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Shallow crustal structure beneath the Juan de Fuca ridge from 2 seismic refraction tomography White, Donald John 1983

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SHALLOW CRUSTAL STRUCTURE BENEATH THE STRAIT OF GEORGIA, BRITISH COLUMBIA by DONALD JOHN WHITE B.Sc. (Physics), University of Toronto, 1981 A THESIS SUBMITTED IN PARTIAL FULFILLMENT OF THE REQUIREMENTS FOR THE DEGREE OF MASTER OF SCIENCE in THE FACULTY OF GRADUATE STUDIES Department of Geophysics and Astronomy We accept this thesis as conforming to the required standard THE UNIVERSITY OF BRITISH COLUMBIA September 1983 © Donald John White, 1983 In presenting t h i s thesis i n p a r t i a l f u l f i l m e n t of the requirements for an advanced degree at the University of B r i t i s h Columbia, I agree that the Library s h a l l make i t f r e e l y available for reference and study. I further agree that permission for extensive copying of t h i s thesis for scholarly purposes may be granted by the head of my department or by his or her representatives. I t i s understood that copying or publication of t h i s thesis for f i n a n c i a l gain s h a l l not be allowed without my written permission. Department of £grpUx|S( CJ) 0—d The University of B r i t i s h Columbia 1956 Main Mall Vancouver, Canada V6T 1Y3 Date 83 DE-6 (.3/81) ABSTRACT In May, 1982 a seismic refraction survey, using a 32 1 airgun and a radio telemetering sonobuoy system with direct digital recording, was carried out in Georgia Strait. The objectives of this experiment were determination of the upper crustal structure beneath the Strait and investigation of the proposed existence of a major fault separating Vancouver Island from the continental mainland. Three reversed profiles across the Strait and reversed lines along i t on either side were shot. In addition, the southernmost line from northern Galiano Island to Point Grey was recorded digitally from a telemetered land-based station on Galiano Island. The refraction data are supplemented by several high resolution reflection profiles from previous experiments. Two-dimensional models of the crustal structure across the Strait have been constructed using a forward modelling ray trace and synthetic seismogram algorithm to match the travel times and amplitude characteristics of the data. Generally, these models consist .of 3 layers. The f i r s t consists of unconsolidated sediments and Pleistocene glacial deposits varying in thickness up to 1 km. The velocity of this layer is assumed to be 1.6 km/s due to lack of velocity information. The second layer represents Upper Cretaceous sediments of the Nanaimo Group and possibly Chuckanut Formation, having velocities ranging from 3.6 km/s to 4.2 km/s. It varies in thickness up to 2 km, thinning toward the mainland side of the Strait. The third layer consists of a thin transitional zone, with velocity 6.0-6.1 km/s at the surface and gradient 0.5 km/s/km, beneath which the velocity gradient is decreased to 0.1-0.15 km/s/km. This layer dips at angles of 2°-16° toward Vancouver Island. This layer, is likely the extension of the Coast. Range intrusives. The Malaspina fault proposed for the Strait of Georgia is not observed, although some uncertainty remains due to termination of seismic coverage short of the mainland. A local fault with calculated throw of 0.55 km has been located approximately 15 km northeast of Galiano Island. The dip and strike of the fault are poorly constrained. Maximum depth of penetration obtained in this study is about 3 km. iv Table of Contents Abstract i i List of tables vi List of figures v i i Acknowledgements ix CHAPTER 1 Introduction and Background 1" 1.1 Geologic Background 3 1.2 Tectonic History 7 1.3 Previous Seismic Studies 11 1.4 Experimental Objectives 13 CHAPTER 2 Data Acquisition and Reduction .......... 15 2.1 The Refraction Experiment 15 2.2 Instrumentation 17 2.3 Data Processing and Corrections 21 CHAPTER 3 The Data: Observations and Modelling Approach . 27 3.1 In i t i a l Observations on the Data Set 27 3.2 Modelling Method 44 CHAPTER 4 Detailed Interpretation ... 52 4.1 The Northern Lines: SB6/SB6R/R6 52 4.2 The Central Lines: SB3/SB3R/R3 63 4.3 The Southern Lines: SB5/SB5R/R5 73 4.4 SB4/SB4R '. 82 4.5 Final Velocity Models 88 CHAPTER 5 Summary and Discussion 92 5.1 Future Studies 99 Bibliography 102 Appendix A Asymptotic Ray Theory vi List of Tables 1.1 Tectonic units of Vancouver Island 6 2.1 Position uncertainties 23 2.2 Timing uncertainties ; 25 3.1 Sonobuoy dri f t calculations 50 v i i List of Figures 1.1 Tectonics of southern British Columbia 2 1.2 Geology of study area 4 1.3 Velocity model of Tseng (1968) 13 2.1 Map of profile locations ... 16 2.2 Refraction recording system flow diagram 18 2.3 Source pulse 20 2.4 Seismic power spectrum 26 3.1 SB6/6R refraction data 28 3.2 R3 and R6 CSP reflection data 29 3.3 SB3/SB3R refraction data 32 3.4 SB5/SB5R refraction data 34 3.5 GAL5/GAL5R refraction data 37 3.6 R5 CDP profile 39 3.7 SB1 refraction data 41 3.8 SB4/SB4R refraction data 43 4.1 SB6: Model A and travel time plot 54 4.2 SB6R: Model A and travel time plot 55 4.3 SB6: Model B and travel time plot 58 4.4 SB6R: Model B and travel time plot 59 4.5 SB6 observed and synthetic seismograms 60 4.6 SB6R observed and synthetic seismograms 61 4.7 SB3 model and travel time plot 65 4.8 SB3R model and travel time plot 66 4.9 SB3 observed and synthetic seismograms 70 4.10 SB3R observed and synthetic seismograms 71 4.11 SB5 model and travel time plot 75 4.12 SB5R model and travel time plot 76 v i i i 4.13 SB5 observed and synthetic seismograms 80 4.14 SB5R observed and synthetic seismograms 81 4.15 SB4 model and travel time plot 83 4.16 SB4 observed and synthetic seismograms 86 4.17 SB4R model and travel time plot 87 4.18 Final cross-strait models 89 5.1 Map showing final models 94 5.2 Map of fault locations 97 ACKNOWLEDGEMENTS To my parents whose love and confidence have motivated me over the years, and to my wife, Ruth, for her patience and understanding. Thank you to Dr. Ron Clowes for his supervision, ideas, and friendship. The technical expertise provided by Bob Meldrum during the experiment is greatly appreciated. Also, thanks to Brad Prager, George Spehce, and Ken Whittall whose computer programs were used extensively in this study. The c r i t i c a l reading of this thesis by Dr. Bob E l l i s , and the reading of the section on geology by Dr. J.E. Muller is appreciated. The people in the Department of Geophysics and Astronomy have made my two year stay an enjoyable one. Thanks especially to Julian, Mark, Linda, and Bob (wherever you are); co-members of the class of 81-83. Also, thanks to my officemate Todd, and the members of the Department who participated in this experiment The CSP profiles provided by Tark Hamilton of the Pacific Geoscience Centre, were extremely useful in this study. The R5 section was obtained from Pan Canadian Oil Co. via Tark Hamilton. The use of the CFAV Endeavour and the cooperation of its personnel, made available by the Canadian Forces through the Defense Research Establishment, is much appreciated. I am grateful to the Science Council of British Columbia who provided financial support through a G.R.E.A.T. award, and to Great Western Petroleum for their cooperation in this matter. The CSEG also provided a scholarship. Funds for the experiment and airgun purchase were made available from NSERC Strategic Operating Grant G0738 and NSERC Strategic Equipment Grant G0739, respectively. 1 Chapter _1_ Introduction and Background The Strait of Georgia, located -between Vancouver Island and the mainland of British Columbia, is a topographic depression. It is part of the Insular Trough which extends locally from Puget Sound in the south to Dix'on Entrance north of the Queen Charlotte Islands. Roddick and Hutchison (1974) point out that this depression actually extends from the head of Lynn Canal in southeast Alaska to the Gulf of California, and suggest that i t is of continental dimensions. Tectonically, the Strait of Georgia is considered to be the locale of the boundary between the two westernmost tectonic provinces of the Canadian Cordillera: the Coast Plutonic Complex to the east and the Insular Belt to the west (Figure 1.1). Geologically, the Strait separates the late Mesozoic-Tertiary crystalline rocks of the mainland from the thick assemblage of middle Paleozoic to Jurassic volcanic, plutonic, and sedimentary rocks of Vancouver Island (Figure 1.2). This latter group is overlain unconformably by the Upper Cretaceous Nanaimo Group which outcrops on eastern Vancouver Island and the Gulf Islands. Because the underlying rocks of the Strait are not directly accessible, i t is particularly useful to consider the geology and tectonic history of the terranes on either side of the Strait. 2 F i g u r e 1.1 T e c t o n i c s of the s outhern B r i t i s h Columbian c o n t i n e n t a l margin. The p l a t e b oundaries are shown with arrows i n d i c a t i n g the p l a t e motions r e l a t i v e to the American p l a t e . The area of i n t e r e s t i n t h i s study i s o u t l i n e d by the b l a c k box. 3 1.1 Geologic Background 1.1.1 Coast Plutonic Complex The Coast Plutonic Complex (CPC) is a long narrow plutonic belt which extends from the Yukon territory to northern Washington state. It varies in width, reaching a maximum of 190 km opposite the north end of Vancouver Island. The Coastal Mountain Range is the major topographic expression of this complex. The geology of the CPC is discussed by Roddick and Hutchison (1974); the following is a brief summary. The CPC consists of a complex matrix of migmatite, gneisses, and foliated plutonic rock with minor zones of schists. High grade metamorphism is indicated in the core of the Coastal Mountains. Within this matrix are discrete and partly discrete plutons of various sizes and composition. The predominant rock types of the CPC are quartz diorite and granadiorite, with lesser amounts of diorite, diorite migmatite and gabbro, gneiss and migmatite, and small amounts of metasedimentary and metavolcanic rocks. In regions adjacent to the Strait of Georgia, granitic type rocks are dominant (Figure 1.2). No Precambrian elements have been positively identified in the CPC in British Columbia, although gneisses of this age have been identified in the northern Cascades (Mattinson, 1972). It is possible that the CPC contains a Precambrian core whose ancient character has been obliterated in the British Columbia area by more recent remobilization. Some evidence exists for Paleozoic and early Mesozoic plutonic activity, but abundant Cretaceous and Early Tertiary cooling ages (from isotopic F i g u r e 1.2 G e o l o g i c a l s k e t c h map of t h e - s t u d y a r e a . T h i s i s a s i m p l i f i e d v e r s i o n of t h e g e o l o g i c a l map o f M u l l e r (1977). The if* M a l a s p i n a f a u l t p r o p o s e d f o r t h e S t r a i t o f G e o r g i a i s shown. 5 d a t i n g ) i n d i c a t e p l u t o n i c a c t i v i t y r e a c h e d i t s peak d u r i n g t h e s e l a t t e r p e r i o d s . M i n o r i n t r u s i o n s c o n t i n u e d a t l e a s t up t o t h e end o f M i o c e n e t i m e . K/Ar d a t e s o f 93-99 m.y. have been r e p o r t e d f o r t h e s o u t h e r n C o a s t M o u n t a i n s n e a r V a n c o u v e r , i n d i c a t i n g a m i d - C r e t a c e o u s age ( W h i t e e t a l . , 1968; B a a d s g a a r d e t a l . , 1 9 6 1 ) . P a l e o l a t i t u d e measurements (Symons, 1973; Beck a nd N o s o n , 1972) on C r e t a c e o u s p l u t o n s w i t h i n t h e CPC i n d i c a t e t h a t i t was 10° t o 20° s o u t h o f i t s p r e s e n t p o s i t i o n r e l a t i v e t o c r a t o n i c N o r t h A m e r i c a when t h e y were e m p l a c e d . 1.1.2 I n s u l a r B e l t S i m i l a r l y , t h e I n s u l a r B e l t t e r r a n e i s a l o n g n a r r o w b e l t w h i c h e x t e n d s f r o m s o u t h - c e n t r a l A l a s k a t o s o u t h e r n V a n c o u v e r I s l a n d . The g e o l o g y o f t h e s o u t h e r n p a r t o f t h e b e l t on V a n c o u v e r I s l a n d i s d e s c r i b e d by M u l l e r ( 1 9 7 7 ; see F i g u r e 1.2 and T a b l e 1 . 1 ) . The o l d e s t r o c k s o f t h e I n s u l a r B e l t a r e a l a t e P a l e o z o i c v o l c a n i c a r c t e r r a n e and a c r y s t a l l i n e basement t h a t i s p r o b a b l y p r e - D e v o n i a n . The m i d d l e t o l a t e P a l e o z o i c r o c k s a r e t h e S i c k e r G r o u p w h i c h c o n s i s t s o f b r e c c i a , t u f f a n d f l o w s o f b a s a l t i c t o r h y o l i t i c c o m p o s i t i o n , u n d e r l y i n g a s e d i m e n t a r y s e q u e n c e o f a r g i l l i t e , s i l t s t o n e , and l i m e s t o n e . M e s o z o i c s t r a t i g r a p h y i s r e p r e s e n t e d by M i d d l e T r i a s s i c t o Lower J u r a s s i c v o l c a n i c and s e d i m e n t a r y r o c k s , • and Lower, t o M i d d l e J u r a s s i c I s l a n d I n t r u s i o n s . The t h o l e i i t i c b a s a l t s o f t h e K a r m u t s e n F o r m a t i o n (Upper T r i a s s i c ) u n d e r l i e most o f c e n t r a l V a n c o u v e r I s l a n d . The K a r m u t s e n v o l c a n i c s a r e o v e r l a i n i n many p l a c e s by Upper T r i a s s i c l i m e s t o n e o f t h e Q u a t s i n o F o r m a t i o n and c l a s t i c -6 c a rbonate sediments of the Parson Bay F o r m a t i o n . The T r i a s s i c sequence u n d e r l i e s the Lower J u r a s s i c Bonanza v o l c a n i c s of b a s a l t i c to r h y o - d a c i t i c f l o w s , b r e c c i a , and t u f f . The I s l a n d I n t r u s i o n s form b a t h o l i t h s and p l u t o n s of q u a r t z d i o r i t e , g r a n a d i o r i t e , and q u a r t z monzonite. S t r u c t u r a l U n i t G e o l o g i c Time T e c t o n i c H i s t o r y S i c k e r V o l c a n i c s m i d d l e - l a t e P a l e o z o i c v o l c a n i c a r c Karmutsen Formation Upper T r i a s s i c r i f t i n g b a s i n Bonanza V o l c a n i c s Lower J u r a s s i c v o l c a n i c a r c I s l a n d I n t r u s i v e s J u r a - C r e t a c e o u s v o l c a n i c - p l u t o n i c a r c P a c i f i c Rim Complex J u r a - C r e t a c e o u s s l o p e - t r e n c h d e p o s i t s deformed to melange Nanaimo Group La t e Cretaceous f o r e - a r c b a s i n T a b l e 1.1 T e c t o n i c s i g n i f i c a n c e of s t r u c t u r a l u n i t s on Vancouver I s l a n d , a f t e r M u l l e r (1977). The P a c i f i c Rim Complex on the west c o a s t of the I s l a n d i s an assemblage of J u r a - C r e t a c e o u s greywacke, a r g i l l i t e , c h e r t , and greenstone which r e p r e s e n t s s l o p e and t r e n c h d e p o s i t s t h a t have been deformed t o melange. On the e a s t c o a s t of the I s l a n d the Upper Cre t a c e o u s Nanaimo Group — c o n s i s t i n g of sandstone, s i l t s t o n e , a r g i l l i t e , and conglomerate — o v e r l i e s p r e - C r e t a c e o u s r o c k s i n c l u d i n g the J u r a s s i c I s l a n d I n t r u s i v e s . The Nanaimo group i s thought to have been d e p o s i t e d i n a f o r e - a r c b a s i n environment ( D i c k i n s o n , 1976; M u l l e r , 1977). Paleomagnetic 7 measurements made on the Karmutsen basalts (Irving and Yole, 1972; Symons, 1971; Schwarz et a l . , 1980; Irving et a l . , 1980), suggest that Vancouver Island has shifted either 13° or 44° north in latitude relative to the interior of North America. The discrepancy is due to the uncertainty in the polarity of the paleopole at the time of magnetization. Irving et a l . (1980) prefer the southerly of the two positions; Muller (1977) favours the northern position. This northward movement of the Insular Belt is also supported by paleontologic evidence (Tozer, 1970; Monger and Ross, 1971). 