Downdip landward limit of Cascadia great earthquake ruptureweb.uvic.ca/~rhyndman/PublicationPDFs/HyndmanCascadia... · 2015. 2. 21. · Downdip landward limit of Cascadia great earthquake

Downdip landward limit of Cascadia great earthquake rupture

R. D. Hyndman1,2

Received 31 January 2013; revised 19 September 2013; accepted 29 September 2013; published 24 October 2013.

[1] This paper examines the constraints to the downdip landward limit of rupture for theCascadia great earthquakes off western North America. This limit is a primary control forground motion hazard at near-coastal cities. The studies also provide information on thephysical controls of subduction thrust rupture globally. The constraints are (1) “locked/transition” zones from geodetic deformation (GPS, repeated leveling, tide gauges); (2)rupture zone from paleoseismic coastal marsh subsidence, “paleogeodesy”; (3) temperatureon the thrust for the seismic-aseismic transition; (4) change in thrust seismic reflectioncharacter downdip from thin seismic to thick ductile; (5) fore-arc mantle corner aseismicserpentinite and talc overlying the thrust; (6) updip limit of episodic tremor and slip (ETS)slow slip; (7) rupture area associations with shelf-slope basins; (8) depth limit for smallevents on the thrust; and (9) landward limit of earthquakes on the Nootka transform faultzone. The most reliable constraints for the limit of large rupture displacement, >10m, aregenerally just offshore in agreement with thermal control for this hot subduction zone, butwell-offshore central Oregon and near the coast of northern Washington. The limit for 1–2mrupture that can still provide strong shaking is less well estimated 25–50 km farther landward.The fore-arc mantle corner and the updip extent of ETS slow slip are significantly landwardfrom the other constraints. Surprisingly, there is a downdip gap between the best otherestimates for the great earthquake rupture zone and the ETS slow slip. In this gap, plateconvergence may occur as continuous slow creep.

Citation: Hyndman, R. D. (2013), Downdip landward limit of Cascadia great earthquake rupture, J. Geophys. Res. SolidEarth, 118, 5530–5549, doi:10.1002/jgrb.50390.

1. Introduction

[2] The downdip landward limit of rupture for the Cascadiasubduction zone off western North America is a primary con-trol for the ground shaking hazard from great earthquakes atnear-coastal cities such as Vancouver, BC; Seattle, WA; andPortland, OR (Figure 1). This paper summarizes and analyzesthe constraints on this downdip limit. It focuses on Cascadia,but many of the limits are applicable to the seismogenic zonesof other subduction zones. It also deals principally with themain part of the 800 km long rupture zone between centralVancouver Island south of the Nootka Fault zone and southernOregon, which generates ~M9 earthquakes. The more com-plex areas to the north and south may be involved in the longM~ 9 ruptures but are also large enough to generate separateevents up to M8. The areas of the Explorer plate off northernVancouver Island and the Gorda region off Northern Californiaare subducting younger ocean lithosphere than the main

central region and may have a shallower downdip limitbecause of the higher inferred temperatures. The mainCascadia rupture appears to occur primarily inM9 events thatrupture much of if not the whole length of the margin basedmainly on the deep-sea record of earthquake-generatedturbidites and the coastal marsh coseismic subsidence [e.g.,Goldfinger et al., 2012; Leonard et al., 2010] and with a gen-erally consistent downdip limit [e.g., Leonard et al., 2010].This is in contrast to many other subduction zones that exhibita wide range of rupture areas and earthquake magnitudes. Themodel of Flück et al. [1997] for the thrust locked zone basedon geodetic data has been used as a reference for comparisonamong the various constraints to the downdip seismogeniclimit (Figure 2). As discussed below, thermal limits to downdipseismic behavior are similar.[3] Understanding the downdip limit provides importantinformation on what controls the areas of rupture on subduc-tion thrust faults globally and the nature of fault movementby seismic and aseismic motions. Also critically importantto this discussion are well-recorded recent M ~ 9 great earth-quakes that provide method calibration, i.e., Tohoku Japan2011, Chile 2010, and Sumatra 2004, although these greatearthquakes occurred where older colder ocean lithosphereis being subducted such that the seismogenic limit may notbe thermally controlled, in contrast to the inferred thermalcontrol for Cascadia. Relevant information comes fromgeological study of deeply exposed thrust faults which showthat subduction thrusts may be very complex shear zones on

1Pacific Geoscience Centre, Geological Survey of Canada, Sidney,British Columbia, Canada.

2School of Earth and Ocean Sciences, University of Victoria, Victoria,British Columbia, Canada.

Corresponding author: R. D. Hyndman, Pacific Geoscience Centre,Geological Survey of Canada, 9860 W. Saanich Rd., Sidney, BC V8L4B2Canada. ([email protected])

©2013. American Geophysical Union. All Rights Reserved.2169-9313/13/10.1002/jgrb.50390

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JOURNAL OF GEOPHYSICAL RESEARCH: SOLID EARTH, VOL. 118, 5530–5549, doi:10.1002/jgrb.50390, 2013

a geological outcrop scale with an active layer of underthrustsediment in a subduction melange [e.g., Bachmann et al.,2009; Fagereng, 2011]. Thrust structure complexity with athick duplex structure, mainly deeper downdip than theseismogenic zone for northern Cascadia has been proposedby Calvert et al. [2011] using multichannel seismic reflectiondata. Some of the constraints have been discussed in previousreviews of the seismogenic zone of subduction faults [e.g.,Tichelaar and Ruff, 1993; Pacheco et al., 1993; Oleskevichet al., 1999; Hyndman, 2007; Hyndman and Rogers, 2010;Heuret et al., 2011].[4] The updip limit of rupture toward the trench also isimportant for estimating earthquake magnitude and espe-cially for modeling tsunami generation. For many other sub-duction zones, it is often concluded that the updip rupturelimit does not reach the trench (e.g., summary by Hyndmanet al. [1997]). The 2011 M9 Tohoku rupture did reach thetrench with large displacement [e.g., Fujiwara et al., 2011]but at such a slow rupture speed that there was little energyin the seismic frequencies. However, the very large tsunamigeneration appears to have been generated mainly in theupdip region. Bilek and Lay [1999] argued that the rigidityis very low and therefore the rupture speed is very lowat the toe of accretionary sedimentary prisms, increasingdowndip perhaps because of clay consolidation. If the updipseismogenic limit is at the depth where stable-sliding claysdehydrate, i.e., 100–150°C [e.g., Moore and Saffer, 2001;Oleskevich et al., 1999], this limit may be near the defor-mation front for the high-temperature Cascadia subductionzone. However, laboratory data by Marone and Saffer [2007]suggest that other physical changes with depth must beresponsible for the updip part of the thrust commonly beingaseismic. The Cascadia updip limit is poorly constrained andis not discussed further in this paper.[5] The downdip seismogenic limit is where the fault zoneno longer exhibits frictional instability [e.g., Scholtz, 2002](Figure 1). Many factors have been suggested for the controlsof the downdip landward extent (e.g., discussions by Tichelaarand Ruff [1993], Pacheco et al. [1993], Oleskevich et al.[1999], and Hyndman [2007]), but two have been argued to

be most important. For hot subduction zones where younglithosphere is being underthrust like Cascadia, there appearsto be a maximum temperature limit for seismic behavior [e.g.,Hyndman and Wang, 1993; Oleskevich et al., 1999]. For themore common cool subduction zones, aseismic serpentiniteand talc downdip in the fore-arc mantle corner may limit thedowndip rupture [e.g., Ruff and Tichelaar, 1996; Hyndmanet al., 1997; Oleskevich et al., 1999; Peacock and Hyndman,1999; Hyndman and Peacock, 2003]. Ruff and Tichelaar[1996] noted that the downdip limit is often near the coast.Other factors that could affect the downdip rupture limit arevariations in the stress regime of the thrust fault zone, porepressure on the fault, fault roughness, and upper and lowerplate structure [e.g., Scholz, 1992; Pacheco et al., 1993;McCaffrey et al., 2008; Llenos and McGuire, 2007; Heuretet al., 2011; Bilek, 2007; Wallace et al., 2009; Wang andBilek, 2011].[6] The location of the seismic source and interpretation ofa number of the downdip constraints depend on the dip angle

Figure 1. Schematic of the downdip limit of Cascadia greatearthquake rupture, the closest approach to near-coastalcities. Episodic tremor and slip (ETS) is well landward of therupture limit and approximately corresponds to the fore-arcmantle corner.

Figure 2. Fully locked (to solid black line) and lineartransition (to dashed gray line) zones based on leveling and tidegauge geodetic data, used as a reference for comparison amongthe various landward seismic limit constraints (modified afterFlück et al. [1997]). The thermal limits of seismic behavior,350 and 450°C, from thermal models are shown with uncer-tainties, as given by Hyndman and Wang [1995]. The limitsfrom this thermal model and geodetic limits are in agreementwithin the uncertainties. ETS is well landward.

HYNDMAN: CASCADIA GREAT EARTHQUAKE RUPTURE LIMIT

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and profile of the subducting slab (Figure 3) [e.g., Flück et al.,1997; McCrory et al., 2012a]. The principal large-scalevariability is the region of shallow dip beneath the marginof northern Washington and southern Vancouver Island. Asdiscussed below, it appears that the downdip rupture limitmay be thermally controlled. Therefore, the region with shal-low dip is expected to reach a critical temperature fartherlandward and have a wider rupture zone than the steeper dip re-gions to the south and north. Most but not all past great eventsappear to have ruptured all of the main part of the Cascadiasubduction fault [e.g., Goldfinger, 2011; Goldfinger et al.,2012; Leonard et al., 2004; 2010] but there have been somethat ruptured shorter extents along the subduction fault. Thedistribution of event sizes and margin-parallel extents is thesubject of ongoing research [e.g., Atwater and Griggs, 2012;Frankel, 2011].[7] We discuss nine constraints to the landward limit ofCascadia great earthquake rupture. These nine constraintsprovide a better understanding of the landward seismogenicdowndip limit and the controlling processes than are avail-able for many other subduction zones, even those which havehad instrumentally recorded great events. Each of the nineconstraints requires different assumptions and has differentuncertainties; we seek a rupture limit that is consistent withall of the constraints. There is a considerable literature onaspects of some of these constraints. An extensive referencelist is given, but far from complete. Emphasis is on thosepapers that provide general results and those that havereference lists themselves that lead to papers with moredetailed information.[8] An important issue is the nature of the landward tap-ering of rupture displacement, i.e., the “transition zone”(Figure 3). This transition is important since 1–2 m ofdisplacement can produce strong ground motions in the maindamaging frequency range, as nowwell recognized, for exam-ple, the recent M9 Tohoku great earthquake off NE Japan[e.g., Kurahashi and Irikura, 2011; Koper et al., 2011] and

for the M9 2004 Sumatra and 2010 Chile earthquakes and anumber of other subduction zones [e.g., Lay et al., 2012]where much of the energy in the damaging frequencies wasfrom the downdip part of the rupture. For some constraints,it is useful to define the area of at least 50% or ~10m of theestimated ~20m average full rupture (see Figure 4). The posi-tion of 10% or 2m, which in the geodetic dislocationmodels isconcluded to be tens of kilometers farther landward, is muchless constrained than 50% because the former depends onthe nature and extent of the transition from full to zero rupturedisplacement. Also, the nature of the actual transition is likelyquite variable among different events. Geodetic locked zonerupture inversions usually give smooth peak rupture andlandward decrease.[9] The degree to which the downdip limit to rupture variesamong individual events remains uncertain. Most Cascadiaevents appear to rupture a reasonably consistent landwardextent based on paleoseismicity coastal marsh subsidencedata [Leonard et al., 2010]. Although there undoubtedly isvariability, the summarized estimates are taken to be themaximum landward rupture extent for most events, as to beexpected if there is a physical control of this limit that doesnot change with time.[10] There are three types of constraints: (a) constraints tothe currently “locked zone” on the thrust that is buildingelastic strain that will be released in future great earthquakes(constraint 1 below), (b) constraints to the landward extentof past great earthquake rupture (constraint 2 below), and (c)physical process/state controls on where seismicity can andcannot occur on the thrust (constraints 3–9 below). It is impor-tant to recognize that none of the constraints directly constrainthe actual rupture limit; all require some assumptions.[11] The constraints are:[12] 1. The “locked/transition” zones on the thrust frommodeling geodetic observations of current deformation(GPS, repeated leveling, tide gauges, etc.).[13] 2. Past great earthquake rupture zone constrained bypaleoseismic coastal marsh subsidence, 1700 and earlierevents, i.e., “paleogeodesy”.[14] 3. Seismic behavior limits from downdip temperatures,loosely described as the “brittle-ductile transition”, approxi-mately 350°C for where the rupture can initiate and there isfull rupture, 450°C for the landward limit of the transition zonefrom full rupture to zero displacement.

Figure 3. Schematic geometry of the Cascadia subductionthrust, showing the simple model of locked or “full rupture”and transition zones (adapted from Flück et al. [1997]). Theshallow thrust angle for northern Washington results in thetemperature limits being farther landward and wider locked/rupture and transition zones.

Figure 4. Locked and transition zone models correspondingapproximately to the full rupture, and linear and exponentialtapering rupture (effective transition zone; adapted fromWang et al. [2003]). The horizontal distance scale and rupturemagnitudes are approximately those for Vancouver Island andnorthern Oregon.


