AN ABSTRACT OF THE THESIS OF

Annette V. deCharon for the degree of Master of Science in Oceanography presented on October 14. 1988. Title: Structure and Tectonics of Cascadia Segment. Central Blanco Transform Zone Redacted for Privacy Abstract approved: Robert W. Embley

Seismic-reflection profiles and SeaBeam bathymetry are used to examine the structural development of the Cascadia Segment, a region of back-tilted blocks flanking a central depression within the Blanco zone. The depressed position of the Cascadia Segment has caused this area to act as a sediment trap for bottom transported material moving through Cascadia Channel (Duncan, 1968; Griggs & Kulm, 1970). The existence of thick turbidite sequences atop tectonic blocks this an ideal place for seismic-reflection imaging. Escarpment slopes and fault dip angles in Cascadia Segment range between 110 and 31°, which are, on average, shallower than values reported from slow-to-intermediate-rate spreading centers. The perched turbidite sequences are assumed to have originated within a central rifting basin. The lack of unconformable sequences within the back-tilted turbidites indicates that the tectonic blocks of Cascadia Segment have been uplifted in a continuous manner. Thus, a quantitative analysis of tilt of turbidite sequences versus distance from the present central rifting basin, Cascadia Depression, is used to understand the structural development of the Cascadia Segment. The high degree of tilt seen in turbidite sequences proximal to the present central rifting basin is consistent with a model in which turbidite flows are deposited within a basin that is being uplifted and spreading away from a narrow extensional zone. Individual tectonic blocks rotate at least 6° as they move through the terraces of Cascaclia Segment. Extension of the eastern terrace has been accommodated by rotation of three large, discrete blocks whereas the western terrace has extended through block rotation and a large number of normal faults with smaller throws. Irregular morphology of the flanking blocks of this spreading center may be attributed to short-term (on the order of 106 years) asymmetric spreading. Turbidite tilt data and fault orientations demonstrate conclusively that the muting of topography from the rift valley to the rift mountains greater than 30 km off the central spreading axis is accomplished by: (1) movement along outward-dipping normal faults in the rift mountains, and (2) back-tilting of tectonic blocks within the terraces. The possibility of reverse motion along pre-existing inward-dipping faults in the rift mountains is not refuted by tilt data from Cascadia Segment. Sediment thickness on the floor of Cascadia Depression varies between less than 50 m and greater than 850 m. Basement doming, evident in the center of the depression, may be caused by the intrusion of sills into the sediments. There is a great disparity in basement slope between the eastern half (average slope = 11.5°) and western half (average slope = 4.2°) of Cascadia Depression. Asymmetry of basement slope seems to be established in the central rifling basin and is propagated outward onto the terraces until isostatic equilibrium is reached. The local slope of Cascadia Depression's floor is the result of interaction between two components: rate of movement along faults within the basin and the rate at which turbidite-derived sediments can bury the faults. At present, movement along low-angle fault surfaces within Cascadia Depression is occurring at a higher rate than sediments are being delivered to the basin. Duncan (1968) has estimated that the rate of turbidite sediment accumulation in this area decreased rapidly about 12,000 years ago. Fault angles and offset values were used to calculate extension rates averaged over this time period. A progressive decrease in opening rate from 5.8 cm/yr to 4.1 cm/yr is evident from the northern to central parts of Cascadia Depression, respectively. This discrepancy in estimated opening rates is probably the result of: (1) "true" short-term asymmetric opening from north-to-south; (2) the masking of surface offset along faults towards the southern part of Cascadia Depression by ponding of turbidites; or (3) a combination of both. The calculated rates of opening support a model in which Cascadia Segment is a "trapped" spreading segment from an episode of rift propagation (Embley et al., in press). The calculated extension rate of Cascadia Depression coupled with the morphologic similarity between Cascadia Segment and slow-to-intermediate mid-ocean ridges support the idea that this is an area of active . The tectonic evolution of Cascadia Segment is compared to existing models of rift valley evolution and crustal generation processes at mid-ocean ridges, especially in the context of steady state versus non-steady state theories of rift valley development (i.e., Davis & Lister, 1977; Kappel & Ryan, 1986). The high subsidence rate of the rift valley floor (1.8cm/yraveraged over 6,660 years) coupled the calculated horizontal extension rate (about 5cm/yraveraged over the past 12,000 years) may indicate that the current magma supply has waned, thus allowing extension to become the dominant mechanism at Cascadia Depression. Therefore, recent tectonics may support the idea that the balance between extension and volcanism at Cascadia Segment may be episodic in nature (after Kappel & Ryan, 1986). Structure and Tectonics of Cascadia Segment, Central Blanco Transform Fault Zone

by

AnnetteV. deCharon

A THESIS submitted to State University

in partial fulfillment of the requirements for the degree of

Master of Science

Completed October 14, 1988 Commencement June, 1989 APPROVED:

Redacted for Privacy

Adjunct Professor of Oceanography in charge of major

Redacted for Privacy

Dean of College of Oceanography

Redacted for Privacy

Dean of GraduateS9i

Date thesis is presented October 14. 1988 Typed by researcher for Annette V. deCharon ACKNOWLEDGEMENTS

During the course of my graduate work, I have come to know and respect many people, both professionally and personally. I would like to express my gratitude to all those who have encouraged me and facilitated the completion of this thesis. I have a great amount of admiration for my major advisor, Dr. Bob Embley, who has continually supported me as a student and a friend. My defense committee, Drs. LaVerne Kuim, Shaul Levi, and Wayne Seim, made a great contribution to content of the thesis and I certainly appreciate their efforts. Much of the seismic-reflection data used in this study was generously provided by Dr. Bill Normark of the U.S. Geological Survey. His quick (yet thorough) edit of the first draft of the thesis hastened its completion and I am grateful for his help. Discussions with Drs. Clive Lister and P.J. Fox helped with the interpretation of the tectonic and structural aspects of the work. I wish to thank Drs. Fred Spiess, John Delaney, and Clive Lister for access to data collected during the 1987 cruise aboard the R/V Melville. Lingyun Chiao spent many hours generating a seismic-reflection profile that was used in this study. Computer-digitization of seismic-reflection records was possible thanks to the help of David Borg-Breen of NOAAIPMEL, Andy Lau and Marijke Van Heeswijk. Figure 4 was drafted and sent to me (at a moment's notice) by Gini Curl of NOAA/PMEL. Three of the people who have had a great influence on my "budding" career in Marine Geology are Drs. Jim McClain, Dallas Abbott, and Jim Franklin. Their good humor, excellent advice, and enthusiasm will remain with me for a long time. My position as a graduate research assistant through the NOAA "Vents" program allowed me to work closely with people from Corvallis and Newport. My first year-and-a- half of graduate school was a pleasant experience thanks to support of the Oceanography faculty, students, and staff. Friends such as Dave M., Collin, Tom, Cindy, Yip, Bruce A., Bruce F., Bruce S., Anne P., Karen, Dana, Beth, and the "Bull Pen" crowd made that time more enjoyable. The Secher Ave. gang will always be remembered for its contribution to my sanity: Jana, Collin, Tom, Heidi, Chris, Lucy, Jack, Rambo, etc. The tolerance and support of Newport folks warrants special recognition. Many thanks to Bob, Steve, Chris, Susan, Bruce, Michele, Maria, Andy, Marijke, Tom, Saskia, Karen D., Anne D., Jessica, Robin, and Pete/Bud. I want to let my family know that I appreciate all the support and sacrifices that they have made for me. The encouragement given to me by my parents, Daye, and Elwood helped me to survive many rainy (and the few sunny) days in ol' Ory-gun. This research was supported by the National Oceanic and Atmospheric Administration grant NA-87-ABH-00018. TABLE OF CONTENTS

Page INTRODUCTION...... 1 The Blanco and its basins...... 2 Cascadia Segment...... 2

PREVIOUSWORK...... 7 Regional studies of the Blanco transform fault zone...... 7 Area-specific studies within the Blanco transform fault zone 10

DATA ANALYSIS AND METHODS...... 15 Seismic-reflection data...... 15 SeaBeam data processing...... 17

STRUCTURE AND SEDIMENTATION OF THE CENTRAL BLANCO TRANSFORM FAULT ZONE BETWEEN 129°20W AND 128°00'W: CASCADIASEGMENT...... 18 Statistical study of Cascadia Segment turbidite sequences...... 18 Fault angles and orientations in Cascadia Segment: comparison with models of slow-to-intermediate spreading mid-ocean ridges...... 20 Tilt measurements of turbidite reflectors about central axis...... 24 Summary of the tectonic evolution of Cascadia Segment...... 30

SMALL-SCALE STRUCTURE WITHIN A SEDIMENTED EXTENSIONAL BASIN: CASCADIA DEPRESSION...... 33 Isopach map of Cascadia Depression sediment...... 33 Cross-sectional profiles across Cascadia Depression...... 36 Classification of acoustic facies using 3.5-kHz reflection records 38 Fault orientations within Cascadia Depression...... 41 Estimated spreading rate of Cascadia Depression over the past 12,000years ...... 47 Summary of the tectonic evolution of Cascadia Depression...... 52 CONCLUSIONS AND PROPOSED FUTHER STUDY...... 54 Models of crustal accretion at spreading centers: steady state versus non-steady state rift valley formation...... 55 Is there hydrothermal activity within Cascadia Depression'...... 59 How crustal accretion may be occurring at Cascadia Depression 60 Futurestudies...... 65

BIBLIOGRAPHY...... 67 LIST OF FIGURES

Figure Page 1. Location map of Blanco transform fault zone within the northeast PacificOcean...... 3 2. SeaBeam map of Cascadia Segment...... 4 3. Seismic-reflection profile from Cascadia Segment...... 5 4. Embley et al.'s (in press) model of possible sequence of events over the past 5 Ma outlining the development of the Blanco transform fault zone...... 8 5. Cores 6609-19 and 6609-20 (from Duncan, 1968)...... 11 6. NOAA watergun seismic record of Cascadia Depression...... 14 7. Plot of ships' tracks within Cascadia Segment showing seismic- reflection data navigation lines...... 16 8. Depth (100 m intervals) of reflectors versus tilt...... 21 9. Fault dips and morphotectonic terminology as applied to Cascadia Segment...... 22 10. Morphotectonic terminology and plot of tilt values in Cascadia Segment assuming turbidites originally horizontal...... 25 11. Macdonald&Atwater's (1978) models of rift valley-to-rift mountain transition...... 29 12. Single-channel seismic-reflection profile showing relative motion alongfaults...... 31 13. Isopach of minimum sediment thickness within Cascadia Depression 34 14. Plot of ships' tracks within Cascadia Depression showing seismic- reflection data navigation lines...... 35 15. Isometric view of seven topographic cross-sections within Cascadia Depression...... 37 16. Examples of four 3.5 kHz echo-character types used in this study from profiles within Cascadia Depression.(based on Damuth, 1975 and 1980; Damuth et al., 1983)...... 39 17. Distribution of four 3.5 kHz echo-character types within Cascadia Depression (based on Damuth, 1975 and 1980; Damuth et al., 1983)...... 40 18. Fault orientations based on SeaMARC II side-scan sonar data, 3.5-kHz reflection data and bathymetry within Cascadia Depression... 42 19. Seismic-reflection records showing approximate fault dips from the northwest corner of Cascadia Depression...... 44 20. SeaMARC II side-scan sonar image of northern Cascadia Depression. 45 21. Unmigrated single-channel airgun seismic records showing some of the fault dips used in calculation of extension within Cascadia Depression ...... 50 22. Map view of faults within Cascadia Depression and total spreading ratesindicated...... 51 23. Mechanisms for the formation of axial-valley topography at slow- spreading ridges (from Deffeyes, 1970)...... 56 24. Models of steady state and non-steady state rift valley formation and the expected distribution of sulfides within one stage of rift valley development...... 58 25. Seismic-reflection profiles from three sedimented spreading centers: Guaymas Basin, Escanaba Trough, and Middle Valley 61 26. Seismic-reflection profile showing drag-folds at margin of volcanic complex at center of Cascadia Depression...... 64 LIST OF TABLES

Table Page Statistical analysis of tilt measurements from eastern blocks 1 & 2... 19 Vertical and horizontal components of displacement rate from surface offsets...... 48 STRUCTURE AND TECTONICS OF CASCADIA SEGMENT, CENTRAL BLANCO TRANSFORM FAULT ZONE

INTRODUCTION

Geophysical data, such as SeaBeam bathymetry and SeaMARC side-scansonar, has been used to understand the complicated structure and morphology of the transformzones of the world's oceans. One common feature within transformzones is the extensional, or pull-apart basin, resulting from divergence along strike-slip faults (Fox & Gab, 1984). On the other hand, Cascadia Segment, a morphologically and structurally complexarea within the Blanco transform fault zone (Plate 1), has evolved as the result ofa long history of active seafloor spreading. This adds another dimension of complexity to whatwe know about the structure of oceanic transform fault zones. The existence ofvery short spreading segment such as Cascadia Segment within a long transform faultzone may force us to re-think the idea that stationary deep-seated processes must control where seafloor spreading occurs. The tectonic evolution of Cascadia Segment is an important factor in deciphering the geologic history of the southern Juan de Fuca and northern Gorda Ridges. Thisarea of the northeast Pacific Ocean has been affected by about 12 millionyears of ridge propagation and transform reorientation. The existence ofa short seafloor spreading segment within the Blanco transform fault zone may help explain the complicated magnetic anomaly patterns found at this area of the ocean floor. The depressed position of the Cascadia Segment has caused thisarea to act as a sediment trap for bottom transported material moving through Cascadia Channel (Duncan, 1968; Griggs & KuIm, 1970). The existence of thick turbiditesequences atop tectonic blocks within Cascadia Segment makes this an ideal place for seismic-reflection imaging. Seismic-reflection data from this area reveals the structure and tectonics associated with seafloor spreading: a difficult task on sediment-free ridges. Cascadia Depression, the largest rhomboidal basin within the Blanco transform fault zone and centrally located Within Cascadia Segment, was proposed to be a site of active seafloor spreading (Embley et aL, 1987). Cascadia Depression is analogousto sedimented spreading centers found on mid-ocean ridge crests. Sedimented spreading centers are considered possible analogues of some ancient sediment-hosted ore deposits. The first section of this paper addresses the overall structure of Cascadia Segment in light of models of rifting processes. The second section featuresan analysis of the 2 detailed bathymetry, structure and tectonics of the current site of active spreading, Cascadia Depression.

