Science, Faculty of
Earth, Ocean and Atmospheric Sciences, Department of
DSpace
UBCV
Littel, Geena
Doctor of Philosophy - PhD
University of British Columbia
The tectonic regime of western British Columbia is largely governed by convergence of the offshore oceanic Pacific and Juan de Fuca plates with the continental North American plate. Offshore British Columbia, the Juan de Fuca plate, a remnant of the Farallon plate, subducts beneath Vancouver Island at the Cascadia subduction zone. The Pacific plate obliquely slips along the Queen Charlotte fault, a major translational fault offshore Haida Gwaii extending to southeast Alaska. These zones are joined by the Queen Charlotte triple junction, a tectonically complex region that is rapidly evolving over geologic timescales. Western mainland British Columbia experiences residual transpressional motion from oblique Pacific plate convergence, and volcanic hot spot activity. I use earthquakes to better understand the underlying tectonic processes of these distinct regions, with each presenting its own unique challenges and scientific insights. Offshore, from analysis of a new earthquake catalog I suggest that the triple junction can be modelled as a migrating pull-apart basin under regional transpression. Furthermore, breakup of the Explorer microplate likely occurs as seismicity along a predetermined strength fabric in the crust that was generated during its formation at a spreading ridge. Western British Columbia is generally seismically quiet, save for a unique concentration of earthquakes near the intersection of a major shear zone and a volcanic hot spot trend called the Anahim Volcanic Belt. I investigate these earthquakes by determining accurate hypocenters, focal mechanism solutions, and principal stresses, and by analyzing their spatio-temporal behavior. I interpret them to be a result of reactivated fault strands under northeast-directed extension in weakened lithosphere of the volcanic belt. In the Cascadia subduction zone, I apply an improved detection method to small earthquakes that accompany slow slip along the plate boundary called low-frequency earthquakes. By synthesizing accurate locations of these earthquakes with previous geophysical observations, I draw a direct link between these earthquakes and material transfer in the subduction zone from the downgoing oceanic plate to the overriding continental crust. Finally, I describe an approach to ascertaining the origin of the interesting frequency characteristics of these small earthquakes using a new seismic array on southern Vancouver Island.
A collection of earthquake studies throughout westernBritish ColumbiabyGeena LittelB. Sc. Geophysics, University of Oregon, 2018a thesis submitted in partial fulfillmentof the requirements for the degree ofDoctor of Philosophyinthe faculty of graduate and postdoctoral studies(Geophysics)The University of British Columbia(Vancouver)April 2024© Geena Littel, 2024The following individuals certify that they have read, and recommend to the Faculty ofGraduate and Postdoctoral Studies for acceptance, the thesis entitled:A collection of earthquake studies throughout western British Columbiasubmitted by Geena Littel in partial fulfillment of the requirements for the degree ofDoctor of Philosophy in Geophysics.Examining Committee:Michael G. Bostock, Professor, Department of Earth, Ocean, and Atmospheric Sciences,UBCSupervisorCharles G. Sammis, Professor Emeritus, Department of Earth, Ocean, and AtmosphericSciences, UBCSupervisory Committee MemberRon Clowes, Professor Emeritus, Department of Earth, Ocean, and Atmospheric Sciences,UBCUniversity ExaminerCarlos M. Hutt, Assistant Professor, Department of Civil Engineering, UBCUniversity ExaminerEmilie Hooft, Associate Professor & Associate Department Head, Department of EarthSciences, University of OregonExternal ExaminerAdditional Supervisory Committee Members:Catherine L. Johnson, Professor, Department of Earth, Ocean, and Atmospheric Sciences,UBCSupervisory Committee MemberiiAbstractThe tectonic regime of western British Columbia is largely governed by convergence ofthe offshore oceanic Pacific and Juan de Fuca plates with the continental North Ameri-can plate. Offshore British Columbia, the Juan de Fuca plate, a remnant of the Farallonplate, subducts beneath Vancouver Island at the Cascadia subduction zone. The Pacificplate obliquely slips along the Queen Charlotte fault, a major translational fault offshoreHaida Gwaii extending to southeast Alaska. These zones are joined by the Queen Char-lotte triple junction, a tectonically complex region that is rapidly evolving over geologictimescales. Western mainland British Columbia experiences residual transpressional mo-tion from oblique Pacific plate convergence, and volcanic hot spot activity.I use earthquakes to better understand the underlying tectonic processes of these dis-tinct regions, with each presenting its own unique challenges and scientific insights. Offshore,from analysis of a new earthquake catalog I suggest that the triple junction can be mod-elled as a migrating pull-apart basin under regional transpression. Furthermore, breakupof the Explorer microplate likely occurs as seismicity along a predetermined strength fabricin the crust that was generated during its formation at a spreading ridge. Western BritishColumbia is generally seismically quiet, save for a unique concentration of earthquakes nearthe intersection of a major shear zone and a volcanic hot spot trend called the AnahimVolcanic Belt. I investigate these earthquakes by determining accurate hypocenters, focalmechanism solutions, and principal stresses, and by analyzing their spatio-temporal behav-ior. I interpret them to be a result of reactivated fault strands under northeast-directedextension in weakened lithosphere of the volcanic belt. In the Cascadia subduction zone, Iapply an improved detection method to small earthquakes that accompany slow slip alongthe plate boundary called low-frequency earthquakes. By synthesizing accurate locationsof these earthquakes with previous geophysical observations, I draw a direct link betweenthese earthquakes and material transfer in the subduction zone from the downgoing oceanicplate to the overriding continental crust. Finally, I describe an approach to ascertaining theorigin of the interesting frequency characteristics of these small earthquakes using a newseismic array on southern Vancouver Island.iiiLay SummaryWestern British Columbia comprises myriad tectonically complex, earthquake-prone re-gions, largely due to oblique convergence of the North American tectonic plate with thePacific and Juan de Fuca plates. In this thesis I explore the earthquakes of the offshore,central coastal mainland, and southwest regions of western British Columbia. With eachregion I focus on different tectonic contexts - oceanic transform faults and spreading ridgesoffshore, volcano-tectonic interaction in the central Coast Mountains, and subduction zoneprocesses in southwest British Columbia. I present new earthquake datasets, devoting sig-nificant effort to generating highly accurate earthquake locations, and analyze their detailedspatio-temporal patterns and other seismological properties. By studying the earthquakesgenerated in these areas, my aim is to better understand the underlying tectonic processesand specific physical conditions that cause them.ivPrefaceThis thesis comprises three research chapters, two of which are published in peer-reviewedjournals and one which is a submitted manuscript. Although I am the lead author foreach of these chapters, each has been a truly collaborative work. I acknowledge the majorresearch contributions of others here.Chapter 2 is published in Tectonics as G. F. Littel, M. G. Bostock, A. J. Schaef-fer, and S. Roecker. Microplate Evolution in the Queen Charlotte Triple Junction &Explorer Region: New Insights From Microseismicity. Tectonics, 42(6), 2023. doi:https://doi.org/10.1029/2022TC007494. As such it is reproduced in this chapter unmodified savefor formatting changes. Michael Bostock provided the earthquake dataset and significantguidance throughout my analysis, interpretation, and writing. I was responsible for theprocessing of the earthquake catalog and performed the analysis, presentation of results,and drafting of the manuscript. Alexandre Plourde and Reid Merrill provided codes andassistance with software that greatly helped me. My co-authors Andrew Schaeffer (Geo-logical Survey of Canada) and Steve Roecker (Rensselaer Polytechnic Institute) and fellowearthquake seismology group members at UBC provided their insights when discussing theresults, which strengthened the final ideas put forth in the paper.Chapter 3 is published in Canadian Journal of Earth Sciences as G. F. Littel and M. G.Bostock. Seismicity at the intersection of the Coast Shear Zone and Anahim Volcanic Beltnear Bella Coola, British Columbia, Canada. Canadian Journal of Earth Sciences, 60(12):1688–1701, 2023. doi:https://doi.org/10.1139/cjes-2023-0057, reproduced here unmodifiedsave for formatting changes. Again I performed the data processing, analysis, presentationof results, and drafting of the manuscript. Michael provided the earthquake dataset andguidance throughout the analysis, interpretation, and writing.Chapter 4 is in review as: G. F. Littel, M. G. Bostock, C. G. Sammis, S. M. Peacock,and A. J. Calvert. Tectonic tremor: the chatter of mafic underplating beneath southernVancouver Island? 2024. doi:https://doi.org/10.21203/rs.3.rs-3909443/v1. The manuscriptis reproduced here unmodified save for formatting changes. As before, I performed theanalysis, presentation of results, and drafting of the manuscript. Michael provided thevtremor detection catalog and several of his codes and played a major role in processingthe data. Charles Sammis is acknowledged for providing significant supporting analysis.Michael, Charlie, Simon Peacock (UBC), and Andrew Calvert (Simon Fraser University)were closely involved in developing the ideas in the paper.Appendix D is an unpublished project undertaken during my PhD. Michael and Charliedesigned the scope of the project and provided guidance throughout my analysis. I collected,processed, and analyzed the data and presented the results. Andrew Schaeffer and ColinPaul (Geological Survey of Canada) brought me to the seismic stations where they taughtme how to service them and collect data. Although this project is unpublished, it motivatedthe design of Chapter 4.viTable of ContentsAbstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . iiiLay Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ivPreface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . vTable of Contents . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . viiList of Tables . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xList of Figures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xiAcknowledgments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xv1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11.1 Tectonic setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11.2 Earthquake location . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51.3 Thesis outline . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 102 Microplate evolution in the Queen Charlotte triple junction & Explorerregion: New insights from microseismicity . . . . . . . . . . . . . . . . . . 122.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 122.2 Tectonic setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 152.2.1 Revere-Dellwood fault and its extension, Explorer ridge, and Winonabasin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 152.2.2 Dellwood Knolls & Tuzo-Wilson volcanic field . . . . . . . . . . . . . 182.2.3 Explorer microplate, Sovanco fracture zone, and Nootka fault zone . 192.3 Data and Methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 212.3.1 Preliminary relocation: Hypoinverse . . . . . . . . . . . . . . . . . . 232.3.2 Re-relocation: HypoDD . . . . . . . . . . . . . . . . . . . . . . . . . 24vii2.4 Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 252.4.1 Revere-Dellwood fault extension, Tuzo-Wilson volcanic field, and Dell-wood Knolls . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 252.4.2 Revere-Dellwood fault, Explorer ridge, and Winona basin . . . . . . 272.4.3 Explorer microplate, Sovanco fracture zone, and Nootka fault zone . 302.4.4 Stress inversions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 342.5 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 352.5.1 Comparison to previous studies . . . . . . . . . . . . . . . . . . . . . 352.5.2 Revere-Dellwood-Queen Charlotte fault migrating pull-apart system 372.5.3 Explorer microplate deformation . . . . . . . . . . . . . . . . . . . . 472.6 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 513 Seismicity at the Intersection of the Coast Shear Zone and AnahimVolcanic Belt near Bella Coola, British Columbia, Canada . . . . . . . . 533.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 533.2 Data & Methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 563.2.1 Microseismicity data . . . . . . . . . . . . . . . . . . . . . . . . . . . 563.2.2 Hypocenter Relocation . . . . . . . . . . . . . . . . . . . . . . . . . . 573.2.3 Magnitude determination . . . . . . . . . . . . . . . . . . . . . . . . 583.2.4 Focal mechanisms & Stress inversions . . . . . . . . . . . . . . . . . 603.3 Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 613.3.1 Relocations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 613.3.2 Source Mechanisms & Stress Inversions . . . . . . . . . . . . . . . . 643.4 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 673.4.1 Consideration of surficial seismic events . . . . . . . . . . . . . . . . 673.4.2 Seismicity at the intersection of the CSZ and AVB . . . . . . . . . . 693.5 Conclusion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 734 Tectonic tremor: the chatter of mafic underplating beneath southernVancouver Island? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 744.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 744.2 Methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 754.2.1 Detection and Location . . . . . . . . . . . . . . . . . . . . . . . . . 754.2.2 False Detections and Location Uncertainty . . . . . . . . . . . . . . 794.2.3 Moment and Moment Magnitude Calculations . . . . . . . . . . . . 794.3 Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 804.3.1 Tremor location clusters . . . . . . . . . . . . . . . . . . . . . . . . . 804.3.2 Tremor layer thickness . . . . . . . . . . . . . . . . . . . . . . . . . . 83viii4.3.3 Spatiotemporal progression and moment estimates . . . . . . . . . . 834.4 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 834.4.1 Comparison of tremor locations with other geophysical observations 844.4.2 A mafic underplating model for tremor . . . . . . . . . . . . . . . . . 864.4.3 Distance-thickness calculation of the E-layer . . . . . . . . . . . . . . 884.4.4 Tremor as diagnostic of material transfer . . . . . . . . . . . . . . . 885 Conclusion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 895.1 Offshore microseismicity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 895.2 Anahim volcanic belt volcano-tectonics . . . . . . . . . . . . . . . . . . . . . 905.3 Cascadia subduction zone source processes . . . . . . . . . . . . . . . . . . . 90Bibliography . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 92A Supporting Materials for Chapter 2 . . . . . . . . . . . . . . . . . . . . . . 119A.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 119A.2 Sensitivity of earthquake locations to station geometry . . . . . . . . . . . . 119A.3 Exploring the Vp/Vs - depth tradeoff . . . . . . . . . . . . . . . . . . . . . . . 120A.4 Vp/Vs and fluid flow . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 120B Supporting Materials for Chapter 3 . . . . . . . . . . . . . . . . . . . . . . 129B.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 129C Supporting Materials for Chapter 4 . . . . . . . . . . . . . . . . . . . . . . 134C.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 134C.2 Positive tremor detection criteria . . . . . . . . . . . . . . . . . . . . . . . . 134C.3 Width of shear layer . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 135D Frequency-dependent array analysis of tremor in northern Cascadia . . 144D.1 Preface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 144D.2 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 144D.2.1 PISA seismic array . . . . . . . . . . . . . . . . . . . . . . . . . . . . 146D.2.2 Tremor data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 148D.3 Method: Cross-station array analysis . . . . . . . . . . . . . . . . . . . . . . 148D.3.1 Relative arrival times . . . . . . . . . . . . . . . . . . . . . . . . . . 150D.3.2 Slowness vector solution . . . . . . . . . . . . . . . . . . . . . . . . . 152D.3.3 Station corrections . . . . . . . . . . . . . . . . . . . . . . . . . . . . 153D.4 Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 153D.5 Discussion and Conclusion . . . . . . . . . . . . . . . . . . . . . . . . . . . . 161ixList of TablesTable 2.1 HypoDD inversion statistics for each subregion. . . . . . . . . . . . . . . 25Table 2.2 Principal stress orientations. . . . . . . . . . . . . . . . . . . . . . . . . . 35Table 4.1 Parameters for each cluster (1, 2, 3, 4) or patch (north, south) for eachdataset (4S or 4S+1P). . . . . . . . . . . . . . . . . . . . . . . . . . . . . 82Table B.1 Principal stress orientations. . . . . . . . . . . . . . . . . . . . . . . . . . 129xList of FiguresFigure 1.1 Western British Columbia tectonic setting. . . . . . . . . . . . . . . . . 2Figure 1.2 Schematic diagram of the double difference earthquake relocation algo-rithm. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8Figure 2.1 Tectonic setting of the Queen Charlotte Triple Junction region. . . . . . 14Figure 2.2 Schematic cartoon of possible triple junction evolution. . . . . . . . . . 16Figure 2.3 Local magnitudes determined for the GSC catalog. . . . . . . . . . . . . 22Figure 2.4 HypoDD relocations. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26Figure 2.5 Temporal patterns of seismicity in the RDFE area. . . . . . . . . . . . . 28Figure 2.6 Relocations and focal mechanisms of the TWVF region. . . . . . . . . . 29Figure 2.7 Seismicity in the Dellwood knolls region. . . . . . . . . . . . . . . . . . . 30Figure 2.8 Seismicity in the RDF region. . . . . . . . . . . . . . . . . . . . . . . . . 31Figure 2.9 Seismicityin the Winona basin. . . . . . . . . . . . . . . . . . . . . . . . 32Figure 2.10 Seismicity in the Explorer microplate. . . . . . . . . . . . . . . . . . . . 33Figure 2.11 Stress inversion results. . . . . . . . . . . . . . . . . . . . . . . . . . . . 36Figure 2.12 Relocation comparisons. . . . . . . . . . . . . . . . . . . . . . . . . . . . 38Figure 2.13 Relocation-hydrophone location comparison. . . . . . . . . . . . . . . . 39Figure 2.14 Schematic diagram illustrating the migrating pull-apart system model. . 45Figure 2.15 Relocations with magnetic intensity from the magnetic anomaly model. 47Figure 3.1 Tectonic setting of the central British Columbia Coast Mountains region. 55Figure 3.2 Results of multiple-linear regression of reported NRCan ML values withWood-Anderson S-wave amplitude and hypocentral distance. . . . . . . 59Figure 3.3 Final earthquake relocations. . . . . . . . . . . . . . . . . . . . . . . . . 62Figure 3.4 2015, 2017 earthquake swarms. . . . . . . . . . . . . . . . . . . . . . . . 63Figure 3.5 Estimated focal mechanisms for events in the Bella Coola strands clusters. 65Figure 3.6 Estimated focal mechanisms for events in clusters northwest of Bella Coola 66Figure 3.7 Stress inversion results. . . . . . . . . . . . . . . . . . . . . . . . . . . . 67Figure 3.8 Three-component waveform alignment results. . . . . . . . . . . . . . . 68xiFigure 3.9 Comparison of temperature and precipitation from the Government ofCanada weather station at Bella Coola Airport (Climate ID: 1060844)during the swarm (top), compared with the daily event detection count(bottom). . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 70Figure 3.10 Map of relocations with hot springs. . . . . . . . . . . . . . . . . . . . . 71Figure 4.1 S and P waveform windows used to detect tremor. . . . . . . . . . . . . 76Figure 4.2 Map of the Cascadia region. . . . . . . . . . . . . . . . . . . . . . . . . . 77Figure 4.3 Map view of final tremor hypocenters. . . . . . . . . . . . . . . . . . . . 81Figure 4.4 4S tremor dataset. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 81Figure 4.5 Distributions of distances and associated statistics from the PCA planesfor the north and south patches. . . . . . . . . . . . . . . . . . . . . . . 84Figure 4.6 Depth section of tremor overlain on other geophysical observations. . . . 85Figure A.1 REST-GSC matched events. . . . . . . . . . . . . . . . . . . . . . . . . . 121Figure A.2 REST and GSC non-matched events from the 4-paired pick culled dataset.122Figure A.3 Velocity models used during preliminary location. . . . . . . . . . . . . 123Figure A.4 Comparison of epicenter locations obtained using three parameters com-binations. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 124Figure A.5 Wadati plots (S-P time versus P travel time). . . . . . . . . . . . . . . . 124Figure A.6 Comparison of Hypoinverse location solutions. . . . . . . . . . . . . . . 125Figure A.7 Hypoinverse relocations of events with 8 paired P and S phase picks alongthe RDF. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 126Figure A.8 Progression of relocations from initial locations to the final hypoDD re-locations. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 127Figure A.9 Depths of final hypoDD relocations. . . . . . . . . . . . . . . . . . . . . 127Figure A.10 Cumulative seismicity detections in the West Valley and Middle Valleyregions of the Juan de Fuca ridge. . . . . . . . . . . . . . . . . . . . . . 128Figure B.1 P-velocity model used for relocation in Hypoinverse and hypoDD andfocal mechanism estimation in HASH. . . . . . . . . . . . . . . . . . . . 130Figure B.2 Wadati plot for best-fitVpVs. . . . . . . . . . . . . . . . . . . . . . . . . . 131Figure B.3 Relocations during the 2005-2006 BATHOLITHS deployment. . . . . . 131Figure B.4 3D inset plots showing spatio-temporal progression of the 2015 and 2017swarms. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 132Figure B.5 Vertical-component waveforms of two events recorded during the tempo-rary BATHOLITHS installment at two stations. . . . . . . . . . . . . . 132xiiFigure B.6 Cumulative event detections over time for the entire relocated datasetwith the BATHOLITHS deployment. . . . . . . . . . . . . . . . . . . . . 133Figure C.1 Best-fit PCA planes to the north and south clusters. . . . . . . . . . . . 136Figure C.2 Comparison of locations from common time detection windows betweenthe narrow and broad band data for each template. . . . . . . . . . . . . 137Figure C.3 Comparison of 4S detections for each ETS episode in 2003, 2004, and 2005.137Figure C.4 Histogram distributions of moment magnitudes and scalar moments foreach cluster. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 138Figure C.5 Gouge thickness, H, as a function of total slip, D. . . . . . . . . . . . . 139Figure C.6 Comparison relocations between LFE template waveforms using the methodpresented in this study and LFE template locations. . . . . . . . . . . . 140Figure C.7 Comparison tremor locations for all six possible datasets. . . . . . . . . 140Figure C.8 Comparison of high- and low Gini index detections. . . . . . . . . . . . 141Figure C.9 Comparison of high- and low Gini index detections. . . . . . . . . . . . 142Figure C.10 Average kurtosis versus width/standard deviation. . . . . . . . . . . . . 143Figure D.1 Discrepant arrivals in varying frequency bands. . . . . . . . . . . . . . . 145Figure D.2 Coherent arrivals in the 4-8 Hz band. . . . . . . . . . . . . . . . . . . . 146Figure D.3 East-component velocity waveforms of on 21 August 2019 show coherentarrivals in each frequency band. . . . . . . . . . . . . . . . . . . . . . . . 147Figure D.4 Map of study and PISA array. . . . . . . . . . . . . . . . . . . . . . . . 149Figure D.5 Arrival time differences at the borehole stations for 6 test earthquakes.Events on the x-axis are given as dates in month-day (MMDD) format. 154Figure D.6 Comparison of slowness and backazimuth values determined using vary-ing window lengths for varying frequency bands. Window lengths shownare [4, 2, 1], [16, 8, 4], and [128, 64, 32] seconds for 1-2 Hz, 2-4 Hz, and4-8 Hz data, respectively. . . . . . . . . . . . . . . . . . . . . . . . . . . 154Figure D.7 Slowness values for 10 hours on August 20 for data of varying frequencybands, plotted with slowness values for known LFE locations recorded atthe same time. Top panel displays raw values, whereas the bottom paneldisplays a running mean. . . . . . . . . . . . . . . . . . . . . . . . . . . 155Figure D.8 Backazimuth values for 10 hours on August 20 for data of varying fre-quency bands, plotted with backazimuth values for known LFE locationsrecorded at the same time. . . . . . . . . . . . . . . . . . . . . . . . . . 156Figure D.9 Mean cross-correlation (CC), mean residuals, slowness, and backazimuthresults for 1-2 Hz, 2-4Hz, and 4-8 Hz frequency data on 08-22-2019. . . 157xiiiFigure D.10 Tremor location density maps for 1-2 Hz, 2-4Hz, and 4-8 Hz frequencydata on 08-22-2019. Red contour on the 1-2 Hz and 2-4 Hz maps indicatesthe 4-8 Hz location contour. . . . . . . . . . . . . . . . . . . . . . . . . . 158Figure D.11 X- and Y- coordinate positions (in km) for 1-2 Hz, 2-4Hz, and 4-8 Hzfrequency data for August 20-22 2019. . . . . . . . . . . . . . . . . . . . 159Figure D.12 Comparison of cross-correlation, slowness, and backazimuth results for4-station subgroups of the PISA array. . . . . . . . . . . . . . . . . . . . 159Figure D.13 Arrival time contours for 4-8 Hz tremor on 08-22-2019. . . . . . . . . . . 160Figure D.14 Arrival time contours for a test regional earthquake on September 3 2019.(a) displays the location of the earthquake relative to the PISA array, aswell as the predicted slowness and backazimuth values. (b), (c), and (d)display the arrival time contours for the 1-2 Hz, 2-4 Hz, and 4-8 Hz data,respectively. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 160xivAcknowledgmentsSeismology is wonderful because it is an inherently open-source field, which fosters collabo-ration and makes science easily accessible to researchers all over the world. From all of thepeople working in the field to those at computers analyzing waveforms and archiving data- thank you. In particular I would like to mention Natural Resources Canada/GeologicalSurvey of Canada, Ocean Networks Canada, and the Incorporated Research Institutions forSeismology.I am fortunate to have benefited from multiple funding sources during my tenure at UBC.I was primarily funded by my supervisor, Michael Bostock, via Discovery grants to himfrom the Natural Science and Engineering Research Council of Canada (NSERC). Thankyou Michael for your commitment to a fair stipend for your students. Additional fundingwas provided by the Department of Earth, Ocean, and Atmospheric Sciences, the Universityof British Columbia, and the PRODIGY (NSERC CREATE - Pacific Rim Oceanic DataMobilization and Technology, https://prodigy.eoas.ubc.ca) program.Aside from supplementing my funding, the PRODIGY program was an extremely funand enriching addition to my PhD studies. Thank you to Philippe Tortell for your initia-tive and enthusiasm for this program, and for eagerly including geophysicists. I’ll alwayslook back fondly on the amazing new things I learned and places I visited through myparticipation in PRODIGY.Thank you to all of my friends in the department and outside UBC for enriching mylife. I especially want to mention my fellow housemates at Casablanca - although there aretoo many to mention here, the times we’ve shared and the bond we’ve developed is so dearto me. It’s been so special getting to know you all so closely and having a place that trulyfeels like home because of the people in it and a shared value of community.I would like to thank Amanda Thomas for jump-starting my path down this road whenI was an undergraduate at the University of Oregon. I will never forget being infected withher enthusiasm for science and research, and will always be grateful for the confidence sheinstilled in me at such a formative time in my life. And to Michael, it has truly been apleasure working with you for the past several years. Thank you for your excitement in myxvdevelopment as a scientist; your generosity and compassion; and your unwavering supportboth in the difficult times and the fun times. Furthermore, travelling to disseminate myresearch and meet with fellow members in the community was so special to me. And finally,thank you for including me on board the M/G Langseth. Although we both reasoned thata life at sea was not fit for either of us, this experience really was incredible for me and I’mnot sure I’ll ever have one like it again.And to the earthquake seismology group at UBC - Doriane Drolet, Raul Mendoza, ReidMerrill, Alex Hughes, Wasja Bloch, Alex Plourde, Charlie Sammis, Simon Peacock - everyweek I looked forward to our meeting of ramblings on various topics. Thank you for yourcompany and unique insights. To Doriane especially, it has been so fun working on ourPhDs concurrently and I’ve sincerely enjoyed your friendship and support over the years.Thank you to Andrew Schaeffer and Colin Paul (GSC) for hosting me at the PacificGeoscience Centre in Sidney and for taking me out to the PISA seismic array. Althoughthese were relatively brief experiences, they were very significant to me and I had a lot offun.Finally, I want to acknowledge my parents, who have always been exceptionally sup-portive and trusting in my life decisions. They nurtured my various ambitions and ideasand gave me the freedom to choose my path in life so long as I could be self-supported. Sothank you, Mom and Dad, for your continued support and interest in me.xviChapter 1Introduction1.1 Tectonic settingSince the 1960’s the west coast of British Columbia (BC) has been the locus of groundbreak-ing work leading to the development of the paradigm-shifting theory of plate tectonics. Forexample, early magnetic studies of the Juan de Fuca ridge (Mason and Raff, 1961, Vine,1966, Vine and Wilson, 1965) offshore BC supported the seafloor-spreading hypothesis (Vineand Matthews, 1963). Wilson (1965) cited the Queen Charlotte Fault offshore Haida Gwaiiin his seminal work on transform faults. Studies of plate motions in this region resultedin one of the early formulations for rigid plate motion kinematics (McKenzie and Parker,1967), as well as the suggestion of Juan de Fuca plate subduction beneath North America(Atwater, 1970). In more recent years the region has been the focus of scientific develop-ments in oceanic plate breakup (Dziak, 2006, Rohr and Furlong, 1995, Savard et al., 2020)and episodic tremor and slip (Dragert et al., 2001, Rogers and Dragert, 2003). The mainpremise of this thesis is to advance the current understanding of regional-scale tectonics inwestern British Columbia through analysis of earthquakes. I begin with a brief introductionto the tectonics of western British Columbia and summarize the theoretical foundations ofearthquake location. Further details regarding the specific tectonics may be found withinthe introductory sections of Chapters 2, 3, and 4.The Juan de Fuca plate subducts beneath North America at the Cascadia SubductionZone, which extends from offshore northern California to northern Vancouver Island (Figure1.1). To the north, the Pacific and North American plates are separated by the QueenCharlotte Fault (QCF), a 1150-km long transform boundary from Yakutat, Alaska, tothe southern end of offshore Haida Gwaii. Notably, the QCF generated Canada’s largestearthquake (Mw 8.1) and has the highest yearly average slip rate of any continental-oceanicstrike-slip fault (like the San Andreas) at ∼50-57 mm/yr (Brothers et al., 2020). Between1−132.0°−132.0°−129.0°−129.0°−126.0°−126.0°−123.0°−123.0°−120.0°−120.0°48.0° 48.0°51.0° 51.0°54.0° 54.0°North AmericanplateCoast Shear ZoneQueen Charlottetriple junctionCascadiasubduction zoneChapter 3Chapter 2Chapter 4Bella CoolaVancouverPacific plateFigure 1.1: Western British Columbia tectonic setting. Major faults are delineatedwith black lines, with the primary focus regions of this thesis labelled (Chap-ter 2: Queen Charlotte triple junction, Chapter 3: Coast Shear Zone, Chap-ter 4: Cascadia Subduction Zone) and encircled in pink. Natural ResourcesCanada stations are indicated with red triangles and additional temporary sta-tions used in this thesis are in blue hatched triangles. Earthquakes from theNatural Resources Canada earthquake database with magnitude >2 from 2000-2023 are plotted as small circles (https://earthquakescanada.nrcan.gc.ca/stndon/NEDB-BNDS/bulletin-en.php). Pacific-North America relative plate motion isshown as a vector arrow at 132◦W, 51.5◦N. Elaboration upon specific tectonicfeatures can be found in the introductory figures of the subsequent chapters.2the QCF and the Cascadia Subduction Zone lies a broad, poorly understood triple junctionof the Juan de Fuca, North America, and Pacific plates, called the Queen Charlotte triplejunction (Figure 1.1). This region has been rapidly evolving for the past few million yearsin adjustment to a clockwise rotation in the prevailing Pacific plate motion (Botros andJohnson, 1988, Riddihough, 1984). The Revere-Dellwood Fault (RDF) is a young (∼2Myr) transform fault that separates the Pacific plate from the actively-deforming Explorermicroplate. There are numerous oceanic spreading centers between the Pacific, Explorer,and Juan de Fuca plates, and dispersed volcanism in the form of the Tuzo-Wilson VolcanicField (TWVF) and