1.1.3 The Strait of Georgia T i f f i n (1969) gives a thorough summary of more recent geologic events in the vicinity of the Strait of Georgia. This summary will not be repeated. It will suffice to mention that there have been a number of episodes of glaciation in this region (Armstrong, 1956) as well as interglacial periods of erosion and deposition. Since Pleistocene time, the Fraser River has been the most important source of sedimentation. T i f f i n (1969) has mapped these shallow deposits over a large part of the Strait using high resolution reflection data. 1.2 Tectonic History The idea that the Canadian Cordillera is a collage of allochthonous or exotic terranes has been recognized by many authors (Jones et a l . , 1977; Coney et a l . , 1980; Jones et a l . , 1982; Monger, 1982). In particular, the exotic terrane 'Wrangellia' (Jones et a l . , 1977) has received considerable attention. The Wrangellia terrane as defined by Jones et a l . is 8 equivalent to the Insular Belt terrane considered by Muller (1977) except that Jones et a l . considers the Alexander terrane (Berg et a l . , 1972) as a separate entity (the Alexander terrane is not considered here). Wrangellia's exotic nature is based on the characteristic sequence of Triassic rocks (Karmutsen and Quatsino Formations on Vancouver Island) which differ from sequences of Triassic rocks to the east. It is generally agreed that the westward migration of the continental margin of North America since early Paleozoic time has occurred via an accretionary process. This process has involved a combination of right-lateral slip motion, collision, and oblique subduction with associated magmatism. Various authors have proposed models for the evolution of this area, each allocating a different importance to each of the above processes. The different models for emplacement of adjacent terranes in the Cordillera are important in this study, as they provide insight into the possible nature of the boundary between the CPC and the Insular Belt. Common to each of the tectonic models compiled by various authors (Monger et a l . , 1972; Dickinson, 1976; Muller, 1977; Davis et a l . , 1978; Monger and Price, 1979; Irving et a l . , 1980; Monger, 1982), is the exotic nature of the Insular Belt. The element of these models which shows the greatest variety is the role that the CPC plays in the accretionary process. Muller (1977) suggests that the Insular Belt was i n i t i a l l y formed as an Early Paleozoic volcanic-plutonic terrane on the edge of the continent, far to the south of its present position. The volcanic arc terrane was then rifted off the continent 9 accompanied by outpouring of the Karmutsen basalts in late Paleozoic-Early Triassic, and shifted northward by transcurrent or transform faulting (Monger and Ross, 1971; Jones et a l . , 1972). Similarly, Muller suggests that the predecessor of the CPC may have been rifted out of the Washington-Oregon gap in Permo-Triassic time. In Early Jurassic time, the tectonic regime of the Insular Belt changed from a r i f t i n g basin to a volcanic arc as evidenced by the Bonanza volcanics on Vancouver Island, although the northward transposition continued. According to Muller (personal communication, 1983), the Insular Belt had arrived at its present position relative to the continent by Late Jurassic-Early Cretaceous. During this time, the Insular Belt formed the arc-trench gap for a subduction _ zone along the continental margin. The Pacific Rim Complex on Vancouver Island represents the corresponding slope-trench deposits, and the associated volcanic arc was in the CPC to the east. Monger and Price (1979) suggest that the earliest appearance of the CPC is documented by the K-Ar ages of 84 to 104 m. y. for granitic rocks on the west side of the CPC (Roddick and Hutchison, 1974). The Upper Cretaceous Nanaimo Group is the remnant of the fore-arc basin. The 'Georgia Basin' at this time was inboard of the emerging Vancouver Island ranges which supplied detritus to the basin. The late Mesozoic arc-trench system along the continental margin became inactive in Late Cretaceous or early Tertiary time. The subduction zone was shifted to the Olympic Peninsula core zone about Late Eocene time (Cady, 1975; MacLeod et a l . , 1977), and then to its present position possibly in the pre-Late 10 Miocene. This tectonic model presented by Muller (1977), would indicate that the boundary between the CPC and Insular Belt is likely an old transform fault. This feature is shown in Figure 1.2 as a steeply dipping thrust fault (Malaspina fault) on the eastern side of the Strait of Georgia. L i t t l e direct evidence for this feature or its location exists. Dickinson (1976) provides a somewhat different account. He ignores the transposition of the tectonic belts and stresses the evolution of this region in terms of an 'Andean' type model. According to Dickinson, tectonic accretion of the exotic Insular Belt occurred along a subduction zone possibly by late Middle to early Late Jurassic. A prospective site for this suture zone is found on the present day continent at an inland location, near the Coquihalla River area, 20 km northeast of Hope in southwestern British Columbia. Here, strata of the Tyaughton-Methow trough — representing an ancient forearc basin (Jeletsky, 1972; Barksdale, 1975) — are in contact with an ophiolitic sequence of ultramafic and mafic rocks of the Coquihalla Belt and deformed oceanic strata (of probable Triassic age) of the Hozameen Group (Anderson, 1976). The CPC evolved later as a result of subduction related arc magmatism predominantly during Cretaceous time. This suggests that the CPC was superimposed on the Insular Belt after the Insular Belt was lodged against the continent. In light of this, Dickinson suggests that the rare Middle Jurassic ages obtained in the western part of the CPC by Roddick and Hutchison (1974) are to be viewed as relics of the Insular Belt. 11 This model supports the proposal of Monger (1975) which suggests that the suture zone between the Insular Belt and the ancient continent passes at an angle through the CPC. The notion that the CPC is a 'superimposed' terrane precludes the necessity that the western edge of the CPC represent an old lithospheric plate boundary. Monger (1982) reconsiders the tectonic history of the Canadian Cordillera from a slightly different perspective. He proposes that the CPC owes its origin not only to subduction related arc magmatism, but also, at least in part, to events related to the collision of Wrangellia with North America. As such, the CPC is a high-grade metamorphic and granitic welt which has been imposed on, and separates, rocks of the Insular Belt and the Intermontane Belt. This model again implies that the western edge of the CPC does not necessarily represent a major crustal discontinuity. Finally, concerning more recent tectonics,- Rogers (1983) .implies that the Georgia Strait-Puget Sound topographic low is to be associated with the modern tectonic regime in this area. He states that this feature is a result of the overriding North American lithospheric plate down-dropping to compensate for the phase-change related volume reduction of the subducting Juan de Fuca plate. The formation of this low would result in abundant normal faults which, at present, are inactive due to the dynamic equilibrium of plate motions in this region. 1.3 Previous Seismic Studies The Strait of Georgia has been the location of several previous seismic studies. As mentioned earlier, T i f f i n (1969) 12 has interpreted an extensive set of continuous seismic reflection profiles (CSPs). The depth of penetration obtained on these profiles is less than 1 km, so the study provides information only about the shallow sediments and depth to bedrock in the Strait. A similar survey was carried out recently (February, 1982), in southern Georgia Strait, by scientists from the Pacific Geoscience Centre (Tark Hamilton, personal communication, 1982). A high frequency energy source was used, again providing only shallow depth penetration. Also, several single and multi-channel CDP surveys have been carried out by o i l companies in Georgia Strait, but access to these data is limited. White (1962) and White and' Savage (1965) present the results of a seismic refraction experiment in the vicinity of Georgia Strait. The data for shorter ranges, which provide the shallow crustal structure, are sparse. They propose a three-layer structure for the Strait along a line running from southern Texada Island to Campbell River. The upper layer is Upper Cretaceous sediments; the second layer has a velocity near 5.9 km/s; the third layer has a velocity near 6.8 km/s. The depth to the third layer is less than 5 km. A velocity of 3.81 km/s was obtained for the Upper Cretaceous sediments using shots detonated between Vancouver Island and Texada Island which were recorded on Hornby Island. The thickness of this layer was calculated to be 1.7-2.3 km. Also, Milne and White (1960) measured a velocity of 4.05 km/s for outcrops of this formation at Galiano Island. Tseng (1968) re-examined an extended set of available 13 r e f r a c t i o n d a t a , a p p l y i n g a t i m e - t e r m a n a l y s i s . He a l s o p r o p o s e s a t h r e e l a y e r s t r u c t u r e f o r t h e upper c r u s t . The p r o p o s e d model f o r t h e S t r a i t of G e o r g i a i s shown i n F i g u r e 1.3. N S 0 i 2.5 -D e p t h ( k m ) 5.0 -7.5 J F i g u r e 1.3 V e l o c i t y s t r u c t u r e of the upper c r u s t f o r n o r t h ana soutfi G e o r g i a S t r a i t ( a f t e r Tseng, 1968). The v e l o c i t i e s shown are compress 1ona1 wave v e l o c i t i e s . B e r r y and F o r s y t h (1975) i n t e r p r e t e d a l o n g range s e i s m i c r e f r a c t i o n p r o f i l e which e x t e n d s from V a n c o u v e r I s l a n d t o t h e i n t e r i o r of s o u t h e r n B r i t i s h C o l u m b i a . B ased on t h e s c a t t e r i n g of s e i s m i c e n e r g y i n the v i c i n i t y o f G e o r g i a S t r a i t f o r e a s t w a r d and westward t r a v e l l i n g waves, t h e y h y p o t h e s i z e t h a t a major f a u l t zone may e x i s t a t t h e e a s t s i d e of t h e S t r a i t of G e o r g i a . 1.4 E x p e r i m e n t a l O b j e c t i v e s The n a t u r e o f t h e boundary between t h e I n s u l a r B e l t and t h e CPC i s u n c e r t a i n . The s p e c u l a t i v e n a t u r e of the boundary, a t p r e s e n t , i s b a s e d on i n f e r e n c e s made m o s t l y from s u r f i c i a l 14 geology and sparse seismic data. In view of this, the objectives of this experiment are: a/ To investigate the nature of this boundary beneath the Strait of Georgia, and, in particular, to explore the proposed existence of a major fault separating Vancouver Island from the continental mainland. b/ To obtain a more detailed model for the upper crustal structure beneath the Strait of Georgia. 15 Chapter 2 Data Acquisition and Reduction 2.1 The Refraction Experiment In May of 1982, a marine refraction survey was conducted in the Strait of Georgia by the University of British Columbia. The area of the survey extends from south of Texada Island to the northern tip of Galiano Island. Three reversed cross-strait profiles as well as a number of profiles along the Strait were shot (Figure 2.1). The orientation of the cross-strait profiles was chosen in order to detect any fault structure that might run along the Strait. The along-strait profiles were designed to intersect these profiles to provide control in any subsequent velocity modelling. A 32 l i t r e airgun was used as the seismic energy source and shots were recorded for a l l of the profiles using a telemetering sonobuoy system with direct digital recording. In addition, the southernmost cross-strait line, from northern Galiano Island to Point Grey, was recorded digitally from a telemetered land-based station on Galiano Island. Profile lengths range from 15 to 30 km with an average shot spacing of about 250 m. The refraction data are supplemented by several high resolution reflection profiles made available by Tark Hamilton from the Pacific Geoscience Centre. The locations of these profiles also are shown in Figure 2.1. F i g u r e 2.1 Map of p r o f i l e l o c a t i o n s . Sonobuoy r e f r a c t i o n l i n e s SB 1 -SB6 and v e r t i c a l i n c i d e n c e r e f l e c t i o n 1ines R3, R5. and R6 are shown. 17 2.2 Instrumentation 2.2.1 General System Figure 2.2 shows a flow diagram outlining the schematic configuration of the instruments used during the refraction experiment. At approximately 2 seconds before firing time, the timing control started the digital recording system (Convert in Figure 2.2). Then a signal was relayed to the firing box to initiate the firing of the airgun, and to the stepper box to cue the X-Y plotter. This shot signal was recorded on channel 1 of the digital tape while the seismic signal was recorded on channel 2. The digital recording system ran for 20 seconds during which the A/D converter digitized at 312.5 sps, per channel, in real time. At the end of 20 seconds, the timing control prompted the controller to write an end-of-file (EOF) on the tape to separate shot f i l e s . The X-Y recorder was used to monitor the incoming seismic signal. After each shot, there was a unit step increase in the d.c. voltage (at the stepper box) added to the seismic signal recorded on the X-Y plotter causing the pen to 'step up' so that the next trace did not overwrite the previous one. An analog FM tape recorder was used as a backup to the primary digital system, recording the shot signal, the seismic signal, and the time code from radio WWVB on channels 1, 3, and 4, respectively, at a tape speed of 15/16 ips. Channel 2 recorded the internal compensation of the recorder. The position of the ship and deployment locations of the sonobuoys were i n i t i a l l y determined using a Loran C navigational system. Shortly after the start of the experiment, the Loran C 18 © Firing B o x Stepper B o x I C \ . F M Recorder * X - Y P lotter Timing Contro l W W V B Sonobuoy Rece iver 5 - 6 2 Hz E O F Conver t Contro l le r 1 2 3 1 2 . 5 s p s Buffered Formatter Tapedr ive F i g u r e 2.2 System flow diagram. The major components of the system used f o r the r e f r a c t i o n experiment a r e shown. 19 coordinates were found to be in error by several kilometres due to a malfunction of the system. As a result this system was abandoned. Instead, range and bearing measurements to landmarks were made by .the navigators of the CFAV Endeavour, approximately every 15 minutes. The seafloor topography was recorded continuously on paper charts by the ship's depth sounder, and time marks were annotated on these charts every 15 minutes. 2.2.2 Instruments A BOLT Associates 32 l i t r e airgun, discharged at a pre-shot pressure of 13.790 MPa (2000 psi), at a depth of approximately 24 m below the water surface, provided the seismic energy source. The nearfield waveform produced by the airgun is shown in Figure 2.3. It was recorded using a hydrophone deployed approximately 10m from the airgun. The sonobuoys used in the experiment were REFTEK 1 expendable sonobuoys which consist of an AGC (Automatic Gain Control) seismic amplifier, radio transmitter, and a hydrophone which is suspended 18 m below the water surface. The AGC level in the buoy is determined by monitoring the background noise level; the gain being adjusted such that the noise level is 15 dB below clipping. The AGC has a 60 second response time, and thus will not normally change during the seismic signal time window. The overall system has a flat acoustic response (±3 dB) from 10 to 2000 Hz. The seismic signal from the sonobuoy was transmitted to an Aquatronics Telseis STR 71-2F receiver onboard the ship. The receiver has an adjustable gain, and bandpasses the incoming signal from 5 to 62 Hz. After completion of each profile, the 20 F i g u r e 2.3 Source P u l s e a/ The n e a r f i e l d s o u r c e p u l s e . N o t i c e the f i r s t bubble p u l s e near 0.16 s. b/ The power spectrum of the source p u l s e to 156 Hz i s shown. The spectrum i s n e a r l y f l a t over t h i s range. c/ The source p u l s e a f t e r bandpass f i l t e r i n g from 5-30 Hz. The f i r s t bubble p u l s e has a magnitude comparable to the o r i g i n a l p u l s e i n t h i s frequency range. 21 crystal in the receiver and the frequency of the next sonobuoy were changed to ensure signal transmission/reception at a different frequency from the previous sonobuoy, which would s t i l l be transmitting. The land-based seismometer on Galiano Island is part of the Lower Mainland Seismic Array, a permanent installation operated by the Seismology Lab at U.B.C.. The signal from the Willmore Mark II seismometer was telemetered to the lab where i t was recorded digitally at a rate of 60 sps. 2.3 Data Processing and Corrections A number of processes were applied to the raw data in order to produce seismic sections, suitable for interpretation. These processes are outlined below. 2.3.1 I n i t i a l Computer Processing The sonobuoy data were recorded digitally on magnetic tape in 'multiplexed' format. Backup copies of the original data tapes were made and the data were demultiplexed. The data for a large number of shots had to be recovered from the backup FM tape due to malfunction of the A/D converter during the original recording. These analog records from FM were digitized, recorded on magnetic tape, and then demultiplexed. The data from the land-based station on Galiano Island were recorded in a format which required no i n i t i a l processing. 2.3.2 Shot-Receiver Positions Ship positions were determined at 15 minute intervals from range-bearing readings. It is d i f f i c u l t to establish the uncertainty in the absolute ship positions determined in this 22 manner, but i t is l i k e l y less than ± 0 . 5 km. This estimate is arr ived at by assuming that the actual pos i t ion determination is good to within 100 m ( i . e. ±50 m), but that the time of measurement has an uncertainty of ±3 minutes. Uncertaint ies for posi t ions taken at the beginning or end of a p r o f i l e are l i k e l y less than t h i s . Average v e l o c i t i e s of the ship between successive measurements are calculated and these are used to infer the locat ion of each of the shots occurring at two minute i n t e r v a l s . The uncertainty in these locations is s l i g h t l y more than for the 15 minute locations ( ± 0 . 6 km). Because the sonobuoy receivers used in th i s experiment are unanchored, they move from the deployed pos i t ion due to l o c a l water currents and wind e f fec ts . This uncertainty in the pos i t ion of the sonobuoy introduces complications in shot-receiver distance determination and in subsequent modelling procedures. I n i t i a l l y , two methods were considered for ascertaining shot-receiver distances for the sonobuoy l i n e s . The f i r s t method, referred to as the 'range-bearing method', i s simply to ca lcu late the difference between the shot locations and i n i t i a l sonobuoy pos i t ion that have been arr ived at from range-bearing measurements. These shot-receiver distances w i l l have minimum uncertaint ies of ±0.8 km assuming that the sonobuoy remains stat ionary re la t ive to land. This uncertainty i s extremely large in re lat ion to the magnitude of the typ i ca l distances in th i s experiment. The second method u t i l i z e s the d irec t water wave a r r i v a l on the seismic traces, and is referred to as the 'water wave method'. Relat ive shot-receiver distances can be ca lculated with uncertaint ies of less than ± 0 . 