The Blanco Fracture Zone and its basins The Blanco Fracture Zone is a 350-km-long transform zone that connects the southern Juan de Fuca and northern Gorda Ridges and includes a fossil transform segmentto the southeast (Figure 1). The active part of the fracture zone, the Blanco transform faultzone, is a series of right-lateral, right-stepping strike-slip faults separated by extensional basins. SeaBeam bathymetry reveals the morphological complexity of each of the basins (Plate 1). From west to east, these basins are identified as the Blanco Trough (or West Blanco Depression), East Blanco Depression, Surveyor Depression, Cascadia Depression and Gorda Depression. Two of these basins, Surveyor Depression and Gorda Depression, have rhomboidal shapes. Rhomboidal basins are also found on continents and result from divergence along strike-slip faults (Aydin & Nur, 1982). The other basins within the Blanco transform fault zone are morphologically more complicated. The Blanco Trough is an elongate "hole" more than 2400 m deep that constitutes the westernmost end of the active transform zone. East Blanco Depression has an overall rectangular shape and is structurally more complex than simple rhomboidal pull-apart basins (Palmer et al., 1987; Embley et al., in press).

Cascadia Segment The structure and tectonics of Cascadia Segment, located in the central part of the Blanco transform fault zone, is the focus of this study (Figure 2). The deepestpart of the basin (greater than 3000 m) has a rhomboidal shape. This deeparea, hitherto referred to as Cascadia Depression, measures 17 km long and 15 km wide. The structure of the Cascadia Segment is complicated by back-tilted blocks that flank the long (N-S) axis of the central basin and extend 50 km beyond it toward the east and west (Figure 3). Incross- section, the Cascadia Segment with its deep central axis and back-tilted flanking blocks, looks very much like a slow spreading mid-ocean ridge axis. The back-tilted blocksare in some cases covered with thick turbidite sequences whose dominant source is the North American Continent (Griggs & Kuim, 1973). Most of these sedimentswere transported by turoidity currents via Cascadia Channel (Plate 1) and other smaller channelsystems (Figure 2: Embley et al., in press; Griggs & Kuim, 1970). The structure of portions of Cascadia Segment was initially summarized by Griggs & Kuim (1973) and Ibach (1981). In this study, a more detailed examination of the structural development of the Cascadia Segment is conducted using seismic-reflection, 3 132° 1300 1280 1260 124° Canada i5 \% Washington cr;i : '.. c - / Oregon - 42° MENDOCINO F. J) 1 -_132° 130° 128° 126° -. 124°California 40° Figure 1: Location map of Blanco transform fault zone within the northeast Pacific Ocean. I?970°00 -I29OOO -129'OOOO -I26OOO -12e10OO. -1263000 -IB2OOO -1281000 128'OO 00 -IIi. 00, 00 -iI*4i-jc------jJi IMLII.

Figure 2: SeaBeam map of Cascadia Segment. Contour interval is 100 m. SeaBeam data collected by NOAA/NOS within the Blanco transform fault zone area in 1981, 1983 and 1984. Line AA refers to cross-section profiles seen in Figures 9 and 10. Dredge tracks are designated by dotted lines. Hydrocast locations are shown as wavy patterns. Locations of cores 6609-19 and 6609-20 are shown as stars at longitudes -12840' and -12829', respectively. +700 gammas -700 gammas Figure 3: Seismic-reflection profile from Cascadia Segment. Vertical exaggeration is werepressurecenteredapproximately used signalat together 200 15: andHz. 1. in aA Anvelocityorder dual-channel electromagnetic to distinguishsignal hydrophone (vertical betweensound geophone). source bottom was These used two with was used that measures both a returns and water signals frequency Segment.thesurface profile returns V.E.). Magneticjust east (Lister Note of anomaly the et ea.sternmostal.,buried 1987).profile back-tilted The is tectonic shown block blockbelow. on the course of Cascadja Channel dissects westernmost flank of Cascadia(directly above the letters Ui SeaMARC IT side-scan sonar, and SeaBeam bathymetric data. On the regional scale, these data demonstrate that the morphology of Cascadia Segment has developed through seafloor spreading. The acoustic penetrability of the turbidites on the flanking ridges of the depression provides a means of calcuLting the amount of tilt these blocks have undergone during the rifting process. Relative motion along faults is determined in the Cascadia Segment, a formidable task at mid-ocean ridges. East-west morphologic asymmetry is evident at Cascadia Segment and may be a result of asymmetric seafloor spreading. The topographic development of Cascadia Segment is compared to models of rift valley evolution and crustal generation processes at mid-ocean ridges. This, in turn, allows examination of current models of topographic evolution at spreading centers (i.e., Macdonald & Atwater, 1978; Davis & Lister, 1977; Deffeyes, 1970). Cascadia Depression is the prese.n locus of spreading within Cascadia Segment. High-resolution 3.5-kHz reflection records and SeaMARC H side-scan sonar data indicate that turbidite sediments, covering the Cascadia Depression floor,are cut by faults. Acoustic basement doming at the center of the basin is revealed in a map of sediment thickness distribution within Cascaclia Depression. A quantitative estimate of spreading rate is made for Cascadia Depression using a method described by Davis & Lister (1977). The relationship between turbidite accumulation and tectonicmovement within Cascadia Depression as it rifts apart is explored and compared to other sedimented spreadingcenters (Guaymas Basin, Gulf of California; Escanaba Trough,southern Gorda Ridge; and Middle Valley, northern Endeavour Ridge). PREVIOUS WORK

Regional Studies of the Blanco transform faultzone Using seafloor magnetic anomaly patterns and the time scale of Heirtzleret al. (1968), Atwater (1970) proposed that the Blanco transform formed approximately10 to 15 Ma. The magnetic anomaly patterns of the southern Juai de Fuca Ridgeare complicated by rift propagation along the ridge system Hey et al., 1980; Wilsonet aI.,1984). The Blanco transform fault zone has extended progressively westwardover the past 12 million years (m.y.) by annexation of propagating offsets (Hey et al., 1980). Wilsonet al. (1984) proposed that the western end of the Blanco transform faultwas established as the southward-directed rift propagation on the ceased about 5 Ma. Shortly after this, in response to a major shift in the Juan de Fuca/ Pacific pole of rotation, northward-directed rift propagation eliminated a short ridge segment located north of the first pseudofault (Wilson et aL, 1984). A more detailed model of Embleyet al. (in press), postulates that the Cascadia Segment is a short remnant of the ridgesegment eliminated in Wilson et al.'s (1984) model (Figures 4A, 4B). Embleyet al. (1987, in press) speculate that Cascadia Segment has been actively spreadingat a full-rate of 6 cm/yr over the past 5 million years. This is consistent with GLORIA side-scansonar data that shows a fault trace, coincident with the southern boundary fault of Cascadia Segment, continuousto at least I 30°30'W (Figure 4E; EEZ Scan 84 Scientific Staff, 1986). A single-channel seismic-reflection record reveals a 14.0 m high faultscarp coincident with the GLORIA side-scan fault trace at 130 lOW longitude. Yet another episode ofsouthward-directed rift propagation along the Juan de Fuca is proposed to have occurredat about 1.5 Ma (Figures 4C, 4D; Embley et al., in press). At this time, the fossil transform locatedat the western end of the Blanco transform fault zone was reactivated. Normal spreadingat the southern Juan de Fuca Ridge, after cessation of propagation, has progressively shiftedthe pseudofaults to the east producing crust with typical ridge-parallelstructures and magnetic anomaly patterns between the ridge axis and pseudofaulttrace (Figure 4E). The Blanco transform fault zone is seismically active (Tobin & Sykes, 1968;Bolt et al., 1968; Chandra, 1974; Couch et al., 1978; Rodgers, 1979), but because ofpoor azimuthal distribution of seismograph stations and poorly constrainedcrustal velocity models, the epicenters determined from P-phase data have routinelyshown offsets of 20- 30 km northward from the topographic expression of the fracturezone (Northrup, 1970). The use of T-phase data greatly improves this fit (Dunnebier & Johnson,1967; Northrup, 1970). A local seismicity study using sea-surface sonobuoyarrays recorded earthquakes Figure 4: Embley et al.'s (in press) model of possiblesequence of events over the past 5 Ma outlining the development of the Blanco transform faultzone. Diagrams based on unpublished magnetic modelling by D. Wilson,SeaBeam bathymetrv and GLORIA side-scan sonar data (EEZ Scan 84 Scientific Staff,1986). A major shift in the Juan de Fuca / Pacific pole of rotationoccurs between 4A and 4B, causing a reorganization of the transform. The CascadiaSegment began at this time as the result of an incompletecapture of the segment by the propagator. A propagator impinged on the inactive extension of the transform fault between diagrams 4C and 4D, reactivating it (about 1.5 Ma). Initiationof small extensional basin at the eastern end of the transform (presently theGorda Depression) may have occurred as the result of the "death' ofa northward- propagating rift before it reached the Blanco Fracture Zone(1-2 Ma ?). The model suggests that the Surveyor and Blanco Depressionsare the youngest of the Blanco extensional basins and have formed bya minor change in the relative motion between the Juan de Fuca and Pacific plates. 6 N N N) N N N NN NN NN N I --- NN N N N/ - N I I I / / N NN N N NN N N -4\ I I N \NN s1, /\\ /NN' - - - N NN N N N /I I I I JI N' N N N N N NN IJi,f $ffrj -41 Ii: ' / N N /NKN / I4 I' / 10

along the northwesternmost Blanco Ridge and within the Cascadia Segment (Plate 1; Johnson & Jones, 1978). During the deployment, a rnicroearthquake swarm of fourteen events occurred within the Cascadia Depression. These events were interpreted as extensional movements consistent with the seafloor spreading. The gravity field over the Blanco transform fault zone is characterized by a negative free-air gravity anomaly (Dehlinger et al., 1970) with the largest anomaly (-30 mgal) located over the Blanco Trough. Melson (1969) notes a strong positive (greater than 1,400 gammas) magnetic anomaly over the Blanco Trough. This magnetic anomaly probably reflects the presence of highly magnetized (Fe-Ti) basalts commonly found at propagating rift tips (Vogt & Bycrly, 1976).