Dellwood Knolls where the RDF meets the QCF. This region is thefocus of Chapter 2.We generally expect much of the seismicity in the offshore region to occur on oceanictransform faults, mid-ocean ridges, and other shallow faults within the oceanic crust.Oceanic transform faults usually accommodate motion between two mid-ocean ridges, andcan extend for hundreds of km along the ocean floor (Hekinian, 1982, Wolfson-Schwehrand Boettcher, 2019). They can rupture quasi-periodically in up to M∼7 events, and alsodeform during aseismic creep events (e.g., Kuna et al., 2019). Seismicity at oceanic trans-form faults can facilitate hydrothermal circulation of fluids containing compounds such asCH4 and H2 from deeper within the mantle through the overlying sediments (Hensen et al.,2019, Kelley et al., 2001), promoting chemosynthetic life at the sea floor (Kelley et al.,2001). Fluid flow at oceanic transform faults also commonly results in serpentinization, ahydration metamorphic reaction that is most prevalent in ultramafic rocks from the uppermantle (Francis, 1981). Indeed, there is evidence for serpentinization in the Nootka faultzone offshore Vancouver Island (Rohr et al., 2018). Transform faults with a componentof extension may exhibit extensional tectonic structures such as pull-apart basins and/orvolcanism, such as the region offshore BC in the Queen Charlotte triple junction (Rohr andFurlong, 1995, Wolfson-Schwehr and Boettcher, 2019).Triple junctions are regions where three tectonic plates meet. The stability of a triplejunction describes its tendency to remain in its present configuration or devolve into a moresustainable arrangement, and is dependent upon the orientation of the plate boundaries(Cronin, 1992). The Queen Charlotte triple junction comprises a spreading ridge (Juan deFuca ridge), trench (Cascadia subduction zone), and transform fault (Queen Charlotte fault)and is believed to have existed offshore Vancouver Island for at least 40 Myr (Riddihough,1982). However these three boundaries do not meet at a single point; rather, their junctionis the broad, rapidly-deforming region between Vancouver Island and Haida Gwaii (Figure1.1. Because of the plate orientations and a change in relative plate motions over the past 20Myr (DeMets and Merkouriev, 2016, Rohr and Furlong, 1995), the region is reconfiguringitself into a more stable orientation. Earthquakes provide one of the best measures of3present-day deformation (Chapter 2).Inland of the triple junction, the deformation zone of the Pacific-North America plateboundary may include the western mainland BC (Mazzotti et al., 2003, 2008). Globalnavigation satellite system (GNSS) velocities relative to stable North America inland ofthe Queen Charlotte Fault decrease but still exhibit northward motion of up to ∼5 mm/yr(Mazzotti et al., 2008). In fact, a major NE-dipping zone of intense ductile deformationcalled the Coast shear zone (CSZ) extends for at least 1200 km along the NW-SE-orientedCoast Mountains of western BC. The CSZ exhibits little to no documented seismicity,although much of it extends through remote, rugged, and poorly instrumented regions.Although its last documented movement was during the Eocene around 56 Ma (Hollister andAndronicos, 2006, Rusmore et al., 2001), as a major margin-parallel structure, the CSZ maybe important for accommodating residual northward Pacific-North America deformation(Hyndman, 2015).Additionally, an ENE-WSW-trending zone of alkaline-peralkaline volcanic centres ex-tends from the central BC coast (at approximately 52◦ latitude) to the interior near Quesnel,called the Anahim Volcanic Belt (AVB). The AVB is the result of Neogene-Quaternary vol-canism since about 15 Ma (Bevier, 1989, Bevier et al., 1979, Souther, 1977). The volcaniccentres exhibit an age progression that becomes younger moving from west to east, at arate consistent with movement of the North American plate over a fixed source (possiblypreviously accreted seamount material in the mantle) (Bevier, 1989, Bevier et al., 1979).Based on this observation and geochemical analyses, a hot-spot origin has been the preferredinterpretation for the AVB (Bevier, 1989, Bevier et al., 1979, Kuehn et al., 2015, Rogers,1981). The focus of Chapter 3 is a unique concentration of earthquakes in the region wherethe CSZ intersects the Anahim Volcanic Belt.Seismic activity in volcanic zones such as the Anahim Volcanic Belt is generated from themovement of magma and hydrothermal fluids within the crust, deformation of the crust itselfin response to stresses imparted by these movements, or tectonic processes in the volcanicplumbing system (e.g., Keir et al., 2021). These earthquakes often occur in swarms, definedas a sequence of (usually) small events lacking a mainshock (Hill, 1997). The spatial andtemporal evolution of events during these swarms and their source characteristics can bediagnostic of whether the events are caused by magmatic intrusion or due to other processessuch as pore-pressure diffusion and aseismic slip facilitated by crustal fluid redistribution(e.g., Cassidy et al., 2011, Russo et al., 2020, Shelley et al., 2013, Waite and Smith, 2002).In the Cascadia subduction zone between the regions of frictionally locked brittle de-formation and slow, ductile slip, accrued stress is periodically released as slow shear slip(e.g., Dragert et al., 2001). This manifests seismically as tectonic tremor and geodetically asseaward motion of the overriding plate (Dragert et al., 2001, Rogers and Dragert, 2003). In4the northern Cascadia subduction zone, this periodic release occurs about every 14 months(Brudzinski and Allen, 2007, Rogers and Dragert, 2003), thus, the phenomenon has beendubbed “episodic tremor and slip” (ETS, Rogers and Dragert, 2003). In Cascadia, ETStakes place between 25 and 40 km depth, just downdip of the region where large megath-rust earthquakes are generated (i.e., the “seismogenic zone) in subduction zones. Variousauthors have suggested that slow slip could transfer stress to the adjacent seismogenic zoneor shift its lower limit landward (Dragert et al., 2004, Wech and Creager, 2011). Thus, ETShas generated considerable interest due to its potential for both elucidating fundamentalprocesses of subduction zones and improving our assessment of the associated seismic haz-ard. This is especially important in earthquake-prone regions, such as Cascadia, which lastexperienced a large megathrust earthquake in 1700 (Satake, 1996).The seismic sources that compose much, if not all, of the tremor signal are called “low-frequency earthquakes” (LFEs) due to their depletion in frequencies > 10 Hz relative toregular earthquakes of comparable magnitude (Katsumata and Kamaya, 2003). The under-lying reason for this depletion is still debated, but it is likely a result of an atypical sourcerupture, as opposed to near-source attenuation (Bostock et al., 2017, Wang et al., 2023).“Tremor” is a broader term referring to the seismic signal of slow slip and “LFE” refers todiscrete source events (Mw ∼ 1-2) that presumably compose the tremor signal (Katsumataand Kamaya, 2003). As such, nearly all studies of tremor restrict analysis to within the 1-10Hz band (e.g., Ghosh et al., 2012, Rubin and Armbruster, 2013, Wech and Creager, 2008).Because LFEs generally coincide spatially and temporally with geodetically-estimated slowslip (Rogers and Dragert, 2003, Wech and Bartlow, 2014), their source depths locate nearthe plate boundary (Brown et al., 2009, Ghosh et al., 2010, LaRocca et al., 2009), andinferred source mechanism is consistent with shear slip, the prevailing notion is that LFEsreflect shear fault slip in the plate boundary zone in many subduction zones such as Nankai,Cascadia, and Chile. However, tremor depths are more difficult to constrain. Thus, somestudies have argued for a much wider depth range of tremors within the forearc crust (Bom-bardier et al., 2023, Kao et al., 2009). In Chapter 4 we address this dilemma by adaptingexisting tremor detection methods and applying them to southern Vancouver Island datato generate a tremor catalog with accurate depths. By combining our results with existinggeophysical observations, we directly connect tremor to a fundamental subduction process,that is, ongoing underplating, in Cascadia.1.2 Earthquake locationIn this section I describe the foundations of earthquake location. Earthquake locations area primary focus of my research and I rely heavily on them to make tectonic interpretations.Earthquakes are usually treated as a point source, with the point being either the origin5of rupture if analyzing first arrivals of body waves, or the centroid of rupture if usingwaveform cross-correlations. Earthquake locations are quantified by four parameters: thelocation coordinates in space, x, y, z, and in time, t (that is, the origin time). For any givenearthquake with n arrival time recordings at i stations, the arrival time ti will be somenon-linear function of the earthquake model parameters x, y, z, and t, and the Earth’sseismic velocity distribution. By estimating the values of the model parameters we mayobtain an arrival time prediction, tpi , for the real arrival time, ti. The “best” model will bethe one that minimizes the residual between the real and predicted arrival times,ri = ti − tpi (1.1)The non-linear function tpi is dependent on Earth structure as well as station geometry.Increasing the number of arrival time recordings lends further constraints for estimating thelocation parameters. As there are 4 earthquake source parameters, at least 4 arrival timerecordings are needed to locate an earthquake.Indeed, minimizing Equation 1.1 is the paramount problem in locating earthquakes.Earthquake location routines commonly linearize the inverse problem following the Geiger(1912) method, using a Taylor series expansion. To minimize arrival time residuals, themethod uses an initial estimate of the location and velocity model and iterates over valuesof the model parameters. This proceeds until the inversion achieves a convergence threshold,which may be an arrival time residual tolerance value. Mathematically, this looks like:t = Ax+ g (1.2)where t is a vector of arrival times, x is a vector of hypocentral coordinates, A is a matrixof partial derivatives of t with respect to the hypocentral parameters x (Equation 1.3), andg is a constant vector containing initial estimates of the hypocentral parameters.Aij =∂ti∂xj(1.3)Suppose one is iteratively approaching a solution. Let t0 and x0 be the arrival time andhypocentral vectors calculated at a previous step. Then we may define the residual at eachstep as:r = t− t0 = A(x− x0) = A∆x (1.4)The hypocentral adjustment vector (x− x0) is represented as ∆x. To minimize r in aleast-squares sense, we solve the normal equations:6∆x = (ATA)−1ATr (1.5)or evaluate the solution using the singular value decomposition (Klein, 2002).In all three projects, I obtain initial location estimates with the Hypoinverse earthquakelocation software (Klein, 2002), which uses the iterative, linearized approach describedabove. Hypoinverse uses initial estimates of earthquake origin times, arrival times, andvelocity model to update earthquake locations. The method does not use initial hypocentrallocation estimates (x, y, z). Rather, it iteratively minimizes the arrival time residuals using aprovided input velocity model and origin time (t). Thus, a priori biases in event location maybe neutralized, such as initial locations generated using different velocity models (Chapters2, 3).For clustered earthquakes within a relatively confined region compared to average station-event distances, relative locations between earthquakes may be determined with muchgreater accuracy than absolute locations. The innovative “double-difference” method ex-ploits the fact that earthquakes closely spaced to one another generate waves that travelnearly the same path to any given station (Waldhauser and Ellsworth, 2000). Residuals be-tween the observed and theoretical travel-time differences (“double-differences,” Equation1.6) are minimized for pairs of earthquakes at each station n. A least-squares solution isfound by iteratively adjusting the vector difference between hypocentral pairs:drijn = (tin − tjn)obs − (tin − tjn)pred (1.6)If the ray path to a station is anomalously slow for one event, then it will be slow forthe other events as well, assuming the events are contained within a local source region thatis small compared to the heterogeneity (Figure 1.2). Hence, residuals caused by velocitystructure heterogeneity not provided in the initial model will be correlated among all of theevents. Therefore, an added advantage of the double-difference method is that it alleviatesthe need for station corrections.Waveform cross-correlation data provide precise double-difference measurements, whichare combined with absolute travel-time data into one system of equations. Thus, inter-event distances within earthquake clusters are determined to the accuracy of the cross-correlation data, while the relative locations between clusters and uncorrelated events aresimultaneously determined to the accuracy of the absolute travel-time data. Thereforeaccurate waveform cross-correlation times become very important because their additionresults in highly accurate relative locations between earthquakes, which are then adjustedto absolute locations in the algorithm based on the preliminary location estimates.Commonly, multiple channels at multiple stations are used to obtain optimal cross-7a) b) c)Figure 1.2: Schematic diagram of the double difference earthquake relocation algo-rithm. a) shows events i and j travelling a similar path to the receiver k(reproduced from Castellanos and van der Baan (2013). b) and c) show beforeand after results of earthquake relocations in the Long Valley Caldera (Califor-nia), reproduced from Waldhauser (2001).correlation times, which is called multi-channel cross-correlation (VanDecar and Crosson,1990). This method obtains cross-correlation delays for all waveform pairs, and solves anoverdetermined system of linear equations (Equation 1.7) for relative arrival times usingleast-squares, subject to the constraint that the mean of the relative arrival times be equalto zero (Equations 1.8, 1.9).Wt =∆t (1.7)W is a coefficient matrix, ti are the relative arrival times at each station, and ∆tij are thecross-correlation derived time delays between seismic arrivals.n∑i=1ti = 0 (1.8)For a system with n = 4 stations, the linear system of equations may look like:1 −1 0 01 0 −1 01 0 0 −10 1 −1 00 1 0 −10 0 1 −11 1 1 1t1t2t3t4 =∆t12∆t13∆t14∆t23∆t24∆t340(1.9)Proper waveform alignment is imperative to obtaining useful relative delay times, as8small errors in relative delay times result in errors in earthquake locations. However, theMCCC approach can be problematic when the signal-to-noise ratio of seismograms is low.In this circumstance the MCCC approach commonly misidentifies the maxima in the cross-correlation functions, resulting in correlating the wrong peaks, known as cycle skipping.Cycle skips ultimately become outliers in the linear system that can severely disrupt least-squares solutions for relative delay times.In chapters 3 and 4 I use an iterative, adaptive alignment procedure based on multichan-nel cross-correlation as described above (Bostock et al., 2022). This method reformulatesand extends the MCCC system (i.e., Equation 1.9) in a principal component framework,minimizing principal-component based objective functions instead of maximizing pair-wisecorrelation coefficients.Furthermore, this method recovers accurate S-wave delay times. In regional studieswhere one may be comparing multiple earthquake sources with different focal mechanismsat common stations to determine relative delay times, the P- and S- waveforms will not bethe same. For events that have different source mechanisms, the P-waveforms will be scalarmultiples of each other, because P-wave polarizations possess one degree of freedom. Thatis, they will either be the same polarity, or flipped. However, the S- waveforms may be evenmore discrepant because they have an additional degree of freedom in their polarization. Byemploying principal components, the adaptive alignment procedure accommodates varyingS- waveforms and delivers highly accurate S-wave delay times (Bostock et al., 2022).After initial relocation with Hypoinverse and calculation of cross-correlation times, I re-locate hypocenters once more using the double-difference method within either the hypoDDor tomoDD FORTRAN routines (Waldhauser and Ellsworth, 2000, Zhang and Thurber,2003). In chapters 2 and 3 I use hypoDD, which only allows a one-dimensional velocitymodel. The reason for using hypoDD in these studies is that no 3D velocity models cur-rently exist for either of the focus regions. Both regions are in extremely remote locationswith sparse station distributions that preclude the possibility of determining our own 3Dvelocity model (which was attempted in Chapter 2). However, in Chapter 4 I use a 3Dvelocity model (Savard et al., 2018) with tomoDD.Each of my three projects highlighted one or more of the primary challenges in earth-quake location, that is, 1) sparse station distribution (in space and time), 2) complicatedand/or unknown three-dimensional velocity structure (and errors associated with existingmodels) and 3) difficulty in correctly identifying P- and/or S- wave arrival time onsets dueto poor data quality, large event-station distances, or low signal-to-noise ratio. For exam-ple, the data of Chapter 2 were recorded with epicentral distances >100 km, compoundingunmodelled velocity errors.Generally, errors introduced during arrival time picking and deviation of the velocity9model from true 3D structure contribute most significantly to errors in hypocentral param-eter estimation. The former may be addressed by 1) ensuring appropriate data processingand filtering are applied, and 2) manual phase arrival review and update if the phase ar-rival picks were determined automatically. The latter may be addressed by incorporatingvelocity structure constraints from previous seismic tomography and refraction work to in-form initial model selection. Lateral velocity structure variations may be modeled by usingseparate initial models for different subregions or for certain stations. Furthermore, testingvarious input models allows one to probe the model space and make an adequate selectionbased on total model residual minimization and hypocentral location errors.1.3 Thesis outlineIn this thesis I present three independent research endeavours, followed by a brief conclusionand thoughts on natural continuations of my work. The appendix includes supplementarymaterials for each chapter (Appendices A-C), followed by a brief, unpublished project un-dertaken during my PhD (Appendix D).In Chapter 2 (Littel et al., 2023), I present a new microseismicity dataset for the offshoreregion of British Columbia, Canada. This dataset was generated by combining the existingGSC catalog with a newly-determined catalog using an automated detection and locationalgorithm. I relocated these data and performed stress inversions to characterize the present-day tectonics of the region. I modified and extended previous geophysical models for platereorganization and deformation in this rapidly-evolving region.In Chapter 3 (Littel and Bostock, 2023), I applied similar methods as in Chapter 2to a much smaller region in the Coast Mountains of western British Columbia near BellaCoola. I investigated the relationship of seismicity to the Coast Shear Zone and the Anahimvolcanic belt, two major geological features in the area. I applied relocation techniques andsolved for focal mechanism estimates to better characterize the spatial and temporal extentof seismicity and its origins.In Chapter 4 (Littel et al., 2024) I shifted my focus to the Cascadia subduction zonebeneath Vancouver Island, where I also applied earthquake relocation techniques to betterunderstand the relationship of tectonic tremor with subduction zone source processes. Incontrast to Chapters 2 and 3, the earthquake detection technique used here includes onlyfour stations to ensure maximally consistent measurements for each tremor event, and athree-dimensional velocity model based on Savard et al. (2018). In result I obtain highlyaccurate tremor locations and demonstrate a direct link between tremor and tectonic un-derplating in the Cascadia subduction zone.In Appendix D I applied a different kind of earthquake location technique based on asmall-scale array of seismic stations to seismic tremor in the Cascadia subduction zone.10My research goal was to answer whether differences in spatiotemporal behavior with datafrequency band could be identified. The working hypothesis was that these differences mayreflect heterogeneity in the source region. Although the method used to answer this questionwas unsuccessful, the project was a valuable experience for me. Realistically, this happensin research and I believe the work should still be documented so that others can learn fromit. Furthermore, unlike the primary research chapters, I was directly involved in the datacollection and instrument service, taking two field trips to Sidney, BC with colleagues atthe Geological Survey of Canada.11Chapter 2Microplate evolution in the QueenCharlotte triple junction &Explorer region: New insights frommicroseismicity2.1 IntroductionThe Queen Charlotte triple junction offshore British Columbia has existed as an unstableconfiguration for about 40 Myr (Riddihough, 1982). Historically, the triple junction itselfis considered the broad, seismically-active region between the Explorer ridge and southernend of the Queen Charlotte fault (that is, it does not occupy a single point) (e.g., Carbotteet al., 1989, Chase et al., 1975, Davis and Currie, 1993, Riddihough, 1982) (Figure 2.1). Aroughly 20◦ clockwise rotation of the relative Pacific-North America tectonic plate motionsover the past ∼20 Myr (c.f., Atwater, 1970, DeMets and Merkouriev, 2016) has causedongoing adjustment and a variety of resultant geological structures (Figures 1, 2; Atwater,1970, Atwater and Stock, 1998, Hyndman et al., 1978). This adjustment is manifest in thecreation of the Explorer microplate at ∼4 Ma through its separation from the larger Juande Fuca plate along the Nootka and Sovanco fault zones offshore northern Vancouver Island(Figure 2.2; Riddihough, 1984). The Queen Charlotte transform fault (QCF) to the northand the Cascadia subduction zone (CSZ) to the south define the Pacific-North America andJuan de Fuca-North America plate boundaries, respectively. The Revere-Dellwood fault(RDF) is a strike-slip fault trending N-NW (N34◦W) which formed around <2 Ma (Davisand Riddihough, 1982) and separates the Explorer microplate from the Pacific plate. TheRDF connects the QCF in the north to the Explorer ridge. The RDF hosts the most12abundant seismicity in the triple junction region, including a Ms 6.8 earthquake in 1992(Cassidy and Rogers, 1995), yet it has received relatively little attention due to its remotelocation.Seismicity offshore BC first garnered attention in the late 1950s-1960s (e.g., Milne andLucas, 1961, Tobin and Sykes, 1968). Event locations prior to 1951 (the installment ofseismometers by the Seismological Service of Canada) have location uncertainties on theorder of 100 km, and through the 1970s, catalogue completeness to magnitude 5 could notbe assumed (Milne et al., 1978). Even with the addition of more stations to the Canadiannetwork in the 1960s, large systematic errors were evident (Milne et al., 1978) and early off-shore locations show a large amount of scatter. Positioning of network stations landward ofoffshore events and the large distances involved (exceeding ∼150 km for events on the RDF;Figure 2.1) combined with the complicated regional velocity structure render hypocentralestimation in this region particularly challenging (e.g., Hyndman et al., 1978, Wahlstromand Rogers, 1992). Currently, routine offshore earthquake location by the Geological Surveyof Canada (GSC) fixes event depths to 5, 10, or 20 km and uses a simple 1D velocity modelwith Moho depth of 36 km.Braunmiller and Na`beˇlek (2002), Ristau et al. (2007), and Kao et al. (2012) computedmoment tensors for the region for larger events (M>4, 4.5, 3.6, respectively) with strike-slip orientations generally consistent with Pacific-North America relative plate motions.The Kao et al. (2012) study employed separate velocity models for different subregionsof offshore Canada. Cleveland et al. (2015) provided the most recent relative offshoreearthquake locations using Rayleigh wave cross-correlations. However, their analysis waslimited to 161 events with M>4.5.Given the rapidly evolving regional tectonics, relocation of offshore seismicity deliversimportant constraints on plate boundary reconfiguration that are not afforded by seafloormorphology or magnetic anomalies (Rohr and Tryon, 2010), which develop over thousands tomillions of years. In this study, we employ an automatic earthquake detection algorithm towaveform data from 1995-2021 to augment the existing Geological Survey of Canada (GSC)earthquake catalogue for the same time period. We relocate these events in a multi-stepapproach to yield high-precision locations and supplement these results with local stressinversions throughout the study area using centroid moment tensor solutions for largerevents (Kao et al., 2012). Our results corroborate previously published models (e.g., Allanet al., 1993, Rohr and Furlong, 1995, Rohr and Tryon, 2010) and allow new insights to bemade into the complex tectonic structure and evolution of the triple junction region, andmicroplate evolution more generally.13−132°−132°−130°−130°−128°−128°−126°−126°48° 48°50° 50°52° 52°54° 54°48 mm/yr41 mm/yrRDFQCFQCTDKDellwood SmtsTWVFPRRExpDeepRiftMRiftNExpRdgSFZCSZHeck SmtsHeckle SmtsSpringfield SmtsWVMVEndvrSeg JdFRdgCobb OSCEndvr OSCNFZEXP MicroplateWBWRHRKRpaleo-QCFBPNootka IslHaida GwaiiRDVNorth AmericanPlateHAKBRDFEJdF Plate−132° −130° −128°48° 48°50° 50°50 km−132° −130° −128°50° 50°52° 52°50 kmPacific PlateRDFTWVFPRRExpDeepRiftMRiftNExpRdgEXP MicroplateWBWRpaleo-QCFRDVNorth AmericanPlateRDFEHRKRDKDellwood SmtsQCFQCT41 mm/yrPacific PlateExpRdgSFZHeck SmtsHeckle SmtsSpringfield SmtsWVMVEndvrSegJdFRdgNFZEXP MicroplateJdF PlateSNFNNFEndvr OSCCobb OSCPRRExpDeepHRKRPacific Plateb)c)b)c)a)100 kmCSZVancouverIslandFigure 2.1: Tectonic setting of the Queen Charlotte Triple Junction region offshoreBritish Columbia, Canada. Stations are indicated with triangles; red areGSC permanent stations, blue are stations used by REST with hatching toindicate temporary deployments, and magenta are common stations usedby the GSC and REST for event detection. Names of relevant tectonicfeatures are abbreviated for clarity in the map. QCF=Queen CharlotteFault, QCT=Queen Charlotte Terrace, RDFE=Northern Revere-DellwoodFault, RDFE=Revere-Dellwood Fault Extension, RDV=Revere-DellwoodVolcanics, TWVF=Tuzo-Wilson Volcanic Field, DK=Dellwood Knolls,RDF=Revere-Dellwood Fault, DS=Dellwood Seamounts, WB=Winona Basin,HAKB=HAKB seismometer, RiftN=North Explorer rift, RiftM=MiddleExplorer rift, ExpDeep=Explorer Deep, PRR=Paul Revere Ridge,WR=Winona Ridge, KR=Kwakiutl Ride, HR=Haida Ridge, BP=BrooksPeninsula, ExpRdg=Explorer Ridge, EXP=Explorer, SFZ=Sovanco FractureZone, WV=West Valley, MV=Middle Valley, NFZ=Nootka Fault Zone,NNF=Northern Nootka Fault, SNF=Southern Nootka Fault, JdFRdg=Juande Fuca Ridge, OSC=Overlapping Spreading Center (black lines point tooverlapping area), CSZ=Cascadia Subduction Zone. Vectors for Pacific-NorthAmerica and Juan de Fuca-North America relative motion are calculated usingthe DeMets and Merkouriev (2016) model. Bathymetry is from the GlobalMulti-Resolution Topography dataset version 4.0 (Ryan et al., 2009).142.2 Tectonic settingThe evolution of the unstable (transform-trench-ridge) Queen Charlotte triple junctionwhere the Pacific (PAC), North America (NAM), and Juan de Fuca (JdF) plates meet,has been the topic of considerable study. The complex tectonics of the northeast Pacificwere first constrained using marine magnetic anomalies (e.g., Mason and Raff, 1961, Riddi-hough, 1982, 1984). However, an absence of visible magnetic anomalies within the Explorermicroplate on the landward side of the triple junction (Davis and Riddihough, 1982) pre-cludes detailed plate reconstructions in this region. Rapid burial of crust due to prolificsedimentation during the last glaciation, hydrothermal alteration, and disorganized volcan-ism are all cited as causes for the absence of marine magnetic anomalies (Carbotte et al.,1989, Davis and Riddihough, 1982, Rohr and Tryon, 2010). Nonetheless, there is generalconsensus that an approximately 20◦ clockwise rotation in Pacific plate motions has causedongoing adjustment within the plate boundary structure (Atwater and Stock, 1998, Barrand Chase, 1974, DeMets et al., 1994, Riddihough, 1982, Rohr and Tryon, 2010), which wesummarize below.2.2.1 Revere-Dellwood fault and its extension, Explorer ridge, andWinona basinThe RDF, as originally identified by Ewing et al. (1968) and named by Srivastava et al.(1971), extends from the southern Dellwood Knoll to the Explorer ridge offshore northernVancouver Island, bounding the Winona basin to the east. The RDF trends 326◦, obliqueto the Pacific-North American relative plate motion which is ∼ 339◦ calculated at 50.5◦N,131◦W using the DeMets and Merkouriev (2016) model. The Winona basin (Srivastavaet al., 1971) is a deep sedimentary basin that is blanketed by up to 8 km of landward-dippingPleistocene turbidite sediments (Davis and Riddihough, 1982). Citing discontinuities ingravity, seismic, magnetic data, and a basement offset visible in side-scan sonar imagery ofthe seafloor, Carbotte et al. (1989) identified a northern extension to the RDF as far northas the Tuzo-Wilson Volcanic Field (TWVF, Figure 2.1, Section 2.2) that they labelled theRevere-Dellwood fault extension (RDFE). In subsequent studies this entire portion from theNorthern Explorer rift to the TWVF has been referred to as the RDF (e.g., Braunmillerand Na`beˇlek, 2002, Rohr and Tryon, 2010, Wahlstrom and Rogers, 1992). Carbotte et al.