Download Downdip landward limit of Cascadia great earthquake rupture

JOURNAL OF GEOPHYSICAL RESEARCH: SOLID EARTH, VOL. 118, 5530–5549, doi:10.1002/jgrb.50390, 2013
Downdip landward limit of Cascadia great earthquake rupture
R. D. Hyndman 1,2
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 the
Cascadia great earthquakes off western North America. This limit is a primary control for
ground motion hazard at near-coastal cities. The studies also provide information on the
physical 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) temperature
on the thrust for the seismic-aseismic transition; (4) change in thrust seismic reflection
character downdip from thin seismic to thick ductile; (5) fore-arc mantle corner aseismic
serpentinite 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 small
events on the thrust; and (9) landward limit of earthquakes on the Nootka transform fault
zone. The most reliable constraints for the limit of large rupture displacement, >10 m, are
generally just offshore in agreement with thermal control for this hot subduction zone, but
well-offshore central Oregon and near the coast of northern Washington. The limit for 1–2 m
rupture 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 landward
from the other constraints. Surprisingly, there is a downdip gap between the best other
estimates for the great earthquake rupture zone and the ETS slow slip. In this gap, plate
convergence may occur as continuous slow creep.
Citation: Hyndman, R. D. (2013), Downdip landward limit of Cascadia great earthquake rupture, J. Geophys. Res. Solid
Earth, 118, 5530–5549, doi:10.1002/jgrb.50390.
1.
Introduction
[2] The downdip landward limit of rupture for the Cascadia
subduction zone off western North America is a primary control for the ground shaking hazard from great earthquakes at
near-coastal cities such as Vancouver, BC; Seattle, WA; and
Portland, OR (Figure 1). This paper summarizes and analyzes
the constraints on this downdip limit. It focuses on Cascadia,
but many of the limits are applicable to the seismogenic zones
of other subduction zones. It also deals principally with the
main part of the 800 km long rupture zone between central
Vancouver Island south of the Nootka Fault zone and southern
Oregon, which generates ~M9 earthquakes. The more complex areas to the north and south may be involved in the long
M ~ 9 ruptures but are also large enough to generate separate
events up to M8. The areas of the Explorer plate off northern
Vancouver Island and the Gorda region off Northern California
are subducting younger ocean lithosphere than the main
1
Pacific Geoscience Centre, Geological Survey of Canada, Sidney,
British Columbia, Canada.
2
School 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 V8L4B2
Canada. ([email protected])
©2013. American Geophysical Union. All Rights Reserved.
2169-9313/13/10.1002/jgrb.50390
central region and may have a shallower downdip limit
because of the higher inferred temperatures. The main
Cascadia rupture appears to occur primarily in M9 events that
rupture much of if not the whole length of the margin based
mainly on the deep-sea record of earthquake-generated
turbidites and the coastal marsh coseismic subsidence [e.g.,
Goldfinger et al., 2012; Leonard et al., 2010] and with a generally consistent downdip limit [e.g., Leonard et al., 2010].
This is in contrast to many other subduction zones that exhibit
a wide range of rupture areas and earthquake magnitudes. The
model of Flück et al. [1997] for the thrust locked zone based
on geodetic data has been used as a reference for comparison
among the various constraints to the downdip seismogenic
limit (Figure 2). As discussed below, thermal limits to downdip
seismic behavior are similar.
[3] Understanding the downdip limit provides important
information on what controls the areas of rupture on subduction thrust faults globally and the nature of fault movement
by seismic and aseismic motions. Also critically important
to this discussion are well-recorded recent M ~ 9 great earthquakes that provide method calibration, i.e., Tohoku Japan
2011, Chile 2010, and Sumatra 2004, although these great
earthquakes occurred where older colder ocean lithosphere
is being subducted such that the seismogenic limit may not
be thermally controlled, in contrast to the inferred thermal
control for Cascadia. Relevant information comes from
geological study of deeply exposed thrust faults which show
that subduction thrusts may be very complex shear zones on
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Figure 1. Schematic of the downdip limit of Cascadia great
earthquake rupture, the closest approach to near-coastal
cities. Episodic tremor and slip (ETS) is well landward of the
rupture limit and approximately corresponds to the fore-arc
mantle corner.
a geological outcrop scale with an active layer of underthrust
sediment in a subduction melange [e.g., Bachmann et al.,
2009; Fagereng, 2011]. Thrust structure complexity with a
thick duplex structure, mainly deeper downdip than the
seismogenic zone for northern Cascadia has been proposed
by Calvert et al. [2011] using multichannel seismic reflection
data. Some of the constraints have been discussed in previous
reviews of the seismogenic zone of subduction faults [e.g.,
Tichelaar and Ruff, 1993; Pacheco et al., 1993; Oleskevich
et al., 1999; Hyndman, 2007; Hyndman and Rogers, 2010;
Heuret et al., 2011].