Area-specific studies within the Blanco transform fault zone The steep northern scarp of Blanco Trough has been dredged and imaged in detail using deep-towed sonr (Delaney et al., 1987). Rock types include hyalloclastites, pillow basalts, diabase, greenstone, tectonic and hydrothermal breccia. Recent hydrothermal activity is noted by SeIk (1978) and M. Leinen (pers. comm.) in the Blanco Trough. Embley et al. (in press; Figure 4E) and Palmer et al. (1987) have proposed that the locus of transform motion near the Southern Juan de Fuca ridge-Blanco transform intersection has recently jumped to the north, opening the Blanco Trough. They interpret the East Blanco and Surveyor depressions as zones of extension that served to "transfer strike-slip motion to the main body of the Blanco Fracture Zone" as the Blanco Trough opened up (Figure 4; Plate 1). The model of Embley et al. (in press; Figure 4E) depicts the Blanco Trough, East Blanco and Surveyor Depressions as young (less than 1 Ma) extensional basins that formed as a result of a minor change in the relative motion between the Juan de Fuca and Pacific plates. The magnetic anomaly over Cascadia Segment is of low amplitude and flattened in nature (Figure 3; Lister et al., 1987). Three mechanisms proposed to account for very low amplitude seafloor magnetic anomalies are hydrothermal leaching of iron-titanium oxides (Levi & Riddihough, 1986), low specific remanence intensity of sills / dikesas compared to extrusive basalts (Irving, 1970; Vogt et al., 1970), and crustal temperatures beneath thick sediment that exceed the Curie points of unoxidized titanomagnetites (Johnson & Atwater, 1977; Johnson & Flail. 1978; Marshall, 1978; Prévot et al., 1979). Core data documents significant vertical tectonism within Cascadia Segment (Duncan, 1968). Piston core 6609-19, from Cascadia Depression, contains turbidite-derived sediments with Mt. Mazama volcanic ash (6,600y. in age) at the base (Figures 2 and 5). Duncan (1968) concluded that the floor of the basin has subsided 120m because Mt. 11

Figure 5: Cores 6609-19 and 6609-20 (from Duncan, 1968). (A): Core 6609-19 is over 8m in length and contains Mt. Mazama volcanic ash (6,600 y. in age) and turbidite-derived sediments(Duncan, 1968). (B): Core 6609-20, was takenat a depth of 2600 m, from a fault-block ridge located east of the Cascadia Depression. It shows a progressive change fromsand turbidite deposits at the base, gradually reducing in numberand thickness upsection, to totally pelagic, foraminiferal sedimentnear the top. Studies of turbidity current flows in thisarea (Duncan, 1968; Griggs & Kuim 1973; Kuim & Fowler, 1974) indicate that sandwas not deposited higher than 250 m above the Cascadia plain or the channel. The estimated upliftof this fault-block ridge is 450 m at a rate of 2 cm/yr (Duncan, 1968). .1 COnC S(D'C'I1 T"S 6609-20 0 -- 14,850!200 '(OP C14 Date A6609-19 a : 1 --- 16,450!750 '(OP C04t SICOINV tPt - -o- 21,950!950 'I'BP 3.3 B (North) B' (South) -- 4 -. 3-5 CISC4O/4 - 1I 4.0 yE. 35 F 0 L 2000.. I :__ ...' --- .. Figure 5 4': .- 0 ,. ,., 13

Mazama ash is present in turbidite sequences bothup and down channel from Cascadia Depression. Core 6609-20, from a fault-block east of Cascadia Depression, showsa progressive change from sand turbidite deposits at its base, gradually reducing in number and thickness upsection, to totally pelagic, foraminiferal sedimentnear the top (Figures 2 and 5). Duncan (1968) estimated the uplift of this fault-block ridgeto be 450 m at a rate of 2 cm/yr. Hydrocasts taken in Cascadia Depression show maxima with respectto methane between depths of 2600 m-3500 m (Embley et al., 1987; Figure 2). Methane anomalies may be hyarothermal exhalations or stem from the heating of organic matter within the turbidite sequences by volcanic sills intruded beneath the basin. Seismic reflection records reveal what may be volcanic sills near the center of the depression (Figure 6). Dredge samples obtained from CascaLlia Depression include mudstone, dolostone and siltstone (Figure 2). Local mechanisms, such as mass-wasting / backcuttingon the south wall of Cascadia Channel triggered by earthquake activity along the Blanco Ridge and turbiditycurrents generated from slumping on the north wall of the Blanco Ridge,are proposed to account for the distinct charmelized terrain located north of the Blanco Ridge (Plate 1; Embley, 1985). The downdropping of Cascadia Depression's floormay be a conthbuting factor to the mass wasting along this section of Cascadia Channel. According to the model of Embley et al (in press; Figure 4D), theeasternmost extensional basin in the Blanco transform fault zone, the Gorda Depression, initiatedas the result of the cessation of a northward-propagating rift before it reached the BlancoFracture Zone (1-2 Ma ?). Hydrothermal activity is noted in the Gorda Depression byHart et al.(1986). Dredge hauls consist of hydrothermally altered basalts, greenschist gabbros and greenschist breccias each representing "distinct stages of hydrothermalmineralization in the evolution of seawater into mature hydrothermal solution and then cooling of these solutions as they exit from " (Hart et al., 1986). 14

.1

.. ... -.. Bk

II BIocki 1

.: ': '.

: ..;!:--..: ;j...

-5km

Figure 6: NOAA watergun seismic record of CascadiaDepression. Vertical exaggeration= 9.7. Subcrustal reflectors, locatednear the center of the depression, may be volcanic sills (Embley et al., 1987). The designations"Block 1" and "Block 2" refer to the statistical study described in Chapter One. 15

DATA ANALYSIS AND METHODS

Seismic-reflection data Seismic-reflection data from Cascadia Segment used in this studycomes from the following sources: (1) Single channel airgun and 3.5-kHz seismic-reflectionrecords taken aboard the RJV S .P. Lee in 1985 by William Normark of the USGeological Survey; (2) Watergun single-channel and 3.5-kHz seismic-reflectiondata collected by Robert Embley in 1984 on the NOAA Ship Surveyor; (3) One 3.5-kHz reflection profile taken within CascadiaDepression on the NOAA Ship Discoverer in August, 1987; and (4) A seismic-reflection profile taken in July, 1987aboard the RJV Melville (Figure 3). Over 140 reflectors were digitized from seismic-reflectionrecords in order to quantify sediment thickness and the tilt of seismic layers. Aninterval velocity of 1.8 km/sec was used (Kuim & von Huene et aL, 1973; Ibach, 1981)and the velocity-depth gradient within these sediments was assumed to bezero. This assumption is within the 95% confidence level for deep-sea sediments shallower than about 1080m (Carison et al., 1986). The calculated tilt of very deep acoustic layers will betoo low because of the onset of the velocity-depth gradient at about 1080m depth. The overall variation in maximum sediment thickness atop the fault blocks at Cascadia Segment isfrom less than 300 m to greater than 1100 m. The largest calculatederror estimate within the Cascadia Segment tilt data is only 2.8°. Individual turbidite reflectors on the back-tilted blockswere selected by eye at an interval of 1-2 reflectors every 100m of depth. The digitized data was stored on a VAX 750 / microVAX 2 operating systemas two-way travel time (depth) versus absolute time. The reflection records were merged with navigationtracks that were likewise keyed to absolute time (Figure 7). The digitized subsurface acousticreflectors were placed into a three-dimensional grid: X (latitude), Y (longitude) and Z(depth as two-way travel time). Cross-sectional profiles (vertical exaggeration= 10:1) generated by a DISSPLA graphics software program were assembledintoa three-dimensional fence diagram model (after Davis & Lister, 1977). The positioning of cross-sectionsrelative to one another was aided by correlation with SeaBeam bathymetry. Using thismethod, records from the various sources at differing scales were reduced to one data base. -t29' 20 00 -129'OO'OO' -12e.sOOo' -126'IOOO' -128'30OO' -1282000 -12810OO' -12800,00. L?\ L°

U' A

S

-y z-

Figure 7: Plot of ships' tracks within Cascadia Segment showing seismic-reflectiondata navigation lines. NOAA Ship Surveyor tracks are solid lines. RJV S.P. Lee tracks are dashed lines. RJV Melville track isa dotted line. 17

SeaBeam bathymetric data vvas a valuable resource in resolving the structure of Blanco transform fault zone. The SeaBeam data base used in this study consists of smoothed ship navigation merged with multi-beam echo sonar data. Usinga VAX 7501 microVAX 2 operating system and a RADIAN CPS-1 ("Contour Plotting System") software program, SeaBeam data files were converted into evenly spaced gridded data sets at various scales. The area of the seafloor ensonified by SeaBeam multi-beamsonar corresponds to a rectangular area subtending angles of 600 by 2 213°. The standard swath width ensonified by SeaBeam sonar is equivalent to 75% of water depth. Within the Blanco transform fault zone, water depth is highly variable: ranging from less than 1600 m to greater than 4800 m (Plate 1). Thus, evenly spaced trackline coverage during the initial SeaBeamsurvey of the Blanco transform fault zone in 1983 left intermittent gaps between adjacent swaths. A field operation designed to fill in some of the more severe gaps in SeaBeamsonar data was carried out on the NOAA Ship Discoverer in 1987. Subsequently, in orderto finalize the bathymetry, it was necessary to edit existing SeaBeam sonar data and append the data collected in 1987 to the earlier survey. This was done usinga computer program that allows blocks of sonar / navigation data to be edited based on absolute time. The final result was a synthesis of SeaBeam bathymetric data from varioussmveys in the Blanco transform fault zone over a five-year period. These datawere used to generate a multitude of graphicalplotssuch as bathymetric maps, cross-sectional profiles and isometric view plots.These various plots proved to be a integral factor in the interpretation ofstructure and morphology of Cascadia Segment. 18

STRUCTURE AND SEDIMENTATION OF THE CENTRAL BLANCO TRANSFORM FAULT ZONE BETWEEN 129°20W AND 128O(YW: CASCADIA SEGMENT

By anaigy with sedimented spreading centers, suchas the Escanaba Trough (Atwater & Mudie, 1973), the occurrence of back-tilted blocks covered with turbiditesequences that "fan out" away from a central depression suggests that the Cascadia Segmentwas formed by seafloor spreading. It was shown by Atwater & Mudie (1973) that turbidite layers perched on back-tilted blocks flanking the Gorda Rise (located north of the Escanaba Trough) were deposited as a more continuous, nearly horizontal body at the time that the underlying crust occupied the valley floor, during and shortly after its formation. Using magnetic intensity and seismic-reflection data they showed that the turbidite-derived sediments occupying the central rifting basin were uplifted and tilted along with the underlying crust as it moved away from the spreading center. Therefore, basedon the assumption that perched turbidite sequences at Cascadia Segment originated withina central rifting basin, seismic-reflection profiles and SeaBeam bathymetrywere used to examine the structural development of the Cascadia Segment. The lack of unconformablesequences within the back-tilted turbidites (Figure 3) indicates that the tectonic blocks of Cascadia Segment have been uplifted in a continuous manner. Duncan (1968) has estimated therate of uplift on one block located east of Cascadia Depression to be 2 cm/yr (Figure 5).

Astatistical analysis of tilt versus depth was conducted on the first two back-tilted blocks located east of Cascadia Depression (Figure 6). These blockswere chosen for the following reasons: (1) the high density of data lines provides good arealcoverage of the first two eastern blocks (Figure 7); (2) one tracklinecrosses the trend of other data over these two blocks (Figure 7); (3) seismic-reflection profiles show better resolution of deep acoustic layers on less tilted blocks (in this case, those blocks located closerto Cascaina Depression). A sampling interval of 100 m was used because of the spacing of reflectors in the original seismic-reflection records (Table 1). Shallow turbiclite layers are easier to image acoustically than deep turbidite layers. The distribution of samples is skewed towards the shallow layers and hence the sample distribution curve is not "normal" (see colunm "N" of Table 1). The statistics presented in Table 1 are still valid and enlightening, however, because trends between acoustic tilt and depth are evident within both block "1" and block "2" and between thetwo blocks. In 19

TABLE 1

STATISTICAL ANALYSIS OF TILT MEASUREMENTSFROM EASTERN BLOCKS 1&2 (assuming an interval velocity of 1.8 km/sec)

BLOCK 1 BLOCK 2

Depth of N X S Range Depth of N X S Range reflector(number(mean) (standardof tilts reflector(number(mean)(standardof tilts (m) of samples) deviation) () (m)of samples) deviation) () 0-100 9 1.590.37 1.08-2.26 0-100 3 3.620.593.07-4.24

100-200 7 2.550.85 1.87-3.96 100-200 6 5.020.973.87-6.19

200-300 5 3.640.682.83-4.38 200-300 4 4.340.693.76-5.33

300-400 5 4.540844.05-6.03 300-400 5 7.01 1.205.75-8.85 400-500 3 5.150.434.78-5.62 400-600 4 7.681.286.57-9.52 500-600 3 5.700.834.80-6.44 > 600 3 6.75 0.266.47-6.99 >600 1 11.02

All Block 1 Data All Block 2 Data number=35 number=23 mean=3.60 mean=5.87 52=3.25 S2=4.21 SS 110.56 SS=92.55

All Data F-TEST = 1.294 T-TEST = 2.337 20

general, the mean tilt values increase with increasing depth intervals (Figure 8). Ranges of tilt values, based on data gathered from different locations on each block (Figure 7) using various types of seismic-reflection equipment, are shown in Table 1. Despite the variation in data sources, the scatter of tilt values within the 100 m sampling intervals is low Two statistical tests are used to compare the data in eastern blocks "1" and "2". Application of the "F-test" shows at a geater than 95% confidence level that the variances of the two data sets are equal (F=l.291; Fcrit=2.1). Results of the two-tailed "T-test" indicate that there is a significant difference (greater than 95% confidence level) between the mean values of blocks "1" and "2" (T=2.337; Tcrit=2.Ol). Together, these findings support the observation that the tilt of individual blocks increases with distance away from Cascadia Depression, the central spreading axis. Also, the equivalent variance of tilt values in "Block 1" and "Block 2" is consistent with the idea that both were formed by thesame controlling process.