(1989) also noted the nearby presence of very recent basalt flows 30 km northwest of theTWVF. Rohr (2015) corroborated the RDF extension, and suggested an additional ∼100km of subsurface extension of the RDF toward the QCF by synthesizing these observationswith new seismicity, seismic reflection and gravity data. In this paper we refer to the originalsegment of the RDF from the Northern Explorer rift to the DK as the RDF and includethe entire portion of the fault from the DK to its most north-westerly extent, as defined15−132°−132°−130°−130°−128°−128°−126°−126°48° 48°50° 50°52° 52°54° 54°100 km−132°−132°−130°−130°−128°−128°−126°−126°48° 48°50° 50°52° 52°54° 54°100 km−132°−132°−130°−130°−128°−128°−126°−126°48° 48°50° 50°52° 52°54° 54°100 kmWVMVENDVRPACJdFEXPEXPEXPNAPACJdFNAPACJdFNANFZCSZQCFSFZ JdF RDGEXP RDGEXP RDGSFZJdF RDGNFZCSZQCFRDFRDFE/RDFQCFEXP RDGSFZJdF RDGNFZCSZ~4 Ma ~2-1 Ma ~0 MaFigure 2.2: Schematic cartoon of possible triple junction evolution over the past 4Ma. Double-bar lines represent spreading ridges, whereas single lines representtransform faults. Shaded regions (SFZ, NFZ) indicate broad fault zones. Reddot indicates the Pacific-Juan de Fuca-Explorer/North America intersection.We note the northwestward migration, clockwise rotation, and segmentationof the Explorer ridge, and generation of the RDF and subsequent northwestpropagation with concomitant retreat of the QCF along the continental slope(Davis and Riddihough, 1982, Riddihough et al., 1980, Rohr and Tryon, 2010).The DK, TWVF, and RDV are interpreted as leaky volcanism in a regionof distributed shear between the advancing RDF and retreating QCF (Allanet al., 1993, Dziak, 2006, Rohr and Furlong, 1995, Rohr and Tryon, 2010). SFZrotated clockwise and broadened followed by counterclockwise rotation in thelast 1 Ma (Barr and Chase, 1974, Botros and Johnson, 1988). The NFZ mayhave originated in preexisting weakness generated at a spreading ridge around4 Ma (Rohr et al., 2018) and then underwent progressive rotation and extension(Savard et al., 2020). The cartoons should be understood as a schematic andhighly general overview of fault zone evolution, and many particular details arestill debated.by Rohr (2015), as the RDFE (Figure 2.1). The RDFE skirts the western edge of a smallplateau west of the QCF, and possibly joins the QCF to the north along a seafloor scarp(Rohr, 2015). Numerous shallow, vertical faults oriented east-northeast between the RDFEand the QCF around -131.5◦W, 51.7◦N are likely conjugate shears; some of these faults cutthrough the sediments to the seafloor (Rohr, 2015).The southern terminus of seismic activity on the RDF lies where it passes the NorthernExplorer rift to become the Paul Revere ridge. The Paul Revere ridge is hypothesized to haveformed from a plate weakness that was uplifted through convergence between the Winona16basin and the North American plate at ∼2 Ma and synchronous with the formation of theRDF (Davis and Riddihough, 1982). Linear magnetic anomalies are present on the SWside of the Paul Revere ridge but are absent on its NE side within the Winona basin, whichexhibits low overall magnetization intensity and hence has a poorly constrained history ofplate motion (Carbotte et al., 1989, Davis and Currie, 1993, Davis and Riddihough, 1982,Mason and Raff, 1961). Nonetheless, a drilled sediment core adjacent to the Paul Revereridge (Deep Sea Drilling Project Site 177, presented in Kulm et al. (1973) and interpretedby Davis and Riddihough (1982) and Riddihough et al. (1980)), dredged basalts (Cousenset al., 1984), heat flow measurements (Davis and Currie, 1993), seismic refraction profiles(e.g., Au and Clowes, 1984, Cudrak and Clowes, 1993, Spence and Long, 1995), seismicreflection profiles (e.g., Davis and Riddihough, 1982, Rohr and Tryon, 2010), and gravitymodeling (Davis and Riddihough, 1982, Rohr and Tryon, 2010, Spence and Long, 1995),have all provided some constraint upon plate configurations.Early seismic reflection profiles suggested that the RDF may form a zone ∼5 km wide(Hyndman and Rogers, 1981), although relocated epicenters form a band 10-15 km widecentered on the seismically active southern segment of the fault, extending from the PaulRevere ridge to the Dellwood Knolls (Wahlstrom and Rogers, 1992). While most events onthe southern RDF have slip vectors aligned with the main fault trace, events to the northhave slip vectors oriented more closely with the PAC-NAM relative motion (Braunmillerand Na`beˇlek, 2002, Ristau, 2004, Rohr and Tryon, 2010). Focal mechanisms are nearly allstrike-slip and inferred dextral (e.g., Braunmiller and Na`beˇlek, 2002, Ristau et al., 2007).Some authors have argued for transpression along the fault on the basis of its obliqueorientation to the PAC-NAM relative motion and evidence in the seafloor structures (Rohrand Furlong, 1995, Rohr and Tryon, 2010), whereas others argue against compression alongit (Braunmiller and Na`beˇlek, 2002, Carbotte et al., 1989). Earthquake depths are ∼3-8km (Cassidy and Rogers, 1995, Hyndman and Rogers, 1981) in the young, warm, oceaniccrust. Not all events seem to locate on the main RDF, however. The largest event in theRDF region, the 6 April 1992 Ms 6.8 earthquake, was composed of multiple subevents andoccurred ∼5-6 km to the SW of the RDF fault and propagated NNW, with mechanismsparallel to the PAC-NAM plate motion (Cassidy and Rogers, 1995).The paleo-QCF extends southward of the active QCF as a steep, narrow, continentalslope bounding the Queen Charlotte Sound, broadening as it progresses south of the TWVF(Davis and Riddihough, 1982, Rohr and Furlong, 1995, Rohr and Tryon, 2010) (Figure 2.1).The paleo-QCF may still host some tectonic motion; however, it has not been possibleto resolve recent deformation due to its steep topography (Rohr and Tryon, 2010). Someearthquakes have been detected on the continental slope where there is proposed convergencebetween the Winona basin and North American plates, however, few events exceed Mw >174 (Braunmiller and Na`beˇlek, 2002, Davis and Riddihough, 1982, Kao et al., 2012). Seismicreflection profiles image a dipping oceanic crust overlain by a wedge of sediments thatthickens landward (Davis and Riddihough, 1982, Rohr and Tryon, 2010). Some near-verticalfaults and a small thrust fault about 35 km west of the continental slope are evident fromseismic reflection profiles. Underthrusting of the Explorer microplate in the last 2 Myr,following the retreat of the QCF, may have slowed or ceased with the offshore plate breakand subsequent formation of the Paul Revere ridge and/or the progressive rotation of theExplorer ridge (Braunmiller and Na`beˇlek, 2002, Davis and Riddihough, 1982).The Explorer ridge is presently divided into short, en echelon ridge segments - thesouthern main ridge and two small rifts to the northeast, herein referred to as the Middleand Northern Explorer rifts (Figure 2.1). Spreading is estimated to have shifted from therift segment called Explorer Deep to the Northern Explorer rift about 0.3 million years ago(Botros and Johnson, 1988). The Middle and Northern Explorer rifts and Explorer Deepare bounded to the northeast by the Paul Revere ridge (PRR).The difference in orientation of the Explorer and Juan de Fuca ridges supports indepen-dence of the Explorer microplate (Botros and Johnson, 1988, Riddihough, 1984). Whereasthe Juan de Fuca ridge has clear magnetic anomaly stripes and the characteristic morphol-ogy of a spreading center, the Explorer ridge only exhibits a clear magnetic signal on thePacific side of the ridge (Botros and Johnson, 1988). Further, the Explorer ridge basaltis geochemically distinct from the Juan de Fuca ridge basalt (Cousens et al., 1984). Anintermittent hot spot plume source may feed the main Explorer ridge (Cousens et al., 1984,Michael et al., 1989, Scott et al., 1990), with magma supply greatest at a broad domecentrally located along the Explorer ridge and diminishing to the northeast and southwest(Macdonald et al., 1984). Since 5 Ma, the Explorer ridge has likely moved northward bymeans of asymmetrical spreading and ridge jumps to the Middle and Northern Explorerrifts (Figure 2.2; Botros and Johnson, 1988).2.2.2 Dellwood Knolls & Tuzo-Wilson volcanic fieldThe Dellwood Knolls (DK) (McManus, 1964) are two small volcanic features on the easternside of the RDF. They are composed of N-type mid-ocean ridge basalt (Allan et al., 1993,Cousens et al., 1984) and are about 15 km and 25 km wide with 800m and 1000m of relief,respectively (Riddihough et al., 1980). Age estimates for volcanism range from < 1 Myr oldto ∼2 Myr (Carbotte et al., 1989, Riddihough et al., 1980). The DK have been interpretedas a site of seafloor spreading (Carbotte et al., 1989, Davis and Currie, 1993), or as aconsequence of volcanic “leaking” in a complex transform zone between the paleo-QCF andthe RDF (Allan et al., 1993, Rohr and Furlong, 1995). Indeed, many tectonic maps of theregion show the former interpretation (e.g., Braunmiller and Na`beˇlek, 2002, Cassidy and18Rogers, 1995, Dziak, 2006). The DK region has experienced more sediment accumulation,a further degree of manganese encrustation, and lower magnetization intensity compared tothe TWVF, indicating that it is likely the older of the two features (Carbotte et al., 1989,Cousens et al., 1984, Riddihough et al., 1980). Extensive sediment-covered volcanic flowsare also present in the region between the DK and TWVF (Carbotte et al., 1989, Rohr andTryon, 2010).The two Tuzo-Wilson seamounts (Chase, 1977) are 500m and 700m high compositevolcanic edifices within a graben structure. This structure contains alkaline basalts witha total volume of ∼12 km3 with the entire area referred to as the Tuzo-Wilson volcanicfield (TWVF) (Allan et al., 1993). The volcanics are very recent, likely erupted in thePleistocene-Holocene; the seamounts have little sediment cover (<0.5 km) and a K-Ar dat-ing indicating an age indistinguishable from zero (Allan et al., 1993, Carbotte et al., 1989,Cousens et al., 1984). The TWVF was initially proposed to be the youngest manifesta-tion of a mantle plume hot spot that formed the Kodiak-Bowie seamount chain, based onbasalt geochemistry and the observation that the Tuzo-Wilson seamounts and Kodiak-Bowieseamounts lie on a circle about the Pacific hotspot Euler pole (Chase, 1977). Subsequentstudies have argued primarily in favour of two alternative models for the TWVF as either aspreading center on the basis of plate reconstructions and tectonic history (Carbotte et al.,1989, Davis and Currie, 1993, Davis and Riddihough, 1982), or as a manifestation of leakytransform tectonics. The first reference to the latter idea is apparently due to R. P. Rid-dihough in 1983 (Cousens et al., 1985). More recent geochemical (Allan et al., 1993) andgeomorphological analyses (Rohr and Furlong, 1995, Rohr and Tryon, 2010) support thehypothesis that the TWVF is the manifestation of leaky transform tectonics along the ∼40km wide transtensional stepover between the QCF and RDFE. Nonetheless, basalts of theTWVF are isotopically consistent with Kodiak-Bowie seamount chain, and its location atthe southern extrapolation of the chain may indicate that the southeast outskirts of thehotspot may just barely extend to reach the lithosphere beneath the TWVF (Cousens,2010).2.2.3 Explorer microplate, Sovanco fracture zone, and Nootka fault zoneSubduction of the Juan de Fuca plate at the Cascadia subduction zone to the south andclockwise rotation of relative plate motions around 4 Ma led to the development and sepa-ration of the Explorer microplate (Botros and Johnson, 1988) along the Nootka fault zone(NFZ) (Hyndman et al., 1979) (Figure 2.2). The Explorer microplate likely behaved inde-pendently for a short time after its formation (Rohr and Furlong, 1995) and is experiencinginternal deformation, as evidenced by distributed seismicity; hence, its status as a tectonicplate is debatable (Davis and Currie, 1993). At the large scale, the NFZ, a broad (12-1819km wide) shear zone, must represent left-lateral transform motion since the Juan de Fucaplate on its southern boundary subducts beneath North America (e.g., Braunmiller andNa`beˇlek, 2002, Hutchinson et al., 2019, Rohr et al., 2018). However, source mechanismsfrom major events within or near the eastern end of the fault zone underneath the continen-tal slope offshore Vancouver Island in 2004, 2011, and 2014 (USGS moment magnitudes 6.4,6.4, and 6.5, respectively) indicate right-lateral strike-slip source mechanisms (Hutchinsonet al., 2019, Savard et al., 2020). Hyndman et al. (1979) first identified fault strands of theNFZ in seismic reflection profiles, and proposed that the NFZ is oriented SW-NE towardsNootka Island (Figure 2.1) and aligned with JdF-North America relative plate motions.Dziak (2006) and Rohr et al. (2018) both confirmed and mapped the Northern (NNF) andSouthern (SNF) Nootka faults (Rohr et al., 2018) in greater detail (Figure 2.1). Becausethe NNF and SNF are 17◦ oblique to the more easterly Juan de Fuca - North America rel-ative plate motion (and interpretations of a more easterly NFZ geometry towards NootkaIsland in earlier studies (e.g., Davis and Currie, 1993, Hyndman et al., 1979)), Rohr et al.(2018) further argue that these strands are Riedel shears. Based on a variety of geophysicalevidence, Savard et al. (2020) instead favour a more northerly orientation of the NFZ (c.f.,Hyndman et al., 1979), which closely follows that of the NNF and SNF offshore, then fansnorthwards towards Brooks Peninsula landward of the subduction front. Thus it representsa zone of high strain accumulation between the Explorer microplate and JdF plate thatbroadens to extend between Nootka Island and Brooks Peninsula (Figures 2.1, 2.2; Savardet al. (2020)) along the margin.The extent of possible subduction or underthrusting of Winona basin beneath NorthAmerica remains an open question, though plate motions, seismic refraction profiles, andmodeling all suggest that convergence is occurring (e.g., Braunmiller and Na`beˇlek, 2002,Davis and Riddihough, 1982, Riddihough et al., 1980, Spence and Long, 1995). As analternative to subduction, the Explorer microplate may be amalgamating to the NorthAmerica continental margin, and a new plate boundary developing within it that connectsthe northern Explorer ridge with the eastern Sovanco fracture zone (Rohr and Furlong,1995). In the future, this fault zone would connect the QCF with the Juan de Fuca ridgeand Cascadia subduction zone where the SFZ, NFZ, and Juan de Fuca ridge meet, orpossibly as far south as the Cobb offset (Dziak, 2006, Rohr and Furlong, 1995).The 125 km-long Sovanco (South of Vancouver Island (Pavoni, 1966)) fracture zone(SFZ) connects the northern end of the Juan de Fuca ridge to the Explorer ridge andNFZ, and experiences an annual slip rate of ∼38 mm/yr (Hyndman and Weichert, 1983),estimated from seismicity. It is roughly orthogonal to the strikes of the Explorer and Juande Fuca ridges at N28◦E and N20◦E, respectively (Botros and Johnson, 1988, Karsten et al.,1986), and has an unusual morphology for a transform offset between two ridge segments.20SeaBeam bathymetry data reveal subdued topography at the northwest end of the SFZand rhomboid-shaped uplifted blocks at its eastern side (Cowan et al., 1986). Bounded byNE- and NW- trending lineaments, the blocks appear to have been rotated 30◦ in responseto the dextral shear within the SFZ, hence the name “bookshelf tectonics” (Cowan et al.,1986). In this view the NE-trending bounding sinistral strike-slip faults are reactivatedridge-parallel normal faults. Magnetic anomalies suggest initiation of the SFZ as an offsetaround 7 Ma in an EW orientation, followed by a clockwise rotation starting at 5 Mawith concomitant lengthening through addition of propagating offsets, a brief period ofcounterclockwise rotation, finally approaching its current orientation of N115◦E around ∼1Ma (Figure 2.2; Botros and Johnson, 1988, Cowan et al., 1986).2.3 Data and MethodsWe have compiled a new dataset of 109,095 automatically detected earthquakes from 1995-April 2021 in the British Columbia offshore using the Regressive ESTimator (REST) al-gorithm, described more fully in the supplementary material of Comte et al. (2019) andbriefly summarized here.The REST algorithm (Comte et al., 2019) combines the autoregressive approach ofPisarenko et al. (1987) and Kushnir et al. (1990) with data windowing procedures suggestedby Rawles and Thurber (2015) to generate detections and onset estimates of earthquake Pand S phase arrivals. The autoregressive functions used to detect signals and generate onsetestimates are insensitive to waveform morphology, relying instead on statistical measures ofsimilarity and predictability between a user-defined window of samples and a representationof background noise. To assist in discriminating probable earthquake related phases, RESTenforces a causality condition by quantifying the asymmetry of the estimation function andby requiring the amplitude of the waveform to increase by a certain (user-defined) percentagefollowing the estimated onset. In order to maximize the number of events detected, RESTis designed to be overly inclusive at the early stages of analysis while iteratively identifyingand winnowing false positives at later stages. Arrival time windows are specified based ontrial hypocenters that are determined by grid search, described in Roecker et al. (2006).The hypocenters are revised as onset estimates are iteratively refined until a final detectionand location is determined.The resulting REST dataset displays a marked increase in the number of earthquakedetections when compared to the standard Geological Survey of Canada (GSC) catalogue.For example, the resulting offshore catalog includes 10,467 events with at least 4 paired Pand S picks in the region 47.5-52.5◦N and 127-132.5◦W, an improvement of a factor ∼1.5over the 6,696 previously known events from the GSC that meet the same criterion. Inaddition to data available from the GSC and IRIS waveform archives, we have coordinated21Figure 2.3: Local magnitudes determined by the GSC for the GSC catalog, dividedinto three subregions of the study area. Events included have at least 4 pairedP and S phase picks.the deployment of a new seismic station at the Hakai Institute on Calvert Island as ofOctober 2020, to help fill a large gap in network azimuthal coverage for the region (HAKB)(Figure 2.1).We merged the REST and GSC catalogs for the same time period. However, the RESTalgorithm was applied with a reduced suite of stations (Figure 2.1) to reduce the computa-tional cost, hence the number of common events we determine between the two catalogs isrelatively low. Common events were identified if they had an initial source location within25 km, initial origin time within 8 s, and at least 6 phase picks with arrival time difference< 2 s. A total of 9,929 common events were identified (Figures A.1, A.2). A further 40,839unmatched GSC events were merged with our catalog for a total of 149,878 events. Wethen culled our dataset to subsets with at least 4-paired picks (63,566 events total beforerelocation - 32,650 REST events, 22,944 GSC events, and 7,972 matched events).Furthermore, the REST algorithm computes magnitudes based on direct P amplitudes incontrast to the GSC magnitudes based on the S-wavetrain, precluding a simple comparisonof magnitudes between the two catalogs. Hence, we omit magnitude information for theREST events and only use magnitudes determined from the GSC in this paper. Localmagnitude (ML) detection thresholds for various subregions throughout the study rangefrom <1 in the NFZ and surrounding vicinity to ∼2 in the far reaches of the Explorer ridgeand SFZ (Figure 2.3). REST detections likely reduce the threshold over the GSC becausemost of the REST additions are probably smaller events, however, quantification is difficultdue to different magnitude measures.22During our initial event detection with the REST algorithm, we employed a relativelyslow velocity model for oceanic crust (Figure A.3). Where the velocity structure is unknown,biasing the model to be slow helps mitigate the effect of outliers from errors in phaseonset pick times or unmodeled velocity structure. However, the slow model introduces alandward bias in our locations. To correct this bias, we first relocate our entire dataset withHypoinverse (Klein, 2002), and then use these relocations as input to hypoDD to obtainprecise double-difference locations (Waldhauser and Ellsworth, 2000).2.3.1 Preliminary relocation: HypoinverseDuring relocation with Hypoinverse we use the oceanic velocity model of Waldron (1982)and a number of trial Vp/Vs ratios. We also attempted various methods of incorporating 2Dvelocity structure to account for the ocean-continent transition, including the Spence et al.(1985) model; however, the best results were found using the simple 1D oceanic model. Evenwithin the ocean basin, there is evidence for considerable structural variability. Moho depthfor the Explorer microplate is estimated to be 8-10 km beneath the sea floor (Hutchinsonet al., 2019, Malecek and Clowes, 1978). A study using three OBS 5 km to the west of theRDF estimated the Moho depth to be about 6.5 km, dipping slightly away from the Explorerridge towards the northwest (Cheung, 1978). Sediment thickness also varies considerablythroughout the study area, from up to 8 km in the Winona basin (Davis and Riddihough,1982) to less than 0.5 m on the Tuzo Wilson seamounts (Carbotte et al., 1989). Further,uncertainty surrounding the Vp/Vs structure contributes to location uncertainty throughoutthe study area.Hypocentral depths are poorly constrained in this region due to the station-event ge-ometry (see Figure 2.1). Additionally, there is a strong tradeoff between hypocentral depthz and Vp/Vs, in that increasing Vp/Vs and/or decreasing z shifts locations landward (FigureA.4d,e). We tested varying Vp/Vs and fixed depth (z) values and finally use Vp/Vs = 1.725,calculated using the Wadati method (that is, Vp/Vs = 1 + the best-fit slope of P-wave traveltime versus S-P time (Wadati, 1928); Figure A.5) and fix depth z = 5 km, which yieldsthe best correspondence of seismicity with relevant tectonic features. Epicenters exhibitmore scatter in the variable depth inversions and fixed-depth relocations exhibit strongercorrelation with tectonic features (Figure A.6), so we prefer these epicenters and focus therest of our discussion on these. Owing to the young age of oceanic crust in the region(< 10 Ma), seismicity is most likely confined to the upper several km of oceanic crust,constrained by the 600◦C (Abercrombie and Ekstrom, 2001, Boettcher et al., 2007, Craiget al., 2014) or 700◦C isotherms (Richards et al., 2018, Weins and Stein, 1983). Further,insulation by overlying sediments may reduce the maximum seismogenic depth (Hyndmanand Rogers, 1981). Best-fit hypocentral depths for regional CMT solutions determined by23the GSC exhibit minima at depths < 10 km (Kao et al., 2012), and a previous study in theregion using local OBS arrays estimated earthquake depths to be ∼ 2− 6 km below the topof the oceanic crust (Hyndman and Rogers, 1981). Cassidy and Rogers (1995) estimatedthe focal depth of the 6 April 1992 earthquake to be 6.7-6.9 km below sea level where thewater depth is about 2 km). Thus, we expect the majority of offshore events to be shallow.Further details regarding the Vp/Vs-z analysis are in Appendix A.2.3.2 Re-relocation: HypoDDFollowing relocation with Hypoinverse, we relocate the 4-pair dataset once more by incor-porating delay-time double differences after culling events that suffer convergence issues inHypoinverse. The final, relocated dataset includes 18,441 events. We use the hypoDD For-tran software that implements the double-difference earthquake location method of Wald-hauser and Ellsworth (2000), which updates preliminary earthquake locations by exploitingthe fact that earthquakes closely spaced to one another generate waves that travel nearlythe same path to any given station (Waldhauser and Ellsworth, 2000). This method yieldsrelative locations between earthquakes with high accuracy. Relative locations are adjustedto absolute locations in the algorithm based on the preliminary location estimates and areimproved by including cross-correlation-derived differential times between events.Catalog differential times for a total of 1,969,790 P and 581,709 S phase arrivals aredetermined with the ph2dt program (Waldhauser and Ellsworth, 2000) for earthquake pairswithin a distance of 10 km. We apply the multichannel cross-correlation algorithm (Van-Decar and Crosson, 1990) to all event pairs within a 10 km search radius recorded at agiven station to obtain waveform cross-correlations, employing a total of 3,805,179 P and1,552,759 S cross-correlation times with cross-correlation coefficient > 0.7 for the entireinitial 4-pair-pick dataset.We relocate events separately for regional subgroups, each with a different set of inver-sion parameters that are determined heuristically so as to maximize the number of retainedevents while minimizing mean traveltime residuals. In total, 18,441 out of an initial 21,157events from the Hypoinverse output are relocated. Naturally, subgroups representing re-gions such as the SFZ are characterized by greater average station-event distances and willhave higher starting and ending average catalog and differential traveltime residuals. Weemploy 6 sets of iterations, first placing more weight on the catalog differential times andthen successively increasing the weighting on cross-correlation differential times and down-weighting the catalog times. RMS residual and linkage distance thresholds are graduallydecreased at each set of iterations. For the south-central RDF region, the mean catalogdifferential time residual is reduced from 0.311 s to 0.106 s (65.92% reduction) and meancross-correlation differential time residual is reduced from 0.369 s to 0.013 (96.48% reduc-24Subregion/Statistic RDF-Extension RDF SFZ NFZ Brooks PeninsulaInitial trial sources 4,574 4,305 3,013 7,876 1,362% retained events 86% 90% 82% 87% 94%Catalog RMS residuals (s) 0.062/0.466 s 0.106/0.311 s 0.204/1.92 s 0.132/0.326 s 0.064/0.314 sCorrelation RMS residuals (s) 0.012/0.439 s 0.013/0.369 s 0.079/1.678 s 0.009/0.496 0.075/1.7 sMaximum RMS residual (s) 0.212 s 0.323 s 0.756 s 0.249 s 0.541 sTable 2.1: HypoDD inversion statistics for each subregion. Catalog and correlationtime RMS residuals are given as fractions of final/initial.tion). A more complete suite of statistics is listed in Table 2.1.2.4 ResultsIn this section, we present the relocation results and highlight new spatiotemporal patternsin various subregions of interest (Figure 2.4). Figure A.8 shows the progressive changefrom initial locations, to Hypoinverse relocations, to the final double-difference hypoDDrelocations. We combine CMT solutions (Mw > 3.4) computed by the GSC to provideadditional insight into regional tectonics and augment our relocation results with stressinversions using the CMT solutions. Note, however, that CMT solutions are not relocatedin hypoDD and are generated by the GSC using different velocity models (Kao et al.,2012). Not all events with CMT solutions are represented within the hypoDD catalog sincewaveforms of larger magnitude events do not generally correlate well with those of smallerones. A subset of 334 from a total 705 events with CMTs in our study area are commonto our catalog with the ≤8 s origin time, 25 km discrepancy threshold. Relocation depthstypically do not deviate significantly from the initial (fixed at z = 5 km) depth (Figure A.9).However, the tradeoff between epicenter and depth is expressed as a slightly deepening trendin depth as locations move farther offshore (Figure A.9).2.4.1 Revere-Dellwood fault extension, Tuzo-Wilson volcanic field, andDellwood KnollsThere is an approximately 40 km gap in relocated seismicity northwest of the TWVFalong the northern RDFE that is interrupted by a cross-cutting suite of events. This suiteextends from the QCF to the RDFE along a nearly orthogonal trajectory, and then appearsto veer southward to a point ∼30 km seaward of the RDFE, although the trajectory ismore scattered seaward of the RDFE (Figure 2.5a). The RDFE cross-cutting event suitecomprises a series of bursts of microseismicity that have occurred regularly, roughly every2 years, since 2007 (Figure 2.5). The largest burst began on 17 November 2009, including25−132°−132°−130°−130°−128°−128°−126°−126°48° 48°50° 50°52° 52°54° 54°100 km20000400006000080000100000Total differential time recordingsFigure 2.4: HypoDD relocations, with stations color-coded according to total num-ber of differential time recordings used (catalog P,S + cross-correlation P,S).Grey dashed lines denote bounds of subregions for separate hypoDD inversions(Section 3.2, Table 1).a Mw 6.45 event on 17 November 2009 followed by 5 Mw > 4.1 events with moment tensorsolutions (Kao et al., 2012, Rohr, 2015). The largest event occurred ∼ 3.7 km west of theRDFE, followed by 3 aftershocks along the RDFE and one ∼15 km farther west. The twomechanisms to the west of the RDFE have nodal planes rotated slightly clockwise relativeto those near the trace, and coincide with aftershocks that extend to the southwest. Theseevents may be related to a fault identified by Rohr (2015) 15 km SW of the RDFE (see theirFigure 7). The NNE-striking lineation of aftershocks suggests that theMw 6.45 main shockmay have ruptured as right-lateral slip on the NNE-striking nodal plane (Figure 2.5). Thisorientation does not clearly align with the prevailing PAC motion, nor the RDF and QCF,but it does parallel the local magnetic intensity lineations of the oceanic crust (Figure 2.15;see Discussion Section 5.2).Seismicity in the TWVF between the RDFE and QCF, although distributed, has amore oblique orientation to the RDFE than that associated with the DK farther south(Figure 2.4), and appears to underlie the trend of submarine channels (Figures 2.4, 2.6).26Notably, three sequences of earthquakes in January 2008 (Figure 2.6b), April-May 2015and early November 2015 (Figure 2.6c) punctuate the temporal occurrence of seismicity inthe TWVF. The 2008 sequence involved two mainshocks separated by 31 km on 5 January11:01:01.92 (Mw 6.52) and 11:01:05.00 (Mw 6.48) (Figure 2.6a). Given their orientation,these events may have ruptured left-laterally on a single NE-trending fault plane. Our4-pair catalog includes one foreshock and more than 30 aftershocks between 5-9 January,although the majority occurred between January 5-6. The aftershocks locate significantlynorth of the main events, underneath the Moresby submarine channel (Figure 2.6b). TheApril-May 2015 sequence comprised 127 events, the largest being a Mw 6.2 event on 24April (Figure 2.6c). The November 2015 sequence comprised 25 events, including a Mw4.49 on 22 November (Figure 2.6c). In both cases, event relocations closely align with thesubmarine channels that cross-cut the RDFE, and extend past the RDFE to the SW. Visualinspection of seismograms from some of the larger events in the TWVF (Figure 2.6) clearlyindicate a tectonic (versus debris flow) origin, suggesting that morphology of the Moresbychannel may be tectonically influenced (see Discussion).Immediately north of the Dellwood Knolls, more diffuse seismicity extends landward,approaching the base of the continental slope (and paleo-QCF, Rohr, 2015) (Figure 2.7).This region is quite active, withMw ∼6 events occurring every ∼six years since 2000 (Figure2.7) and regular events in the Mw 2.5-4 range since at least 2002. Notably, little seismicityis detected within the Dellwood Knolls valley east of the RDF.2.4.2 Revere-Dellwood fault, Explorer ridge, and Winona basinThe RDF exhibits the densest seismicity concentration in the Queen Charlotte triple junc-tion region (Figure 2.4). Locations (4-paired-pick) are concentrated on a roughly 12 km-wideband along the fault that extends from the Northern Explorer rift northwest to an abrupttermination at ∼50.7◦ latitude (Figure 2.8). Within this band, microseismicity shows evi-dence for a persistent, near-normal cross-cutting fabric with prominent excursions SW intothe Pacific plate that occurred as sequences in 2012, 2014, 2015, and 2019 (Figure 2.8).These lineations generally coincide with strike-slip CMTs that may represent left-lateralstrike-slip faulting in some instances.The Paul Revere ridge extends from the southeastern end of RDF between the NorthernExplorer rift past the Explorer Deep to the southeast (Figures 2.1,2.8). Although this stretchhas been previously regarded as inactive (Braunmiller and Na`beˇlek, 2002), seismicity isclearly present (Figure 2.8). In particular, a sequence of 58 events from 23-29 December2019 and then later events in 2020 constitute most of the relocated events in this region(Figure 2.8). There are 7 GSC CMT solutions for the 2019 sequence; six on 23 Decemberand one on 25 December. Four lie near the Paul Revere ridge trace and 3 are several27All relocations and CMTs 2009-2010 SequenceYYYY-MM-DDDateYear−132.5° −132° −131.5° −131°51.5°51.75°52°25 km 19982001200420072010201320162019QCF−132.5° −132° −131.5° −131°25 km2009-11-172010-03-032010-06-112010-09-19RDFERDFEQCFMw 6.45 mainshocka) b)c) d)Figure 2.5: Temporal patterns of seismicity in the RDFE area. a) Relocations andCMT solutions from 1996-2020, color-coded by time. b) Relocations from thelate November 2009 sequence through 2010. c) Cumulative seismicity of theRDFE region from 1996-2020. d) GSC magnitudes (ML) for the >=4-pairrelocations from November 2009-December 2010 and CMT solutions.km further to the east. Moment magnitudes are between Mw 4.42-6.22; three events haveMw >6. The seventh CMT solution in the sequence was the largest with magnitude Mw6.22 at 3:36:00.00 UTC on 25 December. Relocations indicate the sequence began at bothnorthwestern and southeastern ends of the PRR where it intersects the Northern Explorerrift and Explorer Deep, respectively. As the sequence proceeded, the southeastern portionremained active, eventually defining a cluster that cuts through the southernmost PRR andinto the Explorer Deep (Figure 2.8).The Explorer ridge is most seismically active on its northeastern end near where itconnects with the Explorer Deep and Middle Explorer rift segments (Figures 2.8,2.10). TheMiddle Explorer rift is less seismically active than the North Explorer rift, with an absenceof events in our catalogue south of 49.5◦N, 130.45◦W. The offsets between the main Explorerridge and the Middle Explorer rift and the Northern Explorer rift segment are highly activeand appear to be oblique to the rifts, yet very few events occur within Explorer Deep (Figure28−132° −131.5° −131° −130.5°51°51.5°25 km1995199820012004200720102013201620192021−132° −131.5° −131° −130.5°51°51.5°25 km 200801-13200802-02200802-22200803-13200804-02200804-22−132° −131.5° −131° −130.5°51°51.5°201504-26201506-25201508-24201510-23201512-22201602-202015 24 April 2015 Mw 6.2, HG4B9 January 2008 Mw 5.92, MOBC5 January 2008 Mw 6.52, BNB22 November 2015 Mw 4.49, HG4B0Time from origin (s)20 40 60 80 100Amplitude (counts)0 20 40 60 80 100Amplitude (counts)0 20 40 60 80 100Amplitude (counts)0 20 40 60 80 100Amplitude (counts)2008Mw 6.482015 Mw 6.22015 Mw 4.49Moresby ChannelMoresby ChannelScott ChannelScott ChannelMoresby ChannelScott Channel2008TWVFTWVFQCFRDFEQCFDateDateYear b)c)a)RDFE25 km2008 Mw 5.92RDFE2008 Mw 6.52 Mw 6.52Mw 5.92Mw 6.48Mw 6.2Mw 4.49QCFFigure 2.6: a) Relocations and focal mechanisms of the TWVF region, b) spatiotem-poral pattern of relocations of the 2008 TWVF sequence and waveforms fromtwo of the larger constituent events, and c) the 2015 April and November se-quences. Beachball sizes are scaled by magnitude. Waveforms are verticalcomponent waveforms with an applied 2-8 Hz Bessel filter of the 2008Mw 6.52,2008 Mw 5.92, April 2015 Mw 6.52 and November 2015 Mw 4.49 events, re-spectively. Red and blue lines denote P- and S- phase picks.2.8). This behavior suggests that most of the active spreading is occurring tectonically atthe North Explorer rift, although the aforementioned 2019 earthquake sequence along thePRR and its association with strike-slip faulting implies some degree of partitioning ofspreading with the Explorer Deep, as suggested by Riddihough et al. (1980).Relocated seismicity along the continental margin adjacent to the Winona basin typ-ically lies ∼0-20 km seaward of the continental shelf (Figure 2.9), striking ∼N45◦W, andterminating near latitude 51◦ in the north and near the NFZ in the south. It may be sig-nificant that this northern termination coincides with the change in seismicity distributionalong strike the RDF from tightly concentrated to more scattered landward through theDK region (see Discussion Section 5.2; Figure 2.4). Most of the continental margin eventsare small (Mw ≤ 4.8). There are only five CMT solutions in the GSC catalog within the1995-2021 time frame with a range of mechanisms including normal, strike slip and thrust.The southernmost CMT is an oblique strike-slip event (Mw 4.8) that occurred on 11 March2005, seaward of Brooks Peninsula (Figure 2.9). While this event did not register withinour culled dataset, a suite of relocated events during the same time period (10-20 March2005) extend in a line towards the continental slope and exhibit similar waveforms (Figure2.9), so they may be associated foreshocks and aftershocks, although neither CMT nodalplane represents a good match with their trend. The NW-trending nodal plane exhibitsbetter alignment with the Kwakiutl ridge orientation (Figure 2.9), possibly reflecting a re-29−131.5° −131° −130.5° −130°50.75°51°25 km19982004201020162021RDFEQCFRDFRDFEQCFRDF−131.5° −131° −130.5° −130°50.75°51°25 kmRDFERDFYearScott ChannelScott ChannelDKDKa)d)c) Yearb)Transpressionalridges19982004201020162021Figure 2.7: a) All relocations in the Dellwood knolls region, color-coded by time.b) Cumulative relocated seismicity in the