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Crustal structure from an ocean bottom seismometer survey of the Nootka fault zone Au, C.Y. Daniel 1981

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CRUSTAL STRUCTURE FROM AN OCEAN BOTTOM SEISMOMETER SURVEY OF THE NOOTKA FAULT ZONE by C. Y. Danie l Au M. Sc. ( P h y s i c s ) , U n i v e r s i t y of A l b e r t a , 1977 B. Sc. (Hons. P h y s i c s ) , U n i v e r s i t y of A l b e r t a , 1975 A THESIS SUBMITTED IN PARTIAL FULFILMENT OF THE REQUIREMENTS FOR THE DEGREE OF DOCTOR OF PHILOSOPHY i n THE FACULTY OF GRADUATE STUDIES (Department of Geophysics and Astronomy) accept t h i s t h e s i s as conforming to the req u i r e d standard The U n i v e r s i t y of B r i t i s h Columbia June, 1981 (c) C. Y. D a n i e l Au, 1981 I n p r e s e n t i n g t h i s t h e s i s i n p a r t i a l f u l f i l m e n t o f t h e r e q u i r e m e n t s f o r a n a d v a n c e d d e g r e e a t t h e U n i v e r s i t y o f B r i t i s h C o l u m b i a , I a g r e e t h a t t h e L i b r a r y s h a l l m a k e i t f r e e l y a v a i l a b l e f o r r e f e r e n c e a n d s t u d y . I f u r t h e r a g r e e t h a t p e r m i s s i o n f o r e x t e n s i v e c o p y i n g o f t h i s t h e s i s f o r s c h o l a r l y p u r p o s e s m a y b e g r a n t e d b y t h e h e a d o f m y d e p a r t m e n t o r b y h i s o r h e r r e p r e s e n t a t i v e s . I t i s u n d e r s t o o d t h a t c o p y i n g o r p u b l i c a t i o n o f t h i s t h e s i s f o r f i n a n c i a l g a i n s h a l l n o t b e a l l o w e d w i t h o u t m y w r i t t e n p e r m i s s i o n . D e p a r t m e n t o f t S g f r ^ H V S ( C 5 ftH!> P > S T t o i ^ o K Y T h e U n i v e r s i t y o f B r i t i s h C o l u m b i a 2 0 7 5 W e s b r o o k P l a c e V a n c o u v e r , C a n a d a V 6 T 1W5 D a t e Qk<** <9 . W DF-6 ( 2 / 7 9 1 i i ABSTRACT The Nootka f a u l t zone i s the boundary between the small E x p l o r e r and Juan de Fuca p l a t e s which l i e o f f western Canada between the America and P a c i f i c p l a t e s . To i n v e s t i g a t e the c r u s t a l s t r u c t u r e i n the r e g i o n , three r e f r a c t i o n l i n e s were shot with e x p l o s i v e s and a l a r g e airgun i n t o three ocean bottom seismometers (OBS's) each equipped with three-component geophone assemblies. The P and S wave p r o f i l e data are analysed p r i m a r i l y by s y n t h e t i c seismogram modelling using the WKBJ alg o r i t h m . The i n t e r p r e t a t i o n gives r e l a t i v e l y c o n s i s t e n t r e s u l t s f o r the upper c r u s t . Below the 1 km t h i c k sediment l a y e r , the P wave v e l o c i t y ranges from 3-7 to 4.7 km/s and incr e a s e s w i t h depth at a moderate v e l o c i t y gradient of 0.5 km/s/km to a depth of 1.9 km. A zone of high v e l o c i t y gradient marks the t r a n s i t i o n from l a y e r 2A to 2B, below which the v e l o c i t y again increases moderately w i t h depth to a range o f 6.0 to 6.4 km/s at the base of l a y e r 2B. The S wave v e l o c i t y a l s o increases w i t h depth i n the upper c r u s t but no d e t a i l i s a v a i l a b l e f o r the v s s t r u c t u r e i n the shallow region of l a y e r 2A. W i t h i n l a y e r 2B, the P and S wave v e l o c i t y values give a Poisson's r a t i o w i t h an unusually low value of 0.24. This may be r e l a t e d to the presence of q u a r t z - r i c h trondhjemites. Both v P and v s are r e l a t i v e l y uniform at the top of l a y e r 3A. Good agreement i s found between the seismic v e l o c i t i e s of l a y e r 3A and the seismic v e l o c i t i e s of rock samples from the corresponding depths of o p h i o l i t e complexes measured i n the la b o a t o r y , r e s u l t i n g i n c o n s i s t e n t values f o r Poisson's r a t i o . Sub-bottom c r u s t a l thickness v a r i e s from 6.4 to 11.2 km among the i i i v a r i o u s p r o f i l e s . Some aspects of t h i s v a r i a t i o n can be explained by co n s i d e r a t i o n of a recent t e c t o n i c model f o r the development of the f a u l t zone. This r e q u i r e s , w i t h i n the past one m i l l i o n years, v a r i a t i o n i n the process of c r u s t a l formation at the r i d g e , o r u s t a l 'maturing', or both. The abnormally t h i c k c r u s t may r e s u l t i n part from the complex i n t e r a c t i o n of the Juan de Fuca and Ex p l o r e r p l a t e s w i t h the l a r g e r and older America and P a c i f i c p l a t e s . Apparent v e l o c i t i e s of compressional waves i n the upper mantle, c a l c u l a t e d f o r a r e a l l y d i s t r i b u t e d ray paths, show s i g n i f i c a n t a n i s o t r o p y . A maximum v e l o c i t y of 8 . 3 km/s i s found i n the i n f e r r e d d i r e c t i o n of p l a t e motion and a minimum v e l o c i t y of 7.5 km/s i s found p a r a l l e l to the spreading r i d g e . This type of v e l o c i t y v a r i a t i o n can be approximated by a mixture of 42% t r a n s v e r s e l y i s o t r o p i c o l i v i n e and an i s o t r o p i c m a t e r i a l with a constant v e l o c i t y of 7.0 km/s. V e l o c i t y of shear waves i n the upper mantle, v a r y i n g from 4.5 to 4.6 km/s, i s i s o t r o p i c w i t h i n the r e s o l u t i o n of the i n t e r p r e t a t i o n . This causes a prominent anisotropy i n the values o f Poisson's r a t i o . The v a r i a t i o n i n v e l o c i t y w i t h depth i n the c r u s t , values of Poisson's r a t i o , the P wave v e l o c i t y a n i s o t r o p y , and the approximately i s o t r o p i c S wave v e l o c i t y are co n s i s t e n t w i t h l a b o r a t o r y measurements on rock samples from o p h i o l i t e complexes. i v TABLE OF CONTENTS Abstract i i Table of Contents i v L i s t of Tables v i L i s t of Figures v i i Acknowledgements i x CHAPTER 1. INTRODUCTION 1 1.1 Tectonic S e t t i n g of Study Area 1 1.2 Outline of Study 3 CHAPTER 2 . ACQUISITION AND PROCESSING OF DATA 5 2.1 The Experiment 5 2.2 Instrumentation and Data D i g i t i z a t i o n 7 2.3 Range and Travel Time Corrections 8 CHAPTER 3 . P WAVE PROFILE INTERPRETATION 16 3.1 Record Sections 16 3.2 Inversion of Travel Time Data 27 3.3 Synthetic Seismograms and Amplitude Modelling 31 3.4 Results 35 3.5 Discussion 42 3.6 Concluding Remarks 58 CHAPTER 4 . INTERPRETATION OF SHEAR WAVE DATA 60 4.1 Introduction 60 4.2 Shear Wave Conversion 61 4.3 Optimization of Shear Wave Observations 63 4.4 Poisson's Ratio of Sediment 75 4.5 Record Sections and Amplitude Modelling 78 4.6 Results and Discussion 83 4.7 Concluding Remarks 99 V CHAPTER 5. A REAL DATA AND ANISOTROPY 101 5.1 Introduction 101 5.2 Areal Data 102 5.3 Interpretation of Areal Data 104 5.4 Discussion 108 CHAPTER 6. SUMMARY 114 References 116. Appendix 1: Linearized Inversion of Tau-P Data 125 Appendix 2: Inversion of P Wave Anisotropy 127 v i LIST OF TABLES Table Page 3.1 Bay of Islands O p h i o l i t e Complex 45 4.1 Poisson's r a t i o of the sediment layer 77 v i i LIST OF FIGURES Figure Page 1.1 Tectonic map of Canada's west coast 2 2.1 Location map of study area 6 2.2 Sound v e l o c i t y p r o f i l e of water column 9 2.3 Parameters f o r topographical c o r r e c t i o n s 12 2.4 P r o f i l e EX3 t r a v e l times 14 3.1 F o u r i e r power spectra of P phase and noise 17 3.2 Record s e c t i o n and s y n t h e t i c s e c t i o n of p r o f i l e EX1 18 3.3 Record s e c t i o n and s y n t h e t i c s e c t i o n of p r o f i l e EX3 19 3.4 Record s e c t i o n and s y n t h e t i c s e c t i o n of p r o f i l e EX3R 20 3.5 Record s e c t i o n and s y n t h e t i c s e c t i o n of p r o f i l e EX2 21 3.6 Record s e c t i o n and s y n t h e t i c s e c t i o n of p r o f i l e EX2R 22 3.7 Record s e c t i o n and s y n t h e t i c s e c t i o n of p r o f i l e AG2N 23 3.8 Record s e c t i o n and s y n t h e t i c s e c t i o n o f p r o f i l e AG3S 24 3.9 Record s e c t i o n and s y n t h e t i c s e c t i o n of p r o f i l e AG1S 25 3.10 C a l c u l a t i o n of tau(p) 29 3.11 V e l o c i t y - d e p t h models f o r a l l p r o f i l e s 36 3.12 Layer r e p r e s e n t a t i o n of v e l o c i t y - d e p t h models 38 3.13 Tectonic model of study area 43 3.14 Ray t r a c i n g and s y n t h e t i c s e c t i o n f o r p r o f i l e EX3 56 3.15 Ray t r a c i n g and s y n t h e t i c s e c t i o n f o r p r o f i l e EX3R 57 4.1 H o r i z o n t a l p r o j e c t i o n o f SV motion 65 4.2 Water wave p a r t i c l e motion p l o t s 67 4.3 Seismic traces f o r 30 kg shot 68 4.4 Seismic t r a c e s f o r 5 kg shot 70 v i i i 4.5 F o u r i e r power spectra f o r PSS and PPS phases 71 4.6 Three dimensional p a r t i c l e motion s e c t i o n f o r p r o f i l e EX3 . . . 73 4.7 Ray paths of seismic phases 74 4.8 SV record s e c t i o n and s y n t h e t i c s e c t i o n f o r p r o f i l e EX3 79 4 .9 SV record s e c t i o n and s y n t h e t i c s e c t i o n f o r p r o f i l e EX2 80 4.10 V e l o c i t y - d e p t h models f o r p r o f i l e s EX2 and EX3 84 4.11 Layer r e p r e s e n t a t i o n of v e l o c i t y - d e p t h models EX2 and EX3 ... 85 4.12 Poisson's r a t i o versus depth 87 4.13 vp-vs paths f o r the whole c r u s t 88 4.14 Vp-vg paths compared with o p h i o l i t e samples 89 4.15 Comparison of EX2 and EX3 w i t h FF2 and FF4 90 4.16 Amplitude modelling t r i a l s f o r EX2 97 5.1 Distance-azimuth d i s t r i b u t i o n 103 5.2 A r e a l v e l o c i t y data and model curves (I) 105 5.3 A r e a l v e l o c i t y data and model curves ( I I ) 109 ACKNOWLEDGEMENTS I wish to express my s i n c e r e g r a t i t u d e to Dr. R. M. Clowes, my t h e s i s s u p e r v i s o r , f o r h i s h e l p f u l guidance and encouragement during the period of t h i s study. His constant support and understanding are deeply appreciated. I acknowledge the cooperation and a s s i s t a n c e of the o f f i c e r s and crew of CFAV Endeavour during the f i e l d survey. C o l l a b o r a t i o n w i t h Dr. R. D. Hyndman of the P a c i f i c Geoscience Centre, EMR, Sidney, B. C , whose ocean bottom seismometers made the r e f r a c t i o n program p o s s i b l e i s g r e a t l y appreciated. L. Haynes and W. Nehring of the F l e e t D i v i n g U n i t , P a c i f i c Maritime Command, CFB Esquimalt, B. C. prepared and detonated the e x p l o s i v e charges. I thank C. H. Chapman f o r p r o v i d i n g a copy of h i s WKBJ s y n t h e t i c seismogram r o u t i n e and G. A. McMechan f o r h i s ray t r a c i n g / s y n t h e t i c seismogram a l g o r i t h m . F i n a n c i a l a s s i s t a n c e f o r t h i s p r o j e c t was provided by a Research Contract and Research Agreement Nos. 167-3-78, 110-3-79 and 172-3-80 from Energy, Mines and Resources (Earth Physics Branch) Canada and by operating grant A7707 from the N a t u r a l Sciences and Engineering Research Council Canada. A d d i t i o n a l f i n a n c i a l support was provided by S h e l l Canada Resources L i m i t e d and Mobil O i l Canada L i m i t e d . I acknowledge r e c e i p t of an H. R. MacMillan Family F e l l o w s h i p from the U n i v e r s i t y of B r i t i s h Columbia during the period of t h i s study. 1 1. INTRODUCTION 1.1 Tectonic S e t t i n g of Study Area The northeast P a c i f i c Ocean o f f the coast of western Canada has played a s i g n i f i c a n t r o l e i n the development of theories r e l a t e d to plate t e c t o n i c s . Since the o r i g i n a l proposals of Wilson (1965) and Vine and Wilson (1965) based on the magnetic anomaly map of Raff and Mason (1961), the tectonics of the region have been studied by many s c i e n t i s t s . The general configuration and approximate r e l a t i v e motions of the plates have been established for some time (e.g. Atwater, 1970). However, c e r t a i n important and c r i t i c a l d e t a i l s have been worked out only rec e n t l y . These are discussed i n an extensive review by Keen and Hyndman (1979). Figure 1.1 outlines the plate configuration i n the region of the continental margin of western Canada. In the northwest, the Queen Charlotte transform f a u l t zone, which separates the P a c i f i c and America plates, has r i g h t - l a t e r a l motion of about 5.5 cm/yr. The small Explorer and Juan de Fuca plates are separated from the P a c i f i c plate by an en echelon series of spreading centres and transform f a u l t zones, from the Tuzo Wilson seamounts i n the north to the Juan de Fuca ridge i n the south. The pattern i s probably formed by a complex process of ridge jumping, r e o r i e n t a t i o n of spreading axes and i n i t i a t i o n of new spreading centres (Keen and Hyndman, 1979; Hyndman et a l . , 1979). F u l l spreading rates vary from 4 to 6 cm/yr. The convergent boundaries between the two small plates and the America plate are approximately coincident with the 2 Figure 1.1. Major t e c t o n i c p l a t e boundaries o f f Canada's west coast. Arrows i n d i c a t e the d i r e c t i o n and magnitude of the r e l a t i v e p l a t e motions with respect to the America p l a t e . The magnitude s c a l e i s shown i n the lower l e f t corner. Area enclosed by dashed l i n e s d e f i n e s the coverage of the l o c a t i o n map i n f i g u r e 2.1. 3 base of the continental slope. Subduction rates for these young plates are low, l y i n g between 1 and 3 cm/yr (Riddihough, 1977). The d i f f e r e n t d i r e c t i o n s and rates of motion of the Explorer and Juan de Fuca plates require d i f f e r e n t i a l motion across some boundary between them. Hyndman et a l . (1979) established the existence of t h i s boundary, naming i t the Nootka f a u l t zone. Motion across i t i s l e f t l a t e r a l at a rate of about 2 cm/yr. The Nootka f a u l t zone forms the core of the area of i n t e r e s t i n the seismic r e f r a c t i o n project discussed i n t h i s t h e s i s . 1.2 Outline of Study The objective of t h i s study i s to obtain d e t a i l e d seismic information on the c r u s t a l structure of the Juan de Fuca and Explorer plates i n the region of the Nootka f a u l t zone, and to r e l a t e the seismic re s u l t s to the petrology of the oceanic crust and upper mantle. Insight concerning various tectonic processes may be gained from the knowledge drawn from t h i s study. The experimental set up and preliminary data processing relevant to a l l data i n t h i s study are described i n Chapter Two. The main body of the study i s divided into three major sections. In the f i r s t section, Chapter Three, the various aspects of i n t e r p r e t a t i o n of P wave p r o f i l e s are discussed. Current techniques of r e f r a c t i o n seismology used include the l i n e a r i z e d tau-p inversion of tra v e l times, ray tracing, and synthetic seismogram modelling of the 4 amplitude data. To r e l a t e the seismic r e s u l t s to the l i t h o l o g y of the oceanic crust and upper mantle, a detai l e d comparison of the r e s u l t s with the observed properties of o p h i o l i t e complexes i s given. For the purpose of reducing the ambiguities i n the p e t r o l o g i c a l i n t e r p r e t a t i o n of seismic r e s u l t s , shear wave information i s extremely important; therefore, considerable e f f o r t has been made i n t h i s study to obtain high q u a l i t y shear wave data. The second section, Chapter Four, deals with the problems of shear wave observation and the i n t e r p r e t a t i o n of two shear wave p r o f i l e s . The t h i r d section i s concerned with v e l o c i t y anisotropy, which i s shown by the re s u l t s of the p r o f i l e i n t e r p r e t a t i o n s to be very prominent i n t h i s study s i t e . The configuration of shots and receivers i n t h i s experiment provided some a r e a l l y d i s t r i b u t e d t r a v e l times which allowed study of azimuthally varying v e l o c i t i e s i n the upper mantle. The i n t e r p r e t a t i o n of t h i s data set and the implications of anisotropy f o r tect o n i c processes are discussed i n Chapter F i v e . F i n a l l y , Chapter Six gives a summary of the conclusions drawn from the r e s u l t s of t h i s study. 5 2. ACQUISITION AND PROCESSING OF DATA 2.1 The Experiment In 1977, four ocean bottom seismometers were deployed near, the Nootka f a u l t zone f o r a study of seismic r e f r a c t i o n and s e i s m i c i t y . The s e i s m i c i t y was analysed by Hyndman et a l . (1979). A n a l y s i s of the r e f r a c t i o n experiment c o n s t i t u t e s the present study. Three l i n e s were shot using a t o t a l of 150 e x p l o s i v e charges ranging i n s i z e from 5 to 270 kg. The l o c a t i o n and o r i e n t a t i o n of the shot l i n e s and the p o s i t i o n s of the OBS's are shown i n f i g u r e 2.1 : EX1 crosses the Nootka f a u l t zone i n an east-west d i r e c t i o n ; EX2 crosses the zone i n a south-north d i r e c t i o n ; and EX3 runs p a r a l l e l to the general trend of the zone. Reversals were achieved by deploying OBS's on both ends of the p r o f i l e s ; however, data from the OBS that was l o c a t e d at the western end of EX1 were l o s t due to instrument problems, so EX1 was unreversed. Maximum s h o t - r e c e i v e r range was approximately 80 km. The c o n f i g u r a t i o n of the OBS array was a compromise between e f f e c t i v e earthquake l o c a t i o n and convenient layout of seismic r e f r a c t i o n l i n e s . In a d d i t i o n to the exp l o s i v e shots, a 16 l i t r e airgun provided the energy source f o r a number of short range r e f r a c t i o n l i n e s . Shot spacing of 300 m f o r the airgun p r o f i l e s provided d e t a i l e d i n f o r m a t i o n f o r the upper c r u s t at each of the r e c e i v e r s i t e s . F i g u r e 2.1. L o c a t i o n map f o r the three ocean bottom seismometers ( s o l i d c i r c l e s ) and the r e f r a c t i o n shot l i n e s (heavy s o l i d l i n e s ) . Arrows i n d i c a t e the forward d i r e c t i o n of the p r o f i l e s . Bathymetric contours are i n meters (from T i f f i n and Seeman, 1975). The approximate extent of the Nootka f a u l t zone, i n d i c a t e d by the s t i p p l e d area, i s from the r e s u l t s of Hyndman et a l . (1979). 7 2.2 Instrumentation and Data D i g i t i z a t i o n The ocean bottom seismometers are of the free f a l l pop-up type s i m i l a r to those described by L i s t e r and Lewis (1976) and Johnson, L i s t e r and Lewis (1977). The same instruments were used i n an e a r l i e r r e f r a c t i o n experiment c a r r i e d out near the Expl o r e r r i d g e r e gion (Cheung and Clowes, 1981). The OBS's record i n continuous d i r e c t AM mode using a 4 channel tape head at a tape speed of 1 mm/s. The o v e r a l l frequency response i s bandlimited between 2 and 100 Hz with a dynamic range of 80 db. The rather wide dynamic range i s achieved by a b i p o l a r square-rooting s i g n a l compression scheme described by L i s t e r and Lewis (1976). However, such s i g n a l compression scheme does in t r o d u c e some d i s t o r t i o n s i n the waveforms recorded ( f o r example, see f i g u r e s 4.3 and 4.4). A 20 Hz time code from an i n t e r n a l c l o c k i s recorded on the f i r s t channel while the outputs from one v e r t i c a l and two h o r i z o n t a l 4.5 Hz geophones are recorded on the other three channels. In converting the o r i g i n a l analog data i n t o d i g i t a l form, a two stage process i s necessary due to the l i m i t e d c a p a c i t y of the a v a i l a b l e instrumentation f o r playback and d i g i t i z a t i o n . The procedures i n v o l v e d have been described i n d e t a i l by Cheung (1978) and are o u t l i n e d below f o r completeness. F i r s t , the data i s t r a n s c r i b e d from AM to FM at a playback speed of 15/16 i n c h per second ( i p s ) and a recording speed o f 15 i p s . The FM tape i s then played back at a speed of 15/16 i p s (approximately 1.5 times 8 r e a l time) f o r d i g i t i z a t i o n at a sampling r a t e of 312.5 samples per second ( s p s ) . The o v e r a l l e f f e c t of t h i s process i s to give a r e a l - t i m e data sampling rate i n the neighbourhood of 200 sps. The a c t u a l sampling rate v a r i e s w i t h time and d i f f e r s between the OBS's due to minor v a r i a t i o n s i n tape d r i v e motor speed and a small degree of tape s t r e t c h i n g . However, accurate estimates of the data sampling ra t e are obtained by decoding the d i g i t i z e d time code on the f i r s t channel. I t i s found that f o r a given record s e c t i o n , the sampling rate can be considered as constant; but between d i f f e r e n t record s e c t i o n s , v a r i a t i o n s i n the order of 5 d i g i t a l samples per second are found. 2.3 Range and T r a v e l Time C o r r e c t i o n s A l l shot l o c a t i o n s and r e c e i v e r p o s i t i o n s are determined from Loran C n a v i g a t i o n supplemented by s a t e l l i t e checks. R e l a t i v e p o s i t i o n accuracy was +200 m with greater u n c e r t a i n t i e s i n absolute p o s i t i o n (Hyndman et a l . 1979). S h o t - r e c e i v e r distances determined by t h i s method have accuracies s i m i l a r to or b e t t e r than those determined from d i r e c t water wave a r r i v a l s . To b r i n g the source and r e c e i v e r to the same depth l e v e l , a l l shots are corrected to an equivalent source depth of 2.5 km by ray t r a c i n g through the water column. The sound v e l o c i t y p r o f i l e of the water column ( f i g u r e 2.2) i s based upon the compilation of the N a t i o n a l Oceanographic Data Center i n Washington, D. C. f o r the study area. The c o r r e c t i o n s are phase v e l o c i t y dependent. For a phase v e l o c i t y of 6.0 km/s, the o f f -set i n the shot range due to the t r a v e l path through water i s 0.64 km 9 Velocity ( X10 m / s ) 148 149 150 151 J L F i g u r e 2.2. Sound v e l o c i t y p r o f i l e o f the water column based upon the the c o m p i l a t i o n o f the N a t i o n a l Oceanographic Data Center i n Washington D. C. f o r the study a r e a . 