Fault angles and orientations in Cascadia Segment: comparison with models of slow- to-intermediate spreading mid-ocean ridges A classification scheme, derived from terminology developed by Macdonald (1986 and 1982) to describe the FAMOUS and TAG sites on the Mid-Atlantic Ridge, is used here to compare the morphology of Cascadia Segment with mid-ocean ridge spreading centers. The classification of geomorphologic features found commonly at seafloor spreading centers is based upon: (1) fault throws; (2) dip of faults (to or away from the spreading axis); (3) direction of relative motion on the faults (normal or reverse). The descriptive terminology as it is applied to the Cascadia Segment is shown in Figure 9. Note that the thick sediment cover on top of the bounding blocks tends to mask the basement topography of Cascadia Segment. Although the surface expression of tectonism at the Cascadia Segment is not perfecdy analogous to less sedimented ridges, the same general processes probably control both types of spreadingareas. Therefore, the relationship between faulting, back-tilting and crustal extension in Cascadia Segment lend insight into how ridge-crest topography is created at other sites. The "inner floor" is the relatively flat and deep central zone where crustal accretion and crustal extension is maximum. The depth of the inner floor of Cascadia Segment varies between 300 m and 560 m below the level of the flanldng inner walls. Corrected for sediment thickness within Cascadia Depression, the basement of the inner floorranges between 500 m and as much as 1060 m below the level of bounding faultscarps. According to the scheme of Macdonald (1982), this rift valley depth puts Cascadia Segment into the slow-to-intermediate-spreading category. The proposed spreading half-rate of 21

DEPTH (m)

C C C C C C -C C C C C C Lr I C I I C C C C CI CI C C C C C C - Lt

2

C/) C) 4 C) C) C) 6

8

10

iL1- 0 BLOCK 1 + =BLOCK2

Figure 8: Depth (100m intervals) of reflectorsversus tilt. A A1 3 31 25 24 17' °

50 40 30' 20' 10' 6 10' 20' 30 40 50km

W. rift Western Inner E. rift mountains Eastern o terrace floor terrace mountains .E Lii Lu V.E.=5

Figure 9: Fault dips and morphotectonic terminology as applied to Cascadia Segment. Vertical exaggeration is 5: 1. This classification scheme it based on morphology used to describe the FAMOUS and TAG sites on the Mid-Atlantic Ridge (Macdonald. 1986 and 1982). The terms used here include: inner floor, inner walls, terraces, outer walls and rift mountains. The inner floor is the relatively

flat and deep central zone where crustal accretion/crustal extension is maximum. Tlie'inner walls are the faults that bound the inner floor. The 'terraces correspond to the topography created through block rotation and dip-slip motion as extension occurs away from the central axis. The "outer walls" are defined as the transition boundary between the relatively rugged terraces and the more topographically subdued rift mountains.

k) 23

approximately 3 cm/yr clearly qualifies Cascadia Segment as an intermediate rate spreading center (Embley et aL, in press; this paper). However, Macdonald (1986) has shown that for spreading rates over 1.7 cnl/yr, a rift valley greater than 300 m deepcan occur within 20 km of a triple junction or transform fault. Therefore, the "anomalously" deep rift valley at Cascadia Segment does not contradict its calculated intermediate spreading rate. The"inner walls" are the block faults that bound the inner floor. The slopes of the inner walls at Cascadia Segment are 26° to the east and 20° to the west: gentler than those reported from the slow-to-intermediate spreading Mid-Atlantic Ridge (Macdonald & Atwater, 197S) and the slow-spreading Cayman Trough (CAYTROUGH, 1979). These mid-ocean ridge escarpments typically have slopes of 50° to vertical. There is a six degree disparity between east and west inner wall slope values. In addition, the western inner wall is wider than the eastern inner wall (western inner wall= 4.7 km; eastern inner wall= 2.1 km). This may be the result of the rotation of tectonic elements and a resulting decrease in slope of the western inner wall toward the northwest corner of Cascadia Depression. The curved trend of Cascadia Depression's western wall proximal to the northern bounding transform fault is apparent in SeaBeam bathymetry (Figure 2). Fox & Gab (1982) attribute rotation of tectonic elements at ridge-transform intersections to the transmission of shear stresses associated with the transform fault into the spreading center domain. The term" terraces" corresponds to the rugged topography created through block rotation and dip-slip motion as extension occurs off the central axis. The angles of inward- dipping faults in the Cascadia Segment terraces range from 1 i° to 31°: generally shallower than those reported from slow-to-intermediate-rate spreading centers. Faultscan be traced at depth through the turbidite sediments. Therefore, the values of fault angles shown in Figure 9 are true fault dips and not merely calculated from exposed scarp measurements. The tops of the outer flanking escarpments (greater than 20 km from the central axis)are less steep than the faults that formed them. Seismic-reflection records show that the sediments on top of these blocks have been subjected to slumping suggesting thatmass wasting processes have tended to decrease the local escarpment slopes within the Cascadia Segment terraces. East-west topographic asymmetly is evident in the terraces. The extension in the Cascadia Segment has resulted in a 5.7 km difference between the length of east (23.6 km) and west (29.3 km) terraces. The expression of extension differson either side of Cascadia Depression. The western terrace is relatively smooth in contrast to the jagged morphology of the eastern terrace. Extension of the eastern terrace has been accommodated by rotation of three large, discrete blocks. The western terrace has extended through block 24 rotation and a large number of normal faults with smaller throws. Morphological asymmetry between opposing flanks of the Mid-Atlantic Ridge has been attributed to short-term (on the order of 106 years) asymmetric seafloor spreading (Macdonald,1986). Asymmetric spreading for crust 0-5 m.y. in the Atlantic has been reported at 45N (Loncarevic & Parker, 1971), 26°N (McGregor & Rona, 1975) and 6S (van Andel & Heath, 1970). Assuming a opening half-rate of 3 cm/yr at Cascadia Segment, five million years of asymmetric spreading could create topographicasymmetry crust up to 150 km off-axis. Morphologic asymmetry and its relationship to seafloor spreading rate variesas reported from the Mid-Atlantic Ridge. Eberhart (1987) concludes that faster spreading towards one flank at TAG sites 1 through 3 has resulted ina wider terrace in that direction. Conversely, at TAG, narrower terraces coincide with slower spreading. Thesame relationship between spreading rate and terrace width is evident at latitude 36°50'Non the Mid-Atlantic Ridge ("Narrowgate" site; Macdonald, 1977). Using similar reasoning, faster short-term spreading may have produced the wider and gentler slope of the Cascadia Segment's western terrace whereas the relatively narrow, rugged eastern flankmay be the result of slower spreading. On the other hand, the disparity between the widths ofeastern and western terraces at latitude 36°30'N on the Mid-Atlantic Ridge ("FAMOUS") doesnot correlate with asymmetric spreading rate (Ramberg & van Andel, 1977). The "outer walls" are defmed as the transition boundary between the relatively rugged terraces and the more topographically subdued rift mountains. At the Cascadia Segment, the outer walls mark the transition between inward-dipping normal faults of theterraces and the flattened topography of the rift mountains. One obvious mechanism of muting the topography within the rift mountains in the Cascadia Segment is theoccurrence of normal faults that dip away from the central axis of extension (Figure 9). The outward-dipping low-angle faults locally correlate with pre-existing block boundaries (i.e. the former "tops" of rifted blocks).

Tilt measurements of turbidite reflectors about central axis The topographic development of the Cascadia Segment has been, in general,an asymmetric process. Figure 10 shows the variation in amount of turbidite tilt within individual fault blocks. A model of topographic evolution east andwest of Cascadia Segment can be developed through examination of turbidite and underlying basement tilt versus distance from the central axis. Tilt values in the western inner floor and inner wall (less than 10.0 km off-axis)vary from -4.2° to 6.1° at Cascadia Segment:. A similar range of tilting (from -3.7°to 6.8°) was 25

Figure 10: Morphotectonic terminology and plot oftilt values in Cascadia Segment assuming turbidites originally horizontal. Shown abovethe cross-sectional profile is a plot of turbidite layer tiltversus distance from the axis of spreading beyond about 3 km of the central spreading axis. Each datapoint marks the position closest to the central axis (X0 km) where that tilted layer is observed on seismic-reflection records. Positive and negative valueson the vertical axis represent those layers that tilt away from and toward the central axis, respectively. "Hidden" tilts are those turbidite tilt valuesthat are not seen away from the central axis The envelopes of "hidden"tilts are shaded. H.

S .... is. It. '.. I

' :.,. i1' ' : I 4.1 S I . I A .1 S . I

:

'

50 40 30 20 10 10 20 30 4050km

W. rift Western Inner Eastern E. rift mountaifls terrace floor terrace o mountaiflS Ui

L 27

measured within eastern inner floor and inner wall, less than 7.0 kmaway from Cascadia Depression's central axis. The high degree of tilt seen in inner wall and inner floor turbidites is consistent with a model in which turbidite flows are deposited withina basin being uplifted and spreading away from a narrow extensionalzone. The turbidite layers are deposited during and shortly after the accretion of the underlying crust andare subsequently uplifted and tilted as the crust moves away from the rift valley floor. There are two trends of tilt within the eastern and western terraces: (1) the envelope of positive tilt values; and (2) the negative tilts that are found at the bases of fault blocks (Figure 10). The positive tilts are bound by distinct upper and lower trends. The maximum tilt values correspond to the deepest acoustic layers and therefore best define the tilt of the underlying basement. The data indicates that, in general, the basement of the terraces increases its tilt away from the central axis in both the east and west directions. The highest tilt achieved within the western half of the inner wall and inner floor (6.1°) is slightly lower than that seen in the eastern innerwalland inner floor sequence (6.8°). In the western terrace the basement rotates approximately 6° within 10 km of the central axis. The basement rotates from 6.8° to 12.8° Within the entire easternterrace, or 6° over 23 km. Fault blocks on slow-spreading ridges are typically tilted 5°to 15° away from the spreading axis (Atwater & Mudie, 1973; Macdonald & Luyendyk,1977). Macdonald (1982) reports that some major blocks such as those bounding theouter walls of the rift valley may be back-tilted in excess of 30°. Tilts of 3° to 5° and occasionallyup to 10° appear to be typical of intermediate-to-fast-spreading crust (Klitgord & Mudie, 1974). The increase in minimum positive tilt values out to at least 50 km from the axiswithin both the east and west terraces supports the contention that the blocks undergo additional rotation away from the axis during the riflingprocess. According to a model proposed by Davis & Lister (1977) the blocks will 'freeze in "aftera state of isostatic equilibrium is reached, wherein the crust and lithosphere are strong enough tosupport any remaining non-isostatic load. Off-axis extension through fault-block rotation and normal faultingare two modifications to the "frozen in" topography within the terraces of Cascadia Segment. "Hidden" tilts are those turbidite tilt values thatare seen away from the central axis (Figure 10, shaded region). The east-west difference in the envelope of "hidden" tiltvalues demonstrates the asymmetry of block rotation in theterraces. The east-west asymmetry of the "hidden" tilt zones is another expression of the asymmetrical extensionat Cascadia Depression. Negative tilt values in the terrace areas are turbidite sediments that have been tilted toward the central axis owing to their juxtaposition against active bounding faults(drag folds). The sense of folds confirms that they abut active normal faults. 4.

Data from Cascadia Segment is examined and compared with three models proposed by Macdonald & Atwater (1978) to account for the transition from the rugged riftvalley and terrace regime into the relatively smooth rift mountainson the Mid-Atlantic Ridge. The models, shown in Figure 11, include: (a) back-tilting of rift valley andterrace walls to approximately horizontal within the terraces; (b)reverse faulting along pre-existing inward- dipping faults in the rift mountains; and(C)movement along outward-dipping normal faults in the rift mountains. According to Macdonald & Atwater (1978), deep-towstudies of the Mid-Atlantic Ridge indicate that the primaiy mechanism for transformation ofthe rift valley topography is normal faulting along fault planes that dipaway from the valley axis (Figure 1 1C). They find the other mechanisms, i.e. regional tilt (Figure 1 1A) andreverse faulting (Figure 1 1B), to be less important in the rift mountains of the Mid-Atlantic Ridge. Figure 1 1A depicts a model in which the rift valley-to-rift mountain transitionat the FAMOUS site (37°N) on the Mid-Atlantic Ridge is accomplished bya modest rotation (5°- 9°) of the entire rift valley half-section as it passes into the rift mountains (Figure 1 1A; Macdonald & Atwater, 1978; Macdonald & Luydendyk, 1977). This quantitativeestimate of regional back-tilting is consistent with data from Cascadia Segment thatdemonstrates a rotation of at least 6° occurs in the terraces. However,at Cascaclia Segment, the regional slope from the rift valley center to the top of the outer walls (1 '-2°),as determined by SeaBeam bathymetry, is significantly lower than that of the Mid-Atlantic Ridgeat 37°N (5°- 9°). The disparity in regional slopes between Cascadia Segment and the Mid-Atlantic Ridge, may suggest that the rotation of tectonic blocksat Cascadia Segment is not necessarily the dominant mechanism of rift valley-to-rift mountain transitionat this site. Figure 1 lB depicts rift valley relief as being subdued byreverse motion along pre- existing inward facing faults (Macdonald & Atwater, 1978). In this model,as new normal faults are created near the rift valley,the relict normal faults are collapsed by reverse motion at the outer wall and beyond into the rift mountains (Deffeyes,1970;Osmaston, 1971; Harrison, 1974). Evidence from the outer walls andoutermost terraces of Cascadia Segment may support the reverse faulting model. The possibility ofreverse faulting and reverse drag folding along the inward-dipping bounding fault of the eastern outer wall is consistent with the very high positive tilt (highest= 12.8°) of the acoustic layers adjacent to this fault (Figure 10). Tilted turbidite sequencesare not discernable within the eastern outer-wall block. However, the maximum tilt within thewestern outer wall, a value that correlates with the tilt of the underlying basement, is 9.8°. This value fails well belowthe maximum tilt trend established in the western terrace. Reverse faultingalong the inward- dipping bounding fault of the western outerwallblock could account for this low maximum tilt.