Dellwood knolls region. Grid linesdenote strong bursts of seismic activity. c) All GSC CMT solutions, color-codedby time. Transpressional ridges that offset the DK on the southwest side of theRDF (Rohr and Furlong, 1995) are indicated. d) Magnitude-time plot showsthat strong bursts of seismic activity occur every ∼6 years since at least 2000.Three ML6 events occurred in the region to the north of the DK, between theRDF and continental slope.lationship with that structure. About 11 hours after the Mw 7.8 27 October 2012 HaidaGwaii thrust event, a Mw 4.5 normal faulting event occurred in the Winona basin (Figure2.9, second focal mechanism from the top), followed by a Mw 3.98 reverse faulting event tothe north (northernmost focal mechanism) in March 2013.2.4.3 Explorer microplate, Sovanco fracture zone, and Nootka fault zoneSeismicity in the Explorer region exhibits a variety of patterns that do not clearly followthe established plate boundaries. The SW corner of the EXP plate boundary, comprisingsouthern Explorer ridge and western SFZ, is largely devoid of seismicity. Seismicity withinthe SFZ occurs primarily in a wide band along the eastern half of the zone (Figure 2.10).Within the Explorer microplate, linear bands of seismicity extend NE from the SFZ with30−131° −130.5° −130° −129.5°50°50.5°199619992002200520082011201420172020−131° −130.5° −130° −129.5°50°50.5°JanFebMarAprMayJunJulAugSepOctNovDecRDFDKPRRExpDeepRiftMRiftNWBDellwood Smts25km−131° −130.5° −130° −129.5°50°50.5° RDFDKPRRExpDeepRiftMRiftNWBDellwood Smts25km−131° −130.5° −130° −129.5°50°50.5°Transpressional ridgeRDFDKPRRExpDeepRiftMRiftNWBDellwood Smts25km−131° −130.5° −130° −129.5°50°50.5°2012 20152014 2019All AllYear Year−131° −130.5° −130° −129.5°50°50.5°Transpressional ridgeRDFDKPRRExpDeepRiftMRiftNWBDellwood [binary].8: Top row shows all relocations from this study in the RDF region (left) andCMT solutions (right), color-coded by time. Bottom two rows show notableearthquake sequences in 2012, 2014, 2015, and 2019, exhibiting cross-cuttingfaulting on the RDF. A transpressional ridge that offsets the DK on the south-west side of the RDF (Rohr and Furlong, 1995) is most evident in the 2012 and2014 subfigures.3125 March 2013 Mw 3.98, BBB28 October 2012 Mw 4.5, PHC11March 2005 Mw 4.8, PHCAmplitude (counts)Amplitude (counts)Amplitude (counts)1e41e21e410.5 0-0.5-1420-2-46420-2-4-6-80 20 40 60 80 100Time from origin (s) 0 20 40 60 80 1000 20 40 60 80 100b)−130° −129.5° −129° −128.5° −128°50°50.5°51°25 km199519982001200420072010201320162019YearMw 4.8Mw 3.98Mw 4.5WRKRPRRa)Figure 2.9: a) Temporal evolution of seismicity in the Winona basin along thecontinental slope with hypoDD relocations overlain by CMT solutions.WR=Winona ridge, KR=Kwakiutl ridge, PRR=Paul Revere ridge. Inferredpaleo-QCF is indicated by a dashed line along the continental margin. b) Ver-tical component waveforms with applied 2-8 Hz Bessel bandpass filter, startingfrom each event’s origin time as 0 s. P and S phase picks are denoted by redand blue lines, respectively. Top: Mw 4.8 oblique thrust event on 11 March2005 offshore Brooks Peninsula recorded at station PHC (Port Hardy, BC).Middle: Mw 4.5 normal faulting event on 28 October 2012 (note there is no Sphase pick) recorded at station PHC. Bottom: Mw 3.98 thrust faulting eventon 25 March 2013, the most northerly of the CMT solutions on the continentalslope, recorded at station BBB (Bella Bella, BC).a trend ∼39◦, approximately parallel to the Explorer ridge and the NFZ (Figure 2.10).Suggestions of these features are seen in the locations of CMTs presented in this study andthe relocations of large earthquakes in this region by Cleveland et al. (2015). The mostprominent examples are defined by sequences of events in 2001 and 2004 (Figure 2.10).Dense concentrations of seismicity exist within and adjacent to the NFZ, especially atits landward end. Many of these events are associated with the 2004 Mw 6.4, 2011 Mw 6.4and 24 April 2014 Mw 6.5 events (Hutchinson et al., 2020, Savard et al., 2020). Many ofour relocated events follow the traces of the NNF and SNF strands, and in particular, theseismicity shows some indication of the southward extension of the latter; see Figures 2.10and 2.15. Both the NNF and SNF were highly active on their northeastern ends in 2014,reflecting the Mw 6.5 earthquake and aftershocks and in smaller part, the presence of theSeaJade2 OBS offshore northern Vancouver Island (for comparison, no noticeable increasein event detections was documented between June-August 2011 during the SeaJade1 OBSdeployment).The swarm of seismicity evidenced by hydroacoustic locations (including 3 events with32−130° −129° −128° −127°48°49°50°50 km−130° −129° −128° −127°48°49°50°−130° −129° −128° −127°48°49°50°50 km−130° −129° −128° −127°48°49°50°JanFebMarAprMayJunJulAugSepOctNovDec−130° −129° −128° −127°48°49°50°50 km−130° −129° −128° −127°48°49°50°−130° −129° −128° −127°48°49°50°50 km−130° −129° −128° −127°48°49°50°50 km 199519982001200420072010201320162019YearMonth2008 20122014 2018AllAllPRRExpDeepExpRdgSFZ CSZHeck SmtsHeckle Smts WVMVJdFRdgCobb OSCEndvr OSCNFZEXPHRKRWRWBSNFNNFPRRExpDeepExpRdgSFZ CSZHeck SmtsHeckle Smts WVMVJdFRdgCobb OSCEndvr OSCNFZEXPHRKRWRWBSNFNNFPRRExpDeepExpRdgSFZ CSZHeck SmtsHeckle Smts WVMVJdFRdgCobb OSCEndvr OSCNFZEXPHRKRWRWBSNFNNFPRRExpDeepExpRdgSFZ CSZHeck SmtsHeckle Smts WVMVJdFRdgCobb OSCEndvr OSCNFZEXPHRKRWRWBSNFNNFPRRExpDeepExpRdgSFZ CSZHeck SmtsHeckle Smts WVMVJdFRdgCobb OSCEndvr OSCNFZEXPHRKRWRWBSNFNNFPRRExpDeepExpRdgSFZ CSZHeck SmtsHeckle Smts WVMVJdFRdgCobb OSCEndvr OSCNFZEXPHRKRWRWBSNFNNFSpringfield SmtsSpringfield SmtsSpringfield SmtsSpringfield SmtsSpringfield SmtsSpringfield SmtsSpringfield Smts2001 200450 km50 km50 [binary] SmtsHeckle Smts WVMVJdFRdgCobb OSCEndvr OSCNFZEXPHRKRWRWBSNFNNFPRRExpDeepExpRdgSFZ CSZHeck SmtsHeckle Smts WVMVJdFRdgCobb OSCEndvr OSCNFZEXPHRKRWRWBSNFNNFFigure 2.10: Top row shows all relocations (left) and CMTs (right) in the Explorermicroplate, SFZ, and NFZ regions. The dashed line encompassing the NNFand SNF is the inferred NFZ. Following rows show notable earthquake se-quences in the Explorer microplate, SFZ, and NFZ regions for selected years2001, 2004, 2008, 2012, 2014, and 2018, exemplifying NE-trending lineationsthrough the Explorer plate and seimicity along the southern NFZ (2012).The OBS from the SeaJade 1 (July 2010-July 2011) and SeaJade 2 (Decem-ber 2013-September 2014) deployments ((Scherwath et al., 2011); see Figure2.1) were used in detecting and locating events near the NFZ in those timeperiods.33Mw = 5.5, 5.7, and 6.0) that migrated northwards along the western edge of Middle Valleyin 2001, as documented by Davis et al. (2004), is represented by a total 184 events withinour catalogue (Figure 2.10). While Dziak (2006) documented considerable seismicity southof the SFZ-NFZ-JDF triple junction near the Heck and Heckle seamounts using SOSUShydrophone arrays, we observe far fewer events in this area, again probably due to thelarge station-event-distances and presumably low magnitudes of these events. We estimatemagnitude detection threshold for the SFZ and surrounding vicinity to be ML ∼ 2 (Figure2.3).2.4.4 Stress inversionsIn this section we present inversions for principal stress directions and relative magnitudesusing the GSC CMT database to facilitate interpretation of our results. Although stressinversions have been previously reported for western Canada (Ristau et al., 2007), that studydid not consider the offshore region south of Haida Gwaii. We employ the StressInversesoftware which performs iterative joint inversion for stress, using Michael’s method (Michael,1984), and fault orientation, using the Mohr-Coulomb criterion, from focal mechanisms(Vavryc˘uk, 2014). In each inversion, we employ 10 stress inversion realizations with randominitializations of faults, 6 iterations, and 200 random noise realizations for estimating theaccuracy of the solution. In addition, to the principal stress orientations, we recover theshape ratio R as:R =σ1 − σ2σ1 − σ3 , (2.1)which describes relative magnitudes of the principal compressive stresses defined by σ1 ≥σ2 ≥ σ3. As σ2 → σ1, R→ 0 and the stress regime approaches uniaxial tension; conversely,as σ2 → σ3, R→ 1 and the stress regime approaches uniaxial compression.Figure 2.11 displays pressure and tension axes of contributing earthquakes, confidencelimits of the principal stress orientations, histogram distribution of R, and the selectedevents used for each inversion for various subsets of the CMT database. Table 2.2 displaysthe principal axes orientations for each subregion. As expected, the stress regime throughmost of the region is characterized by an approximately vertical orientation of σ2 (plungesare 70.9◦ for the bulk stepover region and 68.9◦ for the Explorer microplate) manifestingthe dominance of strike-slip faulting, aside from the Juan de Fuca ridge where σ1 is closestto vertical (Table 2.2). The σ1 trend exhibits a general counter-clockwise rotation movingsouth through the region from 36.1◦ in the RDFE, 16.2◦ in the TWVF, to 10.1◦ in the DK.It increases slightly to 13.7◦ in the RDF, before shifting due north along the NFZ (3.4◦) and34σ1Subregion Trend Plungesouthern HG 215.0 2.7RDFE 36.1 4.3TWVF 16.2 7.7DK 10.1 32.4southern RDF 13.7 10.8WB 356.7 35.5All stepover 15.5 7.7Explorer 356.8 6.6NFZ 3.4 20.1JDF 196.5 70.2σ2Subregion Trend Plungesouthern HG 311.5 67.1RDFE 292.0 72.8TWVF 245.0 78.5DK 231.1 50.0southern RDF 246.9 72.4WB 91.0 5.9All stepover 262.7 70.9Explorer 249.3 68.9NFZ 222.0 65.0JDF 6.9 19.6σ3Subregion Trend Plungesouthern HG 123.9 22.8RDFE 127.4 16.6TWVF 107.3 8.6DK 114.1 20.8southern RDF 106.4 13.8WB 189.1 53.8All stepover 108.0 17.4Explorer 89.2 19.9NFZ 98.8 14.4JDF 98.0 3.1Table 2.2: Principal stress orientations (trend, plunge) from regional stress inversions.in the Explorer (356.8◦) region. Maximum compression σ1 is oriented at 62◦ to the RDFE(average RDFE strike trends ∼334◦) but adopts a progressively smaller angle to the southof 50◦ in the TWVF and 47◦ in the southern RDF, where the RDF strikes 326◦. In general,R values are close to but slightly below the neutral value of R = 0.5, reflecting a slightpreference for tension over compression. The one exception is the continental shelf regionimmediately NW of Vancouver Island, where R = 0.5. Stress is poorly constrained by thefew (5) CMTs but σ3 is closest to vertical (plunge 53.8◦, and σ1 is directed northward. Rvalues are greatest in the TWVF and southern Haida Gwaii regions at R = 0.86 and 0.88,respectively. In the RDFE, R falls to 0.68, whereas the DK region exhibits a significantlylower value R = 0.15. At the northern JDF ridge, σ3 is horizontal, trending at 98◦, and σ1,σ2 form a vertical girdle as evidenced by low R = 0.02.2.5 DiscussionHere we compare our relocations with previous work on seismicity offshore BC, contextualizeour results in the broader tectonic regime, and discuss implications for its present state andfuture evolution.2.5.1 Comparison to previous studiesWe use the same 8-second origin time, 25-km location discrepancy criteria to comparecommon events as described in Section 2.3. The time duration of our catalog overlapsslightly with Braunmiller and Na`beˇlek (2002) from 1995-1998 and, though our 4-pairedpick criterion eliminates many events pre-2000, there remain 9 common events (out of theirtotal 76 in that time period; Figure 2.12a). Locations for these events are similar, with amean shift of 9.42 km with no clear bias. Our catalog also overlaps that of Cleveland et al.35DKTWVFR=0.68R=0.86R=0.45All stepoverWBsouthern HGR=0.88RDFER=0.38R=0.35R=0.15R=0.43R=0.46R=0.02JDFNFZEXPR=0.50R=0.88RDFEnorthernRDFFigure 2.11: Results from stress inversion for subregions in the RDF-QCF systemlisted from north-south (southern Haida Gwaii (HG), northern RDFE (alsoin Figure 5), TWVF (also in Figure 6), DK (also in Figure 7), RDF (also inFigure 8), WB (also in Figure 9), entire RDF-stepover region, Explorer mi-croplate, NFZ, and JDF ridge (also in Figure 10)). First columns are principalstresses and P/T axes with azimuths of σ1, σ2, and σ3 given below, secondcolumn shows the confidence limits of the stress axes, third column shows theshape ratio distribution and the fourth column shows the selected focal mech-anisms used in the respective inversions. The southern Haida Gwaii inversionuses only events before October 2012 and along the QCF (to exclude the largeMw7.8 thrust event offshore Haida Gwaii).36(2015) from 1995-2012, with 31 common events (out of 152 total relocations in their studyduring that time period; Figure 2.12b). Again, our relocations are similar, with a mean shiftof 8.52 km and no clear bias. Note that Cleveland et al. (2015) estimated precise relativelocations using surface waves for events with magnitude > 4.5, and then applied a commonshift, based on the Cassidy and Rogers (1995) 1992 Ms 6.8 earthquake location, to obtainabsolute locations, in contrast to the approach outlined in Section 2.3.1. The mean shiftsquoted above for common events are larger than that for the GSC CMT catalog of 7.23 km(Figure 2.12c) which employs tailored velocity models for SFZ, RDF, QCF and VancouverIsland subregions (Kao et al., 2012).In addition, the SOSUS event catalog of Dziak (2006), based on offshore hydrophonearray recordings, affords another opportunity for comparison with our event relocations(Figures 2.12d, 2.13). Their catalog overlaps ours from 1995-2008, during which we find146 common events. Location discrepancies are larger for these events (mean 11.83 km).The SOSUS locations are more scattered, especially throughout the Sovanco and southwestExplorer microplate region, although seismicity concentrations along the Juan de Fuca ridgeand SFZ are obviously associated with major tectonic features (Figure 2.13). Around theRDF and northern Explorer ridge, the SOSUS locations appear to be displaced furtheroffshore compared to our relocations. The 1992 Heck seamount earthquake sequence is alsoclearly delineated but took place prior to the modern CNSN station upgrade and so is absentfrom our catalog. Of note, Dziak (2006) resolves earthquake locations within the SW cornerof the Explorer microplate along the western SFZ and around the southern Explorer ridge,where our study and others do not. A lack of seismicity in this area has been interpretedto manifest the initiation of a new transform fault cross-cutting the Explorer plate (e.g.,Braunmiller and Na`beˇlek, 2002, Rohr and Furlong, 1995). SOSUS locations are derivedfrom recordings of earthquake T phases that travel through the oceanic sound fixing andranging channel which generally allow increased detections at lower magnitudes. Nominalquoted errors in the SOSUS locations are within 0.002◦ latitude, 0.04◦ longitude, and 0.04-2.3 s of origin time (Dziak, 2006). However, lack of close correspondence with concentratedrelocated seismicity along, for example, the north and western Explorer ridges raises somequestions as to their veracity and the presence of seismicity in the SW Explorer microplatecorner.2.5.2 Revere-Dellwood-Queen Charlotte fault migrating pull-apartsystemIn this section, we consider our observations in the context of the view that the RDF-QCFsystem around the TWVF and DK involves transtensional stepover tectonics (Allan et al.,1993, Braunmiller and Na`beˇlek, 2002, Rohr and Furlong, 1995, Rohr and Tryon, 2010).37a) b)c) d)Figure 2.12: Comparison of relocations from this study with a) Braunmiller andNa`beˇlek (2002) (9 events; top left), b) Cleveland et al. (2015) relocations(31 events; top right), c) GSC CMT locations (Kao et al., 2012, 334 events;bottom left), and d) SOSUS hydrophone locations (Dziak, 2006, 146 events;top right)38−131°−131°−130°−130°−129°−129°−128°−128°−127°−127°48°49°50°50 km−131°−131°−130°−130°−129°−129°−128°−128°−127°−127°48°49°50°50 kmSOSUSThis studySOSUSThis studyRDFPRRExpDeepExpRdgSFZCSZHeck SmtsHeckle SmtsSpringfield SmtsWVMVJdFRdgCobb OSCEndvr OSCNFZEXP WRHR KRWB RDFPRRExpDeepExpRdgSFZCSZHeck SmtsHeckle SmtsSpringfield SmtsWVMVJdFRdgCobb OSCEndvr OSCEXPWRHRKRWBEndvrSegEndvrSegNFZFigure 2.13: Comparison of relocations from this study with the hydrophone loca-tions from Dziak (2006). The hydrophone locations are culled to those withat least 4 hydrophone recordings. The left and right figures display the samelocations; the left displays the hydrophone locations on top and the right dis-plays our relocations on top.An additional important consideration is a counteracting component of compression thatarises through Pacific plate motion obliquity relative to the Pacific-North America plateboundary (e.g., Carbotte et al., 1989, Rohr, 2015). Pacific plate motion obliquity alongthe QCF system to the north (i.e. north of ∼52◦ latitude) is known to play an importantrole in deformation style (Brothers et al., 2020, Hyndman, 2015, Rohr et al., 2000, Trehuet al., 2015). In particular, as obliquity increases from north to south, deformation styletransitions from the nearly pure strike-slip transform off the Alaska Panhandle, to slip that ispartitioned between strike-slip faulting along a crustal sliver (the Queen Charlotte Terrace,Figure 2.1) and Pacific plate underthrusting off southern Haida Gwaii. Pacific plate motionobliquity in the latter case is variously estimated between 5.6◦ to >15◦ (c.f., Brotherset al., 2020, Hyndman, 2015, Trehu et al., 2015). Although absolute values are debated,comparable levels of convergence must persist south into the RDFE-QCF transtensionalsystem; for example, Pacific motion is ∼19◦ oblique to the RDF trace at 50.7◦N, 130.6◦Was calculated from the DeMets and Merkouriev (2016) model. The southward change toa transtensional stepover regime is marked by focal mechanisms that are overwhelminglystrike-slip, as indicated by near-vertical σ2. This orientation contrasts with that in southernHaida Gwaii where mixed thrust and strike-slip partitioning is manifest by σ2 ≈ σ3 (R =0.88) with the σ2, σ3 axes defining a vertical girdle (Figure 2.11, Ristau et al., 2007).39Nonetheless, σ1 orientation in southern Haida Gwaii and the northern RDFE region remainseffectively the same.Our documentation of seismicity extending from the QCF to the RDFE, 40 km NWof the TWVF, supports the arguments of Rohr (2015), that the RDFE is active to 51.9◦latitude (c.f., Carbotte et al., 1989). Moreover, the orientation of this seismicity as roughlyorthogonal to the RDFE, and its association with basalt flows (Allan et al., 1993, Daviset al., 1987, Rohr, 2015) implies that this region is analogous to the TWVF in an earlystage of development. It also supports the migrating transtensional stepover system model(Allan et al., 1993, Rohr, 2015, Rohr and Tryon, 2010) and, in particular, an extension tothat model that is detailed later in this section.The general paucity of normal faulting events in both the TWVF and DK rift valleys issomewhat at odds with the general transtensional tectonic setting. The few non-strike-slipevents within the northern RDFE and TWVF tend to have reverse mechanisms (Figure2.6), as reflected by shape ratios R > 0.5. This observation suggests that Pacific plate com-pression presently dominates over transtension in this region, although not to the degreeobserved to the north in southern Haida Gwaii (Figure 2.11, Ristau et al., 2007). Eventswithin the DK, while again dominantly strike-slip, yield R = 0.15 indicating a larger com-ponent of extension than the TWVF and RDFE to the north. Relocated seismicity in theDK region is unevenly distributed, with a larger number and broader distribution of eventswithin and to the northwest of the northern Dellwood Knoll. In contrast, only a few scat-tered events lie within the southern Dellwood Knoll with a far denser concentration to thesouthwest along the RDF (Figure 2.7). This pattern is also suggested in the early OBSstudy of Hyndman and Rogers (see their Figure 5 1981), although these authors attributedit to either faulting on two sides of a rift, or the development of a NE-trending transformoffset (Riddihough et al., 1980). The DK valley separating the northern and southern DKmay have formed more recently through faulting across a once contiguous volcanic edifice(Allan et al., 1993, Rohr and Furlong, 1995). Rohr and Furlong (1995) note that a slipvector coinciding with PAC-NA relative motion restores this edifice from the current DKlocations within ∼0.3 Ma. As they argue, these observations favor the interpretation thatextension in the DK has occurred in a transtensional setting, after their formation fromleaky volcanism along the RDF in a migrating stepover between the QCF and RDF. Weconsider the origins of the DK valley further below.Southeast of 50.7◦ latitude, a sharp transition occurs from the intermittent, diffuseseismicity characterizing the RDFE, TWVF and northern Dellwood Knoll to concentratedseismicity along the RDF from the southern Dellwood Knoll to the southeast. This along-strike change coincides with the inboard location of the Winona basin and the onset ofsparse seismicity along the base of the continental slope to the southeast (Figure 2.4). The40stress regime along the continental slope is poorly defined but suggests that σ3 rather thanσ2 is closest to vertical (σ3 plunge 53.82◦ versus σ2 plunge 5.90◦, with standard deviationsof 10.86◦ and 10.47◦, respectively) (Figure 2.11), implying compression. The abrupt changein character of seismicity affords insights into the nature of transtensional tectonics in thepresence of regional compression as previously proposed and developed by Allan et al.(1993), Rohr (2015), Rohr and Furlong (1995), Rohr and Tryon (2010).We consider first the region of distributed seismicity that extends, along-strike, betweenthe northern Dellwood Knoll and the northern limit of seismicity across the RDFE and,normal-to-strike, between the RDFE and continental slope (e.g., Figure 2.4). We suggestthat this region can be viewed as the actively-deforming component of a migrating pull-apart system (Figure 2.14). By definition, a pull-apart basin is a topographic depressionresulting from tectonic extension. Importantly, however, transtension between the RDFEand QCF appears to be accommodated mostly via strike-slip faulting and volcanism, with-out significant subsidence. This style contrasts with the Winona basin to the south, thenow inactive component of the migrating pull-apart system, which exhibits clear evidenceof previous tectonic subsidence (e.g., Davis and Riddihough, 1982). We therefore term thecomponent of the migrating pull-apart system northwest of 50.7◦ the active “pull-apartzone.” The stark difference in deformational styles may be due to a number of factorsdiscussed further below.A lack of widespread tectonic subsidence in the northern, active component of the mi-grating pull-apart system (northern DK/TWVF/RDFE/QCF, Figures 2.14, 2.15) may beexplained by a mechanically weak lithosphere. A weak lithosphere is implied by high heatflow values (average of 220 mWm−2 around the TWVF (Hyndman et al., 1978)), extensivevolcanism and shear faulting, along with a balancing component of compression arising fromPacific plate obliquity as indicated by comparatively high shape ratios (northern RDFE,R = 0.68; TWVF, R = 0.86). The thermal regime may reflect a mantle plume influence(Chase, 1977, Cousens et al., 1985) which is supported by a basalt composition in theTWVF that falls at the enriched end of the mid-ocean ridge basalt (MORB) spectrum andoverlaps isotopically with the Kodiak-Bowie seamount chain (Allan et al., 1993, Cousens,2010, Cousens et al., 1985).Thermomechanical modelling of pull-apart basin development in continental settings(Petrunin and Sobolev, 2006, 2008, van Wijk et al., 2017) provides a starting point fordiscussion and insight into factors that control deformation in the present context. In thesestudies, the subsidence rate in pull-apart basins can be viewed as the result of competingbrittle extension and compensating ductile flow. The results of Petrunin and Sobolev (2008)indicate that pull-apart length, sediment thickness, and deformational pattern are governedby the strike-slip displacement, thickness of the brittle crustal layer beneath the basin,41viscosity of the underlying ductile lithosphere, which are dependent upon temperature,composition, and material softening. Extensional thinning, which mostly occurs within thebrittle crust, is proportional to its initial thickness. Models with weaker rheologies/higherheat flow exhibit lower subsidence rates and slightly greater Moho uplift (Petrunin andSobolev, 2008).Although these modelling studies have concerned continental lithosphere, we anticipatethe qualitative behaviour, in particular the effect of increased temperatures in reducing sub-sidence, to hold in the current setting which may be closer to oceanic in nature (Spence andLong, 1995). The efficacy of the “brittle-brick stretching” model employed by Petrunin andSobolev (2008) in explaining deformational patterns across a range of rheological parameterssupports this contention. Accordingly, we hypothesize that a lithosphere weakened throughthe influence of Kodiak-Bowie plume interaction (including mantle upwelling, decompres-sion melting and concomitant volcanism) in this active component of the migrating pull-apart system leads to little or no subsidence and near isostatic compensation. This oceanicsetting therefore represents an end-member case in which a thin, brittle oceanic crust issupported by underlying warm and ductile mantle lithosphere. Volcanism throughout theactive pull-apart zone, as evident at the TWVF and nearby sediment-covered volcanic flows(Carbotte et al., 1989, Rohr and Tryon, 2010), is also observed in many pull-apart zonesworldwide. For example, plutonic emplacement has been observed to fill dilational gaps ina transtensional stepover in the Sierra Nevada (Tikoff and Teysier, 1992); similarly, exten-sional faults in the Salton sea pull-apart basin may have provided pathways for volcanism(Brothers et al., 2009).The length of the active pull-apart zone (Figure 2.15) is approximately l ≈ 158 kmalong-strike (points 50.7056◦N, 130.5648◦W to 51.8674◦N, 131.8673◦W); width near thenorthern DK is w ≈ 49 km measured from the RDFE to the paleo-QCF at the base ofthe continental slope (Rohr and Tryon, 2010). These values yield a l/w ratio of ∼3 forthe active pull-apart zone that agrees well with the results of Aydin and Nur (1982) intheir comprehensive analysis of 70 pull-apart basins within transtensional settings acrossthe globe, representing scales from m to hundreds of km. Similar l/w scaling has beenreported in numerous subsequent observational and modeling studies, and is found to belargely independent of lithology (e.g., Aydin and Schultz, 1990, Basile and Brun, 1999,de Joussineau and Aydin, 2009, Fossen and Rotevatn, 2015, Long and Imber, 2011). Theindependence of this scaling across a wide range of environments lends further credence tothe notion that fundamental aspects of pull-apart basin development are independent ofdetailed rheology or crustal type.Several mechanisms can potentially explain the scaling of pull-apart basins. For exam-ple, the global average l/w may reflect the apparent limited lifespan of overlapping strike-slip42fault systems (< 10 Myr) (Mann, 2007, Mann et al., 1983, Xie and Heller, 2009). Alterna-tively, Aydin and Nur (1982) suggested that a roughly constant l/w is maintained throughcoalescence of adjacent pull-apart basins into a larger, wider composite basin with increas-ing overlap and/or the formation of fault strands that parallel the bounding faults. A morecompelling explanation for constant l/w scaling in the present context is provided by thepull-apart basin migration mechanism described by ten Brink and Avraham (1989) andWakabayashi et al. (2004) (their Figure 8c) wherein scaling is maintained post basin incep-tion through migration of an (approximately constant area l × w) active step-over regionas one of the bounding faults advances and the other recedes. We posit that the overlaplength l between an advancing RDF/RDFE and retreating paleo-QCF (Rohr, 2015, Rohrand Tryon, 2010) has remained approximately constant in its recent northwestward migra-tion (see Figure 2.14). During stepover propagation, a zone of minimal principal strain atthe tip of the propagating fault, within the stepover region, is predicted to form a local basinhigh (Nabavi et al., 2018). Local uplift of oceanic basement underneath and immediatelyadjacent to the RDFE within the stepover region (inboard) (Rohr, 2015) may be evidenceof this process.One interpretation of the migrating pull-apart system model (as we will refer to it,after Wakabayashi et al. (2004)), to distinguish it from the more general transtensionalmigration model proposed by Rohr and Tryon (2010) and Rohr (2015) is that strike slipmotion must take place on the bounding RDFE and QCF. Yet, interestingly, there is alack of defined seismicity along the RDFE and QCF in our study period; most observeddeformation is distributed between them. The QCF is only evident in multichannel seismicreflection imagery as far southeast along the margin as the TWVF (Rohr and Tryon, 2010),and seismicity is nearly absent along the margin inboard of the DK (Figure 2.9). First, it isplausible that, given the high sediment flux from the continental shelf, basement offset of theQCF may be obscured (Carbotte et al., 1989). Second and more generally, it is conceivablethat the absence of seismicity, as well as subdued and discontinuously distributed seismicityalong both the QCF and the RDFE to the northwest, are consequences of aseismic creep.Oceanic transform faults commonly exhibit low seismic coupling, that is, most of the slip inthe upper crust is thought to be relieved by aseismic creep (Boettcher and Jordan, 2004).Indeed, slow aseismic slip transients and fault creep are commonly observed at oceanictransforms (e.g., Beroza and Jordan, 1990, Liu et al., 2020, Materna et al., 2018), which maybe due to fluids within the crust and mantle serpentinization (Boettcher and Jordan, 2004).Moreover, aseismic creep and slow fault slip are also observed at continental transformsworldwide, including the transpressive Alpine fault in New Zealand (e.g., Chamberlainet al., 2014, Wech et al., 2012) and central San Andreas fault in California (e.g., Carpenteret al., 2011, Rolandone et al., 2008, Steinbrugge et al., 1960).43We interpret the RDF, paleo-QCF and Winona basin to be now former componentsof the migrating pull-apart system model (Figure 2.14). The inception and initial evo-lution of pull-apart basin formation are likely to have been complex, controlled by theridge-transform-trench configuration at the original triple junction location off northernVancouver Island (Atwater, 1970, Barr, 1972, Chase et al., 1975, Riddihough, 1982) and itsresponse to the change in Pacific plate motion over the last 20 Myr (DeMets and Merk-ouriev, 2016). We posit that the Winona basin formed through subsidence accompanyingtranstension starting at ∼ 1.8 Ma during inception of the RDF from a fracture zone off-set and retreat of the QCF away from the former triple junction off Brooks Peninsula,within a context otherwise similar to that portrayed by Davis and Riddihough (1982, seetheir Figure 13). Pull-apart basins commonly experience rapid subsidence on the order ofhundreds of meters to several kilometers per million years, controlled by the fault systemgeometry, boundary fault friction, heat loss, and rheology of the crust and upper mantle,as discussed above (Petrunin and Sobolev, 2006, 2008, ten Brink and Avraham, 1989, vanWijk et al., 2017, Xie and Heller, 2009). Since this is the inactive component of the mi-grating pull-apart system, strike-slip deformation has now been entirely transferred to theRDF from the paleo-QCF. Positive flower structures underlying the Winona ridge, whichlikely formed after the onset of subsidence (Davis and Riddihough, 1982), may be a resultof the regional transpression taking over in the inactive basin (Wakabayashi et al., 2004).Uplifted and translated ridges of the DK on the southwest side of the RDF identified byRohr and Furlong (1995) (Figure 2.8) provide further evidence of transpression. Structuraland rheological differences between the Winona basin and the active pull-apart zone to thenorth likely contribute to the stark differences in deformation described above. We positthat a mechanically stronger Winona basin (Davis and Clowes, 1986, Davis and Riddihough,1982) is responsible for the increased subsidence relative to the active pull-apart zone tothe northwest. Significant volcanism within the active pull-apart zone, plausibly related tothe Kodiak-Bowie hotspot as previously mentioned, may further emphasize these structuraland rheological differences. Furthermore, continental margin events (Figure 2.9) appear torepresent a form of slip-partitioning guided by the paleo-QCF that now accommodates theresidual component of Pacific plate convergence, although the small number of mechanismsprevent us from drawing definitive conclusions here. The change in pattern of seismicity atlatitude 50.7◦N implies that the southern Dellwood Knoll is the most recently deactivatedcomponent of the formerly active Winona basin and suggests the possibility that formationof the DK valley (Allan et al., 1993, Rohr and Furlong, 1995) may have been a response tothis deactivation.In addition to a pronounced rheological contrast, the punctuated distribution in seis-micity within the active pull-apart zone may exhibit the development of conjugate Riedel44RDF RDFQCF2-1 Ma 0 Mapaleo-QCFWB WBPAC-NAPAC-NADKRDFEQCFEXP EXPTWVFDKFigure 2.14: Schematic diagram illustrating the migrating pull-apart system modelat time snapshots 2-1 Ma, during formation of Winona basin and DellwoodKnolls (left), and present-day (right). The shaded regions represent activeportions of the transtensional pull-apart zone. Open triangles indicate vol-canic centers. The active pull-apart zone dimensions, length l and width w,are preserved through activation and deactivation of the basin as the stepovermigrates. Details of the Explorer ridge system and incipient subduction north-east of the WB and the northern RDFE/QCF have been omitted for simplicity.Modified from Wakabayashi et al. (2004).