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A seismic refraction study of the Queen Charlotte fault zone Dehler, Sonya Astrid 1986

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A S E I S M I C R E F R A C T I O N S T U D Y O F T H E Q U E E N C H A R L O T T E F A U L T Z O N E by S O N Y A A S T R I D D E H L E R B.Sc.(Geophysics), The University of Calgary, 1983 A THES IS S U B M I T T E D IN P A R T I A L F U L F I L L M E N T O F T H E R E Q U I R E M E N T S F O R T H E D E G R E E O F M A S T E R O F S C I E N C E in T H E F A C U L T Y O F G R A D U A T E STUD IES Department of Geophysics and Astronomy We accept this thesis as conforming to the required standard T H E U N I V E R S I T Y O F BR IT ISH C O L U M B I A Apr i l 1986 (£> Sonya Astr id Dehler, 1986 7 I n p r e s e n t i n g t h i s t h e s i s i n p a r t i a l f u l f i l m e n t o f t h e r e q u i r e m e n t s f o r an advanced degree a t t h e U n i v e r s i t y o f B r i t i s h C o l u m b i a , I agree t h a t t h e L i b r a r y s h a l l make i t f r e e l y a v a i l a b l e f o r r e f e r e n c e and s t u d y . I f u r t h e r agree t h a t p e r m i s s i o n f o r e x t e n s i v e c o p y i n g o f t h i s t h e s i s f o r s c h o l a r l y purposes may be g r a n t e d by t h e head o f my department o r by h i s o r h e r r e p r e s e n t a t i v e s . I t i s u n d e r s t o o d t h a t c o p y i n g o r p u b l i c a t i o n o f t h i s t h e s i s f o r f i n a n c i a l g a i n s h a l l n o t be a l l o w e d w i t h o u t my w r i t t e n p e r m i s s i o n . Department o f oKOPHYSTr.s & ASTRONOMY The U n i v e r s i t y o f B r i t i s h Columbia 1956 Main Mall Vancouver, Canada V6T 1Y3 E-6 (3/81) i i A B S T R A C T The margin between the continental North American and oceanic Pacific plates west of the Queen Charlotte Islands is uniquely marked by an active transform fault zone. The region is the locus of oblique convergence between the two lithospheric plates. West of the fault zone the absent continental shelf is replaced by a 25 km wide scarp-bounded terrace at 2 km depth which separates the oceanic and continental crust. A n onshore-offshore seismic refraction survey was carried out in 1983 across the Queen Charlotte Islands region. Th i r ty - two explosive charges and several airgun lines were recorded on eleven land-based and six ocean-bottom instruments. A subset of the resulting data set was chosen to study the structure of the Queen Charlotte Fault zone and adjacent terrace. Two-dimensional ray tracing and synthetic seismogram modell ing produced a veloc-ity structural model of the fault region. Underlying the deformed terrace sediments is an upper 3 km thick faulted unit with velocity 3.8 km/s and a high gradient. The lower crustal region is on average 10 km thick and has velocity 5.3 km/s and a slightly lower gradient. Beneath this unit the Moho increases in dip from 5° to 19° eastward. The terrace velocities are anomalously low compared to the adjacent oceanic and continental crustal structures. Velocities of the oceanic crust are consistent with those observed in ophiolite sequences. The velocity structure of the continental crust is not well-defined; however, vertical offset of 1.1 km is seismically observed on the Rennell-Louscoone Fault on Moresby Island. Two tectonic mechanisms are proposed to explain the anomalous terrace structure. Subduction of the oceanic lithosphere beneath the terrace would accrete sediments to the seaward edge of the terrace and subduct material beneath it. Upthrust ing of terrace iii material along near-vertical fault planes would result from compression at the transform fault above an inactive subduction zone. A combination or alternation of the two mechanisms would explain the observed fault zone structure. iv T A B L E O F C O N T E N T S A B S T R A C T ii L I ST O F T A B L E S vi L I ST OF F I G U R E S vii A C K N O W L E D G E M E N T S ix C H A P T E R I I N T R O D U C T I O N 1 1.1 Regional Geotectonic History 1 1.1.1 Terrane accretion 1 1.1.2 Plate interactions 4 1.2 Present Tectonic Regime 7 1.2.1 Geologic setting 7 1.2.2 Geophysical studies 10 1.3 Exper imental Objectives 16 C H A P T E R II D A T A A C Q U I S I T I O N A N D P R O C E S S I N G 17 2.1 The Refraction Experiment 17 2.2 Instrumentation 19 2.3 Da ta Processing 20 2.3.1 Initial data reduction 20 2.3.2 Shot and receiver positioning 21 2.3.3 Depth determinations 23 2.3.4 T iming calculations and corrections 23 2.3.5 Fi l ter ing 25 2.3.6 Ampl i tude corrections 29 C H A P T E R III D A T A A N A L Y S I S 30 3.1 Init ial Observations and Interpretation 30 3.1.1 The airgun data set 30 3.1.2 The three-layer case 31 3.1.3 The explosion data set 33 3.1.4 First break analysis 34 3.2. The Modell ing Method of Interpretation 35 3.2.1 The modell ing algorithm 36 3.2.2 Model l ing techniques 36 V C H A P T E R IV I N T E R P R E T A T I O N O F A I R G U N D A T A 38 4.1 Initial Constraints 38 4.2 Profile O B S W 2 40 4.3 Profile O B S W l E A S T 44 4.4 Profile O B S W l W E S T 48 4.5 The F inal Model . . . 51 C H A P T E R V I N T E R P R E T A T I O N O F E X P L O S I O N D A T A 52 5.1 Initial Constraints 52 5.2 Receiver 010 53 5.3 Receiver 020 58 5.4 Receiver 030 62 5.5 Receiver 060 • 66 5.6 Receiver 100 67 5.7 Shot 13 70 5.8 Shot 25 76 5.9 The F ina l Velocity Model 79 C H A P T E R V I D ISCUSS ION A N D C O N C L U S I O N S 82 6.1 Stratigraphic Interpretation 82 6.2 Tectonic Implications 85 6.3 Conclusions 93 B I B L I O G R A P H Y 94 A P P E N D I X A Ai rgun Data 98 A P P E N D I X B Explosion Data 102 vi L I S T O F T A B L E S I. Receiver numbers and locations of seismographs 20 II. T im ing uncertainties 25 III. Velocity and depth estimates for terrace region 33 IV. Stratigraphic interpretation of velocity units 83 Vll L I S T O F F I G U R E S 1. Plate tectonics off Canada's west coast 2 2. Suspect terranes of the QCI region 3 3. Past plate interactions off western Canada 6 4. Physiographic location map of the QCI 9 5. Line drawing of CSP profile across terrace region 11 6. Lithospheric flexural model of the QC I region 13 7. Bathymetry west of the QC I 14 8. Velocity structural model from previous study 15 9. Location of receivers and survey lines 18 10. Periodograms for receiver 010 26 11. Periodograms for receiver 100 27 12. Comparison of unfiltered and filtered data 28 13. Three-layer interpretation model 32 14. F i rst arrival curves for explosion data 35 15. F inal velocity model from airgun data 39 16. O B S W 2 ray tracing model and traveltimes 41 17. O B S W 2 seismograms 42 18. O B S W l E A S T ray tracing model and traveltimes 45 19. O B S W l E A S T seismograms 46 20. Detai l from O B S W l E A S T data section 47 21. O B S W l W E S T ray tracing model and traveltimes 49 22. O B S W l W E S T seismograms 50 23. F inal velocity model 54 v i i i 24. Receiver 010 model and traveltimes 56 25. Receiver 010 seismograms 57 26. Receiver 020 model and traveltimes 60 27. Receiver 020 seismograms 61 28. Geological cross section of Moresby Island 63 29. Receiver 030 model and traveltimes 64 30. Receiver 030 seismograms 65 31. Receiver 060 model and traveltimes 68 32. Receiver 060 seismograms 69 33. Receiver 100 model and traveltimes 71 34. Receiver 100 seismograms 72 35. Shot 13 model and traveltimes 74 36. Shot 13 seismograms 75 37. Shot 25 model and traveltimes 77 38. Shot 25 seismograms . 78 39. Seismic profiles across the Aleutian Trench 88 40. Tectonic model of the Q C Fault zone 89 41. Conceptual diagram of an upthrust welt 90 42. Proposed tectonic processes at the Q C Fault 92 ix ACKNOWLEDGEMENTS To my parents with gratitude for their years of love, support, and encouragement. I would like to thank my supervisor, Dr. Ron M. Clowes, for his support and guidance over the years and the many helpful discussions which kept me "on track". I would also like to thank Dr. Robert M . El l is for critically reviewing this thesis. I am greatly indebted to Bob Meldrum for his patient and thorough explanations of the more technical (and mystical) aspects of this study. My thanks also go to David Mackie for his invaluable assistance with the processing, and to many others who helped with the thesis, in particular Joane Berube, Jeff Drew, Chris Pike, Alex von Breymann, and Bar ry Zelt. The members of the department (faculty, staff, and students) have made this an enjoyable place to study. M y thanks to al l of them for their friendship. Financial support for data acquisition and analysis was provided by N S E R C through Operat ing Grant A7707 and Strategic (Oceans) Grant G0738, and by D.S.S. contracts 05SU.23235-3-1089 from the Ear th Physics Branch, E.M.R., and 06SB.23227-4-0904 from the Pacific Geoscience Centre, E.M.R. Also, E.M.R. Research Agreement 287 provided support in 1983/84. Addit ional support was provided by Chevron Canada Resources L t d . and Shell Canada Resources L td . 1 CHAPTER I INTRODUCTION The Queen Charlotte Fault zone marks the active boundary west of central Br i t ish Co lumbia, Canada, between the North American and Pacific lithospheric plates. Co-incident with the fault zone between 52° and 54° is the western margin of the Queen Char lotte Islands (figure l ) . The transition across the margin is structurally complex, a product of the long and complicated geologic history of the region. A brief review of this history will provide a tectonic framework for newly developed structural models. 1.1 Regional Geotectonic History 1.1.1 Terrane accretion Much of the North American Cordi l lera is composed of "suspect" terranes of un-certain paleogeograp hie origins (figure 2). Most of these terranes are allochthonous to North Amer ica and were accreted to the cratonic margin as a result of plate interactions (Coney et al., 1980). The Queen Charlotte Islands and immediate vic inity are underlain by two such terranes, Alexander and Wrangell ia. The Alexander terrane is a continental edge fragment (Coney et al., 1980) consisting of a complex assemblage of volcanic, clastic, and limestone rocks of primari ly Late Precambrian to Late Paleozoic age (approximately 650 to 250 Ma) . Wrangellia is a submarine arc complex composed of Middle to Upper Triassic tholeiitic basalts and calcareous sedimentary rocks and Early to Midd le Jurassic terrigenous elastics and volcanics (Coney et al., 1980). Wrangellia terrane rocks are represented on the Queen Char lotte Islands by the Karmutsen, Kunga, Maude, and Yakoun Formations (Yorath and Chase, 1981; Sutherland Brown, 1968). 2 F i g u r e 1. P late tectonics off Canada's west coast (after Riddihough, 1982a). 3 J' ' | CHUGACH TERRANE — Cretaceous sediments and melange I 1 TRACY ARM TERRANE — Central Gneiss complex: Late MesozoJc I 1 and Cenozolc plutons of Coast Plutonic Complex • GRAVINA - NUTZOTIN BELT — Jurrasslc and Cretaceous arc volcanlcs and sediments HHH TAKU TERRANE — Upper Paleozoic to Trlassic volcanlcs, sediments WRANGELLIA — Late Trlassic to Jurassic volcanlcs, sediments. Paleozoic basement STIKINE TERRANE — Late Paleozoic to Jurassic volcanlcs, sediments ALEXANDER TERRANE — Paleozoic volcanlcs, sediments, plutonic rocks -CD Faults — 1. Chatham Strait Fault 2. Clarence Strait Fault 3. Queen Charlotte Fault 4. Beresford Bay and Langara Faults S.Sandspit Fault 6. Rennell Sound Fault 7. Louscoone Inlet Fault 8. Grenville Channel Fault 9. Principe Laredo Fault 10. Kltkalta Fault H.Holberg Fault F i g u r e 2. Suspect terranes of the Queen Charlotte Islands region (after Yorath and Chase, 1981). 4 Both terranes are believed to have moved northwards from their places of origin before accreting to North America. Paleomagnetic studies (van der Voo et al., 1980) indicate that Alexander terrane moved 1800 km northwards with respect to cratonic North Amer ica during Late Carboniferous and Triassic time (300 to 180 Ma) . Wrangellia is believed to have originated within approximately 15° of the paleo-equator (in either the northern hemisphere or the southern hemisphere, the latter more favoured by Irving, 1979, and Hilde et al., 1977) and subsequently travelled between 3000 and 6300 km northwards. Wrangell ia collided with the western edge of Alexander terrane during Middle to Late Jurassic t ime (170 to 130 Ma) along a suture zone marked by the present Sandspit Fault on the Queen Charlotte Islands and Rennell Sound Fault through Hecate Strait (figure 2). Dur ing the collision the terranes were uplifted and intruded by plutons dated at 144 M a (Young, 1981). The amalgamated superterrane collided with the westward moving North Amer ica plate in the Late Cretaceous or Early Tertiary period (90 to 40 Ma) resulting in plutonic uplift of the Coast Mountains and a westward jump of the active continental margin to the seaward edge of the terrane (Yorath and Cameron, 1982). 1.1.2 Pl a t e interactions The concept of seafloor spreading has enabled researchers to reconstruct plate mo-tions and interactions off western North Amer ica using magnetic anomalies, hot spot traces, and paleomagnetism (eg. Atwater,1970; Coney, 1970; Stone,1977; and R id-dihough, 1982b). A l l models propose convergence and strike-slip faulting along the 5 margin for the past 100 Ma , although the details of the interactions do not always agree. A generalized plate motion reconstruction presented by Riddihough (1982b) shows the types and rates of past plate interactions off western Canada (figure 3). Three plates interacted with the Amer ica (A) plate: the Pacific (P) plate, the Ku la (K) plate which once existed north of the Pacific plate, and the large Farallon (F) plate east of the Pacific plate. Dur ing the Middle to Late Cretaceous (100 to 80 Ma) the Farallon plate interacted in a generally north to northeast direction with the Amer ica plate at an estimated rate of 60 to 120 mm/a . By Late Cretaceous to Eocene time (80 to 40 Ma) the rate had increased to 120 to 170 mm/a in a northeast direction. The Ku la plate was obliquely converging at 100 to 150 mm/a with the Amer ica plate at this time and the K - F - A triple junct ion was migrating either southwards along the coast from Alaska (Stone, 1977) or northwards from southern Cal i fornia (Atwater, 1970) to a position near Vancouver Island. In Late Eocene time (45 to 40 Ma) the Pacific plate changed its direction of motion by 50°, as indicated by the bend in the Emperor-Hawai i chain, resulting in a reduction of al l interaction rates across the North American margin by one-half (Riddihough, 1982b). Dur ing the Oligocene to Early Miocene (40 to 20 Ma) a number of plate adjustments occurred as a southern part of the P - F spreading ridge started subducting beneath Cal i fornia, and transcurrent motion was initiated between 30 and 20 M a along the San Andreas Fault (Atwater, 1970). The northern part of the Farallon plate broke off and began to move independently as the Juan de Fuca (J) plate around 20 M a ago. The K - F - A triple junct ion had meanwhile continued migrating rapidly northwards along the margin and the remainder of the Ku la plate was subducted during this time. 6 F i g u r e 3. Plate interactions off western Canada during the last 100 M a (Riddihough, 1982b). P - Pacific plate, A - America plate, K - Ku la plate, and F - Farallon plate. Arrows indicate relative directions of motion; rates of motion in mm/yr are shown. 