30

Macdonald & Atwater's (1978) third model, depicted in Figure 1 1C,proposes an overprinting of the nft valley topography by normal faults that dipaway from the valley axis (Macdonald & Luydendyk, 1977; Needham & Francheteau, 1974). Accordingto this model, these may be new fault planes or reactivated faults that previously had small offsets. These faults obtain much of their throw beyond the outer wall and thus accomplish the transformation of rift valley into rift mountains. Theouter walls / rift mountains of Cascadia Segment are cut by outward-dipping faults (Figure 9). Thesense of motion on these faults is determined in two ways: (1) drag-folding along outward- dipping normal faults has resulted in the negative and shallow positive values of turbidite tilt in the eastern rift mountains (figure 10); and (2) drag-folded turbidite-derived layers, deposited within a former course of Cascadia Channel, record motion along shallow outward-dipping normal faults in the eastern rift mountains (X=41 km, 44 km. Figure 10; and Figure 12). The seismic-reflection profile shown in Figure 12 reveals whatmay be a deeply buried block on the westernmost flank of Cascadia Segment. This blockmay have been dropped down along an outward-dipping fault. Turbidite tilt data and fault orientations demonstrate conclusively that the muting of topography from the rift valley to the rift mountains of Cascadia Segment is accomplished by: (1) movement along outward-dipping normal faults in the rift mountains, and (2) back- tilting of tectonic blocks within the terraces. The possibility ofreverse motion along pre- existing inward-dipping faults in the rift mountains is not refuted by tilt data from Cascadia Segment. These conclusions are in partial agreement with the observations from the Mid- Atlantic Ridge which indicate that the primary mechanism for transformation of the rift valley topography is normal faulting along fault planes that dipaway from the valley axis (Macdonald & Atwater, 1978). However, analysis of turbiditesequences atop tectonic blocks at Cascadia Segment indicates that regional tilt is another significant factor in the transition from rift valley-to-rift mountain topography.

The structure of Cascadia Segment is the result of long-term, active seafloor spreading. The tectonic evolution of Cascadia Segment is revealed by examination of tilted turbidite sequences lying atop the back-tilted blocks. The variation in tilt values with distance away from the Cascadia Segment's central spreading axis demonstrates the following: (1) At Cascadia Segment, fault escarpment slopes are generally shallower than those reported from the Mid-Atlantic Ridge (Macdonald & Atwater,1978) and Cayman Trough (CAYTROUGH, 1979);

32

(2) The high degree of tilt achieved within the inner floors / inner walls at Cascaclia Segment is consistent with a model in which turbidite flows are deposited in an actively rifting basin; (3) Off-axis extension through further fault-block rotation (at least 6°) and normal faulting are two modifications to the topography within the terraces of Cascadia Segment; (4) The asymmetric expression of extension between the eastern and western flanks of Cascadia Segment may be attributable to short-term (on the order of 106 years) asymmetric seafloor spreading. Extension of the eastern terrace has been accommodated by rotation of three large, discrete blocks whereas the western terrace has extended through block rotation and a large number of normal faults with smaller throws; (5) The muting of topography from the rift valley to the rift mountains of Cascadia Segment is accomplished by two mechanisms including back-tilting of tectonic blocks within the terraces and movement along outward-dipping normal faults in the rift mountains. The possibility of reverse motion along pre-existing inward-dipping faults in the rift mountains is not refuted by tilt data from Cascadia Segment. SMALL-SCALE STRUCTURE WITHIN A SEDIMENTED EXTENSIONAL BASIN: CASCADIA DEPRESSION

The gross morphology of Cascadia Segment has been examined and attributedto long- term seafloor spreading from a central axis. The present spreading axis within Cascadia Segment is Cascadia Depression, a small rhomboidal basin. The short-term(on the order of iø - lO years) vertical and horizontal tectonics within CascadiaDepression were investigated using single-channel and 3.5-kHz reflection, SeaMARC IIside-scan sonar, and piston core data. Also, because Cascadia Channel has beena source of turbidite sediment to Cascadia Depression, it is possible to calculate extension of the basinover the past 12,000 years using a method described by Davis & Lister (1977).

In order to better define the subsurfaceprocesses occurring at Cascadia Depression, a sediment isopach map has been constructed. Figure 13 shows the contoured minimum depth to acoustic basement using values from single-channel seismic-reflectionrecords taken within Cascadia Depression. The data consists often tracklines that are oriented semi-parallel to the short axis of the basin, two tracklines thatcross the northwest corner of the basin, and a singular south-to-north tracklineacross the central zone of extension within Cascadia Depression (Figure 14). The available data base is sufficientto contour the basin area with reasonable certainty. The greatest detail, however, is achieved in the centralpart of the basin where tracklinecoverage is densest. A relatively thin cover of sediment (less than 250 m) is found throughoutthe central and northeast sections of Cascadia Depression. The thin sedimentcover corresponds in a general way to a wedge-shaped topographicarea delineated on Figure 2 by the 3500 m depth contour. The 0" isopach is centeredon the "V"-shaped topographic high at 43°46.5'N / 128°41 .5W (Figure 13). This centralarea of high acoustic basement can be considered the present focus of extension in Cascadia Depression. The slope of acoustic basement away from the central domedarea within Cascadia Depression is calculated taking into account bathymetric depth and depthto basement. There is a great disparity in basement slope between theeastern and western halves of the basin. The gentler slope is toward the west (average= 4.2°) whereas the eastern slope of acoustic basement is relatively steep (average= 11.5°). This asymmetry of basement slope between the west and east flanks of Cascadia Depression isconsistent with data presented in the previous section. Asymmetry of basement slopeseems to be established in the -43° 50'

3;(/Q

D Lfl -43° 40' rj

Figure 13: Isopach of minimum sediment thickness within Cascadia Depression. Contour Interval = 100 m; Shading Interval = 200 m. 35

__) (ID cx: (N

Figure 14: Plot of ships tracks within Cascadia Depression showing seismic-reflection data navigation lines. NOAA Ship Surveyor tracks are solid lines. RJV S.P. Lee tracks are dashed lines. RJV Melville track is a dotted line. NOAA Ship Discoverer track is a dot-dashed line. 36 central rifting basin and is propagated outward onto the terraces until isostatic equilibrium is reached. Another notable feature of the isopach map is the great depth to acoustic basementat the northwest corner of Cascadia Depression (Figure 13). Thiscorner of Cascadia Depression may be analogous to the deep "nodal" basins foundat ridge-transform intersections (Fox & Gallo, 1984). The depth to acoustic basementat the southeastern corner is unavailable because of the lack of data there. More seismic-reflection profiles are needed from the southeast corner to test the idea that thecorners of Cascadia Depression are structurally analogous to ridge-transform intersections.

Cross-sectional profiles across Cascadia Depression In order to gain some perspective on the overall structure of Cascadia Depression,a series of NW-SE stacked profiles were generated from the SeaBeam data. Theseare shown in Figure 15 as viewed from the southeast corner of the basin. Themost noticeable topography on the eastern half of Cascadia Depression is the dome-like feature whose southernmost extent is evident in the third cross-sectional profile from the bottom (labelled "doming" on Figure 15). The linear trend of the doming is parallelto the long axis of the basin. The feature seems to "spread out" towards the north ina general "V" shape. This reflects the shape of the subcrustal "doming" seen on the isopachmap (Figure 15). The western half of Cascadia Depression is shallower and shows muchmore fine- scale topographic variation than the eastern half (Figure 15). The topography here consists of turbidite-derived sediments that have been faulted upward toward thewestern inner wail in a series of step-like faults. The faults have throws of 5m to 168 m. The variation in turbidite tilt within the largest (westernmost) sediment bench is -O.Tto 3.5°. The local slope of Cascadia Depression's floor is the result of interaction betweentwo components: rate of movement along faults within the basin and the rate at which turbidite- derived sediments can bury the faults. At present, relative uplift of the topographic highs (or relative sinking of topographic lows) in Cascadia Depression is occurringat a higher rate than sediments are being delivered to the basin. Core 6609-19, taken within the deeper part of the basin (Figure 2), contains only three turbidite sequences within the 86Ccm of section younger than 6,600y (Duncan, 1968). Thissuggests that the recent number of turbidite flows into Cascadia Depression has been small. The southeastern quadrantof Cascadia Depression has apparently acted asa local catchment basin for turbidites travelling down Cascadia Channel (Figure 15). These relatively thick turbidites (3sequences in 860 cm of section) are probably of limited areal extent. Furthermore, Duncan (1968) found Mt. Mazama ash in turbidite sequences both up and down channel from Cascadia Depression 37

V.E.10 South

Figure 15: Isometric view of seven topographic cross-sectionswithin Cascadia Depression. Spacing between the seven cross-sections is 2 km and theyare aligned perpendicular to the long axis of the basin. Vertical exaggeration= 10. 38

and concluded that the floor of the basin has subsided 120 m. Therefore, the floor of Cascad.ia Depression has dropped at an average rate of about 1.8 cm/yrover 6,600 years.

Classification of acoustic facies using 3.5-kHz reflection records In order to examine the recent tectonics of Cascadlia Depression, 3.5-kHz echo- character types have been categorized. Four 3.5-kHz echo-character typesare found within Cascadia Depression. These include: (1) sharp bottom echo with parallel sub-bottom reflectors; (2) irregular overlapping hyperbolic reflectors; (3) sharp bottom echo with parallel sub-bottoms that exhibit sediment thickening against faults (growth faulting); and (4) sharp flat bottom echo with undulating sub-bottom reflectors (after Damuth, 1975 and 1980; Damuth et al., 1983; See Figure 16). The distribution of these four echo-character types within Cascadia Depression is shown in Figure 17. The data base used in this categorization are the same as those used in preparation of the isopachmap (Figure 14). Based on the structural setting of Cascadlia Depression, the present base level ofa deep-sea channel, it is assumed that the echo types reflect various degrees of tectonic deformation of turbidites. Layered turbidite-derived sediment gives sharp bottom echoes with parallel sub-bottom reflectors (Figure 16, line AX). This echo-character type is prevalent in the shallower western half of Cascadia Depression where step-like faults cut turbidite sediment. Likewise, turbidite-derived sediment is found in the northwestcorner of the basin. The sediment in this corner is cut by steep faults that are oblique to the regional direction of spreading. Growth faulting is not evident in these turbiditesequences implying that turbidite sedimentation and active faulting are not penecontemporaneous in this half of the basin. Irregular overlapping hyperbolic reflectors are associated with morphology controlled by acoustic basement that either crops out or is buried by thin conformable sediments (Figure 16, line BB'). The hyperbolic reflections in the middle of Cascadia Depressionare centered on the topographically high area whicl:, in turn, coincides with acoustic basement doming. A submersible investigation of the domedarea in Cascadia Depression did not locate extensive outcrops; however, small, sediment-covered step-faults with subtle outcrops were present (R. Embley, pers. comrn). The 3.5-kHz reflection records from eastern Cascadia Depression reveal sharp bottom echoes with parallel sub-bottoms that exhibit sediment thickening against faults ("growth faulting"; Figure 16, line CC'). Growth faults are characteristically found wherever unconsolidated sediment is deposited in basins that are actively growing in breadth and depth. Sediment thickness is greater on the down-dropped side of the active fault. The 39

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Figure 16: Examples of four 3.5 kHz echo-charactertypes used in this study from profiles within Cascadia Depression (based on Damuth, 1975 and1980; Damuth etal.. 1983). Location of example records is shown in Figure17. Line AA' corresponds to layered turbidite-derived sedimentchat give sharp bottom echoes with parallel sub-bottom reflectors Line BB' corresponds to irregular overlapping hyperbolicreflectors associated with morphology controlled by acoustic basement that eithercrops out or is buried by thin conformable sediments Line CC' corresponds to sharp bottom echoes withparallel sub-bottoms that exhibit sediment thickening against faults ('growth faulting"). Line DD' corresponds to sharp flat bottom echoes withundulating sub-bottom reflectors. This type of surface reflection may be associated withbedforms from a turbulent flow regime, tectonic deformation of sediment,or sediment that has undergone differential compaction, or some combination of these. 43050 ( / J (\%2*%4\

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Figure 17: Distribution of four 3.5 kHz echo-character types within Cascadia Depression (based on Dantuth, 1975 and 1980; Darnuth et aL, 1983). Examples of echo- character types are shown in Figure 16. Lines AA', BB, CC', and DD' designate where examples of echo-characters shown in Figure 16 are located. 41

sediment thickness is consistently greater on the western side of the individualfaults. This indicates that the blocks on the east of these growth faultsare being uplifted relative to the western side. This agrees with the expected sense of motionon the eastern "half' of a basin that is being extended and upliftedaway from the central axis. The limited areal extent of surficial growth faults indicates that active faulting with concomitant sedimentation has been restricted to the eastern half of Cascadia Depressionin post-glacial times. The southeastern corner of Cascadia Depression shows sharp flatbottom echoes with undulating sub-bottom reflectors on 3.5-kHz records (Figure 16, lineDD'). This type of surface reflection may be associated with bedlorms froma turbulent flow regime, tectonic deformation of sediment, sediment that has undergone differential compaction,or perhaps some combination of these factors (W. Normark, pers. comm.; C. Lister,pers. comm.). The exact origin of this echo-character typeat Cascadia Depression is unclear.