(R’) shears. In physical models and natural examples of developed, rhomboidal-shapedpull-apart basins, R’ shears that connect the bounding faults usually develop at ∼ 65◦ tothese faults during early stages of shear deformation (Basile and Brun, 1999, Tchalenko,1970, van Wijk et al., 2017). Seismicity at the northern DK, TWVF, and RDFE generallytrends at highly acute angles to the RDFE/QCF; for example, the northern DK seismicitytrends about 70◦ to the RDFE (Figure 2.7). Slip between the master faults may exploitpreexisting faults that form at a slightly higher angle than 65◦ instead of creating newfaults, and we posit that such preexisting faults may have been generated at the Explorer45ridge. Moreover, some of the seismicity lineations define faults that extend outboard of theRDFE and RDF (Figures 2.5,2.6). Deformation has been observed to occur outside of thebounding faults of transtensional stepover systems as wing-tip or horsetail fractures at theactively propagating end of the fault (Mann et al., 1983). Recent finite-element modelingof transtensional stepover systems predicts the formation of these features parallel to σ1(Nabavi et al., 2018). In general, the RDFE and TWVF seismicity that extends outboardof the RDFE is not parallel to the local σ1 trend (Figures 2.5, 2.8). The most extensiveof the observed lineations outboard of the RDFE crosses the fault trace and changes trendto follow that of the local Pacific plate magnetic anomalies (Figure 2.15a). We hypothesizethat these events represent deformation controlled by preexisting weaknesses in the oceaniccrust created during seafloor generation. Reactivated normal faults within translated litho-sphere have also been cited to explain structures imaged by seismic reflection north of HaidaGwaii, extending off the QCF (Trehu et al., 2015).The presence of translated oceanic crust with preferential directions of weakness mayplay a role in submarine canyon development. Large earthquakes (Mw ∼ 7) have been shownto influence submarine channel evolution by triggering debris flows that modify the canyonenvironment (Mountjoy et al., 2018). Earthquakes with magnitudes up to Mw ∼6.52 occurregularly in the TWVF and north of the DK; in the latter case with a ∼6-year recurrence(Figure 2.7). Given the strong correlation between relocations and submarine channelsnear the TWVF (Moresby channel, Figure 2.6) and to a lesser extent, the DK (Scottchannel, Figure 2.7), and the regular occurrence of Mw ∼ 6 earthquakes we posit that theformation and development of these channels is tectonically controlled, and augmented byheavy sedimentation during the Pleistocene (Carbotte et al., 1989, Davis and Riddihough,1982).As for other oceanic transforms (Boettcher and Jordan, 2004), earthquakes with mag-nitudes 6-7 have occurred regularly along the RDF, DK, and TWVF and, in particular,during our study period (Figures 2.6, 2.7). Kuna et al. (2019) formulated a model foroceanic transform seismic behavior based on the Blanco Transform fault wherein the uppercrustal portion of the fault is fully seismically coupled but the mantle portion shallowerthan the 600◦ isotherm, creeps. They suggest that mantle creep may promote the genera-tion of large seismic events quasi-periodically and that foreshocks to these events may be aresponse to shallow mantle creep preceding the event. The ∼6-year recurrence of moderate(Mw ∼6) earthquakes in the DK area may represent such a periodic seismic cycle. Futurework could test this hypothesis through improved station geometry (e.g., ocean bottom seis-mometer deployments), allowing lower magnitude detection thresholds and more accuratehypocentral locations.46−134°−134°−132°−132°−130°−130°−128°−128°−126°−126°48° 48°50° 50°52° 52°50 km−6000600Magnetic intensity (nT)Cobb OSCEndeavour RidgenTa) b)Nodata48 mm/yr60-60 nTSETBNFZ−130°−130°−128°−128°48° 48°50° 50°50 kmENDSETBWVMVSNF extension?41 mm/yrActive pull-apart zoneNFZESDZESDZb)41 mm/yrFigure 2.15: Relocations with magnetic intensity from the EMAG2 model (Mauset al., 2009). Interpreted linear features, many aligned with the magneticanomaly fabric, are indicated with dashed black lines. SETB = SouthernExplorer Transform Boundary. ESDZ = Eastern Sovanco Deformation Zone.a) Entire study region, with the active pull-apart zone portion indicated. b)Inset of the Explorer-NFZ-SFZ region. Possible extension of the SNF is indi-cated by a dotted line. NFZ interpretation (dotted line) is from Savard et al.(2020) and Merrill et al. (2022)2.5.3 Explorer microplate deformationLike previous work (Braunmiller and Na`beˇlek, 2002, Cleveland et al., 2015, Dziak, 2006),we document seismicity throughout the Explorer plate south of the Winona basin, but witha more pronounced NE trending linear fabric (Figure 2.15). Seismicity is concentratedalong the NFZ and eastern SFZ, and spills into the southern Explorer microplate. Linearbands of seismicity appear to originate in the SFZ, and extend northeastward over distances≥ 100 km, approximately parallel to the dense seismicity associated with the NFZ. Principalstresses in the Explorer region, like those to the north, are again dominated by near-verticalσ2, with R < 0.5 indicating relative extension. At the Juan de Fuca ridge, σ3 is horizontaland perpendicular to the spreading axis while σ1 ∼ σ2 form a vertical girdle, marking theshift to a tensional, spreading regime. The systematic counterclockwise rotation in σ1 fromN36◦E at the northern RDFE to north-south in the Explorer region reflects the competinginfluences of large-scale interactions between the Queen Charlotte plate boundary in thenorth and spreading along the Juan de Fuca ridge to the south. It is also consistent withthe observation that earthquake slip vectors rotate counterclockwise from NNW along theRDFE between the TWVF and DK to NW in the eastern SFZ (Braunmiller and Na`beˇlek,2002, Ristau, 2004).Seismicity lineations emanating from the SFZ trend at ∼39◦, near-parallel to the NE-47striking nodal plane of CMT solutions along the SFZ presented in this study and in Braun-miller and Na`beˇlek (2002, ∼37◦). They are also aligned with average slip vectors along theNFZ (Braunmiller and Na`beˇlek, 2002, 37◦) and the active Northern Explorer rift (Figure2.1, ∼42◦). Based on alignment of the NW-striking nodal plane with NW-trending scarpsof blocks within the SFZ, Braunmiller and Na`beˇlek (2002) reasoned that larger events rup-ture the NW-plane, but acknowledged that there must be some motion on the NE-trendingscarps to account for Pacific-Explorer motion. An alternative explanation, motivated bythe seismicity patterns in e.g., Figures 2.10, 2.15, is that larger events along the SFZ areslipping left-laterally on reactivated, NE-trending normal faults that extend into the Ex-plorer microplate interior. A magnetic anomaly pattern within the Explorer microplatethat cannot be clearly tied to spreading signature weakens the correlation with seismicity(Figure 2.15, Maus et al., 2009). Previous authors have attributed these distorted mag-netic anomalies to hydrothermal alteration and/or prolific sedimentation (e.g., Davis andRiddihough, 1982) or internal deformation (Rohr and Furlong, 1995). While not availablefor the Explorer microplate, multi-channel seismic reflection images of the adjacent JdFplate reveal steeply-dipping faults in the sedimentary layering and larger basement scarpsthat may be reactivated normal faults generated at the ridge, extending to Moho depths(Nedimovic et al., 2009). Moreover, in the Gorda plate, Chaytor et al. (2004) have inter-preted high-resolution bathymetric data along with moment tensor solutions and seismicitydata, and suggest that internal deformation of the Gorda plate has been guided by strike-slip reactivation of oceanic-spreading fabric normal faults. Oceanic spreading fabric is alsodocumented to exert control on the geometry of outer rise faulting at multiple subductionzones (e.g., Masson, 1991). This interpretation supports previous conclusions regarding theformation of the SFZ (Cowan et al., 1986) and NNF, SNF (Rohr et al., 2018) as originat-ing from the preexisting oceanic spreading fabric. Furthermore, it supports the conclusionof Barr and Chase (1974) that the SFZ no longer functions as a right-lateral transformfault. In Figure 2.15 we refer to the eastern seismically active portion of the SFZ as theEastern Sovanco Deformation Zone, following Braunmiller and Na`beˇlek (2002). Accord-ingly, a significant component of relative Explorer-JdF and Explorer-Winona basin motionis likely taken up within the Explorer microplate, as originally suggested by Riddihoughet al. (1980). In particular, we suggest that internal deformation within the Explorer mi-croplate is strongly influenced by the preexisting oceanic spreading fabric. Moreover, due tothe heavy sedimentation that obscures geomorphological features of the oceanic basementand the distorted magnetic signature, seismicity patterns provide the strongest constraintson Explorer deformation to date.Our observation of recent strike-slip activity along the PRR (Figure 2.8) suggests thatpartitioned spreading is occurring between the Explorer Deep and the Northern Explorer48rift (Riddihough et al., 1980), although there are no examples of normal faulting mechanismsalong either of the two latter structures.Together with a lack of recent volcanic activity (Michael et al., 1989), the absence ofseismicity at the southwestern end of the Explorer ridge may be the consequence of anorthward migration of spreading to the Middle and Northern Explorer rifts (Botros andJohnson, 1988, Davis and Currie, 1993, Hyndman et al., 1978). Weak volcanism at thesoutherwestern end of the Explorer ridge has also been explained as a result of along-axisflow of magma from an intermittent deep mantle source farther to the northeast along theridge (Michael et al., 1989). Alternatively, Deschamps et al. (2013) suggest that currentspreading is occurring tectonically, with waxing and waning in tectonic-volcanic cycles last-ing over 11,000-14,000 years due to a discontinuous magmatic source (possibly intermittentepisodes of magmatism from the deep mantle suggested by Michael et al. (1989)). Althoughseismicity is not observed near the southern Explorer Ridge, tectonic spreading could beoccurring aseismically. The seismic coupling coefficient (which describes the proportion ofseismic to aseismic motion) has been found by numerous authors to be inversely propor-tional to the spreading rate at oceanic ridges (e.g., Cowie et al., 1993, Mark et al., 2018).This relation may result from reduced seismogenic thickness of the crust at fast spreadingridges due to warmer temperatures (Cowie et al., 1993, Huang and Solomon, 1988, Market al., 2018).The northward shift of the Explorer ridge is enabled by ongoing, crosscutting deforma-tion along a diffuse NNW trend of seismicity between the eastern Explorer ridge and thewestern SFZ (named the southern Explorer Transform Boundary or SETB by Braunmillerand Na`beˇlek (2002), Figure 2.15), which may manifest a developing transform fault (Barrand Chase, 1974, Braunmiller and Na`beˇlek, 2002, Rohr and Furlong, 1995). In addition,predominantly strike-slip deformation in the deep (∼2900 m) and narrow (6 km) trough(Michael et al., 1989) between the Explorer ridge and the Middle rift (“RiftM” in Figure2.8) provides additional evidence for a developing transform fault. The idea of a transformfault connecting the Explorer ridge to the Juan de Fuca ridge has been suggested by anumber of authors with varying orientations. Barr and Chase (1974) originally suggestedthat the QCF may extend fully from the Northern Explorer rift to the Juan de Fuca ridge(their Figure 15). Rohr and Furlong (1995) proposed that the future evolution of a faultsimilar to that of Barr and Chase (1974) will eventually lead to complete extinction of theExplorer ridge and amalgamation of the western and eastern portions of the Explorer platewith the Pacific and North American plates, respectively. The eventual southern extent ofa composite QCF/RDF transform boundary is uncertain, but may reach as far south as theCobb offset (Dziak, 2006) (see Figure 2.15). Here we note that the southward-extrapolatedtrends of the SETB and NFZ intersect near a seismicity concentration associated with the49Endeavour ridge segment (Figure 2.15) and may point to a future triple junction.Within the NFZ, the densest concentrations of relocated seismicity follow the NNF andSNF. The seismicity patterns suggest the SNF may merge to the southwest with bathymet-ric structures east of Middle ridge, although there is a short, southerly excursion farthereast that appears to follow the trend of the local magnetic anomaly (Figure 2.15b). Accord-ing to the Savard et al. (2020) model, NFZ development has resulted in an unstable triplejunction configuration that has evolved into a landward fanning of the NFZ, consistent withobservations of normal faulting mechanisms within the NFZ (Figures 2.2,2.10,2.11; Hutchin-son et al. (2019)). We note that Savard et al. (2020) employed a rigid plate assumptionin their plate kinematic reconstruction, though seismicity within the Explorer microplate(e.g., Figures 2.10, 2.15) indicates that this assumption is not strictly valid.Seismicity at the northern end of the Juan de Fuca ridge is strongly expressed at MiddleValley and includes larger Mw ∼ 6 events with normal faulting mechanisms (Figure 2.10),whereas it is much more subdued within West Valley, as is the case in the Explorer Deepand Northern Explorer rift. The 2001 spreading event in Middle Valley (Davis et al., 2004)was the most significant in the past 20 years, but smaller seismicity bursts regularly occur(Figure A.10, Golden et al., 2003). Magmatic spreading accompanied by aseismic motionmay currently dominate at the active West Valley, whereas Middle Valley is in a tectonicspreading period (e.g., Sauter et al., 1991) as a result of a waning magmatic source.The Endeavour ridge segment to the south exhibits a dense concentration of seismicity,representing 260 events in our catalog. Previous detailed studies reveal that these eventsoften occur in swarms (Davis et al., 2001, Hooft et al., 2010), which may represent tectonicextension (Johnson et al., 2000) or dike propagation associated with magmatic intrusion(Ark et al., 2007, Bohnenstiehl et al., 2004, Hooft et al., 2010). For example, Davis et al.(2001) analyzed hydrological time series and swarm seismicity associated with the 8 June1999 spreading event along the northern Endeavour ridge segment. They interpreted thisswarm to be the result of an aseismic spreading episode with 12 cm of dilatation. Johnsonet al. (2000) also interpreted this swarm to represent primarily tectonic extension, due tothe lack of observed volcanic activity. A swarm in February 2005 that was recorded on alocal OBS array may, however, have been the response to magmatic intrusion and associateddike propagation (Hooft et al., 2010, Weekly et al., 2013). Seismicity was mainly expressedby strike-slip faulting (Hooft et al., 2010). The 1999 and 2005 swarms are thought to beassociated with a 6-year-long spreading episode at the Endeavor ridge segment (Weeklyet al., 2013). The 1999 swarm is not represented in our final catalog; the 2005 swarmincludes 30 events, compared to >6000 events relocated by Hooft et al. (2010) using localOBS.Focal mechanisms in the Middle Valley and Endeavor ridge segment (Figure 2.11) include50numerous strike-slip events in addition to normal faulting events. While the bulk shaperatio for the combined JDF ridge is very low (R=0.02), ridge segments within the JDFridge exhibit behavior that departs from pure ridge-parallel extension. For example, usingdata from an OBS deployment at the Endeavour ridge, Wilcock et al. (2002) computed focalmechanisms that exhibit north-south compression but are not consistent with ridge-parallelextension. Rather, these authors infer that seismicity on the western flank of the Endeavoursegment is affected by reorganization of the Explorer microplate.2.6 ConclusionsWe have amassed the most extensive seismicity catalog to date for the BC offshore region bycombining the existing GSC catalog with seismicity automatically detected using the RESTalgorithm over the period 1995-2021. We apply double-difference relocation to these eventsand complement microseismicity relocations with stress estimates based on moment tensorsolutions for larger events from the GSC catalog. The larger catalog and better resolvedlocations yield clearer associations with bathymetric structures and allow a number of newinferences concerning regional tectonics to be drawn.We note in particular the stark transition from focused to more diffuse seismicity alongthe RDF at its intersection with the DK. Building upon previous models for this area oftranstensional tectonics (Allan et al., 1993) and migrating transtension under prevailingPacific compression (Rohr, 2015, Rohr and Tryon, 2010), we interpret the area encom-passing diffuse seismicity and distributed volcanism extending from the northern DellwoodKnoll, TWVF and northern RDFE to collectively represent the actively-deforming com-ponent of a migrating pull-apart system. Despite an absence of a significant topographicdepression, the length-width ratio of this zone falls in good agreement with the scaling ob-served for pull-part basins across the globe at a range of scales (Aydin and Nur, 1982), andis plausibly maintained through a migrating pull-apart system mechanism (ten Brink andAvraham, 1989, Wakabayashi et al., 2004). A lack of tectonic subsidence in this area canbe explained by a weak lithospheric rheology that is plausibly associated with the nearbyKodiak-Bowie hotspot and extensive volcanism, along with balancing compression due toPacific plate obliquity. Accordingly, the Winona basin is inferred to represent the nowinactive component of this migrating system along which slip due to Pacific plate motionobliquity is partitioned between strike-slip on the seaward bounding RDF and compressionon the landward bounding paleo-QCF.In addition, we observe a general correspondence between lineations in seismicity andocean floor magnetic anomalies and/or nearby ridge segments suggesting, as in previouswork in this region (Cowan et al., 1986, Rohr et al., 2018) and elsewhere (e.g., Chaytoret al., 2004, Masson, 1991, Trehu et al., 2015), that deformation preferentially develops51along reactivated oceanic ridge fabric. In particular, seismicity within the Explorer mi-croplate region is characterized by northeast-trending lineations that emanate from theSFZ and parallel the NFZ and Explorer ridge system. The general similarity in orientationof these structures suggests that break-up of the Explorer microplate as defined by princi-pal deformation zones is controlled primarily by an anisotropic strength fabric imparted atspreading ridges. Hence, seismicity affords the best constraints on deformation where themagnetic signature and detailed seafloor morphology are obscured by heavy sedimentation.52Chapter 3Seismicity at the Intersection ofthe Coast Shear Zone and AnahimVolcanic Belt near Bella Coola,British Columbia, Canada3.1 IntroductionThe Pacific-North American margin in mainland British Columbia was formed by terraneaccretion from the Mesozoic through the Paleocene in a transpressional regime (e.g., theInsular and Intermontane superterranes (Coney et al., 1980, Gehrels and Saleeby, 1987,Jones et al., 1977, Mihalynuk et al., 1992, Monger and Price, 1979, Monger et al., 1982,Tempelman-Kluit, 1979)). By the Paleocene, major terrane accretion ended, with trans-form motion between the North American and Pacific plates ceasing thereafter during theEocene (Hollister and Andronicos, 2006). Accretion was followed by transtension due toa change in relative plate motions that resulted in crustal exhumation, thinning, and vol-canism generated from melt from within the mantle and from partial melting of the crust(Andronicos et al., 1999, Crawford et al., 1987, Hollister and Andronicos, 2006, Klepeisand Crawford, 1999). Intense metamorphism and arc magmatism associated with theseprocesses generated the Coast Mountains Batholith (CMB), the world’s largest continentalbatholith exposure (Figure 3.1; Hollister and Andronicos, 2006, Monger et al., 1982). Itsintermediate-to-felsic roots were exhumed following the end of magmatism in the regionaround 50 Ma (Hollister and Andronicos, 2006, Rusmore et al., 2005) with an estimated10-20 km of uplift and subsequent extension (Evenchick et al., 1999, Hollister, 1982). Forgreater than 1200 km along its strike, the CMB is split by a steep, NE-dipping zone of53intense ductile deformation, called the Coast shear zone (CSZ) (Figure 1; Crawford andCrawford, 1991, Crawford and Hollister, 1982, Rusmore et al., 2001).The CSZ has a complex history, characterized most recently by predominantly reversefaulting, with its last documented movement during the Eocene around 56 Ma, concurrentlywith the intrusion of sill-shaped plutons (Hollister and Andronicos, 2006, Rusmore et al.,2001). In late Cretaceous to early Tertiary time, strain partitioning in the form of right-lateral shear and compression would have dominated due to oblique convergence of thePacific/Juan de Fuca plates with North America offshore (Rusmore et al., 2001). In thelater Tertiary from about 36 to 20 Ma, volcanic and tectonic evidence indicate a subsequenttranstensional period with concurrent volcanism at the offshore Queen Charlotte Fault(QCF), the Pacific-North American plate boundary (Figure 3.1), and adjacent mainland(Hyndman and Hamilton, 1993). Significant dextral strike-slip motion and convergence inmainland BC has been inferred from geological evidence (e.g., Crawford and Hollister, 1982,Crawford et al., 1987, Gabrielse, 1985, Miller and Bowring, 1990, Price and Carmichael,1986). In particular, the Pootlass High Strain Zone (PHSZ, Figure 3.1) may have been animportant feature accommodating dextral slip (Demerse, 2008, Wang et al., 2023).In the present day, inland of the Queen Charlotte Fault, global navigation satellite sys-tem (GNSS) velocities relative to stable North America decrease but still exhibit northwardmotion of up to ∼5 mm/yr (Mazzotti et al., 2008). Mazzotti et al. (2003) and Mazzottiet al. (2008) further suggested that the deformation zone of the Pacific-North America plateboundary may include the western portion of the Canadian Cordillera. It is thought thatthe CSZ may be an important structure for accommodating this deformation (Hyndman,2015). In fact, geodetic data indicate a dextral strike-slip rate of about 2.8 mm/yr alongthe CSZ farther north in southeast Alaska (Elliott and Freymueller, 2020).Neogene-Quaternary volcanism since about 15 Ma generated an ENE-WSW trendingzone of alkaline-peralkaline volcanic centres extending from Bella Bella, BC to the BCinterior near Quesnel, called the Anahim Volcanic Belt (AVB; Figure 3.1; Bevier, 1989,Bevier et al., 1979, Souther, 1977). The volcanic centres exhibit an age progression thatbecomes younger moving from west to east, at a rate consistent with movement of the NorthAmerican plate over a fixed source (Bevier et al., 1979). A hot-spot origin has been thepreferred interpretation for the AVB (Bevier et al., 1979, Kuehn et al., 2015, Rogers, 1981).Two other contending hypotheses are shear melting of the upper mantle as an edge effectof the subducting Juan de Fuca oceanic plate (Madsen et al., 2006, Souther, 1977, Stacey,1974) and continental rifting (Bevier et al., 1979). However, the oblique and distant positionof the AVB relative to the subducting Juan de Fuca oceanic plate renders the edge-effectinterpretation problematic (Bevier et al., 1979). Moreover, body-wave seismic imaging ofa low-velocity zone extending into the upper mantle down to ∼400 km possibly implies a54−130.0°−130.0°−128.0°−128.0°−126.0°−126.0°−124.0°−124.0°52.0° 52.0°54.0° 54.0°0 100AnahimVolcanic BeltCoast Shear ZoneCNSN seismic stationVolcanic centreCoast mountain batholitheastern extentBATHOLITHS seismic stationDike swarmsSilverthroneWaddingtonNazkoRainbowIlgachuzItchaBella CoolaPrince RupertKitimatPacific PlateNorthAmericanPlateQCFNew HazeltonAnahim LakePHSZFigure 3.1: Tectonic setting including the Coast Mountain Batholith (CMB), CoastShear Zone (CSZ), Anahim Volcanic Belt (AVB), and Pootlass High StrainZone (PHSZ), from Bevier et al. (1979), Cassidy et al. (2011), Demerse (2008),Runsmore et al. (2013), Rusmore et al. (2001). Black dots indicate seismic-ity detections with at least 3 paired P and S phase picks from the combinedREST-NRCan dataset, before relocation. Triangles indicate seismic stationsused; permanent stations are in solid red and temporary BATHOLITHS sta-tions (June 2005-November 2006) are in red hatched triangles. Bella Coola isindicated with a white dot within the AVB. Map projection is Albers conicequal-area.55magma source (Mercier et al., 2009), which provides support for the hot-spot hypothesis.The town of Bella Coola, BC is located ∼100 km from the estimated southern extent ofthe CSZ and within the western portion of the AVB. It is also the location of a concentrationof seismicity immediately westward and up to ∼70 km northwest of the town with localmagnitudes from about 1 to 3, including one ML 4.2 event in 2004. This concentrationis unique in that the Coast Mountains generally exhibit low levels of seismicity elsewhere(Figure 3.1). Within the AVB, little to no seismicity is observed aside from the NazkoCone, its youngest volcanic feature and where a ∼2-month long swarm of several hundredrecorded microearthquakes occurred in 2007 (Cassidy et al., 2011). An earthquake swarmlasting about 3 months from late 1943 to early 1944 comprising about 40 earthquakes nearBella Coola may also have been associated with the AVB, however little is known aboutthis swarm. Shaking was felt by the residents of Bella Coola, however, the events were notrecorded at the then nearest seismic station in Victoria, BC, hence they are thought to havebeen small events with M < 4.5 (Milne, 1956, Rogers, 1983). The CSZ is considered to bea seismically-inactive feature, and no previous studies have documented earthquakes alongit.We aim to better characterize the volcano-tectonic context of seismically active struc-tures at the intersection of the CSZ and AVB in the Bella Coola region. To do so, we aug-ment the Natural Resources Canada (NRCan) catalog with a new microseismicity datasetgenerated using an automatic detection and location algorithm (Comte et al., 2019, Regres-sive ESTimator (REST)). We then determine precise double-difference locations, estimatefocal mechanisms for select events, and compute stress inversions to gain insight into thelocal patterns of deformation and stress. Our analysis of this distinctive cluster of seismic-ity provides further constraints on the source mechanisms and volcano-tectonic processesrelated to the CSZ and AVB, two major geological features in western British Columbia.3.2 Data & Methods3.2.1 Microseismicity dataIn our investigation of the Coast Mountains seismicity, we use continuous seismic wave-form data from the Canadian National Seismic Network (CNSN) and BATHOLITHS, atemporary broadband station array from 2005 to 2006 (Calkins et al., 2010). There areapproximately 100 CNSN stations within 300 km of the target area (Figure 3.1). However,most of these stations are located near the coast, with 4 stations in the BC interior east ofBella Coola close enough to be useful for our study. We employ an additional 5 broadbandstations that were deployed in the Nechako basin from 2006 to 2009 as part of the POLARISproject (Portable Observatories for Lithospheric Analysis and Research Investigating Seis-56micity). The BATHOLITHS project collected passive broadband data along two arraysbetween June 2005 and November 2006, each including 23 stations with 12-15 km spacing.The southern line extended through the Burke Channel past Bella Coola and farther inlandnear Anahim Lake, and the northern line extended through Douglas Channel past Kitimat,BC, following the Skeena river to New Hazelton (Figure 3.1).The catalog includes 681 events in this area with at least 3 paired P and S wave picksfrom 1995 to 2022. Because this catalog does not use the temporary BATHOLITHS sta-tions, we apply the Regressive ESTimator (REST) algorithm (Comte et al., 2019) usingboth Canadian National Seismic Network (CNSN) regional stations and the temporaryBATHOLITHS stations to improve the number of detected events in the region (51.6◦-53.5◦ latitude, 128.5◦W-126◦W longitude).The REST algorithm generates seismic event detections and estimates of earthquakeP and S phase arrivals by combining the waveform data windowing procedures of Rawlesand Thurber (2015) with autoregressive functions to detect signals and estimate arrivals(Kushnir et al., 1990, Pisarenko et al., 1987). These functions use statistical measures ofsimilarity between a user-defined window of samples and a representation of backgroundnoise to detect signals. To maximize the number of events detected, REST is designedto be overly inclusive early on in the procedure and then successively identifies and cullsfalse positives. The arrival time windows are specified based on trial hypocenters that aredetermined by grid search (Roecker et al., 2006). The hypocenters are updated as phasearrival estimates are iteratively refined until a final detection and location is determined.The REST detections in the same region provide significantly more detections duringthe BATHOLITHS recording period. However, outside of this period, the REST detectionsdo not show a significant improvement to the NRCan detection catalog. Events commonto both datasets are determined as having at least 6 phase pick arrival times within 2s,epicenters within 15 km distance, and origin times within 8 s. There are a total of 1,351unique REST detections, 433 unique NRCan detections, and 277 detections common toboth datasets for a total of 2,061 events in the region with at least 3 paired P and S picks.Events common to both catalogs are assigned the NRCan origin and any additional RESTarrival time picks are included for relocation.3.2.2 Hypocenter RelocationWe first relocate all events using Hypoinverse (Klein, 2002) and a 1-D P-velocity modelthat is provided in Figure B.1. We use Vp/Vs = 1.77, which we determine using the Wadatimethod (Vp/Vs = 1 + the best-fit slope of P-wave travel time versus S-P time (Wadati,1928); Figure B.2). All events are assigned an initial trial depth of 7 km. After cullingevents with poor convergence, 1,953 events out of the initial 2,061 events are used as input57for subsequent double-difference relocation using hypoDD.We relocate the 3-pair dataset once more using hypoDD software that implements thedouble-difference earthquake location method of Waldhauser and Ellsworth (2000), andyields relative locations between earthquakes with high accuracy. Within hypoDD, thefinal, absolute locations are obtained by adjustments of the preliminary location estimatesand are improved by including cross-correlation-derived differential times between events.We use the ph2dt program within hypoDD (Waldhauser and Ellsworth, 2000) to determinecatalog differential times for a total of 35,354 P and 32,255 S phase arrivals for earthquakepairs within a distance of 10 km. We determine waveform cross-correlations for all eventpairs within a 20 km search radius recorded at a given station, resulting in a total of22,540 P and 3,867 S cross-correlation times with cross-correlation coefficient > 0.7 for theentire initial 3-pair-pick dataset. We further constrain the cross-correlation measurementsby establishing a threshold based on the ratio between inter-event separation and averagesource-station separation for pairs of events, as a measure of the relative path differencebetween the two events. Let ∆r = ∥r2 − r1∥ and r = ∥r1 + r2∥/2, where r1 and r2 aresource-station separation vectors for events 1 and 2. We require the common path ratioC = ∆r/r to be ≤0.1.We relocate the entire dataset using the conjugate gradients (LSQR) method of min-imizing the weighted least-squares problem within hypoDD (Paige and Saunders, 1982,Waldhauser and Ellsworth, 2000). More weight is initially placed on the catalog differentialtimes relative to the cross-correlation times, and then we successively increase the weight-ing on cross-correlation differential times and downweight the catalog times throughout 4iterations. Root-mean-square (RMS) residual and linkage distance thresholds are graduallydecreased at each set of iterations. Catalog RMS residuals are reduced from 420 ms to242 ms and cross-correlation RMS residuals are reduced from 854 ms to 4 ms. The final,relocated dataset includes 837 events comprising about 60% NRCan events and 40% RESTevents.3.2.3 Magnitude determinationThe NRCan magnitudes are calculated based on the S-wave amplitude whereas the RESTalgorithm computes magnitudes based on direct P-wave amplitudes. To enable direct com-parison of event magnitudes between the two datasets, we empirically assign REST eventmagnitudes using the peak-to-peak S-wave amplitude and calibrate them to NRCan mag-nitudes. We measure the Wood-Anderson S-wave amplitude of 100 NRCan events at acommon station (BBB at Bella Bella, BC) and calculate calibration constants by multi-ple linear regression between reported NRCan ML values and the Wood-Anderson S-waveamplitude and hypocentral distance for NRCan events. Wood-Anderson amplitudes are58a) b)Figure 3.2: a) Results of multiple-linear regression of reported NRCan ML valueswith Wood-Anderson S-wave amplitude and hypocentral distance comparedagainst a perfect 1:1 fit. b) NRCan (red) and REST (blue) magnitudes overtime.determined using the Python ObsPy package. Predicted ML values are calculated as:ML = C1 ∗A+ C2 ∗D (3.1)where A is the S-wave Wood-Anderson amplitude (converted using the Wood-Anderson in-strument response), and D is the hypocentral distance. The constants are determined