zero in the eastern part of Dixon Entrance.Bouguer anomalies along the Pearse Canal and the Portland Canal, reported by Banks (1969), decrease rapidly toward the Coast Mountains. Average free-air anomalies over the eastern part of Dixon 5C -j C 050 -oo-50 -150 5 j-In 0 138 103 136 - 134 2.I0____132 130 128 -5 05 216 - 05 20 15 20(5 3O25 XON ENT CE 25 35 FREE-AIR AND BOUGUER GRAVITY ANOMALY PROFILERUST AND SUSCRUSTAL CROSS SECT 3.32 3530 SEISMIC REFRACTION LINES COMPUTED068ED GRA(A ONAV1TY 0000 / - 40 OREGONDENSITES STATE OFUNWERSITY LAYERS IN - (AFTER SHOR. 9a G$4C&13 I HORIZONTAL SCALE 40 7, MARCH 0 50 00 - 45 1968 45 Figure 22. / Dixon Entrance crustal and subcrustal cross section FF'. / KILOMETERS 50 117 Entrance and isostatic anomalies along the Portland Canal (Banks, 1969) are near zero indicating that the area about the eastern end of the cross section is approximately in isostatic equilibrium. Shor (1962) reported the results of seismic refraction measure- rnents along a profile southwest from Dixon Entrance into the Alaskan Abyssal Plain.The profile coincides closely with the Dixon Entrance section and provides references for the crustal and subcrustal structures in the Alaskan Plain and in Dixon Entrance. Mime (1964) reported a reversed profile shot in a north-west,southeast direction off the north tip of the Queen Charlotte Islands.These results provide data on the structure of the upper crust near the edge of the continental shelf. The Dixon Entrance section indicates the Moho shoals to approximately 9 km beneath the continental rise immediately west of the continental slope off Dixon Entrance and then descends to approximately 25 km beneath Dixon Entrance. The densities assumed for the crust at the east end of the section result in a Moho depth near 40 km beneath the Coast Mountains. The combined gravity and seismic observations require a decrease in the density of the upper mantle beneath the continental rise. The gravity anomaly near 1 36. 5° W long, is due to Hodglcins Ridge and seamount. The two-dimensional analysis inaccurately portrays the three-dimensional structure; however, a moat can be 118 noted along the ridge flanks, and regional isostatic compensation is suggested. Figure 4 shows magnetic anomalies determined from magnetic measurements made along the Dixon Entrance gravity profile.The results of a correlation of the magnetic anomalies and the short wave length gravity anomalies between Hodgkins Ridge and the continental slope were inconclusive. Shor (1962) reports that an unreversed seismic profile at the edge of the continental shelf west of Graham Island indicates a sedimentary layer 3 km thick.The cross section indicates sediment 3 km to 5 km thick near the edge of the continental shelf.The large negative anomaly over the continental slope is attributed to a thick sedimentary wedge along the slope and associated faulting of the crust with the ocean side down relative to the continental shelf.The maximum sediment thickness actually occurs landward of the axis of the Queen Charlotte trough.The short wave length gravity anomaly near 134.2° W long, superimposed on the large negative anomaly may mark the location of a fault near the base of the slope. A high density crustal block beneath the outer shelf is probably related to the substructure of the Queen Charlotte Islands.Both seismic refraction (Shor, 1962) and gravity measurements suggest extensive faulting and structural changes in the upper crust of the continental shelf and Dixon Entrance.The negative anomaly near 119 131.8° W long, is associated with an extension of the topography of Clarence Strait in southeast Alaska.

The Continental Margin Structures off Washington and British Columbia

The five crustal and subcrusta]. cross sections are effectively connected on the western end by a common section; consequently, the structures outlined on the sections may be compared and, with the aid of Figures 6 and 7, their lateral extent estimated. The central Washington and south Vancouver sections indicate a Moho depth of 7 to 9 km and a relatively light material in the upper mantle beneath Juan de Fuca Ridge. The Scott Islands section also suggests a Moho depth of 9 km and a similar low density material in the upper mantle beneath the Explorer Ridge.In the Scott Islands section the low density upper mantle is a consequence of the higher topography of Cobb Rise and Explorer Ridge without an associated increase in the free-air gravity anomaly, The region, bordered on the northeast by the Scott Island fracture zone and on the southwest by Eickelberg Ridge and Cobb Seamount, exhibits numerous seamounts and an average depth less than the surrounding abyssal regions whereas the free-air gravity anomaly averages near zero over both the region of seamounts and the abyssal plains.This implies that the entire region of approximately 135, 000km2may have a relatively 120 shallow Moho and light material in the upper mantle. If, however, a shallow Moho and light upper mantle materials are confined to the immediate vicinity of Explorer and Juan de Fuca ridges Moho depths in the rest of the region will be approximately 12 to 13 km.In this event the crustal thickness, exclusive of the water, will be 9 to 10 km compared to a crustal thickness of approximately 6 km beneath the surrounding abyssal plains. All five cross sections show the upper mantle, in the region extending from the coastline out to a line which closely approximates the 3500 m (1800 fm ) isobath to be slightly less dense than the upper mantle of Tufts and Alaskan abyssal plains farther west.The region of low density mantle apparently extends from the Mendocino escarpment (Dehlinger etal., 1967a; Dehlinger etal., 1968) to and possibly including the continental rise off southeast Alaska. A deep trough is located near the base of the continental slope off British Columbia and extends from Dixon Entrance to the Strait of Juan de Fuca.Off the Queen Charlotte Islands the trough is associated with the Queen Charlotte trough, a bathymetric depression; and west of Vancouver Island it is exemplified by the Scott Islands fracture zone, with an estimated sediment thickness of 6 km or more. West of Washington a smaller sediment-filled trough occurs along the base of the slope.Gravity measurements suggest a filled trough is also located on the continental slope parallel to the trough 121 near the base of the slope. A relatively high density ridge or anti- clinal structure just landward of the edge of the continental shelf apparently extends the length of the Insular Belt of British Columbia. This structure is absent off Washington. The central Washington and Oregon sections (Dehlinger et al., 1968) indicate that the !igranitic layer of the continent appears to be absent in southwestern Washington and northwestern Oregon. The central Washington sections shows a Moho depth of 19. 5 km under western Washington. A vertical offset of 8 to 10 km occurs in the Moho in the vicinity of the Strait of Juan de Fuca. The cross sections of British Columbia show mantle depths under the Insular Belt of approximately 25 to 28 km. The east end of the sections are based on little structural control,but they suggest that Moho depths are approximately 35 to 40 km beneath the coast mountains of Canada and Cascade Range of northern Washington. 122