10 w h i l e the t r a v e l time c o r r e c t i o n i s 1.74 s. The r a p i d sedimentation ra t e i n t h i s region has r e s u l t e d i n a very subdued s e a f l o o r topography which r e q u i r e s only minor c o r r e c t i o n s ; however, s u b s t a n t i a l c o r r e c t i o n s are re q u i r e d f o r the v a r i a t i o n s i n basement topography. Since no r e f l e c t i o n p r o f i l e s along the shot l i n e s are a v a i l a b l e , the basement topography over most of the study area i s i n t e r p r e t e d from the r e f l e c t i o n p r o f i l e s of Davis and L i s t e r (1977). Their r e s u l t s give d e t a i l e d i n f o r m a t i o n on basement depths over a g r i d with spacing of approximately 10 km i n the region of Juan de Fuca r i d g e and v i c i n i t y . A d d i t i o n a l i n f o r m a t i o n on basement topography i s provided by Hyndman et a l . (1979) i n the area not covered by Davis and L i s t e r (1977). The s e a f l o o r and basement topography f o r the three shot l i n e s are p l o t t e d i n f i g u r e s 3-2c, 3 -3c , and 3.5c. Major g e o l o g i c a l features can be deli n e a t e d from the basement topography. At distances between 15 and 30 km along p r o f i l e EX1 ( f i g u r e 3-2c) where the Nootka f a u l t zone i s crossed, the basement shows a subdued expression. The prominent basement r e l i e f between 40 and 50 km i s Middle Ridge of the Juan de Fuca r i d g e system. A basement rid g e a s s o c i a t e d with the Sovanco f r a c t u r e zone emerges beyond 70 km. Along p r o f i l e EX3 ( f i g u r e 3 . 3 c ) , the major features of the basement are East Ridge at 40 km and Middle Ridge beyond 50 km. For p r o f i l e EX2 ( f i g u r e 3-5c), the basement v a r i a t i o n i s l e s s abrupt: East Ridge i s loc a t e d between 5 and 30 km and the Nootka f a u l t zone extends beyond t h i s to 40 km. To c a l c u l a t e the topographical c o r r e c t i o n s f o r a given i n t e r f a c e , a 11 datum l e v e l i s f i r s t chosen. C o r r e c t i o n s f o r the v a r i a t i o n i n the depth of t h i s i n t e r f a c e are obtained simply by r e p l a c i n g m a t e r i a l above or below t h i s reference l e v e l with the appropriate m a t e r i a l . This procedure i s c a r r i e d out f o r the water-sediment i n t e r f a c e and f o r the sediment-basement i n t e r f a c e . Figure 2.3 i l l u s t r a t e s the s i t u a t i o n f o r an exaggerated topographical bump on the sediment-basement boundary. At, the c o r r e c t i o n to be added to the t r a v e l time, i s given by At = Ah • q(p) (2.1) where q(p) = v S E D - 1 { i _ ( p v s E D )2j-1/2 " V B A S E - 1 { J " ( P V B A S E > 2 } " 1 / 2 (2.2) The ray parameter p i s determined from the slope of the uncorrected t r a v e l time curve while y^ED a n c* VBASE a r e ^he s e i s m i c v e l o c i t y of the sediment and the basement r e s p e c t i v e l y . This approach assumes that a l l s t r u c t u r e s at lower l e v e l s are h o r i z o n t a l (Whitmarsh, 1975). While the magnitude of t h i s c o r r e c t i o n i s i n s e n s i t i v e to the value of p, i t i s very s e n s i t i v e to the values chosen f o r the v e l o c i t i e s ( D e t r i c k and Purdy, 1980). Another approach f o r c o r r e c t i n g topographical v a r i a t i o n s i s to remove the t r a v e l time from the shot to the entry p o i n t at the i n t e r f a c e concerned (A to B i n f i g u r e 2.3). In t h i s approach, i t i s assumed that a l l s t r u c t u r e s at lower l e v e l s are p a r a l l e l to the topography of t h i s boundary. This c a l c u l a t i o n does not i n v o l v e the v e l o c i t y immediately below the i n t e r f a c e , but i t i s very s e n s i t i v e to the value of p. 12 Figure 2.3. Schematic diagram illustrating the various parameters used in calculating topographical corrections to the travel time data. The' depths are in units of km and the velocities are in units of km/s. 13 There are two main reasons why the f i r s t approach i s favoured i n the present study.. F i r s t , due to l a r g e s c a t t e r s i n the t r a v e l time data points (see f o r example f i g u r e 2.4), the estimate of the ray parameter i s not very r e l i a b l e , rendering the second method unstable. Second, the v e l o c i t i e s of the sediment and the basement are known s u f f i c i e n t l y w e l l i n the study area f o r the f i r s t method to give reasonable c o r r e c t i o n s . Although no r e f r a c t e d a r r i v a l s are observed from w i t h i n the sediment, i n f o r m a t i o n f o r the sediment v e l o c i t y i s taken from Davis et a l . (1976) who obtained a number of v e l o c i t i e s f o r the sediment at the northern end of the Juan de Fuca r i d g e by using bottom sources and ocean bottom seismometers. The basement P v e l o c i t y of 4.5 km/s i s obtained from the average of the apparent v e l o c i t i e s of the f i r s t branch of the r e f r a c t e d a r r i v a l s on the airgun p r o f i l e s . S v e l o c i t y f o r the basement i s obtained from the P v e l o c i t y by assuming a Poisson's r a t i o of 0.34 (Spudich and Orcutt, 1980a). Values of the various parameters used i n making the topographical c o r r e c t i o n s are shown i n f i g u r e 2.3. In making c o r r e c t i o n s f o r the d i f f e r e n t r e c e i v e r depths, a v e r t i c a l t r a v e l path i s assumed. This i s acceptable because both the water and basement depths at the OBS s i t e s d i f f e r e d only s l i g h t l y from the reference l e v e l s and the a c t u a l angles of emergence are estimated to be near v e r t i c a l . A comparison of the raw t r a v e l time data and the correc t e d data i s shown i n f i g u r e 2.4 f o r p r o f i l e EX3. The s e a f l o o r and basement topography are a l s o p l o t t e d . The obvious t r a v e l time advance due to the basement ridge at 40 km i s much reduced i n the correc t e d t r a v e l time 14 C3i * C O Q I H I E X 3 m + + + + + + + + + + + °D £ D + + + * + + +++ + +++ + + D u n c o r t i m e + c o r t i m e + + 0 10 20 30 ~~40 50 60 Dist (km) Sea Floor Basement Figure 2.4. An example showing the travel time data before and after applying topographical corrections for profile EX3. The sea floor and basement reliefs are plotted at 6 times vertical exaggeration. 15 curve; however, some degree of scatter s t i l l e x ists i n the data. Maximum deviation of the basement depth from the datum l e v e l i s approximately 1.0 km which corresponds to a correction of 0.31 s for P a r r i v a l s and 0.10 s f o r S a r r i v a l s . For va r i a t i o n s i n the water-sediment i n t e r f a c e , the t y p i c a l t r a v e l time correction i s much smaller i n magnitude, i n the order of a few milliseconds. As with other marine r e f r a c t i o n studies i n areas of considerable topography (Spudich and Orcutt, 1980a; Detrick and Purdy, 1980), the topographical correction i s the largest uncertainty i n the t r a v e l time data due to possible errors i n the v e l o c i t y estimates of the upper layers and uncertainties about the ray entry points. Errors asssociated with the topographical corrections are estimated to be i n the order of +0.03s. Further uncertainties i n timing a r i s e from factors such as shot o r i g i n time determination, shot depth estimate, and OBS clock d r i f t s . The contribution to the timing error from these factors are estimated to be +0.03 s, giving a maximum t o t a l error of +0.06 s. 16 3. P WAVE PROFILE INTERPRETATION 3.1 Record Sections V e r t i c a l component record sections, corrected for topography and the t r a v e l path through water, are shown i n figures 3-2a to 3.6a f o r the explosion p r o f i l e s and figures 3.7a to 3.9a for the airgun p r o f i l e s . Trace amplitudes on a l l sections are scaled with an r-2 spreading factor to enhance the weak a r r i v a l s at greater distances. To compensate f o r the d i f f e r e n t charge s i z e s i n the explosion p r o f i l e s , an amplitude normalization factor of weights/3 j [ s applied. This normalization i s based on empirical r e s u l t s r e l a t i n g the si z e of the charge with the amplitude of seismic energy i t generates (O'Brien, 1960; Kanestrtfm and 0 v r e b 0 , 1978). Figure 3-1 shows the Fourier power spectra for a refracted P a r r i v a l from an explosion and a segment of the noise i n the data p r i o r to any seismic a r r i v a l s . Both spectra are calculated from a sign a l duration of one second. It i s apparent that much of the noise i s concentrated i n the frequency range of 23 to 27 Hz and at 2 Hz while the P a r r i v a l has a well defined peak at 6.5 Hz and a small secondary peak at 19 Hz. This i l l u s t r a t e s that the s i g n a l and noise are quite d i s t i n c t i n frequency content and improvement i n data q u a l i t y may be gained by d i g i t a l f i l t e r i n g . The seismic traces i n the record sections have a l l been bandpassed with a 5 to 20 Hz eight pole zero phase Butterworth f i l t e r . F i l t e r i n g of the data does improve the appearance of the seismic traces thus Relative Power Relative Power 0.0 0.2 0.4 0.6 0.8 1.0 0 0.2 0.4 0.6 0.8 1.0 0 4r 1 1 1 1 1 L • I 1 1 1 1 F i g u r e 3.1. F o u r i e r power s p e c t r a , normalized to the maximum power peak, f o r a segment of n o i s e i n the data p r i o r to any seism i c a r r i v a l s and f o r a r e f r a c t e d P phase (see tr a c e (a) of f i g u r e 4.3). The s i g n a l d u r a t i o n i s one second f o r both s p e c t r a . Note the low power of the noise peaks r e l a t i v e to s i g n a l i n the P wave spectrum. 18 co o 0 0 o i n O L U rx K— co a o m o cm-e 2 i o t- e 2 i o i- z c * . (33S) 9/1SI0 - i • (33S) 9/iSIO - 1 WW Hld3Q Figure 3.2. (a) Record section along profile EX1 corrected for topographical variations of the seafloor and the basement. Amplitudes of the traces have been corrected for charge size and normalized by a spherical spreading factor of r 2 . Travel times and ranges have been corrected for the one-way travel path through water. First arrival picks are noted by horizontal tick marks. The travel time curves are transferred from the synthetic seismogram section below, (b) Best f i t synthetic seismogram section for profile EX1 calculated by the WKBJ approach of Chapman (1978) and using the velocity-depth model of figure 3.11a. An r 2 spreading factor is included, (c) Topography of the seafloor and the basement along profile E X 1 . Vertical exaggeration is 3X. The profile runs approximately westward from the origin. 19 Figure 3.3. (a) Record section, (b) synthetic seismogram section and (c) topography along profile EX3. Description is the same as that in figure 3.2. Synthetic seismograms are calculated using the velocity-depth model of figure 3.11c. Profile direction is from NE to SW. 20 r— 1 1 1 h o i 1 1 i f e z t o i - e z i o i -(339) 9/1SIQ - 1 (03S) 9/1910 - 1 F i g u r e 3.4. (a) Record s e c t i o n and (b) s y n t h e t i c seismogram s e c t i o n f o r p r o f i l e EX3R. S y n t h e t i c seismograms are c a l c u l a t e d u s i n g the v e l o c i t y -depth model of f i g u r e 3.11c. O r i g i n i s l o c a t e d at the range of 45 km on p r o f i l e EX3- P r o f i l e d i r e c t i o n i s from SW to NE. 21 C\J X U J LU O CL . CD GO .o co .0 ID -CDLU •z. r r t— e n Q - CD e n - CD CM 1 1 1 1 T O e z i o i-(339) 9/iSIQ - 1 I 0 I-9/1SIQ - 1 Z £ P (WM) Hld3Q Figure 3.5. (a) Record section, (b) synthetic seismogram section and (c) topography along profile EX2. Description is the same as that in fiqure 3.2. Synthetic seismograms are calculated using the velocity-depth model of figure 3.11b. Profile direction is from south to north. Figure 3.6. (a) Record section and (b) synthetic seismogram section for profile EX2R. Synthetic seismograms are calculated using the velocity-depth model of figure 3.11b. Origin is located at the range of 53 km on profile EX2. Profile direction is from north to south. 23 (33S) 9/1SIQ - 1 £ 2 I (33S) 9/1SIQ 0 i Figure 3.7. (a) Record section and (b) synthetic seismogram section of the airgun profile AG2N shot along the same direction as EX2. The synthetic seismograms are calculated to a distance of 18 km using the same velocity-depth model as EX2 but with a denser trace spacing. Some subcritical reflection travel time branches are included, though the amplitudes of these arrivals are small. 24 0 J U W ^ w ^ A j ^ w LCNJ r 21 i l l O U J CO hoc hCD (339) 9/1SIQ - 1 — ^ w v V v V -—NA^VWV^ • — w v s A / * -W-VvV*" >A^VV\/\-^A/\/y\/^ -—WVV • — V V — % —# —w —w vV ——w O b J — C J C E I— C O I—I Q hco hCD \ (339) 9/1SIQ - 1 Figure 3.8. (a) Record section and (b) synthetic seismogram section for airgun profile AG3S shot along the same direction as profile EX2R. The extended coda on traces between 12 and 14 km is well modelled by the synthetic seismograms. The large amplitude secondary arrivals between 7 and 8 km, probably due to multiply reflected phases, are not modelled. The velocity-depth model used in the calculation of the synthetic seismograms is the same as for EX2R. 25 LL_ O Q _ ^ A A / ^ / L-rsj O b J — (_) ~ZL CT I— CD •—i C D hco hcD •£ 2 , 1 0 (339) 9/1SIQ - 1 — # • — v \ A -—vv^-— i j t y -»M -'WVN - A N LCNJ O L U — C_) rx i— C O «—« Q hco £ 2 1 0 (339) 9/1SIQ - 1 Figure 3.9. (a) Record s e c t i o n and (b) s y n t h e t i c seismogram s e c t i o n f o r a i r g u n p r o f i l e AG1S, shot south of the OBS t h a t recorded p r o f i l e EX1. The v e l o c i t y - d e p t h model used i n the c a l c u l a t i o n of the s y n t h e t i c seismograms i s the same as f o r EX1. The secondary a r r i v a l s between 8 and 10 km are not modelled. l 26 f a c i l i t a t i n g comparison with s y n t h e t i c seismograms f o r the purpose of amplitude i n t e r p r e t a t i o n ; however, i t a l s o smooths out the f i r s t breaks of the P a r r i v a l s which i s u n d e s i r a b l e . U s u a l l y the impulsive f i r s t breaks of c r u s t a l a r r i v a l s can be1 picked more a c c u r a t e l y from the u n f i l t e r e d data w h i l e weak Pn a r r i v a l s contaminated by severe noises would b e n e f i t from f i l t e r i n g . The a r r i v a l p i c k s are i n d i c a t e d by h o r i z o n t a l t i c k marks on the data t r a c e s i n the explosion record s e c t i o n s . F i r s t a r r i v a l breaks are not marked i n d i v i d u a l l y on the airg u n s e c t i o n s since the breaks are obvious i n most cases. Superimposed on the data are t h e o r e t i c a l t r a v e l time curves t r a n s f e r r e d from the best f i t s y n t h e t i c record s e c t i o n s which are p l o t t e d immediately below the data (see f i g u r e s 3.2b through 3 . 9 b ) . There are a number of common features i n the explosion record s e c t i o n s ( f i g u r e s 3.2a to 3.6a): (1) l a r g e amplitude a r r i v a l s over a short distance i n t e r v a l associated with the t r i p l i c a t i o n caused by a v e l o c i t y g radient immediately above the crust-mantle boundary; (2) few prominent secondary a r r i v a l s ; and (3) small amplitude Pn phases. There are al s o a number of important d i f f e r e n c e s among the record s e c t i o n s . The di s t a n c e range over which the l a r g e amplitude t r i p l i c a t i o n occurs d i f f e r s from p r o f i l e to p r o f i l e and even between forward and reverse p r o f i l e s . For example, the t r i p l i c a t i o n i s near 25 km i n EX2 but i s near 35 km i n EX2R ( f i g u r e s 6a and 7a). This i n d i c a t e s that there are some l a t e r a l v a r i a t i o n s i n the s t r u c t u r e s at depth and that the c r u s t a l t hickness v a r i e s between the d i f f e r e n t l o c a t i o n s . The t r a v e l time curves of EX2 and EX2R a l s o show no t i c e a b l e d i f f e r e n c e s from those of the other 27 p r o f i l e s . The slopes of the T-X curves of EX2 and EX2R reach an apparent v e l o c i t y of 7.5 km/s at a much c l o s e r distance than the other p r o f i l e s and t h i s slope i s maintained without change to beyond 75 km. Als o note the sudden decay of amplitude along EX2 and EX2R at distances g r e a t e r than 45 km. The observed amplitudes are approximately four times smaller than the headwave amplitudes p r e d i c t e d by the s y n t h e t i c seismograms. Some of the f i r s t a r r i v a l s on EX2R are too poor to be picked with any confidence even though the shot s i z e s are as l a r g e as 270 kg. The moderately strong secondary a r r i v a l s w i t h apparent v e l o c i t y of 6.5 km/s on p r o f i l e EX2 are a l s o attenuated beyond 45 km. I n t e r p r e t a t i o n s of these anomalous features of EX2 and EX2R w i l l be discussed l a t e r . Airgun p r o f i l e AG2N ( f i g u r e 8a), shot along the same d i r e c t i o n as EX2, i s a good q u a l i t y one; c l e a r f i r s t a r r i v a l s can be seen to distan c e s of 18 km. P r o f i l e s AG3S and AG1S ( f i g u r e s 3-8a and 3-9a), w i t h a maximum range of approximately 14 km, are more t y p i c a l of the other airgun p r o f i l e s recorded. The constant source signature of the airgun and the c l o s e shot spacing a l l o w the s u b t l e v a r i a t i o n s i n amplitude w i t h distance to be observed. By modelling the airgun amplitude and t r a v e l time data, much b e t t e r c o n s t r a i n t s on the upper c r u s t a l s t r u c t u r e s to depths of 2 km subsediment at each r e c e i v e r s i t e can be obtained than would be p o s s i b l e w i t h the e x p l o s i v e sources alone. 3.2 I n v e r s i o n of T r a v e l Time Data In the f i e l d of s e i s i m i c t r a v e l time i n v e r s i o n , a number of s i g n i f i c a n t developments have taken place i n the l a s t few years. Most of 28 the new advances are based on the reparameterization of t r a v e l time data, T(X), to the delay time function t(p)=T(p)-pX(p), where p i s the ray parameter. There are several advantages i n using the function x. For example, since r i s a single-valued, monotonically decreasing function of p, d i f f i c u l t i e s associated with t r i p l i c a t i o n s i n the t r a v e l time versus distance function are avoided. Also f i r s t order errors i n p r e s u l t i n only second order errors i n tau which are of the same order as errors i n T(X); therefore, accurate estimate of r(p) can be obtained from the data T(X). Perhaps the most important aspect of the tau function i s that the inverse problem can be formulated as a l i n e a r problem with a simple change of variable (Garmany, 1979). In the present study, the i n t e r p r e t a t i o n of t r a v e l time data u t i l i z i n g the tau-p function proceeds i n two ways. F i r s t , extremal bounds on the velocity-depth function are obtained from the bounds of x(p) using the methods of Bessonova et a l . (1974). Second, a p a r t i c u l a r v-z model which s a t i s f i e s the tau-p data to within the errors i n the observations i s found by l i n e a r programming techniques (Garmany et a l . , 1979). Before carrying out the tau inversion procedures, the tau-p data must f i r s t be calculated from the T-X t r a v e l time points. There are a number of methods for doing t h i s reparameterization (eg. Kennett, 1976; Bates and Kanasewich, 1976) but a f t e r experimentation with the e x i s t i n g data, the technique described by Kennett and Orcutt (1976) was chosen for i t s superior numerical s t a b i l i t y and ease of a p p l i c a t i o n . Figure 3.10 i l l u s t r a t e s t h i s method schematically. F i r s t , a smooth cubic-spline X I : S c h e f f l 3 t i c diagram i l l u s t r a t i n g the c a l c u l a t i o n of tau(p) t^TJ n I m a x i m u i n o f t n e reduced t r a v e l time curve i s the value of , n , f P i ; The procedure i s repeated f o r the range of values of P l c o n s i s t e n t w i t h the data. 1 30 i s f i t t e d to the reduced t r a v e l time data. Then the tau value f o r a given value of p, say p i ? i s obtained by f i n d i n g the maximum of the reduced t r a v e l time curve T-p iX. the e r r o r a s s o c i a t e d w i t h t h i s estimate of tau i s the same as the u n c e r t a i n t y i n T(X) and i s set at +0.06s. The upper and lower bounds of *c(p) from the combined data set of airgu n and e x p l o s i v e l i n e s are then mapped i n t o bounds of the v e l o c i t y -depth f u n c t i o n by the methods of Bessonova et a l . (1974). An example of the extremal bounds s o l u t i o n i s shown i n f i g u r e 3.11d. The bounds de l i n e a t e a change i n the v e l o c i t y gradient at approximately 3.5 km depth; however, no d e t a i l s of the v e l o c i t y - d e p t h s t r u c t u r e are d e f i n e d . Extremal bounds f o r the other p r o f i l e s are s i m i l a r l y broad and bracket the f i n a l v e l o c i t y - d e p t h models i n much the same manner as f o r EX2. This l a c k of r e s o l v i n g power i m p l i e s that very few r e s u l t s of geophysical i n t e r e s t can be obtained from the extremal bounds. S i m i l a r views have been expressed by others i n the course of i n t e r p r e t i n g a c t u a l data by the tau method (eg. Spudich and Orcutt, 1980a). Recently a number of authors have shown that greater c o n s t r a i n t s on the extremal bounds can be achieved w i t h the a d d i t i o n of X(p) data (Orcutt, 1980 and J u r k e v i c s et a l . 