to beC1 = 0.3255046 and C2 = 0.0008653. Predicted ML values obtained using the calibrationconstants are plotted against NRCan-reported ML values in Figure 3.2. All points wouldplot along the 1:1 line if we achieved a perfect fit. However, errors inevitably arise due todifferences in the S-wave amplitude measurement, calibration constants, and station sub-sets employed. Our predicted magnitudes for lower-magnitude events reported by NRCantend to be slightly overestimated and magnitudes >3 tend to be slightly underestimated.Nevertheless, the general trend exhibits a satisfactory fit to the reported magnitudes.The magnitude of REST events is calculated using (3.1) with the same C1 and C2.Since not all of the REST events were recorded at BBB or their amplitudes were belowthe noise floor at BBB, amplitudes were measured at different stations for such events,but because amplitudes are converted to a Wood-Anderson response, any variability ininstrument response is corrected for. REST magnitudes contribute most to the datasetfrom 2005 to 2006, with a lower limit near ML ∼1.1 (Figure 3.2). In general, magnitudesare small for all events, between 1-3.593.2.4 Focal mechanisms & Stress inversionsWe compute focal mechanisms for events with the highest quality polarity measurementswith the HASH focal mechanism estimation software that uses P-wave first-motion polar-ities (Hardebeck and Shearer, 2002). We employ the same velocity model as that usedfor relocation (Figure B.1). The method represents uncertainty in the mechanism nodalplane solution as a suite of acceptable mechanisms that are calculated by perturbing theearthquake location and velocity model, and returns the most likely (preferred) mechanism.Due to the sparse station geometry, many events have few polarity measurements or are toosmall for a confident polarity estimate to be made at distant (>70 km) stations. Thereforewe adopt the preferred mechanism, which is the average of the acceptable solutions afteroutliers have been removed (Hardebeck and Shearer, 2002). The RMS uncertainty of themechanisms presented here is the RMS difference of the acceptable plane solutions from thepreferred plane.We then aggregate these focal mechanisms in stress inversions using the StressInversesoftware Vavryc˘uk (2014). This algorithm uses focal mechanisms to perform iterative jointinversion for principal stress directions following the method of Michael (1984), with faultplane assignment using the Mohr-Coulomb criterion. In each inversion, we set 10 stressinversion realizations with random initializations of faults, 6 iterations, and 200 randomnoise realizations for estimating the accuracy of the solution. Mean fault plane deviation isset to 30◦ to incorporate the large uncertainties in focal mechanism solutions. In additionto pressure (P) and tension (T) axes, we recover the shape ratio R, defined as:R =σ1 − σ2σ1 − σ3 , (3.2)where R describes the relative magnitudes of the principal stresses defined by σ1 ≥ σ2 ≥ σ3.As σ2 → σ1, R → 0 and the stress regime approaches uniaxial tension; conversely, asσ2 → σ3, R→ 1 and the stress regime approaches uniaxial compression.To aid in P-wave polarity estimation and assess variability of earthquake mechanismswithin clusters of seismicity, we apply the adaptive multichannel cross-correlation (MCCC)waveform alignment method of Bostock et al. (2022) for clusters of seismicity northwest ofBella Coola. This method uses minimization of a singular value–based objective function todetermine optimal lag times between waveform traces for multiple events at a given stationwithin a principal-component-based MCCC framework. We choose these clusters becausethe method requires that the source-station distance be large enough that one can assumeall events travel approximately the same path to the station. Clusters nearer to Bella Coolaare too close to the BATHOLITHS stations for the method to be useful. Using this method,we can determine the polarities of all waveforms recorded at a given station and associated60with a given cluster by assigning the polarity of the first motion of the first singular wave-form vector. Once this quantity is assigned, polarities of all contributing waveforms areautomatically determined by the sign of their singular value expansion coefficients.3.3 Results3.3.1 RelocationsRelocations reveal several quasi-linear strands of events oriented N-NW in the immediatevicinity of Bella Coola and three dense clusters about 70 km to the northwest (Figure 3.3).Most of the linear strands near Bella Coola occurred during the BATHOLITHS deploymentand were detected using REST. The clusters northwest of Bella Coola are mostly comprisedof events that occurred as swarms during August-December 2015 and 2017, with a minorburst of activity in September 2022. We quantitatively group the relocated events intoclusters using k-means clustering (the generalized Lloyd’s algorithm via the Python Scipyimplementation; Figure 3.3c-d), for ease of reference and comparison. There are a total of10 clusters that we report on in this paper, each assigned an arbitrary number label: 0, 1,3, 4, 5, 7, 11, 14, 15, and 17. Other clusters with sparse and few events (i.e., clusters 6, 9,and 10) are not considered in this section.Our dataset includes 306 relocated events recorded during the BATHOLITHS deploy-ment. The median depth of the events closest to Bella Coola and the BATHOLITHS arrayis about 10 km, with a median Hypoinverse depth error of about 2.75 km and median hori-zontal error of about 2.48 km. A moment tensor already exists for one of these events, aML4.2 event on 6 October 2004 (Figure 3.3b; Ristau et al., 2007). The strands associated withclusters 1, 4, 5, 7, 11, and 15 are on the order of 10-15 km in length; they do not appearto correspond with any obvious topographic features (Figure 3.3; Figure B.3). There areinsufficient detections to confidently discern detailed spatio-temporal patterns within thestrands. The strands generally appear to be slightly oblique (clockwise) to and locatedeast/northeast of the mapped trends of the CSZ (Figure 3.3; Rusmore et al., 2001). Weobserve no seismicity associated with the PHSZ (Figure 3.3). No large earthquakes wererecorded in this area during the BATHOLITHS deployment and estimated magnitudes fallbetween 0.9 and 2.7 in that time frame.About 70 km northwest of Bella Coola, there are three distinct clusters of seismicity(clusters 14, 17, 0; Figures 3.3c-d and 3.4). Reported depths are slightly deeper than theevents nearest to the BATHOLITHS stations in 2005-2006 (10-20 km), but are more poorlyconstrained here, where the nearest station is either BUTB (in operation from 2017-2020)or BBB, both ∼70 km away. The median Hypoinverse depth error of all events in theseclusters is 5.75 km before input to hypoDD. Station BBB is the the only station that61−128.0° −127.5° −127.0° −126.5°52.0°52.5°53.0°BC0 12.5 25 199619992002200520082011201420172020Yearkm2004-10-04 ML 4.2 eventCSZPHSZ−128.0° −127.5° −127.0° −126.5°52.0°52.5°53.0°BC0 12.5 25010203040Depth (km)CSZPHSZkma) b)c) d)−128.5° −128.0° −127.5° −127.0° −126.5°52.0°52.5°53.0°013456791011141517BC0 12.5 25−128.5° −128.0° −127.5° −127.0° −126.5°52.0°52.5°53.0°013456791011141517BC0 12.5 25 020406080100120140Number of events in clusterkm km−128.5° −128.0° −127.5° −127.0° −126.5°52.0°52.5°0 3456791011141517BC0 12.5 25−128.5° −128.0° −127.5° −127.0° −126.5°52.0°52.5°53.0°013456791011141517BC0 12.5 25 020406080100120140Number of events in clusterkm kmFigure 3northwest clusterslinearstrandsBUTBBBBBCBCBUTBBBBBCBClinearstrandsFigure 3.3: All relocations colored by a) depth and b) time, with size scaled to mag-nitude in b). The 2004 ML 4.2 event is indicated with its focal mechanism(Ristau et al., 2007). c) and d) display K-means clusters of events, colored bycluster assignment and number of events per cluster, respectively.operated continuously throughout the entire study period. Station BCBC in Bella Coolawas in operation between 2007-2014, however it is still just under 70 km away from thenorthwest group of clusters. We consider the majority of these events to have occurredwithin the upper crust, but address the possibility of surficial events, such as landslides, ina later section.Nearly all of the northwest clustered events occurred as swarms during August-Decemberin both 2015 and 2017. These clusters (clusters 0, 14, 17) appear to delineate NE-trendingfeatures that are about 6 km long and highly oblique to the general NW trend of the CSZ.They also reside in the location of an Eocene-age pluton that disrupts the mapped trace of62a)b)0 km 10 km5835583958435847585158555860578 582 587 592 596 601 606 08-1009-0810-0611-0312-0112-29UTM Y (km)UTM X (km)Date (MM-DD)ML 3.31a)b)5835583958435847585158555860578 582 587 592 596 601UTM Y (km)0 km 10 kmUTM X (km)606 08-0812-2309-0510-0210-2911-26Date (MM-DD)ML 3.4a)b)2015 2017c) d)e) f) g)Figure 3.4: a) Vertical-component displacement waveforms of the largest events ofthe 2015 (black, ML 3.31) and 2017 (red, ML 3.40) events recorded at stationBBB in Bella Bella, BC, 70 km WSW. Waveforms are highpass-filtered witha corner of 1 Hz. Time is in seconds from the origin. b) Same as a) but withthe 2017 event (red) flipped and shifted by the lag time corresponding to themaximum in the cross-correlation function between the two waveforms. Thisaligns the events on the S- and surface waves and clearly demonstrates thecontrasting amplitudes, polarities, and S-P times of the two events. Reversedpolarity can be seen in the initial P-wave. c) Map showing cluster comprisingthe 2015 sequence, with colors scaled by time and size are scaled by magnitude.d) Same as for c) but for the 2017 sequence. e) Magnitudes over the duration ofthe swarms in 2015 and 2017, f) Magnitude-frequency plot of the 2015 swarm.Best fit b-value=1.30, g) Magnitude-frequency plot of the 2017 swarm. Best fitb-value=1.40.63the CSZ (Rusmore et al., 2001). Both swarms included ∼50 events during the first 2 monthsof the swarm, followed by a brief pause during November, and then resumed in December(Figure 3.4e-g). The 2015 swarm was localized to a single cluster (cluster 14) whereas the2017 swarm moved from cluster 14 to cluster 17 (August-October), then finally to cluster0 in December (Figures 3.4c-d, B.4b-c). The largest event for the 2017 swarm (ML 3.4) isdetermined to be 16.35 km deep by hypoDD, however, the amplitude of the surface wavesuggests it may be shallower (Figure 3.4). The largest event for the 2015 swarm (ML 3.31)occurred in nearly the same location but slightly to the north of the 2017 event, and itsP-waveform is more emergent in character (Figure 3.4). Their waveforms at a commonstation, BBB (Figure 3.4), exhibit opposing P polarities and different relative excitations ofS- and surface waves, indicating that these two events clearly had different mechanisms. Yet,there is sufficient similarity in the coherence of their P, S and surface waveforms to confirmthat they occurred in close proximity. Neither swarm exhibits a characteristic mainshock-aftershock pattern of magnitude decay; rather, the magnitudes appear to slightly increasewith time (Figure 3.4e). The best-fit of the magnitude-frequency distribution yields a b-value of 1.30 for the 2015 swarm and 1.40 for the 2017 swarm (Figure 3.4f-g).3.3.2 Source Mechanisms & Stress InversionsHere we present the preferred focal mechanism solutions determined using HASH. Sourcemechanisms frequently exhibit high uncertainties owing to the generally sparse and unevenstation distribution and the low signal quality for many P phase arrivals. We tested oursolutions using two alternate velocity models, one biased fast and one biased slow from theoriginal model in Figure B.1. The deviations in final solutions are minor. RMS uncertaintieson nodal planes are between 10◦-40◦ (Figures 3.5; 3.6). Nonetheless, in aggregate, the 84total mechanisms tend to show systematic patterns. Principal stress directions and shaperatios are displayed in Figure 3.7 and corresponding trend and plunge values are in TableB.1.We observe a general predominance of strike-slip and normal faulting events both withinthe linear strands near Bella Coola (Figure 3.5) and within the clusters to the northwest(Figure 3.6). In the northwest clusters 0, 17, and 14, strike-slip faulting appears to beslightly more prevalent, with NE-trending nodal planes generally aligned with the northeast-trending clusters. Clusters 0 and 17 are more suggestive of right-lateral strike-slip faulting,whereas cluster 14 exhibits a variety of differing orientations (Figure 3.6). This is alsoreflected by a near-vertical σ2 (Figure 3.7a). Minimum compressive stress (σ3) axes aredirected northeast.Waveform alignment of events within the northwest clusters shows opposing P-wavepolarities of events within the same cluster, further indicating variability in source mecha-64−127.5° −127.4° −127.2° −127.1° −126.9° −126.7° −126.6°51.9°52.0°52.2°52.3°52.5°52.6°14150 12.5 2510152025303540451175BCBCBCTLCBBBBUBRBLLLBNLLBFSBGDRHOPBCRAGPTBS09BS06BS10BS05BS04MBS03BS15BS15ABS17BS19VI32BBBPTBS06BS05BS07BS04MBS03BS12MBS13BS15BS15ABS17BBBBS19BS21BS22BN01VI10PTBS06BS03BS01BBBBS19BN21BS05BS04MPTBBBALRBTALBTHMBRUBBMASBHOLBMAYBPTBS06BS05BS04MBS03BS17BBBBS19PTBS04MBS06BS03BS02BS01BBBBS197MAYBPTID:55180ID:54871ID:51373ID:24561 ID:47046 ID:00327 ID:43203ID:13825ID:44440ID:07824RMS (deg)BS06BS04MBS03BS02BS01BS10BBB BS05BS19HOLB PTBBBBUTBMBLBHWKBKITBUBRBFSJBPHCHOLBPTBS05BS04MBS03BS06BS07BS10BBBBS21PTFigure 3.5: Estimated focal mechanisms for events in the Bella Coola strands (BCstrands in Figure 6) clusters, with color scaled to their RMS uncertainty esti-mated from HASH, in degrees. Examples of detailed focal mechanisms borderthe map and include their event ID. This figure includes the mechanism of the2004 ML 4.2 event from Ristau et al. (2007) in black.65−128.0° −127.8° −127.7° −127.5° −127.4° −127.2° −127.1°52.6°52.8° 01114170 12.5 2515202530354045BS05BS06BS04MBS03BS10BBBBN04BN11BN23VI02VI05VI06VI089VI10MAYBBNABPTBBBGRIBKITBFSJBHAKBRUBBPTID:259346RMS (deg)ID:20893BS03BS05BS06BBBBS14BN04BN07BS17BS22BN112BN13BN14BN17BN20BN2223TXBPHCWOSBPTPTBUTBBBBHWKBKITBBNABFSJBPHCPTBUTBBBBBNABGRIBGRNB FSJBPHCWOSBNCRBPTBBBPHCTLCBBUTBKITBHWKBUBRBBNKBGRNBFSBGDRPACBPTID:09626ID:18178ID:14486 ID:08983 ID:09102 ID:12275BUTBBBBGRIBHWKBKITBGRNBBNKBBNABPHCPTBUTBBBBGRIBKITBBNKBRUBBHWKBPACBPTID:17017BBBHWKBKITBBNKBFSJBBPTFigure 3.6: Estimated focal mechanisms for events in clusters northwest of BellaCoola (NW Clusters in Figure 6) with color scaled to their RMS uncertaintyestimated from HASH, in degrees. Examples of detailed focal mechanisms bor-der the map and include their event ID. This figure includes the mechanism ofthe 2004 ML 4.2 event from Ristau et al. (2007) in black.nisms. We display results for waveform alignment of events recorded at stations BBB andBUTB for cluster 14 (Figure 3.8). This figure reveals good alignment of the P-wave, visibleon the vertical component BHZ between 8.3-8.8 s (second column), and the first principalcomponent vector displays a clear P arrival (third column). There does not appear to be anysignificant spatial pattern in the events with positive/negative polarities (fourth column).The strands of seismicity near Bella Coola (clusters 11, 15, 4, 1, 7, and 5) generallyexhibit slightly more normal than strike-slip faulting (Figure 3.5), although mechanismsfrom cluster 7 exhibit mostly strike-slip faulting. Given the correspondence between NNW-striking nodal planes and the trend of seismicity, we expect these to reflect right-lateralstrike-slip faulting. The largest of these events (ID: 00327, ML 3.45) is one of the better-constrained events, with a 22◦ RMS uncertainty. To the south, cluster 5 exhibits morenormal faulting. Clusters 1, 4, 11, and 15 are more variable, exhibiting both strike-slip andnormal faulting. These clusters of mechanisms yield a near-vertical σ1, in contrast to theclusters to the northwest, where σ1 and σ2 are more similar in magnitude and form a girdle,or continuous circle (Figure 3.7). Both suites of mechanisms exhibit NE-directed minimumcompressive stress σ3 (Figure 3.7). Shape ratios are somewhat poorly constrained but aregenerally <0.5, which suggests extension.66a) NW Clustersb) BCStrandsc) AllR=0.34R=0.36R=0.27(0, 17, 14)(1, 4, 5, 7, 11, 15)Figure 3.7: Stress inversion results for a) events from clusters 0, 14, and 17, b) eventsfrom clusters 1, 4, 5, 7, 11, and 15, and c) all events. Blue crosses and redcircles are Tension and Pressure axes, respectively. All inversions indicate aNE-directed σ3. The northwest clusters indicate that σ1 and σ2 define a girdle.The Bella Coola strands indicate a near-vertical σ1.3.4 Discussion3.4.1 Consideration of surficial seismic eventsWe consider the majority of the events in the Bella Coola area to be tectonic and not theexpression of surficial processes, such as landslides, avalanches and/or alpine glacier de-formation, for several reasons. First, this area of seismicity is unique within the broaderCoast Mountains region of British Columbia. The region crossed by the northern array ofBATHOLITHS stations extending northeast from Prince Rupert exhibits similar physiog-raphy but not nearly the same level of seismicity (Figure 3.1). Therefore, if the majorityof the events near Bella Coola were related to surficial processes, we might expect similarlevels of seismicity near the northern BATHOLITHS line. Secondly, the nominal depths of∼5-20 km of events closest to the BATHOLITHS stations, along with general similarity of67−128.6°−128.4°−128.3°−128.1°−128.0°−127.8°−127.7°−127.5°−127.4°52.2°52.3°52.5°52.6°52.8°52.9°53.1°BBB0 12.5 25−101−128.6°−128.4°−128.3°−128.1°−128.0°−127.8°−127.7°−127.5°−127.4°52.2°52.3°52.5°52.6°52.8°52.9°53.1° BUTB0 12.5 25−101PolarityFirst principal component vector (BBB, BHZ)First principal component vector (BUTB, BHZ)PolarityMCCC Adaptive alignment (BBB)MCCC Adaptive alignment (BUTB)a) BBBb) BUTBBHE BHN BHZBHE BHN BHZInset waveform section (BBB, BHZ)Inset waveform section (BUTB, BHZ)Figure 3.8: Three-component waveform alignment results for cluster 14 using anadaptive alignment procedure (Bostock et al., 2022). First row a) shows re-sults for station BBB and second row b) shows results for station BUTB. Firstcolumn shows three-component waveform alignment, with the P-wave arrivalclearly visible on the BHZ channel but with opposing polarities for subgroupsof events within the cluster. Second column shows a zoomed-in section of thealigned P-arrival on the BHZ component, which clearly displays opposing po-larities. Third column shows the first principal vector describing the systemof waveforms. The green line is the expected P-arrival time. Fourth columnshows a map of the events within the cluster recorded at the respective station,colored by polarity (positive in red, negative in blue).waveform character of the northwest clusters with the BATHOLITHS seismicity, suggeststhat all of these events reflect similar source processes. Although there is variability in thewaveforms, they resemble those of earthquakes with identifiable phase arrivals (e.g., FigureB.5). Landslides, on the other hand, usually produce an emergent, “cigar shaped” waveformas higher-frequency energy builds up gradually, and with unclear phase arrivals (Allstadt,2013, Dammeier et al., 2011). We note that the waveform of the main 2015 event (ID: 12275,Figure 3.4) is more emergent in character at station BBB. Its estimated focal mechanismindicates strike-slip faulting, albeit with an RMS uncertainty of 25◦. Our estimated focalmechanism for the largest event in the 2017 swarm (ID: 08983; Figure 3.6), suggests normalfaulting, but is only based on 6 polarity measurements. AML 2.67 event the day before (ID:09102, Figure 3.6) also yields an extensional mechanism but in an approximately orthogo-nal orientation. Both the 2015 and 2017 swarms occur in distinct periods over a ∼3-monthlong time period, longer than would be expected for a singular landslide event (e.g., the2020 Bute inlet, British Columbia slide (Geertsema et al., 2022)). Finally, we compare the68daily seismic detections over time during the 2017 swarm with temperature and precipita-tion data from the Government of Canada Bella Coola airport weather station (Figure 3.9.Although this station is over 70 km away from these events, it provides the best estimate ofweather conditions at the time. We do not observe any obvious correlation between spikesin temperature/precipitation and spikes in seismicity. Event detections for the entire studyarea outside of the swarm periods do not exhibit any clear seasonality, although there aredecreased detections between March-June (Figure B.6), which likely reflects the influenceof the BATHOLITHS deployment from June 2005- November 2006. Unfortunately, thereis insufficient satellite data of the area to observe any changes in morphology in the moun-tains directly above the relocations before and after the swarms. We therefore interpret themajority of these events as related to tectonic/volcanic deformation, but do not exclude thepossibility that there are some surficial events contained within our dataset.3.4.2 Seismicity at the intersection of the CSZ and AVBWe now consider the hypothesis that the occurrence of seismicity near Bella Coola is relatedto the intersection of the CSZ with the AVB. Seismicity is nearly absent elsewhere alongthe CSZ, including where it intersects the northern BATHOLITHS line, although it isestimated to slip about 2.8 mm/yr in the northern portion above latitude 54◦ (Figure 3.1;Elliott and Freymueller, 2020)). Clusters of seismicity (e.g., clusters 1, 4, 5, 7, 11, 15) areoriented NW, but oblique to the mapped CSZ. As the most recently documented faultingon the CSZ was reverse motion (Rusmore et al., 2001), the extensional and strike-slip eventsmay be occurring on reactivated reverse fault structures. We note that the regional stressregime of western Canada is one of NE-directed compression from residual Pacific plateconvergence (Mazzotti et al., 2003, Ristau et al., 2007). Moreover, the mapped CSZ inthis area is significantly wider than other areas along its strike (Figure 3.1; Rusmore et al.,2001). Thus, local extension in this area may plausibly be related to the AVB.In addition to the spatiotemporal trend in volcanism, the area is characterized by highheat flow values (>80 mWm−2, Majorowicz and Grasby (2010)) and a relatively high inci-dence of hot springs (Figure 3.10; Woodsworth and Woodsworth, 2014). We do not observeany obvious spatial correlation between hot spring and earthquake locations (Figure 3.10).Yet, hot springs are commonly observed in fault systems (e.g., Finley et al., 2022); in par-ticular, strike-slip faults are especially conducive to hot spring formation as meteoric watertends to follow σ2 and therefore vertically upwards (Sibson, 1996). Active faulting in acomplex network of small-scale fault strands along the CSZ may facilitate the generation ofpermeable pathways for fluids to the surface (Curewitz and Karson, 1997, Faulds and Hinz,2015), and therefore may be responsible for the several hot springs present in the area.The northwest clusters of seismicity (clusters 0, 14, and 17, Figure 3.3) are situated at69Figure 3.9: Comparison of temperature and precipitation from the Government ofCanada weather station at Bella Coola Airport (Climate ID: 1060844) duringthe swarm (top), compared with the daily event detection count (bottom).70−128.0° −127.5° −127.0° −126.5°52.0°52.5°53.0°0 12.5 25CSZ PHSZSeismic stationHot springBCFigure 3.10: Map of relocations (black circles) with hot springs (magenta stars) fromWoodsworth and Woodsworth (2014). “BC” stands for Bella Coola.71the locus of an abrupt change in the thickness of the mapped CSZ trace (Figures 3.1 and3.4; Rusmore et al., 2001). At this location, the CSZ is intruded by a late Eocene-age pluton(Rusmore et al., 2001). The intrusion may reflect a structural weakness or, alternatively,the change in fault morphology may give rise to elevated local stresses, either of which playa role in the origin of these clusters. We note that the elevated b-values for the 2015 and2017 swarms (1.3, 1.4, respectively) are not as high as the Nazko cone swarm (1.9), whichlasted for about 2 months and is believed to be a result of magma injection into the lowercrust (Cassidy et al., 2011). The largest event recorded during the 2007 Nazko swarm wasdetermined to be a normal-faulting event (Cassidy et al., 2011), similar to our observationsof normal faulting for the largest event in 2017 (ID:08983, in Figure 3.5).Because the AVB is proposed to be the result of a hot spot or mantle plume track (e.g.,Bevier et al., 1979), we compare results here with the more extensively studied Yellow-stone hot spot in the United States. We note the following similarities: swarm behaviour,normal/strike-slip mechanisms indicating low shape ratios, and comparable b-values. Asobserved here, the 1985 Yellowstone swarm lasted for about 3 months (Waite and Smith,2002). Moreover, the 1985 and 2010 earthquake swarms at Yellowstone exhibited predomi-nantly strike-slip faulting with some normal faulting (Russo et al., 2020, Shelley et al., 2013,Waite and Smith, 2002), and are believed to reflect hydrothermal fluid migration. The lowshape ratios documented here (R ∼0.3, Figure 3.7) are similar to those measured for the1985 Yellowstone swarm (Waite and Smith, 2002), and our estimated b-values of 1.3 and1.4 are comparable to those estimated for swarms within the Yellowstone caldera, whichare between 1.3 and 1.5 (Farrell et al., 2009).In the geomechanical model of Russo et al. (2020), brittle deformation may result fromoverpressurized fluid flow in response to magma diking. In this model, normal faultingoccurs on planes emanating upwards from where a vertical dike propagates, and strike-slipfaulting occurs on faults oriented obliquely to a horizontally-propagating dike (their Figure11). Therefore, normal and strike-slip faulting in these clusters could reflect brittle failureresulting from hydrothermal fluid flow in response to vertical or horizontal dike propagation,respectively. However, denser instrumentation in the Coast Mountain region is required tobetter distinguish source processes.One notable difference between the seismicity observed here and that associated withthe Yellowstone swarms is that it occurs along an older portion of the inferred hotspottrack. Situated between the Gale Passage/Bella Bella dike swarms (Souther, 1986) and theRainbow range (Figure 3.1); Bevier, 1978), this portion may be between 14 and 9 Ma. AtYellowstone, much of the swarm activity is confined to immediately beneath or adjacentto the caldera overlying the main magma chamber (e.g., Miller and Smith, 1999, Russoet al., 2020)). This observation suggests that the seismicity documented here is due to72enduring interaction between elevated mantle temperatures associated with the passage ofan AVB mantle plume and lithospheric weakness at the CSZ developed in an earlier era ofdeformation.3.5 ConclusionA section of the CSZ, approximately 100 km in length, and centered near the town ofBella Coola, BC exhibits increased levels of microseismicity relative to adjacent portionsof the Coast Mountains. We augment the existing NRCan catalog using an automaticdetection and location algorithm (REST) for the region immediately surrounding the townof Bella Coola over the period from 1995 to 2022. The new dataset provides a significantimprovement in seismic detections due in large part to the 2005-2006 BATHOLITHS passivearray deployment. We apply double-difference relocation to these events and documenttheir spatio-temporal behavior, estimate focal mechanisms and compute stress inversionsto provide new insights into the roles of tectonics and volcanism. This study provides thefirst detailed analysis of earthquakes in the Bella Coola region and motivates future detailedstudy.While hypocenters and focal mechanisms are characterized by relatively large uncertain-ties due to the sparse, uneven, and temporally variable station distribution, we are able todefine previously unrecognized patterns and identify two significant swarm events in 2015and 2017. Double-difference relocation reveals small-scale (∼10 km) linear strands orientedNNW and slightly oblique to the mapped CSZ fault trace and three distinct NE-strikingclusters to the NW of Bella Coola that occurred as swarms in 2015 and 2017. Focal mecha-nisms are variable but indicate a predominance of extensional and strike-slip faulting, and alocal stress regime dominated by NE-directed extension. The latter contrasts with the NE-directed compressional stress and observations of right-lateral transpressional earthquakemechanisms documented at more regional scales (e.g., Elliott and Freymueller, 2020, Ristauet al., 2007). Moreover, the character of the swarm seismicity reported here shares a numberof similarities with that documented for the Yellowstone hotspot to the south (Farrell et al.,2009, Russo et al., 2020, Shelley et al., 2013, Waite and Smith, 2002). We hypothesize thatthis concentration of seismicity within the CMB manifests reactivation of a weakened CSZthrough interaction with elevated mantle temperatures associated with the passage of themantle plume hypothesized to be responsible for the AVB (Bevier et al., 1979, Kuehn et al.,2015).73Chapter 4Tectonic tremor: the chatter ofmafic underplating beneathsouthern Vancouver Island?4.1 IntroductionSince the discovery of tectonic tremor in Cascadia (Rogers and Dragert, 2003), significanteffort has been devoted to understanding its relationship to episodic slow fault slip (i.e.,episodic tremor and slip, ETS). Tectonic tremor (or simply, “tremor”) is a low-amplitude,seismic signal bandlimited between ∼1-10 Hz (Katsumata and Kamaya, 2003, Obara, 2002)that usually temporally coincides with slow slip in Cascadia. While it occurs in othersubduction zones and strike-slip faults worldwide (e.g., Brown et al., 2009, Nadeau andDolenc, 2005, Schwartz and Rokosky, 2007), it is best documented in warm subduction zonessuch as Cascadia and Nankai (southwest Japan). Tremor is widely regarded to comprise asuperposition of individual low-frequency earthquakes (LFEs) and is commonly used to inferdetailed migration patterns of slow slip (e.g., Ghosh et al., 2010, Rubin and Armbruster,2013).The source mechanism responsible for LFEs has been widely debated. Contending hy-potheses include: (1) shear slip in the plate boundary zone, with focal mechanisms consistentwith thrust faulting (Bartlow et al., 2011, LaRocca et al., 2009, Royer and Bostock, 2014,Shelly et al., 2005), (2) slip along multiple surfaces that are distributed across ∼40 km indepth (Kao et al., 2009), (3) rapid fluid transients or pore pressure waves (e.g., Cruz-Atienzaet al., 2018, Shapiro et al., 2018), or (4) local shear instabilities in a granular channel (e.g.,Sammis and Bostock, 2021). Nonetheless, the sensitivity of tremor and LFE activity toEarth tides, and the presence of a zone of elevated Vp/Vs and depressed shear-wave velocity74in the tremor source region suggest that fluids play a significant role by lowering the ef-fective stress (Audet et al., 2009, Bostock et al., 2012, Hawthorne and Rubin, 2010, Royeret al., 2015, Thomas et al., 2009).A primary challenge in ascertaining the source process of LFEs lies in their locationin depth. Depth is difficult to determine accurately because of the low amplitudes of P-waves arising from their generation at a shallowly dipping or steeply dipping structures(in subduction zones or transform faults, respectively). Radiation patterns at subductionzones favor the observation of S-waves at nearby stations (e.g., Ide et al., 2007). Therefore,most previous studies (e.g., Bombardier et al., 2023, Ghosh et al., 2012, Kao et al., 2009,Obara, 2002, Rubin and Armbruster, 2013, Wech and Creager, 2008) used only S-waves todetermine locations, frequently assuming a slab model to fix locations in depth.Several authors have noted associations between the occurrence of tremor and under-plating in warm subduction zones (Bassett and Watts, 2015, Behr and Bu¨rgmann, 2021,Kimura et al., 2010). However, a direct link between tremor and underplating has yet tobe confirmed, largely due to challenges in obtaining precise locations necessary to iden-tify signatures of underplating. Southern Vancouver Island provides an exceptional settingin which to address this problem because minimal crustal scattering yields comparativelyclean seismic waveforms dominated by direct arrivals (e.g., Figure 4.1; Bostock et al.,2015, Royer and Bostock, 2014) (c.f., Plourde et al., 2015). Moreover, stations from thetemporary POLARIS array (2003-2005; Nicholson et al. (2005); Figure 4.2) are situatednear tremor sources and are sufficiently closely spaced to yield high cross-station correla-tions over short (4 s) windows for tremor arising in three major ETS episodes (Armbrusteret al., 2014, Rubin and Armbruster, 2013, Savard and Bostock, 2015). We focus on threeLFE template locations near the axis of the POLARIS array, 053, 065, and 070 (as definedby Bostock et al. (2015); see Figure 1), which have not previously been analyzed in de-tail. Combined with detailed seismic reflection and tomographic imaging, accurate tremorhypocenters allow us to address the relationship of tectonic tremor to warm subduction andunderplating.4.2 Methods4.2.1 Detection and LocationWe employ the P- and S-waveforms for LFE templates 053, 065, and 070 at 5 stationsto determine a) delays for S- and P-wave arrivals corresponding to the nominal templatelocation, computed using the alignment procedures in Bostock et al. (2022); b) the splittingparameters that best reduce the S-wave particle motions on the two horizontal coordinatesto rectilinear motion isolated to a single channel (Peng et al., 2015); and c) the expected P-750 1 2 3 4 5 6Time [s]PGC-SSILB-SSSIB-SKELB-SSSIB-PSNB-P0.0 1.0 2.0 3.0 4.0 5.0 6.0Time [s]1.000.750.500.250.000.250.500.751.00 PGC-SSILB-SSSIB-SKELB-SSSIB-PSNB-Pa)b)Figure 4.1: a) S and P waveforms for template 053 used to locate 4 s tremor win-dows. S-waves have been