CONCLUSIONS

Near zero free-air anomalies west of Washington and British Columbia indicate the entire region is approximately in isostatic equilibrium and that isostatic adjustment is more rapid than tectonic upheaval.The regional gravity low of Puget Sound and the Olympic Peninsula, which suggests isostatic inequilibrium, extends westward to the edge of the continental shelf.Isostatic compensation of individual searnounts in the region appears to occur over an area approximately 7 times the basal area of the seamount. Except for isolated seamounts, free-air anomalies, as elsewhere in the Pacific Ocean (Dehlinger, 1969), are not correlative with topography.Over Juan de Fuca Ridge free-air anomalies show no apparent relation to the ridge nor to the magnetic anomalies which define the ridge. A free-air anomaly of approximately -100 mgalis associated with the partially sediment-filled Queen Charlotte trough west of the Queen Charlotte Islands.Off the north end of Vancouver Is land a free-air anomaly greater than -150 mgaloccurs over the Scott Island fracture zone. An anomaly less than -50 mgal occurs along the base of the continental slope off Washington. The negative anomaly along the base of the continental slope off British Columbia and Washington is caused by the dip of the Mohorovicic discontinuity, lateral density 123 differences in the upper mantle and a sediment-filled trough. In the region between the coastline and a line which approximately follows the 3500 m isobath, a mass deficiency must occur in the upper mantle where densities are slightly less than in the Tufts and Alaskan abyssal plains and in the adjacent continent.Still lower densities appear to occur in the upper mantle beneath the Explorer and Juan de Fuca ridges. Explorer Ridge, proposed to be an active locus of sea-floor spreading, exhibits moderate s eismicity, asymmetric wall structures and magnetic anomalies, and a median trough with approximately two kilometers of relief and little sediment fill. Moho depths are near 25 to 28 km in the Insular Belt of British Columbia and 19 km in western Washington. Moho depths appear to be near 20 km in western Oregon (Dehlingeretal., 1968; Thiruvathukal, 1968) and 21 to 23 km in central California (Tatel and Tuve, 1955; Thompson and Taiwani, 1964).The relatively thin crust in the region between the continental shelf and coast mountains of British Columbia, the Cascade Range in Washington and Oregon and the Sierra Nevada in California, which includes the coast ranges of California, Oregon and Washington, is apparently characteristic of the transition from oceanic to continental structure in western North America. 124

REGIONAL TECTONICS

Two tectonic processes active in the region west of Washington and British Columbia are:1) a geotectonic process causing the formation of the mid- ocean, ridge-rise system, and 2) isostatic adjustment tending to reduce hydrostatic imbalances caused by the lateral transport of material. Morgan (1968b) and LePichon (1968) interpreted the placement and configuration of the mid-ocean ridges and their related faults and magnetic anomaly patterns as a consequence of the rotation, displacement and interaction of rigid crustal blocks moving about the earth's surface on a deeper, more fluid layer.Each of the rigid blocks is bounded by rises where new crustal material is formed as the blocks separate, trenches or young fold mountains where crustal materials are being destroyed, and great faults along which the blocks slide past one another.The seismological studies of Isacks etal. (1968) give impetus to this interpretation, particularly with respect to the destruction of crustal material in the vicinity of the deep trenches. West of Washington and British Columbia the Juan de Fuca and Explorer ridges correspond to the boundary of separation between the north Pacific Ocean block and the North American continent block. The Queen Charlotte fault is a great fault along which the northeast Pacific Ocean is moving northwestward relative to British Columbia 125 and southeast Alaska eventually underthrusting southeast Alaska and the Bering Sea in the vicinity of the Aleutian Trench. An analysis of fault plane solutions of earthquakes in Washington and off British Columbia (Hodgson and Milne, 1951; Hodgson and Storey, 1954; Algermissen and Harding, 1965; Tobin and Sykes, 1968; and Gallagher, 1969) indicates the strike of the minimum horizontal stress axis is approximately N 90° W in western Washington and the Strait of Georgia and N 300 W to N 600 W off the north end of Vancouver Island, in Hecate Strait and north of Graham Island.This suggests the entire region, north of the Strait of Juan de Fuca, is subject to tensional stress in a general northwest-southeast direction. The graben-like structures, indicated by the gravity and seismic reflection studies, which occur in the north end of the Cascadia Plain, along the base of the continental slope and on the continental slope and shelf support the interpretation of the fault-plane solutions as indicating regional tension. The separation of crustal blocks along Juan de Fuca Ridge causes the formation of new crust along the ridge axis and consequently an extension of the crustal blocks.Deep-tow magnetometer measurements (Larson and Spiess, 1969) and model analysis (Emi.lia, 1969) suggest that along active ridges new crust may form as narrow intrusions such as dikes (Bodvars son and Walker, 1964) or diapirs (Maxwell, 1968).The formation of new crust along Juan de Fuca 126 Ridge implies either a net transfer of crust out of the area or destruction of crust in an unfamiliar manner, i. e. without earthquakes, compressive features along continental margins or prominent gravity anomalies.The patterns of stress and related faulting of the region are interpreted here to indicate that no underthrusting of the continent is occurring along the continental margin of Washington and British Columbia.This implies northwestward migration of Juan de Fuca Ridge.The asymmetrical patterns of topography, gravity and heat flow with respect to the ridge axis are not inconsistent with the concept of a migrating ridge.The ridge substructure, shown by the crustal and subcrustal cross sections extends into the upper mantle, indicating the upper mantle is also part of the crustal block in motion. Vine (1966, 1968) proposed a 3 cm/yr spreading rate perpendicular to the axis of Juan de Fuca Ridge.The trend of the topographic structures and gravity anomalies of the region and the strike of the Queen Charlotte fault suggest that separation along the ridge axis is slightly oblique and at a rate near 3. 2 cm/yr.If the ridge is migrating, the crust on the east side is stationary, whereas the crust on the west side is moving northwest with a velocity near 6. 4 cm/yr. The distance between the two magnetic anomalies identified as occurring on the ridge crest 10 million years B, P. (Vine, 1968) indicate an average velocity of separation parallel to the Queen Charlotte fault near 8 cm/yr. The additional crustal extension (or higher velocity 1 27 of separation) is apparently accommodated by strike-slip faulting oblique to the ridge axis, clockwise rotation of the spreading axis, and crustal extension by normal faulting. The magnetic ages combined with the concept of a migrating ridge indicate the ridge axis was located along the continental slope off Washington and Oregon 10 million years ago.Vine (1966) noted a marked change in the character of the magneticanomalies near the 10 m. y. magnetic date.Ewing and Ewing (1967), in a study of the sediment distribution on the mid-ocean ridges, noted a large change in sediment thickness near anomalies with a 10 m. y. magnetic date. They indicate a period of relative quiescence in sea-floor spreading between late Mesozoic or early Cenozoic and 10 m. y. ago. LePichon (1968) suggests each cycle of the intermittent spreading is marked by a reorganization of the global pattern of motion. Menardand Atwater (1968) interpret changes in strike of transform faults in the North Pacific as indicators of changes in direction of block separation. Stonely (1967) noted that Mesozoic and Cenozoic sediments depoisited in a trough along a rapidly subsiding shelf in southeast Alaska were deformed in late Cenozoic time by continuous movement of the Pacific Ocean floor relatively toward the continental margin.These observations suggest a change in geotectonic motion and a consequent change in the regional stress pattern of the northeast Pacific in late Tertiary time; Juan de Fuca Ridge, therefore, could have formed near the base of the continental slope off Washington and Oregon approximately 10 m y. ago in response to a changing stress pattern. 128 The continental slope between northern California and the Strait of Juan de Fuca presents a natural point of rifting between the nearly en echelon San Andreas and Queen Charlotte faults.The uplift of the Coast Range and formation of the high Cascades (Baldwin, 1959) are contemporary with the formation of the ridge 10 m. y. B. P.The period of quiescence between the late Mesozoic and the occurrence of rifting of the continental and oceanic blocks near the slope base is contemporary with the filling of the Coast Range geosyncline. The direction of relative motion between the continental and oceanic blocks is oblique to the ridge axis; consequently, shear--predominantly NW trending right-lateral strike slip and NE trending step faulting- -may have occurred along the continental margin prior to and during the formation of the ridge and separation of the crustal blocks.This hypothesis does not preclude pre-Tertiary thrusting along the continental margin. The apparent rotation of Juan de Fuca Ridge is due to the tendency of active ridges to orient themselves normal to ridge terminating transform faults (Menard and Atwater, 1968).Ridge formation, by block separation as megascale tension or cleavage fractur- ing and intruding or by crustal thinning and diapiric injection, is expected to occur so as to involve a minimum of deformation energy. Ridges tend toward an orientation normal to the trend of the regional tensional stress vector and will rotate by growth so as to minimize 129 shear stress and shear fracture along the spreading axis. Explorer Ridge, which formed approximately 4 m. y. B. P., is now normal to the crustal velocity vector, the Queen Charlotte transform fault and the continental margin of British Columbia.Prior to the formation of Explorer Ridge the Scott Islands fracture zone acted as a sector of the north-bounding transform fault of Juan de Fuca Ridge. Analysis of magnetic anomalies (Vine and Wilson, 1965; Morgan, 1968a, b) suggest Juan de Fuca Ridge is an active locus of sea-floor spreading.Hamilton (1967), based on the observation of undisturbed turbidities, suggests the Juan de Fuca portion of the East Pacific Rise has been quiescent since middle Tertiary time.Por- tions of the ridge axis appear covered with sediment (McManus,