7 During the Miocene (20 to 10 Ma) the Juan de Fuca plate continued subducting beneath the Amer ica plate at 40 to 60 mm/a in a northeastward direction and the Pacific plate moved northwestward at 50 to 60 mm/a, producing transcurrent motion along a paleo- Queen Charlotte Fault at the Amer ica plate margin. The P - J - A triple junction migrated southeastwards along the margin during this time and may have been responsible for the init iat ion of rifting in Queen Charlotte Sound 17 M a ago which eventually shifted and rotated the Queen Charlotte Islands along reactivated faults to their present position (Yorath and Chase, 1981; Horn, 1982). Alternatively, the America plate may have overridden a hot spot or mantle plume which caused the rifting and produced the Anahim volcanic belt of central Br i t i sh Columbia (Yorath and Cameron, 1982; Yorath and Chase, 1981). Final ly, the Pliocene and Pleistocene (10 to 0 Ma) saw the stabilization of the P - J - A triple junction off northern Vancouver Island and a decrease in the J - A convergence rate to 40 mm/a as the spreading direction at the Juan de Fuca ridge changed (Riddihough, 1977). New spreading ridges were formed at Dellwood Knolls 1 M a ago and Tuzo Wi lson Knol ls 0.5 M a ago and joined to the stable triple junction by transform faults (Riddihough et al., 1980). Spreading at the Tuzo Wilson Knolls caused the Queen Char lotte Fault to jump landwards to its present location where it is currently active (Hyndman and El l is, 1981). 1.2 P r e s e n t Tec ton i c R e g i m e 1.2.1 Geo l og i c s e t t i ng The Queen Charlotte Islands were divided by Sutherland Brown (1968) into three physiographic units: the Queen Charlotte Ranges, Skidegate Plateau, and the Queen 8 Charlotte Lowlands (figure 4). The regions differ not only in physiographic features but also in geologic composition, as the 1100 m high Ranges are underlain by granitic and Tertiary volcanic rocks, the 600 m high Plateau by sl ightly-t i l ted Jurassic volcanic, Cretaceous sedimentary, and Early Tertiary basaltic rocks, and the 150 m high Lowland by complex sequences of rocks from Jurassic to Pleistocene age (Sutherland Brown, 1968). The boundaries of the physiographic units generally follow two major northwest-trending fault systems, the Rennell Sound—Louscoone Inlet Fault and the Sand spit Fault (figure 2). The Rennel l-Lou scoone Fault strikes N25°W in the south and N50°W north of Louise Island. The fault zone is composed of several associated strands that dip between 75° northeastward and 90° vertical. Total fault displacement since Late Jurassic is estimated as 600 to 3000 m downwards and 19 to 93 km southward for the east block relative to the west block (Sutherland Brown, 1968). The Sandspit Fault strikes uniformly N37°W and dips at 65° northeastward. The fault zone separates older rocks in the west block from younger formations in the east and cuts through plutons of Cretaceous age. Total fault displacement is estimated as thousands of metres downward for the east block relative to the west and unknown but presumably considerable horizontal offset (Sutherland Brown, 1968). The major faults on the islands are colinear with the Queen Charlotte Fault located at the foot of the continental shelf just offshore. The continental shelf is generally less than 5 km wide west of the Queen Charlotte Islands and is little more than a notch in the wide, steep continental slope. The slope is composed of three parts: a steep upper scarp, a middle terrace plateau, and a lower scarp. The upper scarp has a gradient of up to 20° and includes the 9 F i g u r e 4. Physiographic location map of the Queen Charlotte Islands (Sutherland Brown, 1968). 10 truncated spurs of the Insular Mountains at the coast. The terrace is at an average depth of 2 km and has variable topography which becomes more rugged towards its termination off southern Moresby Island. The lower scarp slopes as steeply as 23° (Chase and Tiffin, 1972). The physiography and geology of the continental slope and shelf appear to be linked to past and present interaction between the Pacific and America plates at the Queen Char lotte Fault zone. 1.2.2 Geophysical studies Present motion at the Queen Charlotte Fault zone is primarily dextral strike-slip at an average rate of about 55 mm/a (eg. Atwater, 1970; Keen and Hyndman,1979; Hyndman and Weichert, 1983). The strike of the fault zone west of Moresby Island, as obtained from its morphology and seismicity, is over 10° different from the more easterly relative plate motion predicted by global plate models (Minster and Jordan, 1978). This suggests that convergence at 10 to 20 mm/a is occurring at the fault zone as subduction of the Pacific plate beneath the Queen Charlotte Islands and/or as crustal deformation of the islands and adjacent region (Hyndman and Ellis, 1981; Yorath and Hyndman, 1983). The active fault has been located by recent microseismicity studies (Hyndman and El l is, 1981; Berube, 1985) as being beneath the upper fault scarp of the continental slope. Composite P-nodal fault plane solutions for earthquake clusters indicate primar-ily strike-sl ip motion along the fault west of Graham Island and some thrusting and vertical faulting with ocean side down west of Moresby Island (Berube, 1985). This is consistent with studies of large earthquakes which determined strike-slip motion with a 11 small thrust component for the 1949 magnitude 8.1 earthquake west of Graham Island and a str ike-sl ip mechanism with a significant thrust component along a fault plane dipping 50° eastward for the 1970 magnitude 7.0 earthquake just south of Moresby Island (Rogers, 1983). Seismic reflection profiles across the fault zone reveal the effect of compression and str ike-sl ip movement on the continental margin (eg. Chase and Tiffin, 1972; Chase et al. , 1975; P.D. Snavely, unpublished data). The terrace region of the continental slope is underlain by grabens and synclines infilled with deformed sediments (figure 5). The steep upper scarp is the currently active Queen Charlotte Fault and the outer scarp is presumed to be the former fault location. Seismic reflections from oceanic basement end abruptly at the outer scarp, as do magnetic lineations from the oceanic crust (Currie et al., 1980). Tectonic faulting, possibly associated with trench formation, could be destroying the remnant magnetization of the crustal rocks at the edge of the terrace (Srivastava, 1973). T 1 1 1 1 1 1 0 KILOMETERS F i g u r e 5. Line drawing interpretation of a C S P profile west of northern Moresby Island (Chase and Tiffin, 1972). 12 Heat flow and gravity profiles of the oceanic crust also show dramatic changes as they cross over the fault zones and terrace to the continent. A sharp decrease in heat flow at the seaward edge of the terrace was observed by Hyndman et al. (1982). They also noted a general decrease in heat flow by a factor of three from the axis of the Queen Char lotte Trough to 40 km inland, consistent with measurements in many subduction zones. The low terrace heat flow was attr ibuted to tectonic thickening of sediments, possibly by accretion, and the underthrusting of cold oceanic lithosphere (Hyndman et al., 1982). Free-air gravity data indicate a large negative anomaly at the base of the continental slope and a positive anomaly on the landward side of the continental shelf (eg. Couch, 1969; Dehlinger et al., 1970; Srivastava, 1973; Curr ie et al., 1980). The gravity low over the terrace may be due to a thick prism of sediments or unusually thick crust (Currie et al., 1980). The gravity profile is again characteristic of subduction zones (Keen and Hyndman, 1979). A lithospheric flexural model for the Queen Charlotte Islands region, developed by Yorath and Hyndman (1983), agrees well with most of the geophysical and physiographic observations (figure 6). Uplift at the edge of the continental lithosphere and depression of the adjacent area inland, a flexural response to underthrusting, have been observed and described by Riddihough (1982a) in the Queen Charlotte Islands region. The corresponding flexural bulging of the oceanic lithosphere is observed on bathymetric maps as the Oshawa Rise (figure 7). The subduction process would produce a thick accretionary wedge of sediments in the vicinity of the continental slope terrace. These sediments would be faulted and deformed by the additional transform motion, as would the crust beneath them. 13 TERRACE ACCRETIONARY OFFSHORE SEDIMENTARY WEDGE (Yorath and Hyndman, 1983). Seismic refraction models developed for the region support underthrusting of the Pacific plate beneath the Queen Charlotte Islands as a possible convergence mechanism (Mackie, 1985; Horn et al., 1984; Horn, 1982; B i rd, 1981). The model of Horn et al. (1984) indicates anomalously low velocities and high gradients in the terrace region separating the oceanic and continental blocks (figure 8); the model is consistent with an accretionary wedge of compressed sediments overlying a subducting oceanic crust. The subduction and transform motion at the continental margin and the variable geotectonic history appear to have created a unique and complex structure at the fault zone. 14 F i g u r e 7. Bathymetry west of the Queen Charlotte Islands. The locations of the C S P linedrawing of figure 5 and the refraction model of figure 8 are shown. 15 Distance (km) 40 60 80 F i g u r e 8. Velocity structural model of Horn et al. (1984) from a seismic profile off southern Moresby Island. 16 1.3 E x p e r i m e n t a l Ob j e c t i v e s As the previous discussion has indicated, the structure at the junction of the Pacific and America plates is very complex. Geophysical observations across the continental margin indicate that a thick wedge of sediments may be accreting to the margin as a result of subduction of the Pacific plate. A t the same time, dextral transform motion along the active fault plane is pushing the accretion zone northwards and deforming the sediments and underlying crust. A seismic refraction experiment was designed to study the structure of the terrace region that marks the active plate boundary west of the Queen Charlotte Islands, and to try to determine the origin and composit ion of the terrace. To achieve these objectives it is necessary to look at the structural transit ion from the oceanic crust west of the terrace to the continental crust beneath the Queen Charlotte Islands. 17 C H A P T E R II D A T A A C Q U I S I T I O N A N D P R O C E S S I N G 2.1 T h e R e f r a c t i o n E x p e r i m e n t A major marine-land seismic refraction survey was carried out in August 1983 by the University of Br i t i sh Columbia (UBC) in collaboration with the Earth Physics Branch, Ottawa, and the Pacific Geoscience Centre, Sidney. A 330 km profile was recorded from the deep ocean across northern Moresby Island and Hecate Strait to the mainland of Br it ish Columbia (figure 9). The experimental objectives were to determine the structure of the sediment, crust, and lithosphere ( l ) of the Queen Charlotte transform fault zone, (2) beneath the Queen Charlotte Islands, and (3) below Hecate Strait. A total of seventeen seismographs were deployed: two ocean bottom seismographs (OBSs) west of the Queen Charlotte Islands, eight land-based seismographs (LBSs) on the Islands, four OBSs in Hecate Strait, and three additional LBSs on the mainland. Two types of energy source were used. The pelleted T N T explosive NITROPEL 1 ^ 1 packed into drums with primacord and C.I.L. Procore Pr imers^ was detonated every 3.4 km along a 110 km line bearing southwest from northern Moresby Island. Twelve 540 kg and twenty 60 kg shots were fired, with two of the small shots following each large shot. Three addit ional refraction profiles, one over the western OBSs and two in Hecate Strait, were shot using a 32 l itre ( 2000 in3 ) airgun with a shot spacing of approximately 0.18 km. This study of the structure of the Queen Charlotte Fault zone will focus on the airgun profile west of the Queen Charlotte Islands and the explosion profile as recorded by the ten westernmost receivers. 18 Figure 9. Location of receivers and survey lines in the 1983 refraction ex-periment. Three of the 540 kg explosion locations ( l , 16, 33) are indicated as are receiver locations 010, 100, and 170. 19 2.2 I n s t r u m e n t a t i o n Four types of seismographs were deployed during the survey (table I). The six OBSs, UBC-mod i f i ed versions of At lant ic Geoscience Centre instruments (Heffler and Barret, 1979), consist of a frame holding a flotation sphere, an instrument pressure case, a hydrophone, recovery devices, and a detachable weight to take the package to the ocean bottom. Inside the pressure case are gimbaled 4.5 Hz horizontal and vertical component geophones, primary and secondary clocks, a microprocessor, a 4-channel slow-speed direct record analog tape recorder, and the necessary power supplies. The outputs from the geophones and hydrophone were recorded together with an internally-generated 10 Hz ampl itude-modulated time code directly onto cassette tapes. Six Earth Physics Branch ( EPB ) Mark II digitally recording "Backpack" seismo-graphs were deployed on Moresby Island and the mainland. The Backpacks sampled the output from a 2 Hz vertical component seismometer at 60 samples per second (sps) and recorded on digital cassettes. A n internal clock provided timing for the data. Four digitally recording U B C Microcorders were deployed on Moresby Island. The output from a 1 Hz vertical component seismometer was sampled at 60 sps and recorded on digital cassette tape. T iming was by an internal clock. The last site on Moresby Island was occupied by a U B C F M analog recorder. It contained 1 Hz vertical and horizontal component seismometers and recorded their output at high and low gain settings together with W W V B time code on five parallel tracks of 7-track analog tape. The seismographs successfully recorded all shots during the experiment. 20 INSTRUMENT TYPE RECEIVER NUMBER LOCATION OBS 010 offshore terrace 020 offshore terrace 110 Hecate Strait 120 Hecate Strait ISO Hecate Strait 140 Hecate Strait MCR 050 Moresby Island 070 Moresby Island 080 Moresby Island 100 Moresby Island FMA 090 Moresby Island EPB OSO Moresby Island 040 Moresby Island 060 Moresby Island 150 mainland region 160 mainland region 170 mainland region T a b l e I. Receiver number and location of seismographs used in the exper-iment. Instrument types are: O B S - ocean bottom seismograph, M C R -microcorder, F M A - F M analog recorder, and E P B - Earth Physics Branch "Backpack" recorder. 