Fault trends, determined using 3.5-kHz and single-channelseismic-reflection records, and SeaMARC II side-scan sonar data,are shown in Figure 18. The orientation of faults located on the eastern side of Cascadia Depression is aboutO30. Recall that these are the only faults within Cascadia Depression that exhibit surficial growthfaulting. This indicates that the eastern side of the neovolcaniczone is the only region within the basin where turbidite sedimentation is occurring simultaneously with active faultmovement. The long- term axis of spreading is slightly east of north, as reflected in the orientation of faultscarps in Cascadia Segment (Figure 2). Therefore, the faultson the eastern half of Cascadia Depression are about 30 degrees oblique to the axis of long-termspreading. The trend of the eastern faults probably reflects the short-term direction andmode of seafloor spreading within Cascadia Depression. Faults of this orientation couldstem from a relatively higher extension rate toward northern Cascadia Depression. A higherrate of opening toward the north is supported by several features within the basin: (1) the"spreading out" character of the bathymetry seen in cross-sectional profiles; (2) the shapeof the subcrustal doming (Figure 13); and (3) the triangular wedge of hyperbolic reflectorsseen in Figure 17. The dominant trend of faults in the regionwest of the extensional axis is slightly west of north (353°-357°). Unlike the faults located in theeastern half of Cascadia Depression, the western step-like faults are located above the levelof turbidite deposition within Cascadia Depression. The orientation of thewestern faults may reflect a relatively older extensional regime, subparallel to the regional strike of fault blockescarpments. The strike of the western faults"bends" gradually towardsthe northwest corner of 41A

Figure 18: Fault orientations based on SeaMARC II side-scan sonar data, 3.5-kHz reflection data and bathymetry within Cascadia Depression. The 3.5-kHz tracklines are those shown in Figure 14. The location of faults observed on the 3.5-kHz are shown by solid line segments (teeth point downward). Dotted lines show inferred fault trends based on 3.5-kHz records. Inferred fault trends between teeth are based on: (1) fault throw values; (2) measured fault dips from single-channel seismic records; and (3) interpretation of detailed bathymetry. The dot-dashed fault trends are taken from SeaMARC H side-scan sonar data and correlated with exact fault locations from seismic records. L.4 28 C.4 43

Cascadia Depression, The observed trend of faults in the northwestcorner ranges between 120°-145° (NW-SE), about 35- 60 degrees oblique to the regional axis of seafloor spreading. Seismic-reflection records indicate that the faultscut through the sediment at relatively steep angles (greater than 60°: Figure 19). The amount of throwon these faults is small and the associated horizontal extension is minor. The northern boundary of Cascadia Depressioncan be divided into distinct zones based on fault orientations, SeaMARC II side-scansonar, and SeaBeam bathymetry data. The northern boundary is divided into two areas: (1) between 128°44'W and 128°48'W longitude ("western north wa1l); and (2) between 128°39'W and 128°44'W ("easternnorth wall"). Faults do not "bend gradually" into the trend of the transform faultat the western north wall of Cascadia Depression. Rather, the faults truncate thewestern north wall at high angles (Figure 18). The western north wall, asseen in SeaMARC II side-scan sonar (Figure 20) and SeaBeam bathymetric data (Figure 18), is highly reflective,steep, and scalloped in nature. Degradation of the eastern north wall of Cascadia Depression is inferred from hyperbolic echoes seen on 3.5-kHz reflection records. The degeneration of the slope is evident in SeaMARC side-scan sonar dataas numerous en echelon fault scarps located at the base of the eastern north wall (Figures 18 and 20). The trend of these faults (1110 to 121°) is subparallelto both the orientation of the boundary itself (108°) and the regional transform fault azimuth (118°). The degradation of theeastern northwallhas resulted in a slope much gentler than that of its westerncounterpart. At ridge-transform intersections one flank of the ridge is accreted against the aseismic limb of a fracture zone without experiencing strike-slip related tectonism (Fox& Gallo, 1984). In the case of Cascadia Depression, the eastern "half' of the north wallshould experience no strike-slip motion. Therefore, the gentle slope anden echelon faulting along the eastern north wall may be the result of mass wasting alonga passive boundary. The western north wall of Cascadia Depression, on the other hand, should accommodate strike-slip motion. The steep, scalloped character of thewestern northwallsupports the idea that it is tectonically active. The active portion of the transform at the / Tamayo Transform intersection is a steep, sediment slope with scallopedscarp surfaces. The scarps are probably produced by gravitational erosion that is triggered by strike-slip faulting of underlying oceanic basement (Tamayo Tectonic Team, 1984). At CascadiaDepression, the spatial correlation between truncating faults and the scalloping of thewestern northwall may indicate that slope failure is triggered by movement along the truncating faults. However, the high reflectivity of the western north wallon SeaMARC II side-scan sonar data suggests that this surface is sediment-free. The available seismic-reflectiondata cannot ji ___i_,__ -- o 0 0 .... ------.;

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Figure 20: SeaMARC II side-scansonar image of northern Cascadia Depression. (A) : Original SeaMARC II side-scansonar image. (B): SeaMARC II side-scansonar record with faults delineated. The northern boundary wall of Cascadia Depression isdivided into two areas: (1) between 128°44'W and 128°48'W longitude ("westernnorth wall"); and (2) between 12839'W and l2844'W ("eastern northwall"). The western north wall is highly reflective, steep, and scallopedin nature. The degeneration of the eastern northwallis evident as numerousen echelon fault scarps. Figure 20 47

support or refute the idea that the northern wall of Cascadia Depression is free of sediment. Further investigation is needed to resolve the fine-scalestructure and tectonics of Cascadia Depression's northern wall. SeaMARC II side-scan sonar data shows that the faults located about 5 kmwest of Cascadia Depression have a trend of 118°. This is equivalentto the azimuth of ridges and troughs that define the Blanco transform fault zoneas a whole. Therefore, at a distance of 5 km west of Cascadia Depression the bearing of faults parallels the azimuth of the principal transform displacement zone or "PTDZ" (after Fox & Gab, 1984). The calculated age of crust at this distance from the spreading center is about 165,000y (half- spreading rate = 3 cm/yr).

The high density of seismic-reflection data across Cascadia Depression is the basis for a quantitative estimate of spreading rate using a method described by Davis & Lister (1977). Using fault throw values from 3.5-kHz reflection records, fault dip angles from seismic-reflection data, estimates of turbidite sedimentation, and geometric relationships, the amount of horizontal extension at the surfacecan be estimated. Turbidite activity is a strong function of the glacial-interglacial-postglacialstate of the terrigenous source in this type of channel-fed basin (Davis & Lister, 1977). Duncan (1968) has estimated that the rate of turbidite sediment accumulation in thisarea decreased rapidly about 12,000 years ago. Prior to l2,000y, high rates of turbidite sedimentation would have maintained a horizontal surface within Cascadia Depression, masking tectonic deformation at the surface. Recall that piston core data indicates that the base level of Cascadia Depression's floor was 120 m higher 6,600years ago (i.e. the deepest part of the basin was at approximately 3400 m depth). Assuming Duncan's (1968) estimatedaverage drop rate of the floor (1.8 cm/yr) can be extrapolated back 12,000years, then the "horizontal surface" which defined Cascadia Depression's floorat the end of the last glacial period would have been at a depth of about 3300m. Figure 18 shows that, with the exception of the major bounding faults (i.e., those faults with greater than 400m of throw), all faults within the central part of Cascadia Depression (south of 43°48'N latitude) are found at depths greater than 3300 m. Therefore, it was assumed for all calculations that the surface displacement along the faults within the centralpart of Cascadia Depression (with throws less than 400 m) is younger than 12,000years. Fault throw values are measuredfrom3.5-kHz records ("Surface offset" column of Table 2). Assuming the faults have been actively displacing the once-horizontal surfaceof the basin for 12,000 years, an average vertical displacementrate is calculated over this 4.07 TOTAL: ------0.20 12 0.04 5 3 0.18 13 0.04 5 3 0.23 20 0.08 10 3 0.24 28 0.13 15 3 0.79 23 0.33 40 3 0.18 29 0.10 12 3 0.54 25 0.25 30 3 0.32 20 0.12 14 3 1.39 40 1.17 140 3 5.42 TOTAL: ------0.18 13 0.04 5 2 1.18 12 0.25 30 2 0.34 20 0.13 15 2 0.43 21 0.17 20 2 0.51 30 0.29 35 2 0.69 21 0.27 32 2 0.14 25 0.07 8 2 0.17 19 0.06 7 2 0.16 20 0.06 7 2 1.62 40 1.36 163 2

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(cni/yr) (crn/yr) south=3) rate (°) rate (m) middie=2, displacement dip offset offset (north=1, Horizontal Fault Vertical Surface Trackline

OFFSETS SURFACE FROM RATE DISPLACEMENT OF COMPONENTS HORIZONTAL AND VERTICAL

2 TABLE period. Thus, the "vertical offset rate" values (Table 2) obtainedare simply the surface fault throws divided by 12,000years. Figure 21 shows the western half ofone of the three seismic-reflection profiles used in these calculations. Fault dip (Table 2)was determined directly from the profiles using an interval velocity value of 1.8 km/sec (Kuhn &von Huene et al., 1973) and a zero velocity- depth gradient because sediment thickness within CascadiaDepression is less than 800 m (Figure 13; Carison et al., 1986). Determination of fault dipusing umnigrated seismic- reflection profiles was made more difficult because of diffractionswhich disrupt the records (Figure 21). However, it was found that the faultscut through the entire section of sediment, are low in angle (1Y-40°), and donot curve with depth. The measured angles match well with the low regional fault dips (1 1°-31°; Figure 9).Using the simple geometric relationship between the estimated vertical offsetrate along each fault and the dip of the fault, the horizontal extension rate (averagedover the past 12,000 years) of each fault is calculated (last column of Table 2). The extension rates on the individual faults within each of thethree tracklines are added together and the total extension rate for each line isshown in Figure 22. The rates obtained, 4.1 cm/yr, 5.4 cm/yr. and 5.8 cm/yr, match extremelywell with the expected full spreadmg rate of approximately 6 cm/yr (Embleyet al., in press). The extension within Cascadia Depression has apparently been accommodatedalong shallow faults. Clearly, the horizontal extension rate is highly dependentupon the dip of the faults (i.e., in general, any increase in fault dip will serve to decrease the calculatedhorizontal rate of extension). However, at Cascadia Depression, it is found that fault dipvalues up to 40 degrees higher than those measured (Table 2), will produce horizontalextension rates well within an order of magnitude of the estimated spreadingrate of 6 cm/yr. A progressive decrease in opening rate from 5.8 cm/yrto 4.1 cm/yr is evident from north to south, respectively (Figure 22). This discrepancy inestimated opening rates is probably the result of: (1) "true" short-term asymmetricopening of Cascadia Depression from north-to-south; (2) the masking of surface offsetalong faults towards the southern part of Cascadia Depression by ponding of turbidites;or (3) a combination of (1) and (2). Short-term asymmetric spreading may be documentedat Cascadia Depression. Averaged over 12,000 years, Cascadia Depression could haveopened up at a higher rate toward the north of the basin. Asymmetrical spreadingmay account for many features in the basin including: the "V"-shaped trends in topography(Figure 15) and acoustic basement doming (Figure 13), and also the obliquity ofsome faults within the basin (Figure 18) to the long-term axis of spreading. The method used to estimate horizontal extensionrates within Cascadia Depression is 50

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Figure 22: Map view of faults within Cascadia Depression andtotal spreading rates indicated. Fault dips (top) were determined directly from seismic-reflection records. Horizontal extension rates (bottom) were estimated usingfault throw values from 3.5-kHz reflection records, fault dip angles fromseismic-reflection data, estimates of turbidite sedimentaiion,and geomethc relationships(Table 2). 52

contingent on the assumption that turbidite accumulation throughout the basin has been insignificant relative to the fault motion over the past 12,000 years. However, this assumption is probably not valid in the southeastern quadrant of Cascarlia Depression where turbidite sediment accumulation has occurred within the past 6,600years (Duncan, 1968). Therefore, the calculated honzontal extension rates at Cascadia Depression should be anomalously low in this area. The values shown in Figure 22 support the idea that sediment ponding has reduced the estimates of horizontal extension within the southeastern quadrant of Cascadia Depression. There is a disparity in the cumulative estimated horizontal extension rates between the eastern and western sides of Cascadia Depression (Figure 22). The western "group" of faults apparently accommodate a greater amount of extension on each of the three seismic- reflection tracklines than their eastern counterparts. This east-west difference in extension is, in part, a consequence of turbidite "masking" of fault offsets in the southeastern quadrant of Cascadia Depression. Recent motion along the eastern boundary fault of Cascadia Depression is suggested by the presence of growth structures at the base of the fault (Figure 16C). Unfortunately, the amount of extension attributable to motion along Cascadia Depression's eastern boundary fault cannot be calculated using the method described above. The additional component of extensionon the eastern boundary fault should compensate for some (if not all) of the asymmetry in calculated horizontal extension rates between the eastern and western sides of Cascadia Depression. Furthermore, any supplemental extension will serve to increase the overall calculated horizontal extension rates within Cascadia Depression.