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[15] 4. Change in seismic reflection character from athrust marked by a thin sharp reflection to a thrust zonemarked by a thick set of reflectors, i.e., brittle to ductileshear deformation.[16] 5. Fore-arc mantle corner, landward of which there isinferred aseismic serpentinite and talc overlying the thrust.[17] 6. Updip limit of interseismic episodic tremor and slip(ETS) slow slip that accommodates most of the long-termplate convergence.[18] 7. Empirical association of rupture area with shelf-slopesedimentary basins defined mainly by gravity anomalies.[19] 8. Landward downdip limit of small thrust earthquakeson the subduction thrust that indicate seismic behavior.[20] 9. Landward limit of earthquakes on the Nootkatransform fault zone as it subducts beneath the margin ofVancouver Island.[21] Very general downdip rupture limits are providedby estimates of tsunami wave amplitudes and runups,although the updip extent and displacement are a moreimportant factor. Only the 1700 tsunami has been at allwell documented for Cascadia so this constraint is notdiscussed here. We have used the Flück et al. [1997]model with a linear downdip transition zone as a simplereference for comparison among the different con-straints. Other models with a linear transition zone givesimilar results [e.g., Mazzotti et al., 2007], as do mostinversion models.

2. Locked/Transition Zones From GeodeticDeformation

[22] The most studied and one of the strongest constraintsto the rupture zone is the definition of the locked zone onthe thrust fault through “back-slip” modeling [e.g., Savage,1983] of geodetic data (GPS, repeated leveling, tide gauges,etc.). This approach is based on the approximation that theearthquake cycle is elastic to a first order, i.e., elastic strainbuilds up nearly linearly with time since the previous eventand is released completely by megathrust earthquakes(Figure 5). The pattern of surface deformation for the ruptureis then close to the mirror image of the interseismic rate timesthe earthquake interval (average 500–600 years for the mainCascadia events). There are two significant limitations.First, there may not be full “coupling”; there may be someaseismic slip between great earthquakes (i.e., discussion byWang and Dixon [2004]). The current geodetic data that areconsistent with a fully locked fault [e.g., Hyndman andWang, 1995; McCaffrey et al., 2013] and the lack of signifi-cant interseismic thrust seismicity suggest that there is cur-rently little interseismic slip. The thrust is fully locked aftersome postseismic transients. Also, the estimated coseismicrupture displacement in past events approximately equals theinterseismic plate convergence slip deficit based on the magni-tude of coseismic coastal subsidence [Leonard et al., 2004;2010]. The one area of uncertainty is off central Oregon wherethe inferred locked/rupture zone is well offshore such that theexpected and observed interseismic uplift rate and coseismicsubsidence are small. From these two types of data, it is there-fore difficult to exclude some degree of interseismic aseismiccreep motion for the thrust in this region with the partiallyseismic zone extending farther landward [e.g., Pollitz et al.,2010; Wang, 2012]. This uncertainty is discussed below.[23] Second, because the Earth is viscoelastic, there aretheoretical problems of uncertain magnitude in the elasticbackslip model for the locked zone, especially postseismictransients [e.g., Hu et al., 2004; Wang 2007; Kanda andSimons, 2010; Hetland and Simons, 2010; Suito andFreymueller, 2009; Wang et al., 2012]. The downdip extentof fully seismic rupture is expected to be somewhat seawardof the limit for the modeled interseismic locked zone becausesome release of the locked zone strain buildup, especially thedeep part, is in postseismic slip [e.g., Wang, 2007; Hu andWang, 2012; Wang et al., 2012], so the limit from the geo-detic data may be slightly landward of the actual coseismicfull rupture limit. Records of local deformation as a functionof time for a few subduction zones that have been adequatelymonitored indicate that the long-term interseismic strainbuildup is approximately linear after a sometimes substantialpostseismic transient that may depend on the earthquakemagnitude [e.g., Wang, 2007]. However, this questionrequires more rigorous analysis and calibration by recent greatearthquakes elsewhere. An example from the M~ 8 1944 and1946 events off southwest Japan is shown in Figure 6 thatindicates good correspondence of the preseismic pattern ofdeformation with that for the coseismic rupture deformation.The downdip rupture extent of defined by dislocation model-ing of leveling data was very close to that for the lockedzone defined by prior geodetic data, within about ±10 km[Hyndman et al., 1995; see also Sagiya and Thatcher, 1999].This agreement suggests that the postseismic transient effect

Figure 5. Schematic of margin deformation through thegreat earthquake cycle illustrating the interseismic upliftand coseismic subsidence of the coastal region. Thecoseismic subsidence for most great events, up to 2m, is about500 years times the uplift rate per year. Most repeated levelingand other vertical geodetic data are landward of the peak upliftand most coastal paleoseismic subsidence data are near themaximum subsidence.


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is quite small for at least these megathrust events. There also isgood correspondence of the modeled locked zone using geo-detic data and the main downdip rupture zone for the 2010M8.8 Chile event [e.g., Delouis et al., 2010; see also Ruegget al., 2009], the 2004 M9 Sumatra earthquake [e.g., Shearerand Burgmann, 2010, and references therein], and the 2011M9 Tohoku, Japan event [e.g., Loveless and Meade, 2011].The interseismic locked zone estimated a few decades afterthe last great earthquake is likely a reasonable estimator ofthe rupture zone for Cascadia, as discussed for past great earth-quakes off SW Japan.[24] Another possible effect on the use of the locked/tran-sition zone as a rupture estimate is that the zone of ETS slowslip downward of the main rupture zone is locked or nearlylocked on a timescale averaging about 14months [e.g.,Rogers and Dragert, 2003]. Strain on this short-term lockedzone is released during a period of several weeks of slow slip.However, if longer-term average geodetic data are used, theeffect of the slow slip zone events on the estimate of the mainlocked zone and the estimate of the subsequent rupture zoneis small [e.g., Holtkamp and Brudzinski, 2010].[25] Both forward models and inversions have beenapplied to the backslip modeling with similar results. Theinversion results generally have a longer taper inland as aconsequence of the smoothing required in the method, unlessthe transition taper width is constrained in the inversion[e.g., McCaffrey and Schmidt, 2012; McCaffrey et al., 2013].Most recent models used profiles with smoothly increasingdip landward [e.g., Hyndman and Wang, 1995], and Flücket al. [1997] developed a 3-D fault model that has beenupdated by McCrory et al. [2006; 2012a] and used in mostsubsequent modeling. A simple fully locked zone and lineartransition zone have been used in the initial models, but a

more appropriate exponential transition zone was proposedby Wang et al. [2003] (Figure 4).[26] Vertical geodetic data give a quite well-defined limit torupture because both coseismic and interseismic vertical dis-placements change from upward to downward approximatelyat the limit of significant rupture on the fault, i.e., hinge line(Figures 5 and 6). Most Cascadia vertical constraints havecome from repeated high-precision leveling and long-durationtide gauges [e.g., Mitchell et al., 1994; Hyndman and Wang,1995; Verdonck, 2005; Burgette et al., 2009]. GPS verticaldata have lower resolution than horizontal data and are justbeginning to have useful accuracy of a few mm/yr [e.g.,Mazzotti et al., 2007; McCaffrey and Schmidt, 2012]. High-resolution gravity and borehole strainmeters also are providingadditional strain data for comparison with dislocation models[e.g., Mazzotti et al., 2007; Henton et al., 2010]. An exampleof vertical releveling data and models depicting the observedand predicted pattern of coastal uplift are shown in Figure 7[Hyndman and Wang, 1995; Burgette et al., 2009]. Note that

Figure 6. SW Japan 1944/1946 earthquakes, comparison ofcoseismic vertical displacements and interseismic verticalmotion from repeated leveling measurement, and dislocationmodels with a linear transition zone (modified from Hyndmanet al. [1995]). The coseismic vertical has been inverted andscaled by the average earthquake cycle duration. The horizon-tal lines at the top show the three widths of the locked (thickline), transition (thin line), and free slip zones for the threevertical models shown, i.e., ±10 km from best fit model (bluecurve). The good agreement between the interseismic lockedzone and the coseismic rupture zone deformations indicatesthat the geodetically determined locked/transition zones aregood estimators for the rupture/transition zone.

Figure 7. Examples of leveling data and corresponding dis-location models across the Cascadia margin (modified fromHyndman and Wang [1995]). The rates of vertical motionderived from leveling data are compared to those predictedby the three different width dislocation models shown; thelabels give the model locked and transition zones widths inkilometers and the average dip angle for the locked part ofthe thrust. Detailed data for the Oregon margin have beenpresented by Burgette et al. [2009].


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the maximum uplift rates of about 1–4mm/yr at the coast fromgeodetic data, multiplied by the average interevent period of500–600 years, predict maximum coseismic subsidences of0.5–2m, in general agreement with those observed in thecoastal marsh data along the coast [e.g., Leonard et al.,2010]. The patterns of the locked and transition zones for allof the studies are very similar to those of the Flück et al.[1997] reference model, with successive improvements frommore accurate data and dislocation parameters. Burgetteet al. [2009] inferred an offset to a shallower locked depth overthe southern end of the coastal Siletz Terrane, but the effectis small.[27] Horizontal GPS data show a clear signal of the coastalregion being squeezed landward [e.g., McCaffrey et al.,2013, Figure 8a]. The vectors also show the strong fore-arc crustal deformation that must be accounted for whenconducting megathrust dislocation modeling. Forward dislo-cation modeling that matches observations of landwardshortening has been presented by many authors with succes-sive improvements from more extensive and accurate dataand dislocation parameters [e.g., Dragert and Hyndman,1995; Flück et al., 1997; Miller et al., 2001; Wang et al.,2003; Mazzotti et al., 2003, 2007]. Wang et al. [2003] used

an effective transition zone that decreased exponentiallylandward to allow for the postseismic transients (Figure 4).The recent models corrected for the effect of the Oregonblock fore-arc block motion, i.e., used the convergenceacross the near-coastal margin rather than the ocean platerelative to North America.[28] Backslip inversions of the GPS and vertical data havebeen presented by Verdonck [2005], Yoshioka et al. [2005],and McCaffrey et al. [2007, 2013] which are generally inagreement with the results of forward modeling. In the mostrecent inversion analysis McCaffrey et al., [2013] found agood fit of both horizontal and vertical data (Figure 8b).They concluded that there is a somewhat wider transitionzone off central Oregon compared to southern Oregon andnorthernmost California but the differences compared to pre-vious models are quite small.[29] From modeling both horizontal and vertical data, theestimated full rupture (locked zone) is offshore everywhereexcept near Cape Mendocino in Northern California, withmodel linear transition zones extending onshore in northernWashington and near the coast in Oregon and VancouverIsland. The recent inversion model from McCaffrey et al.[2013] is shown in Figure 8b. This 50% locked location isvery close to that for the Flück et al. [1997] reference model,i.e., middle of the transition zone. It is not yet clear howmuchof the small-scale variability from this type of inversion isreal or due to data uncertainty. Smooth inversion transitionzones have rupture downdip displacement of 2m or ~10%of full rupture near the landward limit of the linear forwardmodel transition zones, but this part of the models is verypoorly constrained.

3. Rupture Zone From Paleoseismic CoastalMarsh Subsidence in Great Earthquakes

[30] The abrupt vertical deformation associated withthe most recent great earthquake in 1700 and earlierevents recorded by subsidence in coastal intertidal marshes(Figure 9) can be used to constrain the area of rupture, i.e.,“paleogeodesy” [Leonard et al., 2004; 2010; Wang, 2012;Hawkes et al., 2011; Peterson et al., 2012]. Deeper-buriedmarsh surfaces are covered by sediments deposited aftersuccessive coseismic subsidence events of up to 2m (e.g.,Figure 9a). Although the coseismic subsidence is approxi-mately recovered through strain buildup in the interseismicperiod, ongoing sea level rise results in progressive burialof subsided marshes [e.g., Atwater, 1987]. On the coast ofVancouver Island, there is ongoing land uplift faster thansea level rise so only the 1700 buried marsh remains[Clague and Bobrowsky, 1994]. The vertical distancebetween the present marsh top and the 1700 marsh is notthe coseismic subsidence because of eustatic sea level rise,the elastic strain buildup since that event, and tectonic verti-cal motions. However, a rough estimate can be obtained fromthis vertical distance after correcting for these effects. Betterdisplacement estimates come from careful calibration of theelevation range of a variety of coastal organisms, allowingaccurate estimation of the paleo-elevation differencesbetween the sediments just above and just below the top ofthe buried marshes. The resolution is commonly ~0.5m butmay be as good as ±0.3m [e.g., Guilbault et al., 1996;Peterson et al., 2012; Hawkes et al., 2011].

Figure 8. (a) Cascadia GPS vectors of current deformation(modified from McCaffrey et al. [2007]). The motion ratesinclude the effect of elastic strain buildup on the subductionthrust and fore-arc crustal motions. (b) The elastic dislocationmodel from inversion of horizontal and vertical data. Thered dashed line is approximately 50% locking (~25mm/yr)(modified from McCaffrey et al. [2013]). A locked or nearlyfully locked thrust and transition zone is indicated for thewhole margin, very similar to the Flück et al. [1997] refer-ence model shown as black solid and dashed lines.


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[31] For individual events, there is some trade-off betweenthe amount of rupture displacement and the landward limit ofrupture, but the change from seaward coseismic uplift tolandward coseismic subsidence, i.e., hinge line, is approxi-mately at the full displacement rupture limit (see Figures 5and 6). Because coseismic coastal subsidence is observedfor most of the Cascadia margin, this hinge line and thefull coseismic rupture are concluded to be well offshore.The only exception is for a small area just north of CapeMendocino where there has been coastal coseismic upliftand full rupture displacement is inferred to extend to beneaththe coast. This hinge line constraint limits the uncertainty inthe landward limit from this trade-off. Where there is only avery small coastal subsidence in central Oregon, there is anadditional uncertainty [e.g.,Wang, 2012; Pollitz et al., 2010].[32] The observed coastal subsidence for the 1700 eventand the average of previous events compared to the pre-dictions of the backslip model from Leonard et al., [2010]for the whole margin are shown in Figure 10a along withthe high-resolution coastal subsidence data from Hawkeset al. [2011] for Oregon shown in Figure 10b. There is goodagreement of the observed coastal subsidence rupture limitswith the estimates from the geodetic data locked and transi-tion zones, within the considerable downdip uncertaintiesof 20–40 km. Wang [2012] modeled the coseismic subsi-dence pattern with rupture that varied along the margin toget a better fit to a section of the coast with high-resolutiondata. A slightly better model fit was found to the data butthe misfit differences compared to those for rupture scaleddirectly to convergence rate along the margin are withinthe uncertainties.