1 km, based 23 on uncertainties in the 'pick' of the water wave arrival and the water velocity. As in the f i r s t case, the absolute position of the sonobuoy at later times is unknown. This method is used to determine shot-receiver distances for the sonobuoy lines, assuming a water velocity of 1.49 km/s. Comparisons of shot-receiver distances determined using the water wave method and those using the range-bearing method, suggest that sonobuoys drifted distances of 0.4-5.0 km, over the course of a single prof i l e . The position of the Galiano station has been determined using 1:50,0.00 topographic maps. It is located at 49° 00' 44'' N and 123° 35' 00'' W, and has an elevation of 7.6 m above sea level. The uncertainty in the position measurement is less than ±25 m. For profiles recorded at this station, the range-bearing method is used to determine shot-receiver distances since there is no direct water wave arrival. Table 2.1 summarizes the uncertainties in shot-receiver distances and positions. Measurement Method of Determination Uncertainty relative shot-receiver di stance range-bearing method ±0.8 km relative shot-receiver distance water wave method ±0.1 km absolute shot position range-bearing method ±0.6 km T a b l e 2.1 P o s i t i o n u n c e r t a i n t i e s . Methods of d i s t a n c e d e t e r m i n a t i o n are d i s c u s s e d i n s e c t i o n 2.3.2. 24 2.3.3 Amplitude Corrections Due to the AGC component of the sonobuoy system (see section 2.2.2) and the variable gain setting of the Telseis receiver, the relative amplitudes of the seismic arrivals are distorted. In an attempt to restore this amplitude information certain assumptions must be made. Any changes in the gain of the system due to the AGC component cannot be detected by monitoring the recorded background noise level. Thus, i t is assumed that the background noise level is constant over the course of a profile, and any changes in the recorded background noise level are due to changes in the system gain at the receiver. Under these assumptions, the recorded background noise level is a direct indicator of the system gain. The energy in a segment of the data containing only background noise is sampled for each shot trace. The data is then scaled by a factor which is inversely proportional to the energy level of the recorded background noise. For the Galiano Island data, the gain was constant during recording so no amplitude corrections have been applied. 2.3.4 Clock Drift The shipboard clock, which controls the shot times, drifted with respect to the Galiano Island clock (WWVB controlled) during the shooting of profiles GAL5 and GAL5R. Corrections for the shot times for these two lines are calculated by assuming a linear d r i f t rate and fit t i n g a straight line through drift times that have been sampled over the course of these profiles. Drift rates for both profiles are near 60 ms/hr. The maximum uncertainty due to clock d r i f t , after correction, is ±10 ms. 25 2.3.5 Timing Uncertainties The factors contributing to uncertainties in timing include shot time determination, clock dr i f t (for Galiano Island data), digital sampling rate fluctuations, and picking errors. The magnitudes of these uncertainties are summarized in Table 2.2. Source of Uncertainty Magnitude shot time clock drift (Galiano Is. station) sampling rate fluctuations travel time picks ±3 ms ±10 ms ±3 ms ±5 to ±40 ms T a b l e 2.2 Timing u n c e r t a i n t i e s . The shot time as determined by the electrical triggering signal was recorded digitally on magnetic tape, and is uncertain by one sampling interval which is ±3 ms. The uncertainty due to the clock d r i f t correction is outlined in section 2.3.4. Digital sampling rate fluctuations occur only for the data that were digitized from the backup FM tapes, and are due to small variations in the analog tapedrive speed and very small amounts of tape stretching. Fluctuations of this type are responsible for timing uncertainties of less than ±3 ms (Waldron, 1982). Uncertainties involved in 'picking' arrival times are dependent on the S/N ratio of the data. For small shot-receiver distances, where S/N ratios are high, the picking uncertainty may be as 26 small as ±5 ms. For larger distances, picking uncertainties may increase to ±40 ms. Generally, the previously mentioned timing uncertainties are small compared to the uncertainty involved in making the arrival picks. 2.3.6 Filtering Figure 2.4 shows the power spectrum for two representative segments of data. The f i r s t contains background noise only and the second is for a segment containing refracted energy. The seismic energy is concentrated in a band from 5 to 30 Hz, while the noise energy is concentrated mainly at a higher frequency with a small component at the low end of the spectrum. An 8-pole 5-30 Hz Butterworth bandpass f i l t e r (Kanasewich, 1981) has been applied to a l l of the data. NOISE SEISMIC Frequency (Hi) Frequency (Hz) F i g u r e 2.4 S e i s m i c Power Spectrum a/ The power spectrum f o r a segment of n o i s e i s shown. The n o i s e energy i s c o n c e n t r a t e d at the h i g h e r f r e q u e n c i e s . b/ The power spectrum f o r a segment of data c o n t a i n i n g r e f r a c t e d energy i s shown. The s e i s m i c energy i s c o n t a i n e d mainly between 5 and 30 Hz. 27 Chapter 3 The Data: Observations and Modelling Approach 3.1 I n i t i a l Observations on the Data Set ~ The seismograms for a l l of the profiles are shown in Figures 3.1 to 3.8. The data have been bandpassed from 5 to 30 Hz, and are plotted using a reducing velocity of '6.0 km/s. Plots use a variable area format with the amplitudes scaled in direct proportion to the shot-receiver distance. Velocities are compressional wave velocities, unless otherwise stated. 3.1.1 The Cross-strait Profiles a/ SB6/6R The profiles SB6 and SB6R (Figure 2.1) are the northernmost reversed, seismic refraction lines of the survey. The observed seismograms for these two profiles are shown in Figure 3.1. First arrivals are picked over the entire length of the SB6 seismogram (out to 19 km), although after 16 km they are uncertain. The curve defined by the f i r s t arrivals is 'bumpy', showing the effects of bottom topography. There are, however, no major offsets in the f i r s t arrivals which would indicate the presence of a fault structure. The f i r s t arrivals define an apparent velocity near 6 -km/s after a short distance along the profile, suggesting a high velocity medium exists at shallow depth. A velocity near 6 km/s persists over the length of the profile. The amplitudes of the f i r s t arrivals, on the SB6 seismogram, generally show only small variations over the length 28 0 2 4 6 8 10 12 14 16 18 DISTANCE (KM) SB6 DISTANCE (KH) SB6R F i g u r e 3.1 SB6/GR r e f r a c t i o n d a t a . Observed seismograms f o r the SBG (upper) and SB6R (lower) p r o f i l e s a re shown p l o t t e d with a r e d u c i n g v e l o c i t y of 6.0 km/s. The dat a are bandpass f i l t e r e d (5-30 Hz), and a s p r e a d i n g f a c t o r p r o p o r t i o n a l to d i s t a n c e has been a p p l i e d to enhance the d i s t a n t a r r i v a l s . A l s o , amplitude c o r r e c t i o n s have been a p p l i e d . The arrows i n d i c a t e the p i c k e d f i r s t a r r i v a l s where the data are no i sy. F i gure f e a t u r e s acoust i c 3.2 R3 and RG CSP r e f l e c t i o n data. The on these c r o s s - s t r a i t p r o f i l e s a re shown, basement, and the sonobuoy p o s i t i o n s are 1 major t o p o g r a p h i c The arrows i n d i c a t e abe1 l e d . 30 of the profile. The most conspicuous amplitude feature on the seismogram is the large amplitude secondary arrival whose onset defines a straight line from the origin to a time of 2.5 seconds, at a distance near 5 km. This is the direct water wave arrival. The only other observed secondary arrivals mimic the f i r s t arrivals, but are offset with respect to time, revealing them as multiples. Note, also, the waveform of the f i r s t arrivals, manifesting the double pulse nature of the source wavelet (Figure 2.3). First arrivals can also be picked over the entire length of the SB6R record section (Figure 3.1). As for the reverse line, there are no large time offsets in the f i r s t arrivals. Unlike the SB6 section, the f i r s t arrivals define an apparent velocity close to 3.4 km/s out to a distance of 4.5 km. At this point, the f i r s t arrivals break over to a velocity of greater than 6.0 km/s which is maintained to 21 km distance. The intermediate and high velocity arrivals ( i . e. 3.4 and 6 km/s) will be referred to as p, and p 2 arrivals, respectively. There are significant amplitude variations of the f i r s t arrivals in this case. The f i r s t arrival amplitudes decrease between 7 and 10 km, increasing again after this distance. Besides the direct water wave arrival, there are other large amplitude secondary arrivals on the SB6R section that are not present on the SB6 section. These arrivals — observed to a distance of 7-8 km — do not clearly define an arrival branch. The only secondary arrivals observed after 8 km distance are multiples of the f i r s t arrivals. 31 b/ R6 The reflection profile R6 (Figure 3.2) is a single channel, continuous seismic profiling (CSP) line, obtained using an array of small displacement airguns. This CSP line is collinear with the refraction lines SB6 and SB6R (see Figure 2.1), and provides information concerning the shallow structure along the profiles. As shown on the section, the topography of the water-bottom interface is variable along this profile. Two major topographic features (Round Ridge and Sangster Ridge) show a direct correlation with the early arrivals in the f i r s t arrival curves of the SB6 and SB6R sections. Acoustic basement (or bedrock) can be identified over most of this profile. c/ SB3/SB3R The profiles SB3 and SB3R (see Figure 2.1) are the middle pair of cross-strait refraction lines. First arrivals are visible to 18 km distance on the SB3 seismogram (Figure 3.3). The gap in the data set after 18 km is due to a compressor breakdown. Firing was continued after repairs were completed, but, as can be seen, no arrivals are observed. As was the case for the SB6 line, no P i arrivals with intermediate apparent velocities are observed as f i r s t arrivals; p 2 arrivals break over to an apparent velocity near 6 km/s almost immediately. An apparent velocity somewhat less than this continues to 16 km where the f i r s t arrivals define a velocity greater .than 6 km/s. Again, there are no major offsets in the arrivals to suggest a fault structure. The amplitudes of the f i r s t arrivals are essentially constant out to 14 km, where they decay for the remainder of the section. Secondary arrivals seem to be 3 2 -—— •—J-—- ^^>J^ A/vA^ A^ A^^ -fc^^wVJV— —1 -—• —-~^ A^/-*^ --^ t^fs^ -tf-<>-'v^ -— 'Z Q'Z S*l 0*1 S ' (33S) 0 * 9 / Q - i •+-CO CM CM CD CNI 00 5 ! cn m co t-CM ro ~~ CQ CO S*2 0*2 S ' l 0*1 S* (33S) 0 ' 9 / Q - l F i g u r e 3.3 SB3/3R r e f r a c t i o n data. Observed seismograms f o r the SB3 (upper) and SB3R (lower) p r o f i l e s a r e shown p l o t t e d with a r e d u c i n g v e l o c i t y of 6.0 km/s. The data a re bandpass f i l t e r e d (5-30 Hz) and a s p r e a d i n g f a c t o r p r o p o r t i o n a l to d i s t a n c e has been a p p l i e d to enhance the d i s t a n t a r r i v a l s . A l s o , amplitude c o r r e c t i o n s have been a p p l i e d . The arrows i n d i c a t e the f i r s t a r r i v a l s where the data a re n o i s y , as well as a secondary a r r i v a l branch on the SB3R s e c t i o n , between 12 and 16 km. 33 dominated by multiples of the f i r s t arrivals. The best example of this can be seen between 10 and 14 km, near 1.35 s. For the SB3R seismogram (Figure 3.3), two groups of intermediate velocity (pi) arrivals can be identified. An apparent velocity of 2.32 km/s is defined by the f i r s t arrivals (p,, arrivals) out to almost 2 km. The f i r s t arrivals ( p i 2 arrivals) have an apparent velocity of 3.68 km/s from 2 to 7 km, where the arrival curve breaks over to velocities greater than 6 km/s for the p 2 arrivals. A secondary arrival branch (p 3 arrivals) can be seen between 12 and 16 km, later than 1.5 s, having an apparent velocity of 4.55 km/s. This is likely the continuation of the arrival branch. The amplitudes of the arrivals between 2 and 7 km ( i . e. p 1 2 arrivals) deteriorate after 4 km. The f i r s t arrival amplitudes after 7 km (p 2 arrivals) show only small variation, and are somewhat smaller than the amplitudes of the secondary arrival branch defined by the p 3 arrivals. d/ R3 R3 (Figure 3.2) is a single channel, CSP line which parallels the SB3 and SB3R refraction lines, but i t is located 1-2 km southeast of these lines (see Figure 2.1). Acoustic basement is readily recognized on the left side of the R3 section, but proceeding to the right i t becomes highly speculative. e/ SB5/SB5R The profiles SB5 and SB5R (Figure 2.1) are the southernmost refraction lines. On the SB5 seismogram (Figure 3.4), the py F i g u r e 3.4 SB5/5R r e f r a c t i o n d a t a . Observed seismograms f o r the SB5 (upper) and SB5R (lower) p r o f i l e s a r e shown p l o t t e d w i th a r e d u c i n g v e l o c i t y of 6.0 km/s. The data have been bandpass f i l t e r e d (5-30 Hz) and a s p r e a d i n g f a c t o r p r o p o r t i o n a l to d i s t a n c e has been a p p l i e d . Amplitude c o r r e c t i o n s have been a p p l i e d . The o f f s e t i n f i r s t a r r i v a l s on the SB5 s e c t i o n i s i n d i c a t e d by the arrows, as are the f i r s t a r r i v a l s on the SB5R s e c t i o n . A secondary a r r i v a l branch between 4 and 6 km i s a l s o i n d i c a t e d on the SB5R s e c t i o n . 35 arrivals i n i t i a l l y define an apparent velocity of 3.2 km/s, to a distance of 5 km. At this point, the arrival curve breaks over to an apparent velocity near 6 km/s, continuing to a distance near 8 km where the p 2 arrivals fade out. For distances between 9 and 12 km, weak arrivals defining a similar velocity can be seen at times close to 1.4 s. This part of the arrival curve is offset in time by 0.3 s relative to the p 2 arrivals up to 8 km distance. Assuming that these later arrivals do represent the continuation of the p 2 arrivals, this time offset indicates the possible existence of a fault structure or low velocity zone. But, due to the low S/N ratio of the data between 8 and 12 km distance, this observation must be viewed with some reservation. The amplitudes of the p, and p 2 arrivals to a distance of .8 km on the SB5 seismogram remain relatively constant. Similar to the SB6R and SB3R seismograms, large amplitude secondary arrivals follow the f i r s t arrivals for the f i r s t 7-8 km, although they do not define readily observable travel time branches. The S/N ratio for a l l of the SB5R data (Figure 3.4) is low making the picking uncertainties large. No p, f i r s t arrivals are observed, but p 2 arrivals can be picked on most traces to a distance of 19 km. They define an apparent velocity near 6 km/s after a very short distance, and this velocity continues to 19 km. The amplitudes of the p 2 arrivals are small over the distance range 6-12 km, and other than multiples, there is a secondary arrival between 4 and 6 km, extending from a time of 1.14 seconds, at 3.64 km, to a time of 1.40 seconds, at 5.65 km. This arrival has an apparent velocity of 3.3 km/s. The travel 36 time picks for this arrival have large uncertainties (±20 ms). f/ GAL5/GAL5R The GAL5 profile is the one for which the land-based station on Galiano Island recorded the same sequence of shots as were recorded by sonobuoy for line SB5. Similarly, GAL5R is a record from the same land-based station for the same shots that were recorded for SB5R. Thus, these lines do not represent a reversed set, but, instead, represent a repeat of the same experiment, to the extent that the positions of shots for SB5 and SB5R coincide. In the absence of a direct water wave arrival at the land-based station, the shot-receiver distances for these two profiles were determined using the range-bearing method (see section 2.3.2). Because of the large uncertainty in the positions determined using this method, as compared to the method used for the sonobuoy lines, the SB5 profile is used for interpretation in this vicinity of the Strait. The effect of these large position uncertainties is seen (Figure 3.5) in comparing the GAL5 and GAL5R seismograms. Because of the near coincidence of the shot lines used for these two profiles, the record sections should be very similar, but comparison of the apparent velocity defined by the f i r s t arrivals for the two seismograms between 7 and 10 km, for instance, shows a substantial difference. The general characteristics of these record sections have been used as an aid in interpreting the SB5 and SB5R profiles. It is unfortunate that the. distances determined for the GAL5/GAL5R data sets are not more accurate. The S/N ratio on these record sections (particularly for GAL5R) is much higher 37 F i g u r e 3.5 GAL5/5R r e f r a c t i o n data. Observed seismograms f o r GALS (upper) and GAL5R (lower) are shown p l o t t e d with a r e d u c i n g v e l o c i t y of 6.0 km/s. The data a re bandpass f i l t e r e d (5-30 Hz) and a s p r e a d i n g f a c t o r p r o p o r t i o n a l to d i s t a n c e has been i n c l u d e d . No amplitude c o r r e c t i o n s have been a p p l i e d as the r e c e i v e r had con s t a n t g a i n . Arrows i n d i c a t e the o f f s e t in the f i r s t a r r i v a l s . 