Summary of the tectonic evolution of Cascadia Depression The relationship between horizontal extension, vertical tectonics, and sedimentation of Cascadia Depression is consistent with the idea that it is a center of seafloor spreading. The temporal and spatial evolution of Cascadlia Depression as itriftsapart is revealed by examination of acoustic basement morphology, determination of faultor itations and quantification of relative motion along faults. The following conclusionsare made: (1) Doming of the acoustic basement is occurring in the center of Cascadia Depression. Subcrustal mechanisms that could account for the shallow acoustic basement in the central basin are: (a) presence of sills; (b) uplift by volcanic doming alonga linear rift; or (c) diapiric intrusions such as serpentine; (2) Asymmetry of basement slope seems to be established in the central rifling basin and is propagated outward onto the terraces until isostatic equilibrium is reached; 53

(3) Extension within Cascadia Depression has apparently been accommodatedalong shallow faults. Quantitative estimates of extension within Cascadia Depression reveal spreading rates of 4.1 cm/yr, 5.4 cm/yr, and 5.8 cm/yr. The calculatedrates of extension support Embley et al.'s (in press) model in which CascadiaSegment is a remnant 'trapped" spreading segment froman episode of rift propagation; (4) Within Cascadia Depression, a progressive decrease in calculatedopening rate from 5.8 cm/yr to 4.1 cm/yr is evident from northto south, respectively. This may be the result of true short-term asymmetric opening of CascadiaDepression, or the masking of surface offset along faults towards the southernpart of Cascadia Depression by ponding of turbidites; (5) Short-term (on the order of iO years) asymmetric spreadingmay be occurring within Cascadia Depression. Asymmetric spreading could explain the"spreading out" character of bathymetry toward the northern depression, the irregularshape of the acoustic basement doming. the thangular wedge of hyperbolic reflectors inthe central depression seen in 3.5-kHz records, and the oblique orientation of growth faults on the eastern side of the depression; (6) The decrease in calculated opening rate towards the southmay be the result of sediment ponding in the southeastern quadrant of Cascadia Depression.Turbidite accumulation in this part of the basin has been insignificant over the past 12,000 years (Duncan, 1968), contraryto one assumption of the method used to calculate horizontal extension. Therefore, the calculated horizontalextension rates should be anomalously low in this area; (7) Recent motion along the eastern boundary fault of Cascadia Depressionis suggested by the presence of growth structures at the base of the fault (Figure 16C). This additional component of extension shouldcompensate for some (if not all) of the asymmetry in calculated horizontal extensionrates between the eastern and western sides of Cascadia Depression. Furthermore,any supplemental extension should increase the overall calculated horizontal extensionrates within Cascadia Depression. 54

CONCLUSIONS AD PROPOSED FURTHER STUDY

The detailed interpretation of the morphologic and structural information derived from SeaBeam bathymetric, single-channel reflection seismic and SeaMARC II side-scan sonar data outlined in the previous chapters has provided important clues to the tectonic evolution of this portion of the Blanco transform fault zone. A tectonic evolution of Cascadia Segment is outlined in a model in which there has been a juxtaposition of several processes occurring over the past several million years. In this model, a basin forms by seafloor spreading, is filled with sediment and subsequently the crust is rafted away from the central axis, through the terraces and out into the rift mountains. The predominance of conformable turbidite sequences on the uplifted tectonic blocks at Cascadia Segment indicates that uplift within and away from the central basin has occurred in a continuous manner. A progressive decrease in calculated opening rate over the past 12,000 years, from 5.8 cm/yr to 4.1 cm/yr, is evident from northern to central Cascadia Depression, respectively. This may be the result of true short-term asymmetric opening of Cascadia Depression, or the masking of surface offset along faults towards the southern part of Cascadia Depression by ponding of turbidites. The high degree of tilt achieved within the inner floors / inner walls of Cascadia Segment support the idea that turbidite flows are deposited into a basin that is spreading away from a central axis. Asymmetry of basement slope, established in the rifting basin, is propagated outward onto the terraces until isostatic equilibrium is reached. Off-axis extension through further fault-block rotation (at least 6°) and normal faulting are two mechanisms by which topographic relief is created within the terraces of Cascadia Segment. Asymmetric seafloor spreading on the order of 106 years may account for the asymmetric expression of extension between the eastern and western flanks of Cascadia Segment. Extension of the eastern terrace has been accommodated by rotation of three large, discrete blocks whereas the western terrace has extended through block rotation and a large number of normal faults with smaller throws. The muting of topography from the rift valley to the rift mountains of Cascadia Segment is accomplished by back-tilting of tectonic blocks within the terraces and movement along outward-dipping normal faults in the rift mountains. The possibiit of reverse motion along pre-existing inward-dipping faults in the rift mountains is not refuted by tilt data from Cascadia Segment. 55

Models of crustal accretion at spreading centers: steady state versus non-steady state rift valley formation In light of the conclusions made about the tectonic development of Cascadia Segment, further speculation about deep-seated processes occurring there can be addressed. The proposed evolution of Cascadia Segment is compared to existing models of rift valley evolution and crustal generation processes at mid-ocean ridges. This allows examination of current models of topographic evolution at spreading centers, especially in the context of steady state versus non-steady state (i.e. episodic) theories of rift valley development (Kappel & Ryan, 1986; Macdonald & Atwater, 1978; Davis & Lister, 1977; Deffeyes, 1970, for example). Atwater & Mudie (1968, 1973) reported that evidence from turbidite deposits and benthic foraminifera from cores taken at the Gorda Rise indicate that faulted blocks were uplifted along normal faults to form the rift valley rather than the valley floor being faulted down. This model requires the inner floor to "ride out" passively into the terraces along a series of normal faults in a steady-state manner. Several steady state models have been proposed to explain rift valley generation on slow-spreading ridges including: (1) hydraulic head loss and viscous drag in a narrow conduit of upward flowing asthenospheric material (Sleep, 1969; Lachenbruch, 1973); (2' extension and necking of the lithosphere near the axis (Tapponnier & Franceteau, 1978); (3) the disproportion in widths of the axial extension zone and axial volcanic zone (Deffeyes, 1970; Palmasson, 1973; Anderson & Noltimier, 1973); and (4) horizontal extensional stresses in a strong, brittle lithosphere that thickens with distance from the ridge axis, the width of the axial valley being controlled by plate thickness near the ridge axis (Morgan et al., 1987). The morphologic resemblance between Cascadia Segment and the northern Endeavour Ridge ("Middle Valley"), i. e. individual back-tilted blocks covered with turbidite sequences that "fan out" away from a central depression, suggests that these two regions are the result of similar deep-seated processes. In order to explain the evolution of individual back-tilted turbidite-covered blocks at Middle Valley, Davis & Lister (1977) modified Deffeyes (1970) steady-state model of ridge-crest topographic generation where the material supply occurs over a broaaer region than extension (Figure 23). Cascadia Depression's central zone of extension may be as wide as 6 km, if defined as where the shallow faults that cut through the sediment extend through to the underlying acoustic basement (i.e., about twice the width of the faulted basement shown in Figure 2'). Although Cascadia Depression's apparent zone of extension is wider than that proposed for Middle Valley (Figure 23C, curve 1), the similarity of topographic profiles between the two 1 SJRF.8CE 11r EXTENSIONAL

2 2 ATE

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Figure 23: Mechanisms for the formation of axial-valley topography at slow-spreading ridges (from Deffeyes, 1970). In each case, curves 1, 2, 3, and 4 represent surface extensional velocity, material supply rate, integrated supply rate, and the resultant topography, as functions of distance from the ridge crest axis. (A) : The general case with the width of the supply zone greater than the surface extension. (B) : A model in which the crustal extension zone is narrow as suggested by deep- tow magnetic data (Atwater & Mudie, 1973). (C) : This model is similar to 23B, but with a rigid crust superimposed. The formation, uplift, and outward rotation of normal-fault blocks from the axial valley are a consequence of the formation of a rigid crust at the origin. Once the blocks have reached a level of isostatic equilibrium, or once the crust and lithosphere are strong enough to support any remaining isostatic load, the blocks will "freeze in" as permanent features of the seafloor topography (from Davis & Lister, 1977). 57 areas suggests that the general relationship between extension and crustal accretion at Middle Valley also holds for Cascadia Depression. A discontinuous topographic profile with roughly equidimensional blocks is caused by the addition of rigid crust to a model in which the crustal extension zone is relatively narrow (Figure 23C). Rotation of blocks away from the axis of crustal accretion occurs because of the change of material supply rate with distance from the origin. Negative tilts, found within the inner floors / inner walls of both Middle Valley and Cascadia Segment, may be the result of a non-steady state factor such as the formation of a new axial valley where material supply is deficient (Figure 10; Davis & Lister, 1977). A non-steady state nature of Cascadia Segment's rift valley is also supported by the fact that the floor of Cascadia Depression has subsided 120 m in 6,600 years (Duncan, 1968). Kappel & Ryan (1986) have proposed a non-steady state model of rift valley development based on SeaMARC I side-scan sonar data from the Juan de Fuca Ridge system (Figure 24). Kappel & Ryan (1986) use the term "elongate summit depression" (ESD) rather than "rift valley" because the latter implies an exclusively tectonic origin for the axial feature. Models representing both steady state and non-steady state origins for ESD's are shown in Figure 24. Figure 24A depicts the ESD as a steady state feature that, through time and in a conveyor belt fashion, transports crust (slice X, for example) upward from the depression floor to the ridge flanks via faulting. Distinctly different is the episodic model of Kappel & Ryan (1986) shown in Figure 24B. In this non-steady state model, slice X, located at the summit of the crestal ridge at times Ti and T2, is collapsed downward to become part of the ESD (at time T3). The ESD widens by further extension during time T4; over the same relative time interval, slice Y moves away from the central axis and down the crestal ridge. During phase Ti and T2 the ridge axis is a neovolcanic province that may have a width many times broader than the axial extrusion zone (Figure 24B). Later, in phases T3 and T4, the axis becomes a dominantly tectonic province. Kappel & Ryan's (1986) non-steady state model is in agreement with an early working hypothesis of Normark (1976) that "volcanic extrusive activity at the rise crest and (or) periods of extension and faulting are episodic". Clearly, volcanic extrusive activity at Cascadia Segment is not as voluminous as that of spreading ridges hundreds of kilometers in length. However, the model of Kappel & Ryan (1986) lends support to the idea that the balance between extension and volcanism at Cascadia Segment may be episodic in nature. Morphological evidence, such as the extreme depth of the inner floor and the predominance of gentle fault dips in Cascadia Segment, indicates that extension is the more dominant factor. Furthermore, the high subsidence rate of the inner floor (1.8 cnilyr averaged over 6,660 years) and the high calculated iorizontal extension rate (possibly greater than 5.8 A B Ti Xr- T 1 x 12 T2 Aj T 3 T3 T4t T4xx

¶IsoclYon 5. Refefence Piece of Seof1oc

C Rift valley axis Elongate summit depression) Sulfides

2- ; --'-z

f400m 5Km Figure 24: Models of steady state and non-steady state rift valley formation and the expected distribution of suffides within one stage of rift valley development. (A) : The steady state model of rift valley topographic development (after Kappel& Ryan, 1986). In this model, the crestal ridge cross-section does not change significantly from Ti to T4, and slices X and Y follow the seafloor profile as they move away from the ridge-axis. Slice X is faulted up from the elongate summit depression (ESD) floor to become part of the flanks from T2 to T3, and then 'unfaulted' during T4. (B) : Kappel & Ryan's (1986) non-steady state model of rift valley evolution. A crestal ridge grows from Ti-to T2, the ESD collapses into the crestai ridge during T3, and there is incipient growth of a new crestal ridge during T4. There is not necessarily an equal amount of nine during each of these stages. Note that the whole flank of the crestal ridge moves away from the ridge axis intact as the ESD widens. (C) : This schematic diagram integrates Stages T2 and T3 of the non-steady state tectonic model (Figure 24B) with a model that predicts the disnibution of sulfide deposits (Kappel & Franidin, in press). During this stage, the ESD is wide and flcrured, and the 400'C isotherm is located within a thickened gabbro cap. The ESD bounding faults provide the pathway for the hotfluids to reach the seafloor. 59 cm/yr averaged over the past 12,000 years taking into account motion along the eastern boundary fault) may indicate that the magma supply has waned in the recent past allowing extension to become the current dominant mechanism at Cascadia Depression. If this is the case, the present configuration of Cascadia Segment fits best between stages T2 and T3 of the non-steady state model (Figure 24B). This scenario suggests that extension at Cascadia Segment is going on regardless of the upwelling magma supply. The presence of proximal, cold truncating transform boundaries may perturb the production of basaltic melt in the rising asthenospheric wedge beneath the rift valley at Cascadia Segment (P. Fox, pers. comm.; Fox & Gallo, 1982 and 1984).