1980). However, the e x t r a c t i o n of X(p) data from the t r a v e l time data T(X) i s a much more unstable process than the es t i m a t i o n of r ( p ) . Since the aim of the present study i s to o b t a i n d e t a i l e d c h a r a c t e r i s t i c s of the v e l o c i t y - d e p t h s t r u c t u r e by modelling the amplitude data, a more d e s i r a b l e approach than o b t a i n i n g refinements to the extremal bounds would be the c a l c u l a t i o n of a s u i t a b l e s t a r t i n g model f o r s y n t h e t i c seismogram c a l c u l a t i o n s . The l i n e a r i z e d i n v e r s i o n 31 formulation of tau-p data, described by Garmany et a l . (1979), i s i d e a l f o r t h i s purpose. By solving a set of l i n e a r tau-p equations using l i n e a r programming techniques, a p a r t i c u l a r velocity-depth model which s a t i s f i e s the tau-p data to within the errors i n tau can be e a s i l y derived (see Appendix 1 ) . The l i n e a r programming sol u t i o n f o r EX2 (LINP) i s shown i n f i g u r e 3-11d along with the f i n a l model (EX2) for comparison. As i n the case of the extremal bounds c a l c u l a t i o n , an error of +0.06 s i n tau i s assumed. The major advantage of the l i n e a r programming approach i s that the computation i s extremely f a s t and economical and a useful s t a r t i n g model fo r amplitude studies i s r e a d i l y obtained. Only minor changes to LINP are needed to a r r i v e at the f i n a l model (figure 3-11d) which i s obtained by amplitude comparisons using synthetic seismogram sections. 3.3 Synthetic Seismograms and Amplitude Modelling In the l a s t section, i t has been shown that by using the f i r s t a r r i v a l t r a v e l time data alone, the extremal bounds thus obtained provide us with very l i t t l e d e t a i l e d information concerning the v e l o c i t y structure of the c r u s t . The l i n e a r programming technique does give models with considerably more structure, but the inherent uncertainty i n trav e l time inversions, expressed c l e a r l y by the poor r e s o l u t i o n i n the extremal bounds, discourages us from putting too much f a i t h i n the LINP solutions. Considerably more information about the c r u s t a l v e l o c i t y structure can be extracted from seismic r e f r a c t i o n data i f the observed amplitudes 32 as well as t r a v e l times of the various seismic phases are considered. Recent advances i n t h e o r e t i c a l and computational techniques have allowed the modelling of amplitude data by synthetic seismograms to become a routine procedure i n the i n t e r p r e t a t i o n of marine r e f r a c t i o n data (eg. Malecek and Clowes, 1978; Spudich and Orcutt, 1980a). In the present study, synthetic seismogram modelling of the data i s one of the most important steps i n obtaining det a i l e d velocity-depth information for the oceanic crust and upper mantle i n the study s i t e . There are a number of methods for computing synthetic seismograms given the v e l o c i t y and density structure of an earth model. Extensive reviews of the various methods can be found i n Chapman (1978) and Richards (1979) while a comparison of the p r a c t i c a l a p p l i c a t i o n of some of the methods i s given by Spudich and Orcutt (1980b). Almost a l l of the techniques require l a t e r a l homogeneity i n the v e l o c i t y structure and they d i f f e r mainly i n t h e i r degree of accuracy and the expense of computation (Aki and Richards, 1980). For the present study, the WKBJ synthetic seismogram algorithm of Chapman (1978) i s used i n the modelling of the amplitude data. Chapman has shown that the WKBJ seismogram c o r r e c t l y predicts a r r i v a l s f o r turning rays, p a r t i a l and t o t a l r e f l e c t i o n s and head waves of r e f l e c t i o n s . There are, however, a number of l i m i t a t i o n s i n using the WKBJ algorithm. For example, WKBJ seismograms are not s t r i c t l y correct for waves turning i n the v i c i n i t y of a low v e l o c i t y zone or a v e l o c i t y d i s c o n t i n u i t y . D i f f r a c t i o n beyond t r i p l i c a t i o n or p r e c r i t i c a l r e f l e c t i o n from v e l o c i t y gradient zones also are not c o r r e c t l y calculated (Spudich 33 and O r c u t t , 1980b). However, numerical experiments c a r r i e d out by the w r i t e r and others (Spudich and Orcutt, 1980b) have shown that WKBJ seismograms d i f f e r from the more accurate r e f l e c t i v i t y method (RM) s y n t h e t i c seismograms by n e g l i g i b l e amounts f o r v e l o c i t y - d e p t h models t y p i c a l of the oceanic c r u s t . S i g n i f i c a n t l y , the expense of computation f o r the r e f l e c t i v i t y method i s an order of magnitude higher than t h a t f o r the WKBJ al g o r i t h m due to the f a c t that v e l o c i t y gradients must be modelled as stacks of constant v e l o c i t y l a y e r s i n the RM approach whereas the WKBJ al g o r i t h m handles v e l o c i t y gradients e f f i c i e n t l y by a n a l y t i c approximations. The r e l a t i v e l y inexpensive WKBJ method allows a la r g e number of t r i a l models i n the process of f i t t i n g the amplitude data. Except f o r some e x p l a i n a b l e l a t e r a l v a r i a t i o n s i n s t r u c t u r e to be discussed i n d e t a i l i n the f o l l o w i n g s e c t i o n s , the WKBJ approximation appears to be v a l i d f o r the present data s e t . The best f i t s y n t h e t i c record s e c t i o n s are p l o t t e d as f i g u r e s 3b through 9b. In generating the s y n t h e t i c seismograms, only the primary ray paths are considered since t h i s approach g e n e r a l l y produces acceptable f i t s to the data. An exception occurred f o r p r o f i l e EX3R ( f i g u r e 3.4) where a m u l t i p l e r e f l e c t i o n from the underside of the sediment-basement i n t e r f a c e generated s i g n i f i c a n t amplitudes at distances from 15 to 20 km. This m u l t i p l e i s modelled i n the s y n t h e t i c s e c t i o n and i t i s found to have n e g l i g i b l e c o n t r i b u t i o n at grea t e r d i s t a n c e s . Coherent a r r i v a l s of t h i s phase are not observed on the other p r o f i l e s . Reverberation w i t h i n the sediment can a l s o cause l a t e r a r r i v a l s which may e x p l a i n the extended wave-train on the traces of p r o f i l e EX2R ( f i g u r e 3-6a). However, s y n t h e t i c seismograms i n c o r p o r a t i n g 34 the sediment layer multiples do not give s i g n i f i c a n t amplitude for the assumed sediment v e l o c i t y structure. An a d d i t i o n a l complication i s the strong e f f e c t s of varying sediment thickness on the a r r i v a l times of these phases with respect to the primary a r r i v a l s . For these reasons, only the primary ray paths are used. The modelling procedure consists of c a l c u l a t i n g a synthetic seismogram section from a s t a r t i n g velocity-depth model (LINP solution) and comparing i t with the data record s e c t i o n . Perturbations to the model are made by a t r i a l - a n d - e r r o r approach and new synthetic sections are calculated u n t i l an acceptable f i t to the amplitudes and t r a v e l times of the observed section i s found. No attempt i s made to match the observed waveforms exactly; the goodness of f i t i s determined s u b j e c t i v e l y by v i s u a l inspection of the r e l a t i v e amplitudes of the traces and using the t r a v e l time data as c o n s t r a i n t s . With the development of automated schemes for synthetic seismogram f i t t i n g , such as the method of Chapman and Orcutt (1980), objective i n t e r p r e t a t i o n of amplitude data w i l l have some advantages over the more tedious t r i a l - a n d - e r r o r approach. As pointed out by Spudich and Orcutt (1980a), the drawback i n the t r i a l - a n d - e r r o r approach i n amplitude modelling i s the i n a b i l i t y to define adequate bounds on the v e l o c i t y -depth function. Though i t i s impossible to sample a l l models within the model space, one may display a range of acceptable models to give an idea of the extent of the uncertainty i n the procedure, as has been done by Spudich and Orcutt. For the sake of s i m p l i c i t y , the writer has chosen to display only the f i n a l preferred models (figures 11a, b, and c) while 35 bearing i n mind that there w i l l be other acceptable models that can also s a t i s f y the data to within the errors of the observations. 3.4 Results With the a c q u i s i t i o n of more data and the improvement i n seismic i n t e r p r e t a t i o n techniques, the simple three layer model of the oceanic crust suggested by R a i t t (1963) has undergone a number of modifications. Fine structures such as sublayering and v e l o c i t y gradients within the layers are commonly found. In addition, there appears to be a systematic v a r i a t i o n of the c r u s t a l structure with age, l o c a t i o n , and tectonic regime of the s i t e . In t h i s section, the P wave velocity-depth models f or each of the three p r o f i l e s and t h e i r reversals, interpreted by modelling the data record sections with WKBJ synthetic sections, are discussed. I t w i l l be shown that the oceanic crust at the study s i t e as defined by the P wave data does contain some of the f i n e structures as well as l a t e r a l v ariations mentioned. Since the present r e s u l t s are derived by using techniques that require l a t e r a l l y homogeneous structure, the existence of l a t e r a l v a r i a t i o n s w i l l have to be explained c a r e f u l l y . The f i n a l velocity-depth models, shown i n figures 11a, b, and c, indica t e that the c r u s t a l structure i n the Nootka f a u l t zone region consists of zones of d i f f e r i n g v e l o c i t y gradients rather than d i s t i n c t i v e l a y e r s . S t i l l , the notion of a layered structure i s useful for the purpose of comparison of seismic r e s u l t s ; therefore, layer 36 VELOCITY (KM/S) 2 4 6 8 VELOCITY (KM/S) 2 4 6 8 10 VELOCITY (KM/S) 2 4 6 8 10 VELOCITY (KM/S) 2 4 6 8 Figure 3.11. (a), (b), and (c): Velocity-depth curves for a l l the profiles. Depths are measured from the seafloor. The sedimentary layer on a l l profiles is constrained to be 1 km thick and has a velocity of 1.8 km/s. The upper crustal models to a subsediment depth of 2 km are derived from the air gun data, (d) Extremal bounds for profile EX2 are denoted by squares. The linear programming solution LINP (dashed line) can be compared with the final model for EX2 (solid line). 37 d i v i s i o n s are assigned to the f i n a l velocity-depth p r o f i l e s i n accordance with the c h a r a c t e r i s t i c s of the v e l o c i t y gradient i n each depth range. Figure 3.12 i s such a schematic representation. Upper Crust A r e l a t i v e l y consistent picture of the upper crust i s shown by the velocity-depth p r o f i l e s i n f i g u r e 3-11- Immediately below the sediment lay e r , the v e l o c i t y ranges from 3-7 to 4.7 km/s and increases with depth at a moderate average v e l o c i t y gradient of 0.5 km/s/km (hereafter abbreviated to /s) to a depth of approximately 1.9 km. Then a zone of very high v e l o c i t y gradient i s found which marks the t r a n s i t i o n from layer 2A to layer 2B. The average gradient of 1.6/s i n t h i s zone i s the highest that has been encountered. Such a zone i s not found i n the velocity-depth p r o f i l e of EX2R; instead, a v e l o c i t y d i s c o n t i n u i t y i s found at a depth of 1.4 km. However, t h i s feature of EX2R i s not well constrained. Below the t r a n s i t i o n zone l i e s layer 2B where the v e l o c i t y gradient decreases to a lower value of 0.3/s except i n EX2 and EX2R, for which the v e l o c i t y gradient remains high at 1.0/s. At the base of layer 2B, v e l o c i t i e s are found to be i n the range of 6.0 to 6.4 km/s. The combined thickness of layer 2A and layer 2B ranges from 2.3 to 3.1 km but a lower value of 1.6 km i s found f o r p r o f i l e EX2R. 38 EX3 EX3R EX1 EX2R EX2 LAYER o -1.8 1.8 1.8 1.8 1.8 1 4.7 4.7 4.4 4.2 3.7 2A CM -5.0 5.0 4.8 4.9 4.4 2B 6.4 6.4 «» - 6.6 f ! 6.1 6.6 6.9 3A 3B sD -7.0 7.5 co— 7.1 7.8 7.5 MANTLE 7.7 o_ 8.0 8.3 8.3 CM Figure 3.12. Schematic representation of the velocity-depth p r o f i l e s of fi g u r e 3.11 i n terms of the conventional layered model of the oceanic c r u s t . The depth i s measured from the s e a f l o o r . The numbers are the v e l o c i t y at the top of each l a y e r i n km/s. The sediments, l a y e r 1, are constrained to have a thickness of 1.0 km and a v e l o c i t y of 1.8 km/s. Layer designations follow those of Christensen and Salisb u r y (1975). The question marks on the layer 2B-3A boundary of EX3R i n d i c a t e that no boundary i s evident i n the v e l o c i t y depth curve; the v e l o c i t y of 6.1 km/s i s that at the depth marked. 39 Lower Crust Beneath layer 2B, a v e l o c i t y d i s c o n t i n u i t y , varying i n i t s degree of sharpness, i s interpreted f o r a l l p r o f i l e s except f o r EX3R where no d i s c o n t i n u i t y i n either the v e l o c i t y or the gradient i s present. Where the d i s c o n t i n u i t y i s observed, i t i s interpreted to be the boundary between layer 2B and 3A. The structure of the lower crust i s s u f f i c i e n t l y d i f f e r e n t among the various p r o f i l e s to warrant an i n d i v i d u a l discussion of each. EX1: As shown i n figure 3.11a, layer 3A s t a r t s at a depth of 3.8 km and a v e l o c i t y of 6.6 km/s, increasing with a small v e l o c i t y gradient (0.1/s) to 7.6 km. At t h i s depth, a sharp increase i n the v e l o c i t y gradient i s found i n a zone approximately 1.8 km t h i c k . This zone marks the t r a n s i t i o n from layer 3A to the upper mantle and has been i d e n t i f i e d as layer 3B. Such a t r a n s i t i o n zone has also been found i n a number of r e f r a c t i o n studies of the oceanic crust (Malecek and Clowes, 1978; Spudich and Orcutt, 1980a). At the depth of 9.4 km, an upper mantle v e l o c i t y of 8.0 km/s i s found. EX3: The velocity-depth p r o f i l e of EX3 (figure 3.11c) indicates a more complex structure f o r la y e r 3 than that of EX1. At a depth of 6.0 km, a 0.6 km thick zone of high v e l o c i t y gradient (0.9/s) i s required to produce the large amplitude a r r i v a l s observed on the EX3 record section at the distance range of 30 km (figure 3-3a). The second occurrence of large amplitude a r r i v a l s at 45 km range on the record section i s caused by the t r i p l i c a t i o n from layer 3B. Layer 3B of p r o f i l e EX3 i s somewhat thicker than that of EX1 and the v e l o c i t y gradient found 40 w i t h i n i s not as high. The v e l o c i t y of 8.3 km/s derived f o r the depth of 11.2 km i s i n t e r p r e t e d as the upper mantle v e l o c i t y . EX3R: As mentioned e a r l i e r , the v e l o c i t y - d e p t h p r o f i l e of EX3R does not show any d i s c o n t i n u i t y between l a y e r 2B and l a y e r 3A. Instead, the v e l o c i t y remains n e a r l y constant at 6.0 km/s to the depth of approximately 6.0 km and between 6.0 and 8.6 km the v e l o c i t y g r a d i e n t begins to increase r a p i d l y ( f i g u r e 3.11c). This zone of high v e l o c i t y gradient gives r i s e to the l a r g e amplitude a r r i v a l s at the dista n c e s of 27 to 35 km on the record s e c t i o n of p r o f i l e EX3R ( f i g u r e 3.4a). These a r r i v a l s even dominate the t r i p l i c a t i o n caused by l a y e r 3B. This d i f f e r e n c e i n amplitude behaviour between EX3 and EX3R i s q u i t e apparent when comparing the two record s e c t i o n s . Such a d i f f e r e n c e can be explained by the f a c t that the two p r o f i l e s are not exact r e v e r s a l s ; that i s , seismic waves from shots at a range of 30 km from the r e s p e c t i v e r e c e i v e r s d i d not sample the same c r u s t a l m a t e r i a l . The common depth region f o r the two p r o f i l e s i s at upper mantle depths, where the same v e l o c i t y of 8.3 km/s i s found f o r the reversed p r o f i l e s EX3 and EX3R. EX2: U n l i k e the other p r o f i l e s , there i s no apparent v e l o c i t y higher than 7.5 km/s observed on p r o f i l e EX2 below a depth of 6.4 km ( f i g u r e 3.11b). A zone of high v e l o c i t y gradient s i m i l a r to l a y e r 3B of the other p r o f i l e s i s found at a much shallower depth of 5.0 km. The shallow depth of t h i s t r a n s i t i o n zone on EX2 r e f l e c t s the f a c t that the l a r g e amplitude t r i p l i c a t i o n occurs at the much c l o s e r distance of 25 km ( f i g u r e 3«5a) than that observed on the other p r o f i l e s . There are no 41 other l a r g e amplitude t r i p l i c a t i o n s observed to distances as f a r as 80 km. This i m p l i e s that the v e l o c i t y must remain constant below 6.4 km to some depth, a f a c t f u r t h e r confirmed by the extremal bounds of EX2 ( f i g u r e 3•11d) which show that the maximum v e l o c i t y below 6.4 km i s w e l l constrained to be l e s s than 7.8 km/s. The 7-5 km/s v e l o c i t y i s i n t e r p r e t e d to be that of the upper mantle along EX2 (compared w i t h 8.0 and 8.3 km/s f o r EX1 and EX3). The maximum v e l o c i t y observed on the reverse p r o f i l e EX2R i s a l s o 7.5 km/s so that a d i p p i n g s t r u c t u r e cannot be used to e x p l a i n the low upper mantle v e l o c i t y . A more l i k e l y e x p l a n a tion may be anisotropy i n the upper mantle, s i m i l a r to th a t observed i n other r e f r a c t i o n s t u d i e s (Snydsman et a l . , 1975; Malecek and Clowes, 1978). This p o i n t w i l l be discussed f u r t h e r . Aside from the low v e l o c i t y of the upper mantle, the second anomalous aspect of p r o f i l e EX2 i s the shallow upper mantle depth of 6.4 km compared to 9.4 km and 11.2 km f o r EX1 and EX3 r e s p e c t i v e l y . The thinner c r u s t of EX2 has r e s u l t e d mainly from the reduced thicknesses of l a y e r s 3A and 3B. The combined thickness of l a y e r s 3A and 3B f o r EX2 i s 4.0 km compared to the average value of 6.7 km f o r EX1, EX3 and EX3R. EX2R: For p r o f i l e EX2R, an upper mantle v e l o c i t y of 7.5 km/s at a depth of 8.1 km i s found. This depth, while shallower than those of EX1 and EX3, i s conside r a b l y greater than the 6.4 km upper mantle depth of EX2. The greater upper mantle depth a l s o means a g r e a t e r Pn cross-over distance - 35 km f o r EX2R compared to 25 km f o r EX2. This discrepancy between EX2 and EX2R r a i s e s a question concerning the v a l i d i t y of the assumption of l a t e r a l homogeneity required by the i n t e r p r e t a t i o n techniques. C e r t a i n l y , part of the d i f f e r e n c e i s due to the f a c t that 42 EX2 and EX2R do not form exact r e v e r s a l s of each other. Another p o s s i b i l i t y i s the e x i s t e n c e of a v e r t i c a l f a u l t along EX2, which means that the r e s u l t s obtained by assuming l a t e r a l homogeneity must be used with care. 3.5 D i s c u s s i o n Hyndman et a l . (1979) presented a model f o r the development of the Nootka f a u l t zone derived from c a l c u l a t i o n s based on the observed magnetic anomalies and other data. They showed that the o r i e n t a t i o n of the f a u l t zone has rotated from an east-west to a northeast-southwest d i r e c t i o n then r e t u r n i n g c l o s e r to east-west, a l l w i t h i n the past 8 Myr. Figure 3.13 shows the three seismic p r o f i l e s superimposed on the Hyndman et a l . (1979) model; two of the most recent o r i e n t a t i o n s of the Nootka f a u l t zone, f o r average ages of 0.5 and 1.5 Myr, are a l s o shown. According to t h i s model, the age of the Juan de Fuca p l a t e i n the study area ranges from 1 to 3 Myr and the adjacent c r u s t on the E x p l o r e r p l a t e across the f a u l t zone i s older by approximately 4 Myr. Much of the l i t h o s p h e r e o l d e r than 6 Myr on both p l a t e s has now been subducted beneath the America p l a t e . I n the d i s c u s s i o n that f o l l o w s , i t w i l l be i l l u s t r a t e d that t h i s model can e x p l a i n some aspects of the seismic r e s u l t s . A l s o , the l i t h o l o g y of o p h i o l i t e complexes, which are b e l i e v e d to be segments of oceanic c r u s t emplaced on l a n d , w i l l be shown to be r e l a t e d to the s e i s m i c v e l o c i t y s t r u c t u r e of the c r u s t i n t h i s r e g i o n . A summary of the petrology and seismic v e l o c i t y s t r u c t u r e of the Bay of I s l a n d s 43 Figure 3.13. Tectonic model of Hyndman et a l . (1979) showing the present configuration of the Explorer and Juan de Fuca plates using spreading parameters from observed magnetic anomalies. The thin solid lines with small numbers show seafloor ages; the two heavier lines with large numbers show fault locations at the time; both number sets in millions of years before the present. The approximate continental margin is indicated by the dashed line. The location of the explosive lines and CBS sites are shown by the heavy solid lines and solid circles. 