4. Downdip Temperatures: The Seismic-AseismicTransition

[33] Above a critical temperature, rocks deform by ductile-type mechanisms and seismic rupture is not expected,as much discussed for continental transcurrent faults likethe San Andreas [e.g., Chen and Molnar, 1983; Wong andChapman, 1990; Ito, 1990]. The transition from brittle toductile behavior is closely related but not identical to the limitof seismic instability [e.g., Scholz, 1998]. For seismic behav-ior, the critical change downdip is from velocity-weakeningto velocity-strengthening slip. With velocity weakening, theslip surface weakens with ongoing displacement and the ef-fective coefficient of friction decreases. Runaway earthquakedisplacement can occur that releases much if not all of thebuilt-up elastic energy. For velocity strengthening, the slipzone strengthens with ongoing displacement so slip speedsare limited, a viscous-like behavior. For felsic crustal rockcompositions appropriate for the continental crust overlyingthe thrust, for accretionary sedimentary prisms, and subduc-tion channel sediments, the velocity-weakening to velocity-strengthening transition from laboratory data is about 350°C[Tse and Rice, 1986; Blanpied et al., 1991; 1995; Bürgmannand Dresen, 2008] which agrees reasonably well with theestimated temperatures at the maximum depth of seismicityin continental fault zones [e.g., Chen and Molnar, 1983;Wong and Chapman, 1990; Ito, 1990; Bonner et al., 2003].The mafic subducting oceanic crust may have a higherseismogenic temperature limit, but the thrust deformation isexpected to occur in the weaker overlying sediments or fore-arc crust. The exception of local areas where there is shearing

Figure 9. (a) Example of a buried intertidal marsh off central Washington attributed to coseismic subsi-dence during the M9 earthquake of 1700 (photo by L. Leonard). A tsunami-driven sand sheet overlies theburied marsh. (b) Comparison of coastal coseismic subsidence estimates for an area of the Oregon coastplotted with distance orthogonal to the coastal trend, and corresponding dislocation models (location shownin Figure 10), for the 1700 and pre-1700 events [after Leonard et al., 2010]. The model results are shownfor a 14m full downdip rupture displacement and for the rupture that releases the total shortening elasticstrain for a 500 year return period, using the variable convergence rates along the margin.


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of mafic seamounts, etc., having a higher temperature limit,is discussed below. Also, as discussed below, aseismicserpentinite/talc may be important downdip of the fore-arcmantle corner.[34] At the landward downdip end of great earthquake rup-ture, there must be a transition zone within which rupturecannot initiate but into which slip can continue downdip ifrupture is initiated updip. This zone represents “conditionalstability” [e.g., Scholz, 2002]. The downdip limit of thetransition zone is expected to be where the instantaneousshear stress for slip increases rapidly downward, for exam-ple, from the effect of increasing temperature. Laboratorystudies suggest this limit is at about 450°C for crustal rocks[e.g., Blanpied et al., 1991; 1995]. Thermal model estimatesof downdip temperatures are then required to define the loca-tion of 350°C and 450°C on the thrust.[35] Recent subduction thermal models use finite elementmodeling, a landward steepening thrust profile, accurateincoming plate thermal regimes that include the effect ofoverlying insulating sediments, and the effect of radioactiveheat generation and of variable thermal conductivity, e.g.,Figures 11a and 11b [Hyndman and Wang, 1993, 1995;Hyndman et al., 1995]. Their map locations of the 350 and450°C isotherms on the thrust are shown in Figure 2. Thelocations are within the uncertainties of the limits of the ref-erence model locked and transition zones from geodetic data.

It has been argued that the effect of frictional heating onthe thrust in the depth range of the seismogenic zone issmall and usually may be neglected [Wang et al., 1995;Wang et al., 1995; Wang and He, 1999] but the magnitudeof heating remains a source of uncertainty [see McKennaand Blackwell, 2002]. Currie et al. [2004] included convec-tion in the back arc in the models but this made no significantdifference to temperatures in the region of the seismogeniczone. Gutscher and Peacock [2003] proposed a flat slab pro-file for the Cascadia subducting slab and estimated about30 km farther landward positions of the 350 and 450°C tem-peratures on the thrust compared to the previous models.However, the compilation of McCrory et al. [2006, 2012a]with more recent data does not show the degree of slab flat-tening used in their Cascadia models.[36] Cooling of the crust beneath the margin by hydrother-mal circulation has been proposed by Cozzens and Spinelli[2012] such that the critical thermal limits could be as muchas 30–55 km landward relative to results from simulationswithout fluid flow. Heat may be transferred from beneaththe accretionary sedimentary prism to the adjacent deep-seabasin through fluid flow in the subducting crust, decreasingthe heat flow beneath the continental slope and increasingthe heat flow from the adjacent deep sea floor. Off SWJapan where the incoming sediment section is much thinner,the effect of hydrothermal circulation on the thermal data isclear [Spinelli and Wang, 2008], with heat flow seaward ofthe deformation front much higher than for a conductivemodel. However, off Cascadia where the sediment sectionoffshore is very thick and there may be little crustal hydro-thermal circulation, there is no indication of high heat flowseaward of the deformation front; the heat flow correspondsto that expected for a cooling plate model after correctionfor the sedimentation effect. Landward of the deformationfront, the best model fit to the heat flow data across theVancouver Island margin is for little or no hydrothermal cir-culation cooling effect (Figure 11). On this profile, the heatflow has been corrected for the thermal effect of accretionaryprism sediment thickening [Hyndman et al., 1993]. The~15% effect on the thermal model is shown in Figure 11a.It is likely that the models with little or no effect from hydro-thermal circulation would also agree with the thermal data offWashington, Oregon, and California if a correction was madefor the sediment thickening reduction effect. However,Cozzens and Spinelli [2012] also argued that the effect maybe important based on the depth to the basalt-eclogite tran-sition indicated in receiver function structure data, whichemphasizes the uncertainties in the thermal model predictedthrust temperatures.

5. Change in Thrust Seismic Ref lection CharacterFrom Thin Sharp Seismic to Thick Aseismic DuctileShear Zone

[37] From the thermal analysis summarized above, it isexpected that there is a thermally controlled downdip changefrom brittle-seismic to ductile-aseismic behavior at about350°C–450°C. This transition roughly corresponds to thetransition as seen in field geology studies from cold brittletype behavior and thin fault zones, to broad mylonitic, andthen ductile shear zones [e.g., Sibson, 1977, 1983; Cole,et al., 2007]. The thin shear represents strain weakening

Figure 10. (a) Along the coast comparison of the mag-nitude of coastal marsh subsidence and predictions fromdislocation models for 200 and 500 years of strain buildupat the convergence rates shown [after Leonard et al., 2010].The best fit is for approximately a 500 year great earthquakereturn period indicating that most of the interseismic plateconvergence is taken up as elastic strain. The ellipse showsthe location of the data in Figure 9. (b) High-resolution1700 coseismic subsidence estimates for Oregon comparedto elastic dislocation models (modified from Hawkes et al.[2011]). The shaded vertical bar has lower accuracy. Thegray band is the mean estimate of coastal subsidence fromLeonard et al. [2010].


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whereas the thick shear likely represents strain strengthening.In the latter, the fault zone continually moves outward fromthe initial shear zone which becomes stronger with defor-mation [e.g., Kirby, 1985] and thickens into weaker as yetundeformed rocks. The transition from thin brittle to thickductile behavior is expected to be at a slightly higher temper-ature than the limit of seismic instability [e.g., Scholz, 2002]and the comparison needs better calibration.[38] As described by Nedimovic et al. [2003], the changefrom a thin strong detachment reflector to a reflective5–10 km thick band in marine and land multichannel seismicreflection data approximately corresponds to the downdiplimit of seismic rupture for M> 8 earthquakes at a numberof subduction zones, e.g., Alaska, Chile, and SW Japan.They concluded that the thin to thick reflection transitionmarks the brittle to ductile transition. They mapped this transi-tion along the margin of southern Vancouver Island (reflectionzone 2 on map of Figure 12) and found it generally lies justseaward of the coast and corresponds reasonably well to themodeled transition zone downdip limit from other constraints,i.e., approximately 450°C. Despite considerable variation inthe reflection character of the subduction thrust on a numberof seismic lines, there is a consistent downdip change from athin reflection zone offshore (reflection zone 1) to a broadreflection band (E layer; reflection zone 3) some 5 to 7 kmthick, near the west coast of Vancouver Island (Figure 12).The thin reflection is less than 2 km thick, but the true bound-ary is probably thinner, the image being thickened by thereflection signature. The transition is a few tens of kilometerswide that corresponds in position very approximately tothe downdip limit of the geodetic and thermal transition(450°C) zones (Figure 12).

[39] This interpretation across the Vancouver Island mar-gin assumes that the reflective E-zone [e.g., Clowes et al.,1987] is the detachment. However, other authors [e.g.,Audet et al., 2009; Hansen et al., 2012] preferred the inter-pretation that the reflective band is the subducting oceaniccrust or the upper crust. In contrast, from detailed seismicstructure analysis, Calvert et al. [2011] concluded that theE-layer dipping reflective band is a shear zone above thesubducting plate as in the original interpretation of Cloweset al. [1987]. This disagreement is not yet resolved. Furthercalibration of this method using the change in reflection char-acter downdip is needed for subduction zones that have hadrecent great earthquakes with a well-constrained downdiprupture limits. Initial results from a seismic structure surveyto examine the downdip change in reflection character forthe Alaska subduction zone were reported by Shillingtonet al. [2011].

6. Fore-Arc Mantle Corner: Aseismic Serpentiniteand Talc on Thrust

[40] The fore-arc mantle corner may limit downdip rupturebecause of the presence of aseismic serpentinite and talc. Thefore-arc mantle is commonly significantly serpentinized byrising fluids driven upward from the underlying dehydratingsubducting crust [e.g.,Hyndman and Peacock, 2003, and ref-erences therein] (Figure 13). The structure of serpentinite isexpected to make it weak where it is in contact with theunderlying fault zone, a conclusion supported by the argu-ments for no significant frictional heating on the thrust, atleast landward of the fore-arc mantle corner [e.g. Wadaet al., 2007]. There is some uncertainty as to the weakness

Figure 11. (a) Heat flow data across the northern Cascadia margin compared to the thermal model pre-diction ofHyndman and Wang [1995]; the dashed red line shows the prism thickening effect that depressesthe surface heat flow. (b) Thermal cross section showing the thermal downdip limits to postulated locked(350°C) and transition (450°C) zones. The 350 and 450°C isotherms and uncertainties are shown on themap of Figure 2, compared to the locked and transition zones from modeling geodetic data. If ocean crustalhydrothermal cooling is important [Cozzens and Spinelli, 2012], these temperature limits may be severaltens of kilometers landward.


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of serpentinite [e.g., Escartın et al., 2001;Moore et al., 1997;Moore and Rymer, 2007; Reynard, 2013], but another factormay ensure that the thrust beneath the fore-arc mantle is usu-ally aseismic. Talc is generated by the silica-saturated fluidsfrom the underlying dehydrating oceanic crust and under-thrust sediments that react with the overlying fore-arc mantleserpentinite [e.g., Peacock and Hyndman, 1999]. Talc is avery weak mineral and fault zones containing talc arevery likely aseismic, as suggested for the San Andreas Fault[Moore and Rymer, 2007]. Brocher et al. [2003], Prestonet al. [2003], and Ramachandran and Hyndman [2011] usingwide-angle seismic structure and seismic tomography studiesshowed that there is a substantial reduction in seismic velocityand increase in Poisson’s ratio in the Cascadia fore-arc mantle,indicative of extensive serpentinization, and Blakely et al.[2005] showed that the fore-arc mantle serpentinization pro-duces a significant magnetic anomaly high.[41] Tichelaar and Ruff [1993], Hyndman et al. [1997],and Oleskevich et al. [1999] concluded that the maximumdepth of great earthquake rupture, commonly defined byaftershocks, often corresponds to the intersection with thefore-arc mantle corner. Unfortunately, the location of thefore-arc mantle corner has not yet been well defined on mostmargins. Ruff and Tichelaar [1996] concluded that for theSouth American subduction zone, the downdip limit of greatearthquake rupture approximately coincides with the fore-arcmantle corner which in turn is approximately at the coast formuch of that margin. Heuret et al. [2011] compiled the max-imum depth of intermediate magnitude thrust events for anumber of M5.5–7.0 thrust earthquakes in a global studyand concluded that the average maximum event depth forsubduction zones where ocean plates descend beneath conti-nental lithosphere was about 40 km. This is somewhat deeperthan the average of 34 ± 11 km for the compiled depth esti-mates of fore-arc Moho depth by Wada and Wang [2009].However, this is not a large difference since the Moho inter-section depth for most subduction zones is very poorlydefined. For example, theMohomay not exhibit a crust-mantledownward velocity increase because of the low-velocity

fore-arc mantle serpentinite [e.g., Bostock et al., 2002].Therefore, this conclusion remains uncertain.[42] Three recent well-studied great earthquakes allowexamination of this proposed downdip rupture limit forM ~ 9 events, Sumatra in 2004, Chile in 2008, and NEJapan in 2011. For the great earthquake of 2004 offSumatra, the rupture extent was to a depth of about 30 km[e.g., Chlieh et al., 2007, and references therein] in agree-ment with the locked zone from geodetic and paleogeodeticdata [e.g., Chlieh et al., 2008]. The depth of ~30 km is withinthe common range of 30–50 km for continental fore-arc man-tle corners, but Dessa et al. [2009] and Simoes et al. [2004]have concluded from seismic structure data that the fore-arcMoho depth is very shallow, 21–25 km, on this margin suchthat the locked/rupture zone extends to beneath the fore-arcmantle. If this unusually shallow Moho depth does apply tothis fore-arc mantle corner, the corner does not constraindowndip rupture on this margin. However, the inference ofan exceptionally shallow fore-arc mantle leaves considerableuncertainty. Also, the depth and the critical temperature forseismic behavior are similar [Hippchen and Hyndman,2008], so the downdip limit may be thermally controlledrather than by the fore-arc mantle corner.[43] For the great earthquake off Chile in 2008, Tong et al.[2010] from InSAR data and Delouis et al. [2010] fromgeodetic data concluded a downdip coseismic rupture limitnear the depth where the subducting Nazca plate intersectswith the continental Moho of the South America plate at40–50 km [Haberland et al., 2006]. This depth is in the rangeof maximum rupture depths estimates for a compilation ofearlier large thrust earthquakes on this margin by Tichelaarand Ruff [1991]. This depth also agrees with the limit of thelocked zone inferred from preseismic GPS data [Ruegget al., 2009; Moreno et al., 2010]. Haberland et al. [2006]found the fore-arc Moho at a depth of about 50 km in goodagreement with the fore-arc mantle corner being the downdiplimit for rupture in this great earthquake.[44] For the third recent great earthquake, Tohoku off NEJapan in 2011, the main large displacement rupture appears

Figure 12. Change in megathrust multichannel seismic reflection character from a thin inferred seismicsignature (Reflection band 1), to a thick inferred aseismic shear zone (Reflection band 3; modified afterNedimovic et al. [2003]), compared to Flück et al. [1997] reference locked and transition zones frommodeling geodetic data, i.e., approximately 350°C and 450°C. The reflection transition (Reflection band2) occurs at slightly cooler than 450°C and seaward of the ETS zone.