1965).The free-air anomalies indicate isostatic adjustment has occurred and no vertical tectonic imbalance is present.Earthquake epicenters are notably absent along the central portion of the ridge but occur along or near the ends of the ridge (Tobin and Sykes, 1968). Duncanetal. (1969) note a small rift aligned with the axis of the ridge on the south end, and Gallagher (1969) has obtained a fault plane solution which suggests normal faulting parallel to the south end of the axis of Juan de Fuca Ridge. These observations suggest that spreading may not occur simultaneously all along the ridge and that part of the ridge--the center portion- -may be temporarily quiescent. The fault or fracture system connecting Juan de Fuca and Explorer 1 30 ridges and the Explorer Ridge and Queen Charlotte fault are, as yet, undefined. An analysis of fault-plane solutions of earthquakes off northern California and southern Oregon (Byerly, 1938; Tobin and Sykes, 1968; Bolt etal., 1968; and Gallagher, 1969) indicates:tbads of minimum horizontal stress trends east-west, whereas the stress axis off British Columbia trends northwest-southeast. The difference in stress axes suggests that a north-south component of tension may exist between the two regions.At present the north-south stress component may be relieved by crustal extension in the form of normal fault components and possibly the formation of a small amount of new crust along the Blanco fracture zone. Vine (1966), Pavoni (1966) and Peter and Lattimore (1969) interpreted the smaller offsets in the linear magnetic anomalies noted by Raff and Mason (1961) as strike-slip faulting.The offsets strike nearly 30° to 60° to major structures of the region (e.g. Juan de Fuca Ridge) and are consistent with regional tension in a NW-SE direction.It is also possible that some of the offsets in the magnetic anomalies may have formed in situ; the offsets may have propagtedin the crust similar to the manner whereby crystal or lattice defects are propagated or extended in crystals during growth. The magnetic offsets or pseudo-faults may have started or been guided along old fractures or crustal weaknesses.It is noted that the "Sovanco fault, 131 a major left-lateral strike-slip fault postulated by Pavoni (1966), based on magnetic offsets and used in the arguments of Peter and Lattimore (1969), passes through and is normal to a graben (Ewing and Ewing, 1968) located near 49.00 N 128.8° W long. The directions of stress necessary to form both features concurrently are irreconcilable.Further the destruction of large quantities of old crust required in their interpretation of the magnetic offsets is incompatible with the structure of the region resulting. from the inter-. pration of the combined gravity and seismic reflection measurements, i. e. there is no evidence of sinking, down warping, overthrusting or superposing of crustal blocks except possibly along the east end of the Mendocino escarpment. Sediment thickness increases eastward in the Cascadia Abyssal Plain consistent with the concept of a migrating ridge or spreading sea floor.In some areas the sediment surface appears to slope away from Juan de Fuca Ridge.This may be in part attributed to uplift of the ridge and in part to an isostatic adjustment or subsidence of the oceanic crust proportional to sediment thickness. The importance of isostatic adjustment is realized when it is noted that Cascadia

Plain contains nearly 68, 000km3of sediment, a sufficient amount (via subcrustal mass displacement) to elevate the Coast Range and continental shelf off Washington and Oregon 1000 meters. The near- zero average free-air anomalies, thin crust and high heat flow of 132 Cascadia Plain suggest that isostatic adjustment is relatively rapid; at least more rapid than tectonic upheaval. Although the present erosion rate per unit area is 4 to 5 times greater along the Coast Range than east of the Puget-Willamette trough, most of the sediment in Gas cadia Plain comes from the Columbia River drainage basin (Kulmetal., 1968).The analysis of gravity anomalies observed in Oregon (Thiruvathukal etal., 1969) indicates a slight mass excess in the Coast Range Mountains, whereas the Columbia Plateau has a mass deficiency.This suggests a possible phase lag in subcrustal displacement between Gas cadia Plain and the continental interior or renewed ridge activity near the continental margin. The Rocky Mountain Trench (Garland and Tanner, 1957; Leech, 1965) closely parallels the large negative free-air gravity anomalies of the Queen Charlotte trough, Scott Islands fracture zone and Puget Sound depression. The Rocky Mountain Trench and the aligned gravity minima flank the east and west sides, respectively, of the Canadian Cordillera.If these two features are related, a late Cretaceous or early Tertiary age is suggested for the structures defined by gravity minima. The relatively thin crust of western British Columbia may be a Mesozoic relic of oceanic underthrusting of the continental margin. 133

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APPENDIX 1

GRAVITY DATA REDUCTION TO FREE-AIR ANOMALY

The gravity measurements used to construct the free-air anomaly maps of the region west of Washington and British Columbia shown in Figures 6 and 7 were made with LaCoste-Romberg gimbal suspended surface ship gravity meter S-9 from 1963 through 1967. Positioning, a critical factor in obtaining gravity measurements at sea, was determined principally by Loran A.Radar fixes were obtained where possible and sunlines, star fixes and radio positions were obtained on occasion as supplementary data.Dead reckoning was used where all other methods were obviated.All positions were replotted in the laboratory from the original data on a working track- line chart.The times at which the ship turned were obtained from the analogue gravity records, and changes in engine rotation rates, stearage and sea condition were obtained from the ships log.All data pertinent to the ships navigation were added to the track line chart, and lines of best fit were sequentially constructed from one point of change in velocity or heading to another.After a preliminary reduction of the gravity measurements the positions of the track line end points were re-evaluated and errors corrected or minor modifica- tions made.Positional accuracy and consequently gravity 149 measurements improved with experience. Measurements were obtained aboard the BERTHA ANN in July, 1963; the USC&GS SURVEYOR in February, 1964; the USCGC YOCONA in August, 1964, and May, 1965; and the R/V YAQUINA in June, 1965, July, 1966, April, 1967, and September, 1967. When using LaCoste-Romberg meter S-9, errors in gravity measurement occur due to a nonlinear response to vertical acceleration (Dehlinger etal.,, 1966a LaCoste, 1967).Corrections to the vertical errors are obtained using the relation

Ve =a(R-R)A2

where the vertical acceleration is given by the relation

k(B) AvT

B peak to peak amplitude of instantaneous beam motion logged on analogue recorder, T = period of beam motion in seconds,

R average beam position, and R, a, and k laboratory determined constants. The constants as determined by LaCoste-Romberg for the era of operation were:

21 Feb 61 R= 40, a2.3 X 10, k = 23 13Nov64R95,a2.0X104,k23 1 50

4 Dec64 R 40, a=1.1 X k=23

17 Mar65 R°=50.4, a2. 3 X 10, k=22

24 Feb66 R =-160, a=17.8 X 10, k=22

3 Mar66 R 54, a=1.8 X 10, k=29 Because of the change in the parameters with time (e. g. Mar 65 to Feb 66) in an unknown manner empirical equations were determined by statistical analysis from data obtained over the Newport marine gravity range.The equations used to determine the vertical correction for the surveys are:

BERTA ANN July,1963 2.94X105B2F2(95-R);

USC&GS SURVEYORFebruary,19642.94X105B2F2(95-R);

USCGC YOCONA August,1964 2.94X105B2F2(95-R);

USCGC YOCONA May,1965 3.09Xl05B2F2(43.9R); R/V YAQUINA June,1965) LBB(3. 79X102F+0.724) R/V YAQUINA July,1966 NONE

R/V YAQUINA April,1967 5.6X105B2F2(1 36. 4-R)

R/V YAQUINA September, 1967 NONE where the frequency of instanteous beam motion F1/T. Modification of the amplitude B in the correctionequ.ation usedin 1965 compensates for the attenuation of the higher frequencies by an RC filter in the phase comparator circuit of the gravity meter. The frequency response of the meter was improved in 1966. An addition of an automatic reader to the gravity meter in 1965 1 51 allowed a higher data sampling rate but extended the response.time of the gravity meter. The time constants of meter 5-9 varied from approximately 1 minute to nearly 5 minutes.Filter or overaging time constants can cause phase shift and variable attenuation of the higher frequency spectral components of the observed gravity anomalies. A time offset of 4 minutes was used to compensate for the meter time delay when operating with the automatic reader. Hanning smoothing (e.g. Blackman and Tukey, 1958, p. 15) was used on automatic reader data sampled at four minute intervals. A mathematical analysis showed that neither the meter response and transfer function nor the hanning smoothing significantly distorted the observed gravity field. A computer program to reduce the gravity measurements to free-air anomalies was written by Mr. M. Gemperle in FORTRAN IV for use on the CDC 3300.The program parameters are: 152

INPUTS

1. Calibration curve - converts gravity meter counter units to milligals. Input is 50 data points - at even l00's of gravity meter counter units F OR MAT (8 F 10.0)

2. Base Station Data, etc. a) Base meter reading (in counter dial units) b) Base gravity value c) Year and month of cruise d)Effective date of results e) Four. letter code for ship f) Four. letter code for base station g) Code for punching output cards (blank no cards; 1 card)' h) Meter drift i) Code for interpolation subroutine to find location of points on track lines with even lO's free air anomaly (blank - no interpolation; 1-interpolation) j) Time constant for averaging circuit FORMAT (2F10.1, 4(lx,A4), 3X, 12, F5.l, IS, F5.1, F5.2)

3. Navigation Cards a) Time and lat and long for beginning of line b) Time, lat, and long for end of line. c)Profile number FORMAT (2 (3F2.0, F3.0, F5.1, F4.0, FS.1, 2X), AS)

4. Gravity Station Cards a) Time of reading b) meter reading (in meter units) c) depth in fathoms d) freq. e) amp f) average reading line g) Browne correction h) quality (poor or blank) i) Test for last card in profile (1-last card; blank-not last card) j) Test for first card in profile (1-firstcard; blank-notfirst card) FORMAT (3F2.0, F9.1, SF5.0, 1X, A4, Ii, 9X, 15) 1 53

C) I TT PT IT S

1. Base station information a) Date effective b) Zero gravity (gravity corresponding to zero meter units) c) base meter reading d) base gravity e)drift f)Filter time constant FORMAT (18H1RESULTS EFFECTIVE, ZX, A4, 3X, 13HZERO GRAVITY=, FlO. 1, Z1H BASE METER READING=, F7. 1, 15H BASE GRAVITY, FlO. 1, 8H DRIFT, F5. 1, 5X, 4HRC, F6. 2)

2. Navigation Data a) Profile number b) Speed c) Course d) Time, lat, and long of first point e) Time, lat, and long of end point FORMAT (12HOPROFILE NO., A5, 2X, 6HSPEED, F5. 1, 2X, 7HCOURSE=, F5.O, 2X, 7HPOfl\TT 1, 2X, 312, F5.O, F6. 1, F5.O, F6.1, 2X, 7HPOINT 2, ZX, 312, F5.O, F6.l,F5.O, F6. 1)

3. Gravity Station Data a) Year and month b) DTG - Day, hour, mm c) Depth in meters d)Latitude and longitude of station e) Gravity meter reading (meter units) f) Eotvos Correction g) Normal (theoretical) Gravity h) Meter reading in milligals i) Gravity corrected for Eotvos effect j) Gravity corrected for Eotvos and vertical nonlinearity k) Browne correction 1) Vertical correction m) Free air anomaly [j) - g)J n) Quality (poor or blank) o) Averaged (filtered) free air anomaly [either 1/3 (FAA1...j+FAAm+FAA1+1) or 1/4 (FAA.j+ 2FAA+FAAm+i)] Outputs 1, 2, and 3 are standard outputs to high speed line printer. Optional outputs are as follows: 1 54

4. Even l0's free air anomaly output on line printer gives each even multiple of 10 mgal free air anomaly interpolated between measurement stations. a) Interpolated free air anomaly b)Latitude and longitude FORMAT (1H, 20X, Fl0.1, F5.0, F6.l, Fl0.0, F6.l)

5. Punched card #1 Col1-10Year, month, day, hour, mm 11-18Latitude 19-28Longitude 29-34Depth (meters) 35-43Observed gravity 44-52Normal gravity 53-57Free air anomaly (58-62Simple Bouguer never )63-68Terrain corrected Bouguer used Isostatic anomaly 1 69-73 ¼75-76Reliability 77-80Date effective

6. Punched Card #2 Col 1-10Year, month, day, hour, mm 11-16Latitude 17-23Longitude 24-2 7Ship (4 letter code) 28-31Base station (4 letter code) 32-37Base gravity value (minus 900 gals) 38-42Base meter reading 43-48Zero gravity (minus 900 gals) 49-51Drift 52-54Course 55-57Speed 58-62Eotvos correction 63-67Meter reading 68-74Eotv.os corrected gravity 75-77Browne correction 78-80Vertical correction The program listing is: PPAN1 MATNX 155 "IN A,KK nT1Fr'lsI'II Y(5fl) .A(.l) '.j''N1.fl ..'JAIc'!'Y1L41 ?- Fc'1Ar(A4) 17fl FPAAF(2 .U,F3.Q,F5.1 ,)i() A' r3 P4AI(1pH0pqFTIF I7'C1'T j,2X,T?,F.,F.1 41 F. 1 1) I A7 (7X,3F?.fl,F7.1,c.(.! x,A4.T I 7 FAT (1H ,ix,A4.1X,T) ,'.().fc.1 ,I. i