44 o p h i l o l i t e complex i s given i n t a b l e 3-1 ( S a l i s b u r y and Christensen, 1978; and Christensen and S a l i s b u r y , 1979). Layer 2 Figure 3-12 shows that the top of the igneous c r u s t has a v e l o c i t y ranging from 3-7 to 4.7 km/s which i s i n good agreement with other s t u d i e s (eg. Cheung and Clowes, 1981; Spudich and O r c u t t , 1980a). Deep sea d r i l l i n g (eg. Hyndman et a l . , 1976), dredging, and s t u d i e s on o p h i o l i t e s ( S a l i s b u r y and Christensen, 1978) have e s t a b l i s h e d that the subsediment m a t e r i a l near spreading r i d g e s i s g e n e r a l l y composed of f r a c t u r e d b a s a l t flows and p i l l o w s . The v e l o c i t y of the upper l a y e r found by seismic r e f r a c t i o n surveys i s u s u a l l y much lower than t h a t observed from v e l o c i t y measurements on l a b o r a t o r y samples (Hyndman and Drury, 1976). This discrepancy i s most l i k e l y due to the f a c t that the a c t u a l c r u s t contains f r a c t u r e s , voids and i n t e r c a l a t e d sediments of a s c a l e much l a r g e r than the s i z e of the l a b o r a t o r y samples (Hyndman et a l . , 1976). These f a c t o r s tend to decrease the seismic v e l o c i t y of the upper c r u s t while the coherent samples used i n l a b o r a t o r y measurements tend to give the maximum v e l o c i t y . The increase i n v e l o c i t y w i t h depth i n l a y e r 2A i s probably a r e s u l t of the c l o s i n g of cracks and f r a c t u r e s w i t h i n the b a s a l t (Malecek and Clowes, 1978). In the t r a n s i t i o n zone from l a y e r 2A to l a y e r 2B, the v e l o c i t y increases from an average of 4.5 km/s to the 6.0 km/s range. This sudden increase i n v e l o c i t y probably marks the change from e x t r u s i v e to i n t r u s i v e l e v e l s of the c r u s t . Spudich and Orcutt (1980a) have argued Table 3.1. P e t r o l o g y and measured seismic v e l o c i t i e s (km/s) of the Bay of Isl a n d s o p h i o l i t e complex (from S a l i s b u r y and Chr i s t e n s e n , 1978 and Christensen and S a l i s b u r y , 1979) DEPTH (km) LAYER LITHOLOGY METAMORPHIC ASSEMBLAGES Vp MEASURED v s CT metabasalt p u m p i i f y ? ! ! 6 ! a c i e s 6.25 3.40 0.29 1 2B greenschist f a c i e s 6.20 3.30 0.30 brecc. d i k e s 2 -3A sheeted dikes (metadolerite) epidote-amphibolite f a c i e s <i> •H O (4 IH •U (0 •H ji (amphibolite o f a c i e s ) c 0) 0) M 00 1 6.75 3.75 0.28 3 4 5 -l a t e d i f f e r e n t i a t e s metagabbro pyroxene gabbro u 0 M H 01 > o 6.35 6.90 7.00 7.10 3.65 3.80 3.80 3.85 0.25 0.28 0.29 0.29 6 3B o l i v i n e gabbro and t r o c t o l i t e I I a 7.40 7.35 7.30 3.90 3.85 3.85 0.31 0.31 0.31 7 o •H U ctf N •H 8.70« * 4.90 0.2711 8 mantle u l t r a m a f i c a •H 4J c P. 8.20 1 0.23 9 — M a) 0) 1 * average v a l u e 0 p a r a l l e l to spreading d i r e c t i o n 1 perpendicular t o spreading d i r e c t i o n 46 that the reduction i n pore space from a porosity of 24% to 2% can only account f o r the v e l o c i t y increase within the top 0.6 km of the igneous crus t . Any further increase i n v e l o c i t y with depth must r e s u l t from other factors such as changing composition or increasing pressure. From the r e s u l t s of Salisbury and Christensen (1978), the increase i n v e l o c i t y from layer 2A to layer 2B observed i n t h i s study would correspond to the change from pillow and flow basalt to greenschist f a c i e s metabasalt and brecciated dikes found i n the upper l e v e l of the Bay of Islands o p h i o l i t e complex i n Newfoundland (see table 3.1). The reduced v e l o c i t y gradient i n layer 2B probably r e f l e c t s the f a c t that the cracks and fractures at t h i s depth are fewer i n number so that t h e i r e f f e c t on the v e l o c i t y gradient i s diminished. The boundary between layer 2B and layer 3A i s highly v a r i a b l e . The v e l o c i t y change across i t ranges from a sharp jump i n p r o f i l e EX1 to a smooth t r a n s i t i o n i n p r o f i l e EX3R. Salisbury and Christensen (1978) suggested that the v e l o c i t y d i s c o n t i n u i t y between layer 2 and layer 3 marks a metamorphic boundary between greenschist f a c i e s and epidote-amphibolite f a c i e s metabasalt and i s related to the downward migration of water along j o i n t s and f r a c t u r e s . L i t h o l o g i c a l l y , i t coincides with the t r a n s i t i o n from brecciated dikes to metadolerite sheeted dikes. From the study of the Chilean o p h i o l i t e , Stern et a l . (1976) proposed a s i m i l a r metamorphic boundary between layer 2 and 3- I f t h i s i s the case, the abnormally deep boundary between layer 2B and 3A (~3 km subsediment) may be a consequence of the complex plate motions that took place i n the region within the past 8 Myr as described by Hyndman et a l . (1979). The intense shear and f r a c t u r i n g associated with the constantly changing 47 plate boundary, exemplified by the present day Nootka f a u l t zone, would allow seawater to migrate deeply i n t o the c r u s t . The v a r i a b i l i t y of the nature of the boundary between layer 2 and layer 3 would be a r e s u l t of the varying degree of water penetration and metamorphism at the d i f f e r e n t l o c a t i o n s . One possible contradiction to t h i s hypothesis i s p r o f i l e EX2R which has the shallowest layer 2 - l a y e r 3 boundary, yet a great number of active f a u l t s are found along the p r o f i l e (Hyndman et a l , 1979). This suggests that the nature of the boundary between layer 2 and layer 3 may be much more complicated than a simple metamorphic boundary associated with water c i r c u l a t i o n . Layer 3A Previous studies have shown that layer 3 of the oceanic crust has a well-defined v e l o c i t y range (6.4 to 7.0 km/s) but i t s thickness varies considerably ( for example, see the compilation of Christensen and Salisbury, 1975). The present study i s no exception. The v e l o c i t y at the top of layer 3A has a narrow range of 6.4 to 6.6 km/s and the t o t a l thickness of layer 3 ranges from 3.1 to 7.4 km. The low 6.1 km/s l a y e r 3A v e l o c i t y of p r o f i l e EX3R can be explained by an i n t e r f a c e with medium v e l o c i t y of 6.4 km/s which dips at an angle of 4° from the horizontal away from the Juan de Fuca ridge. Such a dip near an active spreading ridge i s reasonable i n view of the f a c t that most models of the oceanic crust indicate sloping structures near the ridge c r e s t . Besides the well-defined v e l o c i t y range, the other d i s t i n c t i v e aspect of l a y e r 3A i s i t s small v e l o c i t y gradient (~0.1/s) i n contrast 48 w i t h the l a r g e r and more v a r i a b l e gradient of l a y e r 2. The u n i f o r m i t y of the v e l o c i t y s t r u c t u r e of l a y e r 3A has l e d to the p o s t u l a t i o n of uniform composition f o r t h i s part of the c r u s t (Malecek and Clowes, 1978). The l i t h o l o g y of the Bay of I s l a n d s o p h i o l i t e shows that the top 1.0 km of l a y e r 3 i s composed of metadolerite sheeted dikes w i t h a compressional wave v e l o c i t y of 6.4 to 6.8 km/s measured at pressures of 0.9 to 1.2 kbar ( S a l i s b u r y and Christensen, 1978). This c e r t a i n l y c o r r e l a t e s w e l l w i t h the r e s u l t s f o r the top of l a y e r 3A. However, l a y e r 3A of the Bay of I s l a n d s o p h i o l i t e does show some changes i n composition with depth ( t a b l e 3-D: a t h i c k l a y e r of i n t r u s i v e metagabbro and gabbro followed by a l a y e r of pyroxene gabbro are found below the sheeted d i k e s . V e l o c i t i e s w i t h i n the metagabbro vary widely between 6.0 and 6.8 km/s where the low v e l o c i t i e s are a s s o c i a t e d w i t h the discontinuous presence of l a t e d i f f e r e n t i a t e s . V e l o c i t i e s i n the pyroxene gabbro increase s l o w l y w i t h depth from 6.9 to 7.1 km/s. S a l i s b u r y and Christensen (1978) pointed out that these d e t a i l e d changes i n the v e l o c i t y s t r u c t u r e are probably too small to be detected by r o u t i n e seismic r e f r a c t i o n surveys, r e s u l t i n g i n the i n t e r p r e t a t i o n of a uniform l a y e r 3A. This can e x p l a i n the simple l a y e r 3A s t r u c t u r e i n t e r p r e t e d f o r p r o f i l e EX1 where some gaps i n shot spacing e x i s t e d between 20 and 40 km due to m i s f i r e d shots. On the other hand, p r o f i l e EX3 had a dense shot spacing and revealed more d e t a i l s w i t h i n l a y e r 3A. S i m i l a r s t r u c t u r e s , however, are not present i n the reverse p r o f i l e EX3R, rendering i t d i f f i c u l t to make any s i g n i f i c a n t i n f e r e n c e s on the f i n e s t r u c t u r e of l a y e r 3A. 49 Layer 3B The existence of a t r a n s i t i o n layer between layer 3A and the upper mantle i s well documented i n both the P a c i f i c (Sutton et a l . , 1971; Malecek and Clowes, 1978; Spudich and Orcutt, 1980a) and the A t l a n t i c (Steinmetz et a l . , 1977; Fowler and Keen, 1979; Detrick and Purdy, 1980). The v e l o c i t y i n t h i s t r a n s i t i o n layer (3B) generally varies from 7.0 to 7.7 km/s, with the thickness ranging from 2.0 to 5.0 km. The velocity-depth p r o f i l e s i n figure 3.11 i n d i c a t e that layer 3B of the crust i n t h i s region of the P a c i f i c i s best described as a zone 1.4 to 2.3 km thick, i n which the v e l o c i t y increases gradually from that of the lower crust to that of the upper mantle, with no sharp d i s c o n t i n u i t y i n v e l o c i t y at the crust-mantle i n t e r f a c e . Christensen and Salisbury (1975) suggested that the t r a n s i t i o n from layer 3A to 3B corresponds to the t r a n s i t i o n from metagabbros to fresh gabbros. In the model of Stern et a l . (1976), t h i s boundary marks the maximum depth of hydrothermal c i r c u l a t i o n and layer 3B i s also composed of fresh gabbros. The observed v e l o c i t y i n the lower h a l f of l a y e r 3B, however, i s higher than that of fresh gabbro and the presence of moderate to strong v e l o c i t y gradients implies that the composition within t h i s layer must be changing gradually. Layer 3B, therefore, can not be composed of fresh gabbro throughout. The zone of i n t e r l a y e r e d o l i v i n e gabbro, t r o c t o l i t e and p e r i d o t i t e found i n the Bay of Islands o p h i o l i t e s (Salisbury and Christensen, 1978) would be more l i k e l y to give the required v e l o c i t y and gradient i n f e r r e d from the data (see table 3.1). 50 Upper Mantle Normal upper mantle v e l o c i t i e s of 8.0 and 8.3 km/s are observed f o r p r o f i l e EX1 and the reversed p r o f i l e EX3,3R. An anomalously low v e l o c i t y of 7.5 km/s has been i n t e r p r e t e d as the upper mantle v e l o c i t y f o r the reversed p r o f i l e EX2,2R. Although these values of upper mantle v e l o c i t y are w i t h i n the range found during other s t u d i e s on the Juan de Fuca p l a t e , some exp l a n a t i o n of the d i f f e r e n c e i s necessary. Davis et a l . (1976) reported the r e s u l t s of an OBS r e f r a c t i o n survey on the Juan de Fuca r i d g e . Their r e f r a c t i o n l i n e 'east' p a r t i a l l y reverses the p r o f i l e EX1. For t h i s l i n e , they i n t e r p r e t e d an upper mantle v e l o c i t y of 8.3 km/s which i s s i m i l a r to the 8.0 km/s found f o r EX1. For l i n e s 'north' and 'south', which are p a r a l l e l to our p r o f i l e EX2, they observed no upper mantle r e f r a c t e d a r r i v a l s . From a r e f r a c t i o n survey c a r r i e d out on Explorer p l a t e near E x p l o r e r r i d g e , i n an area northwest of the Nootka f a u l t zone, Malecek and Clowes (1978) found a v e l o c i t y of 7.9 km/s f o r the upper mantle along a reversed p r o f i l e that ran perpendicular to the spreading r i d g e . On the p r o f i l e that ran p a r a l l e l to the r i d g e , the upper mantle v e l o c i t y was found to be 7.3 km/s. Anisotropy of the upper mantle was invoked to e x p l a i n the d i f f e r e n c e i n the v e l o c i t i e s . Figure 3.13 shows that p r o f i l e EX2 i s p a r a l l e l to the Juan de Fuca rid g e and i s only 23° from the magnetic l i n e a t i o n on the Exp l o r e r p l a t e . I t i s along t h i s p r o f i l e that we found the upper mantle v e l o c i t y to be low (7-5 km/s). For p r o f i l e s EX1 and EX3, the s t r i k e s of the magnetic anomaly pattern vary from 55° to 80°. Here, we found the upper mantle 51 v e l o c i t i e s to be high. This apparent anisotropy agrees w e l l w i t h the observations of Malecek and Clowes (1978). Anisotropy of the upper mantle has been w e l l documented by other s t u d i e s and i s summarized by Christensen and S a l i s b u r y (1975). G e n e r a l l y , i t i s found that the upper mantle v e l o c i t y i s high i n a d i r e c t i o n perpendicular to the r i d g e and low p a r a l l e l to i t ( R a i t t et a l . , 1969; Keen and B a r r e t t , 1971; Snydsman et a l . , 1975). Hess (1964) suggested that mantle anisotropy i s produced at the r i d g e c r e s t by the alignment of o l i v i n e c r y s t a l s i n the d i r e c t i o n of flow. Subsequently, a number of authors proposed s i m i l a r mechanisms which required a p r e f e r r e d alignment of the a-axis of o l i v i n e c r y s t a l s to create the anisotropy observed. I t i s suggested that the 10% v a r i a t i o n (defined by (1 - Vmin/Vmax) x 100%) i n upper mantle v e l o c i t y observed i n t h i s study i s due to the anisotropy of the upper mantle. Since the c r u s t a l ages are so young, t h i s i m p l i e s that the process which causes formation of the anisotropy e f f e c t takes place while new c r u s t i s being created at or near the r i d g e and t h i s e f f e c t i s preserved t h e r e a f t e r . C r u s t a l Thickness Except f o r p r o f i l e EX2, the upper mantle depths shown i n f i g u r e 3-12 are unusually deep f o r c r u s t l e s s than 6 Myr o l d . The r e s u l t i s i n l i n e with other c r u s t a l s t u d i e s undertaken on the Juan de Fuca and E x p l o r e r p l a t e s . For example, Davis et a l . (1976) i n t e r p r e t e d a c r u s t a l thickness i n excess of 11.2 km i n t h e i r r e f r a c t i o n study on the Juan de Fuca r i d g e . Malecek and Clowes (1978) reported abnormally t h i c k c r u s t (8-10 km) on the E x p l o r e r p l a t e w i t h i n 50 km of the spreading r i d g e 52 c r e s t . 'Bunching up' of the young crust on ' c o l l i s i o n ' with the America plate was suggested by Malecek and Clowes to explain the anomalous crust i n the region. On inspection of f i g u r e 3.12, i t can be concluded that whatever the mechanism i s f o r producing such thick crust, i t must do so by the thickening of layer 3, the layer that i s associated with metagabbros and cumulate gabbros. I t was remarked e a r l i e r that the differences i n the upper mantle depths interpreted f o r the p r o f i l e s must be reconciled so that the i n t e r p r e t a t i o n using l a t e r a l l y homogeneous models would be meaningful. This i s e s p e c i a l l y true f o r the reversed p r o f i l e s EX2,2R and EX3,3R- The following arguments are offered to explain the differences i n c r u s t a l thickness observed f o r these two p r o f i l e s . I t i s asserted that the velocity-depth structure of EX2 (figures 3-11b, 3.12) i s more representative of the 1 Myr old Juan de Fuca plate and i t s r e v e r s a l , EX2R, i s more i n d i c a t i v e of the c r u s t a l structure of the 5 Myr old Explorer p l a t e . The differences of the upper c r u s t a l structures and the thicknesses are r e a l . They can be a t t r i b u t e d to differences i n c r u s t a l formation and/or evolution with time, as described l a t e r , or to,the Nootka f a u l t where a 1.6 km v e r t i c a l o f f s e t i n the upper mantle depth may be present. Such an o f f s e t i s i n f e r r e d from the difference between the interpreted depths to Moho for p r o f i l e s EX2 and EX2R. Consider the tectonic model of Hyndman et a l . (1979) i n figure 3.13. For p r o f i l e EX2, shots from distances to 30 km would have t r a v e l l e d through only the 1 Myr old crust of the Juan de Fuca plate while f o r EX2R shots over the same distance range would have t r a v e l l e d 53 through only the 5 Myr old Explorer p l a t e . Interpretation of the data up to the f a u l t would produce the velocity-depth models of f i g u r e 3•11b f o r the respective plates. What about crossing the f a u l t zone? Ray t r a c i n g through a f a u l t e d model consisting of the two l a t e r a l l y homogeneous models of figure 3.11b, joined by a v e r t i c a l f a u l t , produces n e g l i g i b l e differences i n the t r a v e l times i n both d i r e c t i o n s . However, the amplitudes of a r r i v a l s that have crossed the f a u l t are expected to be somewhat attenuated. This i s consistent with the amplitude c h a r a c t e r i s t i c s of EX2 and EX2R where the Pn amplitudes at distances greater than 40 km are s i g n i f i c a n t l y lower than those predicted by the synthetic seismograms. Additional support for the i n t e r p r e t a t i o n of d i f f e r e n t c r u s t a l thicknesses on e i t h e r side of the Nootka f a u l t i s provided by EX1 which also crosses the f a u l t zone. The layer 2 structure of EX1 interpreted from trav e l paths through the 2 Myr old part of Juan de Fuca plate i s s i m i l a r to that of EX3, as we would expect. For distant a r r i v a l s , the ray paths t r a v e l p r i n c i p a l l y through the Explorer plate and Nootka f a u l t zone. Thus the lower c r u s t a l structure should be s i m i l a r to that interpreted f o r EX2R. As shown i n f i g u r e 3.12, t h i s i s the case. The f i n a l r e s u l t i s a c r u s t a l thickness intermediate between that of EX2R and EX3. I t can be noted that there i s some degree of attenuation of energy beyond 56 km, a distance for which ray paths are bottoming i n the f a u l t zone. Unlike p r o f i l e EX2, the reversed p r o f i l e EX3,3R traverses only the Juan de Fuca plate; however, the c r u s t a l age along the p r o f i l e varies 54 from 1 to 3 Myr. The obvious differences between the velocity-depth structures of EX2 and EX3 imply that there must be v a r i a t i o n of the crust with age. Since the upper mantle depths of both EX3 and EX3R are the same at 11.2 km, i t i s reasonable to assume that t h i s depth i s representative of the crust at the midpoint of the reversed p r o f i l e s where the c r u s t a l age of the Juan de Fuca plate i s 2 Myr. A comparison with p r o f i l e EX2, where the 1 Myr Juan de Fuca plate has a c r u s t a l thickness of 6.4 km, would imply that the crust thickens by 4.8 km within a time i n t e r v a l of only 1 Myr. Such a rapid rate of ' c r u s t a l maturing' would mean that, except f o r very young regions such as that of EX2, differences i n c r u s t a l thicknesses would be small between crusts of d i f f e r e n t ages on adjacent sides of transform f a u l t s . This i s confirmed by other studies on fracture zones of the Mid-Atlantic ridge (eg. Detrick and Purdy, 1980) where no change i n structure i s detected on e i t h e r side of the f a u l t . A l t e r n a t i v e l y , the differences i n c r u s t a l thickness could represent r e a l v a r i a t i o n s with time of the processes of c r u s t a l formation at the ridge. Then the inferences are that the formation process i s time varying on a time scale of l e s s than 1 Myr and the v a r i a t i o n s are established at the ridge and c a r r i e d away with the spreading process. Of course, the e f f e c t s of both time v a r i a t i o n s i n the formation process and subsequent changes as a r e s u l t of ' c r u s t a l maturing' with age could be involved. Variations of 4.8 km i n c r u s t a l thickness over distances of tens of kilometers could c a l l i n t o question the use of the WKBJ synthetic seismogram algorithm which assumes a l a t e r a l l y homogeneous earth model. A recently developed approximate method f o r c a l c u l a t i o n of synthetic 55 seismograms, combined with ray tr a c i n g , i n l a t e r a l l y varying structures (McMechan and Mooney, 1980) has enabled a check of the consistency of the i n t e r p r e t a t i o n against the observed seismic sections. The l a t e r a l l y varying model (figure 3.14) i s constructed from the homogeneous models derived from p r o f i l e s EX3 (range 0 to 15 km), EX3R (range about 30 km) and EX2 (range 45 to 60 km), with smooth va r i a t i o n s to provide model co n t i n u i t y . This model gives good agreement i n both the amplitudes and t r a v e l times of the forward and reverse p r o f i l e s EX3,3R. Although head wave contributions are not calculated by t h i s approximate method, r e s u l t i n g i n the f a c t that traces beyond 45 km have zero amplitude, the l a t e r a l l y varying model of figure 3.14 and 3.15 does o f f e r some q u a l i t a t i v e explanation to the observed amplitudes at f a r distances on both p r o f i l e s . The pronounced change i n upper mantle depth at the distance of 45 km on p r o f i l e EX3 implies that propagation of Pn phases would be hindered. The sharp drop-off i n amplitude beyond 50 km on the record section of EX3 (figure 3.3a), which the WKBJ synthetic seismograms were unable to model, could be the manifestation of t h i s l a t e r a l v a r i a t i o n i n structure. For p r o f i l e EX3R, however, the geometry of the ray paths indicates that the Pn phases would not be affected i n a s i m i l a r manner. This i s consistent with the r e l a t i v e l y large Pn amplitudes on p r o f i l e EX3R at distances ranging from 50 to 60 km (figure 3-4). No attempt i s made to match p r e c i s e l y the synthetic sections from the l a t e r a l l y varying models with the observed data. The main purpose of t h i s exercise i s to show that the observed c h a r a c t e r i s t i c s of the record 56 x x C O X - 5 T X X X x x_ -v\/vv-- # 5 x X O CD O LO O o C O g o a (0 o C\2 C Z I 0 I-(s) HA/X-I 2 t 9 9 Figure 3.14. Synthetic seismogram section and ray paths for profile EX3 calculated by the methods of McMechan and Mooney (1980). The model is constructed from the laterally homogeneous models of EX3 (range 0 to 15 km), EX3R (about 30 km) and EX2 (range 45 to 60 km), with smooth interpolation of boundaries and velocities. Generally, no velocity discontinuity existed across the boundaries. The crosses are observed travel times. VR, the reducing velocity, is equal to 6 km/s. Figure 3.15. Same as figure 3.14 for the reverse profile EX3R. Note the differences in both the ray paths and the amplitude characteristics of the synthetic seisomogram section as compared with profile EX3. 