[4] The updip limit of rupture toward the trench also is
important for estimating earthquake magnitude and especially for modeling tsunami generation. For many other subduction zones, it is often concluded that the updip rupture
limit does not reach the trench (e.g., summary by Hyndman
et al. [1997]). The 2011 M9 Tohoku rupture did reach the
trench with large displacement [e.g., Fujiwara et al., 2011]
but at such a slow rupture speed that there was little energy
in the seismic frequencies. However, the very large tsunami
generation appears to have been generated mainly in the
updip region. Bilek and Lay [1999] argued that the rigidity
is very low and therefore the rupture speed is very low
at the toe of accretionary sedimentary prisms, increasing
downdip perhaps because of clay consolidation. If the updip
seismogenic limit is at the depth where stable-sliding clays
dehydrate, i.e., 100–150°C [e.g., Moore and Saffer, 2001;
Oleskevich et al., 1999], this limit may be near the deformation front for the high-temperature Cascadia subduction
zone. However, laboratory data by Marone and Saffer [2007]
suggest that other physical changes with depth must be
responsible for the updip part of the thrust commonly being
aseismic. The Cascadia updip limit is poorly constrained and
is not discussed further in this paper.
[5] The downdip seismogenic limit is where the fault zone
no longer exhibits frictional instability [e.g., Scholtz, 2002]
(Figure 1). Many factors have been suggested for the controls
of the downdip landward extent (e.g., discussions by Tichelaar
and 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 young
lithosphere is being underthrust like Cascadia, there appears
to be a maximum temperature limit for seismic behavior [e.g.,
Hyndman and Wang, 1993; Oleskevich et al., 1999]. For the
more common cool subduction zones, aseismic serpentinite
and talc downdip in the fore-arc mantle corner may limit the
downdip rupture [e.g., Ruff and Tichelaar, 1996; Hyndman
et 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 are
variations in the stress regime of the thrust fault zone, pore
pressure on the fault, fault roughness, and upper and lower
plate structure [e.g., Scholz, 1992; Pacheco et al., 1993;
McCaffrey et al., 2008; Llenos and McGuire, 2007; Heuret
et al., 2011; Bilek, 2007; Wallace et al., 2009; Wang and
Bilek, 2011].
[6] The location of the seismic source and interpretation of
a number of the downdip constraints depend on the dip angle
Figure 2. Fully locked (to solid black line) and linear
transition (to dashed gray line) zones based on leveling and tide
gauge geodetic data, used as a reference for comparison among
the various landward seismic limit constraints (modified after
Flück et al. [1997]). The thermal limits of seismic behavior,
350 and 450°C, from thermal models are shown with uncertainties, as given by Hyndman and Wang [1995]. The limits
from this thermal model and geodetic limits are in agreement
within the uncertainties. ETS is well landward.
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Figure 3. Schematic geometry of the Cascadia subduction
thrust, showing the simple model of locked or “full rupture”
and transition zones (adapted from Flück et al. [1997]). The
shallow thrust angle for northern Washington results in the
temperature limits being farther landward and wider locked/
rupture and transition zones.
and profile of the subducting slab (Figure 3) [e.g., Flück et al.,
1997; McCrory et al., 2012a]. The principal large-scale
variability is the region of shallow dip beneath the margin
of northern Washington and southern Vancouver Island. As
discussed below, it appears that the downdip rupture limit
may be thermally controlled. Therefore, the region with shallow dip is expected to reach a critical temperature farther
landward and have a wider rupture zone than the steeper dip regions to the south and north. Most but not all past great events
appear to have ruptured all of the main part of the Cascadia
subduction fault [e.g., Goldfinger, 2011; Goldfinger et al.,
2012; Leonard et al., 2004; 2010] but there have been some
that ruptured shorter extents along the subduction fault. The
distribution of event sizes and margin-parallel extents is the
subject of ongoing research [e.g., Atwater and Griggs, 2012;
Frankel, 2011].
[7] We discuss nine constraints to the landward limit of
Cascadia great earthquake rupture. These nine constraints
provide a better understanding of the landward seismogenic
downdip limit and the controlling processes than are available for many other subduction zones, even those which have
had instrumentally recorded great events. Each of the nine
constraints requires different assumptions and has different
uncertainties; we seek a rupture limit that is consistent with
all of the constraints. There is a considerable literature on
aspects of some of these constraints. An extensive reference
list is given, but far from complete. Emphasis is on those
papers that provide general results and those that have
reference lists themselves that lead to papers with more
detailed information.