2.3 D a t a P r o c e s s i n g 2.3.1 I n i t i a l da t a r e d u c t i o n The init ial stage in the data reduction process was to standardize the data recorded by the different seismograph systems. First, the analog data from the OBSs and F M recorder were converted to digital format on 9-track magnetic tapes using a P D P 11/34 -based digitization and editing package. Digit ization of the O B S data was a two-step process involving the transfer of field data, recorded at 0.2 mm/s, to high quality 1/4 inch F M tape at 6.0 mm/s. The 21 resulting 30-t ime speed-up was necessary to overcome digitization system limitations for the slow-speed data recordings. The 4-channel F M tapes of OBS data were then digit ized at 120 sps using the 10 Hz time code signal recorded on one of the tracks as a trigger to avoid t iming problems associated with tape stretch and recorder speed variations. A detailed description of the OBS digitization procedure may be found in Mackie (1985). The F M recorder analog data tapes were digitized directly at 120 sps using an external frequency generator as a sampling trigger. T iming errors associated with tape stretch and tape speed variations were negligible. Both the OBS (explosion only) and F M analog data were later reduced to 60 sps recordings for plotting compatibi l ity with the digitally recorded data from the other seismographs. Data from the E P B "Backpack" systems were transfered from digital cassettes to 9-track magnetic tapes by personnel at the Earth Physics Branch. A similar transfer of U B C Microcorder data was done at U B C . Once the entire data set was in digital format on magnetic tapes, editing of shot recordings and demultiplexing was done using the P D P 11/34 system. A l l explosion recordings and randomly-selected airgun shots were plotted to monitor data quality before proceeding with further processing. 2.3.2 Shot and receiver positioning The position of the ship was determined from Loran C navigation with a relative accuracy of ± 200 m and an absolute accuracy of ± 300 m (Hyndman et al., 1979). The accuracies were increased by updating the Loran C system with satellite fixes when available. The hyperbolic Loran C measurements were converted to geographic 22 coordinates and a linear interpolation between measurements was used to determine the ship's position for each airgun shot. The location of the explosive shots was determined from the ship's position at drop time. Similarly, the O B S positions were measured at the time of deployment and assumed a vertical fall of the instrument to the ocean floor. The L B S positions were directly obtained from 1:50,000 scale topographic maps with absolute accuracies of approximately 150m. Shot-receiver distances were computed from the geographic coordinates of the vari-ous shots and receivers. A geometric factor was applied to the airgun data to correct for the offline location of O B S W l (receiver 010) and shift the O B S position to the shooting line. This correction ranged from 2.4 km for the closest shot to 0.2 km for the shot 15 km distant. A similar correction incorporating receiver depth as well was applied to the corresponding travel times. A positioning error was noticed on recordings of explosive shots at O B S W l . Hy-drophone component plots showed that direct water wave arrivals did not pass through the origin but instead intersected the distance axis at a point 1.78 km away. When the bathymetric profile was compared to the depth profile of the overlapping airgun survey a discrepancy of approximately 2 km was observed for major physiographic features. The result was a 1.78 km correction applied to shot-receiver distances for the O B S W l explosion data set. The error may have resulted from an incorrect log entry during data acquisit ion; the above correction has hopefully restored the true positions and distances. 23 2.3.3 Depth determinations Water depth was continuously recorded through the ship's depth sounding system on an E P C line scan recorder during the survey. The depth of each OBS was determined from these recordings at the time of deployment. The elevations of the LBSs were read from the topographic maps on which their positions were located. Two methods were used to calculate explosion shot depths. The geometrical method, described in Horn (1982), uses the ship speed, ditch time, water depth, and the time lag between the direct and bottom reflected water waves as recorded on a trai l ing hydrophone to compute shot depths. This method provided a good estimate of depth for the more distant shots beneath which the seafloor was relatively flat and horizontal. The bubble pulse method relies on the relationship between explosion depth and the period of first oscillation of an expanding explosion-generated gas bubble which is forced to collapse due to hydrostatic pressure. The Rayle igh-Wi l l i s bubble formula, based on work by Lord Rayleigh (Strutt, 1917) and Wil l is (1941), was used to compute depths for shots over the irregular topography of the terrace. Both methods have a maximum estimated error of ± 30 m. The airgun maintained a relatively constant depth of 17 m throughout the survey as controlled by the length of the towing cable and the ship speed, thereby obviating the need for complicated depth calculations. 2.3.4 Timing calculations and corrections The W W V B time signal received aboard ship was used as absolute time throughout the survey. Readings from other clocks and instruments were converted to W W V B time during processing to standardize time measurements. 24 Shot origin times for explosions were calculated from the shot depth and distance previously computed, using where d =shot depth, t<i =arrival time of direct water wave, x —ship-to-shot dis-tance, and Vw ^water velocity at 1.49 km/s (Horn, 1982). The airgun firing time was controlled by a trigger clock for which the signal was recorded with W W V B code on both a 2-channel chart recorder and a 4-channel tape recorder. A linear clock drift correction was subsequently applied to airgun shot origin times based on these recordings. A l l instruments with internal clocks (OBSs and Microcorders) were rated against either W W V B (OBSs) or W W V (Microcorders) immediately before and after use. In-ternal clock readings in code were recorded directly on to the tapes. Recording times could thus be related to W W V B after applying any necessary corrections for the linear drift of the internal clocks. The F M analog system recorded W W V B directly onto a parallel channel thereby avoiding timing problems. Clock drift for the E P B Backpacks is unknown as no information was provided by E P B personnel. Corrections applied to times of first sample include the geometrical offline corrections mentioned in section 2.3.2 and shot depth corrections. The explosions and airgun shots were vertically shifted to the ocean surface using the shot depths and a water velocity of 1.49 km/s. The maximum uncertainties due to shot origin times, clock drift, shot depths and travel time picks are summarized in table II. The time errors due to uncertain travel t ime picks are considerably larger than other t iming uncertainties in most cases. 25 CORRECTION OBS (s) MCR (s) FMA (s) EPB (s) Origin Time (EX) O.OSO O.OSO O.OSO O.OSO Origin Time (AG) 0.010 0.010 0.010 0.010 Clock Drift 0.0 IS 0.006 0.006 0 * Shot Depth 0.020 0.020 0.020 0.020 Time Picks 0.01-0.05 0.01-0.15 0.01-0.15 0.01-0.05 Tab l e I I . T im ing uncertainties due to various corrections and traveltime measurements. E X - explosives, A G - airgun; * = value assumed. 2.3.5 F i l t e r i n g Spectral analysis of airgun and explosion data was performed for representative signal and noise recordings. Figure 10 shows periodograms for an O B S recording of an explosion. The seismic signal of the O B S extends from 0.5 to 15.0 Hz with significant noise in the 18 to 27 Hz range. Periodograms for a Backpack recording of the same explosion (figure 11) indicate the presence of noise below 2 Hz and strong signal above 2.5 Hz. The characteristics of these periodograms are typical of all L B S recordings which show low frequency signal and a clear noise-signal separation leading to relatively good signal-to-no ise ratios. A n eight-pole zero-phase Butterworth bandpass filter was applied to the explosion data to reduce the noise content. A 0.5-15.0 Hz bandpass was used for the OBSs and a 2-12 Hz bandpass filter was applied to most of the L B S data. Fi ltering visibly enhanced the data especially on receivers 090 and 100 where noise had obscured first arrival energy (figure 12). The filtered data sections were used to approximately locate the first arrivals; more accurate time picks were then made on the unfiltered sections. 26 NOISE (010) o_ c o . d Frequency (Hz) SIGNAL (010); o , ', c o . d Frequency (Hz) F i g u r e 10. Periodograms computed for segments of noise and signal + noise recorded on the vertical geophone of receiver 010, an OBS. 27 NOISE (100) o o . 6 0 15 20 25 30 1Frequency (Hz) SIGNAL (100) '0 5 10 15 20 25 30 Frequency (Hz) F i g u r e 11. Periodograms computed for segments of noise and signal + noise recorded on receiver 100, a microcorder on Moresby Island. 28 200 180 160 140 Receiver 100 (2 120 12 Hz) 100 F i g u r e 12. A comparison of unfiltered and filtered data sections for receiver 100 on the eastern side of Moresby Island. The horizontal axis indicates shot-receiver distance in km. 29 Spectral analysis of the airgun data was performed and a suitable bandpass filter was tested. No significant improvement in the signal resolution of the data was achieved, however, so the data remain unfiltered. 2.3.6 Amplitude corrections Ampl i tude variations from trace to trace due to different charge sizes and receiver responses were expected. O'Brien (i960) determined that the amplitude of a TNT -generated wave varies as W2IZ where W is the weight of the charge in pounds. Ac-cordingly the trace amplitudes were multipl ied by W~ 2/ 3 to account for the charge size differences. This process overcorrected large shot amplitudes for close (<30 km) shots so an addit ional correction was applied to make large shot amplitudes consistent with the amplitudes of adjacent small shots for near distances. The explosion data were also corrected for spherical spreading by mult iplying the trace amplitudes by d2 where d is the shot-receiver distance normalized to 200 km. This correction enhanced the farther arrivals with respect to the closer ones. A spreading correction of d1 where d is normalized to 20 km gave best results for the airgun data. Ampl i tude variations due to different receiver responses were corrected by relating the velocity sensitivities for the various instruments and computing the necessary gain factors to apply to the data. The application of the above processing methods resulted in data of generally high signal-to-no ise ratio in a format ready for interpretation. 30 C H A P T E R I I I D A T A A N A L Y S I S 3.1 I n i t i a l O b s e r v a t i o n s a n d I n t e r p r e t a t i o n Common receiver gathered sections of the airgun and explosion data are shown in appendices A and B. The filters and amplitude corrections mentioned in the previous chapter have been applied as noted with each section. Further discussion of these as well as other explosion data from the experiment may be found in an open file report (Clowes, 1984). 3.1.1 T h e a i r g u n d a t a set The three airgun sections shown in appendix A represent a single profile west of receiver 010 ( O B S W l W E S T ) and a reversed profile between receivers 010 and 020 (OBSW1 E A S T and OBSW2) . The data have been plotted with a reducing velocity of 4.0 km/s; hence a line parallel to the distance axis is equivalent to an apparent velocity of 4.0 km/s. The oscillatory waveforms heading diagonally away from the origin on all three sections are the direct water wave arrivals. The non-linearity of their traveltime characteristic is a result of the highly irregular topography in the region and causes them to interfere with early refraction arrivals. A t least two refraction branches can be seen on each section. A n early branch which interferes with the irregular direct wave arrivals on O B S W l W E S T intersects sharply wi th a strong second refraction branch that continues linearly across the section to the max imum shot-receiver distance of 15 km. The transition between refraction branches is much smoother for the other two sections and the gentle curvature of OBSW2 ' s arrivals suggests several ever steeper branches are involved. The refraction arrivals for reversed profiles O B S W 2 and O B S W l E A S T undergo a severe amplitude loss just past 3 km 31 and 10 km model distance respectively; this corresponds to 10 km shot-receiver in both cases, suggesting a common loss mechanism. The location of arrivals after this distance is uncertain. Addit ional ly the arrivals on O B S W l E A S T show amplitude fluctuations over their entire length suggesting structural complications. 3.1.2 T h e t h r e e - l a y e r case A preliminary velocity-depth interpretation of the airgun data sections was done using three-layer horizontal interface equations (Sheriff and Geldart, 1982) which have been rederived to consider a surface source and a receiver positioned at the first interface (figure 13). Refraction along first interface: hi x — hi tan 0\ t = + -V^cosflx V2 x h\ x — 1 COS V\ — h t ] V2 Vi V2 1 Vxtx hy -cos 9i Refraction along second interface: t -hi + 2h2 x — hi tan <j>i — 2h2 tan 02 Vi cos 4>i V2 cos 92 x hi h2 x — + — cos 4>i + 2 — cos 62 = rr + t2 2cos f l 2 ^ 2 - ^-COS^!^ 32 where t =travelt ime of ray, hn =thickness of layer n, Vn ^velocity of layer n, tn =inter-cept of n-th refraction branch at zero offset, x = horizontal component of shot-receiver distance, and 0n ^crit ical angle for layer n. F i g u r e 13. Ray paths for critically refracted waves in a three-layered medium travelling from a surface source to a receiver at the first interface. The equations assume horizontal interfaces; this assumption is invalid for interface 1 (water—sediment) because of the known irregular topography. However, the depth estimates wil l provide at least a general structure which can later be refined through additional interpretation. Layer thicknesses were computed for all three data sections by measuring the veloc-ities and times on the sections and using Vt = 1.49 km/s as water velocity. The results 33 PROFILE Vi hy v2 h2 v3 OBSWl WEST 1.49 2.24 2.29 1.7S 4.4S OBSWl EAST 1.49 1.14 1.64 1.20 4.00 0BSW2 1.49 0.44 2.05 1.01 8.84 Table III. Velocity and layer thickness estimates from three-layer interpre-tation of airgun data. are summarized in table III. These values wil l form the preliminary model for further interpretation. 