corrected for splitting and P-waves for polarity. Noteprominence and similarity of direct arrival across channels. b) Same waveformssuperposed and aligned according to optimal delays Only stations PGC, KELB,SSIB, and SILB are used for S-wave detection and only SSIB or SNB are usedfor P-wave detection.76-128.0˚-128.0˚-126.0˚-126.0˚-124.0˚-124.0˚-122.0˚-122.0˚42.0˚ 42.0˚44.0˚ 44.0˚46.0˚ 46.0˚48.0˚ 48.0˚50.0˚ 50.0˚VancouverSeattlePortlandEugene48.0°49.0°-124.0° -123.5° -123.0°48.5°KLNBPGCSILBSNBSSIB~ 41 mm/yrJuan de Fuca plate Cascadia subduction zoneBlanco Fracture ZoneGorda RidgeNFZNorth American plateSan JuanFault−123.5° −123.0°kmVancouverIsland84-020 12.5 25Figure 4.2: Map of the Cascadia region with the purple line denoting the outline oftremor epicenters from 2008-2019 (Wech, 2010) and major tectonic boundaries(black lines) and Cascade volcanoes (triangles) indicated. Map inset shows thestudy region. Large blue circles show the three LFE templates used in thisstudy; smaller blue circles are other LFE template locations (Savard et al.,2018). Red triangles are the five recording stations used and black line is theseismic reflection profile 84-02 (Clowes et al., 1987). The San Juan Fault, aforearc block boundary fault, bisects the study region. NFZ = Nootka FaultZone.77polarity for P-waveforms on the vertical component. These quantities are used to normalize24 hour-long waveforms for the 4 S-wave stations (KELB (2003)/KLNB (2004 and 2005),PGC, SILB, SSIB) and 1 P-wave station (SNB) employed in this study (stations KELBand KLNB differ in location by ∼40 m). PGC is designated as the time-stamp referencestation (0 s delay) with template-dependent relative delays applied to the remaining 4channels. Moreover, a single suite of station-specific splitting parameters is employed forall 3 templates to maintain the same S-waveforms (with different relative delays) for eachcase. Following Rubin and Armbruster, the normalized 24-hour waveforms are divided into86396 4 s windows with 3 s overlap starting at midnight (PGC time) and cross-correlationsare computed for all (4) pairs of S-stations with a common time stamp to maximum lags of+/− 0.4 s for a given template to mitigate against cycle skips (Figure 4.1). Thus, a tremorburst originating at the nominal template epicenter would register maximum correlation at0 s lag for each pair, whereas non-zero lags characterize epicenters away from the templateepicenter. The chosen +/− 0.4 s lag allows some overlap across the 3 template epicentralregions. A prospective detection is declared if 2 conditions are met relating to thresholds onthe values of a) the 4 possible 3-station delay time circuits (i.e. |tij + tjk − tik| for stationsi, j, and k), and the mean cross correlation coefficient (see Appendix C.1; Rubin andArmbruster, 2013). In the event of a prospective detection, the first principal componentwaveform of the aligned S-waveforms is cross-correlated with the single-station P-waveformand a detection is declared upon meeting a second correlation coefficient threshold. Thelag at maximum correlation enables computation of an S-P time (and therefore hypocentraldepth) and lags are again restricted to lie between +/− 0.4 s.The computations above are performed for two different frequency bands: a narrowband of 1.5-6 Hz like that employed by Rubin and Armbruster (2013), and a broaderband of 1-8 Hz which yields fewer detections, but reduced likelihood of cycle skips dueto the wider range of frequencies represented. There is therefore the possibility for upto 6 duplicate detections per time stamp (2 frequency bands for each of 3 templates).We employ several thresholds and statistics (described in detail in Appendix C.1) to cullthe tremor catalogue to a maximum of 1 detection per time stamp to emphasize tremorhypocentral patterns but minimize scatter arising from false detections. We do not attemptto eliminate repeated arrivals across consecutive overlapping waveforms, thereby allowingfor a continuously evolving tremor wavefield. Each detection is thus characterized by 4 Stimes and 1 P time allowing for location with only 1 degree of redundancy (5 constraintson 4 hypocentral parameters). This approach yields high relative precision in location(dependent upon the precision of the relative delays), but an absolute accuracy that isdependent upon the validity of the velocity model given that stations are dominantly toone side of the source region. The choice of a suitable P-wave station is a compromise in78SNR between epicentral distance and relative P-to-S radiation (assessable from templatewaveforms) and is best met for this region by station SNB (Armbruster et al., 2014).We obtain initial locations using Hypoinverse (Klein, 2002) with a 1D velocity modelbased on the Savard et al. (2018) 3D model at this location. After culling detections (de-scribed in Appendix C.1), we determine final hypocenters using double-difference relocationin tomoDD (Zhang and Thurber, 2003) with the 3D velocity model of Savard et al. (2018).We also compare the resulting relocations using SSIB as the P-wave detection station in-stead of SNB. The relocations preserve a very similar relative pattern, although due toinaccuracy of the velocity model, absolute locations are shifted slightly deeper by about 1.5km. However, relocation of the same template waveforms as Savard et al. (2018) using themethod presented here shows good agreement with the Savard et al. (2018) study. Further-more, our final tremor catalog locations show good agreement with the Savard et al. (2018)LFE family template locations, indicating that our method yields robust absolute locations(Figure C.6). The final catalog represents 4,851 locations determined for a total of 13 daysspread over 3 ETS episodes, as determined from the LFE catalogs of Bostock et al. (2015).4.2.2 False Detections and Location UncertaintyAll detections scattered away from the main patches were visually inspected for cycle-skipping in the waveform alignment or false detections. Only about 15% of the scattereddetections were deemed true detections, and most of these events occurred shallower than37 km depth. We do not observe any significant tremor patches distributed in depth oranywhere significantly above or below about 39 km depth. We therefore cull the datasetfurther to include only the events between 37 and 39.5 km depth. Visual inspection of manyof the waveforms contributing to the detections composing the primary layer structuresindicates that most of these detections do not suffer from cycle-skipping.We estimate the location uncertainty by comparing the locations of common timestampdetections between the narrow & broad band data in 3 different templates (each gives adifferent location measurement, resulting in 6 different locations to compare). Epicentraluncertainty is estimated to be about 250 m, and depth uncertainty about 300 m (FiguresC.2, C.7). We note that the nominal horizontal and depth errors from Hypoinverse are 1.15km and 2.3 km, respectively.4.2.3 Moment and Moment Magnitude CalculationsFollowing Rubin and Armbruster (2013), we assign a consistent, coherent radiated energymetric for each 4 s, narrow-band S-detection window as79E(t) =(SA(t)SB(t′B) + SA(t)SC(t′C) + SB(t′B)SC(t′C))3(4.1)where S(t) is the S-wave seismogram, A, B, and C denote stations PGC, SILB, SSIB, andt′ is the time offset between the tremor arrival at stations B or C and station A. To assignmagnitudes from energy, we identify those 4 s windows (a total of 391) that correspond tomatched filtered LFE detections with magnitudes from the catalog of Bostock et al. (2015),and perform a power-law regression to establish the appropriate calibration for momentmagnitude:Mw = 0.238 ∗ log10(E) + 0.319 (4.2)yielding the scalar moments for all available S-wave windows:M0 = 10(1.5Mw+10.7) (4.3)4.3 Results4.3.1 Tremor location clustersIn map view, the tremor is localized in four discrete, approximately planar clusters, sepa-rated by about 2 km, and labelled 1, 2, 3, and 4 in Figures 4.3 and 4.4. Each cluster exhibitsa different dip orientation that we quantify using principal component analysis to definebest-fit planes (Table 4.1). We also fit one plane to all the events and two planes (one to the“northern patch” (clusters 1 and 2), and a second to the “southern patch” (clusters 3 and4) (Table 4.1). Division into northern and southern patches was motivated by their distinctgeometries (Figure 4.3b) and their independent spatiotemporal rupture patterns (Figure4.4c). Comparing the one, two, and four patch solutions, we found that two planes (north-ern and southern patches) best fit the data. After removing outliers beyond 3 standarddeviations, we find that the northern plane dips 11.2◦ at a 66.8◦ azimuth and the southernplane dips 0.6◦ at a 106.9◦ azimuth (see Figure C.1). The average dip of the CSZ in thislocation is about 11◦ at a ∼N50◦E azimuth. We note that the San Juan Fault, a majorforearc block-bounding fault (Figure 4.2), divides the northern and southern patches. Thenumber of events in each cluster and its area are given in Table 4.1.80−123.8° −123.7° −123.6° −123.5°48.6°48.7°KLNB404484-02065053070ABC0 10km12340 2 4 6 8 10Along profile (km)37.538.038.539.039.540.040.541.0Depth (km)Profile A1234N/ACluster assignment0 2 4 6 8 10Along profile (km)37.538.038.539.039.540.040.541.0Depth (km)Profile B1234N/ACluster assignment0 2 4 6 8 10Along profile (km)37.538.038.539.039.540.040.541.0Depth (km)Profile C1234N/ACluster assignmenta) Profile Ab) Profile Bc) Profile CNW SESW NEW EFigure 4.3: Map view of 4S+1P tremor hypocenters with lines A, B, and C indicat-ing the profile sections to the right. White dots with number labels indicaterespective LFE template locations. Thin black contours are Moho depths ofthe Bloch et al. (2023) slab model. a-c) Depth sections of tremor hypocentersalong lines A, B, and C. Hypocenters are colored by patch number.442.5 445.0 447.5 450.0 452.5UTM X (km )538253845386538853905392539453965398UTM Y (km)2005-09-1600:0004:4809:3614:2419:1223:45Time (HH:MM)00:00 04:48 09:36 14:24 19:12 23:45Tim e on 2005-09-15 (HH:MM)0.02.04.06.08.010.012.014.016.0Along-profile distance (km)2005-09-16442.5 445.0 447.5 450.0 452.5UTM X (km )538253845386538853905392539453965398UTM Y (km)4S 4-circuit detect ions1234b) c)a)Figure 4.4: a) Map view of “4S” dataset epicenters and four clusters defined forslip/moment calculations. b) Temporal progression of tremor hypocenters (4Sdataset) color-coded by time on a particularly active day of the 2005 ETSepisode (16 September). Diagonal black dashed line indicates the profile usedfor panel c), which shows the along-profile distance of locations for the sametime-period (16 September 2005).81Cluster/Patch statisticsCluster/PatchDataset Number of Events Area (km2)2003 totalmoment2004 totalmoment2005 totalmoment2003-2005cumulativemomentAzimuth (deg.) Dip (deg.) Standard deviation (m)1 4S 2468 10.23 3.74e14 3.16e14 2.23e14 9.13e14 - - -2 4S 5234 16.83 5.73e14 6.19e14 6.03e14 1.79e15 - - -3 4S 1219 8.34 8.35e13 1.47e14 2.15e14 4.45e14 - - -4 4S 571 2.92 5.46e13 6.51e13 9.24e13 2.12e14 - - -1 4S+1P 567 4.55 - - - - 46.5 10.5 3062 4S+1P 2236 11.99 - - - - 69.6 10.5 2513 4S+1P 1005 7.50 - - - - 156.0 6.8 2024 4S+1P 215 1.73 - - - - 40.0 333.9 145North 4S+1P 3231 - - - - - 66.8 11.2 243South 4S+1P 1442 - - - - - 106.9 0.6 219Table 4.1: Parameters for each cluster (1, 2, 3, 4) or patch (north, south) for each dataset (4S or 4S+1P). Momentcalculations are carried out with 4S data only; cluster/patch dip orientations are only determined for 4S+1P datawhich have depth resolution. 4S data exhibit a slight spread in cluster areas when compared to the more confined4S+1P detections. This results in the slightly different cluster areas in the table and is shown in Figure 4.1.824.3.2 Tremor layer thicknessIt is evident in Figure 4.3 (profile B) that tremor is distributed in bands that are up to ∼400m thick. The question arises: how much of this apparent thickness is due to location errorsin depth and how much reflects the actual width of the shear zone? A comparison of absolutedepths calculated from data at common timestamps but filtered in a relatively narrow (1.5-6Hz) versus broad (1-8 Hz) frequency band exhibit a range of ∼300 m (see Methods, FigureC.2), implying that the actual maximum width of the shear-zone is probably on the order of100m. A constraint on relative depths comes from the distribution of the distances of eventsfrom the best-fitting plane. These distributions are near normal with standard deviationsof 243 m and 219 m and kurtoses of 5.3 and 4.2 for the northern and southern patchesrespectively (Figure 4.5). The kurtosis of a normal distribution is 3. The observation ofvalues significantly greater than 3 implies abnormally broad tails that probably reflect someseismicity outside a narrow shear zone. Our best estimated width of the shear zone is lessthan or equal to 400 m, depending on the error in depth (see Appendix C.2).4.3.3 Spatiotemporal progression and moment estimatesTo assess spatio-temporal behavior of epicenters and determine the moment release, weused the larger 4S dataset. Most of the 15,986 detections are readily associated with one ofthe four clusters and therefore provide a more complete measure of tremor excitation (seeFigure 4.4, Table 4.1).Spatio-temporal progression of the tremor hypocenters on 16 September 2005 (Figure4.4) clearly indicates that the north and south clusters behave independently. Similarbehavior is observed for the ETS episodes in 2003 and 2004 (Figure C.3).We estimated the scalar moment and moment magnitude from the energy for each tremordetection. As in Sammis and Bostock (2021), we observe a narrow normal distribution ofmagnitudes Mw = 1.60± 0.1, implying a log-normal distribution of scalar moments (FigureC.4). Details on the moment and magnitude calculation are provided in Methods. Totalmoment released in each of the four clusters is given in Table 4.1.4.4 DiscussionAccurate depth determination of the individual LFEs that constitute tremor beneath Van-couver Island provide two constraints on their source: 1) their location relative to knownsubsurface structures and, 2) the width of the shear-zone that produces them.83− 2.0 − 1.5 − 1.0 − 0.5 0.0 0.5 1.0 1.5 2.0Depth dista ce from PCA pla e (km )0100200300400500600CountsN = 3159σ = 243 mskew = -0.8kurtosis = 5.3North patch (clusters 1, 2)− 2.0 − 1.5 − 1.0 − 0.5 0.0 0.5 1.0 1.5 2.0Depth dista ce from PCA pla e (km )050100150200250N = 1411σ = 219 mskew = 0.6kurtosis = 4.2South patch (clusters 3, 4)UTM X (km )447 448 449 450 451 452UTM Y (km)53815383538553875389Depth (km)38.839.039.239.439.639.840.040.2South patch (clusters 3, 4) abovebelowUTM X (km )442 444 446 448 450UTM Y (km)53885390539253945396Depth (km)38.038.539.039.540.040.5North patch (clusters 1, 2) abovebelowa) b)c) d)Figure 4.5: North (a) and south (b) patch data colored by direction from the best-fitPCA plane. c) and d) show distributions of distances from the PCA plane forthe north (c) and south (d) patches, along with the number of events, standarddeviation, skew, and kurtosis.4.4.1 Comparison of tremor locations with other geophysicalobservationsConstraints on subsurface structure in the study region are provided by Lithoprobe seismicreflection profiling (Clowes et al., 1987), regional double-difference tomography incorpo-rating LFE templates (Savard et al., 2018), and receiver function studies (Audet et al.,2010, Bloch et al., 2023). Figure 4.6 shows tremor hypocenters from this study overlainon a northward extrapolation of the Lithoprobe 84-02 seismic reflection profile. Panel Aincludes Vp while panel B includes Poisson’s ratio, (i.e. σ =(Vp/Vs)2−22[(Vp/Vs)2−1]). A dipping zoneof quasi-parallel reflectors, dubbed the “E-layer” (Clowes et al., 1987), was interpreted torepresent underplated imbricate oceanic sediments and volcanics. However, tomographic840612182430364248200 400 600 800CDPDepth (km)EFSouth NorthWAR (Moho?)10 kmSeismic Line 84-024.0 4.5 5.0 5.5 6.0 6.5 7.0 7.5 8.0Velocity (km/s)0612182430364248200 400 600 800CDPDepth (km)EFSouth NorthWAR (Moho?)10 kmSeismic Line 84-02Poisson’s Ratio0.21 0.22 0.23 0.24 0.25 0.26 0.27 0.28 0.29slab ? slab ?Figure 4.6: Depth section of tremor (green dots) overlain on P-wave velocity (left)and Poissons’ ratio (right) (Savard et al., 2018) and the seismic reflection profile84-02 (Clowes et al., 1987), modified from Calvert et al. (2020). LFE templatelocations are in blue circles and forearc crustal seismicity are in black circles.Grey line is a wide-angle reflector interpreted by Calvert et al. (2020) to rep-resent the oceanic Moho. F = reflection horizon interpreted by Clowes et al.(1987) to represent the Moho.studies constrain the P-wave velocities at the depth of the E-layer below southern Vancou-ver Island to be ≥7 km/s (Figure 4.6; Savard et al., 2018) suggesting, instead, that thismaterial is predominantly mafic (Calvert et al., 2011, 2020). All of the E-layer, tremor andLFE template locations lie within a zone of unusually high Poisson’s ratio anomaly (∼0.28)that dips landward parallel to the slab, and is interpreted to indicate near-lithostatic fluidpressures in the tremor source region (Audet et al., 2010, Calvert et al., 2020) .Our tremor locations indicate quasi-planar, segmented layers in the plate boundaryregion just below the E-layer, at a depth near 39 km and are consistent with those of nearbyLFE template locations (Calvert et al., 2020, Savard et al., 2018) projected into the profile.Our observation that the E-layer lies above the active locus of seismic deformation (Figure4.6) is contrary to reports from previous studies in this region (Bombardier et al., 2023,85Kao et al., 2009). Although we cannot rule out the possibility that some tremor mappedaway from the four principal patches represents true detections, almost all detections arerelatively tightly constrained to layers less than 400 m thick.4.4.2 A mafic underplating model for tremorBased on our locations, we interpret tremor beneath Vancouver Island to represent maficunderplating, wherein each LFE represents shear failure within mixed brittle-ductile defor-mation occurring in the top few hundreds of meters of crystalline oceanic crust (commonlyreferred to as layer 2A (Houtz and Ewing, 1976)) of the subducting Juan de Fuca plate.This layer is expected to be rich in free fluids (to ∼4%; Hyndman, 1988, Peacock et al.,2011) as it undergoes active prograde metamorphism at tremor depths (e.g., Angiboustet al., 2022, Peacock et al., 2011). These fluids are produced by metamorphic reactions atlithostatic pore pressures (e.g., Condit et al., 2020) and promote material weakening andlocalized deformation. We expect that fluid distribution within the layer 2A will be hetero-geneous as governed by the distribution and persistence of fault-controlled fluid pathwaysconnecting to the seafloor prior to subduction (Fisher and Becker, 2000), thus influencingwhere deformation is partly brittle versus where it is purely ductile. Simple shear inducedthrough ongoing subduction causes the basaltic material to be continuously eroded leadingto comminuted wear products with an increasingly anisotropic fabric, that are graduallyplated onto the overriding lithosphere. Weaker volumes with higher water contents trappedby a highly anisotropic permeability are elongated by shear to produce the seismic reflec-tors that characterize the E-layer (Calvert and Clowes, 1990), which is mostly aseismic (e.g.,Figure 4.6). Localized areas where material transfer is occurring within the subduction zonemay manifest the distinct tremor patches as seen in Figures 4.3 and 4.4.Support for this interpretation is multifold. Based on analyses of multiple exhumedaccretionary complexes, Kimura and Ludden (1995) argued that exhumed, underplatedbasalt occurs as thin (≤300 m) layers derived exclusively from layer 2A, consistent withthe model interpretation laid out above. In a study of the Arosa zone, plate boundaryrocks exhumed from tremor depths, French and Condit (2019) noted that plate bound-ary slip occurs as frictional deformation in chlorite and talc schists surrounding blocks ofmetabasalt, consistent with our inference that tremor is hosted within layer 2A. Moreover,the remarkably coherent, coast-parallel distribution of tremor epicenters along the entireCascadia margin seems unlikely were tremor associated exclusively with sediments giventhe variable sediment input along the margin. This observation provides additional supportfor a basaltic layer 2A origin of tremor beneath southern Vancouver Island. As previouslynoted by Bassett and Watts (2015), this epicentral distribution also generally mirrors thehigh coastal topography associated with warm subduction zone settings, and is readily ex-86plained if tremor is a manifestation of crustal underplating (Angiboust et al., 2022, Menantet al., 2020). Moreover, in the limited locations where crustal seismic profiling has beenundertaken in the Cascadia forearc, notably from the Strait of Juan de Fuca through tocentral Vancouver Island (Brocher et al., 1990, Clowes et al., 1987, Nedimovic´ et al., 2003)and in central Oregon (Keach et al., 1989, Trehu et al., 1994), a highly reflective E-layerabove the inferred slab has been identified. This suggests that it, like tremor, is presentalong the full Cascadia margin and is a key element of the subduction complex.Although Kimura and Ludden (1995) suggested a “peeling” of layer 2A in the transferof metabasalt from subducting to overriding plates, the nature of tremor suggests an origininvolving significant cataclasis (Kirkpatrick et al., 2021, Oncken et al., 2022). For example,Sammis and Bostock (2021) observed log-normal distributions of LFE moments that theyargue can be explained via a model wherein shear failure at contacts between rigid grainsjammed within a viscous channel generates tremor. In the current context, we interpretthe granular and viscous elements of layer 2A to be associated with less altered tracts ofmetabasalt surrounded by a more intensely hydrated and overpressured matrix, respectively.The lognormal distribution of moments originates from the lognormal distribution of contactareas within jams expected as larger competent clasts are gradually broken down intosmaller ones. As the clasts decrease in size, they become less prone to jamming, and wesuggest that a scale-dependence set by layer 2A thickness contributes to the band-limitation(∼1-10 Hz) of tremor and the limited range of magnitudes observed in Sammis and Bostock(2021) and in this paper.As comminution proceeds, we expect increasing shear strain, ductile deformation, andgradual material transfer/transformation to the E-layer because of decreased density andstrength imparted by the release of fluids (Gerya and Meilick, 2011, Menant et al., 2020).We note that tremor and LFE template hypocenters lie on average 2-3 km below the base ofthe reflectivity that defines the E-layer in Figure 4.6. We discount the possibility of locationbias since the same velocity model (Savard et al., 2018) is used to locate hypocenters andmigrate reflections. Rather, we argue that, at some point in the comminution and shearingprocess, a permeability anisotropy “percolation” threshold is reached wherein fluids becomesegregated within horizontally contiguous “lenses” producing the pronounced reflectivityhorizons and abrupt base that define the E-layer, perhaps through changes in dihedralangle as suggested by Hyndman (1988). Although our estimates of slip within tremorgenicvolumes based on Kostrov strain significantly exceed those previously reported for tremorwithin the ETS zone more generally (Kao et al., 2010), they nonetheless fall far short ofthe plate motion budget (∼ 3mm versus ∼ 4-5cm per ETS episode) indicating that ductiledeformation must still dominate. It is likely then that the slow slip of ETS representsductile shear persisting well into the lower reaches of the E-layer at steadily diminishing87levels, both where tremor is well expressed and where it is not (Wech and Bartlow, 2014).4.4.3 Distance-thickness calculation of the E-layerWe assess the feasibility of E-layer assembly over realistic time periods using calculationsmodified from Clowes et al. (1987). The E layer is roughly 100 km long and about 5 kmthick beneath Vancouver Island, corresponding to a volume of 500 km3/km along strike.Roughly 1800 km of plate have been subducted beneath Vancouver Island over the last 40Ma (Clowes et al., 1987). If a thickness H of the basalt layer was eroded away during thistime to form the E-layer, the eroded volume is 1800H km3/km along strike. If the E-layerwas formed through basal accretion of layer 2A, then 1800H = 500, or H = 0.28 km, inrough agreement with the thickness of the layer 2A pillow basalts (Houtz, 1976, Houtz andEwing, 1976). Moreover, the total relative displacement of the Juan de Fuca plate D andthe thickness of the E-layer are consistent with the observed relation between displacementand thickness of crustal fault zones (Figure C.5; Hull, 1988, Scholz, 1987).4.4.4 Tremor as diagnostic of material transferFinally, we note the additional occurrence of tectonic tremor beneath accretionary prismsin subduction zones such as Nankai (Obana and Kodaira, 2009), where underplating is alsoimplied, and at major strike-slip faults (such as the Alpine Fault (Wech et al., 2012) or SanAndreas Fault (Nadeau and Dolenc, 2005)). We suggest that the occurrence of tremor inthese environments, as in the deep plate boundary of subduction zones, may be diagnosticof granular flow and/or material transfer in zones of high pore-fluid pressure (e.g., Berozaand Ide, 2011, Saffer and Tobin, 2011).88Chapter 5ConclusionIn each of the previous chapters, I discuss the conclusions and implications of the associatedresearch. Each chapter is a relatively independent work, yet as a collection this thesisis primarily concerned with the regional study of earthquakes to better understand theirtectonic context. Here, I discuss future research directions engendered by each chapter.5.1 Offshore microseismicityChapter 2 focused on a region that has been largely neglected in terms of new geophysicaldata acquisition ventures since the 1980’s. The most significant outcomes of this work arethe earthquake catalog including accurate epicenters, detailed documentation and analysesof seismicity in specific regions, refinement of an overarching model for the triple junctionevolution, and the recognition of seismicity in the Explorer microplate with an orientationparallel to the crustal spreading fabric. Our detailed analysis of seismicity in the regionhas laid the ground for more focused research projects. Recognizing this, Michael Bostockand others have secured funding through the NSERC ALLIANCE program for a multi-yearocean-bottom-seismometer experiment in the region. Specific target areas were identifiedbased on regions of interest in Chapter 2 and field deployments are currently on-going.Eventually, data from each of the deployments will be used to answer questions such as:• What is the nature of deformation at the advancing RDFE, and how is oblique slippartitioned between the RDFE and QCF?• How is local structure beneath the TWVF influencing deformation? Is there anyfurther evidence for/against compression in this area? Could local tomography providenew insights here?• What is the local structure beneath the Winona basin and palaeo-QCF? What is thenature of deformation at the continental margin?89• What is the local structure beneath the Explorer microplate? Local tomography,active seismic imaging, and earthquake studies would help answer the hypothesis ofdeformation along preexisting strength fabric in the region.• Is the Sovanco fracture zone acting as a region of initiating left-lateral faults into theExplorer plate as hypothesized in Chapter 2? Is there right-lateral motion?5.2 Anahim volcanic belt volcano-tectonicsSimilarly to Chapter 2, Chapter 3 addressed a tectonic area of British Columbia that hasbeen largely unstudied in recent years. This work has presented the first analysis of accurateearthquake locations and their focal mechanisms, leading to a tectonic interpretation basedon two major geological features in the region. Furthermore, our study resulted in the firstdocumentation of swarm seismicity in the area and comparisons with other volcanic regionsin North America. This study partly motivated the decision to profile the sea bottom ofthe Burke Channel/Bentick Arm on a recent Natural Resources Canada seagoing missionin order to investigate geomorphological evidence of surface faulting.Future temporary seismic deployments on land and aerial surveying would help to con-firm the hypothesis that these earthquakes are related to volcanic activity and structuralweakness from the Coast shear zone. A targeted deployment near the swarms to the north-west of Bella Coola would provide better constraints on their locations and mechanisms,allowing a more specific interpretation of their origin. However, this region comprises ex-tremely remote, rugged, heavily forested and glaciated terrain, which all make instrumentdeployments and surveying challenging and costly.5.3 Cascadia subduction zone source processesBoth Chapters 4 and Appendix D are concerned with the nature of the tectonic tremorsource and its relationship to processes in the Cascadia subduction zone. While the projectin Appendix D was ultimately not successful in addressing the hypothesis, it resulted in afew valuable outcomes as mentioned. Data from this seismic array is freely available andcan be used by scientists to tackle the mystery of tectonic tremor in Cascadia.Chapter 4 posed a novel connection between tectonic tremor and the process of maficmaterial transfer in the Cascadia subduction zone. This work was rendered possible througha fortuitous tremor source-station geometry that was not considered in the original arraydesign, but which can now be considered in other regions of Cascadia. 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Bulletin of the Seismological Society ofAmerica, 93(5):1875–1889, 2003. doi:http://dx.doi.org/10.1785/0120020190. → pages9, 79118Appendix ASupporting Materials for Chapter 2A.1 IntroductionThis supporting information includes 3 text sections and 10 figures. The first text A.1 de-scribes the approach to understanding sensitivity of relocations to stations. The second textA.2 describes analysis of Vp/Vs and depth z. The third text A.3 describes the Vp/Vs estimate.The figures exhibit the REST and GSC seismicity data, velocity models used during ini-tal location and relocation, maps showing the sensitivity of relocations to input parameterselection, comparison of initial and final relocations, magnitude detection threshold in theSFZ region, and comparison of seismicity detections in Middle and West Valley of the Juande Fuca ridge.A.2 Sensitivity of earthquake locations to station geometryGiven the suboptimal station geometry, we test the sensitivity of relocations on the mainRDF by relocating the events leaving one station out at a time. We do this for the fivestations with the most differential time recordings for this subregion (listed in descendingorder): MAYB, HG3B, PHC, BNB, BBB. Median location shifts are 0.239 km, 0.185 km,0.700 km, 0.600 km, and 1.58 km, respectively. Relocation without station BBB retainedthe fewest events (88%, compared to 90-91%). MAYB and PHC are on northern VancouverIsland whereas HG3B and BNB are on Haida Gwaii, thus BBB fills a critical gap in thestation distribution. The larger relocation shift and lower percentage of relocated eventswhen BBB is removed reflects the stronger dependence of relocation on the BBB locationcompared to those on Vancouver Island and Haida Gwaii, where additional stations areavailable.119A.3 Exploring the Vp/Vs - depth tradeoffIncreasing Vp/Vs and/or decreasing z shifts locations landward (Figure A.4d,e). We first setVp/Vs = 1.71 estimated using shallow earthquakes from the Dellwood Knolls area (Hyndmanand Rogers, 1981), and find that z = 3 best aligns epicenters with the offshore morphology.Second, we fix Vp/Vs = 1.725 as the best fit value for our traveltime data set as determinedusing the Wadati method (that is, Vp/Vs = 1 + the best-fit slope of P-wave travel time versusS-P time (Wadati, 1928); Figure A.5). In this case z = 5 km provides the best alignmentwith structure. Third, we test a higher Vp/Vs = 1.74 as might be anticipated for propagationpaths that include substantial portions of ultramafic lithology (Christensen, 1996), and findz = 7 km yields the best correspondence with seafloor structures. Epicenter solutions forthe three scenarios are very similar (Figure A.4a-c).We also attempt relocation by fixing Vp/Vs = 1.725 and allowing hypocentral depth tovary from an initial trial depth on an event-by-event basis. Figure A.6 compares the epi-center solutions obtained for z = 5 km and varying depth using Vp/Vs = 1.725. Epicentersexhibit more scatter in the variable depth inversions. In the variable-depth inversion, depthsdisplay some tendency for increasing depth with distance to the southwest which may indi-cate bias introduced by the velocity model. Both fixed and variable-depth inversions yieldgenerally similar locations. However, fixed-depth relocations exhibit stronger correlationwith tectonic features, so we prefer these epicenters and focus the rest of our discussionon these. In particular, we select the z = 5, Vp/Vs = 1.725 as best representative of ourdata set on the basis of the Wadati fit and use the locations obtained with this parametercombination for subsequent relocation with hypoDD.As the paired-pick threshold increases, epicentral scatter decreases but the key tectonicfeatures become more sparsely defined. For example, 8-paired-pick events fall more narrowlyon the RDF, suggesting that the fault zone seismicity is narrowly concentrated along thesurface trace of the fault, rather than the broader distribution suggested by the 4-paired-pick data (Figure A.7). Comparison of patterns in the seismicity as the threshold increasesenables a qualitative assessment of the error associated with Hypoinverse locations in the4-pair dataset.A.4 Vp/Vs and fluid flowOur bulk estimate for Vp/Vs = 1.725 for the offshore region is comparatively low for maficoceanic crust, which is typically around 1.8-1.9 (Hyndman, 1979). Inferred anisotropy inthe upper oceanic mantle (Malecek and Clowes, 1978) may contribute to depressing Vp/Vsvalues, as lower Pn velocity has been observed parallel to the Explorer ridge axis (7.3km/s) relative to the perpendicular direction (7.8 km/s) (Malecek and Clowes, 1978). Low120Figure A.1: REST-GSC matched events.Vp/Vs inferred from our Wadati plot may be biased by most of the earthquakes originatingwithin the fractured, presumably hydrothermally-altered fault zones - RDF, SFZ, Explorerridge, Dellwood Knolls and TWVF. Serpentinization and subsequent volume increase isinvoked to explain the uplift of crustal blocks composing the SFZ (Cowan et al., 1986) andthe NFZ (Rohr et al., 2018). Low Vp/Vs is also observed in the NFZ (Hutchinson et al.,2019), which they attribute to fracturing and subsequent seawater infiltration, which wouldcause serpentinization within the uppermost oceanic mantle. Furthermore, talc, a mineralcommonly produced during hydrothermal alteration, has a low Vp (6.59 km/s) and Vp/Vs(1.59) (Mainprice et al., 2008). We expect Vp/Vs to vary considerably within the entirestudy region, and given the 1-D Vp/Vs structure constraint in hypoDD, an averaged valuesuggested by the data is justified.121Figure A.2: REST and GSC non-matched events from the 4-paired pick culleddataset.122Figure A.3: Velocity models used during preliminary location with REST (red), GSC(black), and relocation (Waldron, 1982), blue).123a) b) c)d) e)Figure A.4: a-c) Comparison of epicenter locations obtained using three parameterscombinations: Vp/Vs = 1.71, z = 3 (yellow), Vp/Vs = 1.725, z = 5 (red), andVp/Vs = 1.74, z = 7 (blue). a) z = 3 locations plotted on top, b) z = 5 locationsplotted on top, c) z = 7 locations plotted on top. d) The effect of varying depthwith constant Vp/Vs = 1.725. e) The effect of varying Vp/Vs with constant z = 3km.Figure A.5: Wadati plots (S-P time versus P travel time) for data with ≥4 pairedphase picks (a) and data with ≥8 paired picks (b). Both yield a best-fit VpVs =1.725.124Figure A.6: Comparison of Hypoinverse location solutions for a) fixed z = 5 km, andb) trial z = 5 km. Black lines denote the RDF and QCF fault traces.125Figure A.7: Hypoinverse relocations of events with 8 paired P and S phase picksalong the RDF exhibit a narrow distribution along the RDF.126a) b) c)Figure A.8: Progression of relocations from initial locations (a, 6,471 events shown)to Hypoinverse relocations (b, 6,414 events shown) to the final hypoDD relo-cations (c, 5,792 events shown).Figure A.9: Left: Final hypoDD relocations color-coded by depth, Right: Histogramof final relocation depths. A minor seaward-deepening trend is evident alongthe RDF; most event depths do not deviate significantly from 5km.127Figure A.10: Cumulative seismicity detections in the West Valley and Middle Valleyregions of the Juan de Fuca ridge.128Appendix BSupporting Materials for Chapter 3B.1 IntroductionThis supporting information includes 8 figures and 1 table.