1 4 10.1 ,FS .0' FQ 1, F9 1 y , A4 FQAT(1H ,lx,A4,1x,3J? I 4FIo.1,F5o,F.I,Fq.1,7.A4,F7.1,1 F'AAT(,F1r1.1 ,1X,A4,1X,A4,I,A4,1X,4,3*,1,',F., 7' c0 F1AT(1RHIFSIJITS FFFCTTV'X.A4,,1H/EC I ?1i 3ASFT 'FtFP kFaDINr,F7.1 ,16Hsc 1 I, 2 H fl F1=,fS.1,,X,4I-IPC,%7) 1'4 F'A1(241TrAI ANMCL'Y AVEAfZ,V1,.1/2 iT"(p( "r TscT' 1MS ,FO) 'O FO'lAT(F1fl.0) Aflt1'7O) Y Fr(1,,00) IPC F V TXy = ciiM = j = F T'I S = 0. T F 3 1'1 ? 7 PA1,1O/IIF Ir0 PFAD(l'g) JF (MP)400,40l ,40() 4r TAVF=SU// lITF(3,1fl4)T\',7 STP 4rfl FXP(..T/C) IMp*00149. fl=T .49

3 PFAr)(1,1?'))TA1 ,TA?,TA.11 ,,2,A?ri'i r?,T-1 ,Ti',T,4I ' '. 1 ki C1,L3?,P TF(TAI )403,100.403 43 A[ ATAL1*60.+Ai2 Al N(AI 1*60. .AL2 RI ATkL1*60,+H LN(3i '60.PLC? .TA2.TA3ThO. TRTF1*,4. +Y,T3/(0. V!AT(ALAT4LAT)*2.9qF_(6/2. Afl IF (ALAIRLAT) 1(1,19,10 19 IF (ALCHLN(1 11., 12,13 11 C15TF r, r c() r,129) 11 C=r.5*PTF 10 TFTA=ATAN(DFPA/(ALAT, T)( IF (THF1P)2l,3,23 'I IF (ALAi_PLAfl?4,?9O,?, '4 CARS(THFTA).2.O*°IFril00 f, CARS(TkETA)+PTF (3y30O '3 TFiALAtRLAfl?7,fl,2 '7 (=T-WT (. TO 3r 156 C) AS(FiFT4).PTF r,Tr3OO ?01DT=iV1c(DFPAk) r,CT3O1 3o .1 pFFn=uS1/lTk-TA) 1 1 41 = 1 A 1 I I 471 1141=141 TI =1 TTP2T7

F (=C * A 0 wPTTF(3jl)l)PN,cPEFr),cflFr,TTA1 ,tT",1TA1.AL1 ,AL,,ALrl,AL)', I TT11 , LIR?, IT1.4L) .L?'q1C r? 2

I ( , AMAr 1X1X1*?4.+1X2.1X3/60. I TX'l =TX T TX ? TX? ITXITX1

TXE 1=XLT/60. XL 1 = I XLI XL2XLAT_XL1*,fl. I XL? = XL? * 10. X'1X CNG/60. XL = I X Cl XLC2=XLrrs4t-XLC1*6O. +0.Oc I XL 0? XL C? * tO. TF5CALF.E:Q.0. 1000,1001 1001 FASTCSTN(C)*SPFED P4T=XLA1*2.9OMR2F04 P.4I72.*PHT FOTCR7.5CCS(P-4T )*EASTC 1F(KK.EQ.l)SO,l qC) IF(TX-IyY.LT.O.11#,7171,lr. 1 TF(ARS(rV).LI.O.1)7(,,77 77 Fv=(ECTCR_ETVCS)*EXPTA_TX)*./r,C) FCTCREoTCP_EV 76 TF(KTFST.EQ,1)7,84 7S FTVCSETCP TX TA 4 T4FCG=978O49.*(J.,,OO524*(cT,(r1))**?_c.9E_()c*(STN(1I4T,),**?) J=rALE.O,O14q PJ,49 CZECc,,Y(J),(SCALE*0.fl1_R)*tvtj+1)_V(.J)) C C P c, PC R , P + ECT CP VEPT-EP(FRFQ,AMP,P) IF ( ik ,E.r,l) 71, 70 70 VEPT=VT_CVERT_VEPT2)*FxP((TXxX_TX)*M)./QC) 71 VT2VcRT TXXXTX IF ARC Wt,GT .400, CR, ARS ( VFRT). tT.1S.} 72 7 72 ItIJALzIDCCR 73 TRt6RgCfPC,R+VPT FANOM:TPUGR-THFCG S UM= SUM F A NCM 77+1. 1000 CALL VELCCITV(OFP,XLAT) flPgDP*1 ,R2AP Ay4(rjFT5.o*AMA( IF fl(K.E0.fl700,701 7r0 KKK=KK 1=1 157 Jjj=o 7fl7 K(IIAL(IpIJ4L YM(T)YOMC JtJP'iK(1'1 )ITX1 JIJNI

A(T,4)XLCI - A(T,5)aYLC2 ..------TF(SCALF.E(3.O.) 1001.1004 1003 ACT ,6)A (I,7):A(I,g)=A(r,q)=A(T .10)=A(I.11 )A(1;-12)*o. A(T,13)AC1,14)=0, G IC 1005 0n4 A(T,$)ccALE: A(T,7)FCTCR A CT ,R) IHECG AC 1,9) BGR ACT ,!0) .CCRGR

ACT,11)TRUGP -

A(1,13)2VERT ... - -- ACT, 14)FANCM

IOnS ACT,15)XLAT . ------A CT ,18),,TX A(I,16)rXLCPIG ----- ACT, i9 .AMAGNET GCIC(800,60,$10)I ------r,0 WRIT 3,7)yMCl,cJNKC1,K,K=l.3,A(1,J) ,J1,14) ,KOUAL(1) ,A(i,IQ) GCIC0 -- ...... -- ...... __ 7n1 IF(KKK.FQ.1)70?,704 7n2 KKKO I GCTC707 74 1z3 GT707 BiO TF(A(3.1R)-A(2,IR) .GI.O.fl7.Cp.A(?,1-A(1,18).Gr,fl.07)900,Q01 9n0 WRIT 3,7)vMC2,cJuNK2,,(,K1,3,A2,J.j1,j L,YC.4tI9.1 JJJ=0 (01C902 C WFIGHTEJ) AVFPAGE 9o1 A(?,17)u(A(1,j4).,*A(2,4).AC3,1.4))/4. - _____ IF (15.E0,O)452,453 453 KKaJJJJJJ.1 ______CALLTNTFR

1 ,A(2.q)

...?IF(K.TEST.EQ.1)8SO.851 . - ______850 WRITC3,7)YM(3),CJUNK(3,K),KI,3),(A(3,J),J1,l4),K0UAL(3),A(.iQ) 851 oCR20L1.19 ______DR20Mal,2 80 A(M,L)aA(M+1,L) ______DCR21NI .2

8I JUNK(N.T)JtJNKCW.1,!) 00R22KZ1 ,2 Kr3UAL(K) KQ(JAL (I(,1) 822YM(K)sVp4(X.1) A° IF(KTEST.EQJ)9,2 FtJPJCTTCI ER(FRFt,AMP, ) t5 C vETT(AL EopCR AS ADJUSTED 7/S/A7 CC.6OE-Q5 .....-. PCC*AMP*AMP*FPEQ*FRF*(13.4. RETURN F Mn

- --- -.- .___ .. SLJRCIJTTNE INTER 159 CMMCN A,KK OTMENSIeNA(3,19) 1Or0 FPMAT(1H ,2OX,F10.1,F5.O,Ff,.,F1fl,o,FA.1) FA!\IOM=Ac2,17) XL A1*A C;, 15)