3.1.3 The explosion data set The gathered sections for receivers 010 to 100 found in appendix B have been plotted w i th an 8.0 km/s reducing velocity. It should be noted that the high amplitudes resulting from large shots at near offsets and close shots in general have been over—compensated for in the correction routines and hence their relative amplitudes are misrepresented in these sections. A l l sections show distinct primary arrivals; some also have lower velocity higher ampl i tude secondary arrivals. Receiver 010 recorded both sets of arrivals across the entire section. Noise obscures arrivals at receiver 020 beyond 100 km shot-receiver distance; both refraction branches are present to that point. Receivers 030 to 100 recorded low amplitude first arrivals for shot-receiver distances greater than about 85 km and much higher amplitude first and/or second arrivals for distances less than 85 km. 34 3.1.4 First break analysis Figure 14 is a plot of first arrivals for shots 1 to 16 recorded by receivers 010 to 100. First arrival times were determined from unfiltered data sections used in conjunction with the filtered sections which more clearly showed overall phase coherency. Shot positions are fixed for all arrivals and receivers are located at various distances to the right (east) of the plot. P lott ing the arrival times above fixed shot positions provides a "common factor" for easy comparison of traveltime curves. Traveltime differences resulting solely from increased receiver offset wil l delay first arrival times but will not alter the shape of the curve. Structural variations encountered by rays travelling to different receivers wi l l however be reflected as changes in the shape of the curves. Several trends between curves immediately become apparent. There is an average time delay of 0.6 s between arrivals from shots 4 and 5 at all receivers. This indicates a major lateral change in subsurface structure beneath these points which are located above the outer edge of the terrace. The first arrival curves are all approximately parallel w i th the exception of receiver 010 ( O B S W l ) which has a much lower apparent velocity. Energy travelling east of O B S W l towards the other receivers appears to be passing through material of different composition. The remaining arrival curves successively increase in time with increasing receiver distance except for arrivals at receiver 020 (OBSW2) which are delayed by a minimum of 0.2 s. This delay suggests the existence of lower velocity material beneath O B S W 2 than beneath the Queen Charlotte Island receivers. The general information provided by these curves wil l be combined with available constraints to construct the preliminary velocity structural model for interpretation. 35 60.0 75.0 90.0 D I S T A N C E 105.0 ( K M ) 120.0 135.0 F i g u r e 14. First arrivals for shots 1—16 at receivers 020 to 100 and shots 1-18 at receiver 010. Shot locations are fixed to their positions in the final model and receivers are located to the right (east) of the plot. Distance corrections have offset the shots for receiver 010 by 1.78 km. (See text for interpretation method.) 3.2 The Modelling Method of Interpretation The general method of interpretation used in this study involves the comparison of vertical component observed data with theoretical seismograms. A ray tracing algorithm 36 is applied to a representative velocity model and synthetic seismograms are computed from the resulting traveltimes and amplitudes. The solutions determined by this forward modell ing technique are non-unique so in this study every attempt was made to generate the most geologically reasonable model within the available constraints. 3.2.1 The modelling algorithm A n algorithm based on asymptotic ray theory (Spence, 1984; Spence et al., 1984) was used to calculate the synthetic seismograms. The ray tracing routine is an extension of the Whi t ta l l and Clowes (1979) algorithm which calculates traveltimes for refracted and postcritically reflected rays and "pseudo" head waves traced through a two-dimensional velocity model. The extension enables tracing of precritically and mult iply reflected rays through the model. Ampl i tudes are calculated by asymptotic ray theory approaches and synthetic seismograms are generated after superimposing the displacements of all arrivals at a particular distance. 3.2.2 Modelling techniques Models to be used with the ray tracing routine are constructed of polygonal blocks, each with its own velocity and linear velocity gradient. The velocity is defined on the uppermost boundary of each block and is constant along the line segment. The gradient is then oriented normal to this boundary. Rays are traced through the model and the computed traveltimes are compared to observed times. The model is changed by repositioning boundaries and blocks or varying the defined velocity and gradient. The procedure is repeated until a suitable traveltime fit is attained. The calculated amplitudes must also agree with the observed data; a 37 model that fits the traveltimes may not provide a good amplitude match. The forward modell ing procedure must thus be continued until satisfactory agreement with both the traveltimes and amplitudes is obtained and the model meets the init ial constraints. 38 C H A P T E R IV I N T E R P R E T A T I O N OF A I R G U N D A T A 4.1 Initial Constraints The airgun line A G 2 was shot over the two OBSs to provide information on the sediment and upper crustal structure of the terrace west of the Queen Charlotte Islands. Several continuous seismic reflection profiles (CSPs) in the region (Chase and Tiffin, 1972; Chase et al., 1975; Davis and Seemann, 1981) indicate that the terrace subsurface is composed of faulted and tightly folded sediments partially infilled with stratified sediments showing increasing deformation with depth. Analyses of terrace sediment velocities south of the present study area (Kirstiuk, 1981; Tjaden, 1981) indicate extreme lateral inhomogeneity in the shallow subsurface above 4 km depth. A structural model developed for a seismic refraction profile west of southern Moresby Island (Horn, 1982; Horn et al., 1984) provides structural and velocity information for the deeper structure of the terrace region. The velocities and layer thicknesses estimated from the data sections (table III) and constrained by these other studies were combined with the bathymetry profiles to form init ia l models for use with the ray tracing routine. Continuity between receivers was provided by the Horn et al. (1982) model. The final composite velocity model was derived from analysis of data recorded at O B S W l for shots west of the O B S location (segment I) and from reversed data recorded at OBSs W l and W2 for shots between the two OBSs (segment II). The model is presented in figure 15 for reference during the following modell ing discussions of the individual profiles. 39 W D I S T A N C E (KM) E o 0 . 0 4.0 8.0 12.0 16.0 20.0 24.0 28.0 D I S T A N C E (KM) \ W ^ 0.0 4.0 8.0 12.0 16.0 20.0 24.0 28 I I I I I I I I F i g u r e 15. F inal velocity model from ray tracing of airgun data. Velocities are in km/s, velocity gradients are in km/s/km. The upper model shows the division into individual profile models. Section I was modelled as O B S W l W E S T and section II was modelled as O B S W 2 and O B S W l E A S T . 40 4.2 P r o f i l e O B S W 2 Sections O B S W l E A S T and O B S W 2 provide a reversed profile between O B S W l on a mid-terrace topographic high and O B S W 2 at the edge of the narrow continental shelf. The major amplitude decrease on both sections severely l imits the extent of the region sampled by both data sets; good control of depth and dip in the central portion of the model is however sti l l expected. Figure 16 shows rays traced through the final O B S W 2 model and compares the resulting traveltimes with picks from the data. The model shows a two-layer structure with low velocity sediment above a higher velocity unit. Lateral velocity variations in both layers are seen across the proposed Queen Charlotte Fault (QCF) location. First arrival time picks indicate that the sediment beneath the continental shelf is of fairly high velocity (2.35 km/s). No clear reflections from the top of the underlying structural unit were observed on the data section, suggesting a partial loss of energy transmitted through the sediments. Several flat-lying units of increasingly higher veloc-ity were thus used to model the sediment layer, although it should be recognized that the energy loss may be from passage through deformed and faulted sediments instead of through several reflecting interfaces. The arrivals between 27 and 23 km in figure 16b are the modelled refractions through the shelf sediments with some later reflected arrivals from the underlying unit. A lower velocity (1.65-1.80 km/s) was required by traveltimes for the deformed sediments of the terrace itself. The complex structure observed on the CSPs was not compensated for in the model. High velocity gradients are required for the sediments in both regions. 41 15.0 21.0 D I S T A N C E F i g u r e 16. Ray tracing model (a) and traveltimes (b) for profile OBSW2 . Computed traveltimes are marked as x's, observed traveltimes are connected points. The distance scale and velocity structure correspond to those given for segment II in figure 15. OBSW2 is at 28 km as indicated. 42 29.0 LINE AG2 OBSW2 (b) F i g u r e 17. Seismograms for profile OBSW2 . (a) synthetic seismograms corresponding to figure 16. (b) vertical component observed data. The source wavelet used for all airgun synthetic seismograms is from data section O B S W l W E S T . Distance scales are model distances. 43 Arr ivals from 23 to 16 km were modelled as refractions through the lower structural unit, which is composed of several blocks with high-angle boundaries modelled as ver-tical faults. A good fit to the undulating observed first arrival curve (figure 16b) was achieved by varying the velocity and depth of each individual block. The faulted unit beneath the continental shelf has a velocity of 4.1 km/s and a fairly high gradient of 0.40 km/s /km. This unit is separated from the main Q C F by a narrow block of much lower velocity (3.4 km/s) and higher gradient. West of the Q C F lies a slightly deeper unit of 3.8 km/s velocity and moderate gradient. Figure 17 compares the synthetic seismograms with the processed data section. The computed amplitudes for refractions between 27 and 23 km are much higher than the observed ones. This is presumably because of interference between these refractions and direct waves which have not been modelled, and attenuation due to near-surface sediments which cannot be modelled with the program used. Modelled refractions through the deeper unit are seen as a smooth line of increasingly higher amplitude arrivals to approximately 17 km model distance. The reason for the sudden loss of amplitude for corresponding first arrivals at about 18 km model distance on the data section is uncertain. The amplitude loss was successfully reproduced with a model that incorporated a low velocity zone (LVZ) east of the Q C F below 3 km depth. Rays travell ing upwards through the LVZ were reflected back at its upper boundary. However, the thickness of the L V Z could not be determined from the available data. This factor would have introduced a large degree of traveltime uncertainty into the explosion model for O B S W 2 and created a local anomaly within the regional structure so the L V Z model was rejected. The Q C F itself remains the most probable cause of the amplitude loss. 44 4.3 Profile O B S W l E A S T The ray tracing model and traveltimes for arrivals at OBSW 1 are shown in figure 18. The model shown is the same as previously discussed for O B S W 2 . The reversed receiver position provides addit ional constraints for the western and central subsurface structure. Arr ivals to 20 km model distance have been modelled as refractions through the upper sediment layer. A single unit of velocity 1.65 - 1.80 km/s and high gradient (0.53 km/s/km) was used to represent the sediment layer. The sediment deformation seen on CSP sections could not be duplicated at the scale of the model but is expected to reduce amplitudes of refractions through the sediments and reflections from the underlying unit. This may explain why the tripl ication branch (a to b in figure 18b) from modelled reflections is not observed on the data section. Refractions through the lower structural unit generate the remaining first arrivals east of 20 km distance. The O B S W l E A S T data section and corresponding synthetic seismograms are shown in figure 19. Ampl i tudes for rays refracted through the upper sediment to 19 km distance are higher than those observed on the data as expected. Arr ivals for rays refracted through the lower unit show uniformly low amplitudes on the synthetic seis-mogram plot. The corresponding arrivals on the data section have moderately high amplitudes to 22.3 km distance followed by a sudden decrease to the modelled ampli-tude level. The arrivals continue with low amplitude to 24.5 km where they appear to stop. Very weak arrivals between 25 and 27 km (figure 20) may be the attenuated continuation of these low amplitude refractions. A band of high amplitude arrivals with concave upward characteristics is seen on the data section between 22.5 and 25.5 km (figures 19b and 20). Their nature suggests 45 F i g u r e 18. Ray tracing model (a) and traveltimes (b) for profile O B S W l E A S T . O B S W l is at 15 km model distance as shown. 46 F i g u r e 19. Seismograms for profile O B S W l E A S T , (a) synthetic seismo-grams corresponding to figure 18. (b) observed data. The receiver is located at 15 km. 47 LINE AG2 OBSWl EAST F i g u r e 20 . Detail from figure 19 showing low amplitude arrivals. that they may be reflections but a suitable reflector position could not be modelled. Similarly, the inclusion of multiply-reflected rays in the model did not provide the necessary traveltime match. The arrivals have tentatively been attr ibuted to a local focussing of rays. 48 4.4 Profile O B S W l W E S T Profile O B S W l W E S T is a westward continuation of the previously modelled data recorded at O B S W l . The profile is unreversed, affording no dip control, but lateral continuity of segment II model features (figure 15) across the receiver location will provide velocity constraints. A two-layer structure has again been used to reproduce the data traveltimes (figure 21). The upper layer is a continuation of the sediment unit from the O B S W l E A S T model and as before does not compensate for sediment deformation. The lower layer has the same velocity as the adjacent unit in the O B S W l E A S T model but is located an average 1 km deeper to satisfy the traveltimes. A slight eastward dip is required for this unit; a horizontal interface and higher velocity would also have reproduced the traveltimes but would have required a lateral velocity discontinuity between sections. Ray tracing has duplicated the observed traveltimes well (figure 21b). Refractions through the upper sediment layer generated first arrivals from 12 to 8 km model distance. A narrow tr ipl ication branch from the modelled interface reflections (a to b in figure 21b) was not observed on the data section; the discrepancy results from the difficulty in representing deformed sediments in the model. The interface may also be of a more gradational nature than the sharp boundary in the model; the data did not enable such fine stratigraphic detail to be determined. A small vertical offset and change in dip of the lower unit reproduce the slight bend at 4.5 km in the smooth first arrival time curve. The synthetic and observed seismograms for this section are shown in figure 22. Upper layer refractions again show larger amplitudes on the synthetic data than on the observed set for the same reasons mentioned earlier. As just discussed, the synthetic 49 DISTANCE (KM) 0.0 2.0 4.0 6.0 8.0 10.0 12.0 14.0 16.0 16.0 F i g u r e 21 . Ray tracing model (a) and traveltimes (b) for profile O B S W l W E S T . O B S W l is at 15 km model distance. The section between 15 and 16 km overlaps the previous model (figure 18). The velocity structure and distances correspond to section I in figure 15. 