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to terminate just seaward of the coast where the thrust depthis 30–40 km deep [e.g., Simons et al., 2011; Loveless andMeade, 2011], a common depth for the fore-arc mantle cor-ner. Aftershocks extend to a depth of 40–45 km just seawardof the coast. However, seismic structure data reported byHino et al. [2000] and Takahashi et al. [2004] just to thenorth suggest that the fore-arc Moho is at an unusuallyshallow depth of about 20 km. If the latter is correct, thefore-arc mantle corner does not limit downdip rupture. Seno[2005] also found the maximum depth of other thrust seis-micity inferred to be on the subduction thrust to extend to over50 km depth on this margin, much deeper than the estimatedfore-arc Moho. However, he found the maximum thrust earth-quake depth to be only ~30 km to the north of the 2011 rupturearea and to the south off SW Japan, in general agreementwith the fore-arc Moho depth. As for Sumatra, rupture offNE Japan appears to extend deeper than the fore-arc mantlecorner, but in both cases, this conclusion depends on thefore-arc Moho being exceptionally shallow at 20–25 km. Ifthe fore-arc Moho corner is actually at the more common35–50 km, there would be agreement with this rupture limit.[45] Lay et al. [2012] in a summary of global great earth-quake behavior found that large earthquake displacementsmost commonly occur to a depth of 35 km. Ruptures of smallerisolated patches on the thrust extend deeper to ~55 km, gener-ating only small magnitude earthquakes. Their results suggestthat some smaller thrust events may extend to greater depthsthan the main displacements in megathrust events. Cautiontherefore should be used in employing the maximum depthof smaller thrust events for a limit to large displacementrupture and to megathrust earthquake hazard.[46] Another complicating factor that may result in thrustearthquakes deeper than the fore-arc mantle corner is the pos-sibility that some thrust earthquake ruptures occur withinsubducted sediments, subducted seamounts, or other sub-duction channel material, isolated from overlying fore-arc

mantle serpentinite and talc. More detailed structural studiesare needed in the region of the Moho intersection where thedowndip rupture limit for both low- and high-frequencyenergy has been well determined for recent great earthquakes.[47] The thickness of the Cascadia fore-arc crust is quitewell determined in several areas by a number of seismic stud-ies, including receiver function analyses, seismic tomography,and wide-angle seismic structure experiments. However, thedip angle of the thrust profile is not so well determined andthere remains some debate, especially for the VancouverIsland margin, as to what represents the thrust in seismic struc-ture data [e.g., Hansen et al., 2012]. McCrory et al. [2006,2012a] provided a review of the thrust constraints. An initialsummary of constraints to the location of the fore-arc mantlecorner has been provided by McCrory et al. [2012b]. Theaccuracy is sufficient to show that the fore-arc mantle corneris generally well landward of the seismogenic limit from mostof the other constraints (e.g., Figure 14a). The differences invarious thrust profile estimates are not significant to a depthof about 30 km [e.g., Wech et al., 2009, their Figure 4]and so are not very important for the seismogenic zone as de-fined by the geodetic, thermal, and coastal marsh constraints.However, should the seismogenic zone extend deeper, atgreater depths, there are significant differences in various esti-mates of the thrust dip profile. A constraint to the location ofthe fore-arc mantle corner may come from the updip limit ofETS tremor (discussed later). Wech et al. [2009] suggestedthat the well-resolved, sharp updip edge of ETS tremor epi-centers reflects a change in plate interface coupling properties.This boundary in coupling may be associated with the fore-arcmantle corner (e.g., Figures 13 and 14a), with ETS occurrencebeing controlled by the position of the fore-arc mantle corner[e.g., Holtkamp and Brudzinski, 2010].

7. ETS Slow Slip Updip Limit

[48] Episodic tremor and slip (ETS) is a remarkable recentdiscovery, with the most detailed studies in Cascadia and SWJapan [e.g., Rogers and Dragert, 2003; Obara et al., 2004;Schwartz and Rokosky, 2007; Peacock, 2009; Gomberget al., 2010; Beroza and Ide, 2011]. A number of authorshave suggested that ETS may provide a great earthquakedowndip rupture limit. The slow slip associated with ETStremor likely occurs mainly on or just above the subductionthrust, and for Cascadia accommodates much if not all ofthe plate convergence in the ETS zone [e.g., Dragert andWang, 2011]. Therefore, there is little or no elastic strainbuildup in the ETS zone to be released in great events. If thislast conclusion is correct, the slow slip zone represents adowndip limit for significant rupture displacement. The larg-est uncertainty as to whether most if not all of the plate con-vergence is so accommodated lies in the amount of slip insmall events between the main ETS events [e.g., Wanget al., 2008]. The physical processes involved in ETS remainuncertain, but the slow slip behavior itself indicates that thefault zone has ductile characteristics at that depth.[49] It cannot be excluded that there is some small elasticstrain buildup in the ETS zone and that great earthquake rup-ture extends into the ETS zone with very small displacement.It is recognized that even a small displacement slip in thisregion could substantially influence the resulting seismicshaking at major cities in Cascadia. If slow slip represents

Figure 13. Cascadia section showing the inferred aseismicserpentinite and talc on the subduction thrust that may limitseismic rupture downdip (modified from Peacock andHyndman [1999]; bru = brucite; serp = serpentinite). TheETS zone approximately corresponds to the fore-arc mantlecorner, whereas the estimated Cascadia seismogenic zone issignificantly shallower.


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conditionally stable frictional behavior, is it possible that theslow slip event zone could be pushed to seismic rupture by alarge earthquake on the seismogenic portion of the interface.However, the intervening zone between ETS and the mainlocked/rupture zone estimated from most constraints alsowould need to rupture coseismically.[50] The horizontal positions of ETS tremor are quitewell determined [e.g., Kao et al., 2009; Wech et al., 2009;Gomberg et al., 2010; Dragert and Wang, 2011; Ghoshet al., 2012] (e.g., Figure 14b) although the depths are some-what uncertain. The location of the slow slip has been esti-mated through dislocation modeling of geodetic data and itslocation is also quite well determined but not quite as wellas the tremor [e.g., Dragert and Wang, 2011; Wech et al.,2009; McCaffrey, 2009; Schmidt and Gao, 2010; Bartlowet al., 2011]. There is some indication that the slip is slightlyseaward of the tremor but this is within the uncertainties(example in Figure 14a).Hawthorne and Rubin [2013] founda short timescale correlation between slow slip and tremor inCascadia that indicates that they are closely related.

[51] Audet et al. [2010] and Ghosh et al. [2012] concludethat the top of the oceanic plate is at a depth of 30–40 kmbeneath the tremor based on various estimates of plate depth.In their review, Gomberg and the Cascadia 2007 andBeyond Working Group [2010] also concluded that the depthrange for Cascadia slow slip is generally 30–40 km. The fore-arc Moho is approximately at 30 km so the tremor appears tobe across and landward of the fore-arc mantle corner but thedepth uncertainties for the fore-arc Moho are substantial. In amore conclusive analysis of slow slip from continuous GPSdata, Holtkamp and Brudzinski [2010] found that the ETSzone is likely located where the subduction thrust underliesthe fore-arc mantle, such that the fore-arc mantle corner isthe updip limit of most of the ETS. Peacock [2009] alsoconcluded that the tremor for SW Japan and Cascadia is justbelow or at the fore-arc mantle corner. A physical composi-tion or structure constraint is supported by the conclusionof Peacock [2009] that tremor occurs at different depths fordifferent subduction zones. He estimated the temperatureon the thrust at the depth of the Cascadia tremor to be about

Figure 14. (a) Slow slip from inversion of GPS data and tremor locations [Dragert and Wang, 2011]for a portion of the Cascadia margin. The reference locked and transition zones are shown for com-parison. The approximate location of the fore-arc mantle corner is from McCrory et al. [2012b].There is a gap of 50–100 km between the updip limit of slow slip and tremor and most estimatesof the downdip limit of significant great earthquake rupture. (b) Distribution of Cascadia tremor2009–2013 (data from Wech et al. [2009] and Wech [2010], http://www.pnsn.org/tremor), comparedto Flück et al. [1997] reference model locked and transition zones. (c) Illustration that the sum ofthe coseismic downdip transition and the updip transition for the slow slip does not equal the totalplate convergence rate over the great earthquake cycle. There must be an intervening region that doesnot move coseismically or in slow slip events.


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510°C, well above the proposed maximum of about 350°Cfor seismic behavior. The slow slip is therefore 30–50 kmlandward of the great earthquake rupture zone from thethermal constraint and most of the other constraints for thedowndip limit of rupture.[52] A possible explanation for the ETS tremor beingassociated with the fore-arc mantle corner is that the fluidsresponsible are channeled upward and seaward beneath thefore-arc mantle to rise above the corner. Very low Poisson’sratio in the crust above this corner has been observed,suggested by Ramachandran and Hyndman [2011] to resultfrom silica precipitation from the rising fluids. Katayamaet al. [2012] also proposed an impermeable barrier at thefore-arc Moho that channels fluids seaward to rise into thecrust above the corner producing tremor.[53] If the conclusion that the slow slip of ETS occurs welllandward of the great earthquake rupture zone is correct(Figure 14), it means that there is an intervening section ofthe thrust where long-term plate convergence is accommo-dated neither in great earthquake rupture nor in slow slipevents, but rather in some form of long-term aseismic creep[e.g., Holtkamp and Brudzinski, 2010]. It should be notedthat Cascadia has an unusually shallow downdip great earth-quake rupture limit, probably because there is a thermalrupture limit for this very hot subduction zone. The more com-mon cool subduction zones often have a much deeper limitthat may be at the fore-arc mantle corner and may not have agap between the great earthquake rupture and ETS slow slip.

8. Geological Associations of RuptureWith Offshore Basins

[54] There is a common correlation between offshore fore-arc basins, defined mainly by gravity, and the regions of seis-mic rupture in past great events globally [Wells et al., 2003;Song and Simons, 2003]. In a related observation, Ruff andTichelaar [1996] found that the landward limit of great earth-quake rupture is often beneath the coast. The coast oftendefines the landward limit of the basins. This correlationneeds to be examined for recent great events such as the2004 great earthquake of Sumatra, the 2008 great earthquakeof Chile, and the 2011 earthquake of NE Japan.[55] Fuller et al. [2006] and Llenos and McGuire [2007]discuss mechanisms to explain this correlation. The largestseismic moment release for the majority of great earthquakesstudied elsewhere tended to occur beneath local lows in the

gravity field along the margin [e.g., Wells et al., 2003].Also, the limits of rupture approximately correspond withincreases or maxima in the trench parallel gravity field. Aphysical mechanism has not yet been identified; however,the deep-sea terraces and basins may evolve not only bygrowth of the outer arc high but also by interseismicsubsidence that is not recovered by uplift during greatearthquakes. There could be a link between subsidence,subduction erosion, and seismogenesis. Kilometers of lateCenozoic subsidence and crustal thinning above some ofthe source zones elsewhere are indicated by seismic profilingand drilling. They are thought to be caused by basal subduc-tion erosion [e.g., Scholl and von Huene, 2007] althoughlarge amounts of erosion are most commonly argued for sed-iment deficient subduction zones with high convergencerates [e.g., Stern, 2011] and may not apply to Cascadia.[56] The source zone for the Cascadia 1700A.D. earth-quake contains five semicontinuous basin-centered gravitylows (Figure 15) that may indicate asperities at depth [e.g.,Wells et al., 2003]. The gravity gradient marking the inferreddowndip limit to large coseismic slip lies offshore, except innorthwestern Washington, where the low extends landwardbeneath the coast. Figure 15 shows the good correlation be-tween the locked and transitions zones defined geodeticallyand thermally and the gravity-defined basins for Cascadia.The basin landward extents are also in general agreementwith the estimated landward limit of rupture in 1700 andearlier great events based on coastal coseismic subsidence[Leonard et al., 2004; 2010]. On other margins, seismic sliphas been concluded to be less beneath along-strike structuralhighs [e.g., Wells et al., 2003] so such highs may delineaterupture variability along the margin with some aseismic slip.