58 sections can be explained by the strong l a t e r a l v a r i a t i o n s i n c r u s t a l structure which have been interpreted along the p r o f i l e . A comparison of figures 3.14 and 3-15 with figures 3.3a and 3.4a exemplifies t h i s consistency. 3.6 Concluding Remarks The following summarizes the conclusions based on the r e s u l t s of the P wave i n t e r p r e t a t i o n presented i n t h i s chapter: (1) As c r u s t a l age increases from 1 to 2 Myr on the Juan de Fuca plate, the c r u s t a l thickness increases from 6.4 to 11.2 km. This difference i s due p r i marily to differences i n the thickness of layer 3. A rapid process of ' c r u s t a l maturing' within a 1 Myr time i n t e r v a l , v a r i a t i o n s with time i n the process of c r u s t a l formation at the ridge, or both could account for t h i s s i g n i f i c a n t change. (2) On average, sub-sediment crust i n t h i s region i s abnormally t h i c k compared to c r u s t a l sequences from other studies near spreading centres. This may be a r e s u l t of the complex i n t e r a c t i o n of the small and young Juan de Fuca and Explorer plates with the l a r g e r and older America and P a c i f i c plates. (3) The velocity-depth curves interpreted for the d i f f e r e n t p r o f i l e s are consistent i n general terms with the model of oceanic crust represented by o p h i o l i t e complexes. Variations i n d e t a i l e x i s t , as would be expected for materials formed i n d i f f e r e n t regions. (4) V e l o c i t y anisotropy within the upper mantle i s found. The v e l o c i t y varies from 7.5 km/s i n a d i r e c t i o n p a r a l l e l to the ridge to 8.3 km/s i n a d i r e c t i o n approximately perpendicular to i t , giving a 59 10% a n i s o t r o p i c e f f e c t . This concurs with previous studies i n the region. (5) Seismic energy i s attenuated by the Nootka f a u l t zone. The three p r o f i l e s which cross the f a u l t zone show a noticeable decrease i n the amplitude of seismic phases at distances corresponding to ray paths traversing i t . In the following chapter, i t w i l l be shown that some of the findings based on the P wave i n t e r p r e t a t i o n w i l l be substantiated by the independent i n t e r p r e t a t i o n of the shear wave data. 60 4. INTERPRETATION OF SHEAR WAVE DATA 4.1 I n t r o d u c t i o n The importance of shear wave i n f o r m a t i o n i n determining the p h y s i c a l p r o p e r t i e s and mineralogy of the oceanic c r u s t has been propounded by.a number of authors (Christensen and S a l i s b u r y , 1975; Hyndman, 1979; and Spudich and Or c u t t , 1980a). S t i l l , very l i t t l e i s known about the S wave v e l o c i t y s t r u c t u r e of the oceanic c r u s t and upper mantle, despite the f a c t that ocean bottom seismometers have been i n widespread use f o r r e f r a c t i o n s t u d i e s i n recent years. The f i r s t part of t h i s chapter o u t l i n e s some of the problems i n observing shear wave a r r i v a l s i n marine r e f r a c t i o n surveys and discusses methods f o r improving the s i g n a l q u a l i t y of these a r r i v a l s . Out of the f i v e r e f r a c t i o n p r o f i l e s i n t h i s study, two have recorded u s e f u l shear wave a r r i v a l s . In the second part of t h i s chapter, the i n t e r p r e t a t i o n of these two S wave p r o f i l e s i s discussed. I t i s shown that the r e s u l t s of the shear wave i n t e r p r e t a t i o n s complement those of the P waves presented i n the previous chapter. On the other three p r o f i l e s , shear a r r i v a l s can be observed; however, due to e i t h e r poor a r r i v a l p i c k s or e r r a t i c amplitude behaviour of the S phases, no i n t e r p r e t a t i o n of the r e f r a c t e d S waves i s attempted. The next s e c t i o n discusses some of the f a c t o r s that a f f e c t the amplitude of the converted S waves. 61 4.2 Shear Wave Conversion In marine r e f r a c t i o n experiments, only compressional energy i s generated i n the water. For shear waves to be observed, the P waves from the source must be converted to S mode at some boundary i n the media. The most l i k e l y l o c a t i o n i n the oceanic c r u s t f o r t h i s conversion to take place i s at the water-sediment or the sediment-basement i n t e r f a c e where the c o n t r a s t i n seismic p r o p e r t i e s across the boundary i s l a r g e . However, the e f f i c i e n c y of conversion and t r a n s m i s s i o n of shear waves i n the c r u s t i s h i g h l y v a r i a b l e , l e a d i n g to the unpredictable nature of the S wave a r r i v a l s . Spudich and Orcutt (1980a) and White and Stephen (1980) have discussed i n d e t a i l some of the f a c t o r s a f f e c t i n g the conversion of P to S waves. A summary of the more important aspects i s given i n the f o l l o w i n g . (a) I f Poisson's r a t i o of the sediment i s high, such as i n unconsolidated marine sediment, there w i l l be very l i t t l e conversion of P to S waves at the water-sediment i n t e r f a c e . However, the conversion w i l l be strong at the sediment-basement i n t e r f a c e (Spudich and Helmberger, 1979). (b) E f f i c i e n c y of conversion i s reduced i f the basement Poisson's r a t i o i s h i g h . Weathering and cracks can i n c r e a s e the Poisson's r a t i o of the igneous basement; t h e r e f o r e l a t e r a l v a r i a t i o n i n weathering or non-uniform d i s t r i b u t i o n of cracks can cause d r a s t i c v a r i a t i o n s i n the amplitude of the converted S waves (White and Stephen, 1980). 62 (c) Phase coherency of the converted S waves i s affected by rough topography on the basement i n t e r f a c e , e s p e c i a l l y i f the scale of the topographical v a r i a t i o n i s i n the order of a seismic wavelength (~ 300 m). Even i n regions where seismic v e l o c i t y contrast i s favourable, i e . Poisson's r a t i o being high i n the sediment and low i n the basement, rough topography can cause poor S a r r i v a l s (Spudich and Orcutt, 1980a). (d) Conversion e f f i c i e n c y i s strongly dependent on the incident phase v e l o c i t y , which i s a function of angle of incidence, as well as the basement P v e l o c i t y . For phase v e l o c i t y equal to the basement P v e l o c i t y , no conversion w i l l take place. Conversion remains low for phase v e l o c i t y higher than the basement P v e l o c i t y (White and Stephen, 1980). (e) I f the P v e l o c i t y of the sediment exceeds the S v e l o c i t y of the basement, no conversion w i l l take place (Spudich and Helmberger, 1979). Since converted S waves are observed i n a l l of the p r o f i l e s i n t h i s study, t h i s l a s t remark probably does not apply to t h i s area. I t remains to be determined where the P and S conversion takes place. Two sets of evidence i n d i c a t e that the main refracted S a r r i v a l s are converted from the down-going P waves at the sediment-basement i n t e r f a c e . F i r s t , i f the conversion had taken place at the water-sediment i n t e r f a c e , the intercept time of the refracted S a r r i v a l s would be 1.5 s l a t e r than 63 that observed due to the low S v e l o c i t y of the sediment (^0.5 km/s). Second, i t i s shown i n s e c t i o n 4.4 t h a t Poisson's r a t i o of the sediment l a y e r i n t h i s region i s high (>0.45) which i m p l i e s that strong conversion of P to S waves i s to be expected from the sediment-basement i n t e r f a c e w i t h l i t t l e or no conversion o c c u r r i n g at the s e a f l o o r . Counteracting the favourably high Poisson's r a t i o of the sedimentary l a y e r i s the rough topography of the basement i n the study area. Large s c a l e (<-<-. 1 km) v a r i a t i o n s i n the basement depth have been shown i n Chapter 2 to have caused l a r g e t r a v e l time anomalies i n the P waves. D e t a i l e d examinations of the r e f l e c t i o n p r o f i l e s of Davis and L i s t e r (1977) and Hyndman et a l . (1979) re v e a l many small s c a l e (~ 300 m) basement topographical disturbances, e s p e c i a l l y i n the v i c i n i t y of the Nootka f a u l t zone. This could e x p l a i n the poor S wave a r r i v a l s i n some of the p r o f i l e s , although except f o r l a r g e features such as the Nootka f a u l t zone, d e t a i l e d c o r r e l a t i o n of topographical f e a t u r e s with the converted S wave amplitude v a r i a t i o n i s d i f f i c u l t to make. 4.3 O p t i m i z a t i o n of Shear Wave Observations 4.3-1 Shear Wave Enhancement Where three component seismogram data are a v a i l a b l e , s e i s m i c a r r i v a l s can be enhanced by the use of p o l a r i z a t i o n f i l t e r s . Cheung (1978) described i n d e t a i l a p a r t i c u l a r a p p l i c a t i o n of such f i l t e r s to OBS r e f r a c t i o n data to improve the S wave a r r i v a l s . However, due to the 64 large uncertainties i n the estimation of the emerging angle and the azimuth of the seismic phases, the improvement i n data q u a l i t y was l i m i t e d . Only one out of four sections that he processed benefited s l i g h t l y from the a p p l i c a t i o n of the two p o l a r i z a t i o n f i l t e r s . No further attempt i s made i n t h i s study to u t i l i z e such f i l t e r s . However, some improvement i n signal-to-noise r a t i o f o r the S wave data can be achieved by simply constructing SV ( v e r t i c a l l y p o l a r i z e d shear wave) seismograms from the two horizontal components. Figure 4.1 i l l u s t r a t e s the proje c t i o n of r e c t i l i n e a r SV motion on the horizontal plane where X and Y are the two orthogonal horizontal components. Knowing the angle oc, SV(t) may be obtained from X(t) and Y(t) as SV(t) = X(t) / C0S(*) + Y(t) / SIN(*) / 2 (4.1) In t h i s process, the coherent signal i s expected to be reinforced while any random noise w i l l be cancelled out. This operation can also be regarded as a rot a t i o n of the horizontal seismograms into the r a d i a l d i r e c t i o n of motion. I t has been assumed that the v e r t i c a l component of the SV motion i s n e g l i g i b l e due to the near v e r t i c a l incident angles which are generally l e s s than 10° from v e r t i c a l . In order to f i n d the angle <*, the o r i e n t a t i o n of the horizontal seismometers i n the free f a l l OBS's must f i r s t be determined. This may be achieved by using the r e l a t i v e amplitude of the two horizontal seismograms f o r a refracted P a r r i v a l . This approach, however, has two disadvantages: the refracted ray path i s strongly affected by near-65 SV Figure 4.1. The horizontal plane projection of SV motion, <*. is the angle!that the incident plane makes with respect to the X component. 66 surface inhomogeneities and the n e a r - v e r t i c a l P a r r i v a l s have very small horizontal components of motion. A better method i s to use the r e l a t i v e amplitude of the d i r e c t water wave a r r i v a l s which are not affected by any l a t e r a l v a r i a t i o n s i n the sub-bottom structures and usually have adequate amplitudes for r e l i a b l e estimates of the seismometer ori e n t a t i o n s . Figure 4.2 shows the p a r t i c l e motion on the horizontal plane for a number of d i r e c t water wave a r r i v a l s on 0BS1. The angle i s found by f i t t i n g a regression l i n e through the locus of the p a r t i c l e motion for a one ha l f second duration. When shots from various azimuths are used to check for the consistency of oc, a variance i n the order of 5° i s found. This i s taken to be the uncertainty i n the estimate. A set of seismograms for shot 23 of p r o f i l e EX3 i s shown i n figure 4.3- The dashed l i n e s denote a number of a r r i v a l s whose ray paths are shown schematically i n figure 4.7. Trace (a) of figure 4.3 i s the v e r t i c a l component while traces (b) and (c) are the horizontal components, Y and X r e s p e c t i v e l y . Trace (d) i s produced from (b) and (c) v i a equation (4.1) and i t represents the r a d i a l component of motion. Two d i s t i n c t i v e S a r r i v a l s can be .seen on the horizontal seismograms. As expected, these a r r i v a l s r e g i s t e r very small amplitudes on the v e r t i c a l component seismogram. The main S phase, PSS, originates from conversion of P to S waves at the sediment-crust i n t e r f a c e below the shot point while the PPS phase t r a v e l s as a P wave through most of i t s t r a v e l path and undergoes a conversion to an S wave at the sediment-crust i n t e r f a c e immediately Figure 4.2. Horizontal particle motion plots of four direct water wave arrivals. On each plot, the •+• symbols define the regression line fitted to the locus of the motion. The orientation of the horizontal seismometers is determined from the angle that this line makes with respect to the X axis. PPS PWW PSS second n«ire U 3 Trace (a) is the vertical component while traces (b) and (c) are the SJSwn^'oSSnento of a 30 kg shot recorded at a distance of 25 km. Trace (d) is the radial component of motion constructed from traces (b and (c). Trace e) is the bandpass filtered version of (d). Bandpass frequency limits are 2 and 10 Hz. 69 below the rec e i v i n g OBS. Both S a r r i v a l s are s l i g h t l y enhanced i n trace (d), although the improvement i s not very obvious due to the already high signal-to-noise r a t i o of the o r i g i n a l data. In contrast, figure 4.4 shows a set of seismograms which have a higher noise l e v e l than the previous example. Here, the main PSS phase i s c l e a r l y seen i n the SV seismogram of (d); whereas, the same a r r i v a l on the two horizontal traces i s i n t e r f e r e d with by the wave coda of e a r l i e r a r r i v a l s . These two examples i l l u s t r a t e that the r o t a t i o n a l process described by equation (4.1) can provide some enhancement of the shear a r r i v a l s . I t should be noted that i n practice, i f the angle i s near either 0 or 90 , as i n the case of p r o f i l e EX2, equation (4.1) i s not used. Instead, the horizontal component with the larger amplitudes i s considered as the SV trace and no rot a t i o n i s necessary. Further improvement i n the appearance of the seismogram can be obtained by bandpass f i l t e r i n g . Trace (e) on both figures 4.3 and 4.4 are the SV seismograms bandpass f i l t e r e d from 2 to 10 Hz. Figure 4.5 shows the r e l a t i v e Fourier power spectra of the two S phases and i t i s evident that they have well-defined peaks i n the frequency band so that f i l t e r i n g would help to reduce the noise. The waveforms i n the f i l t e r e d traces are smoothed and most of the high frequency noise has been removed. This w i l l f a c i l i t a t e the comparison with synthetic seismograms for amplitude i n t e r p r e t a t i o n . However, as i n the case of the P wave data, bandpass f i l t e r i n g usually does not help i n the picking of the shear wave a r r i v a l times. Figure 4.4. Same description as figure 4.3 for a 5 kg shot recorded at a distance of 45 km. The main PSS phase shows clearly in (d) and (e), whereas the arrival about 0.4 s earlier interferes with PSS in (b) and (c). o 0.0 o c o o o c n o Relative Power 0.2 0.4 0.6 0.8 1.0 0 Relative Power 0.2 0.4 0.6 0.8 _____ r GO 1.0 < — _ _ _ _ _ _ _ r > T co GO Figure 1.5. Relative Fourier power spectra of the PSS and PPS phases (see trace (c) of figure 4.3). Spectra are calculated from signal durations of 1 second. 72 4.3.2 Three Dimensional P a r t i c l e Motion Section One of the major d i f f i c u l t i e s i n observing shear waves i n marine r e f r a c t i o n experiments i s the f a c t that the refracted shear waves often a r r i v e i n coincidence with some of the P phases, most noticeable of which are the water multiples of the main refracted P wave. This makes i t d i f f i c u l t to i d e n t i f y the onset of the S a r r i v a l s . In order to discriminate between the various a r r i v a l s , i t i s necessary to compare a r r i v a l s on the horizontal component section with those on the v e r t i c a l component. This can be a tedious process. A possible solution proposed here i s to construct a three dimensional p a r t i c l e motion section on which information from the three components are displayed simultaneously. An example of such a section i s given i n f i g u r e 4.6 f o r p r o f i l e EX3. In i s o t r o p i c media, p a r t i c l e motion i s r e s t r i c t e d to the incident plane; therefore, only two components, the v e r t i c a l and the r a d i a l , are needed to define the p a r t i c l e motion. The construction of the r a d i a l component from the two orthogonal ho r i z o n t a l seismograms has been described i n section 4.3.1. In f i g u r e 4.6, the time axis i s horizontal and increasing to the r i g h t . V e r t i c a l motions are plotted up and down with respect to t h i s axis while horizontal motions are plotted at a perspective of 45° with respect to t h i s a x is. Some of the major a r r i v a l s are l a b e l l e d on the section and the ray paths corresponding to these a r r i v a l s are shown i n f i g u r e 4.7. I d e n t i f i c a t i o n of the various phases i s f a c i l i t a t e d i n two ways by the p a r t i c l e motion section. F i r s t , d i s t i n c t i o n between compressional and shear waves can be made by considering the sense of motion of an PROFILE EX3 DIST/6 (SEC) F i g u r e 4.6. Three dimensional p a r t i c l e motion s e c t i o n for profile EX3. The reduced t r a v e l time a x i s i s h o r i z o n t a l and i n c r e a s i n g to the r i g h t . V e r t i c a l motions are p l o t t e d up and down w h i l e h o r i z o n t a l motions are p l o t t e d a t a p e r s p e c t i v e of 45 w i t h respect to the time a x i s . The d i s t a n c e a x i s i s a l s o a l i g n e d at 45 w i t h r e s p e c t to the time a x i s . Topographical c o r r e c t i o n s to the t r a v e l times are c a l c u l a t e d u s i n g the phase v e l o c i t i e s of the PSS a r r i v a l s . See f i g u r e 4.7 f o r d e f i n i t i o n of the ray paths of the various phases. 74 Figure 4.7. Schematic diagram showing the ray path of the various seismic phases that can be seen in the section of figure 4.5. 75 a r r i v a l , i e . motion would be predominantly v e r t i c a l f o r P waves versus the predominantly horizontal motion f o r S waves. Second, due to the perspective p l o t , the coherency between traces i s enhanced so that the differences i n the phase v e l o c i t i e s of the a r r i v a l s are more obvious. This serves to further i d e n t i f y the d i f f e r e n t phases. One good example of these two points i s i l l u s t r a t e d by the converted shear phase PPS. This phase lags behind the f i r s t P a r r i v a l by an almost constant amount of time which i s equal to the difference i n the one way t r a v e l time through the sediment f o r P and S modes. The PPS a r r i v a l s have the same phase v e l o c i t y as the P, yet the p a r t i c l e motion i s c l e a r l y h o r i z o n t a l , i d e n t i f y i n g them as shear waves. The PPS phases are important i n defining the Poisson's r a t i o of the sedimentary l a y e r . This w i l l be discussed i n the next s e c t i o n . 4.4 Poisson's Ratio of the Sediments I t w i l l be shown i n t h i s section that an accurate estimate of Poisson's r a t i o i n the sediment layer can be obtained from the t r a v e l times of the PPS phases, without e x p l i c i t l y c a l c u l a t i n g the P and S v e l o c i t i e s . This new procedure assumes constant v e l o c i t i e s within the sediment and the Poisson's r a t i o thus determined i s an average value. Consider that the sedimentary layer beneath the OBS has a thickness of h and that the P and S v e l o c i t i e s , v P and v s , are constant within the l a y e r . Let tp and t$ be the one way v e r t i c a l t r a v e l times through the sediment f o r the P and S waves r e s p e c t i v e l y . We have 76 tp = h/vp t s = h/v s I f we define At = t s - tp then vp/vg = tg/tp = 1 + at/tp (4.2) Now Poisson's r a t i o i n terms of Vp/v^ i s a = ( ( v p / v s ) 2 - 2} / { 2 ( v P / v s ) 2 - 2 } (4.3) Substituting (4.2) into (4.3) we obtain cr = ( 1 - p / 2 (4.4) where | = {2( t / t p ) + ( t / t P ) 2 } -1 (4.5) So Poisson's r a t i o of the sediment can be calculated by knowing only the values of _ t and t p . i f we assume that the P and PPS phases have v e r t i c a l t r a v e l paths, then _ t can be e a s i l y measured from the t r a v e l time lag between the two phases. The error introduced by t h i s assumption i s i n the order of 1$ which i s n e g l i g i b l e compared to the uncertainties of other f a c t o r s . The v a r i a b l e t p i n equation (4.5) can be obtained from the r e f l e c t i o n p r o f i l e s over the OBS s i t e , and i n t h i s study, the values of tp for the three s i t e s are interpreted from the r e f l e c t i o n p r o f i l e s of Davis and L i s t e r (1977) and Hyndman et a l . (1979). Table 4.1 tabulates Poisson's r a t i o for the sedimentary layer 77 beneath the three OBS s i t e s , c a l c u l a t e d with the value of At and tp given i n the same t a b l e . N denotes the number of observations averaged to give At. TABLE 4.1 Poisson's r a t i o of the sediment l a y e r . OBS N A t (s) t p (s) Poisson's R a t i o 1 15 1.50 ± 0.01 0.67 i 0.05 0.45 + 0.01 2 7 1.22 + 0.01 0.44 + 0.05 0.46 + 0.01 3 15 1.42 + 0.01 0.56 + 0.05 0.46 + 0.01 The values of o given i n t a b l e 4.1 f o r the three s i t e s are s i m i l a r and are c o n s i s t e n t l y h i g h . This i n d i c a t e s that the sediment i n t h i s region i s uniform and poorly c o n s o l i d a t e d , a f a c t a l s o v e r i f i e d by other s t u d i e s i n the area. For example, Cheung (1978) found values of a i n the neighbourhood of 0.49 f o r the sediment column i n the E x p l o r e r Ridge region. Based on the f i n d i n g of low P v e l o c i t i e s w i t h i n the sediment l a y e r from an OBS r e f r a c t i o n survey, Davis et a l . (1976) concluded that the sediment i n the Juan de Fuca Ridge region i s only s l i g h t l y compacted. Rapid sedimentation r a t e s (up to 1 cm/yr) were given as the cause of the poor c o n s o l i d a t i o n . So f o r t h i s near c o a s t a l region of the P a c i f i c Ocean, high values of o" are to be expected. 