50 16.0 (b) F i g u r e 22. Seismograms for profile O B S W l WEST , grams corresponding to figure 21. (b) observed data, km distance as indicated. (a) synthetic seismo-The receiver is at 15 51 reflection branch is not seen on the data section but could not be avoided in the model. Refractions through the deeper layer are seen as a smooth set of increasingly higher ampl i tude arrivals extending to the end of the synthetic section and match the observed amplitudes well. 4.5 T h e F i n a l M o d e l The composite model combining the results for the entire airgun data set was shown in figure 15. The two-layer structure was interpreted as a variable sediment layer over-ly ing a complex lower unit. Sediment thicknesses range from 0.4 to 1.6 km and velocity increases rapidly with depth. The lower unit shows a graben- and horst-l ike feature and has different velocity structure on either side of the Queen Charlotte Fault. Am-plitude loss on profiles O B S W 2 and O B S W l E A S T suggest that an unmodelled, highly attenuative feature may exist in the inner terrace region. An attenuation zone located at depths greater than 3 km along the vertical fault would satisfy the requirements. Such a zone might be due to extensive shearing and deformation along the fault plane and could permit passage of energy from lower frequency explosive sources while atten-uating the higher frequency airgun energy. Model l ing of the explosion data set should provide a look at the deeper structure and inhomogeneities of the terrace region. 52 C H A P T E R V I N T E R P R E T A T I O N O F E X P L O S I O N D A T A 5.1 I n i t i a l C o n s t r a i n t s The explosive shots were detonated along a 110 km profile extending into the Pacific Ocean from the western edge of the terrace. Energy generated by the 33 shots was recorded by the colinear array of seismographs after it passed through the ocean crust, terrace, and continental material beneath the Queen Charlotte Islands. Initial constraints for a velocity model of the sampled region were obtained from numerous sources. Several CSPs (Chase and Tiff in, 1972; Chase et al., 1975; Davis and Seemann, 1981) and a nearby high density reflection seismic profile (P.D. Snavely, unpublished data) provided estimates of ocean sediment thickness and seismic basement dip. Velocities of ocean sediment and upper crust were approximated from previous studies to the north (Snavely et al., 1981) and south (Clowes and Knize, 1979) of the present study area. The model developed for the airgun data (Chapter IV) defined the shallow structure of the terrace region. Constrai nts for the deeper structure were based on the Horn et al. (1984) refraction model. A study by Mackie (1985) provided initial estimates of depth to the Mohorovicic discontinuity beneath the survey region. A somewhat iterative procedure was used to develop the final model. An initial model was constructed and modified through forward modelling to fit the first common receiver data section (receiver 010). This model was then expanded and altered to fit the second data section (receiver 020) and then further modified to provide a good fit for both data sections. Modell ing continued in this manner with each subsequent model having to satisfy previous data sections within reasonable limits until the final model satisfied all sections. Model reversal was then achieved by tracing rays from various shots to all receivers through the previously modelled regions and comparing 53 the traveltimes and amplitudes with common shot record sections that show the data recorded by all receivers for each particular shot. The development of a single model degraded the computed traveltime and amplitude fits for each indiv idual data section; however, the final model satisfies the total data set so its features are considered more representative of the entire region than those of the indiv idual models. Five common receiver gathered sections were used to develop the final model: the two O B S data sections, and three of the generally similar L B S sections recorded on Moresby Island (see appendix B ) . Two common shot gathered sections for large explo-sions were also used. The final velocity model is shown in figure 23 for reference during the following discussions of individual record section modelling. The main components of the model are: (a) an oceanic segment from 0 to 114 km; (b) the terrace region from 114 to 135.8 km; and (c) the Queen Charlotte Islands from 135.8 to 200 km distance. 5.2 Receiver 010 The first data modelled were the arrivals at receiver 010 ( O B S W l ) on the terrace. The initial model was constructed as a 130 km long segment extending from the ocean to receiver 010 and incorporated the model for airgun profile O B S W l W E S T between 115 and 130 km distance. West of the terrace a uniform oceanic structure was inserted based on the velocities and depths of Horn et al. (1984). This structure was extended beneath the terrace region to the end of the model. Several problems with this init ial model became apparent when ray tracing com-menced. A much lower velocity was required for the lower terrace and an eastward d ip on the adjacent oceanic segment was necessary to sufficiently delay the computed w _o.o PA plate 54 terrace OISTRNCE IKMB-2 u 8 0 . 0 100.0 ° 120.0 O 1 4 0 . 0 J i y__i * i NA plate 160.0 ^ 160.0 200.0 VE 1:1 . 0 . 0 O 20 .0 4 0 . 0 JL 60.0 _ l 0 <*> 1 — D^ISTANCE (KM) 8 0 . 0 100.0 120.0 _ l O y o CM O T 140.0 1_ O O O O CO <o o o y y r 180.0 2 0 0 . 0 160.0 VELOCITY (km/s) ; GRADIENT (km/s/km) 1| | 1 .49; 0 001 2jj 2.20 ; 0.530 3S 2.35; 0 700 3.80 ; 0 350 5 6 3.80 ; 0.350 3.40 ; 0.500 7|p] 4.10 ; 0.400 8g 5.00 ; 0.180 9 + + 6 .70 ; 0 012 10 ft 5 30 ; o 200 11& 6 50 ; o 030 6 10 ;o 095 CO 7 90 ;o 005 F i g u r e 23 . The final velocity model derived from airgun and explosion data modell ing. The positions of shots and receivers used for modell ing are shown. The upper 5 km between 115 and 143 km correspond to figure 15. The upper figure shows the model with no vertical exaggeration. 55 traveltimes to match the observed values. These and other parameters were varied until satisfactory traveltime and amplitude fits were attained and ultimately a final model was produced. Figure 24 shows the final ray tracing model for receiver 010 and compares the com-puted traveltimes with the observed times. The earliest arrivals at near offsets are refractions and reflections from the sediment and upper crust of the terrace. The cal-culated traveltimes for these rays are slightly slower than the observed times. However, the structure was well defined by airgun modelling so alterations were not made to the model to improve the traveltime fit. The boundary between the upper and lower crustal units was positioned so as to l imit the lateral extent of upper crustal refractions to the region observed on the data. The oceanic boundary (unit 4 in figure 23) is vertically downdropped 2 to 3 km with respect to the upper terrace (unit 5) to match the ob-served time delay of 1 s at 110 km. Traveltimes for refractions through the oceanic crust require a thickening and eastward dip of the upper crust and overlying sediment between 80 and 113 km. This approximately corresponds to the location of the Queen Char lotte Trough observed on bathymetry (figures 24a and 17). A low gradient (0.012 km/s /km) is required for the lower oceanic crust (unit 9 in figure 23) to permit refrac-tions to travel across the entire region; the velocity was adjusted accordingly to provide the best traveltime fit. The refractions appear as secondary arrivals (a to b in figure 24b) of apparent velocity 6.8 km/s across most of the section. Reflections from the Moho (arrivals c to d) constrained its depth between 70 and 113 km distance. A dip of 2.3° to 4.5° is required in this region. No depth or dip infor-mat ion is available west of 60 km so the Moho and shallower interfaces were extended horizontal ly to the end of the model. The effect of adding a component of eastward dip to these units, especially the Moho, would be an increase in the apparent velocity of 56 57 140 Receiver 010 (0.5-15.0 Hz) (b) F i g u r e 25. Seismograms for receiver 010. (a) synthetic seismograms cor-responding to figure 24. (b) observed data. The three-cycle source wavelet used for the synthetic seismograms was taken from the data. The receiver is located at 130 km model distance. Low amplitude mantle refractions are indicated with arrows in (b). 58 mantle refraction arrivals. These arrivals (e to f in figure 24b) presently agree well with the observed arrivals of apparent velocity 7.9 km/s. The synthetic seismograms and observed data for receiver 010 are shown for compar-ison in figure 25. Rays which travelled through the upper terrace region have generally moderate amplitudes that are slightly lower on the data section for presumably the same reaons mentioned earlier. The amplitudes of ocean crust refractions gradually increase across the section from their initially moderate level near the outer fault; amplitude fluctuations occur near 40 km as Moho reflections interfere with the crustal refractions. Mant le refractions have very weak amplitudes on the synthetic seismograms; they are barely visible on the data west of 40 km because of the noise level. The amplitudes of the arrivals on the synthetic seismograms increase slightly with increasing shot-receiver distance. This trend cannot be confirmed on the data section. 5.3 R e c e i v e r 020 Receiver 020 (OBSW2) was situated on the continental shelf just east of the terrace and across the Queen Charlotte Fault (QCF) from O B S W l . The initial model from receiver 010 was extended across the remaining terrace region by inserting the reversed airgun profile O B S W l E A S T — O B S W 2 (segment II in figure 15) between 130 and 143 km and assuming starting velocities for the region beneath it. Crustal velocity estimates and thicknesses were later refined by modell ing. The velocities of the sediments and upper terrace crust between 130 and 144 km were determined from the airgun data (figure 15). Ray tracing through the upper terrace region determined approximate locations for upper/lower crustal boundaries. Lower crustal velocities are not well constrained as these units represent only a small portion 59 of the rays' travelpaths. A velocity of 5.3 km/s was chosen for the block west of the Q C F (block A in figure 26) to be consistent with the adjacent terrace blocks (figure 23). This required a rather high velocity of 6.5 km/s for the block east of the Q C F (block B in figure 26) to achieve a good traveltime fit. The vertical boundaries seen within this block are used only to define velocity gradients and do not represent structural features. Three groups of rays were used to model the observed traveltimes and amplitudes. The arrivals extending from 123 to 110 km as first arrivals represent refractions through the upper terrace. Secondary arrivals of apparent velocity 7.2 km/s (a to b in figure 26b) represent crustal refractions through the lower terrace and ocean crust. The time delay at 110 km again clearly shows the transition from terrace to oceanic structure. The refractions through the ocean crust in the model end abruptly at 69 km when the rays encounter the Moho. Reflections from the Moho are focussed into a narrow group of secondary arrivals between 102 and 112 km (c to d in figure 26b). The arrivals indicate a dip of 19° for this portion of the Moho. Refractions through the mantle are the first arrivals west of 100 km and extend across the modelled section with an apparent velocity of 8.1 km/s (arrivals e to f in figure 26b). The synthetic seismograms and observed data for receiver 020 are presented in figure 27. The large amplitudes of the crustal refractions terminate abruptly at 69 km on the synthetic seismograms. A fairly low signal-to-noise ratio makes it difficult to locate the refraction arrivals on the data section beyond 90 km but they appear to continue to perhaps 55 km. Extending the modelled refractions to this distance would require a deeper Moho or lower terrace gradient, both of which would degrade the fit of the model to other data sets. Refractions through the mantle are of equally low amplitude on both sections. A n increase in amplitude with increasing shot-receiver distance is observed for the synthetic 60 D I S T A N C E (KM) o 0 0 20 0 40.0 60.0 80.0 100.0 120.0 140.0 160.0 (a) F i g u r e 26. Ray tracing model (a) and traveltimes (b) for receiver 020, located at 143.1 km model distance. Numerous vertical lines in the model define velocity gradient directions and are not structural features. 61 o oo" U ° LlJcd" o C O o rsi" 0.0 20,0 T 40.0 60.0 80.0 D I S T A N C E (KM) 100.0 120.0 140.0 (a) 20 40 60 80 100 Receiver 020 (0.5-10.0 Hz) 120 140 (b) F i g u r e 27. Seismograms for receiver 020. (a) synthetic seismograms corre-sponding to figure 26. (b) observed data. The receiver is located at 143.1 km model distance. The arrows in (b) indicate low amplitude mantle refractions. 62 arrivals; the noise-laden signal on the data section unfortunately limits observations of ampl i tude characteristics at far offsets. 5.4 Receiver 030 Receiver 030 is a Backpack recorder located near the western shore of Moresby Island. Rays travell ing to the receiver from all shots must pass through the terrace and the continental crust beneath the island. The velocity structure of the island is poorly denned as few studies have sampled the region. Extensive mapping of surficial geology provided the basis for a stratigraphic and structural analysis of the Queen Charlotte Islands (Sutherland Brown, 1968). Several geologic cross sections were constructed in a southwest-northeast direction across the islands and one of these is fortunately located within 3 km of the LBSs on Moresby Island. This profile (figure 28) can be divided into three major sections. West of the Louscoone Fault the subsurface is composed primari ly of Triassic lavas of the Karmutsen Formation overlain in regions by Upper Triassic limestones of the Kunga Formation. Severe deformation and faulting has occurred between the Louscoone and Rennell Sound Faults. Near-surface units within the fault zone are primarily Jurassic and Cretaceous in age and generally consist of volcanics of the Yakoun Formation, calcareous siltstone of the Longarm Formation, and sandstones of the Haida and Honna Formations. East of the Rennell Sound Fault lie volcanics of the Yakoun Formation which are overlain by Haida Formation sandstones in the western half. Vertical offset across the Rennel l-Louscoone Fault zone is estimated as 600 to 3000 m with the east block dropped down (Sutherland Brown, 1968). The stratigraphic interpretation of the profile was used to obtain crude estimates of upper crustal thickness and velocity for the model. 