38 than for the corresponding sonobuoy profiles. On both of these record sections, f i r s t arrival picks can be made out to distances of close to 30 km. On the GAL5R section, the offset in the f i r s t arrivals that was suggested by the SB5 data, is very clear (between 10 and 14 km). A comparison of the positions of this offset on the GAL5, GAL5R, SB5, and SB5R seismograms, localizes this feature suggesting that i t is likely due to a fault structure rather than a low velocity zone. 9/ R5 R5 (Figure 3.6) is a 12-fold stacked reflection profile which crosses Georgia Strait within a maximum distance of 2.5 km from the SB5/SB5R refraction lines (see Figure 2.1). The f i r s t 0.4 s of data are muted on this section. A prominent reflector beneath flat-lying layers is identified over the f i r s t half of the profile . This reflector appears to end abruptly toward the right side of the upper part of the section, corresponding to a position in Georgia Strait 1.6±0.8 km from the position of the offset in the SB5 record section/Deeper reflectors on this part of the section are obscured by multiples. On the left half of the lower part of the section, acoustic basement is obscured by multiples and diffraction patterns from shallow reflectors. On the right half of this part of the section, acoustic basement is again observed as well as other deeper reflectors. h/ Summary of General Characteristics Generally, the quality of the data for the cross-strait profiles is good. Except for the SB5 and SB5R lines, f i r s t arrival picks can be made over the f u l l length of the profiles. R 5 C D P P R O F I L E i— j- i. i i Water Depth (ft) —j—aoo ooo NE • SB5R •°°Ml j- "I j l-L.J:-L:-J. -i-rHftf f• t ^ r - r i f f o.oi,—j-. ! ! J.--j • I- Fraser Delta • I j I 4. ._4_ J.. J - j f • j TWO-WAV REFLECTION TIME (S> MAINLAND 0 SCALE: 0 L. 5 KM -J F i g u r e 3.6 R5 CDP p r o f i l e . T h i s i s a 1 2 - f o l d coverage r e f l e c t i o n p r o f i l e with the f i r s t 0 . 4 s of dat a muted, and the dashed l i n e showing the water depth. The v e r t i c a l e x a g g e r a t i o n i s much, l e s s than f o r the CSP p r o f i l e s . The lower s e c t i o n i s a c o n t i n u a t i o n of the upper s e c t i o n . The bedrock s u r f a c e i s i n d i c a t e d by the arrows on the upper s e c t i o n , and the 'F' i n d i c a t e s the p o s i t i o n of the proposed f a u l t . On the lower r i g h t s e c t i o n , bedrock and a deeper r e f l e c t o r a re i n d i c a t e d by the arrows. The p o s i t i o n s of the sonobuoys a r e l a b e l l e d . co 40 The lines running across Georgia Strait from the mainland to Vancouver Island are characterized by a break over from the water wave arrival directly to a velocity slightly less than 6.0 km/s. p, arrivals are not observed on these profiles except on the SB5 profile where they appear as secondary arrivals. Thrs suggests that a high velocity medium exists at shallow depth on the mainland side of the Strait, and any intermediate velocity sediments, i f present, are thin. In contrast, the lines running from Vancouver Island to the mainland are characterized by arrivals out to distances of 5-6 km which define an intermediate apparent velocity of 2.3-3.4 km/s. The p, arrivals are succeeded by p 2 arrivals for greater distances, defining velocities exceeding 6.0 km/s. The difference in apparent velocities for the forward and reverse lines, across the Strait, indicate that the higher velocity structure may dip from the mainland to Vancouver Island. Also, the profiles running toward the mainland have large amplitude secondary arrivals for the f i r s t part of each of the profiles, which are not observed for the reverse profiles. For the two northern lines — SB6/SB6R and SB3/SB3R - there are no offsets in the f i r s t arrivals that suggest the presence of a fault structure. On the southernmost line, an offset in the data is present indicating an anomaly in the velocity structure in this part of the Strait. 3.1.2 The Along-strait Profiles a/ SB1 The SB1 refraction line (Figure 2.1) runs northeast along the Strait of Georgia. It is situated approximately 10 km from 41 hc\i l 1 1 1 1 h CD 9'Z O'S S'l O'l S* 0 (33S) 0*9/0-1 F i g u r e 3.7 SB 1 r e f r a c t i o n data. The observed seismogram f o r SB 1 is shown p l o t t e d w ith a r e d u c i n g v e l o c i t y of 6.0 km/s. The data have been bandpass f i l t e r e d (5-30 Hz) and a s p r e a d i n g f a c t o r p r o p o r t i o n a l to d i s t a n c e has been a p p l i e d . No amplitude c o r r e c t i o n s have been a p p l i e d . The arrows i n d i c a t e the p r i m a r y and secondary a r r i v a l s to a d i s t a n c e near 6 km. 42 the mainland coast. The observed seismogram for this line is shown in Figure 3.7. First arrivals with an apparent velocity of 5.5 km/s can be picked out to 6 km, but are not observed for greater distances, and a secondary arrival branch with velocity 3.6 km/s is also observed to a distance near 6 km. Two other profiles shot adjacent to SB1 yielded similar poor results and are therefore not presented. Furthermore, the navigation problems (mentioned in section 2.2.1) were discovered after the SB1 profile was complete, making the absolute position measurements questionable. Because of this, and the short distance over which arrivals are observed, no detailed modelling is done for this profile. The lack of coherent seismic energy received in this area merits, further investigation in future studies. b/ SB4/SB4R SB4 and SB4R (Figure 2.1) are a partially reversed pair of refraction lines running along Georgia Strait, on the Vancouver Island side. First arrivals can be picked out to 12 km on the SB4 seismogram (Figure 3.8). p, arrivals with an apparent velocity of 3.9 km/s are observed over the f i r s t 3 km distance, after which the p 2 arrivals define a velocity of near 6 km/s. This velocity is maintained to 12 km except between 6 and 10 km where i t is somewhat less than this. Like the cross-strait profiles which started on the Island side of the Strait, large amplitude secondary arrivals follow the f i r s t arrivals for some distance. The most prominent amplitude characteristic of the f i r s t arrivals is the sudden decrease in amplitude after 9.5 km. On the SB4R seismogram (Figure 3.8), p, arrivals define an rt Ol —^ -J TJ Cl Oi TJ r < — CO ID - O — ID a i rt • N —1 0) zr 3 ro a 3 D) 01 "5 Ul "J TJ O "5 £ ID (/) (1) 01 CD a fe — TO - Zi 3 IO. ui a (B —• -h n o oi r+ ni n r+ r+ o - o zs zs 1 • io TJ -> -ti o - TJ 3 O U) "J o 0) 3 1 01 1 — < 01 rt — o -j cn ID CD Q. fe -n C O ~ ( Q c c (Q TJ ro (D -J < w CO ID — oi oo O 3 o a < < / ! < / > 03 CD fe O J J \ -ti fe XI cn — • o -J 0 £ ro <D -ti n T 3 w 0) \ n 01 rt • TJ - • " J O H O D ZS -ti ID — a — 01 Q. ro rt ai u 01 rt 01 01 o zr s a oi ro ui ro "5 Ul < 3- ro o a £ 3 Ul IB 3- ro ID O 3 -h 01 rt, U) Ul ro rt a ro - ro 3 3 0" 01 TJ Ul 3 - 3 a o o TJ rt IQ 01 rt -j Ul ID 01 Ul a 3 VI -h — £ -ti - O r+ rt 1 ID 3 "J rt ro 3 a oi ro LO CM' C D CM' LLl co LD C O \ i L O CM* CD C M C_> U J C O L O C O C3 o -in 6 8 .10 DISTANCE (KM) 14 14 DISTANCE (KM) SB4R 44 apparent velocity of 3.6 km/s out to 4 km distance, where they are succeeded by p 2 arrivals with apparent velocity near 6 km/s. The p 2 arrivals seem to fade out near 6 km, although arrivals are visible at greater distances at later times. Due to the low S/N ratio of the data near 6 km, the data trends are obscure, but the signature of the SB4R data in this region is very similar to that for SB5, indicating once again the possible presence of a fault structure in this vicinity of Georgia Strait. No arrivals are observed past 12 km. 3.2 Modelling Method 3.2.1 General Procedure Originally, i t was intended that the refraction data set be interpreted using the techniques of slant stacking and wavefield continuation (Clayton and McMechan, 1981), which are applicable to one-dimensional velocity structures ( i . e. where velocity varies only with depth). Shot spacings were deliberately small during shooting in order that the data have high spatial density, suitable for application of this inversion technique. I n i t i a l observations, however, made on the reversed refraction and reflection lines made i t apparent that a two-dimensional inversion (allowing for lateral inhomogeneities due to topography) of the data was required to represent the velocity structure. When dealing with a two-dimensional velocity medium, there are many variables to be considered: depths to layers, velocities, vertical and horizontal velocity gradients, and topography of layers. To make this problem tractable, as many constraints as possible are needed. The topography of the water-45 bottom interface is provided by depth sounding records, and constraints on the depth of shallow sediments and topography of the underlying bedrock are provided by the reflection profiles. For the reversed cross-strait lines, the dip and velocity of deeper layers are constrained. Some constraints on velocity gradients are provided by the amplitude characteristics of the data, in the absence of clearly defined secondary travel time branches. Horizontal velocity gradients are avoided. Preliminary estimates of depths to velocity interfaces,, not observed on the reflection profiles, are obtained by assuming a model of horizontal, constant velocity, plane layers. Comparison of these depths from reversed profiles starting on opposite sides o'f the Strait, as well as apparent velocities, provide estimates of the dips of these layers. This information, along with information from the accompanying reflection profile, provides" a starting model to begin the modelling procedure. Modifications to these starting models are made by attempting to match the travel times of rays traced through the model to those observed on the data sections. This is done.via a forward modelling ( i . e. t r i a l and error) ray tracing technique, based on the algorithm of Whittall and Clowes (1979), which calculates travel times for rays traced through a two-dimensional velocity model. Having f i t the travel times to the desired 'goodness of f i t ' , an attempt is made to reproduce the amplitude characteristics of the data sections by calculating synthetic seismograms for the proposed velocity models. Using these two modelling procedures requires a trade-off between matching of travel times and amplitude characteristics for the observed and 46 synthetic seismograms. The Asymptotic Ray Theory (ART) algorithm of Spence et a l . (1983) is used to calculate synthetic seismograms for the two-dimensional models. A brief description of ART and the algorithm of Spence et a l . is provided in Appendix A. Because of sonobuoy dri f t uncertainties and non-coincidence of reversed lines, i t is not possible to obtain a single model which satisfies the data for the forward and reverse lines, and thus final velocity models are constructed by combining features from the separate models for the forward and reverse lines. Where the models overlap and do not completely agree, either a compromise between the structures of the two models is included, or the structure from the model which is best constrained at that point is included. Generally, in these instances, the structure is taken from that model in which this segment of the model is closest to the start of the profile since the effects of sonobuoy drift should be reduced near the start of the profile.. 3.2.2 Travel Time Modelling Travel time modelling involves construction of a velocity model for which calculated travel times of rays propagating through the model match the travel times picked for the observed data. Deciding how well to f i t the travel times, and what features of the travel time curve should be reproduced by the model, is a subjective matter. In this kind of seismic analysis, the data provide information about the bulk or average properties of the shallow crust in the study area. Although there may be velocity variations on a small scale (relative to 47 the path length) along the travel path through a particular rock unit, information is obtained only about the average seismic velocity of this unit. Thus, in modelling, the objective is to produce a simplified velocity structure which represents the general velocity characteristics of the crust. To this end, several assumptions and guiding objectives have been used in this part of the model construction. Velocity models have been constructed in terms of layered structures with linear vertical velocity gradients. Vertical gradients are included — rather than homogeneous layers — to produce turning rays because of the amplitude characteristics of the data. Cerveny and Ravindra (1971, p. 147) state that the amplitude of the pure head wave decreases as 1/r2 at large distances from the source. Braile and Smith (1975) show on several synthetic seismogram sections that the amplitude of the pure head wave is at least one order of magnitude smaller than the amplitude of wide angle reflections at distances of 2-3 times the c r i t i c a l distance. Both of these properties of pure head waves are contradictory to i n i t i a l observations made on the data set. For instance, on both the SB3R and SB6R seismograms, where f i r s t arrivals are observed to 3-4 times the c r i t i c a l distance, amplitudes near the ends of these profiles are not appreciably smaller than at points close to the c r i t i c a l distance, suggesting a 1/r dependence of amplitude on distance. Also, the secondary arrivals on the SB3R seismogram between 12 and 16 km, which are likely wide angle reflections, have amplitudes comparable in size to the f i r s t arrivals. These characteristics imply that the observed f i r s t arrivals are 48 better modelled as turning rays. Cerveny (1966) has shown that even small velocity gradients produce interference head waves with amplitudes 1-2 orders of magnitude greater than the pure head wave, and the simple turning ray is the principal component of the interference head wave at large distances. The slight curvature in the arrival curves due to velocity gradients are not likely to be observed due to uncertainties in the travel time picks. Generally, an attempt is made to f i t only those features of the observed travel time curve which are due to large scale structure in the velocity model. Small perturbative features in the observed travel times, due to localized rock structure, are considered in the velocity modelling only i f they can be attributed to structure indicated on the reflection profiles. For reversed profiles, the same velocities and velocity gradients are used for corresponding layers of the forward and reverse profiles. 3.2.3 Amplitude Modelling The procedure used to compensate the amplitudes of the observed data (section 2.3.3) is only approximate. Obviously, there will be trace to trace variations in amplitude due completely to the simple-mindedness of this correction procedure; such variations are not representative of the true amplitude characteristics. However, the general 'group' amplitude characteristics ( i . e. for a group of traces) should demonstrate true amplitude variations of the observed seismogram. There are also trace to trace amplitude fluctuations, which are genuine, due mostly to local focusing, 49 but such small scale variations are not accounted for in the velocity model. Considering these factors, an attempt is made to reproduce only the 'group' amplitude signatures of the observed data when generating synthetic seismograms for the velocity models. Also, because of the uncertainty in the amplitude correction procedure, the fitt i n g of the travel times is emphasized in the trade-off between fitting travel times and amplitudes. 3.2.4 Bottom Topography and Reflection Profile Constraints As mentioned previously, the topography of the water-bottom interface is provided by depth sounding records. Because time marks were annotated roughly every 15 minutes on the original records, there is some uncertainty (probably ±1 or ±2 shot spacings) in the positions of topographic features indicated on these records. Allowance for this uncertainty is made in the modelling procedure; the position of a topographic feature which results in the best f i t t i n g of the data is chosen, within the uncertainty limits. The reflection profiles used to constrain the shallow structure do not completely coincide in position with the refraction profiles. Thus, although the shallow structure given by the reflection profiles is used in the velocity models, a certain degree of freedom is allowed in determining depths and position of topographic features on interfaces. Positions (relative to the refraction profiles) on the reflection profiles are estimated by correlating the water-bottom topography of the reflection and refraction profiles. 