78 4.5 Shear Wave Record Sections and Amplitude Modelling Despite the considerable e f f o r t s made i n enhancing the shear wave a r r i v a l s , only two out of the f i v e p r o f i l e s give s u f f i c i e n t data f o r the i n t e r p r e t a t i o n of the PSS phases. The SV record sections of p r o f i l e s EX3 and EX2 are plotted with a reducing v e l o c i t y of 4 km/s i n f i g u r e s 4.8a and 4.9a. The timings of the traces have been corrected f o r topographical v a r i a t i o n s using the phase v e l o c i t y of the PSS phases derived from the slope of the uncorrected t r a v e l time curves. Amplitude normalization f o r the d i f f e r e n t charge s i z e s and the a p p l i c a t i o n of an r 2 spreading factor follow the same procedures as those discussed f o r the P wave record sections. The data traces are bandpass f i l t e r e d between 2 and 10 Hz. The choice of frequency l i m i t s lower than those chosen f o r the P wave data stems from the fac t that the S a r r i v a l s generally have lower frequency content than the P waves. By glancing at the two record sections of figures 4.8a and 4.9a, one can e a s i l y detect the obvious diff e r e n c e i n t h e i r amplitude behaviour; namely, the large amplitude a r r i v a l s associated with the near c r i t i c a l r e f l e c t i o n from the gradient zone immediately above the cr u s t -upper mantle boundary are found at a distance range of 45 km for p r o f i l e EX3 versus 23 km f o r p r o f i l e EX2. This difference i s a consequence of the d i f f e r e n t c r u s t a l thicknesses, already detected by the P wave data, f o r the two p r o f i l e s . A s i m i l a r contrast i n the amplitude c h a r a c t e r i s t i c s of the two p r o f i l e s can be seen also i n the P wave record sections (figures 3.3a and 3«5a). Other points to be noted about the S wave record sections are the 79 i 1—;—i h o i 1 1 r 9 9 P £ 9 S £ (339) W19IQ - 1 (339) W19IQ - 1 Figure 4.8. (a) SV record section and (b) best f i t synthetic seismogram section for profile EX3. Synthetic seismograms are calculated using the S velocity-depth model of figure 4.10a. 80 Figure 4.9. (a) SV record s e c t i o n and (b) best f i t s y n t h e t i c seismogram s e c t i o n f o r p r o f i l e EX2. S y n t h e t i c seismograms are c a l c u l a t e d u s i n g the S v e l o c i t y - d e p t h model of f i g u r e 4.10b. 81 emergent character of the f i r s t breaks, and the low amplitudes of the a r r i v a l s at near distances and at distances beyond the cross-over point of the Sn phases. The emergent f i r s t breaks introduce greater uncertainties i n t o the picking of the S a r r i v a l s . Travel time errors are i n the order of +0.08 for the S waves versus +0.06 s for the P waves. The lack of energy f o r the near distance a r r i v a l s means that the airgun p r o f i l e s , so v i t a l l y important i n providing the constraints on the upper c r u s t a l structures, are unfortunately useless f o r shear wave i n t e r p r e t a t i o n . A d d itional complication f o r the observation of near distance S a r r i v a l s i s the interference from the large amplitude water waves which a r r i v e i n approximately the same time window. Having i l l u s t r a t e d some of the l i m i t a t i o n s of the S wave data, i n t e r p r e t a t i o n of the two p r o f i l e s i s presented i n the following. Inversion of the t r a v e l time data, using the tau(p) function, proceeds i n the same way as f o r the P waves. Due to the greater uncertainty i n the t r a v e l times, the extremal bounds calculated f o r the S v e l o c i t y -depth functions are even more lac k i n g i n resolving power than the P v e l o c i t y bounds. I t i s found that constraints on the S v e l o c i t y structure come mainly from the amplitude data; therefore, the discussion that follows w i l l concentrate on the modelling of the S wave amplitude, using synthetic seismograms calculated by the WKBJ algorithm. The WKBJ algorithm allows inter-mode conversion at any boundary of the model so that the PSS phases can be modelled d i r e c t l y both i n t r a v e l time and amplitude. This i s p a r t i c u l a r l y important since the conversion of P to S at the sediment-basement i n t e r f a c e i s strongly dependent on 82 the phase v e l o c i t y and the v e l o c i t y contrast present. Certain assumptions concerning the S v e l o c i t y structure at shallow depths have to be made since no refracted S a r r i v a l s are observed from these depth regions. A good estimate of the S v e l o c i t y i n the sediment comes from the PPS phases (see discussion i n the previous s e c t i o n ) . With a Poisson's r a t i o of 0.45, the P v e l o c i t y of 1.8 km/s i n the sediment would give 0.5 km/s f o r the S waves. For t h i s reason, the sediment layer i s constrained to have a constant v e l o c i t y of 0.5 km/s and a uniform thickness of 1.0 km. The existence of converted shear waves places the lower l i m i t of the basement S v e l o c i t y at 1.8 km/s, equal to the sediment P v e l o c i t y . Basement S v e l o c i t y of 2.2 km/s i s chosen to give reasonable amplitudes f o r the observed PSS phases, which are p a r t i a l l y c o n t r o l l e d by the v e l o c i t y contrast at the sediment-basement i n t e r f a c e . This value of basement S v e l o c i t y i s s i m i l a r to those found f o r the Gaudalupe s i t e by Spudich and Orcutt (1980a). With the exception of the constraints mentioned above for the shallow structures, the S wave data are i n i t i a l l y modelled independently of the P wave data. The t r i a l - a n d - e r r o r approach employed i n modelling the P wave amplitude data i s used also f or the S waves. I t i s found from the independent modelling that the i n t e r f a c e s marking zones of d i f f e r e n t v e l o c i t y gradients coincide to within 200 m with those found f o r the P velocity-depth functions. This difference i s i n s i g n i f i c a n t considering that the t y p i c a l wavelength of c r u s t a l shear waves i s approximately 0.8 km, so i t i s deemed j u s t i f i a b l e to constrain the i n t e r f a c e s to be at 83 the same depths for the ease of i n t e r p r e t i n g the r e s u l t s . The best f i t synthetic seismogram sections are plotted below the data sections i n figures 4.8b and 4.9b. As with the P wave presentation, the t r a v e l time curves from the synthetic sections are transferred onto the data s e c t i o n . Generally, the synthetic record sections give excellent f i t to the observations. An exception i s found i n p r o f i l e EX3 where the traces at distances beyond 48 km show a sudden decrease i n amplitude which i s not modelled adequately by the WKBJ seismograms. This phenomenon i s s i m i l a r to that observed i n the P wave record section of p r o f i l e EX3 and i t lends support to the in f e r r e d l a t e r a l v a r i a t i o n i n cr u s t a l thickness at approximately 45 km distance range, as discussed i n the previous chapter. 4.6 Results and Discussion In t h i s section, the shear wave velocity-depth models for p r o f i l e s EX2 and EX3, interpreted by modelling the t r a v e l time and amplitude data with the WKBJ synthetic sections, are discussed. The r e l a t i o n s h i p between the seismic v e l o c i t y of oceanic crust and the l i t h o l o g y of o p h i o l i t e complexes, expounded upon i n d e t a i l i n section 3.5, i s re-examined here i n l i g h t of the added information from the shear wave data. Figure 4.10 shows the f i n a l vp and vg models f o r p r o f i l e s EX2 and EX3. The zones of d i f f e r i n g v e l o c i t y gradients, more obvious i n the vp curves, are also evident i n the vg curves. A schematic representation of the v e l o c i t y structures, s i m i l a r to figu r e 3.12 of the l a s t chapter but X H-l_lg CD 1 3 (D Q . W • X -» CO o •X3 0 E X 3 Velocity (km/s) 2 4 6 8 i CO o O. 3 CD 0) 3 a < a cu co 3 * CD a. a CD s o a. CD t— u o T •a i o s cvH coH Q CO O J CD W 0 2 85 EX3 EX2 v P V S V P v s LAYER o -1.8 0.5 1.8 0.5" 1 4.7 2.2 3.7 2.2 s 2A CM -5.0 2.7 4.4 2.5 2B 6.4 3.6 - 6.6 3.6 3A vO -6.9 3.7 3B 7.5 4.5 00 -7.7 4.1 MANTLE O _ i-l CM 8.3 4.6 Fi g u r e 4.11. Schematic r e p r e s e n t a t i o n of the v e l o c i t y - d e p t h models of f i g u r e 4.10. The depth i s measured from the s e a f l o o r . The numbers are the v e l o c i t y at the top of each l a y e r i n km/s. Layer 1, the sediment l a y e r , i s co n s t r a i n e d to have the given v e l o c i t i e s and a t h i c k n e s s of 1 km. 86 w i t h vg inf o r m a t i o n added, i s shown i n f i g u r e 4.11 to f a c i l i t a t e the d i s c u s s i o n of the r e s u l t s . In f i g u r e 4.12, the apparent Poisson's r a t i o versus depth f o r the two p r o f i l e s are p l o t t e d . The values of Poisson's r a t i o are only apparent because the r e l a t i o n s h i p between o and Vp/vg (equation 4 .3) used i n the c a l c u l a t i o n i s appropriate only f o r i s o t r o p i c media. In the l a s t chapter, an a n i s o t r o p i c e f f e c t of 10% i s found f o r the upper mantle P v e l o c i t y ; t h e r e f o r e , i n the neighbourhood of upper mantle depths, values of Poisson's r a t i o i n f i g u r e 4.12 are f o r the purpose of comparison only, and they do not represent the a c t u a l e l a s t i c parameter at those depths. Further d i s c u s s i o n of v e l o c i t y anisotropy w i l l be given l a t e r . Spudich and Orcutt (1980a) have suggested a method of d i s p l a y i n g s e i s m i c v e l o c i t y models i n which the vp(z) and vg(z) are p l o t t e d as a continuous path through the Vp-vg plane, parameterized by depth z. The main advantage of t h i s p r e s e n t a t i o n i s that r e s u l t s from l a b o r a t o r y samples can be conveniently p l o t t e d on the same diagram. In f i g u r e 4.13 the Vp-vg paths of EX3 and EX2, w i t h subsediment depths l a b e l l e d along the paths at 1 km i n t e r v a l s , are shown. Paths of constant Poisson's r a t i o which form s t r a i g h t l i n e s i n the vp-vg plane a l s o are p l o t t e d . An enlarged s e c t i o n of the two paths i s shown i n f i g u r e 4.14 along w i t h some l a b o r a t o r y measurements of o p h i o l i t e samples. The l a b o r a t o r y data are compiled by Spudich and Orcutt (1980a) who have adjusted the o r i g i n a l measurements of Christensen (1978) and S a l i s b u r y (1974) on o p h i o l i t e samples to pressure c o n d i t i o n s appropriate f o r comparison w i t h t h e i r seismic models i n t e r p r e t e d f o r the Gaudaliipe s i t e . These same la b o r a t o r y data can be used f o r comparison with the r e s u l t s of the 87 0.2 o Poisson's Ratio 0.3 0.4 0.5 C\2-Q CO O J CM 1 • • • \ \ \ / M — EX 3 / — EX 2 M ' Figure 4.12. Apparent Poisson's r a t i o versus depth f o r p r o f i l e s EX2 and EX3 c a l c u l a t e d from the velocity-depth models of f i g u r e 4.10. The question mark emphasizes the uncertainty of the estimate of Poisson's r a t i o i n l a y e r 2A. The symbol 'M' denotes the crust-mantle i n t e r f a c e . 2 3 4 5 S Velocity (km/s) Figure 4.13. v p - v s paths for EX2 and EX3. Subsediment depths in 1 km intervals are marked along the paths with numbers above the paths for EX3 and numbers below the paths for EX2. Paths of constant Poisson's ratio, which form straight lines in the Vp-vg plane, are plotted at intervals of 0.05. 89 x E • S A G C D P • 0 x T o M S Velocity (km/s) F i g u r e 4.14. v p - v s paths f o r EX2 and EX3 compared w i t h o p h i o l i t e sample v e l o c i t i e s . The rock types are: E - e p i d o t e - a m p h i b o l i t e f a c i e s m e t a d o l e r i t e and metagabbro, S - p a r t i a l l y s e r p e n t i z e d p e r i d o t i t e , G -g r e e n s c h i s t f a c i e s m e t a d o l e r i t e , P - pyroxene gabbro, 0 - o l i v i n e gabbro, T - trondhjemite, and M - metabasalt. 90 * H — r - , 1 2 3 4 5 S Velocity (km/s) F i g u r e 4.15. A comparison of EX2 and EX3 w i t h the s e i s m i c models, FF2 and FF4, o f Spudich and Orcutt (1980a). 91 present study since the small changes i n v e l o c i t y (~0.01 km/s for Vp) due to the s l i g h t v a r i a t i o n s i n depths between the seismic models are n e g l i g i b l e . S i m i l a r i t i e s between the seismic models interpreted by Spudich and Orcutt f o r the Gaudalupe s i t e (FF2 and FF4) and p r o f i l e s EX2, EX3 can be seen i n figure 4.15 where the Vp-vg paths coincide i n the vP range of 6.0 to 7.0 km/s. In the following, a discussion of the seismic r e s u l t s of the present study and t h e i r r e l a t i o n s h i p with the laboratory r e s u l t s i s given i n order of increasing depths. Layer 2A There are no d i r e c t constraints on the S v e l o c i t i e s i n layer 2A (1.0 to 2.0 km sub-bottom depth). The choice of 2.2 km/s for Vg (see fi g u r e 4.11) at the top of t h i s layer i s based on the f a c t that i t gives reasonable conversion of P to S energy at t h i s i n t e r f a c e . By a s i m i l a r argument, Spudich and Orcutt (1980a) also interpreted a basement Vg of 2.2 km/s f o r the Gaudalupe s i t e . However, changing the value of basement Vg by +0.3 km/s produces only minor differences i n the amplitudes of the PSS phases. Since no strong s u b c r i t i c a l r e f l e c t i o n s are observed from the top of t h i s l a y e r , the existence of an S v e l o c i t y gradient within the layer i s required. However, the exact nature of t h i s gradient i s not defined by the data so that the vg models for t h i s depth region are constructed i n a rather ad hoc manner to give reasonable intercept time for the deeper a r r i v a l s . Due to the uncertainty of vg, the value of Poisson's r a t i o i n t h i s layer i s not r e l i a b l e ; a question mark i s placed on f i g u r e 92 4.12 to emphasize t h i s uncertainty. The inference that t h i s layer i s composed of fractured basalt flows and pillows i s consistent with the Vg r e s u l t s . T y p i c a l laboratory measurements of basalt samples dredged from the seafloor give values of v s i n the neighbourhood of 2.5 km/s (Hyndman and Drury, 1976). The presence of large scale cracks and voids i n the i n s i t u oceanic crust can be expected to reduce Vg to the order of 2.2 km/s. Layer 2B In the t r a n s i t i o n zone from layer 2A to 2B, both vp and Vg increase r a p i d l y with depth. This increase i n v e l o c i t i e s i s thought to correspond to the change from pillow and flow basalt to greenschist f a c i e s metabasalt and brecciated dikes found i n the upper l e v e l of the Bay of Islands o p h i o l i t e complex (Christensen and Salisbury, 1978; table 3-D-Within layer 2B, the most s t r i k i n g aspect of the seismic r e s u l t s i s the very low values of Poisson's r a t i o which are determined there (see figure 4.12 for example). Upon comparing the vp-vg paths with the re s u l t s from o p h i o l i t e samples (figure 4.14), we see that the correspondence between the vp-vg paths and the metabasalt samples i s only marginal, since the values of o* f o r the metabasalts are i n the neighbourhood of 0.28 while EX2 and EX3 have values of a near 0.25. Only towards the lower depths of layer 2B i s a rock type, the trondhjemites from the Bay of Islands o p h i o l i t e complex, found that has s i m i l a r vp-vg values. This i s rather s u r p r i s i n g since trondhjemites are seldom found i n ocean bottom dredge samples, and i n the o p h i o l i t e s u i t e s , 93 trondhjemites are u s u a l l y found beneath the sheeted d i k e s , at depths w e l l below that of l a y e r 2B. Spudich and Orcutt (1980a) a l s o found low values of Poisson's r a t i o (o"< 0.27) at depths of 0.8 to 1.5 km subsediment which they were not able to e x p l a i n . They suggested that the observed low values of o could be due to the presence of pore f l u i d ; however, they have done t h e o r e t i c a l c a l c u l a t i o n s which i n d i c a t e d that pore f l u i d e f f e c t s are not s u f f i c i e n t to lower Poisson's r a t i o from approximately 0.28 measured i n l a b o r a t o r y samples to a value of 0.24 observed i n l a y e r 2B. The r e s u l t s remain unexplained. Layer 3A In the t r a n s i t i o n from l a y e r 2B to 3A, the Vp-vg paths i n f i g u r e 4.14 move i n t o accordance w i t h the l a b o r a t o r y measurements. Although there i s some overlapping i n the l a b o r a t o r y r e s u l t s f o r the greenschist f a c i e s and epidote-amphibolite f a c i e s m e t a d o l e r i t e s , a t r a n s i t i o n i n the Vp-vg plane between the two groups can be seen i n the v i c i n i t y of vP=6.6 km/s and vg=3.6 km/s. The t r a j e c t o r y of the vp-vg paths f o r both EX2 and EX3 i n t h i s range of v e l o c i t i e s supports the hypothesis that the i n t e r f a c e between l a y e r 2B and 3A marks a metamorphic boundary between the g r e e n s c h i s t f a c i e s and epidote-amphibolite f a c i e s m e t a d o l e r i t e s (Stern et a l . , 1976). V e l o c i t y gradients w i t h i n l a y e r 3A are much more moderate than those at shallower depths. Poisson's r a t i o , however, inc r e a s e s w i t h depth to a value of approximately 0.29 as the v p - v s paths move from the 94 epidote-amphibolite f a c i e s m etadolerite and metagabbro to pyroxene gabbro, then to o l i v i n e gabbro ( f i g u r e 4.14). The good agreement between the v e l o c i t i e s of the l a b samples and the observed seismic r e s u l t s a l l o w s us to i n f e r the composition of l a y e r 3A from the l i t h o l o g i c a l sequence of the Bay of I s l a n d s o p h i o l i t e complexes, where the corresponding l a y e r 3A c o n s i s t s of m e t a d o l e r i t e sheeted dikes u n d e r l a i n by a l a y e r of i n t r u s i v e metagabbro and gabbro, followed by a l a y e r of pyroxene gabbro ( t a b l e 3-1). For t h i s p a r t i c u l a r study s i t e on the Juan de Fuca p l a t e , the correspondence between the i n s i t u oceanic l a y e r 3A and the l i t h o l o g y of the Bay of Islands o p h i o l i t e s i s e x c e l l e n t . Layer 3B and Upper Mantle In l a y e r 3B, v e l o c i t i e s increase from those of the base of l a y e r 3A to those of the upper mantle, at v e l o c i t y g r a d i e n t s that are higher than those found i n l a y e r 3A. There are no sharp d i s c o n t i n u i t i e s i n e i t h e r v or v at the crust-mantle i n t e r f a c e . The i n c r e a s e i n v e l o c i t i e s w i t h depth i n t h i s l a y e r i s accompanied by a decrease i n the apparent Poisson's r a t i o ( f i g u r e 4.13). These are the only common features of the lower c r u s t between p r o f i l e s EX2 and EX3. In the r e s u l t s d i s p l a y e d i n f i g u r e s 4.10 to 4.14, a number of d i s p a r i t i e s between the two p r o f i l e s , at depths near the crust-mantle boundary, can be detected. The most obvious one i s the 4.6 km d i f f e r e n c e i n the i n t e r p r e t e d upper mantle depths. This aspect has been discussed i n d e t a i l i n the l a s t chapter and i t s u f f i c e s here to say that the independent i n t e r p r e t a t i o n of S wave data confirms the l a r g e v a r i a t i o n 95 i n c r u s t a l thicknesses between EX2 and EX3. The other major d i s s i m i l a r i t y between the p r o f i l e s i s the differe n c e i n the values of apparent Poisson's r a t i o at upper mantle depths (see figure 4.12). The term 'apparent' has to be stressed here since, as mentioned e a r l i e r , equation 4.3 used i n the c a l c u l a t i o n of 0* i s appropriate only for i s o t r o p i c media. E f f e c t s of anisotropy f o r upper mantle P waves i s s i g n i f i c a n t l y large (10%) but the interpreted upper mantle shear wave v e l o c i t i e s , 4.5 and 4.6 km/s re s p e c t i v e l y f o r EX2 and EX3, give an anisotropic e f f e c t of only 2%. This combination of high degree of P wave anisotropy and low degree of S wave anisotropy r e s u l t s i n the unusually low apparent C of 0.22 for p r o f i l e EX2, versus the more common value of 0.28 for p r o f i l e EX3. The r e s u l t s given so far are based on the best f i t synthetic seismogram modelling of the data. I t i s reasonable at t h i s point to ask the question of how s i g n i f i c a n t i s the observed difference between the two apparent Poisson's r a t i o s . I f we consider the extremal bounds calculated from the t r a v e l time data, the r e s u l t i n g bounds on o" are c e r t a i n l y broad enough to encompass both values of apparent C. A l l t h i s means i s that t r a v e l time data alone can not resolve the difference i n 6. Modelling the amplitude data, on the other hand, involves a t r i a l -and-error procedure so that a rigorous answer can not be given to the above question. However, a numerical experiment discussed below w i l l shed some l i g h t on the possible uncertainties of the S wave v e l o c i t y models which are expected to dominate the errors i n the c a l c u l a t i o n of o*. 