XLNG5A(2,j6) ------______IF (l(PC.E0,1)300,301 3r'l IFA=FANM/10. JFA*pA2,1O, K=IABSyFA-JFA) IF (ARS(FANOM-F4?) .GT.ABS (FAN3M.FA?) ) 30P,303 3o (K.1 3n3 IF(K.EQ.0.3R.K.r,T.10)300,04 3n4nC2001a1 ,K tFCFA2.T.FANCM) 310,305 35 TFCFA2.LT.O,)3Q,3O7 3rF, A?JM=(JFA,j-1).1O r,010315 ____ 3o7 APJHCJFA,)*1 G3T0315 310 TF(FA2.IT.0.)311,312 3i1 ANOMJFA-fle10 (CT315 32 AJOM(JFA-I+1)*1Q ------3iS OELTAsARS( (ANCM-FA2RFANM-FA,fl

CLAr.XLAT2+(XLAT-XLAT2)*DELTA - CL3N(IXLCN2+ (XLCNG-XLON2, *DELT4 ILAT!CL&T/60, DLAT*ILAT CMIN!CLAT-I)LAT!AQ, I LNG.CLONG/6O PjGRIJCNG CLMI WsC,.CNG-DLCNG60, 2ñ0 WRTTEC3i1000)ANCMLAT,CMTN,flLCNr,.CLMtN 3o0 F42.FAN,M XLAT2.XLAT - xXLN RTUPN______P

StIRR3)1TNE vECITY(DEP,xIAT r SrUNr) VFLOCyTY CORRFCTICNS FP DFP1'4S r rCRPECIICNc FORNORTHEAST PAC!f1._.. TF(PEP.i T.IOO.)1.2 ?

1 RFIIJRN F Nfl 161

APPENDIX 2

SUPPLEMENTARY GRAVITY DATA

The bathymetry of region of the northeast Pacific Ocean west of Washington and British Columbia exhibits large topographic contrasts between physiographic provinces similar to those observable in the diverse geologic structures of the adjoining continental coastal regions.The greatest problem in studying this region gravimetrically was obtaining an areal density of measurements sufficient to allow detection and analysis of the major structures of the region.Surface ship gravity measurements along the track lines indicated in Figure 5 provided approximately 15, 400 data points.However, these points are concentrated along track lines leaving large unsurveyed areas which could contain major structural features.Additional valuable data were obtained from the literature and from personal files. Harrison etal. (l957) reported 10 pendulum measurements obtained in a submarine west of the Columbia River and seven measurements southwest of Vancouver Island.He estimated an rms uncertainty of ± 3.8 mgal for these measurements. Worzel (1965) has reported the values of 28 pendulum measurements made in sub- marines in the area west of Washington and British Columbia.He estimated a rms uncertainty of ± 3. 6 mgal for these measurements. 162 These predominantly deep ocean pendulum measurements provided data between track lines and also an independent check of the surface ship gravity meter measurements. The location of the pendulum stations of Harrisonetal. (1957) and Worzel (1965) are shown on Figure 5. Approximately 280 underwater gravimeter measurements were made available by the gravity Division of the Dominion Observatory of Canada (Stacy, 1967 personal communication). No estimation of the accuracy of these measurements has been reported; however these underwater measurements are expected to be three to five times more accurate than the available surface-ship data.The underwater meter measurements are spaced at 10-15 km intervals in portions of Dixon Entrance, Hecate Strait and Queen Charlotte Sound. The areas in which these measurements were made are indicated in Figure 5 by hatchure lirie. The U. S. Coast and Geodetic Survey check the performance of their surface-ship gravity meter at sea by retracing a previously measured track line called asignature line." The USC&GS (Ericson, 1967 personal communication) made available approximately 90 gravity measurements along a signature line off Cape Flattery, Washington. Considerable assistance in extending gravity contours shoreward was obtained by using available land gravity measurements made along 163 shore lines.Land gravity data in the Queen Charlotte Islands and along some of the fjords of British Columbia was made available by the Gravity Division of the Dominion Observatory of Canada (Stacy, 1967 personal communication). A Bouguer gravity map of southern British Columbia by Walcott (1967) was of considerable help in defining the gravity anomalies along the continental shelf off Vancouver Is land.The sea values west of Washington were joined to the contours of an unpublished Bouguer anomaly map by Sauers

(1966).The rms uncertainties of land gravity measurements are near ± 0. 3 mgal; approximately an order of magnitude betterthan sea gravity measurements. In contouring the reduced measurements to form Figures 6 and 7 two considerations were applied:1) the gravity field is a potential field and is therefore everywhere smooth and continuous, 2) the high density contrast between the water and the material of the ocean bottom cause gravity gradients to follow bathymetric gradients.The first consideration indicates that when measurements are made at a distance from the anomaly producing source (i. e. in deep water) gravity gradients are less.This suggests that both limited smoothing and sympathetic contouring of scattered data generally result in the best graphic approximation of the gravity field.The second consideration suggests that in continuing a gravity anomaly contour from a point defined by measurements to a second similar point, the most 164 likely path follows the isobath or isobathic trend between the two points.This is not always possible in detail but generally the gravity anomaly trends may be considered coincident with bathymetric trends. In some areas, however, bathymetric contours are defined by insufficient data and both the gravity and bathymetry outline only the largest of features. In certain areas, although the bathymetry is reasonably well known, the gravity field is unknown and cannot reasonably be extrapolated from available data.Most notable are areas containing seamounts which have closed contours.Ships tracks usually do not pass over searnount peaks, and therefore anumber of closed contours remain which are not tied to any gravity measurement. To help define gravity contours about seamounts the gravity anomaly value over the seamount summit was calculated for 44 seamounts in the region mapped in Figures 6 and 7.Seamounts for which calculations were made were selected on the basis of size, symmetry and distance from areas in which the free-air gravity anomaly about the base of the seamount would be greatly different from zero. The equation, from which the expected anomaly may be determined, is based on the assumption that seamounts may be approximated by a cone structure.The dimensions of the cone are indicated in Figure Z3 and the derivation of the equation is as follows: 165

Figure 23.Diagram illustrating equation parameters for computing gravity over a cone structure. For the solid cone in Figure 23 the vertical acceleration of gravity at point o is given by the equation

Zir R D+h zr dzdrdO (1) = S 2 3/2 ooD(r+ z where D apex depth

h cone height

a =radius of cone base gravitational constant

p = cone density and R = (z-D)

Integrating with respect toe yields

R D+h zr g =2 rr dz dr (2) z SS 2 2 3/2 oD(r +z 166 Integrating with respect toryields

D+h z a1 = _Idz hSa [{[i]h z -2Dz+D2 21/2} hj a (3) letting h2 c [(i;) +1]

Equation (3) reduces to

D+h zdz D+h a dz (4) 2 $D (cz -2Dz+D21/2-S

Integrating with respect to z and simplifying yields

= 2v [{DZ(c1)}2 {D2(c1)+2Dh(ci)+ch2}l/2]

1/2 2 1/2 D D(c-l)+c [D (c-i)] +a (5) +ln 1/2 2 21/2 c D(c_1)i-ch+c [D (c_.I)-i-2Dh(c_1)+ch]

Equation (5) was written in Fortran IV with minor adaptations for running on a CDC 3300 computer.