50 3.2.5 Sonobuoy Drift: An Uncontrolled Variable As mentioned in section 2.3.2, the relative shot-receiver distances can be calculated to within ±0.1 km. However, the absolute position of the sonobuoy is unknown, and for the case of a two-dimensional medium, the absolute position of the receiver is important. To obtain an estimate of the magnitude of sonobuoy d r i f t , a comparison of profile lengths calculated using the range-bearing method and the water wave method (see section 2.3.2) is made for each of the profiles. The results are tabulated in Table 3.1. Profile Length Difference (km) SB3 0.8±0.7 SB3R 0.4±0.7 SB4 5.0±0.7 SB4R 0.9±0.7 SB5 1.7±0.7 SB5R 3.6±0.7 SB6 1.4±0.7 SB6R 0.9±0.7 T a b l e 3.1 Sonobuoy d r i f t . P r o f i l e l e n g t h s a re c a l c u l a t e d u s i n g the r a n g e - b e a r i n g and water wave methods ( d i s c u s s e d in s e c t i o n 2.3.2). and the l e n g t h d i f f e r e n c e s between the two methods are l i s t e d i n the above t a b l e . T h i s d i f f e r e n c e p r o v i d e s an e s t i m a t e of the minimum sonobuoy d r i f t d i s t a n c e s , f o r each of the p r o f i l e s . The results show differences in length of about 1.0 km for most profiles, but differences range from 0.4-5.0 km. These distances directly represent drift distances if the drift occurs parallel 51 to the profile length. If drift is perpendicular to the profile length, a much larger dr i f t distance is required to account for these differences in profile length. The importance of 1 km of sonobuoy drift, with respect to a profile length of 20 km is small in most cases. 52 Chapter $ Detailed Interpretation 4.1 The Northern Lines; SB6/SB6R/R6 This profile was chosen as the f i r s t for interpretation since i t is a reversed profile, f i r s t arrivals can be picked almost over the entire length of both profiles, sonobuoy dri f t appears to be small, and the R6 reflection profile corresponds very closely in position with the refraction lines. The i n i t i a l observations made on the SB6/SB6R data set suggest a two-layer structure (excluding the shallow unconsolidated sediments) as a starting model: a layer of intermediate velocity material (3.6 km/s), and a higher velocity (near 6.0 km/s) medium below, dipping from the mainland toward Vancouver Island. Also, the bumps, in the f i r s t arrival curve suggest the presence of topographic effects. The R6 reflection profile (Figure 3.2) provides information concerning the depth of the shallow sediments. The major topographic features along this line — as given by T i f f i n (1969) — are labelled on this profile. Its interpretation is based essentially on the interpretation of similar profiles by T i f f i n (1969). Shallow ponds (<200 m) of unconsolidated sediments are seen in Ballenas Basin and the other smaller troughs along the profile. Based on correlation with Pleistocene ridges on the mainland, T i f f i n suggests that the sequence of well stratified reflectors of McCall Ridge represent Pleistocene sedimentary deposits of glacial t i l l s , d rifts, and outwash. 53 Acoustic basement is readily identified from Vancouver Island to McCall Ridge, although i t is uncertain what happens beneath McCall Ridge. T i f f i n discusses the differing stratification characteristics of the bedrock across the Strait. Stratification within bedrock terminates just offshore along the mainland, and the bedrock of the mainland slope is unstratified. T i f f i n proposes that although no clear contact between these units is visible on the profiles, the stratified bedrock probably overlies the crystalline rocks of the mainland slope. He suggests, further, that an unconformable contact between these two units likely exists beneath Sechelt Basin. T i f f i n also indicates that the stratified bedrock on the mainland side of Georgia Strait is seismically similar to that of the Island Slope. On this basis, acoustic basement is considered, here, to represent a single morphologic unit from the Island Slope, to McCall Ridge. To the northeast of this, bedrock on the R6 profile presumably represents the extension of the Coast Range intrusives from the mainland. 4.1.1 Travel Time Modelling The preliminary models obtained for SB6 and SB6R by travel time modelling, using the ray trace algorithm, are shown in Figures 4.1 and 4.2. The following is a description of how these models were constructed. No velocity information concerning the unconsolidated sediments or Pleistocene sediments is obtained from the refraction data, so a velocity of 1.6 km/s is assumed in subsequent velocity modelling for these units. This velocity is considered to be representative of these sediments (McKee, 54 F i g u r e 4.1 A/ I n i t i a l v e l o c i t y model f o r SB6 p r o f i l e . The v e l o c i t y (km/s) and v e l o c i t y g r a d i e n t (km/s/km) f o r each l a y e r a re shown on the l e f t and r i g h t , r e s p e c t i v e l y . D i s t a n c e i s - measured from the SB6 sonobuoy l o c a t i o n . The top l a y e r i s sea water . B/ Rays t r a c e d th rough the v e l o c i t y mode l . On ly t u r n i n g rays in Layer 2 a re i n c l u d e d . C/ Obse rved and c a l c u l a t e d ( f rom ray t r a c i n g ) t r a v e l t imes a re p i o t t e d . 55 SB6R: M O D E L A DISTANCE (Km 8 10 12 14 16 L_ 18 i_ 20 co-4.0: .55 1.6 6.1: .25 20 B <-> CZ>" LU . co "~ co' \ O LO") + + o b s e r v e d -c a l c u l a t e d -8 10 12 DISTANCE (KM) 14 16 ~18~ 20 F i g u r e 4.2 A/ I n i t i a l v e l o c i t y model f o r SB6R p r o f i l e . The v e l o c i t y (km/s) and v e l o c i t y g r a d i e n t (km/s/km) f o r each l a y e r a re shown. D i s t a n c e i s measured from the SB6R sonobuoy l o c a t i o n . The top l a y e r i s sea water. B/ Rays t r a c e d through the v e l o c i t y model. T u r n i n g rays i n L a y e r s 1 and 2 a r e i n c l u d e d as well as r e f l e c t i o n s from the i n t e r f a c e between these l a y e r s . C/ Observed and c a l c u l a t e d t r a v e l times are p1ot ted . 56 1966; Texaco, 1969). The P i arrivals for the SB6R refraction data have an apparent velocity of 3.37 km/s. From depth sounding records, the Island Slope descends at an average angle of 5°±1°, for the f i r s t few kilometres along the SB6R profile. Assuming that the p, arrivals are head wave arrivals from this interface, the real velocity of this unit — referred to from now on as Layer 1 — is 4.2±0.2 km/s (the uncertainty in the velocity is due to the uncertainty in the slope of the interface). Travel times from ray tracing, which model the p, arrivals as head waves from the bedrock with velocity 4.2 km/s, compare well with the observed pt travel times for SB6R. But because of the amplitude characteristics of the data (as discussed in section 3.2.2), i t seems more reasonable to model the arrivals from Layer 1 as turning rays, rather than pure head waves. To do this, a velocity gradient of 0.55 km/s/km is assigned within this unit. This velocity gradient is chosen because the high amplitude secondary arrivals on the SB6R seismogram are only observed out to 8 km distance. This gradient produces turning rays in Layer 1 such that the ray which just grazes the bottom of this layer reaches a maximum distance close to 8 km. A l l deeper angle rays (which do not penetrate deeper than Layer 1) are reflected from the bottom of this layer, arriving at the surface at distances less than that of the grazing ray. Also, the slope of the interface is taken to be 4.3° ( s t i l l within the uncertainty limits) which reduces the velocity of Layer 1 to 4.0 km/s, at the surface. The average velocity of the p 2 arrivals on the SB6R record section is 6.5 km/s. On the SB6 record section, the average 57 velocity of the p 2 arrivals is 5.8 km/s. Assuming that these arrivals are from the same velocity unit, henceforth referred to as Layer 2, this suggests that this unit is dipping toward Vancouver Island, and that the true velocity is somewhere between these two apparent velocities. Taking the real velocity to be 6.1 km/s, and modelling the p 2 arrivals as turning rays in Layer 2, a dip of 3.2° toward Vancouver Island produces ray tracing travel times which satisfy the observed p 2 arrival times for SB6R. A velocity gradient of 0.25 km/s/km has been arbitrarily assigned within this layer, although this seems somewhat large for a material with a high velocity. The early arrivals in the p 2 arrival curves for SB6R and SB6 are attributed to Sangster Ridge and Round Ridge, which are topographic features of Layer 1. Because p, arrivals are not observed on the SB6 seismogram, and since the stratified bedrock (associated with Layer 1 on the reflection profile) is thought to terminate off the mainland, Layer 1 has been terminated in the model at a short distance from the beginning of the SB6 profile. This corresponds to the vicinity of McCall Ridge where acoustic basement is not clear on the R6 profile. Using essentially the same shallow structure as was used for the SB6R model, and using 6.1 km/s as the velocity for Layer 2 (with gradient 0.25 km/s/km), a dip of 3.0° toward Vancouver Island satisfies the observed p 2 travel times between 3 and 19 km, for SB6. The surface of this layer is taken to be horizontal from 0 to 3 km at a depth of 0.575 km. The dip of 3.0° compares well with that obtained for the SB6R model. 58 F i g u r e 4.3 A / R e v i s e d v e l o c i t y model f o r SB6 p r o f i l e . The v e l o c i t y (km/s) and v e l o c i t y g r a d i e n t (km/s/km) f o r each l a y e r are shown. D i s t a n c e i s measured from the SB6 sonobuoy l o c a t i o n . The t h i n t r a n s i t i o n l a y e r at the top of Layer 2, i s i n c l u d e d to s a t i s f y a m p l i t u d e c h a r a c t e r i s t i c s of the observed seismogram. The top l a y e r i s sea water. B/ Rays t r a c e d through the v e l o c i t y model. Only t u r n i n g r a y s i n Layer 2 are i n c l u d e d . C/ Observed and c a l c u l a t e d (from ray t r a c i n g ) t r a v e l times are p l o t t e d . 59 SB6R: MODEL B DISTANCE (KM) 8 10 12 14 i _ 16 L_ 20 LU o r" 1.6 4.0: .55 6.1: .5 6.2: .1 8 10 12 14 16 18 20 NE B LO LU • LO ^~ CD \ a L o _ i o b s e r v e d - » » » » c a l c u l a t e d - + + + + I 10 12 DISTANCE (KM) tt ~2Q F i g u r e 4 . 4 A/ R e v i s e d v e l o c i t y model f o r S B 6 R p r o f i l e . The v e l o c i t y (km/s) and v e l o c i t y g r a d i e n t (km/s/km) f o r each l a y e r a r e shown. D i s t a n c e i s measured from the SBGR sonobuoy l o c a t i o n . The top l a y e r i s sea water. B/ Rays t r a c e d through the v e l o c i t y model. T u r n i n g rays i n La y e r s 1 and 2 a r e i n c l u d e d as well as r e f l e c t i o n s from the i n t e r f a c e between these l a y e r s . C/ Observed and c a l c u l a t e d t r a v e l times are p i o t t e d . LO d O CNI O LU i n c n o i LO 4 6 8 10 12 D I S T A N C E (KM) S B 6 14 16 18 8 10 12 D I S T A N C E (KM) A . I f 8 10 12 14 D I S T A N C E (KM) 16 18 F i g u r e 4.5 A/ Observed SB6 seismogram. B/ S y n t h e t i c seismogram f o r model SB6: A. C/ S y n t h e t i c seismogram f o r model SB6: B. I n s e t s show ampl i t u d e v e r s u s d i s t a n c e f o r : 1=turning ray in Laver 2. Note the l o g a r i t m i c s c a l e . F i r s t water-bottom m u l t i p l e t u r n i n g r a y i n Layer 2 i s i n c l u d e d i n s y n t h e t i c s . 61 C N CM G LU CO LO CD CO i— LO-I LO CN CM CJ UJ LO CO —• CO \ a i CD a -* LO-l 8 10 12 14 DISTANCE (KM) SB6R D I S 12 14 flNCE (KM) 16 18 20 B a io 20 10 12 14 16 18 20 DISTANCE (KM) F i g u r e 4.6 A/ Observed SB6R seismogram. B/ S y n t h e t i c seismogram f o r model SB6R: A. C/ S y n t h e t i c seismogram f o r model SB6R: B. Inset s show ampl i t u d e v e r s u s d i s t a n c e f o r : 1=turning ray in Layer 2, 2=turning ray i n Layer 1, 3 = r e f l e c t e d ray from bottom of Layer 1. Note a l s o , f i r s t water bottom m u l t i p l e t u r n i n g ray i n Layer 2 has been i n c l u d e d i n s y n t h e t i c s . 62 4.1.2 Amplitude Modelling The synthetic seismograms computed for the i n i t i a l travel time models of SB6 and SB6R (models SB6: A and SB6R: A) are shown along with the actual seismograms in Figures 4.5 and 4.6. Like the SB6R seismogram, the SB6R synthetic seismogram has large amplitude secondary arrivals out to 8 km distance. The synthetic data in this region is complex even though only four types of rays are included in the calculation: turning rays in Layer 1, reflected rays from the bottom of Layer 1, the f i r s t water-bottom multiple turning rays in Layer 1, and turning rays in Layer 2. The complexity of the real data in this region, which obscures any secondary travel time branches, is not surprising. Reverberations in the sediments of Layer 1 probably add to this complex pattern. Unlike the real data for SB6R, the p 2 arrivals on the synthetic seismogram show a general increase in amplitude from 5 to 20 km distance. On the real seismogram, the p 2 amplitudes are reduced between 7 and 12 km distance. Also, the amplitudes of the p 2 arrivals on the real data section are comparable in size to the secondary arrival amplitudes between 4 and 7 km, while the secondary arrivals are much larger than the p 2 arrivals for the synthetic data. The amplitudes of the p 2 arrivals on the SB6 seismogram remain relatively constant over the length of the profile, while on the SB6 synthetic section (Figure 4.5 b), the p 2 arrival amplitudes increase with distance. To increase the p 2 amplitudes at shorter distances relative to the amplitudes at greater distances, the velocity gradient in Layer 2 is increased to 0.5 km/s/km. This gradient, though, has been restricted to the upper 63 0.2 km of this layer for two reasons. First, a compressional velocity of 6.1 km/s is representative of granitic type rocks which are unlikely to have a high velocity gradient. Second, the models proposed by Tseng (1968) for Georgia Strait have a velocity of 6.6 km/s at depths of 4.3-6.0 km. With a velocity gradient of 0.5 km/s/km for a l l of Layer 2, a velocity of 6.6 km/s would be reached as shallow as 1.6 km in this part of the Strait. Thus, an a r t i f i c i a l boundary at 0.2 km below the surface of Layer 2 has been included, beneath which the velocity gradient decreases to 0.1 km/s/km. The velocity in the revised models (models SB6: B and SB6R: B) reaches 6.6 km/s at a minimum depth of 4.8 km, which is more consistent with Tseng's model. The effect of this slight change in the model has a negligible effect on the computed travel times. The revised models and computed travel times for SB6 and SB6R are illustrated in Figures 4.3 and 4.4. The synthetic seismograms for these models are shown in Figures 4.5 and 4.6. As can be seen, the p 2 amplitudes of the model SB6: B synthetic seismogram remain relatively constant over , the length of the profile as desired. The small amplitudes of the p 2 arrivals after 7 km on the SB6R seismogram have been reproduced on the model SB6R: B synthetic seismogram. This decrease in amplitude is a result of the decrease in velocity gradient from 0.5 km/s/km to 0.1 km/s/km in Layer 2. Also, on the SB6R: B synthetic seismogram the p 2 amplitudes between 5 and 7 km are more comparable in size to the secondary arrivals. 4.2 The Central Lines; SB3/SB3R/R3 The main topographic features on the R3 profile (Figure 64 3.2) are labelled following T i f f i n (1969). Ponds of unconsolidated sediments (<200 m in thickness) are found in Ballenas Basin and Sechelt Basin. According to T i f f i n , Halibut Ridge and McCall Ridge represent Pleistocene deposits, several hundred metres thick in some places. A velocity of 1.6 km/s is assigned to these units. From the Island Slope, acoustic basement can be followed beneath Ballenas Basin and Halibut Ridge, but becomes uncertain beneath McCall Ridge. The acoustic basement interface from the Island Slope to the east side of Halibut Ridge is taken to represent a single morphologic unit. After the descent of the Island Slope, this unit levels out at a depth of approximately 0.6 km. 4.2.1.Travel Time Modelling The models for profiles SB3 and SB3R, along with observed and calculated travel times, are shown in Figures 4.7 and 4.8. I n i t i a l observations made on the SB3/SB3R data set suggest a two layer starting model, with the deeper higher velocity layer dipping from the mainland to Vancouver Island. The p,^ and p 1 2 arrivals on the SB3R seismogram have apparent velocities of 2.32 km/s and 3.68 km/s, respectively. This suggests the existence of two distinct layers. However, the R3 reflection profile shows that the Island Slope is very steep beneath the f i r s t few shot positions of SB3R, attaining a shallower angle of descent as i t slopes under the sediments of Ballenas Basin. This indicates that the p!, and p 1 2 arrivals are likely refracted by the same velocity unit (Layer 1), and that the difference in apparent velocities is due to the change in slope of the surface of this layer. From depth sounding data for SB3R, the steep part of the SB3 M O D E L F i g u r e 4 . 7 A/ V e l o c i t y model f o r SB3 p r o f i l e . The v e l o c i t y (km/s) and v e l o c i t y g r a d i e n t (km/s/km) f o r each l a y e r a re shown on the l e f t and r i g h t , r e s p e c t i v e l y . D i s t a n c e i s measured from the SB3 sonobuoy l o c a t i o n . The top l a y e r i s sea water . B/ Rays t r a c e d th rough the v e l o c i t y mode l . On ly t u r n i n g rays in Layer 2 a re i n c l u d e d . C/ Observer! and c a l c u l a t e d ( f rom ray t r a c i n g ) t r a v e l t imes a re p l o t t e d . 