9. Downdip Limit of Small SubductionThrust Earthquakes

[57] The downdip limit of small to intermediate size thrustearthquakes on the subduction thrust provides an importantseismogenic zone constraint for many other subductionzones [e.g., Tichelaar and Ruff, 1993]. These thrust eventstend to concentrate near the downdip limit of rupture for greatearthquakes elsewhere [e.g., Tichelaar and Ruff, 1993; Ruffand Tichelaar, 1996].[58] However, on the Cascadia margin, only a very fewsmall such events have been detected and in a limited area[Tréhu et al., 2008; Williams et al., 2011] (Figure 16), but

Figure 15. Gravity anomalies along the Cascadia margin that may delineate the great earthquake ruptureextent downdip and possible along-strike segmentation defined by sedimentary basins A–E [after Wellset al., 2003], compared to Flück et al. [1997] reference locked.


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they are at the estimated depth of the subduction thrust andsome have been shown to have rupture mechanisms at thedip angle of the thrust. They may be related to local stressconcentrations associated with features such as subductedseamounts [Tréhu et al., 2012]. Some are accurately locatedwith ocean bottom seismographs and they may provide anestimate of the downdip seismogenic limit. The estimatedsource depths are 9–11 km for a M4.9 event and 15–17 kmfor a M4.8 event. Other small earthquakes are in the rangefrom 10 km to 18 km. These events are within the locationuncertainties of the subduction thrust and probably representseismogenic motion on the Cascadia megathrust. The deeperevents are within the transition zone for our reference model[Flück et al., 1997] but well updip of where episodic tremorand slip (ETS) has been documented andwell updip of the fore-arc mantle corner. Williams et al. [2011] concluded that theevents occurred at the downdip edge of the locked/transitionzone which may be some 30 km farther landward than theFlück et al. model (Figure 16) and suggested that the transitionzone extends approximately to the coast. However, an alterna-tive explanation is that these events may have occurred onsheared-off subducted seamounts [Tréhu et al., 2012] or localsmall blocks of crustal material in the subduction channel witha mafic composition. Such material could have a higher limit-ing temperature than the normal more felsic sediments andfore-arc crustal rocks on the main thrust contact [e.g., Chenand Molnar, 1983; He et al. 2007] and explain why thesethrust earthquakes occur where the temperature is higher thanthe concluded common seismogenic limit of about 350°C.[59] As noted above, Lay et al. [2012] in a summary ofglobal great earthquake behavior found that large earthquakedisplacements most commonly occur to a depth of 35 km butthat the rupture of smaller isolated patches often extendsdeeper to ~55 km. Their results suggest that some smallerthrust events may extend to greater depths than the main

displacements in megathrust events and therefore that themaximum depth of smaller thrust events may not define thelimit to large megathrust rupture displacement.

10. Landward Limit of Earthquakeson the Nootka Transform Fault

[60] The left-lateral Nootka transform fault zone lies offcentral Vancouver Island approximately orthogonal to themargin. The landward downdip extent of the strike-slip seis-micity on this fault zone (Figure 17) may provide a limit tothe downdip extent of seismic behavior on the overlyingsubduction thrust. The strike-slip motion is between the mainJuan de Fuca plate to the south and Explorer plate complex tothe north [Hyndman et al., 1979]. The fault zone is concludedto be the result of quite recent ~3Ma breakup of the Juande Fuca plate [e.g., Riddihough, 1984; Braunmiller andNábělek, 2002]. There is a continuous change in the orienta-tion of this fault zone, resulting in a broad irregular shearzone. The epicenter uncertainties for the Nootka fault zoneearthquakes are significant because most seismograph sta-tions are landward. However, a few events have been locatedwith ocean bottom seismographs [Hyndman and Rogers,1981] and the location bias near the coast is small. Strongseismicity on the Nootka fault zone extends approximatelyto the coast well landward of the deformation front. Likethe small thrust earthquakes off Oregon, the Nootka faultearthquakes extend approximately to the landward limit ofthe transition zone of the Flück et al. [1997] reference model(Figure 17), where the estimated temperature is about 450°C.[61] It is likely that the Nootka seismicity is thermally lim-ited landward, in a similar manner to the postulated thermallimit for continental strike-slip faults and for the subductionthrust. The seismicity on the subducted Nootka transformfault zone is slightly deeper than the megathrust and thereforemay be slightly hotter. The result is a landward seismogenic

Figure 16. Small thrust earthquakes that may define thedowndip limit of seismic behavior on the subduction thrustoff Oregon [afterWilliams et al., 2011]. They occurred wherethermal model temperatures on the subduction thrust inter-face are 400–450°C, so they may represent rupture of maficseamounts that have a higher seismogenic temperature limitthan most of the subduction thrust. The downdip limits tothe locked (black solid line) and transition (dashed line)zones are from the Flück et al. [1997] reference model.

Figure 17. Seismicity on the Nootka transform faultextending off Vancouver Island that may provide a landwardlimit of seismic behavior (from Geological Survey of Canadadata file), compared to the Flück et al. [1997] referencelocked and transition zones. These earthquakes extend towhere thermal model temperatures are about 450°C. Thishigher than normal seismogenic temperature limit may beexplained by rupture of mafic ocean crust material that hasa higher temperature seismogenic limit than most of thesubduction thrust.


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limit that is slightly seaward. However, the transform seis-micity is argued to be only in the upper few kilometers ofthe oceanic crust [Hyndman and Rogers, 1981; Willoughbyand Hyndman, 2005] so not much deeper and not much hot-ter than the subduction thrust. However, this fault cuts maficoceanic crustal rocks that likely have a higher maximumseismogenic temperature than for the more felsic composi-tion of the rocks and sediments in contact with the subductionthrust, so there may be a higher maximum seismogenic tem-perature and therefore the landward limit may be farther land-ward. For example, Chen and Molnar [1986] estimated thetemperature at the base of the seismogenic zone for the frac-ture zones cutting oceanic lithosphere to be 600–800°C andHe et al. [2007] from laboratory data found that seismicbehavior could extend to over 510°C for gabbroic oceaniclower crust rocks. This limiting temperature is in generalagreement with the landward limit of Nootka seismicity zone(Figure 17) at about 450°C and similar to that inferred for thesmall thrust earthquakes off Oregon discussed earlier. Thisconstraint provides evidence that significant thrust ruptureat least does not extend far beneath the Vancouver Islandcoast. From this constraint, the maximum megathrust ruptureextent may be seaward but is unlikely to be landward.[62] To the south of the Nootka fault zone, there is incom-ing plate seismicity that extends several tens of kilometerslandward of the coast (Figure 17). This seismicity does notcontinue to the south off Washington and Oregon. It couldresult from extensional faulting generated by plate bendingand steepening landward or due to dehydration fluid pres-sures [e.g., Abers et al., 2009]. Like the Nooka seismicityand the small events off Oregon, these earthquakes extendto where the estimate temperature is about 450°C which isapproximately that expected for mafic ocean crustal rocks.Like the Nootka fault constraint, the maximum megathrustrupture extent may be seaward but is unlikely to be landwardof this subducting plate seismicity.

11. Discussion

[63] The following is a summary of constraint results:[64] 1. The geodetic constraints for the current locked andtransition zones have excellent data and the highest resolution,about ±10 km differences among the various models. Thegreatest uncertainty is from the degree to which the locked/transition zones represent the full rupture and the taperingto zero rupture displacement. Recent great earthquakes else-where indicate that this is a good approximation. The seismicrupture limit may be slightly seaward of the locked/transitionzones because of the postseismic transients that are notincluded in most locked zone models.[65] 2. Coastal coseismic subsidence provides the leastambiguous constraint to the rupture zone but has quite lowresolution of about ±30 km downdip [see Leonard et al.,2010], and there is some ambiguity trade-off of location ofrupture and amount of rupture displacement, especially offOregon where the subsidence is small. There also is the ques-tion of how much of the displacement occurs in postseismicdeformation. The coastal subsidence constraint agrees wellwith the geodetic constraint.[66] 3. The thermal constraint has significant uncertaintiesin both the thermal models and in the critical temperatures(see uncertainties shown in Figure 2), but the results generally

agree well with the geodetic constraint. If hydrothermalcooling in the incoming crust is significant, the critical temper-atures may be a few tens of kilometers farther landward.[67] 4. The transition in thrust reflection character fromthin seismic to thick aseismic has quite high resolution [seeNedimovic et al., 2003] but has a large uncertainty becauseof the lack of a good calibration where there has been a recentmegathrust earthquake. From laboratory and field data, ductilebehavior in crustal rocks is expected at roughly 450°C. Thereflection transition agrees well with the transition zonethermal limit and the geodetic constraint transition zone limit.[68] 5. The fore-arc mantle corner or Moho limit hasconsiderable uncertainty, about 20 km [e.g., McCrory et al.,2012], but is about 50 km landward of the main rupture indi-cated by the geodetic models. It roughly agrees with the ETSlimit. This seismic rupture limit has been concluded for anumber of cold subduction zones elsewhere but there havebeen several interpretations of thrust earthquakes extendingdeeper. This constraint is only a maximum depth, and forCascadia the rupture limit appears to be thermally controlledat a shallower depth than the fore-arc Moho intersection.[69] 6. The ETS slow slip is quite well located to about±10 km (references given above). The ETS updip limit iswhere the thrust is at a depth of about 30 km, which corre-sponds quite well with the fore-arc mantle corner, althoughthe latter is not yet well constrained in some areas of themargin. The ETS updip limit is 30–50 km landward of thegeodetic locked/transition zone limit. This constraint is onlya maximum depth and allows a shallower rupture limit. ForCascadia, the rupture limit appears to be thermally controlledat a shallower depth.[70] 7. The center of gravity lows beneath the slope andshelf gives a moderate resolution constraint of about ±30 km.There is agreement with the maximum rupture for some butnot all great earthquakes elsewhere, but the reason is not yetclear. There is general agreement with the geodetic, coastalpaleoseismic, and thermal constraints.[71] 8. The small thrust earthquakes off Oregon that areinterpreted to be on the subduction thrust are located withan accuracy of a few kilometers, near the downdip limit ofthe transition zone of the geodetic and thermal models,deeper than expected for initiation of rupture in felsic rocks.Such events are very rare and may be within seamounts assuggested by Tréhu et al. [2012] or other local mafic materialin this area, with a higher temperature limit of about 450°C.[72] 9. The earthquakes on the Nootka transform fault zoneare located with an accuracy of a few kilometers. Theyextend to about the coast of Vancouver Island, i.e., near thelandward limit of the transition zone of the geodetic models.The temperature at this location is too hot for seismic initia-tion in felsic composition rocks. However, the fault cutsmafic oceanic crust that likely has a higher temperature limitthan the normal thrust seismic temperature limit.[73] Three of the downdip limits to great earthquake ruptureagree within the uncertainties with reasonable resolutionof ±10 to ±20 km, i.e., geodetic, coastal paleoseismic, andthermal, and are close to our reference model (e.g., Figures 2and 8–10). The other limits allow the reference model limit.For major rupture, i.e., ~10m or 50% of the maximumdisplacement, the uncertainty is about ±20 km, with a smallpossibility that the constraints are systematically incorrectlyinterpreted. The limit for the landward tapering 1–2m rupture


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that can still provide strong shaking is less well estimated but ap-pears to be 25–50 km farther landward. This latter limit extendssome distance inland of the coast for northern Washington.[74] The small thrust earthquakes off Oregon and the earth-quakes on the Nootka fault are interpreted to be subject to ahigher temperature limit, ~450°C, than the subduction thrust,~350°C, because they are interpreted to rupture more maficrocks, i.e., seamounts and ocean crust, respectively. Thechange in thrust seismic reflection character from a thin faultto a broad shear zone also occurs at about 450°C as expectedfrom laboratory and field data. The remaining constraints areconsistent with thermally limited rupture mainly offshore forthis especially hot subduction zone.[75] There are two constraints that provide a fairly strictmaximum rupture depth if they are correctly interpreted,albeit with significant uncertainty in position. These are theaseismic fore-arc mantle corner and the updip limit of ETSslow slip that accommodates most if not all plate con-vergence. The two constraints agree within the large uncer-tainties. They are a little deeper than the transition zonedowndip thermal limit of ~450°C. Both allow a shallowerrupture limit and are likely deeper than the actual maximumdepth for significant rupture displacement for this hotsubduction zone.[76] A surprising result from the analyses of the differentdowndip rupture limits is that there appears to be a gap of50–100 km between the downdip limit of significant rupturedisplacement in great earthquakes and the region of slow slipof ETS. The analysis of slow slip from continuous GPS databy Holtkamp and Brudzinski [2010] supports this conclu-sion., i.e., a gap between the great earthquake interseismiclocked zone and the deeper inter-ETS event locked zone thatslips in ETS slow slip events. It is not yet clear how and whenthis intermediate depth zone moves. Does it move continu-ously with slow creep with plate convergence accommodatedby episodic rupture displacement above and slow slip eventsbelow? As an alternative, are there smaller very frequentslow slip events that are too small to be observed geodeticallyand do not result in tremor? This intermediate depth zoneresult suggests that the ETS zone does not provide a downdiplimit of great earthquake coseismic rupture. Rupture is lim-ited to a shallower depth, although more work is needed toconfirm this conclusion. It is noted that for at least some coolsubduction zones, the downdip great earthquake rupture limitmay coincide with updip limit of ETS slow slip, since bothlimits may be at the fore-arc mantle corner.[77] The mechanisms responsible for the downdip limitof seismogenesis on subduction thrust faults globally arestill not well understood. The thermal limit of subductionthrust rupture appears to apply to hot subduction zones likeCascadia, SW Japan, and Mexico. However, there are onlya few subduction zones that are sufficiently hot for the criticaltemperatures to be reached at a shallow depth to test thishypothesis. The downdip limit of serpentinite/talc at thefore-arc mantle corner agrees with data from a number ofsubduction zones but is questioned by a number of greatearthquakes, including the Sumatra and NE Japan TohokuM9 earthquakes. Both these earthquakes have been con-cluded to have some rupture deeper than the fore-arc mantlecorner. However, in both cases, this corner is inferred to beexceptionally shallow, 20–25 km, and it is recognized thatthe depth of the fore-arc Moho tends to be especially difficult

to determine. The location of the fore-arc mantle corner is animportant need for future study. There also is the question ofsmaller thrust earthquakes that occur at depths greater thanthe main rupture displacements. They may represent ruptureof seamounts or other mafic crustal material with a higherseismogenic temperature than the main thrust.