63 S o l l i t Point BOTTLE HARBOUR MITCHELL INLET OOUGLAS INLET UUOGE INLET 4 000 ' -TM KHo "EEL INLET 040 x"RKu "RKA Louscoone F a u l t ; GILLATT ARM 060 070 050 " R K u . J Y K L K L K L K H o KHA R e n n e l l Sound F a u l t SKIOEGATE LAKE C O P P E R R I V E R T M "RJKo U KA R J K U "RKA Sandspit Fault COPPER BAY HECATE STRAIT 100 - 4 0 0 0 S e a L e v e l J Y 2 m i T E R T I A R Y TM Ma s s e t t Fm: b r e c c i a s , columnar b a s a l t s , r h y o l i t e s C R E T A C E O U S A N D T E R T I A R Y KT Post T e c t o n i c P l u t o n s : d i o r i t e , q u a r t z , g r a n i t e C R E T A C E O U S Queen C h a r l o t t e Group KS S k i d e g a t e Fm: s h a l e , s i l t s t o n e , sandstone KHo Honna Fm: conglomerate, f e l d s p a t h i c sandstone KHA Haida Fm: s h a l e , sandstone KL Longarm Fm: c a l c a r e o u s s i l t s t o n e , l i t h i c sandstone J U R A S S I C JY Yakoun Fm: v o l c a n i c sandstone, t u f f , s i l l T R I A S S I C A N D J U R A S S I C JK (, Kunga Fm: f l a g g y a r g i l l i t e T/jKfj Kunga Fm: f l a g g y and massive l i m e s t o n e T R I A S S I C TpKA Karmutsen Fm: l i m e s t o n e , p i l l o w l a v a s , f l o w s F i g u r e 28. Geologic cross-section across Moresby Island near the receiver locations (after Sutherland Brown, 1968). Approximate receiver positions for this experiment are indicated. 64 D I S T A N C E (KM) (a) oo F i g u r e 29. Ray tracing model (a) and traveltimes for receiver 030, located at 154.5 km model distance. 65 o 00 140.0 0 20 40 60 80 100 120 140 Receiver 030 (1.5-8.5 Hz) F i g u r e 30. Seismograms for receiver 030. (a) synthetic seismograms corre-sponding to figure 29. (b) observed data. The receiver is located at 154.5 km model distance. 66 The ray tracing model and traveltime plots are shown in figure 29. Three groups of rays were again involved in the interpretation. Refractions through the lower crust of the terrace are seen as first arrivals from 122 km until the outer fault offset at 110 km. Arr iva ls from 108 to 87 km with apparent velocity 7.6 km/s (a to b in figure 29b) are from rays that continued through the fault into the lower oceanic crust until they were stopped at the Moho. Reflections from the Moho are again limited to a branch of secondary arrivals from 110 to 114 km (c to d) and confirm the previously determined Moho dip of 19°. Deeper travell ing rays were refracted at the Moho and travelled through the mantle to form the 8.2 km/s apparent velocity branch (e to f) that intersects the lower crustal refractions at 103 km on the traveltime plot. The synthetic and observed seismograms are shown in figure 30. The high amplitude arrivals seen on the data between 110 and 90 km correspond well to the modelled refractions through the lower oceanic crust. Addit ional energy is provided by the slightly later Moho reflections between 103 and 113 km. The much weaker mantle refractions continue across both sections from 100 km and increase slightly in amplitude with distance. 5.5 Receiver 060 Receiver 060 is a Backpack recorder located midway across Moresby Island in the Rennel l-Louscoone Fault zone (figure 28). The surface geology indicates that vertical displacement of units adjacent to the fault zone has certainly occurred. However, rays arriving at receiver 060 are apparently not affected so there is no major velocity contrast across the fault at the depth concerned. 67 The ray tracing model and corresponding traveltimes are shown in figure 31. The only change from the previous model is the extension of the island structure to the receiver 060 location. The fault and vertical offset shown at 172 km have no effect on ray paths or traveltimes and were included in the final model to satisfy the receiver 100 data discussed in the next section. The results of the ray tracing are very similar to those for receiver 030. Crustal refractions, Moho reflections, and mantle refractions are again the predominant ray groups involved. The high terrace gradient and Moho depth have further restricted the range to which crustal refractions can occur and all arrivals modelled west of 100 km are mantle refractions. The eastward dip of the Moho between 60 and 110 km is clearly evident in these refraction arrivals as they show an unrealistically high apparent velocity (8.8 km/s) in this region before assuming the more reasonable 7.9 km/s for the horizontal region. The seismograms (figure 32) show the concentration of high amplitude crustal re-fraction arrivals east of 98 km and the continuous low amplitude mantle refractions extending westward across the section. Arr ivals on the data section are very clear to large shot-receiver distances, especially for the big shots, and show well an amplitude increase with distance of the theoretical mantle refractions. 5.6 Receiver 100 The last common receiver gather modelled was for receiver 100, a microcorder lo-cated near the eastern shore of Moresby Island. Rays travelling to the receiver from the shots wil l be passing through lower crustal material east of the Rennell-Louscoone Fault zone and will be sensitive to the thickness of the crustal blocks. Combining the F i g u r e 31 . Ray tracing model (a) and traveltimes (b) for receiver 060, located at 174.8 km model distance. 69 0.0 20.0 60.0 D I S T A N C E 80.0 (KM) 100.0 120.0 140.0 (a) 60 80 Receiver 060 (2.0-12.0 Hz) 140 (b) F i g u r e 32. Seismograms for receiver 060. (a) synthetic seismograms corre-sponding to figure 31. (b) observed data. The receiver is located at 174.8 km model distance. 70 constraints of this model with those of receiver 060 should provide an estimate of vertical displacement at depth across the R L F zone. The island structure of the receiver 060 model was extended uniformly to receiver 100; the upper crustal structure east of the R L F was then varied in subsequent mod-elling. The model that provides the best overall fit is shown with corresponding travel-times in figure 33. Traveltimes for crustal refractions, seen on the time plot between 115 and 104 km, are satisfied by downdropping the easternmost crustal block 0.4 to 2.0 km. The best fit for all arrivals was achieved with a vertical displacement of 1.1 km. The remaining arrivals are attributed to mantle refractions and a narrow band of Moho reflections between 105 and 116 km (c to d in figure 33b). The refraction arrivals (e to f) have an apparent velocity of 11.1 km/s for the dipping Moho region and 8.06 km/s through the horizontal boundaries to the west. The amplitudes of the synthetic seismograms (figure 34) match the observed seis-mograms well. The crustal refractions yield high amplitudes which are augmented by the Moho reflections. Low amplitude mantle refractions lead away from the the high ampl i tude arrivals starting at 105 km and continue westward across the section with slightly increasing amplitudes. 5.7 Shot 13 Shot 13 was a 540 kg explosion detonated approximately 48 km west of O B S W l at 81.7 km model distance. The ray tracing model and traveltimes are presented in figure 35. Rays were traced from the shot position to the surface through the previously F i g u r e 33 . Ray tracing model (a) and traveltimes (b) for receiver 100, located at 194.3 km model distance. 72 CO o o J n 1 1 i i i i 0.0 20.0 40.0 60.0 80.0 100.0 120.0 140.0 D I S T A N C E (KM) (a) 140 Receiver 100 (2.0-120 Hz) (b) F i g u r e 34. Seismograms for receiver 100. (a) synthetic seismograms corre-sponding to figure 33. (b) observed data. The receiver is located at 194.3 km model distance. 73 derived model. O B S positions were time-corrected to the surface for easy comparison of observed and computed traveltimes. The terrace region was well sampled at all depths by the various ray groups, which consist of refractions through the oceanic crust and terrace, reflections from the oceanic Moho, and refractions through the mantle that surface through the terrace and island crust. The time plot shows good agreement between the theoretical and observed trav-eltimes. The computed first arrivals at all receivers are from rays refracted through the mantle. The secondary arrivals picked for receivers 010 and 020 (the OBSs) correspond to Moho reflections and crustal refractions respectively. The most distant reflection from the Moho surfaced at approximately 138 km so O B S W l is the only receiver in this model that could record Moho reflections from this shot. Figure 36b shows the observed data recorded by the ten receivers. Ampl i tude cor-rections were made for the different types of seismographs (Chapter II) but did not consider the siting of the instruments. While the LBSs were carefully sited on bedrock, the OBSs were presumably dropped into soft ocean bottom sediments. The resulting poor coupling would reduce the amplitudes of the recorded signals and make theoretical amplitude comparisons difficult. The synthetic seismograms are presented in figure 36a. The high amplitude arrivals between 120 and 135 km are from Moho reflections. No corresponding high amplitudes were observed on the data recorded at receiver 010 ( O B S W l ) . Similarly, the moderate ampl i tude crustal arrivals seen on the synthetic seismograms between 140 and 150 km were not observed on the OBSW2 data. Presumably the generally low amplitudes of these two observed seismograms are a result of their location on the soft ocean bottom. The mantle refractions observed across the remainder of the section generally match the arrivals recorded by the LBSs. 74 F i g u r e 35. Ray tracing model (a) and traveltimes (b) for shot 13, located at 81.7 km model distance. Receiver locations are indicated. 75 LUio" LT) O CO a . _ ! ( ! ! 120.0 130.0 140.0 150.0 160.0 170.0 D I S T A N C E (KM) 180.0 190.0 200.0 (a) 120 130 140 150 160 170 Shot 13 (2.0-12.0 Hz) 180 190 200 •(b) F i g u r e 36. Seismograms for shot 13. (a) synthetic seismograms correspond-ing to figure 35. (b) observed data, from receiver 010 at left to receiver 100 at right. The shot is located at 81.6 km model distance. 76 5.8 Shot 25 Shot 25 was a more distant 540 kg explosion detonated approximately 91 km west of O B S W l at 38.4 km model distance. Rays traced from this shot to the surface near the 10 receivers sampled the ocean crust west of the terrace, the terrace itself, and the upper mantle and island crust. The model and traveltime plot are shown in figure 37. Three groups of rays were again used to reproduce the observed traveltimes. The first group was refracted through the oceanic crust and travelled near-horizontally before being refracted sharply upwards to the surface by the high velocity gradient within the terrace. These rays are seen as secondary arrivals between 124 and 138 km on the traveltime plot (a to b in figure 37b). The second ray group was reflected from the shallowly dipping Moho west of the terrace and travelled upwards through the oceanic and terrace regions. The corresponding traveltime arrivals begin at 80 km and become successively later to join the crustal refraction arrivals at 133 km (c to d). The final group represents rays that were refracted into the mantle at the oceanic Moho and were subsequently refracted back to the surface through the various crustal structures at the dipping Moho beneath the terrace. These rays are the first time arrivals between 100 and 200 km (e to f) and fit the observed traveltimes well. The "V" shape of the arrivals is due to the dip of the Moho and the velocity structural contrast between the terrace and the island. The synthetic and observed seismograms are shown in figure 38. The low overall ampl itudes and the noise on the receiver 010 and 020 traces makes amplitude compar-ison difficult but there appears to be good general correlation between the sections. Crustal refractions and Moho reflections are the source of the high amplitude secondary arrivals between 120 and 135 km. The reduced amplitude on the synthetic trace at 130 F i g u r e 37. Ray tracing model (a) and traveltimes (b) for shot 25, located at 38.4 km model distance. 78 LlJlD" CD O 00 — l 140.0 ~l 190.0 120.0 130.0 150.0 160.0 D I S T A N C E (KM) 170.0 180.0 200.0 (a) W q C O G 120 130 140 150 160 170 Shot 25 (2.0-12.0 Hz) 180 190 200 F i g u r e 38. Seismograms for shot 25. (a) synthetic seismograms correspond-ing to figure 37. (b) observed data, receivers 010 to 100. The shot is located at 38.4 km model distance. 79 km is the result of destructive interference between the two types of arrivals. Low am-plitude mantle refractions can be followed across the section and reproduce the observed amplitudes quite well. 5.9 T h e F i n a l V e l o c i t y M o d e l The final model based on the complete ray tracing interpretation was presented in figure 23. The model can be divided into three major structural components (ocean, terrace, and island) which will be discussed individually. The region extending from the beginning of the model to the edge of the terrace at 114 km comprises the oceanic segment. A n average 1 km of sediments (unit 2) overlie a 6.2 km thick oceanic crust which has been subdivided into a lower velocity 1.3 km thick upper unit (unit 4) and a higher velocity lower unit (unit 9). A low velocity gradient (0.012 km/s/km) is required for unit 9 to enable crustal refractions from the distant shots to reach all the receivers. Beneath the crust lies the oceanic mantle (unit 13) with a velocity of 7.9 km/s. West of 60 km a horizontal Moho at 10 km depth satisfied the available data. East of 60 km the Moho begins to dip towards the mainland, first at a gentle 2.3° and then at 4.5° to the edge of the terrace. The depth of the Moho in this region is well constrained by reflections and the cessation of crustal refractions. The terrace region extends from 114 km to the Q C F at 135.8 km. Al though it too is composed of three major units, it is structurally much more complicated than the oceanic region. The uppermost layer consists of low velocity deformed sediments of irregular surface topography and variable thickness. The upper crustal unit has been separated into three blocks (unit 5) by vertical faults to satisfy the airgun data, although 80 a more complicated structure may be present. The vertical extent of the faults into the adjacent units could not be determined; surface topography however suggests that the fault at 130 km extends to the ocean floor. The lower crustal unit (unit 10) has an anomalously low velocity and fairly high gradient as required by the crustal refractions. It has been modelled as a homogeneous unit although faults may complicate the structure. The Moho dips beneath the terrace at 19° as determined from mantle refractions and a few unreversed Moho reflections. The Moho dip and mantle depth are dependent on the velocity structure of the terrace and would vary if it were altered. Good agreement is found, however, with the 20° dip determined by Horn et al. (1984) off southern Moresby Island, although their model was more poorly constrained at this level than that of the present study. Mackie (1985), in a seismic refraction study of the Queen Charlotte Islands/Hecate Strait region, de-termined a Moho dip of 5° extending from the outer edge of the terrace to the mainland of Br i t ish Columbia and a 2° dip beneath the oceanic crust to the west. The simplified velocity structure that he used for the