Is there hydrothermal activity within Cascadia Depression? Most seafloor hydrothermal deposits have been identified using deep-towed cameras and submersibles while surveying mid-ocean ridge crests (i.e., RISE Project Group, 1980; Ballard et al., 1981, 1982; Normark et al., 1983). Kappel & Ryan's (1986; Figure 24B) non-steady state model of rift valley development provides the basis for explaining the distribution of hydrothermal circulation and massive sulfide deposits based on regional tectonic setting of northeast Pacific spreading centers (Kappel & Franidin, in press). Confined hydrothermal circulation, restricted by impermeable sediment, results in high upper crustal temperatures and intense sediment diagenesis, causing localized hydrothermal deposits that are much larger than those found in unsedimented spreading centers (Lonsdale & Becker, 1985; Holmes & Morton, 1986; Karlin & Lyle, 1986; Davis et al., 1987). The conditions necessary to host hydrothermal activity at seafloor spreading centers include (Kappel & Franidin, in press): (1) the availability of high temperature fluids (400°C; Bischoff et al., 1983; Seyfried & Janecky, 1985); (2) focussing of fluids into conduits; (3) transportation of fluids to the seafloor with a minimum of conductive cooling or mixing with cold seawater; and (4) in order to preserve sulfide deposits, some form of covering, such as sediment must be present (Kappel & Franidin, in press).

Recall that the present configuration of Cascadia Segmentfallsbest between stages T2 and T3 of Kappel & Ryan's (1986) non-steady state model (Figure 24B). At this stage of the non-steady state model (Figure 24C), volcanism has waned and the rift valley widens by tectonic stretching. The model preseated in Figure 24C depicts the 400°C isotherm to mirror the shape of a subcrustal magma chamber. In the case of Cascadia Segment and in lieu of a distinct magma chamber, the insulating quality of turbidite sediment contained within the rift valley basin (wherever greater than 300 m in thickness) is sufficient to contain the heat necessary to attain the required temperature for formation of metalliferous hydrothermal solution (after Kappel & Franidin, in press). The major bounding faults of the rift valley should provide a pathway to the surface for hot and buoyant fluids. An indirect indicator of hydrothermal activity at Cascadia Segment is the low amplitude, flattened nature of the magnetic anomaly signature (Figure 3). The absence of significant magnetic anomalies at sedimented spreading centers may be a positive indicator of a closed system of hydrothermal activity. Levi & Riddihough (1986) propose that the absence of magnetic anomalies at sedimented spreading centers is a result of hydrothermal leaching of iron-titanium oxides, a consequence of a long fluid residence time in the basalt. Longer residence time of the hydrothermal fluid should result in more complete alteration of basalt than is expected at relatively more "open" (i.e., non-sedimented) hydrothermal systems(Levi& Riddihough, 1986). Two other mechanisms have been put forward to explain the flattened magnetic anomalies at sedimented spreading centers. Irving (1970) and Vogt et al. (1970) propose that slower-cooling magma dikes / sills have larger grain sizes than their extrusive counterparts, and thus correspondingly lower specific remanence intensities and lower stabilities. The second model addresses the thermal conditions at the base of sediment column where temperatures in the basalts may exceed the Curie points of unoxidized titanomagnetites (Johnson & Atwater, 1977; Johnson & Hall, 1978; Marshall, 1978; Prévot et al., 1979). This situation would decrease the spontaneous magnetization of the underlying rocks and reduce observed magnetic anomalies.

How crustal accretion may be occurring within Cascadia Depression Sedimented spreading centers have become the focus of considerable interest as possible modern analogs of some ancient sediment-hosted ore deposits. The mechanism(s) by which crustal accretion is occurring within Cascadia Depression can be addressed by comparing its small-scale seismic-reflection characteristics to other sedimented spreading centers such as the Guaymas Basin within the Gulf of California, Middle Valley located at the northernmost end of Endeavour Ridge, Northeast Pacific and Escanaba Trough at the southern end of Gorda Ridge, Northeast Pacific. Sediment-hosted hydrothermal activity exists at each of these locales. Volcanism within the Guaymas Basin and Middle Valley occurs as intrusive sills (Gieskes et al., 1981). Escanaba Trough, on the other hand, has numerous volcanic domes and edifices that protrude through the sediment cover along the axial rift in a linear trend (Morton et al., 1987a, 1987b). The Guaymas Basin, located within the Gulf of California, has total spreading rate of 6cm/yr. Intrusion is by emplacement of thin sills and thicker bysmalith-like bodies, which episodically raise hills of folded and faulted semiconsolidated turbidites (Figure 25A, cartoon; Lonsdale & Lawyer, 1980; Lonsdale et al., 1980). Einsele et al. (1980) report that sill intrusions into highly porous sediments in the Guaymas Basin cause large-scale Figure 25: Seismic-reflection profiles from three sedimented spreading centers are shown. (A) : Cartoon (on left) suggesting mechanism of intrusion, faulting and folding, and morphogenesis in axial rift valley of Guaymas Basin (from Lonsdale & Lawyer, 1980). Note that intrusion of thin sills has little effect on relief, because they made space for themselves by expelling pore water from intruded sediments (Einsele et al., 1980). The seismic-reflection profile (on right) crosses near a hydrothermal site at the Guaymas Basin. Note confused diffracted echoes from the sill complex that underlies horizontal basin-floor strata (from Lonsdale et al., 1980). (B): Seismic-reflection profile near seafloor dome at the Escanaba Trough (from Morton et al., 1987b). This record shows upward bowing of the basement reflector and side-echoes associated with the volcanic edifice beneath the seafloor dome. Marginal faulting of the sediment fill isalsoshown. Pictured below is an interpretive cross-section showing schematic fluid circulation paths through the volcanic edifice and overlying sedimentary rocks (from Morton et

al.,1987b). (C) : Seismic-reflection profile from Middle Valley, northern Endeavour Ridge (from Davis et al., 1987). Small topographic mound (enclosed by box) is a site of active hydrothermal venting (J. M. Franklin, pers. comm.). The sediment thickness here is about 300 m and the mound itself is located about 2 km from the eastern bounding normal fault scarp. Both the airgun seismic- reflection profile and the 3.5-kHz profile (inset) show the mounds to be, in general, acoustically transparent, with occasional interbedded strong reflectors (Davis et al., 1987). 62

NW SE VOLCANIC IPINNACLE BASIN FLOOR HYOROTHRMAL S

------B

MASSIVE SULFIDE /

------_. C

SW

Figure 25 63

expulsion of heated pore fluids, thus creating space for the intruding magma which can, in turn, lead to initiation of a hydrothermal systems. The type of uplift depicted schematically in Figure 25A (left) should induce discernable drag-folding of turbidite sequences at the margins of the growing intrusions. Single-channel seismic-reflection data from the near the hydrothermal site show a confused pattern of diffracted echoes from the sill complex (Figure 25A, right). The structure of the slow-spreading (about 2.3cm/yr half-rate) Escanaba Trough has been described in detail by Morton et al. (1987a, 1987b; Figure 25B). Sulfide deposition at two sites in the Escanaba Trough is associated with volcanic edifices that have been faulted upward through the axial valley sediment. In order to account for the small amount of turbidite sequence deformation associated with the uplift of these volcanic edifices, Morton et al (1987b) concluded that "these edifices either predate much of the sedimentary fill or have grown more-or-less contemporaneously with sedimentation". At the margins of small domes located on the tops of the large edifices, Morton et al (1987b) note an abrupt change in 3.5-kHz echo-characteristics from gentle pelagic drape to a highly deformed and uplifted sequences. This is similar to the transition seen at the central domed area of Cascadia Depression from growth faulted turbidite sediment to hyperbolic echoes (Figures 16 and 17). The characteristic structure of Middle Valley, northern Endeavour Ridge, is shown by a seismic-reflection profile in Figure 25C (from Davis et al., 1987). This profile crosses a topographic mound which is a site of active hydrothermal venting (J.M. Franklin, pers. comm.). The sediment thickness here is about 300 m and the mound itself is located about 2 km from the eastern bounding normal fault scarp. Both the airgun seismic-reflection profile and the 3.5-kHz profile (Figure 25C) show the mounds to be, in general, acoustically transparent, with occasional interbedded strong reflectors. The character of the acoustic reflectors at the center of Cascadia Depression is revealed in Figure 26. The subsurface "doming" at Cascadia Depression is defined by highly reflective, chaotic hyperbolic returns on seismic-reflection records. This is in contrast to the acoustically transparent nature of seismic-reflection records taken at Middle Valley (Figure 25C). It is unclear, however, whether this "doming" is caused by shallow volcanic sills (i.e., similar to the Guaymas Basin), the uplift a volcanic hill (i.e., analogous to the Escanaba Trough), or perhaps, serpentine diapirism. Examination of seismic-reflection records at Cascadia Depression reveals small-scale drag folding in the turbidite sequences around the subsurface margins of the subcrustal high-reflectivity zones. Subsurface drag-folding at Cascadia Depression does not show increased deformation with depth thus indicating that diapiric intrusion is jjthe cause of .,H;;;II:;I::;,,.u,::,,,..

::.,:;: . H : I

I. i. I II I I - II

I I I. 3 II 11 II 2 I(' I

I

Figure 26: Seismic-reflection profile showing drag-folds at margin of volcanic (?) complex at center of Cascadia Depression. Vertical exaggeration is 3:1. 65 subsurface doming. The extent of the drag-folding in turbidite reflector sequences around the subsurface margins of the volcanic complexes at Cascadia Depression indicates that subsurface processes generally postdate the sedimentary fill (Figure 26). This is in contrast to the observations of Morton et al. (1987b) at the Escanaba Trough where a general lack of extensive drag-folding at the margins of volcanic sequences indicates that the volcanic edifices predate much of the sedimentary fill (Figure 25B, interpretive cross-section). The type and extent of drag-folding seen in seismic-reflection records (Figure 26) indicaies that the mechanism of crustal accretion at Cascadia Depression is probably by growth of volcanic sills, perhaps in a bysmalith-like fashion (i.e., similar to that depicted in schematically in Figure 25A, cartoon). Therefore, the mechanism(s) of crustal accretion at Cascadia Depression may be directly analogous to those processes occurring at the Guaymas Basin (Figure 25A, cartoon; Lonsdale & Lawyer, 1980). Lonsdale & Lawyer (1980) report that "detached" faults at the Guaymas Basin intersect volcanic sifis and serve as conduits for hydrothermal circulation (Figure 25A, cartoon). By analogy, the fault surfaces at Cascadia Depression that intersect the subsurface "doming" (i.e., faults shown in Figure 21) may serve to focus upwelling hydrothermal fluids from a volcanic sill complex (Lonsdale & Lawyer, 1980; Kappel & Franklin, in press; Figure 24C).

Future Studies This study has addressed the detailed structure and tectonics of Cascadia Segment, using a variety of methods and many types of geophysical data. It has been demonstrated that active seafloor spreading is the probable mechanism by which the morphology and structure of Cascadia Segment has evolved. In order to test this hypothesis directly, further exploration of the current site of active spreading, Cascadia Depression, is required. By analogy to other sedimented spreading centers, specific target areas for investigation within Cascadia Depression have been proposed. Investigative surveys of Cascadia Depression can be carried out using a number of research methods. A first-order field operation should include a detailed investigation of the structures within Cascadia Depression using deep-towed side-scan sonar (i.e.. SeaMARC I). These type of data should indicate the true character and orientation of the faults seen in seismic-reflection data, detect whether any extrusive basalts exist on the floor of the basin, and reveal the presence of high-density, hydrothermally-altered sediment (perhaps in the shape of of mounds) within this extensional basin. Detailed sonar records, coupled with the available SeaBeam bathymetry, wifi serve to delineate the structures that warrant further investigation. Of key interest is the detection of sediment-hosted hydrothermal activity and its associated sulfides withi Cascadia Depression. Water-sampling and salinity-temperature- density measurements should reveal the presence of hydrothermally-derived constituents within the water column such as radon, manganese, iron, and associated high concentrations of methane. Heat-flow measurements within the floor of Cascadia Depression should define the areas where hydrothermal activity may be occurring. Seafloor photographic surveys could outline the extent of hydrothermal activity through mapping of features such as talc deposits, metalliferous sediments, and high concentrations of hydrothermally-associated organic matter such as bacterial mats. Material obtained in piston cores and dredge hauls should show similar mineralogy to that found at known sediment-hosted hydrothermal areas. Mineralogical assemblages from sediment-hosted hydrothermal areas include pyrrhotite, sphalerite, pyrite, marcasite, barite, chalcopyrite, talc, and chlorite (i.e., Einsele et aL, 1980; Lonsdale et al., 1980; Lonsdale & Lawyer, 1980; Davis et al., 1987; Morton et al., 1987a, 1987b). On the regional scale, a first-order problem is to confirm that a progressive age increase of turbidite-derived sediment in fact exists away from the central depression. This can be done using a systematic program of dating core material from the individual back- tilted tectonic blocks that flank Cascadia Depression. The availability of reliable dating methods (i.e., carbon-14) and detailed SeaBeam bathymetric data makes this a worthwhile endeavor. The above discussion gives a general idea of the types of data needed to further our knowledge of this area. Cascadia Segment's position as a possible center of crustal accretion that has subject to long-term turbidite sedimentation provides an excellent opportunity to understand the structure of spreading centers in general. Furthermore, the existence of very short spreading segment within a long transform fault zone gives new insight as to where and how crustal accretion can occur. Deciphering the tectonics and structure of this small, enigmatic spreading segment will help explain the complicated geologic history of this part of the ocean basin and, perhaps more importantly, may tell us more about the structural and tectonic processes associated with seafloor spreading. 67

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