12. Seismic Ground Motion From LandwardLimits of Rupture

[78] For the earthquake shaking at near-coastal cities,models for ground motion with landward distance are needed[e.g., Gregor et al., 2002; Olsen et al., 2008; Atkinson andMacias, 2009]. The landward limit of rupture is one of themost important control of shaking inland for these models.The lack of significant historical earthquakes in Cascadiameans it is very difficult to test existing ground motionattenuation models. There are several significant problemsin converting the available landward limit constraints forCascadia to ground motion:[79] 1. The ground motion attenuation with distance rela-tions so far presented used poorly defined distances fromthe earthquake rupture. Usually, they have used the landwardrupture extent for great earthquakes elsewhere as defined byaftershocks or by seismic rupture models which have alarge uncertainty.[80] 2. The nature of rupture displacement landward taper,i.e., the transition zone, may be very important, since, asdiscussed earlier, 1–2 m of displacement can produce verydamaging ground motions. This landward taper is not incor-porated in most ground motion models. Most great earth-quake rupture inversions and models have a transition zonewith decreasing rupture displacement landward, correspond-ing roughly with the estimated tapering of displacement inobserved recent great events. For Cascadia with approxi-mately 20m average full rupture, the 50% or 10m rupturelimit can be estimated with reasonable assurance to withinabout ±20 km, but the position of the taper to ~2m (~10%)is less accurately defined 25–50 km farther landward.[81] 3. Within the great earthquake rupture zone, there maybe a frequency dependence of seismic energy that varies withdowndip distance. More damaging high frequencies may begenerated near the landward limit of rupture, even thoughthe rupture displacement may be small there.

[82] Acknowledgments. The author would like to thank scientists at thePacific Geoscience Centre, Geological Survey of Canada for many useful dis-cussions, and a number of authors whose work is discussed in this article forclarifications and additional information.

ReferencesAbers, G. A., L. S. MacKenzie, S. Rondenay, Z. Zhang, A. G. Wech, andK. C. Creager (2009), Imaging the source region of Cascadia tremor andintermediate-depth earthquakes, Geology, 37, 1119–1122.

Atkinson, G. M., and M. Macias (2009), Predicted ground motions for greatinterface earthquakes in the Cascadia subduction zone, Bull. Seismol. Soc.Am., 99, 1552–1578, doi:10.1785/0120080147.

Atwater, B. F. (1987), Evidence for great Holocene earthquakes along theouter coast of Washington State, Science, 236, 942–944.

Atwater, B. F., and G. B. Griggs (2012), Deep-sea turbidites as guides toHolocene earthquake history at the Cascadia Subduction Zone—Alternative views for a seismic-hazard workshop: U.S. GeologicalSurvey Open-File Report 2012–1043, 58 p., http://pubs.usgs.gov/of/2012/1043/.


5545

Audet, P., M. G. Bostock, N. I. Christensen, and S. M. Peacock (2009),Seismic evidence for overpressured subducted oceanic crust andmegathrust fault sealing, Nature, 457, 76–78.

Audet, P., M. G. Bostock, D. C. Boyarko, M. R. Brudzinski, and R. M. Allen(2010), Slab morphology in the Cascadia fore arc and its relation toepisodic tremor and slip, J. Geophys. Res., 115, B00A16, doi:10.1029/2008JB006053.

Bachmann, R., O. Oncken, J. Glodny, W. Seifert, V. Georgieva, andM. Sudo (2009), Exposed plate interface in the European Alps revealsfabric styles and gradients related to an ancient seismogenic couplingzone, J. Geophys. Res., 114, B05402, doi:10.1029/2008JB005927.

Bartlow, N. M., S. Miyazaki, A. M. Bradley, and P. Segall (2011), Space-time correlation of slip and tremor during the 2009 Cascadia slow slipevent, Geophys. Res. Lett., 38, L18309, doi:10.1029/2011GL048714.

Beroza, G. C., and S. Ide (2011), Slow earthquakes and nonvolcanic tremor,Annu. Rev. Earth Planet. Sci., 39, 271–296, doi:10.1146/annurev-earth-040809-152531.

Bilek, S. (2007), Influence of subducting topography on earthquake rupture,in The Seismogenic Zone of Subduction Thrust Faults, edited by T. Dixonand C. Moore, pp. 123–146, Columbia Univ. Press, New York.

Bilek, S. L., and T. Lay (1999), Rigidity variations with depth along inter-polate megathrust faults in subduction zones, Nature, 400, 443–446.

Blakely, R. J., T. M. Brocher, and R. E. Wells (2005), Subduction-zonemagnetic anomalies and implications for hydrated forearc mantle,Geology, 33, 445–448, doi:10.1130/G21447.1.

Blanpied, M. L., D. A. Lockner, and J. D. Byerlee (1991), Fault stabilityinferred from granite sliding experiments at hydrothermal conditions,Geophys. Res. Lett., 18, 609–612, doi:10.1029/91GL00469.

Blanpied, M. L., D. A. Lockner, and J. D. Byerlee (1995), Frictional slip ofgranite at hydrothermal conditions, J. Geophys. Res., 100, 13,045–13,064,doi:10.1029/95JB00862.

Bonner, J. L., D. D. Blackwell, and E. T. Herrin (2003), Thermal con-straints on earthquake depths in California, Bull. Seismol. Soc. Am., 93,2333–2354.

Bostock, M. G., R. D. Hyndman, S. Rondenay, and S. M. Peacock (2002),An inverted continental Moho and serpentinization of the forearc mantle,Nature, 417, 536–538, doi:10.1038/417536a.

Braunmiller, J., and J. Nábělek (2002), Seismotectonics of the Explorerregion, J. Geophys. Res., 107(B10), 2208, doi:10.1029/2001JB000220.

Brocher, T. M., T. Parsons, A. M. Tréhu, C. M. Snelson, and M. A. Fisher(2003), Seismic evidence for widespread serpentinized forearc uppermantle along the Cascadia margin, Geology, 31, 267–270, doi:10.1130/0091-7613(2003)031<0267:SEFWSF>2.0.CO;2.

Burgette, R. J., R. J. Weldon II, and D. A. Schmidt (2009), Interseismic upliftrates for western Oregon and along-strike variation in locking on theCascadia subduction zone, J. Geophys. Res., 114, B01408, doi:10.1029/2008JB005679.

Bürgmann, R., and G. Dresen (2008), Rheology of the lower crust andupper mantle: Evidence from rock mechanics, geodesy, and fieldobservations, Annu. Rev. Earth Planet. Sci., 36, 531–567, doi:10.1146/annurev.earth.36.031207.124326.

Calvert, A. J., L. A. Preston, and A. M. Farahbod (2011), Sedimentaryunderplating at the Cascadia mantle-wedge corner revealed by seismicimaging, Nat. Geosci., 48, 545–548.

Chen, W., and P. Molnar (1983), Focal depths of intracontinental andinterplate earthquakes and its implications for the thermal and mechanicalproperties of the lithosphere, J. Geophys. Res., 88, 4183–4214.

Chlieh, M., J.-P. Avouac, V. Hjorleifsdottir, T.-R. A. Song, C. Ji,K. S. A. Sladen, H. Hebert, L. Prawirodirdjo, Y. Bock, and J. Galetzka(2007), Coseismic slip and afterslip of the great Mw 9.15 Sumatra-Andaman earthquake of 2004, Bull. Seismol. Soc. Am., 97, 152–173,doi:10.1785/0120050631.

Chlieh, M., J. P. Avouac, K. Sieh, D. H. Natawidjaja, and J. Galetzka (2008),Heterogeneous coupling of the Sumatran megathrust constrained bygeodetic and paleogeodetic measurements, J. Geophys. Res., 113, B05305,doi:10.1029/2007JB004981.

Clague, J. J., and P. T. Bobrowsky (1994), Evidence for a large earthquakeand tsunami 100–400 years ago on western Vancouver Island, BritishColumbia, Quat. Res., 41, 176–184, doi:10.1006/qres.1994.1019.

Clowes, R.M.,M. T. Brandon, A. G. Green, C. J. Yorath, A. Sutherland Brown,E. R. Kanasewich, and C. Spencer (1987), LITHOPROBE—SouthernVancouver Island: Cenozoic subduction complex imaged by deep seismicreflections, Can. J. Earth Sci., 24, 31–51.

Cole, J., B. Hacker, L. Ratschbacher, J. Dolan, G. Seward, E. Frost, andW. Frank (2007), Localized ductile shear below the seismogenic zone:Structural analysis of an exhumed strike-slip fault, Austrian Alps,J. Geophys. Res., 112, B12304, doi:10.1029/2007JB004975.

Cozzens, B. D., and G. A. Spinelli (2012), A wider seismogenic zone atCascadia due to fluid circulation in subducting oceanic crust,Geology, 40,899–902, doi:10.1130/G33019.1.

Currie, C. A., K. Wang, R. D. Hyndman, and J. He (2004), The thermal effectsof steady-state slab-driven mantle flow above a subducting plate: TheCascadia subduction zone and backarc, Earth Planet. Sci. Lett., 223, 35–48.

Delouis, B., J. M. Nocquet, and M. Vallée (2010), Slip distribution of theFebruary 27, 2010 Mw=8.8 Maule earthquake, central Chile, from staticand high-rate GPS, InSAR, and broadband teleseismic data, Geophys.Res. Lett., 37, L17305, doi:10.1029/2010GL043899.

Dessa, J.-X., F. Klingelhoefer, D. Graindorge, C. Andr, H. Permana,M.-A. Gutscher, A. Chauhan, and S. C. Singh (2009), Megathrust earth-quakes can nucleate in the forearc mantle: Evidence from the 2004Sumatra event, Geology, 37, 659–662, doi:10.1130/Gs5653A.1.

Dragert, H., and R. D. Hyndman (1995), Continuous GPS monitoring ofelastic strain in the northern Cascadia subduction zone, Geophys. Res.Lett., 22, 755–758.

Dragert, H., and K. Wang (2011), Temporal evolution of an episodic tremorand slip event along the northern Cascadia margin, J. Geophys. Res., 116,B12406, doi:10.1029/2011JB008609.

Escartın, J., G. Hirth, and B. Evans (2001), Strength of slightly serpentinizedperidotites: Implications for the tectonics of oceanic lithosphere,Geology, 29,1023–1026, doi:10.1130/0091-7613(2001)029<1023:SOSSPI>2.0.CO;2.

Fagereng, A. (2011), Geology of the seismogenic subduction thrust inter-face, Geol. Soc. London Spec. Publ., 359, 55–76, doi:10.1144/SP359.4.

Flück, P., R. D. Hyndman, and K. Wang (1997), Three-dimensional disloca-tion model for great earthquakes of the Cascadia subduction zone,J. Geophys. Res., 102, 20,539–20,550, doi:10.1029/97JB01642.

Frankel, A. D. (2011), Summary of November 2010 meeting to evaluateturbidite data for constraining the recurrence parameters of greatCascadia earthquakes for the update of the national seismic hazard maps,U.S. Geologic Survey Open-File Report 2011–1310, 13 p., http://pubs.usgs.gov/of/2011/1310/.

Fujiwara, T., S. Kodaira, T. No, Y. Kaiho, N. Takahashi, and Y. Kaneda(2011), The 2011 Tohoku-Oki earthquake: Displacement reaching thetrench axis, Science, 334, doi:10.1126/science.1211554.

Fuller, C. W., S. D. Willett, and M. T. Brandon (2006), Formation of forearcbasins and their influence on subduction zone earthquakes, Geology, 34,65–68, doi:10.1130/G21828.1.

Ghosh, A., J. E. Vidale, and K. C. Creager (2012), Tremor asperities in thetransition zone control evolution of slow earthquakes, J. Geophys. Res.,117, B10301, doi:10.1029/2012JB009249.

Goldfinger, C. (2011), Submarine paleoseismology based on turbiditerecords, Annu. Rev. Mar. Sci., 3, 35–66, doi:10.1146/annurev-marine-120709-142852.

Goldfinger, C., et al. (2012), Turbidite event history: Methods and implica-tions for Holocene paleoseismicity of the Cascadia subduction zone, inUSGS Professional Paper 1661-F, pp. 184, U.S. Geol. Surv., Reston, VA.

Gomberg, J., and the Cascadia 2007 and Beyond Working Group (2010),Slow-slip phenomena in Cascadia from 2007 and beyond: A review,Geol. Soc. Am. Bull., 122, 963–978, doi:10.1130/B30287.1.

Gregor, N., W. Silva, I. Wong, and R. Youngs (2002), Ground-motion atten-uation relationships for Cascadia subduction zone megathrust earthquakesbased on a stochastic finite-fault model, Bull. Seismol. Soc. Am., 92,1923–1932.

Guilbault, J. P., J. J. Clague, and M. Lapointe (1996), Foraminifera evidencefor the amount of coseismic subsidence during a late Holocene earthquakeon Vancouver Island, west coast of Canada, Quat. Sci. Rev., 5, 913–937,doi:10.1016/S0277-3791(96)00058-3.