96 Since the Sn head waves have extremely small amplitudes, t h e i r a r r i v a l times do not provide s u f f i c i e n t c o n s t r a i n t s to the upper mantle v s . The p r i n c i p a l f eatures of the data set which a l l o w the upper mantle S wave v e l o c i t y to be determined w i t h adequate r e s o l u t i o n are the r e l a t i v e amplitudes and p o s i t i o n s of the n e a r - c r i t i c a l - a n g l e r e f l e c t i o n s from the crust-upper mantle i n t e r f a c e . F i g u r e 4.16 shows three s y n t h e t i c seismogram s e c t i o n s c a l c u l a t e d by the WKBJ method and the three v e l o c i t y - d e p t h curves used f o r the c a l c u l a t i o n , f o r p r o f i l e EX2. The b e s t - f i t s e c t i o n i s produced with an upper mantle S-wave v e l o c i t y of 4.53 km/s (model b ) . The s y n t h e t i c seismograms give l e s s s a t i s f a c t o r y f i t s (Figure 4.16a, c) when the upper mantle v e l o c i t y deviates by 0.2 km/s from the p r e f e r r e d value, a f a c t demonstrated more c l e a r l y by the p l o t s of amplitude versus distance i n F i g u r e 4.l6e. These observations i n d i c a t e that the probable e r r o r i n the estimate of the upper mantle S wave v e l o c i t y i s i n the order of +0.1 km/s. U n c e r t a i n t i e s i n the upper mantle v p are somewhat l e s s than that of vg, but even assuming e r r o r s of +0.1 km/s f o r both v p and Vg, the d i f f e r e n c e i n cr of 0.06 i s detectable and can not be a t t r i b u t e d wholly to u n c e r t a i n t i e s i n the i n t e r p r e t a t i o n a l techniques. Apparent Poisson's r a t i o of the two p r o f i l e s not only d i f f e r s f o r the upper mantle, but from f i g u r e 4.13, we see that the Vp-vg paths a c t u a l l y s t a r t to diverge w i t h i n l a y e r 3B, i n d i c a t i n g t h a t the e f f e c t s of anisotropy are present, but to a smaller degree, i n the lower crust a l s o . This may have important i m p l i c a t i o n s f o r the p o s s i b l e cause of anisotropy i n the oceanic upper mantle. Further d i s c u s s i o n of t h i s 97 SYNTHETIC DATA Distance (km) 2 4 6 Depth (km) Figure 4.16. (a), (b), and (c) Synthetic seismogram sections modelling the n e a r - c r i t i c a l - r e f l e c t i o n s from the crust-mantle i n t e r f a c e , c a l c u l a t e d from the respective v e l o c i t y models of ( f ) . Model b i s the preferred model, (d) Relevant portion of the data record section of p r o f i l e EX2. (e) Plot of r e l a t i v e amplitude versus distance f o r the three model synthetic sections, compared with the data. The numbers denote the upper mantle S v e l o c i t y of the given v e l o c i t y model. 98 aspect i s given i n the next chapter on anisotropy. In summarizing the seismic r e s u l t s of layer 3B and the upper mantle, i t i s noted here that any postulated p e t r o l o g i c a l model of the oceanic crust for t h i s study s i t e would have to be consistent with the following observations: (1) v e l o c i t i e s i n layer 3B increase with depth at a r e l a t i v e l y pronounced gradient; (2) no sharp v e l o c i t y d i s c o n t i n u i t i e s are present at the crust-upper mantle i n t e r f a c e ; (3) Poisson's r a t i o decreases with depth i n layer 3B; and (4) anisotropy, mainly a f f e c t i n g the P wave v e l o c i t i e s , i s present i n the upper mantle and possibly i n the lower h a l f of layer 3B. Again, the o p h i o l i t e analogy provides a p l a u s i b l e model to account f o r the above observations. Corresponding to l a y e r 3B, the zone of interlayered o l i v i n e gabbro, t r o c t o l i t e , and p e r i d o t i t e found i n the Bay of Islands o p h i o l i t e complex (table 3.1) has the required medium v e l o c i t i e s (vp~7.4 km/s, Vg~3-9 km/s) compared to those observed i n layer 3B. The v e l o c i t y gradient and decreasing o i n t h i s layer can be explained by an increasing o l i v i n e content with depth. A gradual t r a n s i t i o n from gabbroic rocks to ultramafic rocks, s i m i l a r to that observed i n the Bay of Islands o p h i o l i t e would cor r e l a t e well with the crust-upper mantle i n t e r f a c e observed here. Laboratory derived v e l o c i t i e s of the o p h i o l i t e ultramafic samples at 25°C are 8.4 and 4.9 km/s r e s p e c t i v e l y for vp and vg which are s l i g h t l y too high, but higher temperatures (~200°C) at depths corresponding to the crust-upper mantle in t e r f a c e are l i k e l y to reduce the laboratory determined v e l o c i t i e s to those observed f o r the upper mantle. P r e f e r e n t i a l 99 alignment of the o l i v i n e c r y s t a l s , on the other hand, may cause the observed anisotropy i n la y e r 3 B and the upper mantle. This point i s discussed further i n the next chapter. 4.7 Concluding Remarks In terms of s t r u c t u r a l d e f i n i t i o n , the shear wave data i s found to have l e s s r e s o l u t i o n than the P wave data. This i s e s p e c i a l l y true f o r shallow depths where l i t t l e i s known about the shear wave v e l o c i t i e s from the present r e f r a c t i o n techniques. Nonetheless, the i n t e r p r e t a t i o n of the shear wave data i s able to confirm the v a r i a t i o n i n c r u s t a l thicknesses, interpreted previously from the P wave data, between p r o f i l e s EX2 and EX3. This adds support to the in f e r r e d tectonic processes discussed e a r l i e r i n Chapter 3 . More important i s the fac t that the added Vg information has greatly reduced the ambiguities i n the i n t e r p r e t a t i o n of the petrology of the oceanic crust. For example, the v e l o c i t y structures of p r o f i l e s E X 2 and EX3 appear to be quite d i f f e r e n t at f i r s t glance due to the varying thicknesses of various l a y e r s . Their Vp-Vg paths, however, show remarkable s i m i l a r i t i e s over much of the crust, implying that the layers have s i m i l a r composition even though t h e i r thicknesses vary. The obvious d i s p a r i t i e s i n the lower crust and upper mantle are explained by invoking v e l o c i t y anisotropy. I t i s shown that the o p h i o l i t e model of the oceanic crust and upper mantle i s consistent with the r e s u l t s presented here. Variations e x i s t , as exemplified by the low Poisson's r a t i o i n layer 2 B which s t i l l lacks 100 s a t i s f a c t o r y explanation. Spudich and Orcutt (1980a) suggested that the low o i n the upper crust i s a phenomenon related to young oceanic c r u s t . The f i n d i n g of low Poisson's r a t i o i n layer 2B i n t h i s study supports t h e i r conjecture. 101 5. AREAL TRAVEL TIME DATA AND VELOCITY ANISOTROPY 5.1 I n t r o d u c t i o n Compressional wave anisotropy i n the oceanic upper mantle has been w e l l documented by numerous marine r e f r a c t i o n s t u d i e s ( R a i t t et a l . , 1969; Keen and B a r r e t t , 1971; Snydsman et a l . , 1975; and Malecek and Clowes, 1978). In a l l cases, the d i r e c t i o n of maximum P-wave v e l o c i t y i s approximately p a r a l l e l to the d i r e c t i o n of s e a - f l o o r spreading as i n f e r r e d from magnetic anomaly p a t t e r n s . In the present study, a s i m i l a r a n i s o t r o p i c e f f e c t i s i n d i c a t e d by the i n t e r p r e t a t i o n of the p r o f i l e data. The p r i n c i p a l e x p l a n a t i o n f o r t h i s anisotropy i s p r e f e r e n t i a l alignment of the a-axis of o l i v i n e c r y s t a l s i n the d i r e c t i o n of flow during formation at the r i d g e c r e s t (Hess, 1964). That o l i v i n e i s s e i s m i c a l l y a n i s o t r o p i c w i t h the d i r e c t i o n of maximum compressional v e l o c i t y p a r a l l e l to the a-axis i s w e l l e s t a b l i s h e d by l a b o r a t o r y experiments (Verma, 1960; and B i r c h , 1960, 1961). In t h i s chapter, f u r t h e r evidence f o r P wave v e l o c i t y anisotropy i n the upper mantle i s g i v e n . The low degree of shear wave ani s o t r o p y , discussed i n Chapter 4, i s r e l a t e d to two recent l a b o r a t o r y s t u d i e s on seismic anisotropy of o p h i o l i t e samples. I m p l i c a t i o n s f o r s e a f l o o r spreading processes are a l s o discussed. 102 5.2 Areal Data Although the present experiment was not designed for the purpose of a study of v e l o c i t y anisotropy, the shot-receiver configuration (figure 2.1) does provide some useful t r a v e l time data for examining the v e l o c i t y v a r i a t i o n as a function of the azimuthal angle. This i s i l l u s t r a t e d i n figure 5.1 where the possible d i s t r i b u t i o n of shot-receiver distances and azimuths i s shown. I t i s c l e a r from t h i s f i g u r e that for distances over 30 km, which correspond to upper mantle a r r i v a l s , the data cover a r e l a t i v e l y wide range of azimuthal angles. This indicates that i t may be f e a s i b l e to obtain information concerning the azimuthal v a r i a t i o n of v e l o c i t y i n the upper mantle from the ' o f f -p r o f i l e ' t r a v e l time data. There are, however, several problems related to the i n t e r p r e t a t i o n of such t r a v e l time data; the most important of which i s the e f f e c t s of l a t e r a l v a r i a t i o n s i n structure which can also produce azimuthally varying apparent v e l o c i t i e s . To overcome the e f f e c t s of large scale v a r i a t i o n i n c r u s t a l thicknesses, the data are f i r s t p a r t i t i o n e d j u d i c i o u s l y into groups according to the d i f f e r e n t c r u s t a l regimes defined broadly by the p r o f i l e i n t e r p r e t a t i o n s (see Chapter 3). Within each group, the v a r i a t i o n i n the shot-receiver azimuth i s l i m i t e d to l e s s than 25 . A f t e r correcting the t r a v e l times for topography, an apparent v e l o c i t y i s calculated f o r each group and t h i s v e l o c i t y i s assigned to the average azimuth of the group. Of course, only upper mantle a r r i v a l s are selected f o r the c a l c u l a t i o n . E f f e c t s of unknown l o c a l v a r i a t i o n s near the surface at the shot or 103 + + + + + + +++ A + I + + + '+ +++++ + + + + + +++ + + + + + + + ++ + + + ++ ^ £ + + + + + + + + + J ++++#+++ + J £ ^  ^4+H^ ++^ ++*+^  CO o lO o O XJ -+-> N O C O o CO o 08 09 Of 02 ( l l l > [ ) 9 0 U 1 9 ^ S T Q 0 F i g u r e 5.1. Distance versus azimuth f o r a l l p o s s i b l e a r r i v a l s given the s h o t - r e c e i v e r c o n f i g u r a t i o n of f i g u r e 2.1. Azimuths are measured i n degrees from North. ^ 104 r e c e i v e r s i t e s can not be r e a d i l y accounted f o r . I t i s assumed that contaminations by such e f f e c t s are random and can be minimized by c a l c u l a t i n g the apparent v e l o c i t y using a s u f f i c i e n t l y l a r g e number of t r a v e l time p o i n t s . This requirement and the l i m i t placed on the range of azimuths w i t h i n the groups are mutually incompatible so that compromises are o f t e n necessary. Given the above l i m i t a t i o n s , only a small number of apparent v e l o c i t i e s of the upper mantle are obtained from the P wave data and no u s e f u l data i s a v a i l a b l e from the S wave data due to poor a r r i v a l p i c k s . The a z i m u t h a l l y v a r y i n g Pn v e l o c i t i e s are p l o t t e d i n f i g u r e 5.2 where Pn v e l o c i t i e s from the p r o f i l e i n t e r p r e t a t i o n s are a l s o i n c l u d e d . In t h i s p r e s e n t a t i o n , the azimuth i s measured with respect to the d i r e c t i o n of spreading, as defined by the magnetic p a t t e r n based on the Hyndman et a l . (1979) model ( f i g u r e 3.13), at the midpoint of a given s h o t - r e c e i v e r path. The s c a t t e r i n the data i s considerable; however, the existence of a v e l o c i t y maximum near 0° azimuth and a minimum near 90° azimuth i s d e f i n i t e l y i n d i c a t e d . The i n t e r p r e t a t i o n of t h i s data set i s discussed i n the f o l l o w i n g s e c t i o n . 5.3 I n t e r p r e t a t i o n of A r e a l Data Backus (1965) has shown that f o r weakly a n i s o t r o p i c media (< 10% a n i s o t r o p i c e f f e c t ) , the square of the P wave v e l o c i t y as a f u n c t i o n of azimuth i s given by v 2(Q) = A + B C0S(29) + C C0S(49) + D SIN(20) + E SIN(40) (5.1) 105 - LSQ FIT - ORTHOR - TRANSV i 3 0 T 6 0 9 0 Azimuth Figure 5 . 2 . Apparent upper mantle P v e l o c i t i e s as a f u n c t i o n of azimuth, determined from a r e a l l y d i s t r i b u t e d data, are p l o t t e d as s o l i d c i r c l e s . V e l o c i t i e s determined from the p r o f i l e s , w i t h s m a l l e r estimated e r r o r s , are p l o t t e d as s o l i d squares. E r r o r s i n v e l o c i t i e s and azimuths are s u b j e c t i v e estimates based on the r e l i a b i l i t y of i n d i v i d u a l data s e t s . The azimuths are measured i n degrees from the s e a f l o o r spreading d i r e c t i o n s i n f e r r e d from the t e c t o n i c model o f Hyndman et a l . ( 1 9 7 9 ) . The s o l i d l i n e i s a l e a s t - s q u a r e s - f i t of the data t o equation ( 5 . 2 ) . ORTHOR represents the v e l o c i t y v a r i a t i o n of a mixture o f 29% orthorhombic o l i v i n e and an i s o t r o p i c m a t e r i a l w i t h a v e l o c i t y o f 7 . 2 km/s. TRANSV represents the v e l o c i t y v a r i a t i o n o f a mixture o f 3 7 $ t r a n s v e r s e l y i s o t r o p i c o l i v i n e and an i s o t r o p i c m a t e r i a l w i t h a v e l o c i t y of 6.8 km/s. 106 where the c o e f f i c i e n t s A to E are l i n e a r combinations of s i x e l a s t i c constants (see Appendix 2 ) . I f the a n i s o t r o p i c media contains at l e a s t one v e r t i c a l symmetry plane, which i s a reasonable assumption f o r the upper mantle where anisotropy i s caused by the alignment of the o l i v i n e c r y s t a l s , equation (5.1) reduces to where - i s measured from the d i r e c t i o n possessing s a g i t t a l symmetry (Crampin, 1977). For the case where the a-axis of the o l i v i n e c r y s t a l i s a l i g n e d i n the d i r e c t i o n of s e a f l o o r spreading and the b- and c-axes are randomly o r i e n t e d , the d i r e c t i o n of s a g i t t a l symmetry i s the spreading d i r e c t i o n . I f the f u n c t i o n a l form of equation (5.2) i s assumed f o r a given set of v e l o c i t y versus azimuth data, the three c o e f f i c i e n t s can be obtained by a least-squares f i t t i n g procedure. This i s done f o r the Pn v e l o c i t y data shown i n f i g u r e 5.2. The s o l i d curve (LSQ) represents the l e a s t -s q u a r e s - f i t s o l u t i o n and i s given by The c o e f f i c i e n t s A, B, and C thus obtained do not a l l o w immediate i d e n t i f i c a t i o n of the composition of the media s i n c e the r e s u l t i n g e l a s t i c constants are only the apparent values of the parameters. Crampin and Bamford (1977) suggested a p o s s i b l e i n t e r p r e t a t i o n a l technique i n which the upper mantle i s assumed to be a mixture of v 2(Q) = A + B C0S(29) + C C0S(H9) (5.2) v 2 ( 0 ) = 62.28 + 5.53 C0S(29) - 2.77 C0S(49) (5.3) 107 i s o t r o p i c m a t e r i a l and a l i k e l y a n i s o t r o p i c m a t e r i a l such as o l i v i n e . The r a t i o of these two components and the parameters of the i s o t r o p i c m a t e r i a l can then be determined from the c o e f f i c i e n t s of equation (5.2) using the e l a s t i c constants of the assumed a n i s o t r o p i c m a t e r i a l (see Appendix 2). Such informa t i o n can be valu a b l e f o r the i n t e r p r e t a t i o n of the petrology of the upper mantle. Using the r e s u l t s of the l e a s t - s q u a r e s - f i t equation (5.3), the v e l o c i t y v a r i a t i o n s of two mixtures are c a l c u l a t e d and are shown i n f i g u r e 5.2. I f the a n i s o t r o p i c m a t e r i a l i s assumed to be orthorhombic o l i v i n e , i t would c o n s t i t u t e 29% of the mixture and the i s o t r o p i c p o r t i o n would have a v e l o c i t y of 7.2 km/s. I f the a n i s o t r o p i c m a t e r i a l i s assumed to be t r a n s v e r s e l y i s o t r o p i c o l i v i n e , a more l i k e l y candidate f o r the c o n s t i t u e n t of the upper mantle than orthorhombic o l i v i n e , i t would make up 37% of the mixture, g i v i n g a v e l o c i t y of 6.8 km/s f o r the i s o t r o p i c p o r t i o n . The d i f f e r e n c e s i n the v e l o c i t y v a r i a t i o n s of the two mixtures are s l i g h t ( f i g u r e 5.2) which i m p l i e s that w i t h t h i s type of data, the two d i f f e r e n t o l i v i n e f a b r i c s are probably i n d i s t i n g u i s h a b l e . However, i t i s evident from f i g u r e 5.2 that the v e l o c i t y v a r i a t i o n s of both mixtures d i f f e r s i g n i f i c a n t l y from that e x h i b i t e d by the data. There are a number of p o s s i b l e causes f o r the m i s f i t and they are discussed i n the next s e c t i o n . For the moment, one p a r t i c u l a r aspect of the m i s f i t i s addressed. Given a two-component mixture of the type described above, equation (5.2) p r e d i c t s that the maximum of the v e l o c i t y should occur at 0° azimuth w h i l e the minimum should occur at 90° sin c e the c o e f f i c i e n t of 108 the COS(49) term i s p o s i t i v e f o r the two o l i v i n e s . The mixture model curves derived from the c o e f f i c i e n t s of the l e a s t - s q u a r e s - f i t equation should, and do, c o i n c i d e w i t h the LSQ curve at these two values of azimuth. However, the maximum of the LSQ curve i s not l o c a t e d at 0° but rath e r at approximately 30°. This suggests that a more reasonable approach i n modelling the data i s to abandon the l e a s t - s q u a r e s - f i t curve and i n s t e a d , take the well-determined maximum and minimum v e l o c i t i e s (8.3 and 7.5 km/s r e s p e c t i v e l y ) from the p r o f i l e i n t e r p r e t a t i o n s to be the end points of the model curves. The r e s u l t s of the second modelling approach are shown i n f i g u r e 5.3« The v e l o c i t y v a r i a t i o n s of the two model mixtures are now more i n l i n e w i t h the trend of the data. The orthorhombic o l i v i n e mixture c o n s i s t s of 33% a n i s o t r o p i c m a t e r i a l and the i s o t r o p i c m a t e r i a l has a v e l o c i t y of 7.4 km/s. For the t r a n s v e r s e l y i s o t r o p i c o l i v i n e mixture, the percentage of a n i s o t r o p i c m a t e r i a l i s 42% while the v e l o c i t y of the i s o t r o p i c m a t e r i a l i s 7.0 km/s. 5.4 D i s c u s s i o n The r e s u l t s obtained from the i n t e r p r e t a t i o n of the P wave v e l o c i t y anisotropy data, using the r e l a t i v e l y simple techniques described i n the l a s t s e c t i o n , can only be regarded as a rough i n d i c a t i o n of the p o s s i b l e composition l i k e l y to be found i n the upper mantle. I t i s important to s t r e s s the f a c t that seismic waves a c t u a l l y sample the media over s e v e r a l wavelengths so that the value of the v e l o c i t y i s only an average over a distance of s e v e r a l k i l o m e t e r s ; whereas, the proposed mixtures 109 02 LSQ FIT ORTHOR TRANSV 30 60 Azimuth 90 Figure 5.3. Description is the same as figure 5.2 except ORTHOR now represents the velocity variation of a mixture of 33$ orthorhombic olivine and an isotropic material with a velocity of 7.4 km/s, and TRANSV represents the velocity variation of a mixture of 42$ transversely isotropic olivine and an isotropic material with a velocity of 7.0 km/s. 110 are modelled a f t e r rock c r y s t a l s much smaller i n dimension and are assumed to have i d e a l i z e d c o n f i g u r a t i o n s . The r e a l earth i s probably much more complex and equation (5.2) i s only an approximation. Complication a l s o a r i s e s from the inherent u n c e r t a i n t i e s i n the data. For example, the determination of the azimuthal angle i s f a r from p r e c i s e due to the l a r g e u n c e r t a i n t i e s i n the i n f e r r e d d i r e c t i o n of s e a f l o o r spreading and the f a c t that the v e l o c i t y c a l c u l a t i o n s are done over a f i n i t e range o f azimuths. Even i f the o r i e n t a t i o n s of the ray paths with respect to the spreading d i r e c t i o n are known e x a c t l y , the alignment of the o l i v i n e c r y s t a l s i n the upper mantle i s not n e c e s s a r i l y p e r f e c t . Given these c o n s i d e r a t i o n s and the s p a r s i t y of data, more s o p h i s t i c a t e d modelling than what has been done i s probably not warranted. Before d i s c u s s i n g i n more d e t a i l the anisotropy r e s u l t s of t h i s and the previous two chapters, a b r i e f summary of the observations i s given i n the f o l l o w i n g . For the upper mantle, P wave v e l o c i t y anisotropy i s pronounced with v e l o c i t i e s of approximately 8.3 and 7-5 km/s p a r a l l e l and perpendicular, r e s p e c t i v e l y , to the d i r e c t i o n of s e a f l o o r spreading. The S wave v e l o c i t i e s , 4.6 and 4.5 km/s, determined from two shear wave p r o f i l e s are n e a r l y i s o t r o p i c . This r e s u l t s i n a prominent apparent anisotropy o f Poisson's r a t i o f o r the upper mantle, w i t h values ranging from 0.22 to 0.28, the high value being i n the d i r e c t i o n of p l a t e motion. The upper mantle P wave anisotropy can be crudely modelled by a mixture of k2% t r a n s v e r s e l y i s o t r o p i c o l i v i n e and an i s o t r o p i c m a t e r i a l w i t h a v e l o c i t y of 7.0 km/s. There are h i n t s that the lower h a l f of 111 l a y e r 3B a l s o possesses v e l o c i t y anisotropy of the type e x h i b i t e d by the upper mantle, but the magnitude i s s m a l l e r . I f the hypothesis that o p h i o l i t e s represent obducted fragments of oceanic l i t h o s p h e r e i s to be accepted, then evidence from l a b o r a t o r y s t u d i e s of o p h i o l i t e samples must be c o n s i s t e n t w i t h the observations of anisotropy i n the oceanic c r u s t and upper mantle reported here and elsewhere. Christensen and S a l i s b u r y (1979) stud i e d the p e t r o f a b r i c s of f i f t e e n u l t r a m a f i c samples c o n s i s t i n g e n t i r e l y of t e c t o n i t e s from three widely spaced tr a v e r s e s i n the Bay of I s l a n d s , Newfoundland o p h i o l i t e complex. T h e i r comprehensive a n a l y s i s y i e l d e d a compressional wave anisotropy of 6% (percentage defined by (1 - Vmin/Vmax) X 100) w i t h the maximum v e l o c i t y of 8.7 km/s i n the plane of the Mohorovicic d i s c o n t i n u i t y and p a r a l l e l to the i n f e r r e d spreading d i r e c t i o n , and the minimum v e l o c i t y of 8.2 km/s perpendicular to t h i s d i r e c t i o n ( t a b l e 3.1). For shear waves, f o r which the propagation v e l o c i t y depends on v i b r a t i o n d i r e c t i o n r e s u l t i n g i n two d i f f e r e n t S-wave v e l o c i t i e s , the r e s u l t s y i e l d e d n e a r l y i s o t r o p i c S wave v e l o c i t i e s f o r both p o l a r i z a t i o n s of the shear wave (average vg = 4.9 km/s). These v e l o c i t y values were c a l c u l a t e d f o r a temperature of 25°C and f o r appropriate upper mantle pressures. As pointed out by Christensen and S a l i s b u r y (1979), i n c r e a s e s i n temperature of 200°C w i l l lower the v e l o c i t i e s i n t o good agreement with observed oceanic upper mantle v e l o c i t y w i t h l i t t l e e f f e c t on anisotropy. They found Poisson's r a t i o ranging from 0.23 to 0.27 w i t h the high values g e n e r a l l y i n the d i r e c t i o n of spreading. The agreement between the r e s u l t s of the present study and those of Christensen and S a l i s b u r y (1979) i s e x c e l l e n t . 