The gravity anomaly expected at point o of Figure 23,i. e. above the seamount summit, depends on the average density of the material within the conic structure.Nayudu (1962) and Engel and Engel (1963) have reported the density of some volcanic rocks obtained by dredging in this region.The densities of these surface 167 rocks, however, cannot necessarily be expected to represent the average density of seamounts of thisregion. Shurbet and Worzel (1955) used a density of 2.84gm/cm3in approximating the free-air gravity anomaly observed over two seamounts in the Gulf of Maine.Dehlinger etal. (1967a)also noted that a density of 2.84gm/cm3could best remove the effect of topography in a two-dimensional analysis in the vicinity of Cobb seamount. Harrison and Brisbin(1959), in conducting a geophysical investigation of Jasper seamount found the observed gravity anomaly could best be approximated using a density of 2. 3gm/cm3. LePichon and Talwani (1964) in investigating a seamount in the north Atlantic Ocean discuss the possible densities to be assumed in forming a Bouguer anomaly which exhibits a minimum relief over the seamount. They suggest a range between 2. 3 and 2.9gm/cm3and use 2. 6gm/cm3 in their analysis of possible seamount structures. Further analysis (Woollard, 1951, McKay, 1966) indicates that to actually fit the observed gravity anomaly with a model a structure and substructure must be assumed for the seamount which consists of masses having large variations in density. In particular seamount slope densities are typically light; 2. 0 to

2. 3gm/cm3while core densities must be heavy; 3. 0 to 3. 3gm/cm3. No attempt was made to generate a seamount model which would give the expected anomaly over the seamount.A cone model having a uniform average density was assumed to enable estimatingthe gravity anomaly to be expected over the peak of the seamount.In this respect, the flank and basal anomalies of seamounts, shown in Figures 6 and 7, are at least partly defined by actual measurement and only the maximum anomaly is dependent on calculated values. To arrive at the best average density for seamounts in the region west of Washington and British Columbia, calculated values were compared with measured values for Cobb and Union seamounts. Cobb and Union seamounts were the only two seamounts where measurements were obtained over or nearly over the seamount peak. The density of best average fit was 2. 43gm/cm3 (ora density contrast between rock and water of 1.4gm/cm3).Peak anomaly values of 44 seamounts were calculated using this density.The dimensions and resulting anomalies of these seamounts are given in Table 3 and their location is shown by the x's of Figure 5.These calculated values were used as an aid in constructing the free-air gravity anomaly contours of Figures 6 and 7. Linear interpolation between gravity measurements was used in areas where no other control was available. pkA1 SFAMNT j-9 mo FOPMAT(14) iril FOPMAT(4F10.2) 2n0 rPMAT(cqH GRAVITY CORRECTION P SEAMOUNT USTN r CCNE AQ!T TON///6I4 DEPTH (F'MS) HEII4T (FMS) PASE (T) RHO FT)O, (M(4L)j) 2o1 FOPMAT(1H PRI'T 2rO REA( jOr,N PTF3.14159 CA.67 COJST2.*PIE*CA I I IT4I READ lOi,AD,Apl,Aw,RHO WXAW*j .53 184 HAH*,OrjR2RR DsAr*,Orj 8288 AW/2. CaH*H/(*A)+l. F*f)* (Ce.) FjPST(j,/C)*( (D*F)*.5-E) x4uM=F+(c*D*F)** .5 XDEN5Fr*H+E(C**.5) XSECAL'G (XNIJM/XDEN) SECONDD/(C**CP) )X5EC Ga C (FIRST.SECCMF+A) PRINT 21,AD,AH,AW,PHC,G (yN) i,2,2 2 CNTINUF FN 3200 rCRTRAN F)!AGNOSTIC RESULTS - FOR SEANT 170

Table 3.Seamount locations, dimensions and gravity

Base Apex Location Diameter Height Depth Gravity No. Name N. Lat. W. Long. (nt. nil.) (lath) (lath) (mgl)

1 Thompson 4602.5 12838.0 5.4 686 764 24 2 Bear 4602.5 13026.5 11.8 1220 280 85 3 4622.0 13037.5 10.8 1120 380 72 4 4631.5 13058.0 12.0 800 600 52 5 4641.8 13120.4 20.0 1505 195 124 6 Cobb 4645.6 13049.0 20.0 1687 13 150 7 4713.0 13129.5 10.0 740 880 38 8 4721.0 13144.0 10.0 560 1090 27 9 4733.9 13019.8 9.5 850 670 46 10 4754.0 12939.0 9.0 595 855 30 11 4754.4 13031.2 7.0 687 863 28 12 4757.0 13158.0 12.0 1000 660 60 13 4758.9 12956.0 8.0 570 880 26 14 4804.1 13011.8 9.0 920 580 50 15 Warwick 4804. 1 13248.6 20.0 1520 280 121 16 4809.0 13025.2 4.0 640 880 15 17 4816.3 13310.4 18.0 1460 390 108 18 4820.0 12914. 5 7. 5 675 625 34 19 Springfield 4822.0 12954. 5 9. 1 626 874 31 20 4824.5 12923.3 11.0 835 565 52 21 Eickelberg 4830.8 13309.0 18.0 1420 432 104 22 4831.5 12937.0 11.0 885 565 55 23 Heck 4838.4 13008.0 8.0 720 780 34 24 4845.0 13148.0 8.6 610 990 28 25 4855.9 13352.4 17.6 1405 495 100 26 4907.9 13217.5 15.0 700 937 44 27 4919.0 13350.0 18.0 880 919 59 28 Union 4932.5 13241.6 18.0 1570 158 127 29 4946.0 13146.0 10.0 861 789 45 30 4948.8 13157.0 7.0 708 942 27 31 4949.0 13327.0 14.0 1050 690 67 32 4954.0 13206.5 8.0 750 745 37 33 5019.5 13127.5 6.5 700 859 27 34 5024.5 13213.0 5.6 853 747 29 35 5035.8 13042.5 10.4 910 490 57 36 5039.5 13106.5 8.0 656 844 30 37 5044.6 13116.5 6.3 745 655 32 38 5045.5 13052.8 9.2 1077 293 69 39 Graham 5222. 1 13406.0 12.0 1010 490 66 40 Bowie 5318.0 13542.0 19.0 1626 24 143 41 Hodgkins 5330. 5 13559. 5 15. 3 1318 432 93 42 Denson 5407.5 13745.0 8.0 1051 599 50 43 Dickens 5432. 8 13658. 4 18. 3 1340 260 108 44 Brown 5453.5 13836.0 7.0 788 812 32 171

APPENDIX 3

TWO-DIMENSTIONAL GRAVITY COMPUTATIONS

The perimeter of a two-dimensional structure can be approximated by an n-sided polygon. Taiwani, et al. (1959b) derived expressions for the vertical component of gravitational acceleration due to an n-sided polygon at any point.These authors obtained for the vertical acceleration at a point P due to the attrac-. tion of the polygon in Figure 2.4.