terrace and QC Islands will permit small varia-tions in these values, but the study nonetheless indicates a shallower regional Moho dip than the 19° value obtained in this study for the limited region beneath the terrace. This suggests that the steeper dip beneath the terrace may be local and the dip may decrease beneath the Q C Islands. The island region east of the Q C F to 200 km represents the edge of the continental structure of North America. Sampling density decreases towards the end of the model and the structure can only be roughly approximated. The region adjacent to the Q C F appears distinct from the other island structure at all model depths. A 1 km thick unit of stratified, 2.35 km/s sediments (unit 3) overlies the crustal units. The upper crust again appears to be faulted and is divided into two distinct velocity units at 137 km. 81 The division has been modelled as a vertical fault but is not so constrained. The upper crust (unit 7) has been modelled with a slightly higher velocity (4.1 km/s) than the oceanic crust (3.8 km/s) with the exception of the 1.2 km wide sliver next to the Q C F (unit 6) which has a lower velocity (3.4 km/s). A higher velocity and lower gradient were determined for the lower crust here (unit 11) than for the adjacent lower continental crust (unit 12). However, these values are not well constrained and the two units may be much more similar. The remaining island structure has been roughly divided into two units based on stratigraphic information. The upper unit (unit 8) has a velocity of 5.0 km/s at the surface and moderately high gradient. There is l ittle velocity difference at the 5 km deep boundary between units but the gradient has been significantly reduced for the lower unit (unit 12). A modelled offset on the Rennell-Louscoone Fault has displaced the eastern half of the upper unit 1.1 km downwards with respect to the adjacent part. The fault dip cannot be resolved to better than ± 30° nor the vertical displacement to ± 0.8 km. The mantle was not sampled beneath the island and total crustal thickness is therefore unknown. A geologic interpretation of the velocity structure is necessary to fit the derived model into the tectonic framework for the region. The next chapter wi l l attempt to do this and answer some of the questions raised at the beginning of the study. 82 CHAPTER VI DISCUSSION AND CONCLUSIONS The application of two-dimensional ray tracing techniques has produced a velocity structural model for the region of the Queen Charlotte Fault zone. The fault—bounded terrace does not appear to be related to either of the adjacent crustal units but seems structural ly distinct. A geologic interpretation of the velocity model wil l help to clarify the relationship between the observed structure and the present tectonic regime. 6.1 Stratigraphic Interpretation The stratigraphic interpretation of the velocity units is summarized in table IV. Unit 2 is the sedimentary layer covering the terrace and ocean crust. Ocean bottom sediments west of Dixon Strait were interpreted as Late Miocene to Quaternary deposits (Snavely et al., 1981). Chase and Tiffin (1972) identified trench infill as Late Tertiary pelagic sediments overlain by well stratified Pl io-Pleistocene turbidites and recent turbidity current deposits. Uni t 3 is a higher velocity sedimentary unit underlying the narrow continental shelf east of the active fault and may be of terrigenous origin. Units 4 and 9 represent upper and lower oceanic crust respectively. Ophiol ite studies indicate that crustal layer 2 (unit 4) consists of an extrusive sequence of pillow lavas and basaltic sheet flows that grades down to dense dike swarms (Christensen and Smewing, 1981). A n associated high velocity gradient is related to the increase of metamorphic grade wi th depth. Layer 2 velocities range from 3.7 to 6.5 km/s and average 5.5 km/s (Christensen and Salisbury, 1975). The difference between ophiolite sequence velocities of 5.0 to 6.0 km/s measured in the laboratory and the unit 4 model velocities of 3.8 to 4.3 km/s suggests the presence of large fractures within the uppermost crustal section which would significantly lower observed velocities (Christensen and Smewing, 1981). 83 V E L O C I T Y UNIT V E L O C I T Y (km/s) G R A D I E N T (km/s/km) S T R A T 1 G R A P H I C I N T E R P R E T A T I O N 1 1.49 0.001 water layer 2 2.20 0.530 pelagic and turbiditic sediments 3 2.35 0.700 shelf sediments 4 3.80 0.350 basaltic pillow lavas and sheet flows grading to dike swarms 5 3.80 0.350 deformed upper ocean crustal basalts or compressed sediments 6 3.40 0.500 sheared sediments or basalts 7 4.10 0.400 sheared sediments 8 5.00 0.180 tholeiitic basalts, granite, and limestone 9 6.70 0.012 interlayered gabbros and peridotites 10 5.30 0.200 sheared gabbros and peridotites ?? II 6.50 0.030 continental silicic crust 12 6.10 0.095 continental silicic crust 13 7.90 0.005 harzburgite and dunite (mantle) Tab l e I V . Stratigraphic interpretation of the velocity units of figure 23. 84 Oceanic crustal layer 3 consists of interlayered gabbros and peridotites of amphibo-lite facies metamorphism. Velocities in the Oman Ophiolite range from 6.7 km/s at the top of the layer to 7.3 km/s at its base (Christensen and Smewing, 1981). The lower crustal velocity determined by modell ing (6.7 km/s for unit 9) agrees well with these laboratoy measurements. Unit 13 represents the upper mantle which is composed primarily of tectonized harzburgite and dunite (Christensen and Smewing, 1981). The modelled mantle velocity (7.9 km/s) for the 8 M a old oceanic lithosphere is lower than average mantle velocity measurements (8.15 km/s) but is in agreement with other observations of young oceanic l ithosphere (Christensen and Salisbury, 1975). Velocity unit 8 represents the upper crustal rocks of the Wrangellia terrane. The Karmutsen Formation, which outcrops at the surface between 190 and 212 km model distance, consists primari ly of Late Triassic tholeiitic pillow basalts (Sutherland Brown, 1968). Estimates of ocean basalt velocities range from 3.7 to 6.5 km/s and a velocity of 5.04 km/s has been measured for basalts at the top of oceanic layer 2 (Christensen and Salisbury, 1975). A n average velocity of 5.0 km/s is thus representative of the volcanics of the Karmutsen Formation and the younger rocks of the Yakoun Formation that outcrop to the east. Unit 12, the lower continental crust, is poorly constrained in the model. Its upper boundary is not a velocity discontinuity but instead marks a reduction in gradient from the upper crust. The reduced velocity gradient required by the ray trace modelling may signify a change in metamorphic facies or composition. The stratigraphic equivalents of the terrace velocity units are not readily apparent. The velocity of unit 5 suggests that it may be upper oceanic crust (unit 4) although de-85 formation by thickening and fault ing has clearly occurred. Alternatively, it may contain highly compressed sediments of oceanic or continental origin. Hyperbolae and disrup-tions, seismic signatures of irregular surface features, make it impossible to distinguish any subsurface structure on the available C D P data. Unit 7 may similarly be composed of sediment derived from the upper continental crust. The low velocity and position of the narrow unit 6 suggests that it has been extensively sheared and fractured by motion along the fault and may belong to either of the above mentioned crustal units. Lower crustal unit 11 east of the fault is again not well defined and is probably related to the lower continental crust, unit 12. Unit 10 , however, has a low velocity and high gradient inconsistent with the other lower crustal units. The most acceptable hypothesis is that extensive deformation by compression and shearing has altered normal lower crustal material beyond easy recognition and masked the boundary between upper and lower crustal material. 6.2 Tectonic Implications The fault structure of the velocity model is consistent with a compression al tec-tonic regime. The graben- and horst-l ike appearance of the upper terrace may in fact represent high angle reverse faults and low frequency compressive sediment folds as the fault angles are Only constrained to ± 10°. The downward vertical displacement of the oceanic crust at the outer fault (114 km distance) agrees with the previously mentioned vertical fault plane solution (Chapter I) of Berube (1985). Hypocentral lo-cations of microearthquakes (Hyndman and El l is, 1981) require a westward dip of 80° on the inner fault, consistent with the velocity model. The remaining major fault is the Rennell-Louscoone Fault at 172 km model distance on Moresby Island. Sutherland 86 Brown (1968) observed dips ranging from 75° northeastwards to vertical on the asso-ciated fault strands. This suggests a tensional mechanism. The rotation and shifting of the Queen Charlotte Islands along this fault zone during the initiation of spreading in Queen Charlotte Sound (Chapter I) would have produced the necessary tensional regime. The ongoing vertical uplift of the western edge of the islands by lithospheric flexure (Riddihough, 1982a) may produce additional vertical offset on this fault in fu-ture. Two tectonic mechanisms wil l be examined to explain the observed terrace structure: accretion associated with active subduction, and structural compression associated with convergence at a transform fault. The sharp increase in model Moho dip from 2.3° west of the terrace to 19° beneath the terrace is supportive of subduction of the Pacific plate beneath the Queen Char lotte Islands. Similar values were obtained from refraction studies for southern Moresby Island (a poorly constrained 20° - Horn et al., 1984) and for the Juan de Fuca Plate beneath Vancouver Island (14-16° - Spence et al., 1985). A lower value of 5° (Mackie, 1985) for the Queen Charlotte Islands region implies that Moho dip decreases beneath the islands but is stil l consistent with subduction processes (see also discussion in section 5.9). The slow rate of convergence across the plate margin (10 to 20 mm/a) may cause weak coupling between the subducting Pacific plate and the America plate. The ex-pected tectonic processes for such a case are: ( l) accretion by deposition of sediment and subsequent tectonic incorporation into an accretionary complex; (2) sediment sub-duct ion of material ponded in a trench axis; and (3) erosion by failure of the margin front through mass wasting and subsequent incorporation into subducting sediment (von Huene, 1984). 87 Trench sediments seen in CSP profiles (eg. Chase and Tiffin, 1972; Chase et al., 1975) are relatively undisturbed except for minor faulting so it is unlikely that process (2) is occurring on a large scale. However, tectonic erosion of the margin has almost certainly occurred. Str ike-sl ip faulting, an important erosion process in zones of oblique compressional stress (von Huene, 1984), appears to have removed some continental material as evidenced by the truncated spurs and triangular seaward-facing facets of the coastal mountains on the Queen Charlotte Islands (Keen and Hyndman, 1979). Sedimentary accretion of material from the subducting ocean crust to the front of the margin could have produced the complex structure of the upper terrace. Seismic records across the landward slope of the eastern Aleutian Trench (von Huene, 1984) reveal the presence of 3 km thick blocks of sediment, separated by steep high angle thrust faults, beneath younger sediments at the continental margin (figure 39). The lower sediments may be subducting with the oceanic crust. Accreted sediments and possibly imbricated oceanic crust were imaged beneath southern Vancouver Island by high density seismic reflection profiling. Similar structures may compose the upper Queen Charlotte terrace. The interpretation of the upper terrace as an accretionary sediment wedge would of course require the presence of subducting oceanic crust beneath it. No internal reflections from within the lower terrace were detected, providing no support for a structural division. It is quite possible, however, that the top of the subducting crust is extremely irregular as a result of strike-sl ip faulting and convergence in the terrace region. Strong reflections would not be generated in such a case and any energy would be severely attenuated by the deformed overlying sediments. The velocity structure of the ocean crust could then have been amalgamated with the overlying terrace structure to produce the single "averaged" velocity unit of the model. 88 SCALES FOR B, C F i g u r e 39. Seismic profiles across the Aleut ian Trench, i l lustrating the structures that result from obduction, imbrication, accretion, and subduction of sediments at a convergent margin (after von Huene, 1984). Alternatively, the complex structure of the terrace may result from compression associated wi th oblique convergence at the fault zone. Hyndman and El l is (1981) pro-posed a model in which faulting occurs on the Queen Charlotte strike-slip fault with the shallow portion of the thrust zone locked. Convergence is taken up either by peri-odically unlocking the upper portion of the thrust or by compressively deforming and shortening the whole oceanic lithosphere beneath the terrace. Hyndman et al. (1982) extended the transcurrent fault into the underlying oceanic lithosphere to be consistent with the microseismicity observed primarily between 10 and 25 km depth. They sug-gested that the faulting must jump seaward at intervals to remain beneath the edge of the continental crust — terrace sediment contact (figure 40). Both models require a sharp downward bend of the oceanic lithosphere beneath the margin, possibly with associated normal faulting at the outer edge of the terrace. 89 0 Deformation Q.C. Trough Front E 10-X r-a. DE 2 0 -Terrace T E R R A C E SEDIMENTS Coast = 7 Q.C. Islands Possible old fault or qc< j discontinuity ~jANlC -— — HUSr CONTINENTAL : : : : C R U S T : 3 0 - No vertical exaggeration Main Transcurrent Fault Fossil Faults ? ^ F i g u r e 40 . A possible tectonic model of the Queen Charlotte Fault zone. Oblique convergence is resolved into str ike-sl ip motion on the fault parallel to the margin and a small component of underthrusting perpendicular to the margin (Hyndman et al., 1982). Numerous earthquakes recorded in the recent microseismicity study of Berube (1985) were located in the Queen Charlotte fault zone along Moresby Island; a few were located on the terrace and outer fault scarp. The focal depths of the events extend to 25 km although most of the activity was well constrained to depths of less than 15 km. Thus, seismic activity at depths greater than 13 km would be located within oceanic crust if subducted crust is assumed to underlie the terrace in the velocity structural model of figure 24. The Queen Charlotte fault would thus have to extend into the oceanic crust, as Hyndman et al. (1982) suggested, and essentially "slice" it as transform motion occurs. The slicing would have to be discontinuous, perhaps by alternating transform 90 motion and subduction, to allow the continued movement of the oceanic material across the fault zone and beneath the continental lithosphere. Lowell (1972) modelled strike-sl ip movement with a small component of conver-gence across the fault zone using a clay cake on moving sheets of t in. The relative motion vector of the model was oriented at 15° to the fault trace, similar to the Queen Charlotte region. The material at the fault zone moved upwards to produce a welt with downward tapering wedges and upthrust margins (figure 41). The steep faults modelled for the terrace region may be upthrust faults formed in a similar manner as opposed to underthrusts formed by accretion. F i g u r e 41 . Conceptual diagram of an upthrust-bounded welt created by convergent strike-slip or transform motion. The welt has downward-tapering wedges and upthrust margins (Lowell, 1972). 