112 In another study, an ha r z b u r g i t e sample from the Ant a l y a o p h i o l i t e complex i n Turkey was analysed i n d e t a i l ( P e s e l n i c k and N i c o l a s , 1978). C a l c u l a t i o n s of P wave v e l o c i t i e s f o r the uppermost mantle at 250°C and 5 kbar, based on the e l a s t i c constants obtained from the sample, gave values of 8.46 km/s i n a d i r e c t i o n normal to the i n f e r r e d r i d g e c r e s t and 8.16 km/s p a r a l l e l to i t . The 3-5% anisotropy of 0.3 km/s agrees with some i n - s i t u experiments ( R a i t t et a l . , 1969), but i s l e s s than the anisotropy u s u a l l y determined i n marine r e f r a c t i o n s t u d i e s . For the two modes of shear wave propagaton, a n i s o t r o p i c e f f e c t s of 2.3% and 1.6% were found. Although there are d i f f e r e n c e s between these two o p h i o l i t e s t u d i e s , the general c h a r a c t e r i s t i c s of a n i s o t r o p i c e f f e c t s , such as the d i r e c t i o n s of maximum and minimum v e l o c i t y w i t h respect to spreading d i r e c t i o n and the higher a n i s o t r o p i c e f f e c t f o r P waves compared w i t h S waves, are shared by the two sets of r e s u l t s . These c h a r a c t e r i s t i c s are a l s o c o n s i s t e n t w i t h the r e s u l t s determined i n the present study. This f u r n i s h e s yet another p o s i t i v e argument f o r the o p h i o l i t e hypothesis. A p u z z l i n g aspect of the present r e s u l t s i s the anisotropy observed i n the lower h a l f of l a y e r 3B where the primary composition i s thought to be cumulate o l i v i n e gabbro (Christensen and S a l i s b u r y , 1975) which g e n e r a l l y i s s e i s m i c a l l y i s o t r o p i c . Christensen (1972) suggested t h a t d i p p i n g of amphibolite f o l i a t i o n planes i n the lower c r u s t may cause some of the observed v e l o c i t y v a r i a t i o n s i n t e r p r e t e d by Keen and B a r r e t t (1971) f o r that part of the c r u s t . However, i t i s not c l e a r what 113 mechanism would cause dipping of the f o l i a t i o n plane of the amphibolite. Another possible explanation for the anisotropy observed i n layer 3B i s given, again, by the o p h i o l i t e samples. Christensen and Salisbury (1979) found that for the gabbros immediately overlying the ultramafics of the Bay of Islands o p h i o l i t e s , the o l i v i n e f a b r i c i s i d e n t i c a l with that i n the ultramafic t e c t o n i t e s , i n d i c a t i n g that t e c t o n i z a t i o n has extended beyond the ultramafics i n t o the overlying gabbros. This r e o r i e n t a t i o n of o l i v i n e o r i g i n a l l y formed by cumulous processes may be caused by t r a n s l a t i o n g l i d i n g presumably induced by the same stress f i e l d that i s responsible for the anisotropy i n the upper mantle. The fa c t that the or i e n t a t i o n of o l i v i n e i s the same i n the lower crust as i t i s i n the upper mantle would suggest the possible existence of anisotropy i n the lower crust of the same type as observed f o r the upper mantle but the lower o l i v i n e content of the lower crust would imply a lower degree of anisotropic e f f e c t . To further substantiate t h i s assertion would require a more extensive data set than that a v a i l a b l e i n the present study. In conclusion, o p h i o l i t e s provide a number of explanations for the observed anisotropy i n the oceanic crust and upper mantle, as discussed above. 114 6. SUMMARY The f i r s t objective of t h i s study i s to investigate the c r u s t a l structure i n the region of the Nootka f a u l t zone through the use of seismic r e f r a c t i o n methods. This has been accomplished and the in t e r p r e t a t i o n s of the data show remarkably deta i l e d changes i n c r u s t a l properties with depth and s i g n i f i c a n t l a t e r a l v a r i a t i o n s i n stru c t u r e . The r e s u l t s i n t h i s study indicate that layers 2A and 2B i n the upper crust are characterized by rapid increases i n seismic v e l o c i t i e s . For example, P wave v e l o c i t y increases from a range of 3-7 to 4.5 km/s at the top of layer 2A, to values of 6.0 to 6.4 km/s at the base of layer 2B. S i m i l a r l y , shear wave v e l o c i t y increases from 2.2 to 3-6 km/s over the same depth i n t e r v a l of approximately 3«5 km; although the exact nature of the shear wave v e l o c i t y gradient i s uncertain. In contrast, v e l o c i t i e s i n layer 3A are r e l a t i v e l y uniform, with v p l y i n g i n the range of 6.4 to 6.8 km/s and v s remaining f a i r l y constant with depth at values near 3-6 km/s. No v e l o c i t y d i s c o n t i n u i t y i s observed at the crust-upper mantle i n t e r f a c e . Instead, the t r a n s i t i o n from crust to upper mantle i s marked by the presence of a v e l o c i t y gradient zone, layer 3B, i n which the v e l o c i t i e s increase gradually from those at the base of layer 3A to those of the upper mantle. P wave v e l o c i t y i n the upper mantle varies s i g n i f i c a n t l y with azimuth from 8.3 km/s i n a d i r e c t i o n p a r a l l e l to the i n f e r r e d d i r e c t i o n of seafloor spreading to 7.5 km/s i n a d i r e c t i o n p a r a l l e l to the spreading ridge. However, corresponding var i a t i o n s i n the observed upper mantle shear wave v e l o c i t i e s , 4.6 and 4.5 km/s, are small and l i e within the 115 uncertainties of +0.1 km/s i n the Sn v e l o c i t y determinations. The i n t e r p r e t a t i o n of shear wave data presented i n t h i s study provides the f i r s t i n s i t u evidence f o r a low degree of shear wave v e l o c i t y anisotropy i n the upper mantle, a phenomenon which has been suggested by laboratory studies of o p h i o l i t e ultramafic samples. On the other hand, upper mantle P wave v e l o c i t y exhibits a 10% anisotropic e f f e c t , r e s u l t i n g i n a prominent anisotropy i n the values of Poisson's r a t i o for the upper mantle. These aspects of v e l o c i t y anisotropy can be v e r i f i e d by experiments s u i t a b l y designed f o r the study of anisotropy i n the crust and upper mantle. D e t a i l s of the c r u s t a l structure outlined i n t h i s study should provide relevant guidelines f o r designing such an experiment for th i s region. The most notable feature of the c r u s t a l structure i n t h i s region i s the sudden increase i n c r u s t a l thickness from 6.4 to 11.2 km over a l a t e r a l distance of approximately 30 km. A rapid process of 'cr u s t a l maturing' within a 1 Myr time i n t e r v a l , v a r i a t i o n s with time i n the process of c r u s t a l formation at the ridge, or both could account f o r t h i s s i g n i f i c a n t v a r i a t i o n i n structure. Such an abrupt change i n c r u s t a l thickness should show c l e a r l y on a gra v i t y anomaly map, but unfortunately, the e x i s t i n g g r a v i t y data f o r the west coast of Canada do not cover t h i s region. Thus an independent check on t h i s aspect of the interpreted r e s u l t s could be provided by a gra v i t y survey over the area. The second objective of t h i s study i s to re l a t e the seismic r e s u l t s to the petrology of the oceanic crust and upper mantle. In t h i s respect, ^ c o n v e n i e n t unifying theme i s supplied by the o p h i o l i t e hypothesis f o r 116 the oceanic lithosphere. Correlations between the seismic r e s u l t s and the properties of o p h i o l i t e complexes, i n p a r t i c u l a r the Bay of Islands o p h i o l i t e s , c o n s i s t e n t l y surface throughout the study. For example, the observed v e l o c i t i e s and gentle v e l o c i t y gradients i n layer 3A agree very well with laboratory measurements of seismic v e l o c i t i e s for the sequence of metadolerite sheeted dikes, metagabbros, and pyroxene gabbros from the Bay of Islands o p h i o l i t e s . Below the pyroxene gabbros, the l i t h o l o g y of t h i s o p h i o l i t e complex shows increasing content of pyroxene and o l i v i n e as t r o c t o l i t e s and o l i v i n e gabbros become more abundant with depth. This has i t s correspondence i n the more pronounced v e l o c i t y gradient i n layer 3B compared with layer 3A as interpreted i n t h i s study. Yet another c o r r e l a t i o n between the properties of the o p h i o l i t e s and the seismic c h a r a c t e r i s t i c s of the oceanic upper mantle i s provided by the observation of v e l o c i t y anisotropy. Petrofabric studies of ultramafic samples from o p h i o l i t e s indicate that the P wave v e l o c i t y i n the upper mantle would exhibit azimuthal v a r i a t i o n s , with maximum and minimum v e l o c i t i e s i n d i r e c t i o n s p a r a l l e l and perpendicular to the d i r e c t i o n s of plate motion, r e s p e c t i v e l y . The p e t r o f a b r i c studies also show that the corresponding S wave v e l o c i t y i s only m i l d l y a n i s o t r o p i c . These r e s u l t s are consistent with those found i n t h i s study. The correspondence between data measured from o p h i o l i t e s and the observations i n t h i s study provide support for the supposition that o p h i o l i t e s are segments of oceanic crust and upper mantle emplaced on land. By comparing the seismic r e s u l t s with the l i t h o l o g y of o p h i o l i t e s , the enormous non-uniqueness i n the p e t r o l o g i c a l i n t e r p r e t a t i o n of seismic v e l o c i t i e s i s thus reduced, so that inferences on the petrology 117 of the oceanic crust and upper mantle are possible. More de t a i l e d determination of the l i t h o l o g i c a l sequences of the i n - s i t u oceanic crust w i l l require high q u a l i t y P and S wave data recorded on three-component instruments with good amplitude c o n t r o l . 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R., 1976. An ocean-bottom seismometer s u i t a b l e for arrays, Deep Sea Res., 13, 113-124. 123 Malecek, S. J . & Clowes, R. M., 1978. C r u s t a l s t r u c t u r e near E x p l o r e r Ridge from a marine deep seismic sounding survey, J . Geophys. Res., 83, 5899-5912. McMechan, G. A. & Mooney, W. D., 1980. Asymptotic ray theory and s y n t h e t i c seismograms f o r l a t e r a l l y v a r y i n g s t r u c t u r e : theory and a p p l i c a t i o n to the Im p e r i a l V a l l e y , C a l i f o r n i a , B u l l . Seism. Soc. Am., 70, 2021-2035. O'Brien, P. N. S., 1960. Seismic energy from e x p l o s i o n s , Geophys. J . R. a s t r . S o c , 3, 29-44. Orcutt, J . A., 1980. J o i n t l i n e a r , extremal i n v e r s i o n of seis m i c kinematic data, J . Geophys. Res., 85, 2649-2660. P e s e l n i c k , L. & N i c o l a s , A., 1978. Seismic anisotropy i n an o p h i o l i t e p e r i d o t i t e : a p p l i c a t i o n to oceanic upper mantle, J . Geophys. Res., 83, 1227-1235. R a f f , A. D. & Mason, R. G., 1961. Magnetic survey o f f the west coast of North America, 40°N to 52°N l a t i t u d e , Geol. Soc. Am. B u l l . , 72, 1267-1270. R a i t t , R. W., 1963. The c r u s t a l rocks, i n The Sea, v o l . 3, M. N. H i l l (ed.), I n t e r s c i e n c e , New York, 85-102. 124 R a i t t , R. W., Shor, G. G. J r . , Francis, T. J . G., & Morris, G. B., 1969-Anisotropy of the P a c i f i c upper mantle, J . Geophys. Res., 74, 3095-3109. Richards, P. G., 1979. Theoretical seismic wave propagation, Rev. Geophys. Space Phys., 17, 312-328. Riddihough, R. P., 1977. A model for recent plate i n t e r a c t i o n s o f f Canada's west coast, Can. J . Earth S c i . , 14, 384-396. Salisbury, M. H., 1974. Investigations of seismic v e l o c i t i e s i n the Bay of Islands, Newfoundland o p h i o l i t e complex f o r comparison with oceanic seismic structure, Ph. D. t h e s i s , Univ. of Washington, Se a t t l e , 197 pp. Salisbury, M. H. & Christensen, N. I., 1978. The seismic v e l o c i t y structure of a traverse through the Bay of Islands o p h i o l i t e complex, Newfoundland, an exposure of oceanic crust and upper mantle, J . Geophys. Res., 83, 805-817. Snydsman, W. E., Lewis, B. T. R., & McClain, J . , 1975. Upper mantle v e l o c i t i e s on the northern Cocos plate, Earth Planet. S c i . L e t t . , 28, 46-50. 125 Spudich, P. K. & Helmberger, D. V., 1979. S y n t h e t i c seismograms from model ocean bottoms, J . Geophys. Res., 189-204. Spudich, P. K. & O r c u t t , J . , 1980a. Petrology and p o r o s i t y of an oceanic c r u s t a l s i t e : r e s u l t s from wave form modelling of seismic r e f r a c t i o n data, J . Geophys. Res., 85, 1409-1433. Spudich, P. K. & Orcutt, J . , 1980b. A new look at the seismic v e l o c i t y s t r u c t u r e of the oceanic c r u s t , Rev. Geophys. Space Phys., 18, 627-645. Steinmetz, L., Whitmarsh, R. B., & Mo r e i r a , V. S., 1977. Upper mantle s t r u c t u r e of the M i d - A t l a n t i c r i d g e north of the Azores based on observation of compressional waves, Geophys. J . R. a s t r . S o c , 50, 353-380. St e r n , C., de Wit, M. J . , & Lawrence, J . R., 1976. Igneous metamorphic processes a s s o c i a t e d with the formation of Chilean o p h i o l i t e s and t h e i r i m p l i c a t i o n f o r ocean f l o o r metamorphism, seismic l a y e r i n g and magnetism, J . Geophys. Res., 81, 4370-4380. Sutton, G. H., Maynard, G. L., & Hussong, D. M., 1971. Widespread occurrence of a high v e l o c i t y basal l a y e r i n the P a c i f i c c r u s t found with r e p e t i t i v e sources and sonobuoys, The S t r u c t u r e and P h y s i c a l P r o p e r t i e s of the Earth's Crust, Geophys. Monogr. Ser., v o l . 14, J . G. Heacock (ed.), AGU, Washington, D. C., 193-209. 126 T i f f i n , D., L. & Seeman, D., 1975. Bathymetry map of the continental margin of western Canada, open f i l e map, Geol. Surv. Can., Vancouver, B. C. Verma, R. K., 1960. E l a s t i c i t y of some high density c r y s t a l s , J . Geophys. Res., 65, 757-766. Vine, F. J . & Wilson, J . T., 1965. Magnetic anomalies over a young oceanic ridge o f f Vancouver Island, Science, 150, 485-489. White, R. S. & Stephen, R. A., 1980. Compressional to shear wave conversion i n oceanic crust, Geophys. J . R. a s t r . S o c , 63, 547-565. Whitmarsh, R. B., 1975. A x i a l i n t r u s i o n zone beneath the. median v a l l e y of the Mid-Atlantic Ridge at 37°N detected by explosion seismology, Geophys. J . R. a s t r . S o c , 42, 189-215. Wilson, J . T., 1965. Transform f a u l t s , ocean ridges and magnetic anomalies southwest of Vancouver Island, Science, 150, 482-485. 127 APPENDIX 1: Linea r i z e d Inversion of Tau-P Data Assuming that the v e l o c i t y i s a monotonically increasing function of depth, the i n t e g r a l expression f o r t i s zCp) •t(p) = 2 $ ( u 2 - p 2 ) 1 / 2 dz (A.1.1) where u(z) = 1/v(z) I t was shown by Garmany (1979) that by changing the dependent variable to Z=Z(u) and defining the model to be m(u)=dZ/du, a l i n e a r r e l a t i o n s h i p e x i s t s between x(p) and m(u). Equation (A.1) becomes -c(p) = J m(u)-h(u,p) du (A.1.2) o where m(u) = dZ/du umax = 1/v(z=0) r - 2 ( u 2 - p 2 ) 1 / 2 u > p and h(u,p) = ^ 0 u „ p I f we p a r t i t i o n (0, u- a x) into M i n t e r v a l s and assume that m^=dZ/du i s constant within each i n t e r v a l (u^_-|, u^), we can rewrite (A.1.2) as (p)= 1 a ^ (A.1.3) 128 where j h(u,p) du wi t h 0 Given N d i s c r e t e data points ("Cj, p j ) , j = 1,N , w i t h e r r o r A T J , then (A.1.3) becomes the 2N i n e q u a l i t i e s : from which the m^'s can be e a s i l y found using l i n e a r programming techniques. The model m(u) i s then i n t e g r a t e d to give Z(u) which i n t u r n i s transformed to V(z) (LINP curve i n f i g u r e 3.11d). The approach taken i n t h i s study i s to f i n d the m^'s by minimizing the depth to the highest v e l o c i t y subjected to the c o n s t r a i n t s of (A. 1.4). This i s s l i g h t l y d i f f e r e n t from the approach of Garmany et a l . (1979) who obtained the bounds on the V(z) f u n c t i o n by a l t e r n a t e l y maximizing and minimizing the depth to each v e l o c i t y i n t e r v a l i n s t e a d of f i n d i n g a p a r t i c u l a r set of mi's which s a t i s f i e s the c o n s t r a i n t s (A.1.4). t j - ATJ £ 2. a i m i J = 1,N (A.1.4) "Cj + A*J > £ Hmi 129 APPENDIX 2: Inversion of P Wave Anisotropy Backus (1965) has shown that for weakly ani s o t r o p i c media (< 10% anisotropic e f f e c t ) , the square of the P wave v e l o c i t y as a function of azimuth i s given by ov 2(9) = { 3 C n + (C-n + ( C 2 1 + { C n + ( C 2 1 1 + 3C2222 + 2 ^1122 + 2^1212^ / '^ 1 + c 2 2 2 2 ) / 2 COS(29) 1 + C 1 2 2 2 ) / 2 SIN(29) 1 + c2222 - 2( c1122 + 2 C 1212^/8 C0S(49) 1 " C 1 2 2 2 ) / 2 SIN(40) (A.2.1) where 9 i s measured from the x-| a x i s . I f the media contains at least one v e r t i c a l plane of symmetry, Crampin (1977) has shown that equation (A.2.1) can be reduced to ov 2(9) = { 3 C HH + 3C 2 222 + 2( C1122 + 2 C1212^1' / 8 + (C-jm + C 2 2 2 2 ) / 2 C0S(29) + { C i r n + C 2 2 2 2 - 2 ( C 1 1 2 2 + 2 C 1 2 1 2 ) } / 8 C0S(49) (A.2.2) where 9 i s measured from the d i r e c t i o n possessing s a g i t t a l symmetry. Rewriting equation (A.2.2) as v 2 ( 9 ) = A + B C0S(29) + C C0S(49) (A.2.3) then from equations (A.2.2) and (A.2.3) we have 130 v-|2(9=0) = A + B + C = C1111/9 (A.2.4) v 2 2 ( 0 = /2) = A - B + C = C2222/9 (A.2.5) Given a set of v e l o c i t y versus azimuth data, a least-squares f i t t i n g of the data to the functional form of equation (A.2.3) would y i e l d the c o e f f i c i e n t s A, B, and C. Following Crampin and Bamford (1977), i f we assume that the observed v e l o c i t y v a r i a t i o n i s due to a mixture of anisotropic material with the following v e l o c i t y behaviour v A2(Q) = A 1 + B' C0S(29) + C C0S(49) (A.2.6) and an i s o t r o p i c material with unknown constant v e l o c i t y v T i n the r a t i o of x/(1-x) ( a n i s o t r o p i c / i s o t r o p i c ) , the v e l o c i t y of the mixture i s then given by v 2 ( 9 ) = x v A 2 ( 9 ) + (1-x) V T 2 (A.2.7) Equating t h i s to the l e a s t - s q u a r e s - f i t equation (A.2.3) and solving f o r x and v j 2 we have x = 2B/2B1 (A.2.8) and v T 2 = A + B + C - x (A' + B' + C ) / (1-x) (A.2.9) where A', B', and C* are the appropriate c o e f f i c i e n t s for the anisotropic material and are assumed to be known from laboratory 131 measurements and A, B, and C are the c o e f f i c i e n t s from the l e a s t -s q u a r e s - f i t equation. In terms of v-| and v 2 we have x = (v-,2 _v 2 2)/2B' (A.2.10) and vj2 = f v-| 2 - x (A' + B' + C')} / (1-x) (A.2.11) For convenience, the density has been assumed to be the same i n a l l cases f o r the above derivation, although t h i s i s not necessary. Using the r e s u l t s of Verma (1960), Crampin and Bamford (1977) calculated A'=75.256, B»=18.953, and C'=3.262 f o r orthorhombic o l i v i n e and A'=79.012, B' = 15.117, and C'=3.31*2 for transversely i s o t r o p i c o l i v i n e (with v e l o c i t i e s measured i n units of km/s). These are the values used i n Chapter 5 for the c a l c u l a t i o n of x and v j . 

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