P -.< x

Figure 24.Polygon illustrating equatiopanameters for. two-dirensicai gravitycomputatthn 172

V = 2Gp z

where G is the gravitational constant, p is the density, and

cos 0, (tan0.-tan.) 1 1 1 z. 1 1 1 1 e cosO (tan 0 -tan)] j+l i+l 1

z. O,tan-1 -1 1 X. 1

z. -z. -1 1+1 1 4.tan -x. 1+1 1

1 0 = tan i+l X1 r i+l iJ J

a.X I 1 i+l i+lI z -z L

The expression forZ.may be simplified as the following cases:

CaseA: IfX.=O 1

z = -a1 sin4cos [0+_ }] 173

Case B:lix. 0 1+1

Z. = a. Sin4 Cos4 [. +

tan4. log {COs 8. (tan 0. - tan4 )} ] 1 e 1 1 1

CaseC: IfZ.Z. 1 i+1

Z. = Z.(O. - 0.) 1 1 i+1 1

Case D:If X. X. 1 i+l

Cos 0 Z. X. log 1 1 e CosO.1+1

Case E:If 0. 1

Z. 0 1

Case F:If X. = Z. = 0 1 1

z.= 0 1

Z. 0 CaseG: IfX.i+1 i+1 z.=0 1 174

The computer program for two-dimensional cross sections which uses the method of Taiwanietal. (1959b) computes the attractions along a profile of a two-dimensional polygonal body.The program, written in FORTRAN IV was used on a CDC 3300 computer.The program was initially written by Mr. W. A. Rinehart and extended and modified by Dr. E. F. Chiburis and Mr. M. Gemperle.The program parameters are: 175 Input Cards 1) General Control a) Cross section identification b)First location where gravity is to be computed c) Kilometers between points of gravity computation d)Total number of points where gravity is to be computed e) Number of blocks in model f) Code to indicate if model check is to be performed (blank - no check; any characters - check) FORMAT (A8, F6.O, F4.O, 215, A3)

2)First card of each block a) Density of block b) Number of corners in block c) Block number FORMAT (F5.O, 13, A8) 3) Corner of block (polygon) a) X-coordinate of corner b) Z-coordinate of corner c)block identification 4) Last card of deck - If model check is specified, a card is required at end of the deck giving limits of model.

Output (High speed line printer)

1)Profile identification FORMAT (1H1, A8l1) 2) Computed gravity a) X-location b) Computed gravity c) Computed gravity of point minus computed gravity at first point using density of 1.0 for all blocks.This serves as a check to make sure no gaps (equivalent to zerodensity) or overlaps (equivalent to doubledensity) exist in model, and helps to locate errors when they occur.All values should be zero except for edge effects. FORMAT (1H, Fl0.l, F15. 2,ZOX, F17.4) 176 3) Model check - residual at three station points is computed. This residual should be zero if the data has been read correctly. FORMAt (28H (model-check),Station 1 ,F8. 2/ 28H (model-check), Station 2 , F8. 2 / 28H (model-check), Station 3 , F8. 2!!')

4)Listing of Block Densities and block numbers FORMAl' (1H, A8, F1O.2) 5) Block identification error - output occurs if block identification is not consistent in input data. FORMAT (41H COORDINATE IDENTIFICATION ERROR IN CASE, A5) The program listing is: 177 pp'r,pAM TAL.wANT rTMFNSIpJ PX(?),x(200),7(2),TCT(2OO),CMK1(2flfl),K2(2rr DT'lFNStN CHV.3(?O0) ,A10C200) ,I'7(PflO) ,TXT(200) 1 FCPP.lAT(6R,F6.O,F4.O,2I,A3) 2 FCPMAT(FS,fl,13,AR) 900 F°MAT(1H,AR,FI0.2) 3fl1 FDMAT(2FR.n,Ac, 35 F1AT(41H COCQrTN4TE TDF'TTTeATIN ERPORIN Cr,A5) 3 FPiAT( BFR.r,) 11 FCPJ44T(4) 200 FMAT(IH ,FIO.1,F15.2.2r)X,F17.4) 21 FcMAr(1H1,AR//) 2n5 FQ4AT (2811 qMO0EL-CHFCK),STATTN la,FR.2/28H (MflFLCHECK) ,STATT 2,FR., / 2I4 3,F&,2 // P1=3. 1'597 100 p4j) l,SAM,PX(l),STEPqM,M,MCWI( IF (M.E.0) øi ,7? 7o1 ST° 7')2 PRTNT2O1 ,SAJ4 TTT (1) r. TT C1)0 flrl12 M PX (1) =Px (I-i) +STEP P7 (J)0, ITT I T)C). o IT(I)=.0 JJj=0 52 1061'4=),M IF (JJJ.0.l)50,703 703 pEAfl2,4',J,S4LT PPT1T 9r)0,SALT,PHQ r 300 T*l,j 3ooPFAp3o1,xn ,Zn oc 32 y.2,j TFCAII)(1) .1E.At0(T-1) )700,302 7o PRTPT305,AT0(1 302CNTTNUF 50 KJ+1 X(K)X() 7(K) zZ ( SUM.0 f)D101I1,M DOSJU1 'K S X(J).X(J)-PX(1) OCl?2Jl ,K IF(X (J) )6,7,8 6 rM.ATANZJ/XCJ).P IF(PI-TH-.00000I)47,4-7,9 47 114-PT 7 TH.5P1 -601C $ r14.ATAN(Z(J)/X(J)) Fcn4-.00O001)45,45,q 178

45 T'40 9 TF(J.NE.l)l0,704 704 T141=TH GC T 17 10 Ti42=TH LJ-1 A7(J)7(L) RX(J)x(L) JJ= IF(X(L).F'.O.)70S,706 JJi 7(Th IF(*(J),E.0.)707.708 o7JJ=2 70R TF(A,EQ.0.)709,710 JJ=3 710 IF(P.E(.0.)711.712 ii JJ4 712 TF(T41.JE.TH2)19,713 7i s=0. r,r102 19 IF(X(L),NE,0.)??,714 714 IF(7(L).NE.0.)2P,715 715 S0. r,rc 102 ? TF(y(J).NE.O,)pc,716 716TF(7CJ).NE.O.)2c,717 7i S. 102 25 GCT(2b,27,28,2q,3cfl,JJ 26 S.(X(J)*X(J)*Z(L)/(A*A+X(J)*(J)))*(TH2,5*PIe(A/X(J) )0(AL(Z(L)) 1-.5*ALO( (XLI)X LI) .Z (J)* (J) 0C10102 27 S(X(L)*X(L)*Z(J)/(A*A+X(L)0X()))*(TH1.5*PI.(A/X(L))*(AL(Z(J)) 1..S*ALOX(L)*XL),Z(L)*nL) S-S G31C102 2 S7(L)*(TH2_TH1) GTQ102

29 VRSQRT C (Z (J)Z CM +X (J) .X Cj) ) / CZ(L) Z CL) .X (L 'X L) ) SX(t_.)*ALOG((XCL)/X(J))*V) GOT1O2 - O TX(J)(7(J)*R/A) u=R*4/ (qeB,A*4) V=71.T2

SzT*tJ*(V.W) 102 SUMSUM+S TH1T2 122CCNTTN'k DQ1O3J1,K - 13 X(J)X(j)+PX(j) TF(MCI4K.EQ.0)101,718 718 IF(T.NE,l)60,710 79 CHK1(IN)=SIJM 179

TF(.JjJ.F().1)10J,Afl 0 IF IT .JtM/2) ,1 ,7?O ?O CK2(I'l)S1JM fl T (i.NE.1)(,2,77 7?1 CIrSIiM F? I S'le13.34 111(1) 111 (I) 1)UM

StiMSUM*4)1*1 3e TCT (I) TCT () +M 101 StJM.() 106 rJTTNtJr IF (,JJJ.FT.0)53,72? 7? !C 1000 T=?,M 1000 ITT (Kr) T11 (KT )-TIT (1) TI! (1) O. PP1T00,(PX(T),TT(I),TTT(T),T1,M) IF (MCHK.FQ,O) lnn,723 723 JJJ1 TC-4K10. TCl4P2O. TC'4K30. I)51 Izi ,N TCHK1=TCHKI+CHK1 (I) TCNK2TCHK2+CNK7 IT) 51 TC3T(HK3+CHKI(I) 307 PF4n3 (xU ,Z(fl ,I1,4) N I J=4 53 PsTCH(ICHK1(1) R2=THK-CHK2 (1 1 P3TCHK3-CHK3 Ill pTNT20c,w1 ,P2,p3 (T'o0