Gutscher, M.-A., and S. M. Peacock (2003), Thermal models of flat subduc-tion and the rupture zone of great subduction earthquakes, J. Geophys.Res., 108(B1), 2009, doi:10.1029/2001JB000787.

Haberland, C., A. Rietbrock, D. Lange, K. Bataille, and S. Hofmann (2006),Interaction between forearc and oceanic plate at the south-central Chileanmargin as seen in local seismic data, Geophys. Res. Lett., 33, L23302,doi:10.1029/2006GL028189.

Hansen, R. T. J., M. G. Bostock, and N. I. Christensen (2012), Nature of thelow velocity zone in Cascadia from receiver function waveform inversion,Earth Planet. Sci. Lett., 337–338, 25–38, doi:10.1016/j.epsl.2012.05.031.

Hawkes, A. D., B. P. Horton, A. R. Nelson, C. H. Vane, and Y. Sawai(2011), Coastal subsidence in Oregon, USA, during the giant Cascadiaearthquake of AD 1700, Quat. Sci. Rev., 30, 364–376, doi:10.1016/j.quascirev.2010.11.017.

Hawthorne, J. C., and A. M. Rubin (2013), Short-time scale correlationbetween slow slip and tremor in Cascadia, J. Geophys. Res. Solid Earth,118, 1316–1329, doi:10.1002/jgrb.50103.

He, C., Z. Wang, and W. Yao (2007), Frictional sliding of gabbro gougeunder hydrothermal conditions, Tectonophysics, 445, 353–362, doi:10.1016/j.tecto.2007.09.008.

Henton, J. A., H. Dragert, K. Wang, H. Kao, and A. Lambert (2010),Investigating northern Cascadia ETS processes with absolute gravity anddeformation measurements near Port Renfrew, British Columbia, Am.Geophys. Un., Fall Meeting 2010, San Francisco, Abstract #S23A-2101.


5546

Hetland, E. A., and M. Simons (2010), Post-seismic and interseismic faultcreep II: transient creep and interseismic stress shadows on megathrusts,Geophys. J. Int., 181, 99–112, doi:10.1111/j.1365-246X.2009.04482.x.

Heuret, A., S. Lallemand, F. Funiciello, C. Piromallo, and C. Faccenna(2011), Physical characteristics of subduction interface type seismogeniczones revisited, Geochem. Geophys. Geosyst., 12, Q01004, doi:10.1029/2010GC003230.

Hino, R., S. Ito, H. Shiobara, H. Shirnarnura, T. Sato, T. Kanazawa,J. Kasahara, and A. Hasegawa (2000), Aftershock distribution of the1994 Sanriku-oki earthquake (Mw 7.7) revealed by ocean bottom seismo-graphic observation, J. Geophys. Res., 105, 21,697–21,710.

Hippchen, S., and R. D. Hyndman (2008), Thermal and structural models ofthe Sumatra subduction zone: Implications for the megathrust seismogeniczone, J. Geophys. Res., 113, B12103, doi:10.1029/2008JB005698.

Holtkamp, S., and M. R. Brudzinski (2010), Determination of slow slipepisodes and strain accumulation along the Cascadia margin, J. Geophys.Res., 115, B00A17, doi:10.1029/2008JB006058.

Hu, Y., and K. Wang (2012), Spherical-Earth finite element model of short-term postseismic deformation following the 2004 Sumatra earthquake,J. Geophys. Res., 117, B05404, doi:10.1029/2012JB009153.

Hu, Y., K.Wang, J. He, J. Klotz, and G. Khazaradze (2004), Three-dimensionalviscoelastic finite element model for postseismic deformation of the great1960 Chile earthquake, J. Geophys. Res., 109, B12403, doi:10.1029/2004JB003163.

Hyndman, R. D. (2007), The seismogenic zone of subduction thrust faults:What we know and don’t know, in The Seismogenic Zone of SubductionThrust Faults, edited by T. Dixon and C. Moore, pp. 15–49, ColumbiaUniv. Press, New York.

Hyndman, R. D., and S. M. Peacock (2003), Serpentinization of the forearcmantle, Earth Planet. Sci. Lett., 212, 417–432.

Hyndman, R. D., and G. C. Rogers (1981), Seismicity surveys with oceanbottom seismographs off western Canada, J. Geophys. Res., 86, 3867–3880,doi:10.1029/JB086iB05p03867.

Hyndman, R. D., and G. C. Rogers (2010), Great earthquakes on Canada’swest coast: A review, Can. J. Earth Sci., 47, 801–820.

Hyndman, R. D., and K. Wang (1993), Thermal constraints on the zone ofmajor thrust earthquake failure: The Cascadia subduction zone, J. Geophys.Res., 98, 2039–2060.

Hyndman, R. D., and K. Wang (1995), Current deformation and thermalconstraints on the zone of potential great earthquakes on the Cascadiasubduction thrust, J. Geophys. Res., 100, 22,133–22,154.

Hyndman, R. D., R. P. Riddihough, and R. Herzer (1979), The Nootka faultzone—A new plate boundary off western Canada, Geophys. J. R. Astron.Soc., 58, 667–683.

Hyndman, R. D., K. Wang, T. Yuan, and G. D. Spence (1993), Tectonicsediment thickening, fluid expulsion, and the thermal regime of subduc-tion zone accretionary prisms: The Cascadia margin off VancouverIsland, J. Geophys. Res., 98, 21,865–21,876.

Hyndman, R. D., K. Wang, and M. Yamano (1995), Thermal constraints onthe seismogenic portion of the southwestern Japan subduction thrust,J. Geophys. Res., 100, 15,373–15,392, doi:10.1029/95JB00153.

Hyndman, R. D., M. Yamano, and D. A. Oleskevich (1997), Theseismogenic zone of subduction thrust faults, Island Arc, 6, 244–260.

Ito, K. (1990), Regional variations of the cutoff depth of seismicity in thecrust and their relation to heat flow and large inland earthquakes,J. Phys. Earth, 38, 5223–5250.

Kanda, R. V. S., and M. Simons (2010), An elastic plate model forinterseismic deformation in subduction zones, J. Geophys. Res., 115,B03405, doi:10.1029/2009JB006611.

Kao, H., S.-J. Shan, H. Dragert, and G. Rogers (2009), Northern Cascadiaepisodic tremor and slip: A decade of tremor observations from 1997 to2007, J. Geophys. Res., 114, B00A12, doi:10.1029/2008JB006046.

Katayama, I., T. Terada, K. Okazaki, and W. Tanikawa (2012), Episodictremor and slow slip potentially linked to permeability contrasts at theMoho, Nat. Geosci., 5, 731–734, doi:10.1038/ngeo1559.

Kirby, S. H. (1985), Rock mechanics observations pertinent to the rheol-ogy of the continental lithosphere and the localization of strain alongshear zones, Tectonophysics, 119, 1–27, doi:10.1016/0040-1951(85)90030-7.

Koper, K. D., A. R. Hutko, T. Lay, C. J. Ammon, and H. Kanamori (2011),Frequency-dependent rupture process of the 2011 Mw 9.0 Tohoku earth-quake: Comparison of short-period P wave backprojection images andbroadband seismic rupture models, Earth Planets Space, 58, 1–4.

Kurahashi, S., and K. Irikura (2011), Source model for generating strongground motions during the 2011 off the Pacific coast of TohokuEarthquake, Earth, Planet., Space, 63, 571–576.

Lay, T., H. Kanamori, C. J. Ammon, K. D. Koper, A. R. Hutko, L. Ye, H. Yue,and T. M. Rushing (2012), Depth-varying rupture properties of subductionzone megathrust faults, J. Geophys. Res., 117, B04311, doi:10.1029/2011JB009133.

Leonard, L. J., R. D. Hyndman, and S.Mazzotti (2004), Coseismic subsidencein the 1700 great Cascadia earthquake: Coastal estimates versus elasticdislocation models, Geol. Soc. Am. Bull., 116, 655–670, doi:10.1130/B25369.1.

Leonard, L. J., C. A. Currie, S. Mazzotti, and R. D. Hyndman (2010),Rupture area and displacement of past Cascadia great earthquakes fromcoastal coseismic subsidence, Geol. Soc. Am. Bull., 122, 2079–2096,doi:10.1130/B30108.1.

Llenos, A. L., and J. J. McGuire (2007), Influence of fore-arc structure on theextent of great subduction zone earthquakes, J. Geophys. Res., 112,B09301, doi:10.1029/2007JB004944.

Loveless, J. P., and B. J. Meade (2011), Spatial correlation of interseismiccoupling and coseismic rupture extent of the 2011 MW = 9.0 Tohoku-okiearthquake, Geophys. Res. Lett., 38, L17306, doi:10.1029/2011GL048561.

Marone, C., and D. Saffer (2007), Fault friction and the upper transition fromseismic to aseismic faulting, in The Seismogenic Zone of SubductionThrust Faults, edited by T. Dixon and C. Moore, pp. 346–369, ColumbiaUniv. Press, New York.

Mazzotti, S., H. Dragert, J. Henton, M. Schmidt, R. Hyndman, T. James,Y. Lu, and M. Craymer (2003), Current tectonics of northern Cascadiafrom a decade of GPS measurements, J. Geophys. Res., 108(B12), 2554,doi:10.1029/2003JB002653.

Mazzotti, S., A. Lambert, N. Courtier, L. Nykolaishen, and H. Dragert(2007), Crustal uplift and sea level rise in northern Cascadia from GPS,absolute gravity, and tide gauge data, Geophys. Res. Lett., 34, L15306,doi:10.1029/2007GL030283.

McCaffrey, R. (2009), Time-dependent inversion of three-component con-tinuous GPS for steady and transient sources in northern Cascadia,Geophys. Res. Lett., 36, doi:10.1029/2008GL036784.

McCaffrey, R., and D. Schmidt (2012), Models for the eastern edge ofrupture zones of great Cascadia earthquakes: Evidence from GPS anduplift data In, Workshop on Update of Pacific Northwest Portion of theU.S. National Seismic Hazard Maps (NSHMs), March 21–22, Univ. ofWashington, Seattle, Wash.

McCaffrey, R., A. I. Qamar, R. W. King, R. Wells, G. Khazaradze,C. A. Williams, C. Stevens, J. J. Vollick, and P. C. Zwick (2007),Fault locking, block rotation and crustal deformation in the PacificNorthwest, Geophys. J. Int., 169, 1315–1340, doi:10.1111/j.1365-246X.2007.03371.x.

McCaffrey, R., L. M.Wallace, and J. Beavan (2008), Slow slip and frictionaltransition at low temperature at the Hikurangi subduction zone, Nat.Geosci., 1, 316–320, doi:10.1038/ngeo178.

McCaffrey, R., R. W. King, S. J. Payne, and M. Lancaster (2013), Activetectonics of northwestern US inferred from GPS-derived surface velocities,J. Geophys. Res. Solid Earth, 118, 709–723, doi:10.1029/2012JB009473.

McCrory, P. A., J. L. Blaire, D. H. Oppenheimer, and S. R. Walter (2006),Depth to the Juan de Fuca slab beneath the Cascadia subduction margin—A 3-D model for sorting earthquakes, U.S. Geol. Surv. Data Series91, Version 1.2, 13.

McCrory, P. A., J. L. Blair, F. Waldhauser, and D. H. Oppenheimer (2012a),Juan de Fuca slab geometry and its relation to Wadati-Benioff zoneseismicity, J. Geophys. Res., 117, B09306, doi:10.1029/2012JB009407.

McCrory, P. A., R. D. Hyndman, and J. L. Blair (2012b), Is there a relation-ship between Cascadia non-volcanic tremor and the forearc mantle wedge,and the downdip limit of seismogenic rupture? Abstract 1497823, FallMeeting, Am. Geophys. Un., December 3, 2012.

McKenna, J. R., and D. D. Blackwell (2002), Heat sources in subductionzones: Implications for slab seismicity and arc volcanism, in TheCascadia Subduction Zone and Related Subduction Systems, SeismicStructure, Intraslab Earthquakes and Processes, and EarthquakeHazards, edited by S. Kirby, K. Wang, and S. Dunlop, pp. 127–131,U.S. Geol. Surv. Open-File Report 02–328, Geol. Surv. Canada OpenFile 4350.

Miller, M. M., D. J. Johnson, C.M. Rubin, H. Dragert, K.Wang, A. Qamar, andC. Goldfinger (2001), GPS-determination of along-strike variation in Cascadiamargin kinematics: Implications for relative plate motion, subduction zonecoupling, and permanent deformation, Tectonics, 20, 161–176.

Mitchell, C. E., P. Vincent, R. J. Weldon, and M. A. Richards (1994),Present-day vertical deformation of the Cascadia margin, PacificNorthwest, United States, J. Geophys. Res., 99, 12,257–12,277.

Moore, D. E., and M. J. Rymer (2007), Talc-bearing serpentinite and thecreeping section of the San Andreas fault, Nature, 448, 795–797,doi:10.1038/nature06064.

Moore, J., and D. Saffer (2001), Updip limit of the seismogenic zone beneaththe accretionary prism of southwest Japan: An effect of diagenetic to low-grade metamorphic processes and increasing effective stress, Geology, 29,183–186, doi:10.1130/0091-7613(2001)029.

Moore, D. E., D. A. Lockner, S. Ma, R. Summers, and J. D. Byerlee (1997),Strengths of serpentinite gouges at elevated temperatures, J. Geophys.Res., 102, 14,787–14,801.


5547

Moreno, M., M. Rosenau, and O. Oncken (2010), Maule earthquake slipcorrelates with pre-seismic locking of Andean subduction zone, Nature,467, doi:10.1038/nature09349.