91 Thrust mechanisms and near-vertical faulting were computed by Berube (1985) for two earthquake clusters at 12 and 9 km depth respectively on the Queen Charlotte Fault zone near the present refraction study area. The shallow focal depths and the clay cake models suggest that thrusting of terrace material above any subducting oceanic crust is occurr ing as a result of compression in the fault zone. The structure of the terrace can thus be attributed to two different tectonic pro-cesses which may occur concurrently or alternate as necessary as shown in figure 42a. Convergent motion may be accommodated by slow, shallow subduction of the oceanic lithosphere beneath the Queen Charlotte Islands. Sediment overlying the oceanic crust would be accreted to the seaward edge of the terrace and/or subducted beneath the terrace. Transform motion is occurring as dextral strike-sl ip on the Queen Charlotte fault and possibly on associated faults on the Queen Charlotte Islands and beneath Hecate Strait. Sedimentary and upper crustal material may be thrusting upwards in the terrace (figure 42b) as subduction is halted at the fault and compression needs to be accommodated. The compressive stress in the lower oceanic lithosphere would con-tinue to build during this phase and may force the resumption of subduction and the subsequent migration of faulting in the downgoing crust. The refraction model developed in this study supports recent subduction but sug-gests that compression and upthrusting of terrace material presently account for a major part of the component of convergence at the active Queen Charlotte fault zone. 92 O c e a n i c C o n t i n e n t a l P l a t e T e r r a c e P l a t e F i g u r e 42 . Proposed tectonic processes in the Queen Charlotte Fault re-gion, (a) subduction - sediments and upper crust deform by accretion and compression as subduction occurs, (b) transform motion - compression in sediments and crust relieved by upthrusting along the fault zone. 93 6.3 C o n c l u s i o n s The structural model developed for the Queen Charlotte Fault zone confirms an anomalous low velocity terrace region separating oceanic crust to the west and conti-nental crust to the east. The terrace crustal structure consists of a 3 km thick upper unit with velocity 3.8 km/s at the top and fairly high (0.35 km/s/km) gradient, and a lower unit with an average thickness of 10 km, velocity 5.3 km/s, and gradient 0.20 km/s / km. The terrace is bounded on the west by an old fault zone and on the east by the currently active near-vertical Queen Charlotte Fault; steep faulting also extends through at least the upper crustal unit of the terrace. Moho dip steepens considerably beneath the terrace from 5° to 19°. The structural complexity of the terrace may result from a combination of accretion and deformation by underthrusting associated with subduction, and compression and upthrust ing related to oblique convergence at the transform fault. Addit ional study is necessary to provide the structural resolution required to distinguish between the two. High resolution multichannel seismic reflection profiling over the terrace would better define the velocities and fault orientations and possibly locate the top of subducted oceanic lithosphere. Similar ly, further sampling is necessary to constrain the vertical offset of 1.1 km modelled on the Rennell-Louscoone Fault. Reversed refraction profiles across and along the islands would provide information presently lacking about regional crustal structure and would better constrain Moho depth and dip. The addition of such future observations to the existing geophysical and geological information will further the understanding of the complicated tectonics and complex structure of the Queen Charlotte Fault zone region. 94 BIBLIOGRAPHY Atwater, T., 1970, Implications of plate tectonics for the Cenozoic tectonic evolution of western North America: Geol. Soc. A m . Bul l . , 81, 3513-3536. Berube, J . , 1985, A seismicity study of the Queen Charlotte Islands/Hecate Strait region: M.Sc. thesis, Univ. of Br i t ish Columbia, Vancouver, 100 pp. B i rd , D.N., 1981, T ime-term analysis using linear programming and its application to refraction data from the Queen Charlotte Islands: M.Sc. thesis, Univ. of Br i t ish Columbia, Vancouver, 79 pp. Chase, R.L. and Tiff in, D.L., 1972, Queen Charlotte Fault-Zone, Br i t ish Columbia: Proc. Int. Geol. Congr. 24th, section 8, 17-27. Chase, R.L., Tiff in, D.L. and Murray, J .W., 1975, The western Canadian continental margin, in Yorath, C.J., Parker, E.R. and Glass, D.J., Eds, Canada's continental margins and offshore petroleum exploration: Can. Soc. Petr. Geol., Memoir 4, 701-721. Christensen, N.I. and Salisbury, M.H., 1975, Structure and constitution of the lower oceanic crust: Rev. Geophys. Space Phys., 13, 57-86. Christensen, N.I. and Smewing, J.D., 1981, Geology and seismic structure of the northern section of the Oman ophiolite: J . Geophys. Res., 86, B4, 2545-2555. Clowes, R.M., 1984, Acquisit ion of a crustal refraction profile across the Queen Charlotte Islands and Hecate Strait: Earth Physics Branch Open File Report 84-22, Ottawa, 55 pp. Clowes, R.M., Brandon, M.T., Green, A .G . , Yorath, C.J., Sutherland Brown, A. , Kanasewich, E.R. and Spencer, C , 1986, Lithoprobe - southern Vancouver Is-land: Cenozoic subduction complex imaged by deep seismic reflections: Can. J . Earth Sci. in press. Clowes, R .M . and Knize, S., 1979, Crustal structure from a marine seismic survey off the west coast of Canada: Can. J . Earth Sci., 16, 1265-1280. Coney, P.J., 1970, The geotectonic cycle and the new global tectonics: Geol. Soc. A m . Bul l . , 81, 739-748. Coney, P.J., Jones, D.L. and Monger, J .W.H. , 1980, Cordilleran suspect terranes: Nature, 288, 329-333. Couch, R. W., 1969, Grav i ty and structures of the crust and subcrust in the southeast Pacific Ocean west of Washington and Br i t ish Columbia: Ph .D. thesis, Oregon State Univ., Corvall is. Curr ie, R.G., Stevens, L.E., Tiff in, D.L. and Riddihough, R.P., 1980, Mar ine geo-physical survey of the Queen Charlotte Islands (abstr.): Trans. A m . Geophys. Union (EOS), 61, 71. Davis, E.E. and Seemann, D.A., 1981, Compilat ion of seismic reflection profiles across the continental margin of western Canada: Geol. Surv. Canada Open File Report 751. 95 Dehlinger, P., Couch, R.W., McManus, D.A. and Gemperle, M., 1970, Northeast Pa-cific structure in Maxwel l , A .E . , Ed., The sea, vol. IV, part II: Wiley-Interscience Pub l . , New York, 133-190. Hefner, D.E. and Barret, D.L., 1979, O B S development at Bedford Institute of Oceanography: Mar. Geophys. Res., 4, 227-245. Hilde, T.W.C. , Uyeda, S. and Kroenke, L., 1977, Evolution of the western Pacific and its margin: Tectonophysics, 38, 145-165. Horn J.R., 1982, A snapshot of the Queen Charlotte Fault zone obtained from P -wave refraction data: M.Sc. thesis, Univ. of Brit ish Columbia, Vancouver, 79pp. Horn J.R., Clowes, R.M., El l is, R .M. and B i rd , D.N., 1984, The seismic structure across an active oceanic/continental transform fault zone: J . Geophys. Res., 89, B5, 3107-3120. Hyndman, R.D. and El l is, R.M., 1981, Queen Charlotte fault zone: microearth-quakes from a temporary array of land stations and ocean bottom seismographs: Can. J . Earth Sci., 18, 776-788. Hyndman, R.D., Lewis, T . J . , Wright, J.A., Burgess, M., Chapman, D.S. and Ya-mano, M., 1982, Queen Charlotte fault zone: heat flow measurements: Can. J . Earth Sci., 19, 1657-166 9. Hyndman, R.D., Riddihough, R.P. and Herzer, R., 1979, The Nootka fault zone -a new plate boundary off western Canada: Geophys. J . Roy. Astr . Soc , 58, 667-683. Hyndman, R.D. and Weichert, D.H., 1983, Seismicity and rates of relative motion on the plate boundaries of western North America: Geophys. J . Roy. Astr. Soc , 72, 59-82. Irving, E., 1979, Paleopoles and paleolatitudes of North America and speculations about displaced terrains: Can. J . Earth Sci., 16, 669-694. Keen, C.E. and Hyndman, R.D., 1979, Geophysical review of the continental mar-gins of eastern and western Canada: Can. J . Ea r th Sci., 16, 712-747. Kirst iuk, S., 1981, A preliminary investigation into the structure of the upper crust in the terrace region of the Queen Charlotte fault zone: B.Sc. thesis, Univ. of Br i t ish Columbia, Vancouver, 77 pp. Lowell, J.D., 1972, Spitsbergen Tertiary orogenic belt and the Spitsbergen fracture zone: Geol. Soc. A m . Bul l . , 83, 3091-3102. Mackie, D., 1985, Subduction beneath the Queen Charlotte Islands? Results of a seismic refraction survey: M.Sc. thesis, Univ. of Br it ish Columbia, Vancouver, 130 pp. Minster, J .B. and Jordan, T .H . , 1978, Present day plate motions: J . Geophys. Res., 83, 5331-535 4. O 'Br ien , P.N.S., 1960, Seismic energy from explosions: Geophys. J . , 3, 29-44. 96 Riddihough, R.P., 1977, A model for recent plate interactions off Canada's west coast: Can. J . Ear th Sci., 14, 384-396. Riddihough, R.P., 1982a, Contemporary movements and tectonics on Canada's west coast: a discussion: Tectonophysics, 86, 319-341. Riddihough, R.P., 1982b, One hundred years of plate tectonics in Western Canada: Geoscience Canada, 9, 28-34. Riddihough, R.P., Currie, R.G., and Hyndman, R.D., 1980, The Dellwood knolls and their role in triple junction tectonics off northern Vancouver Island: Can. J . Earth Sci., 17, 577-593. Rogers, G.C., 1983, Seismotectonics of Br i t ish Columbia: Ph.D. thesis, Univ. of Br i t ish Columbia, Vancouver, 247 pp. Sheriff, R.E. and Geldart, L.P., 1982, Explorat ion seismology, vol. I: Cambridge Univ. Press, Cambridge, 253 pp. Snavely, P.D., Wagner, H.C., Tompkins, D.H. and Tiffin, D.L., 1981, Prel iminary geologic interpretation of a seismic reflection profile across the Queen Charlotte Island fault system off Dixon Entrance, Canada - United States: U.S. Geol. Surv. Open File Report 81-299. Spence, G.D., 1984, Seismic structure across the active subduction zone of western Canada: Ph.D. thesis, Univ. of Br it ish Columbia, Vancouver, 191 pp. Spence, G.D., Clowes, R.M. and El l is, R.M. , 1985, Seismic structure across the active subduction zone of western Canada: J . Geophys. Res., 90, B8, 6754-6772. Spence, G.D., Whit ta l l , K.P. and Clowes, R.M., 1984, Practical synthetic seismo-grams for laterally varying media calculated by asymptotic ray theory: Bul l . Seism. Soc. Am. , 74, no. 4, 1209-1223. Srivastava, S.P., 1973, Interpretation of gravity and magnetic measurements across the continental margin of Br i t ish Columbia, Canada: Can. J . Earth Sci., 10, 1664-1667. Stone, D.B., 1977, Plate tectonics, paleomagn etism, and the tectonic history of the N.E. Pacific: Geophysical Surveys, 3, 3-37. Strutt, J.W., 1917, On the pressure developed in a l iquid during the collapse of a spherical cavity: Philosophical Mag., 34, 94-98. Sutherland Brown, A., 1968, Geology of the Queen Charlotte Islands: Br i t . Col. Dept. of Mines and Petr. Res., Bul let in 54, 226 pp. Tjaden, G., 1981, Seismic modelling of refraction data from the shelf off the Queen Charlotte Islands: B.Sc. thesis, Univ. of Br i t ish Columbia, Vancouver, 50pp. van der Voo, R., Jones, M., Gromme, C.S., Iberlein, G.D. and Chuckin, M., 1980, Paleozoic paleomagn etism and northward drift of the Alexander terrane, south-eastern Alaska: J . Geophys. Res., 85, 5281-5296. 97 von Huene, R., 1984, Tectonic processes along the front of modern convergent mar-gins — Research of the past decade: Ann. Rev. Earth Planet. Sci., 12, 359-381. Whitta l l , K.P. and Clowes, R.M., 1979, A simple simple, efficient method for the calculation of traveltimes and ray paths in laterally inhomogeneous media: J . Can. Soc. Explor. Geophys., 15, 21-29. Wil l is , H.F., 1941, Underwater explosions, time interval between successive explo-sions: Brit ish Report WA-47-21. Yorath, C.J . and Cameron, B.E.B., 1982, Oi l on the West Coast?: GEOS , 11, 13-15. Yorath, C.J . and Chase, R.L., 1981, Tectonic history of the Queen Charlotte Islands and adjacent areas - a model: Can. J . Earth Sci., 18, 1717-1739. Yorath, C . J . and Hyndman, R.D., 1983, Subsidence and thermal history of Queen Charlotte Basin: Can. J . Earth Sci., 20, 135-159. Young, I.F., 1981, Geological development of the western margin of the Queen Charlotte basin: M.Sc. thesis, Univ. of Br i t ish Columbia, Vancouver, 380 pp. 98 A P P E N D I X A A i rgun data. Vertical component data sections for airgun line AG2. Distance scales match the distances used in the modelling; shot-receiver distances are also indicated. Spread-ing corrections have been applied to the data; relative trace amplitudes are otherwise preserved (refer to chapter II). 13.0 Increasing shot-receiver distance (km) 0.0 14.0 LINE AG2 OBSWl EAST 00 Increasing shot-receiver distance (km) 14.0 102 A P P E N D I X B Explosion data. Vertical component data sections for explosion line E X l . Fi ltering parameters are given with each section. Shot numbers are the central 3 digits of the 5—digit code number plotted with each trace. Distance scales are model distances; shot-receiver distances are also indicated. Spherical spreading corrections and charge size compensations have been applied to the data as described in chapter II. The data plots show relative trace amplitudes after application of these corrections. c o < \ j — 01 oo r» I D I / I T i n M - o o i n r - u ) in ^ f o r v j - o u i a j ^ a j m ^ r n r v i Receiver 020 (0.5-10.0 Hz) 143.1 Increasing shot-receiver distance (km) 8.1 (n C M — o o o m D r - cc i n 7 c i <\i C M ( M ( M C M C M ( M ( M C\J o o o o o o o o o c n o o r ^ t D i n ^ n M - o o m r - i o m v i n i M -O M C M — — — — — o o o o o o o o o o o o o o o o o o o o o o o o o o o o o o t n m r o r n r o r o r o m r n r o r o r o r o r o r o r n r o r o r n m m O 0 154.5 45 60 75 90 Receiver 030 (1.5-8.5 Hz) Increasing shot-receiver distance (km) 105 120 135 19.5 C 0 15 30 45 60 75 90 105 120 135 Receiver 060 (2.0-12.0 Hz) 174.8 Increasing shot-receiver distance (km) 39.8 m c\j — o o o 00 oo co co r-c\i <M o o oo oo ID i n (M <M o o CO 00 rn C M — o C M C M C M C M C M o o o o o OO CO CO CO oo ifl in co N o o o CO 00 00 — — o o o o o o o o o o o o o o o o o o o o o o o c o c o c o o o o o a o o o o o o o o o c o o o o o c o 0 15 184.0 30 45 60 75 90 Receiver 080 (2.0-12.0 Hz) 1 ncreasin g shot receiv er distance (kin) 105 120 135 49.0 135 Receiver 090 (1.5-8.5 Hz) 189.5 Increasing shot-receiver distance (km) 54.5 


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