66 F i g u r e 4.8 A/ V e l o c i t y model f o r SB3R p r o f i l e . The v e l o c i t y ( k m / s ) a n d v e l o c i t y g r a d i e n t (km/s/km) f o r e a c h l a y e r a r e shown. D i s t a n c e i s m e a s u r e d f r o m t h e SB3R s o n o b u o y l o c a t i o n . The t o p l a y e r i s s e a w a t e r . B/ R a y s t r a c e d t h r o u g h t h e v e l o c i t y m o d e l . T u r n i n g r a y s i n L a y e r s 1 a n d 2 a r e i n c l u d e d a s w e l l a s r e f l e c t i o n s f r o m t h e i n t e r f a c e b e t w e e n t h e s e l a y e r s . C/ O b s e r v e d a n d c a l c u l a t e d t r a v e l t i m e s a r e p l o t t e d . 67 Island Slope descends at an angle of 14°±3° from horizontal. As it slopes beneath the sediments of Ballenas Basin, the average slope (from the R3 reflection profile) is 5°±2°. Using slopes within these limits in the ray tracing model, a velocity of 3.9 km/s with a velocity gradient of 0.2 km/s/km satisfies the observed p,^ and p 1 2 arrival times. The f i t t i n g of the p,, and p 1 2 arrivals provides constraint on the surface of Layer 1 to a distance of 6.5 km along profile SB3R. For the remainder of the profile, this surface is constrained by the acoustic basement on R3, and the fact that p, arrivals are not observed as f i r s t arrivals on the SB3 record section. Also, matching the p 3 arrivals (for the SB3R section) as wide angle reflections and turning rays in Layer 1 to 16 km distance provides constraint on the surface of Layer 1. The R3 profile indicates that the acoustic basement slopes upward from Ballenas Basin to a depth near 0.5 km beneath Halibut Ridge. To the east of Halibut Ridge, the identification of acoustic basement is not as certain, and i t is not clear whether that which has been identified east of Halibut Ridge represents the same morphologic unit. Because of this, the surface of Layer 1 has been continued to McCall Ridge at a depth of 0.5 km. As can be seen in Figure 4.8, the calculated travel times match the observed p 3 travel times quite well. Some constraint on the velocity gradient in Layer 1 is provided by the requirement that turning rays in Layer 1 be observed at least out to the distance of the observed p 3 arrivals. On the SB3R seismogram (Figure 3.3), the p 2 arrivals have an apparent velocity of 6.0 km/s between 7 and 14 km, after 68 which the apparent velocity increases to 8.7 km/s out to 20 km distance. This extremely high apparent velocity is obviously due to dip. On the SB3 seismogram, the p 2 arrivals have an apparent velocity of very close to 6.0 km/s out to 3 km. After this, the average apparent velocity of these arrivals is generally less than 6.0 km/s out to a distance of 16 km, where the velocity is again close to 6.0 km/s. Also, no arrivals ( i . e. intermediate velocity arrivals) are observed as fi r s t arrivals on the SB3 record section, indicating that Layer 1 is either absent or very thin toward the mainland side of the Strait. There are no secondary arrivals observed that would indicate the presence of Layer 1, but they may be obscured. In any event, Layer 1 has been squeezed out in the SB3 and SB3R models a short distance from the beginning of the SB3 profile. A horizontal layer (Layer 2) with velocity 6.0 km/s at the surface of the layer, and at a depth of 2.3 km, is incorporated into the model for SB3R. This layer continues out to 12.5 km from the start of SB3R, and travel times for turning rays in this layer satisfy the p 2 arrivals out to 14 km. From 12.5 km distance, the surface of Layer 2 slopes up toward the mainland at an angle of 12.9°. This slope is responsible for the high apparent velocity of the SB3R p 2 arrivals, after 14 km. This interface reaches a depth of 0.35 km at 21 km distance where i t flattens out as required by the SB3 model. I n i t i a l l y , a velocity gradient of 0.25 km/s/km was assigned to this layer. The flat-lying bedrock on the mainland side of R3 may represent the surface of Layer 2. To f i t the p 2 arrivals out to 3 km distance, for the SB3 data, using a velocity of 6.0 km/s 69 for Layer 2, a horizontal interface at a depth of 0.35 km is required. To f i t the p 2 arrivals between 3 and 8 km on the SB3 seismogram, the ^ surface of Layer 2 slopes down toward Vancouver Island at an angle of 16°. After 7.5 km, the slope is reduced to 10.7° until 12 km distance where Layer 2 becomes horizontal at a depth of 2.35 km. The same velocity gradient for Layer 2 is used for SB3 and SB3R. 4.2.2 Amplitude Modelling The same 'transition layer' at the surface of Layer 2 that is included in the SB6/SB6R models, is added to the travel time models constructed for SB3 and SB3R. The effect of this layer on the travel times is negligible. It has been added for the same reasons as in the case of SB6 and SB6R. The synthetic seismograms obtained for the SB3 and SB3R models of Figures 4.7 and 4.8, are shown along with the observed seismograms in Figures 4.9 and 4.10. The main amplitude characteristic of the SB3 data is the drop in amplitude of the f i r s t arrivals after 14 km distance. On the synthetic seismogram, there is only a small decrease in amplitude after 14.5 km with a larger decrease occurring after 16 km. This decrease in amplitude is caused by the slope of the Layer 1-Layer 2 interface starting at 12 km in the SB3 model. The lesser decrease in amplitude at 14 km on the synthetic section as compared to the decrease at 16 km, is due to a 'channelling' effect of the 'transition zone' in Layer 2, in the vicinity of the corner at 12 km. This is likely an artifact of the model. Because the algorithm used to produce the synthetic seismograms is based on the zero-order ART solution, no 0 2 4 6 8 10 12 14 16 18 20 22 24 DISTANCE (KM) SB3 F i g u r e 4.9 A/ Observed SB3 seismogram. SB3 model. Inset shows amplitude (on d i s t a n c e f o r : 1=turning ray i n Layer 2. B/ S y n t h e t i c 1ogar i thmic seismogram f o r s c a l e ) v ersus 71 0 2 4 6 8 • 10 12 14 16 18 20 22 24 DISTANCE (KM) F i g u r e 4.10 A / Observed SB3R seismogram. B/ S y n t h e t i c SB3R seismogram. Inset shows amplitude (on l o g a r i t h m i c s c a l e ) v ersus d i s t a n c e f o r : 1= t u r n i n g ray i n Layer 2, 2=turning i n Layer 1, 3 = r e f l e c t i o n from the bottom of Layer 1. 72 diffraction effects are included which may be caused by the corner. The p 2 arrivals on the SB3R synthetic seismogram show only small variations between 8 and 22 km, as is the case for the observed SB3R data. However, the relative amplitudes of the p 2 arrivals, as compared to the secondary arrival amplitudes, on the synthetic section are small, whereas the relative sizes of these arrivals are not as disproportionate in the observed data. Attempts to correct this by adjusting velocity gradients leads to unacceptably large velocity gradients in Layer 2. Three suggestions may be considered to explain this discrepancy. First, the synthetic seismogram algorithm does not take into account attenuation. Layer 1, having velocities characteristic of sandstone or shale, is likely to have a much lower Q, value than Layer 2, which has velocities characteristic of granitic type rocks. Although both p, and p 2 arrival.s travel through Layer 1, most p, arrivals have a proportionately longer travel path through Layer 1, and thus experience greater attenuation. Secondly, stratified reflectors within the sedimentary Layer 1 will produce reverberations within the sediments that are not modelled. This effect will tend to spread out the energy in the seismic wavelet, reducing the amplitude of the i n i t i a l arrival, and producing a longer train of arrivals following i t . This is likely the cause of the complicated pattern of secondary arrivals observed on the SB6R, SB5, and SB3R record sections, which is not observed on the reverse lines where the sediments are thin or absent. This effect is more significant for the p, arrivals which have a longer travel path in the sedimentary 73 Layer 1. The third alternative is that the f i r s t order discontinuity in the model between Layers 1 and 2 is actually a second order discontinuity. This final alternative is least favoured, since (as will be seen) the R5 reflection profile provides evidence for the f i r s t orcler interface between Layers 1 and 2. The region of small amplitudes in the p, arrivals between 4 and 6 km on the real SB3R seismogram is likely due to a local feature in the shallow topography not observed on R3, and thus it is not included in the model. 4.3 The Southern Lines; SB5/SB5R/R5 Although the f i r s t 0.4 s of data are muted on the R5 reflection profile (Figure 3.6), acoustic basement can be identified over the length of the upper section. It ranges in depth from 0.35 km to 0.75 km, sloping away from Vancouver Island. Material above i t , but below the water-bottom surface, has been assigned a velocity of 1.6 km/s. To the right of the point where the acoustic basement fades out, marked on the R5 section by the 'F', there is a thick sequence of flat-lying sediments. These are interrupted by the Fraser Ridge (on the left side of the lower section) which is a Pleistocene landform of glacial sediments, identified by the stratified and chaotic nature of the internal reflectors. Acoustic basement cannot be identified in this region. On the right half of the lower section of R5, acoustic basement is identified near 0.9 s, and a second reflector is observed near 1.5 s at the left of this part of the profile. This second reflector slopes upward to 1.0 s at the right side of the profile. These two reflectors converge on 74 the right side of the profile, squeezing out the upper sequence represented by acoustic basement. The depth to acoustic basement varies from 0.6 to 0.7 km. Whether or not acoustic basement on this side of the profile represents the same rock unit as on the upper part of the profile is unknown. The unconsolidated sediments overlying acoustic basement are the thick sediments of the Fraser Delta. Using a velocity of 1.6 km/s for these sediments, they are up to 0.13 km thick in this region, and are thicker to the southwest. 4.3.1 Travel Time Modelling The velocity models for SB5 and SB5R along with observed and calculated travel times are shown in Figures 4.11 and 4.12. The apparent velocity of the p, arrivals on the SB5 seismogram is 3.2 km/s. Using the bedrock surface obtained from R5 for Layer 1, a real velocity of 3.6 km/s is found to f i t the observed travel times out to 5 km. A velocity gradient of 0.25 km/s/km has been used in Layer 1. To satisfy the p 2 arrival times out.to 8 km distance, with a velocity of 6.1 km/s at the upper surface, Layer 2 is located at a depth of 1.8 km at the start of SB5, and i t dips toward Vancouver Island. The offset in f i r s t arrivals observed on the SB5 seismogram is modelled as a vertical fault, since the dip of the fault is not well constrained by the data. Because the identification of acoustic basement on R5 becomes uncertain after the position corresponding to the proposed fault, i t is assumed that the thickness of Layer 1 is the same on either side of the fault. Also, the interface of Layer 1 and the overlying unconsolidated sediments is assumed to be horizontal on the down-dropped side 75 SB5 MODEL DISTANCE (KM) in CJ LU CO CO o i o b s e r v e d - » • — c a l c u l a t e d - * * • 8 10 12 14 DISTANCE (KM) i r 16 18 20 22 F i g u r e 4 .11 A/ V e l o c i t y model f o r SB5 p r o f i l e . The v e l o c i t y (km/s) and v e l o c i t y g r a d i e n t (km/s/km) f o r each l a y e r are shown. D i s t a n c e i s measured from the SB5 sonobuoy l o c a t i o n . The proposed f a u l t i s l o c a t e d at 7.5 km. The top l a y e r i s sea water. B/ Rays t r a c e d through the v e l o c i t y model. T u r n i n g rays i n Layers 1 and 2 are i n c l u d e d as well as r e f l e c t i o n s from the i n t e r f a c e between these l a y e r s . C/ Observed and c a l c u l a t e d t r a v e l times are p l o t t e d . 76 SB5R MODEL DISTANCE (KM) A 0 2 4 6 8 10 12 14 16 18 20 -I 1 1 1 1 1 i i i i i 0 2 4 6 8 10 12 14 16 18 20 DISTANCE (KM) F i g u r e 4.12 A/ V e l o c i t y model f o r SB5R p r o f i l e . The v e l o c i t y (km/s) and v e l o c i t y g r a d i e n t (km/s/km) f o r each l a y e r are shown on the l e f t and r i g h t , r e s p e c t i v e l y . D i s t a n c e i s measured from the SB5R sonobuoy l o c a t i o n . The top l a y e r i s sea water. B/ Rays t r a c e d through the v e l o c i t y model. T u r n i n g rays i n Layers 1 and 2 are i n c l u d e d , as well as r e f l e c t i o n s from the bottom of Layer 1. C/ Observed and c a l c u l a t e d (from ray t r a c i n g ) t r a v e l times are p l o t t e d . 77. of the fault. Under these assumptions, a fault with throw of 0.55 km f i t s the observed offset travel times. The model for SB5 after 11 km is not based on observed travel times, but continues Layer 1 and Layer 2 to the depths indicated by the SB5R model, on the mainland side of the Strait. The thin transition zone has been included in Layer 2, where the velocity gradient is 0.5 km/s/km, to be consistent with the other models. Beneath this thin zone, the velocity gradient is reduced to 0.15 km/s/km. There are no p, arrivals observed as f i r s t arrivals on the SB5R seismogram (Figure 3.4) to suggest the presence of Layer 1 on this side of the Strait, but the secondary arrivals observed out to almost 6 km, having an apparent velocity of 3.6 km/s, suggest the presence of at least a thin layer of intermediate velocity material. Using the acoustic basement identified on the lower right side of profile R5 (Figure 3.6), and a velocity of 3.6 km/s (with gradient 0.25 km/s/km), travel times for turning rays in this layer match the observed travel times. The acoustic basement on R5 provides the depth to the top of Layer 1 out to a distance of 5 km. From this point, the surface of this layer is arbitrarily continued down to the depth of Layer 1 for the SB5 model, at the fault location. Using a velocity of 6.1 km/s at the upper surface of Layer 2, a depth of 0.9 km is required to match the p 2 arrivals out to 3.5 km. The interface between Layer 1 and Layer 2 then dips toward Vancouver Island, at an angle of 14° from 3 to 5.5 km, and then at a shallower angle of about 2° out to 19 km distance, where i t reaches a depth of 1.94 km. These slopes are chosen to f i t the travel times using the velocity of 6.1 km/s for Layer 2. 78 Using the velocities obtained from the refraction data, the depth to the second reflector on the R5 profile is approximately 1.1 km at the start of SB5R and 1.6 km at a distance of 5 km from the start of SB5R. These compare well with depths of 0.9 and 1.5 km at the corresponding distances in the velocity model. The thin transition zone with velocity gradient of 0.5 km/s/km, has been included in both models. Below i t the velocity gradient is reduced to 0.15 km/s/km. The transition layer for the SB5 model is 0.3 km thick as compared to 0.2 km for the other models. This variation is included to eliminate amplitude variations caused by this layer which are considered a r t i f i c i a l . Using the distance to the proposed fault indicated by the SB5 model, the effects of the fault would be expected to appear on the SB5R seismogram at a distance of 14 -km. But arrivals are observed before and after 14 km with no apparent offset in arrival times. When the.length of profile SB5R is calculated using the range-bearing and water wave methods (see section 3.2.5), a difference of 3.6±0.7 km is found. When the same calculation is made .for profile SB5, a length difference of 1.7+0.7 km is found. The fact that the effect of the fault is not observed at the expected position is due mostly to sonobuoy d r i f t . It is also due, in part, to the non-coincidence of the forward and reverse lines. There are strong arrivals observed on the SB5R seismogram to near 19 km beyond which they are no longer observed. This truncation of the f i r s t arrivals suggests that this may be the position of the fault on this line. Comparison of the shot numbers where the effect of the fault is fi r s t manifested on the SB5R and GAL5R seismograms, indicates 79 that this is probably the case. 4.3.2 Amplitude Modelling The observed and synthetic seismograms for SB5 and SB5R are shown in Figures 4.13 and 4.14. The synthetic section for SB5 has been calculated for the entire length of the velocity model for completeness, but is only valid out to 7.5 km. Obviously, the ART algorithm used to generate the synthetics cannot reproduce any diffraction, attenuation or scattering effects due to the fault zone. What the latter part of the synthetic section does show, are the small amplitudes of the arrivals for the proposed model at greater distances, neglecting these factors. Including"these effects would likely reduce the amount of coherent seismic energy arriving at the surface, even more. Thus, the lack of observed coherent seismic energy after 7.5 km is accounted for by the model. The large, amplitude secondary arrivals observed to 7.5 km on the SB5 seismogram are reproduced by turning rays and reflections in Layer 1. The discussion of section 4.2.2, concerning the disproportionate ratio of f i r s t arrival amplitudes to secondary arrival amplitudes for SB3R, applies to this case as well. The predominant amplitude characteristic of the SB5R seismogram is the region of small amplitude arrivals from 6 to 12 km distance. These have been reproduced on the synthetic seismogram by the bend in the surface of Layer 2 at a distance of 5.5 km. The slight increase in amplitudes between 6 and 9 km on the synthetic section is due to channelling of the rays in the small transition zone included at the top of Layer 2. The p, secondary arrivals are observed out to 5.5 km on the synthetic 80 F i g u r e 4 .13 A/ Observed SB5 se ismogram. B/ S y n t h e t i c SE5 se ismogram. Inse t shows a m p l i t u d e (on l o g a r i t h m i c s c a l e ) v e r sus d i s t a n c e f o r : 1=turn ing ray in Layer 2, 2=turn ing ray in Layer 1 , 3 = r e f l e c t i o n from the bottom of Layer 1. 