σ1Subregion Trend PlungeNW Clusters 147.2 40.0BC Strands 342.3 64.0All 336.8 60.8σ2Subregion Trend PlungeNW Clusters 319.2 50.1BC Strands 163.0 26.0All 156.6 29.2σ3Subregion Trend PlungeNW Clusters 53.9 3.9BC Strands 73.0 0.3All 246.6 0.1Table B.1: Principal stress orientations (trend, plunge) from each stress inversion.129Figure B.1: P-velocity model used for relocation in Hypoinverse and hypoDD andfocal mechanism estimation in HASH. It is a generic continental model basedon that used for western Canada by the Geological Survey of Canada..130Figure B.2: Wadati plot (S-P time versus P travel time) yields a best-fitVpVs= 1.77.−127.5° −127.0° −126.5°52.0°52.5°BC0 10 20051015202530Depth (km) −127.5° −127.0° −126.5°52.0°52.5°BC0 10 20 2005-072005-102006-012006-042006-072006-10Date (YYYY-MM)a) b)cszcszPHSZPHSZkm kmFigure B.3: Relocations during the 2005-2006 BATHOLITHS deployment. a) Showsrelocations colored by depth and b) shows relocations colored by time, withsize scaled to magnitude. Map projection is Universal Transverse Mercator.131a) b) c)Figure B.4: 3D inset plots showing spatio-temporal progression of a) the 2015 swarm,b) the first part of the 2017 swarm, and c) the second part of the 2017 swarm(December).−128.0° −127.5° −127.0° −126.5°52.0°52.5°BS06BS10BC0 12.5 25kmCSZPHSZ4320342943Figure B.5: Vertical-component waveforms of two events recorded during the tem-porary BATHOLITHS installment at stations BS06 (top) and BS10 (bottom).Data have been highpass-filtered above 1 Hz. Red and blue lines denote P andS phase picks, respectively. Map shows each event as blue dots, labelled withtheir respective event IDs in white, and recording stations as red triangles.Stations BS06 and BS10 are labelled. Map projection is Universal TransverseMercator.132a) b)BATHOLITHSdeployment2015 swarm2017swarmFigure B.6: a) Cumulative event detections over time for the entire relocated datasetwith the BATHOLITHS deployment indicated by the shaded orange regionand 2015 and 2017 swarm onsets indicated by dashed black lines. b) Eventdetections binned by month, excluding the events occurring during the 2015and 2017 swarms.133Appendix CSupporting Materials for Chapter 4C.1 IntroductionThis supporting information includes 2 text sections and 11 figures.C.2 Positive tremor detection criteriaWe cull detections using four criteria: 1) four S delay-time circuits meeting the threshold≤ 0.0375 s, 2) mean S-wave cross-correlation coefficient 0.60 for the broad band data, 3)S-P cross-correlation coefficient 0.45 for all data, and 4) require the Gini coefficient of thedetection window to be 60 for all data. The Gini coefficient was originally used in socialscience research to measure wealth inequality, but in more recent applications it has beenused as a measure of sparsity (Hurley and Rickard, 2009). Detection windows with moreimpulsive wave arrivals are more “sparse” and yield a higher Gini coefficient (Figure C.8).These metrics were determined heuristically to emphasize hypocentral patterns but reducethe scatter resulting from false detections or real detections emanating from regions of lowtremor density.The detection and initial location procedure is applied to each template 053, 065, and 070using both the narrow and broad band data. We merge the narrow and broad band datasetsby assigning the narrow-band data detection for all common times. Common detectionsbetween templates are determined as those with common detection window timestampsand hypocentral locations within 200m. The selected location is chosen as the contributingdetection from the template the matched pair is closest to.134C.3 Width of shear layerIn Figure 4.2 of the main text, the seismogenic layers are about 500 m thick. We take thisfigure as an upper limit, yet uncertainty in depth determination implies that the actualthickness must be less. Comparison of locations of common events between narrow andbroad-band data in three different templates yields an absolute depth uncertainty of about300 m (Figure C.2). The layer width can be better estimated from the distribution of eventsabout the best fitting planes for one set of stations shown in Figure 4.5. The standarddeviations are 243 m and 219 m for the north and south patch respectively. The kurtosesof 5.3 and 4.2 indicate broader tails than a normal distribution which has a kurtosis of 3.0.If the events are uniformly distributed in a zone of width w, the kurtosis is always lessthan 3 and depends on the ratio wstd where std is the standard deviation of the measurementerror. To quantify this relation, we initially place 2000 points at random depths between−w/2 and w/2. We then move each point by a random distance taken from a normal dis-tribution with zero mean and an assigned standard deviation that varies from std=w tostd= w10 . For each normal distribution, we run 1000 test cases (moving the points differently1000 times), and calculate the kurtosis of the moved data points each time. The averagekurtosis and its standard deviation over 1000 runs are shown as a function of wstd in FigureC.10. Note that the kurtosis = 3 for wstd = 1 and that a kurtosis of 3 is possible at the onestd level for 1<std<2. Thus, based on the kurtosis values 3 for the north and south clusters(Figure 4.5), the width of the seismogenic layer must not be much larger than the standarddeviation (243 and 219 m for the north and south, respectively).135UTM X (km)444446 448 450452UTM Y (km) 53845386538853905392Depth (km)3436384042Figure C.1: Best-fit planes to the north and south clusters (clusters 1,2, and 3,4, re-spectively) exhibiting the steeper dip of tremor in the northern clusters (about10 degrees). Northern plane is orange and southern plane is blue.136Figure C.2: Top row: Comparison of locations from common time detection win-dows between the narrow and broad band data for each template. Bottomrow: Depth discrepancies for each matched detection do not show a significantshallowing or deepening bias.442.5 445.0 447.5 450.0 452.5UTM X (km)53825384538653885390539253945396UTM Y (km)2003442.5 445.0 447.5 450.0 452.5UTM X (km)53825384538653885390539253945396UTM Y (km)2004442.5 445.0 447.5 450.0 452.5UTM X (km)53825384538653885390539253945396UTM Y (km)2005Figure C.3: Comparison of 4S detections for each ETS episode in 2003, 2004, and2005. To varying extents, each tremor patch is activated in each year.1371.0 1.2 1.4 1.6 1.8 2.0Mw0200400600800100012001400CountsCluster Moment Magnitude DistributionsC1C2C3C40.0 0.2 0.4 0.6 0.8 1.0 1.2M0 1e1202505007501000125015001750CountsCluster Moment DistributionsC1C2C3C4a) b)Figure C.4: Histogram distributions of a) moment magnitude, Mw, and b) scalarmoment, M0, for each cluster.138Log H (Thickness, m)Log D (Slip, m) 1 km100 km10,000 km1 kmSubduction D=8000 kmE-layer thickness = 5 kmFigure C.5: Log-log plot of gouge thickness H as a function of total slip D redrawnfrom Scholz (1987). Point SAF was estimated by Scholz for the San Andreasfault in central California. The line D = 10−3H is extended to the slip and E-layer thickness appropriate for subduction beneath Vancouver Island (Calvertet al., 2020).1395380 5385 5390 5395 5400UTM Y (km)32343638404244464850Depth (km)Y-Depth053Savard 053065Savard 065070Savard 070440 445 450 455UTM X (km)32343638404244464850Depth (km)X-Depth053Savard 053065Savard 065070Savard 070435 440 445 450 455UTM X (km)5380.05382.55385.05387.55390.05392.55395.05397.55400.0UTM Y (km)X-Y053Savard 053065Savard 065070Savard 070Figure C.6: Comparison relocations between LFE template waveforms using themethod presented in this study (black, blue, red circles), and the LFE templatelocations of Savard et al. (2018) (black, blue, red stars) superposed on the fulltremor catalog (blue dots) of this study.447 448 449 450 451 452 453UTM X (km)5388538953905391539253935394UTM Y (km)SNB common timestamp detections (339)053 N053 B065 N065 B070 N070 BFigure C.7: Comparison of 339 detection time windows that yielded tremor locationsfor all six possible datasets, that is, the broad- (1-8 Hz) and narrow- (1.5-6 Hz)band data for each of the three templates (053, 065, 070) used for detection.Average epicentral discrepancy is about 250 m.140Figure C.8: Comparison of high- and low Gini index (1-8 Hz) detections (Gini=89,left; Gini=61, right; respectively). Top panel shows the S-waveforms alignedaccording to template, middle panel includes S-waveforms after fine-scale align-ment and includes their first principal component (black dashed line), and bot-tom panel compares the fine-scale aligned P-waveform with the first S-waveprincipal component.141442.5445.0447.5450.0452.5UTM X (km)538253845386538853905392539453965398UTM Y (km)4S1P 4-circuit detections4S1P4SFigure C.9: The slight difference in dense cluster areas for the 4S1P (blue lines) and4S (red lines) data sets. Cluster outlines are overlain on the 4S1P data tohighlight the varying area extent of the dense clusters for the 4S and 4S+1Pdata.142Figure C.10: Average kurtosis versus width/standard deviation for events with ran-domly assigned initial depths between −w/2 and +w/2 that are moved randomlybased on a normal distribution defined by a standard deviation that variesfrom std=w to std = w/10. The average kurtosis and its standard deviationcalculated over 1000 test cases for each std assignment.143Appendix DFrequency-dependent arrayanalysis of tremor in northernCascadiaD.1 PrefaceIn this section I include a report on a brief unpublished project undertaken during myPhD. This project was valuable both as a learning experience for me and my co-authors,and because through it I provided a new tremor dataset to the community. Furthermore,the work from this project motivated the design of Chapter 4.D.2 IntroductionAs discussed in Chapter 4, slow slip events are usually accompanied by seismic tremor, anoisy, bandlimited signal (∼1-10 Hz) that exhibits preferential excitation of S-waves (Kat-sumata and Kamaya, 2003). Tremor is believed to be composed of low-frequency earth-quakes, small (moment magnitude ∼ 1-2) earthquakes that are depleted in frequencies>∼10 Hz relative to regular earthquakes of comparable magnitude, the reason for whichremains unclear. Due to its comparatively bandlimited nature, analysis of tremor is typi-cally restricted to the ∼1-10 Hz frequency band (e.g., Armbruster et al., 2014, Ghosh et al.,2010, Rubin and Armbruster, 2013). The goal with this work is to determine if distinctspatiotemporal behavior can be discerned between different frequency bands within, andeven above, the typical tremor bandwidth of 1-10 Hz. Distinct spatiotemporal patternsobserved in different frequency ranges may reflect heterogeneity in the source region andwould aid in our understanding of the tremor source and its relation with LFEs.1441-2 Hz 2-4 Hz 4-8 Hz Velocity [nm/s]1-2 Hz 2-4 Hz 4-8 Hz Velocity [nm/s]PA09 HHE on 08-22-2019Time (seconds)Figure D.1: Comparison of tremor waveforms on 22 August 2019 filtered in 1-2 Hz, 2-4 Hz, and 4-8 Hz frequency bands exhibits variations in wave arrivals betweenthe different data frequencies.In particular, we are motivated by preliminary observations of tremor waveforms thatexhibit distinct arrivals in higher frequencies that are not observed in data filtered in lowerfrequency bands and are coherent along an array of seismic stations (Figures D.1, D.2).Yet, arrivals for known LFEs are consistent across the three frequency bands (Figure D.3).We therefore hypothesize that structural heterogeneity in the plate boundary zone causesradiation of tremor at different dominant frequencies during the slow slip process, and thatLFEs emanate from localized regions of more pronounced and discrete heterogeneity. To testthis hypothesis, we use a newly-deployed, densely-spaced seismic array on Vancouver Islandto study spatiotemporal characteristics of tectonic tremor filtered in varying frequencies todetermine if different parts of the plate interface preferentially radiate seismic energy atcertain frequencies.Southern Vancouver Island provides an ideal setting to test this hypothesis because its1454-8 Hz wavefront on HHE for 7 PISA stations Velocity [nm/s]Time (seconds)Figure D.2: East-component velocity waveforms of tremor arrivals in the higher-frequency (4-8 Hz) band on 22 August 2019 are coherent along an array ofseismic stations (7 out of 12 stations are shown). Time window is the same asFigure D.1.geology is dominated by the Wrangellia terrane, which comprises remnants of an accretedoceanic plateau (basalts and marine sedimentary rocks (Jones et al., 1977)) and exhibitslow scattering and intrinsic attenuation of seismic waves. These geologic conditions alsoenable the study of low-amplitude, emergent tectonic tremor signals and constituent low-frequency earthquakes with higher resolution and greater fidelity than has been achieved atother locations along the Cascadia subduction zone margin (e.g., Plourde et al., 2015).D.2.1 PISA seismic arrayIn anticipation of an ETS event in August 2019, we deployed 12 seismic stations to forma small-aperture seismic array called the Pacific Geoscience Centre & Institute for OceanSciences Seismic Array (PISA). The sites are situated on land on and around the Geological146August 21 2019 LFE #240 1-2 Hz 2-4 Hz 4-8 Hz 1-8 Hz Velocity [nm/s]Time (seconds)Figure D.3: East-component velocity waveforms of on 21 August 2019 show coherentarrivals in each frequency band. Dashed contours are Juan de Fuca slab depthsfrom the McCrory et al. (2012) model. Red triangle indicates the PISA arraylocation and magenta star with “240” label indicates the general low-frequencyearthquake template location. The arrival is consistent across all frequencybands.Survey of Canada Pacific Geoscience Centre (PGC) in Sidney, BC, Canada (Figure D.4).This location is above the source region of tremor, which, assuming originate at the plateboundary (c.f., Kao et al., 2009), is between 30-40 km depth in this region according tovarious Cascadia slab models (e.g., Bloch et al., 2023, Bostock et al., 2019). Proximity to thetremor source is essential because, aside from the signals experiencing a lower degree of pathattenuation, it reduces the likelihood of free-surface interactions distorting the waveforms(Savard and Bostock, 2015).The array aperture is 667m x 664m, with a maximum inter-station spacing of 672m(Figure D.4). There are 12 sites associated with the PISA array that can be easily ac-cessed on foot or by vehicle from PGC. The sites were originally installed in 2012 to study147tremor and slow slip beneath southern Vancouver Island, yet have been quiescent since theirinstallation. Thus, this is the first study to employ these instruments.The stations are Nanometrics Trillium 120 Broadband 3-component sensors sampling at100 Hz. At PGC, the stations complement the existing Canadian Network surface stationPGC, as well as 3 Plate Boundary Observatory (PBO) stations B009, B010, and B011.These borehole stations are situated ∼198m below the ground. Each PISA site is situatedon a concrete block with a GPS antenna and metal casing for the equipment. The metalcasing contains the seismic instrument, a digital recorder, four 12V batteries, and a solarcontroller. Atop the metal casing are, for all but one site, 2 solar panels. The exception,station PA01, is connected via AC power to PGC and therefore only needs one battery.There is not currently an established radio link for all of the stations so that data may betelemetered to the data collection centre at PGC, so I collected data and swapped memorycards in September 2019 and February 2020.The 2019 ETS episode in southern Vancouver Island began on 11 August in the Juan deFuca Strait, and appeared to have bifurcated, proceeding both to the northwest to Victoriaand to the south towards Seattle, stopping on 25 August (Malone, 2019). This event wasunusual in that its duration was only a couple of weeks, with little to no observable GPSdisplacement (Malone, 2019).D.2.2 Tremor dataThe data are velocity time series that are demeaned, detrended, corrected for instrumentresponse and filtered in 1-2 Hz, 2-4 Hz, 4-8 Hz passbands. We use a zero-phase, 2-poleButterworth filter. While higher-order filters would allow less spectral leakage into otherfrequencies, frequency dispersion becomes more significant with increasing filter order. Thisarises when the filter corners are near a change in the instrument’s frequency response, andtherefore, the phase offset is not constant. The Butterworth filter has a slightly non-linearphase response, resulting in minor dispersion when applied. This frequency dispersion effectcan be exacerbated by increasing the number of poles used, which generally causes widerfilter group delays (Goldstein et al., 2003). In this study a zero-phase filter is used topreserve phase information, imperative when comparing relative arrivals of traces.D.3 Method: Cross-station array analysisWe use a cross-station approach to locate tremor, following the slowness analysis method fora small-aperture array used by McCausland et al. (2010). This approach is similar to thatof Rubin and Armbruster (2013) and Armbruster et al. (2014) in that we are comparingwindows of data at the same time at a set of stations (in this case, more closely-spaced148−128˚−128˚−126˚−126˚−124˚−124˚−122˚−122˚46˚ 46˚48˚ 48˚50˚ 50˚PGCVancouverSeattleFigure D.4: Map of study area, indicating the Pacific Geoscience Center in Sidney,BC. Purple line denotes the contour of tremor from 2009-2018 (Wech, 2010).Inset shows a Google Earth image of the PISA array (blue diamonds) and ex-isting stations in the area (PGC: pink hexagon, PBO stations: red diamonds).149stations than those used in Rubin and Armbruster (2013) and Armbruster et al. (2014)),instead of comparing waveforms from different time windows at a single station. Because ofthe dense PISA station distribution, waveforms are coherent between stations and envelopefunctions are not necessary (e.g., Wech and Creager, 2008).We solve for the apparent slowness of tremor (i.e., the ray parameter) in sliding windowsof tremor data. In each window, the cross-correlation-derived relative arrival times of thetremor wavefront between station pairs of the array are computed and used along with thestation spacings to determine the x- and y- components of slowness. We tested multiplewindow lengths, using both constant window lengths for all frequencies, and variable-lengthwindows for the different frequencies. For the variable-length windows, the window lengthfor each passband was halved as the frequency band increased by a scale factor of 2, so as tocompare an equal number of cycles between frequency bands. Henceforth, “short windows”are 4, 2, 1, 0.5 s long windows for 1-2, 2-4, 4-8, and 8-16 Hz, respectively. Long windows are16, 8, 4, 2- s long windows, and extra-long windows are 64, 32, 16, 8- s long. Short windowlengths have the advantage that cross-correlations between waveforms are generally higher,however, this can easily include highly-correlated noise, resulting in locations of spuriousnoise. Long windows are more likely to have high correlations for meaningful signals andlow correlations for noise, however, do not capture spatiotemporal behavior with the sameresolution as short windows (e.g., Armbruster et al., 2014). Thus, there is a tradeoff betweenlong and short window lengths.D.3.1 Relative arrival timesWe first solve for the relative arrival times of the tremor at all stations using the multichannelcross-correlation algorithm (MCCC) (VanDecar and Crosson, 1990), in which the relativearrival times are determined by maximizing the cross-correlation functions for all stationpairs.For a section of n traces, we solve for relative arrival times of the wavefront at station iusing MCCC with L1-norm minimization of the systemWt =∆t (D.1)as outlined in the Introduction. Again, W is a sparse coefficient matrix of size n(n−1)2 + 1x n, t is the solution vector containing the relative arrival times ti for each trace i, and∆tij = ti − tj is the vector containing cross-correlation-derived delay times (i.e., the lagcorresponding to the maximum of the correlation function for each pair of traces ϕij). Thissystem includes the zero-sum constraint,∑ni=1 ti = 0, which forces the mean of the optimalarrival times to be zero. In practice, we compute the cross-correlation function in the150frequency domain using the fast Fourier transforms F of the ith and jth seismograms.ϕij = Re(F−1(F(Xi) · F(Xj)∗)) (D.2)The cross-correlation coefficient r for station pair ij is then the cross-correlation functionevaluated at its maximum, τijmax, normalized by the standard deviations σ of the twotraces (Equation D.3). We assume the data follow a standard normal distribution (i.e.,zero-mean and variance = 1), so σ is computed as the square-root of the maximum of theautocorrelation function of station i or j.rij =ϕij(τmaxij )σiσj(D.3)We restrict the search window for the cross-correlation maximum to ±0.25 seconds toensure physically reasonable delays times based on the station spacings and the requirementthat incoming waves have pre-critical incidence angles. Further, delays occurring betweentime samples were obtained by fitting a parabola to the values nearest the maximum of thecross-correlation function and then using the time lag corresponding to the maximum of theparabola’s peak. These delay times are then used to generate a linear system of equationswhose solution yields the optimal set of relative arrival times.We deviate from the original VanDecar and Crosson (1990) method and McCauslandet al. (2010) by solving for the L1-solution to the system of equations instead of the classicL2-solution. The L1-solution is known to be more robust to outliers and noise in the data(Scales and Gersztenkorn, 1988). One disadvantage to the L1-norm is that it requires aniterative solver to converge to a solution, making its minimization more computationallyintensive than the L2 solution. Yet, the L1-solution is especially advantageous when usingrelatively high frequency data (e.g., 4-8 Hz). We find that for lower frequencies 1-2 and 2-4Hz, the L2 solution yields satisfactory results for the relative arrival times, however, in the4-8 Hz passband it suffers from cycle skips and noise present in the data. While Shearer(1997) showed that using the L1-norm in waveform cross-correlation improved earthquakelocations, it is not commonly seen in the more recent literature.Following the initial waveform alignment, we perform a second fine-scale alignment,employing the L2 solution and allowing very small time lags (+/ − 0.03125 s - half theperiod of an 8 Hz signal) to search for maxima in the cross-correlation function.Along with relative arrival times, we keep track of timing uncertainties in the formof mean residuals σ and mean cross-correlations in each sliding window. We develop anadditional parameter to diagnose windows that exhibit coherent energy. For each window,we perform a singular value decomposition (SVD) to obtain the singular values. If the151system of shifted waveforms is well-aligned, the first singular value will be much larger thansubsequent values, indicating a primary “mode” in the window. However, if the windowcontains incoherent noise, the singular values will be more similar in magnitude, representingmore “modes” present in the window. Thus we parametrize the singular values by takingthe ratio of the Euclidean sum of the singular values σi, i = [2, n] for n singular values,normalized by the largest (first) singular value. The ratio is large when the waveforms areincoherent and exhibit a large degree of variability within the window, and the ratio is smallwhen the waveforms exhibit a coherent signal that is well-aligned.D.3.2 Slowness vector solutionWith the optimal relative arrival times ti, we follow the approach of McCausland et al.(2010) and solve for the horizontal slowness vector s (x and y components of the ray pa-rameter; inverse of velocity) by solving Equation D.4 using least-squares. We analyze datain three different passbands, whereas (McCausland et al., 2010) use a single passband, 1.5-4Hz. Other array analyses of tremor in northern Cascadia have used passbands of 2-8 Hz(Rocca et al., 2010) or 3-9 Hz (Ghosh et al., 2010).Bs = δt (D.4)Here, δt is a vector containing the relative arrival time difference between station i andstation j for all station pairs, ti−tj , and B contains the corresponding station-pair spacingsin the X- and Y- directions in km.We weight the equation according to the errors, using the diagonal covariance matrixC:Bw = C−1/2B (D.5)δtw = C−1/2δt (D.6)Bws = δtw (D.7)From Equation D.7, we obtain the observed slowness s. By raytracing, we can obtainthe predicted S-wave slowness as the x- and y- components of the modeled S-wave rayparameter, ps (where ϕ is the modeled backazimuth).ss = [pssinϕ, pscosϕ] (D.8)We obtained backazimuth and slowness values in moving windows of data during the152strongest periods of tremor in August and November 2019. Window lengths are varied byfrequency, such that we use longer windows for lower frequency bands to capture the samenumber of cycles. In principle, longer windows yield more stable, smooth results at the costof finer temporal resolution. We empirically determine optimal window lengths to be be 16,8, and 4 s long for the 1-2 Hz, 2-4 Hz, and 4-8 Hz data, respectively (Figure D.6).D.3.3 Station correctionsTo correct for inconsistencies in wavefront arrival times at the station array, we introducea station correction term, τ , as the difference between the observed arrival time differencesand the modeled travel time differences for station pairs ij:τij = δtobs − δtmodel = δt−B · ss (D.9)We can again employ the zero-sum rule to obtain optimal τi for each trace. Then thefinal, corrected relative arrival times tif will be:tif = ti + τi (D.10)We initially intended to employ station correction terms with the PBO stations about100m below ground. To do so, we selected 6 regional earthquakes and calculated theirpredicted arrival time at the PBO station locations as if they were at the surface. We thentook the difference between this arrival time and their observed arrival time as the stationcorrection term (Figure D.5). However, we found that the arrival time differences werehighly variable for different earthquake events (Figure D.5), and therefore we were not ableto apply a stable station correction value to each station. We omitted the PBO stationsfrom further analysis.D.4 ResultsI first display results for select windows of data during the 2019 ETS episode in August andNovember. In general, 1-2 Hz and 2-4 Hz data yield similar slowness results (Figures D.7,D.8,D.9), whereas the 4-8 Hz data yield slightly reduced slowness magnitudes and elevatedbackazimuths (Figure D.9). This results in 4-8 Hz detections locating generally closer toand northwest of the array on each day between 20-22 August (Figure D.10). X- and Y-positions of the tremor locations along strike for each of the days follow similar patternsacross the three frequency bands, with the 4-8 Hz locations showing a lower range in values,probably due to reduced higher frequency detections at larger distances (Figure D.11).I also test the stability of the solutions by solving for slowness using 4-station subgroups153Figure D.5: Arrival time differences at the borehole stations for 6 test earthquakes.Events on the x-axis are given as dates in month-day (MMDD) format.Figure D.6: Comparison of slowness and backazimuth values determined using vary-ing window lengths for varying frequency bands. Window lengths shown are[4, 2, 1], [16, 8, 4], and [128, 64, 32] seconds for 1-2 Hz, 2-4 Hz, and 4-8 Hzdata, respectively.154Figure D.7: Slowness values for 10 hours on August 20 for data of varying frequencybands, plotted with slowness values for known LFE locations recorded at thesame time. Top panel displays raw values, whereas the bottom panel displaysa running mean.of the PISA array. Although the PA01-PA04 subgroup exhibits lower cross-correlationvalues, slowness (and therefore backazimuth) values are comparable (Figure D.12).However, a closer look at the arrival time contours determined using the tremor windowsreveals a relatively planar field for the 1-2 and 2-4 Hz data that is significantly degraded forthe 4-8 Hz data (Figure D.13). This is also seen for the regional earthquakes, which shouldbe the most well-constrained (Figure D.14). Like the PBO stations, station corrections forthese events is apparently temporally-variable, precluding application of a simple correctionterm.155Figure D.8: Backazimuth values for 10 hours on August 20 for data of varyingfrequency bands, plotted with backazimuth values for known LFE locationsrecorded at the same time.156Figure D.9: Mean cross-correlation (CC), mean residuals, slowness, and backazimuthresults for 1-2 Hz, 2-4Hz, and 4-8 Hz frequency data on 08-22-2019.157Figure D.10: Tremor location density maps for 1-2 Hz, 2-4Hz, and 4-8 Hz frequencydata on 08-22-2019. Red contour on the 1-2 Hz and 2-4 Hz maps indicatesthe 4-8 Hz location contour.158Figure D.11: X- and Y- coordinate positions (in km) for 1-2 Hz, 2-4Hz, and 4-8 Hzfrequency data for August 20-22 2019.Figure D.12: Comparison of cross-correlation, slowness, and backazimuth results for4-station subgroups of the PISA array.159Figure D.13: Arrival time contours for 4-8 Hz tremor on 08-22-2019.Figure D.14: Arrival time contours for a test regional earthquake on September 32019. (a) displays the location of the earthquake relative to the PISA array,as well as the predicted slowness and backazimuth values. (b), (c), and (d)display the arrival time contours for the 1-2 Hz, 2-4 Hz, and 4-8 Hz data,respectively.160D.5 Discussion and ConclusionWithout consistent arrival time contours across the three frequency bands, we cannot re-liably solve for slowness/backazimuth values for tremor and compare them between thedifferent bands. This is because a planar incoming wavefield is a key assumption of themethod. For this reason, we conclude that this method is not suitable for frequency-dependent tremor locations, at least at this location. While we are not able to report onrobust frequency-dependent variations in tremor source characteristics, there are severalother useful outcomes of this study:1. Discrimination of simple array analysis as a method for determining frequency-dependenttremor source locations. Further work in other regions or using other types of signalsmay be useful.2. Revival of a small-aperture array in an ideal setting for the study of tremor in theCascadia subduction zone or other regional earthquakes. Data is openly available atIRIS (https://ds.iris.edu/gmap/#network=C8&station=PA*&planet=earth).3. Personal development from field maintenance of the PISA stations and collection ofthe data.161
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Bostock, Michael G.
A collection of earthquake studies throughout western British Columbia
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