Nedimovic, M. R., R. D. Hyndman, K. Ramachandran, and G. D. Spence(2003), Reflection signature of seismic and aseismic slip on the northernCascadia subduction interface, Nature, 24, 416–420.

Obara, K., H. Hitoshi, F. Yamamizu, and K. Kasahara (2004), Episodicslow slip events accompanied by non-volcanic tremors in southwestJapan subduction zone, Geophys. Res. Lett., 31, L23602, doi:10.1029/2004GL020848.

Oleskevich, D. A., R. D. Hyndman, and K. Wang (1999), The updip anddowndip limits to great subduction earthquakes: Thermal and structuralmodels of Cascadia, south Alaska, SW Japan, and Chile, J. Geophys.Res., 104, 14,965–14,991.

Olsen, K. B., W. J. Stephenson, and A. Geisselmeyer (2008), 3D crustalstructure and long-period ground motions from a M 9.0 megathrust earth-quake in the Pacific Northwest region, J. Seismol., 12, 145–151.

Pacheco, J. F., L. R. Sykes, and C. H. Scholz (1993), Nature of seismiccoupling along simple plate boundaries of the subduction type, J. Geophys.Res., 98, 4,133–14,159, doi:10.1029/93JB00349.

Peacock, S. (2009), Thermal andmetamorphic environment of subduction zoneepisodic tremor and slip, J. Geophys. Res., 114, B00A07, doi:10.1029/2008JB005978.

Peacock, S., and R. D. Hyndman (1999), Hydrous minerals in the mantlewedge and the maximum depth of subduction earthquakes, Geophys.Res. Lett., 26, 2517–2520.

Peterson, C. D., K. M. Cruikshank, M. E. Darienzo, G. C. Wessen,V. L. Butler, and S. L. Sterling (2012), Coseismic subsidence andpaleotsunami run-up records from latest Holocene deposits in theWaatch Valley, Neah Bay, northwest Washington USA: Links to greatearthquakes in the northern Cascadia Margin, J. Coastal Res., 29,157–172.

Pollitz, F. F., P. McCrory, D. Wilson, J. Svarc, C. Puskas, and R. B. Smith(2010), Viscoelastic-cycle model of interseismic deformation in the north-western United States, Geophys. J. Int., 181, 665–696, doi:10.1111/j.1365-246X.2010.04546.x.

Preston, L. A., K. C. Creager, R. S. Crosson, T. M. Brocher, and A.M. Tréhu(2003), Intraslab earthquakes: Dehydration of the Cascadia slab, Science,302, 1197–1200.

Ramachandran, K., and R. D. Hyndman (2011), The fate of fluids releasedfrom subducting slab in northern Cascadia, Solid Earth Discuss., 3,943–962.

Reynard, B. (2013), Serpentine in active subduction zones, Lithos, 178,171–185, doi:10.1016/j.lithos.2012.10.012.

Riddihough, R. P. (1984), Recent movements of the Juan de Fuca platesystem, J. Geophys. Res., 89, 6980–6994.

Rogers, G., and H. Dragert (2003), Episodic tremor and slip on the Cascadiasubduction zone: The chatter of silent slip, Science, 300, 1942–1943,doi:10.1126/science.1084783.

Ruegg, J. C., A. Rudloff, C. Vigny, R. Madariaga, J. B. de Chabalier,J. Campos, E. Kausel, S. Barrientos, and D. Dimitrov (2009), Interseismicstrain accumulation measured by GPS in the seismic gap betweenConstitución and Concepción in Chile, Phys. Earth Planet. Int., 175,78–85, doi:10.1016/j.pepi.2008.02.015.

Ruff, L. J., and B. W. Tichelaar (1996), What controls the seismogenic plateinterface in subduction zones?, in Subduction Top to Bottom, Geophys.Monogr. Ser., vol. 96, edited by G. E. Bebout, W. Scholl, H. Kirby, andP. Platt, pp. 105–112, AGU, Washington, D.C.

Sagiya, T., and W. Thatcher (1999), Coseismic slip resolution along a plateboundary megathrust: The Nankai Trough, southwest Japan, J. Geophys.Res., 104, 1111–1129.

Savage, J. C. (1983), A dislocation model of strain accumulation and releaseat a subduction zone, J. Geophys. Res., 88, 4984–4996, doi:10.1029/JB088iB06p04984.

Schmidt, D. A., and H. Gao (2010), Source parameters and time-dependentslip distributions of slow slip events on the Cascadia subduction zonefrom 1998 to 2008, J. Geophys. Res., 115, B00A18, doi:10.1029/2008JB006045.

Scholl, D. W., and R. von Huene (2007), Crustal recycling at modernsubduction zones applied to the past—Issues of growth and preservationof continental basement crust, mantle geochemistry, and supercontinentreconstruction, in 4-D Framework of Continental Crust, edited by R. D.Hatcher, Jr., M. P. Carlson, J. H. McBride, and J. R. Martínez Catalán,vol 200, pp. 9–32, Geol. Soc. Am. Memoir, Boulder CO doi:10.1130/2007.1200(02).

Scholz, C. H. (1992), Paradigms or small change in earthquake mechanics,Int. Geophys., 51, 505–517.

Scholz, C. H. (1998), Earthquakes and friction laws, Nature, 391, 37–42.Scholz, C. H. (2002), Mechanics of Earthquakes and Faulting, CambridgeUniv. Press, New York.

Schwartz, S. Y., and J. M. Rokosky (2007), Slow slip events and seismictremor at circum-Pacific subduction zones, Rev. Geophys., 45, B00A07,doi:10.1029/2006RG000208.

Seno, T. (2005), Variation of downdip limit of the seismogenic zonenear the Japanese islands: Implications for the serpentinization mech-anism of the forearc mantle wedge, Earth Planet. Sci. Lett., 231,249–262.

Shearer, P., and R. Burgmann (2010), Lessons learned from the 2004Sumatra-Andaman megathrust rupture, Annu. Rev. Earth Planet. Sci., 38,103–131.

Shillington, D. J., et al. (2011), Constraints on the Aleutian subduc-tion zone from the Shumagin gap to Kodiak asperity from newMCS and OBS data of the ALEUT Project, Abstract T33A-2387presented at 2011 Fall Meeting, AGU, San Francisco, Calif., 4–9December, 2011.

Sibson, R. H. (1977), Fault rocks and fault mechanisms, J. Geol. Soc., 133,191–213, doi:10.1144/gsjgs.133.3.0191.

Sibson, R. H. (1983), Continental fault structure and the shallow earthquakesource, J. Geol. Soc., 140, 741–767, doi:10.1144/gsjgs.140.5.0741.

Simoes, M., J. P. Avouac, R. Cattin, and P. Henry (2004), The Sumatrasubduction zone: A case for a locked fault zone extending into the mantle,J. Geophys. Res., 109, B10402, doi:10.1029/2003JB002958.

Simons, M., et al. (2011), The 2011 magnitude 9.0 Tohoku-Oki earthquake:Mosaicking the megathrust from seconds to centuries, Science, 332,1421–1425, doi:10.1126/science.1206731.

Song, T.-R., and M. Simons (2003), Large trench-parallel gravity varia-tions predict seismogenic behaviour in subduction zones, Science, 301,630–633.

Spinelli, G. A., and K. Wang (2008), Effects of fluid circulation insubducting crust on Nankai margin seismogenic zone temperatures,Geology, 36, 887–890, doi:10.1130/G25145A.1.

Stern, C. R. (2011), Subduction erosion: Rates, mechanisms, and its role inarc magmatism and the evolution of the continental crust and mantle,Gondwana Res., 20, 284–308, doi:10.1016/j.gr.2011.03.006.

Suito, H., and J. T. Freymueller (2009), A viscoelastic and afterslip postseismicdeformation model for the 1964 Alaska earthquake, J. Geophys. Res., 114,B11404, doi:10.1029/2008JB005954.

Takahashi, N., S. Kodaira, T. Tsuru, J.-O. Park, Y. Kaneda,K. Suyehiro, H. Kinoshita, S. Abe, M. Nishino, and R. Hino(2004), Seismic structure and seismogenesis off Sanriku region,northeastern Japan, Geophys. J. Int., 159, 129–145, doi:10.1111/j.1365-246X.2004.02350.x.

Tichelaar, B. W., and L. J. Ruff (1991), Seismic coupling along the Chileansubduction zone, J. Geophys. Res., 96, 11,997–12,022, doi:10.1029/91JB00200.

Tichelaar, B. W., and L. J. Ruff (1993), Depth of seismic coupling alongsubduction zones, J. Geophys. Res., 98, 2017–2037.

Tong, X., et al. (2010), The 2010Maule, Chile earthquake: Downdip rupturelimit revealed by space geodesy, Geophys. Res. Lett., 37, L24311,doi:10.1029/2010GL045805.

Tréhu, A. M., J. Braunmiller, and J. L. Nabelek (2008), Probable low-anglethrust earthquakes on the Juan de Fuca-North America plate boundary,Geology, 36, 127–130, doi:10.1130/G24145A.1.

Tréhu, A. M., R. J. Blakely, and M. C. Williams (2012), Subductedseamounts and recent earthquakes beneath the central Cascadia forearc,Geology, 40, 103–106, doi:10.1130/G32460.1.

Tse, S. T., and J. R. Rice (1986), Crustal earthquake instability in relation tothe depth variation of frictional slip properties, J. Geophys. Res., 91,9452–9472.

Verdonck, D. (2005), An inverse dislocation model of surface deformationin western Washington, Tectonophysics, 395, 179–191.

Wada, I., and K. Wang (2009), Common depth of slab-mantle decoupling:Reconciling diversity and uniformity of subduction zones, Geochem.Geophys. Geosyst., 10, Q10009, doi:10.1029/2009GC002570.

Wada, I., K. Wang, J. He, and R. D. Hyndman (2007), Weakening of thesubduction interface and its effects on surface heat flow, slab dehydrationand mantle wedge serpentinization, J. Geophys. Res., 113, B04402,doi:10.1029/2007JB005190.

Wallace, L. M., et al. (2009), Characterizing the seismogenic zone of a majorplate boundary subduction thrust: Hikurangi Margin, New Zealand,Geochem. Geophy. Geosyst., 10, doi:10.1029/2009GC002610.

Wang, K. (2007), Elastic and viscoelastic models of crustal deformation insubduction earthquake cycles, in The Seismogenic Zone of SubductionThrust Faults, edited by T. Dixon, pp. 540–575, Columbia Univ. Press,New York.

Wang, P.-L. (2012), Rupture models of the great 1700 Cascadia earthquakebased on microfossil paleoseismic observations. MSc Thesis, Univ. ofVictoria, Victoria, B.C., 94p.

Wang, K., and S. L. Bilek (2011), Do subducting seamounts generate or stoplarge earthquakes?, Geology, 39, 819–822, doi:10.1130/G31856.1.


5548

Wang, K., and T. Dixon (2004), “Coupling” semantics and science inearthquake research, EOS, Trans. Am. Geophys. Un., 85, doi:10.1029/2004EO180005.

Wang, K., and J. He (1999), Mechanics of low-stress forearcs: Nankaiand Cascadia, J. Geophys. Res., 104, 15,191–15,205, doi:10.1029/1999JB900103.

Wang, K., T. Mulder, G. C. Rogers, and R. D. Hyndman (1995), Case forvery low coupling stress on the Cascadia subduction fault, J. Geophys.Res., 100, 12,907–12,918.

Wang, K., R. Wells, S. Mazzotti, R. D. Hyndman, and T. Sagiya (2003), A re-vised dislocation model of interseismic deformation of the Cascadia subduc-tion zone, J. Geophys. Res., 108(B1), 2026, doi:10.1029/2001JB001227.

Wang, K., H. Dragert, H. Kao, and E. Roeloffs (2008), Characterizing an“uncharacteristic” ETS event in northern Cascadia, Geophys. Res. Lett.,35, L15303, doi:10.1029/2008GL034415.

Wang, K., Y. Hu, and J. He (2012), Deformation cycles of subductionearthquakes in a viscoelastic Earth, Nature, 484, 327–332, doi:10.1038/nature11032.

Wech, A. G. (2010), Interactive tremor monitoring, Seis. Res. Lett., 81(4),664–669, doi:10.1785/gssrl.81.4.664.

Wech, A. G., K. C. Creager, and T. I. Melbourne (2009), Seismic andgeodetic constraints on Cascadia slow slip, J. Geophys. Res., 114,B10316, doi:10.1029/2008JB006090.

Wells, R. E., R. J. Blakely, Y. Sugiyama, D. W. Scholl, and P. A. Dinterman(2003), Basin-centered asperities in great subduction zone earthquakes:A link between slip, subsidence, and subduction erosion? J. Geophys.Res., 108(B10), 2507, doi:10.1029/2002JB002072.

Williams, M. C., A. M. Tréhu, and J. Braunmiller (2011), Seismicity at theCascadia plate boundary beneath the Oregon continental shelf, Bull.Seismol. Soc. Am., 101, 940–950, doi:10.1785/0120100198.

Willoughby, E. C., and R. D. Hyndman (2005), Earthquake rate, slip rate,and the effective seismic thickness for oceanic transform faults of theJuan de Fuca plate system, Geophys. J. Int., 160, 855–868.

Wong, I. G., and D. S. Chapman (1990), Deep intraplate earthquakes in thewestern United States and their relationship to lithospheric temperatures,Bull. Seismol. Soc. Am., 80, 589–599.

Yoshioka, S., K. Wang, and S. Mazzotti (2005), Interseismic lockingof the plate interface in the northern Cascadia subduction zone, in-ferred from inversion of GPS data, Earth Planet. Sci. Lett., 231,239–247.


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Downdip landward limit of Cascadia great earthquake ruptureweb.uvic.ca/~rhyndman/PublicationPDFs/HyndmanCascadia... · 2015. 2. 21. · Downdip landward limit of Cascadia great earthquake