81 - r — 1 I I i * 1 i t • 11 • i • i — , 0 2 4 6 8 10 12 14 16 18 20 OISTRNCE (KM) F i g u r e 4.14 A/ Observed SB5R seismogram. B/ S y n t h e t i c SB5R seismogram. Inset shows amplitude (on l o g a r i t h m i c s c a l e ) versus d i s t a n c e f o r : 1=turning ray i n Layer 2, 2=turning ray in Layer 1. 3 = r e f l e c t i o n from the bottom of Layer 1. 82 section. The maximum distance at which they are observed is limited by the thickness of Layer 1 at this point. A velocity gradient of 0.25 km/s/km in Layer 1 produces the cut-off in arrivals at the appropriate distance. 4.4 SB4/SB4R Originally, the SB4 and SB4R profiles were intended to form a reversed pair of lines, but arrivals are not observed on either profile over distance segments common to both profiles. Thus, these profiles are considered separately. 4.4.1 SB4: Travel Time Modelling There is no reflection profile that coincides with the SB4 refraction profile. The only control is provided by the point where i t intersects the SB3/SB3R line. Because of this, this profile is modelled in terms of flat-lying interfaces. The resulting model and travel times for SB4 are shown in Figure 4.15. The depth of the water-bottom interface, is provided by the depth sounding records. Turning rays in a layer (Layer 1) 0.67 km deep, with a velocity of 3.9 km/s, satisfy the p, arrivals out to 3 km distance. A velocity gradient of 0.2 km/s/km has been used in this layer, and the velocity assigned to the overlying sediments is 1.6 km/s. Both the velocity and depth of this layer are consistent with those of the SB3R model, at the point where they intersect. A velocity of 6.0 km/s at the surface of Layer 2, produces arrivals fi t t i n g the p 2 arrivals out to a distance of 5 km. The surface of Layer 2 is 1.3 km deep. The observed arrivals define a somewhat lower apparent SB4 MODEL F i g u r e 4.15 A / V e l o c i t y model f o r SB4 p r o f i l e . The v e 1 o c i t y '(km/s) and v e l o c i t y g r a d i e n t (km/s/km) f o r each l a y e r are shown. D i s t a n c e i s measured from the SB4 sonobuoy l o c a t i o n . The top l a y e r i s sea water. B/ Rays t r a c e d through the v e l o c i t y model. T u r n i n g rays in Layers 1 and 2 are i n c l u d e d as well as r e f l e c t i o n s from the i n t e r f a c e between these l a y e r s . C/ Observed and c a l c u l a t e d t r a v e l times are p l o t t e d . 84 velocity between 5 and 10 km, reaching an apparent velocity close to 6 km/s again after 10 km. This change in the arrival times is modelled as being due to a down-sloping of the interface between Layer 1 and Layer 2. This interface slopes from a depth of 1.3 km at 4.5 km distance, to a depth of 2.35 km at a distance of 7.5 km. From this point on, this interface is horizontal. There is support for this feature of the model from the two bordering cross-strait profiles. The depths of Layer 1 and Layer 2 - at the point where SB4 crosses SB3R - on the SB3R model are 0.68 km and 2.3 km, respectively. The depths to these same layers on the SB6R model (where the extension of the SB4 profile would intersect it) are 0.59 and 1.16 km. This suggests that although the depth to Layer 1 in Ballenas Basin between SB6R and SB3R changes l i t t l e , the depth to Layer 2 increases toward SB3R. Such an increase in depth from northwest to southeast is consistent with the SB4 model. The SB4 sonobuoy drifted at least 5.0±0.8 km (the most of any of the sonobuoys) toward the southeast end of the profile, over the course of the line. If the drift was predominantly along the profile, i t would have l i t t l e affect on the seismic results, provided the structure along the Strait is flat-lying as depicted in the velocity model. The only feature of the model that would be uncertain is the location of that part of Layer 2 which slopes downward. If, in fact, the sonobuoy drifted toward Vancouver Island where Layer 2 tends to become deeper, the dipping feature of the SB4 model may be a product of this. The support of the cross-strait lines suggests that this slope is probably real. 85 4.4.1 SB4: Amplitude Modelling The observed and synthetic seismograms for SB4 are shown in Figure 4.16. The sudden drop-off in amplitudes near 10 km, on the original SB4 seismogram, has been reproduced on the synthetic seismogram. This decrease in amplitude is due to the down-sloping interface at a distance of 7.5 km, in the model. This decrease in amplitude further supports the reality of the down-sloping feature of the model. 4.4.3 SB4R: Travel Time Modelling As is the case for the SB4 profile, the SB4R profile is an unreversed refraction profile with no reflection data to supplement i t , although, some control is provided by the SB5 model where i t intersects the SB4R model. Because of this, the SB4R profile is modelled using planar horizontal layers. The SB4R model with calculated travel times is shown in Figure 4.17. Also, due to the similar data signature of SB4R, as compared to SB5, a vertical,fault has been included in the model. Layer 1 at 0.5 km depth, with a velocity of 3.6 km/s, accounts for the p, arrivals observed to 4 km distance. A velocity gradient of 0.25 km/s/km is used in Layer 1. Layer 2, at a depth of 1.4 km and having a velocity of 6.1 km/s at its upper surface satisfies the p 2 arrivals to 5.5 km distance, where they fade out. Where SB4R intersects SB5, the depths to Layer 1 and Layer 2 on the SB5 model are 0.7 and 1.5 km, respectively. These compare reasonably well with the depths for the SB4R model at this point (0.5 and 1.4 km). Also, the velocities and velocity gradients used in both models are the same. 86 DISTANCE (KH) F i g u r e 4 .16 A/ Observed SB4 se ismogram. B/ S y n t h e t i c SB4 se ismogram. Inset shows a m p l i t u d e (on l o g a r i t h m i c s c a l e ) ve r sus d i s t a n c e f o r : 1=turn ing ray in Layer 2. 2=tu rn ing ray in Layer 1. 3 = r e f l e c t i o n from the bottom df Layer 1. SB4R MODEL D I S T A N C E (KM) 8 A 10 E 3.6: .25 r\H 6.1: .5 LU • CO (=3 CO \ a L O ' 1.6 3.6: .25 + + ++4 observed-calculated- + + 2 4 6 D I S T A N C E (KM) 8 10 F i g u r e 4 .17 A/ V e l o c i t y model f o r SB4R p r o f i l e . The v e l o c i t y (km/s) and v e l o c i t y g r a d i e n t (km/s/km) f o r each l a y e r are shown. D i s t a n c e i s measured from the SB4R sonobuoy l o c a t i o n . The proposed f a u l t i s l o c a t e d at a d i s t a n c e of 5.5 km. The top l a y e r i s sea water. B/ Rays t r a c e d through the v e l o c i t y model. T u r n i n g rays in Layers 1 and 2 are i n c l u d e d . C/ Observed and c a l c u l a t e d t r a v e l times are p l o t t e d . 88 A vertical fault is located at 5.5 km distance in the SB4R model. Assuming that Layer 1 has the same thickness and the same velocity on either side of the fault, a throw of 0.55 km satisfies the late p 2 arrivals, agreeing with the throw of the SB5 model. The down-dropped side of the fault is the northeast side. It should be noted that the picks for the late p 2 arrivals, between 8 and 10 km distance, have large uncertainties (±40 ms). No amplitude modelling has been done for this profile due to the small segment of data which provides reliable amplitude information. 4.5 Final Velocity Models The final models for the cross-strait profiles are shown in Figure 4.18. The final models for SB4 and SB4R remain unaltered from those shown in Figures 4.15 and 4.17, since these lines are unreversed. 4.5.1 SB6/SB6R The models SB6: B and SB6R: B (shown in Figures 4.3 and 4.4) are very similar. The only significant difference between these models is the depth of Layer 2 on the Vancouver Island side of Georgia Strait. The difference is less than 0.15 km at its maximum. Because this part of the model is closest to the SB6R sonobuoy position, the effect of sonobuoy dr i f t on the SB6R model in this vicinity should be small. Hence, the depth to Layer 2 on this side of the Strait is taken from the SB6R: B model. This model is only constrained by ray tracing to a depth near 2 km. S B 5 / 5 R MODEL 0 2 4 6 8 10 12 14 16 18 20 C F i g u r e 4.18 F i n a l c r o s s - s t r a i t v e l o c i t y models. V e l o c i t i e s (km/s) and v e l o c i t y g r a d i e n t s (km/s/km) are shown ( l e f t and r i g h t , r e s p e c t i v e l y ) . The d i s t a n c e s in A, B, and C are measured from the sonobuoy p o s i t i o n s of SB6R, SB3R and SB5, r e s p e c t i v e l y . These models combine the models f o r each r e v e r s e d set of p r o f i l e s to o b t a i n a s i n g l e v e l o c i t y model f o r that p a r t of the S t r a i t . Broken l i n e s in models i d e n t i f y s p e c u l a t i v e b o u n d a r i e s . 90 4.5.2 SB3/SB3R The SB3 and SB3R models are again very similar, but there are some small differences. The depth to Layer 2 on the mainland side of the Strait for both models is in agreement, and the depth to Layer 2 on the opposite side of the Strait is within 0.05 km for the two models. The two models differ by less than 1.5 km as to the position where Layer 2 begins dipping toward Vancouver Island. This feature in the final model has been taken from the SB3 model since this segment of the model is closest to the start of SB3. Also, the small bend in the sloping segment of the Layer 1-Layer 2 interface, indicated by the SB3 model, has been included due to its proximity to the start of the SB3 profile, although, i t is a minor feature of the model. The topography of Layer 1 differs slightly for the-SB3 and SB3R models, but the differences are insignificant. The geometry of this interface is taken from the SB3R model for the f i r s t 14.5 km, and from the SB3 model for the remainder of the distance. This interface, between 14.5 and 20 km, is uncertain as i t is not based directly on either, reflection data or refraction arrivals bottoming in this layer. 4.5.3 SB5/SB5R The structure for the f i r s t 11 km of the final SB5/5R model is taken from the SB5 model. The SB5R model indicates that Layer 2 dips toward Vancouver Island, reaching, a depth close to that of the SB5 model on the northeast side of the fault. But, because of the drift distance of the SB5R sonobuoy, the actual dip indicated by the SB5R profile is uncertain. Thus, the structure between 16 and 21 km in the final SB5/5R model is 91 taken from the SB5R model, and the structure of the final model between 11 and 16 km is arbitrarily assigned to match the model at 11 and 16 km distance. This segment of the final SB5/5R model is consistent with the SB5R model in that Layer 2 is dipping toward Vancouver Island, although, at a slightly steeper angle. 92 Chapter 5 Summary and Discussion The maximum depth of penetration obtained in this survey, for profiles of up to 24 km in length, is about 3 km, with some uncertainty due to lack of constraint on the velocity gradients. No true velocities greater than 6.25 km/s are directly observed. This is consistent with the results obtained by White (1962) and Tseng (1968), who obtained velocities in the neighbourhood of 6.5 km/s only for shot-receiver distances greater than 30 km. The model of Tseng (Figure 1.3) shows depths of 4.2 and 6.1 km to the layer of 6.6 km/s for north and south Georgia Strait, respectively. The gradients used for Layer 2, in this study, are such that a velocity of 6.6 km/s is reached at a minimum depth of 3.8 km. The depth at which a.velocity of 6.6 km/s is reached in Tseng's model is used in this study as an additional constraint, since the depth penetration in this study was not sufficient to detect a deeper, higher velocity layer. In the models of both Tseng (1968) and White (1962), there is a f i r s t order discontinuity at the point where a velocity of 6.6 km/s (actually 6.67 km/s in the model of White) is reached. Although this velocity is reached in this model using a constant velocity gradient, this is not to imply that there is no f i r s t order discontinuity; the velocity gradients within Layer 2 are not well enough constrained. But, the amplitude analysis of this study does suggest that a small velocity gradient within Layer 2 93 does exist which would reduce the magnitude of the fi r s t order discontinuity at depth. Since the previous studies of Tseng (1968) and White (1962) included only conventional travel time analyses in terms of homogeneous layers, velocity gradients were - not considered. The velocity models proposed for the shallow crust of the Strait of Georgia are shown in Figure 5.1 in their respective positions in the Strait. Generally, they consist of three layers (excluding the sea water). The f i r s t combines the ponds of unconsolidated sediments and Pleistocene glacial deposits into a single velocity unit with an arbitrarily assigned velocity of 1.6 km/s. This layer is relatively thin along the northern and central lines, but reaches a thickness of at least 500 m where the southern line crosses the Fraser Delta. The second layer, referred to in the detailed interpretation as Layer 1, thins in a northeast direction ( i . e. toward the mainland), and the velocity at the upper surface of this layer increases from 3.6 km/s, .for the southern line, to at least 4.0 km/s for the northern line. These velocities compare well with velocities obtained in previous studies: 4.05 km/s at Galiano Island by Milne and White (1960); 3.81 km/s by White (1962) and 4.0 km/s by Tseng (1968) near Hornby Island; 4.5 km/s near Galiano Island by Tseng (1968). The velocity gradients used for Layer 1, ranging from 0.2-0.55 km/s/km, are weakly constrained due to the lack of clearly defined secondary travel branches in this analysis. Some constraint, however, is provided by the observed cut-off distance for the large amplitude secondary arrivals. 9 F i g u r e 5 . 1 The f i n a l v e l o c i t y models a r e shown i n t h e i r r e s p e c t i v e p o s i t i o n s i n the S t r a i t of G e o r g i a . V e l o c i t i e s g i v e n i n the legend a r e v e l o c i t i e s a t the upper s u r f a c e of the i n d i c a t e d l a y e r . The models have been i n c l u d e d to a depth of 3 km, but because of the u n c e r t a i n t y of the v e l o c i t y g r a d i e n t s , the a c t u a l depth of p e n e t r a t i o n i s not well d e f i n e d . For the northernmost l i n e , the depth of p e n e t r a t i o n i s c l o s e r to 2 km. 95 L a y e r 1 r e p r e s e n t s t h e c o n t i n u a t i o n of t h e C r e t a c e o u s Nanaimo Group, from V a n c o u v e r I s l a n d , b e n e a t h G e o r g i a S t r a i t , and a t t h e n o r t h e a s t end of t h e SB5/5R model i t l i k e l y r e p r e s e n t s t h e c o n t i n u a t i o n of t h e Chuckanut assemblage from t h e m a i n l a n d . T h e s e s e d i m e n t a r y u n i t s a r e a s s e m b l a g e s of s a n d s t o n e , s i l t s t o n e , c o n g l o m e r a t e , and a r g i l l i t e . The c o n t i n u i t y of L a y e r 1, e x c e p t f o r t h e f a u l t , a c r o s s t h e S t r a i t i n t h e SB5/5R model s u g g e s t s t h a t t h e Nanaimo and Chuckanut a s s e m b l a g e s a r e e x t e n s i o n s o f t h e same m o r p h o l o g i c u n i t . L a y e r 1 i n c r e a s e s i n t h i c k n e s s t o w a r d t h e s o u t h e a s t , h a v i n g a maximum t h i c k n e s s of o v e r 2 km a t t h e s o u t h w e s t end of t h e SB3/3R l i n e . The s i m p l i f i e d model of Tseng (1968) f o r t h e S t r a i t of G e o r g i a ( F i g u r e 1.3), a l s o i n d i c a t e s a t h i c k e n i n g o f L a y e r 1 t o w a r d t h e s o u t h e r n p a r t o f t h e S t r a i t , a l t h o u g h he a l s o i n d i c a t e s an i n c r e a s e i n t h e v e l o c i t y o f L a y e r 1 i n t h e s o u t h e r n S t r a i t w hich i s c o n t r a d i c t o r y , t o what i s f o u n d i n t h i s s t u d y . Because T s e n g ' s v e l o c i t i e s f o r t h e s o u t h e r n S t r a i t of G e o r g i a a r e b a s e d on u n r e v e r s e d r e f r a c t i o n d a t a o n l y , t h e v e l o c i t i e s o b t a i n e d i n t h i s s t u d y a r e t h o u g h t t o be more r e p r e s e n t a t i v e of t h i s l a y e r due t o t h e c o n t r o l p r o v i d e d by t h e r e f l e c t i o n p r o f i l e s . The t h i r d l a y e r , r e f e r r e d t o as L a y e r 2 i n t h e d e t a i l e d i n t e r p r e t a t i o n , d i p s toward V a n c o u v e r I s l a n d a t a n g l e s of 2° t o 1 6 ° . I t i s v e r y s h a l l o w on t h e m a i n l a n d s i d e of G e o r g i a S t r a i t (0.35-0.90 km), and r e a c h e s d e p t h s n e a r 2.3 km toward V a n c o u v e r I s l a n d . T h i s l a y e r has a c h a r a c t e r i s t i c v e l o c i t y of 6.0-6.1 km/s a t i t s upper s u r f a c e . In a l l of t h e models, a s m a l l t r a n s i t i o n zone, w i t h h i g h v e l o c i t y g r a d i e n t , has been i n c l u d e d a t t h e t o p o f L a y e r 2. T h i s zone i s i n c l u d e d t o s a t i s f y t h e a m p l i t u d e 96 requirements of the observed data (see section 4.1.2). An a r t i f i c i a l boundary at the bottom of this thin zone is included, beneath which the velocity gradient is reduced so that the depth at which a velocity of 6.6 km/s is reached is consistent with the model of Tseng (1968), and also to prevent the amplitudes of the p 2 arrivals from becoming extremely large at greater distances. The velocity gradient for the transition zone is 0.5 km/s/km, and beneath i t the velocity gradients range from 0.1-0.15 km/s/km. These velocity gradients are poorly constrained. On the basis of seismic velocity alone, i t is not possible to ascertain whether Layer 2 represents the granitic type material observed on the mainland, or whether i t represents Karmutsen Volcanics which presumably underlie the Nanaimo Group on Vancouver Island. However, correlation of the surface of Layer 2 on the mainland side of Georgia Strait with reflection profiles, starting less than 2 km from shore, suggests that Layer 2 is the extension of the Coast Range i-ntrusives to beneath the Strait. The identification of bedrock as the continuation of the granitic rocks of the mainland is based on the observed absence of internal stratification on the reflection profiles (Tiffin, 1969). A single characteristic velocity of 6.0-6.1 km/s for the extent of Layer 2 is consistent with the data. A local fault structure has been located approximately 15 km northeast of the northern tip of Galiano Island. The positions of this feature, as indicated on a number of different profiles, are shown in Figure 5.2. This fault is modelled as a 123 50' 40' 30' 123 20' 123 50' 40' 30' 123 20' F i g u r e 5.2 The l o c a t i o n s of the v e r t i c a l f a u l t p roposed f o r the S t r a i t of G e o r g i a , as de te rmined from G d i f f e r e n t p r o f i l e s , a re shown. The c i r c l e s encompass the minimum u n c e r t a i n t i e s in any of the measurements ( i . e. the e f f e c t s of sonobuoy d r i f t a re not i n c l u d e d ) . 98 vertical fault (due to lack of constraint on the dip), and a throw of 0.55 km is indicated by the models with the down-dropped side of the fault toward the mainland. The strike of the fault is not determined, although, it is observed on perpendicular profiles which run northeast-southwest and northwest-southeast, respectively. The Malaspina fault (see Figure 1.2), previously proposed for the northeastern side of Georgia Strait by Muller (1977), is not observed in this study. Because the refraction profiles only extend to within 5-6 km of the mainland shoreline, the position of the fault can be adjusted so that i t lies closer to the coast, beyond the extent of the proposed models. But, as mentioned earlier, correlation of the surface of Layer 2 with horizons on the reflection profiles indicate that Layer 2 is an extension of the mainland slope. Also, E l l i s et a l . (1983) include a profile which crosses the proposed fault north of the area of this study. This line runs from Jervis Inlet on the mainland, across Texada Island, to the west coast of Vancouver Island, and has a depth of penetration of approximately 13 km beneath Georgia Strait. The model proposed for this profile shows no indication of a major crustal fault structure in the Strait of Georgia. Thus, both shallow and deeper seismic interpretations discriminate against the proposed Malaspina Fault. These results have implications concerning the tectonic models discussed in section 1.2. In the model adopted by Muller (1977), the Insular Belt has been shifted northward to its present position relative to the Coast Plutonic Complex by 99 transcurrent or transform faulting, suggesting that the boundary between these tectonic provinces is an old transform or transcurrent fault (i.e. the Malaspina Fault in Georgia Strait). The absence of this fault indicated by the seismic evidence suggests that revision of this aspect of Muller-'s model is necessary. Other models such as those of Dickinson (1976), Monger (1975), Monger and Price (1979), and Monger (1982) which do not require that the western edge of the CPC represent a major crustal discontinuity, are compatible with the results of this analysis. Although this study provides evidence against a major crustal fault forming the boundary between the Coast Plutonic Complex and the Insular Belt in Georgia Strait, the nature of this boundary remains vague. Since the northeast end of Layer 2 has been associated with the granitic mainland rocks, and because a single characteristic velocity over the extent of Layer 2 is consistent with the data, the granitic rocks of the mainland may extend most of the distance to Vancouver Island. Because this unit is not observed on Vancouver Island, and since i t is younger than the Karmutsen Formation, i t must terminate beneath the Cretaceous Nanaimo Group. How and where the granitic rocks terminate, and how far the Karmutsen Formation extends under the Nanaimo Group, are questions that remain to be answered. 5.1 Future Studies An extension of the reflection and refraction profiles onto the mainland, and onto the Karmutsen Formation on Vancouver Island would help to further resolve the essence of this 100 boundary between the CPC and the Insular Belt in this area. In particular, twoaspects concerning this boundary would be cla r i f i e d . First, although the evidence presented in this study strongly suggests that the Malaspina fault is non-existent, there is s t i l l uncertainty due to the termination of seismic coverage short of the mainland coast. An extension of refraction and reflection lines onto the mainland, would remove this uncertainty. Due to the rugged physiography of the mainland, a suitable site for this could be one of the inlets (such as Howe Sound) that penetrate the coastline. The profile of E l l i s et a l . (1983), mentioned earlier, actually does run from Jervis Inlet across the Strait of Georgia, but due to the shot-receiver spacing of over 5 km, the shallow structure is not well defined. A series of shorter overlapping refraction lines runing from Howe Sound across the Strait adjacent to the SB3/3R line would be suitable to extend the shallow structure of this survey to the mainland, and to further probe the existence of the Malaspina Fault. An extension of the refraction and reflection lines of this study onto the Karmutsen Formation on Vancouver Island, would be useful in determining how far the Karmutsen Formation continues beneath the Cretaceous Nanaimo Group, and the distance to which the Coast Range intrusives extend across the Strait. With the shallow structure well determined, longer range refraction profiles across Georgia Strait, such as that of E l l i s et a l . 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Canadian Journal of Earth Sciences, 14, pp. 2565-2577. Jones, D.L., Cox, A., Coney, P.J. and Beck, M.E. 1982. The Growth of Western North America. Scientific American, 247, (5), pp. 70-84. Kanasewich, E.R. 1981. Time Sequence Analysis in Geophysics , Third Edition, University of Alberta Press, Edmonton, 480 pp.. • MacLeod, N.S., T i f f i n , D.L., Snavely, P.D., Jr. and Currie, R.G. 1977. Geologic interpretations of magnetic and gravity anomalies in the Strait of Juan de Fuca, U.S.-Canada. Canadian Journal of Earth Sciences, 14, pp. 223-238. 104 Marks, L.W. 1980. Computational topics in ray seismology. Unpublished Ph.D. thesis, Department of Physics, University of Alberta, 170 pp.. Mattinson, J.M. 1972. Ages of zircons from Northern Cascade Mountains, Washington. Geological Society of America Bulletin, 83, pp. 3769-3783. McKee, A.H. 1966. CSP and shallow single end refraction lines. British Columbia Department of- Mines and Petroleum Resources Open File No. 1275. McMechan, G.A., and Mooney, W.D. 1980. Asymptotic ray theory and synthetic seismograms for laterally varying structures: theory and application to the Imperial Valley, California. Bulletin of the Seismological Society of America, 70, pp. 2021-2035. Milne, W.G., and White, W.R.H. 1960. A seismic survey in the vicinity of Vancouver Island, B.C.. Publication of the Dominion Observatory, 24, pp. 145-154. Monger, J.W.H. and Ross, C.A. 1971. Distribution of fusulinaceans in the western Canadian Cordillera. Canadian Journal of Earth Sciences, 8, pp. 259-278. Monger, J.W.H., Souther, J.G., and Gabrielse, H. 1972. Evolution of the Canadian Cordillera: a plate-tectonic model. American Journal of Science, 272, pp. 577-602. Monger, J.W.H. 1.975. Correlation. of eugeosynclinal tectonostratigraphic belts in the North American Cordillera. Geoscience Canada, 2, pp. 4-10. Monger, J.W.H. and Price, R.A. 1979. Geodynamic evolution of the Canadian Cordillera — progress and problems. Canadian Journal of Earth Sciences, 16, pp. 770-791. Monger, J.W.H. 1982. Tectonic accretion and the origin of the two major metamorphic and plutonic welts in the Canadian Cordillera. Geology, 10, pp. 70-75. Muller, J.E. 1977. Evolution of the Pacific Margin, Vancouver Island, and adjacent regions. Canadian Journal of Earth Sciences, v. 14, pp. 2062-2085. Roddick, J.A. and Hutchison, W.W. 1974. Setting of the Coast Plutonic Complex, British Columbia. Pacific Geology, 8, pp. 91-108. Rogers, G.C. 1983. Seismotectonics of British Columbia. Unpublished Ph.D. thesis, Department of Geophysics and Astronomy, University of British Columbia, 247 pp.. 105 Schwarz, E.J., Muller J.E., and Clark, K.R. (1980). Paleomagnetism of the Karmutsen basalts from southeast Vancouver Island. Canadian Journal of Earth Sciences, 17, pp. 389-399. Spence, G.S., Whittall, K.P., and Clowes, R.M. (1983). Practical synthetic seismograms for laterally varying media calculated by asymptotic ray theory. Manuscript in preparation. To be submitted to Bulletin of the Seismological Society of America. Symons, D.T.A. 1971. Paleomagnetism of Jurassic Intrusions, B. C . Geological Survey of Canada Paper 70-63, pp. 1-17. Symons, D.T.A. 1973. Concordant Cretaceous paleolatitudes from felsic plutons in the Canadian Cordillera. Nature,,241, pp. 59-61. Texaco Georgia Straits 1969 Project 1200% marine reflection data (NMO/Stack velocities). British Columbia Department of Mining and Petroleum Resources Open File No. 1530. Ti f f i n , D.L. 1969. Continuous Seismic Reflection Profiling in the Strait of Georgia, British Columbia. Unpublished Ph.D. thesis, Department of Geophysics, University of British Columbia, 177 pp.. Tozer, E.T. 1970. Marine Triassic faunas. I_n Geology and Economic Minerals of Canada (R.J.W. Douglas, ed.). Geological Survey of Canada, Economic Geology Report 1, pp. 633-640. Tseng, K.H. 1968. A new model for the crust in the vicinity of Vancouver Island, B.C.. Unpublished M.Sc. thesis, Department of Geophysics, University of British Columbia, 83 pp.. Young, G.B., and Braile, L.W. 1976. A computer program for the application of Zoeppritz's amplitude equations and Knott's energy equations. Bulletin of the Seismological Society of America, 66, pp. 1881-1885. Waldron, D.A. 1982. Structural characteristics of a subducting oceanic plate. Unpublished M.Sc. thesis, Department of Geophysics and Astronomy, University of British Columbia, 121 pp.. White, W.R.H. 1962. The structure of the earth's crust in the vicinity of Vancouver Island as ascertained by seismic and gravity observations. Unpublished Ph.D. thesis, Department of Physics, University of British Columbia, 97 pp.. White, W.R.H., and Savage, J.C. 1965. A seismic refraction and gravity study of the earth's crust in British Columbia. Bulletin of the Seismological Society of America, 55, pp. 463-486. 106 White, W.H., Harakal, J.E., and Carter, N.C. 1968. Potassium-argon ages of some ore deposits in British Columbia, Can. Inst. Min. Met. Bull., 61, pp. 1326-1334. Whittall, K.P., and Clowes, R.M. 1979. A simple, efficient method for the calculation of traveltimes and raypaths in laterally inhomogeneous media. Journal of the Canadian Society of Exploration Geophysicists, 15, pp.. 21-29. 107 Appendix A Asymptotic Ray Theory The algorithm used to produce synthetic seismograms for the two-dimensional velocity models in this study is that of Spence et a l . (1983), which is based on asymptotic ray theory (ART). A brief description of ART (taken from Cerveny and Ravindra, 1971) and of the algorithm of Spence et a l . (1983) is in order. The linearized equation of motion governing the propagation of elastic waves in an inhomogeneous, perfectly elastic, and isotropic medium i s : o52u = ( X + M ) V ( V - U ) + M V 2 U + V X ( V - U ) + V * X X ( V X U ) + 2 ( V M « V ) U (1) where, u(x",t) is the particle displacement vector, X(x) and u(x) assumed that the solution, u, can be expressed as an asymptotic expansion; where, T ( X ) is a phase function, and u(x) (k=0,1,2 ,.. .) are the k amplitude coefficients of the asymptotic expansion. This series solution is referred to as a 'ray series'. Strictly, this asymptotic expansion is valid only where T ( X ) is analytic. Roughly speaking, in terms of the travel time curve associated with a wave propagating through an elastic medium, this means that the expansion is invalid in the neighbourhood of end points, cusps, points where the curve is tangent to the travel time branch of another wave, or more generally, wherever the travel time curve is discontinuous. Such 6 t 2 are Lame paramters, and p(x) is the density of the medium. It is 108 points are referred to as singular points. Because the solution (2) of the equation of motion is expressed as an asymptotic expansion in terms of inverse powers of frequency, i t is most useful for high frequencies. In applying this approximation, it is assumed that the velocity gradients of the medium are much smaller than the frequencies of the propagating wave. Also, the radii of curvature of any interfaces within the medium are assumed to be much larger than the wavelength of the propagating wave. At points outside the neighbourhood of any singular points, the error in including only the f i r s t few terms of the series (2) tends to zero as frequency becomes infinitely high. For determining the most prominent kinematic and dynamic properties of reflected and refracted waves, i t is often sufficient to consider only the f i r s t term in (2); u = exp{iw(t-r)}u0 (3). This solution is the 'zero-order' solution and corresponds to geometrical optics. Other waves, such as head waves, diffracted waves, and inhomogeneous waves, are included in the higher order terms of the expansion, and are excluded in the zero-order solution. ART allows the amplitudes of propagating rays to be calculated. If the amplitude of a reflected or refracted ray at a point x 0 is A(x 0), then the amplitude of this ray at a point x according to the ART zero-order solution i s : 109 A(x) = M x J fv(x«)fl(xn)| V 2 TT /v' (x;)o' (x;)\ 1/2. C- (4) R(x,x0)l'v(x) p(x) / | | |v(x^r p(x^)vJ * i " where Cj is the plane-wave reflection/refraction coefficient of the ray at the jth interface, and x^  (j=1,2,3,...,k) is the point of incidence on the jth interface. The primed quantities are evaluated on the side of the interface from which the ray emerges. Also, k R ( x , x 0 ) = f d a ( x ) ] V2 l I / d a ( x ; ) \ 7 2 (5) d a ( x 0 ) j I I ^ d a ' ( x - ) ] is the geometrical spreading function for the ray. The quantity do(x) is the elementary ray tube area at the point x. The product term on the right of (5) represents the effects that interfaces have on geometrical spreading along the ray path, and the term on the left represents the change in ray tube area between interfaces. If the velocity model is parametrized in terms of blocks of constant velocity gradient, the left side term in (5) can be expressed as dg(x ) = 5jr cosfl(x) y __v0: rj (6) da(x 0) 860 v 0 t£t sin^oi where N is the number of blocks through which the ray passes, v 0^ is the velocity on entering the jth block, 9Q^ is the ray angle on entering the jth block measured with respect to the velocity gradient, and r^ is the distance travelled within the jth block. In the algorithm of Spence et a l . (1983), ray tracing is performed using a modified version of the original ray trace algorithm of Whittall and Clowes (1979). The amplitudes of reflected and refracted waves are based on the ART zero-order 110 solution. This algorithm estimates the quantity 5r/50o at each departure angle 60 by shooting two rays, at departure angles 60 and 60+i\d0l and uses the difference in epicentral distance Ar to calculate Ar/A0o. This procedure differs slightly from that of Marks (1980) who splined seven rays at each receiver to estimate 6r/60Of and that of McMechan and Mooney (1980) who used two rays at successive departure angles of 0o. The density associated with a compressional wave velocity is approximated by, p = 0.252 + 0.3788v (7) , and a value of Poisson's ratio of 0.25 is used for a l l velocity layers in the models. The plane-wave reflection/refraction coefficients used in (4) are those of Cerveny and Ravindra (1971, p. 63), a-nd a routine described by Young and Braile (1976) and Cassell (1982) is used to calculate them. Once the amplitudes of the reflected and refracted rays have been calculated, synthetic seismograms are produced using the same algorithm as McMechan and Mooney (1980). Displacements corresponding to the different arrival types are superposed, and amplitude and travel times are interpolated to the desired distances. A phase-shifted impulse is constructed using a linear combination of a unit impulse and its Hilbert transform, and an estimate of the source wavelet taken directly from the data is convolved with the calculated impulse synthetic seismogram. No head waves, diffractions, dispersion or attenuation effects have been included in the calculation of the synthetic seismograms. 

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