[8] An important issue is the nature of the landward tapering of rupture displacement, i.e., the “transition zone”
(Figure 3). This transition is important since 1–2 m of
displacement can produce strong ground motions in the main
damaging frequency range, as now well recognized, for example, 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 a
number of other subduction zones [e.g., Lay et al., 2012]
where much of the energy in the damaging frequencies was
from the downdip part of the rupture. For some constraints,
it is useful to define the area of at least 50% or ~10 m of the
estimated ~20 m average full rupture (see Figure 4). The position of 10% or 2 m, which in the geodetic dislocation models is
concluded to be tens of kilometers farther landward, is much
less constrained than 50% because the former depends on
the nature and extent of the transition from full to zero rupture
displacement. Also, the nature of the actual transition is likely
quite variable among different events. Geodetic locked zone
rupture inversions usually give smooth peak rupture and
landward decrease.
[9] The degree to which the downdip limit to rupture varies
among individual events remains uncertain. Most Cascadia
events appear to rupture a reasonably consistent landward
extent based on paleoseismicity coastal marsh subsidence
data [Leonard et al., 2010]. Although there undoubtedly is
variability, the summarized estimates are taken to be the
maximum landward rupture extent for most events, as to be
expected if there is a physical control of this limit that does
not change with time.
[10] There are three types of constraints: (a) constraints to
the currently “locked zone” on the thrust that is building
elastic strain that will be released in future great earthquakes
(constraint 1 below), (b) constraints to the landward extent
of past great earthquake rupture (constraint 2 below), and (c)
physical process/state controls on where seismicity can and
cannot occur on the thrust (constraints 3–9 below). It is important to recognize that none of the constraints directly constrain
the actual rupture limit; all require some assumptions.
[11] The constraints are:
[12] 1. The “locked/transition” zones on the thrust from
modeling geodetic observations of current deformation
(GPS, repeated leveling, tide gauges, etc.).
[13] 2. Past great earthquake rupture zone constrained by
paleoseismic coastal marsh subsidence, 1700 and earlier
events, i.e., “paleogeodesy”.
[14] 3. Seismic behavior limits from downdip temperatures,
loosely described as the “brittle-ductile transition”, approximately 350°C for where the rupture can initiate and there is
full rupture, 450°C for the landward limit of the transition zone
from full rupture to zero displacement.
Figure 4. Locked and transition zone models corresponding
approximately to the full rupture, and linear and exponential
tapering rupture (effective transition zone; adapted from
Wang et al. [2003]). The horizontal distance scale and rupture
magnitudes are approximately those for Vancouver Island and
northern Oregon.
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2. Locked/Transition Zones From Geodetic
Deformation
Figure 5. Schematic of margin deformation through the
great earthquake cycle illustrating the interseismic uplift
and coseismic subsidence of the coastal region. The
coseismic subsidence for most great events, up to 2 m, is about
500 years times the uplift rate per year. Most repeated leveling
and other vertical geodetic data are landward of the peak uplift
and most coastal paleoseismic subsidence data are near the
maximum subsidence.
[15] 4. Change in seismic reflection character from a
thrust marked by a thin sharp reflection to a thrust zone
marked by a thick set of reflectors, i.e., brittle to ductile
shear deformation.
[16] 5. Fore-arc mantle corner, landward of which there is
inferred 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-term
plate convergence.
[18] 7. Empirical association of rupture area with shelf-slope
sedimentary basins defined mainly by gravity anomalies.
[19] 8. Landward downdip limit of small thrust earthquakes
on the subduction thrust that indicate seismic behavior.
[20] 9. Landward limit of earthquakes on the Nootka
transform fault zone as it subducts beneath the margin of
Vancouver Island.
[21] Very general downdip rupture limits are provided
by estimates of tsunami wave amplitudes and runups,
although the updip extent and displacement are a more
important factor. Only the 1700 tsunami has been at all
well documented for Cascadia so this constraint is not
discussed here. We have used the Flück et al. [1997]
model with a linear downdip transition zone as a simple
reference for comparison among the different constraints. Other models with a linear transition zone give
similar results [e.g., Mazzotti et al., 2007], as do most
inversion models.
[22] The most studied and one of the strongest constraints
to the rupture zone is the definition of the locked zone on
the 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 the
earthquake cycle is elastic to a first order, i.e., elastic strain
builds up nearly linearly with time since the previous event
and is released completely by megathrust earthquakes
(Figure 5). The pattern of surface deformation for the rupture
is then close to the mirror image of the interseismic rate times
the earthquake interval (average 500–600 years for the main
Cascadia events). There are two significant limitations.
First, there may not be full “coupling”; there may be some
aseismic slip between great earthquakes (i.e., discussion by
Wang and Dixon [2004]). The current geodetic data that are
consistent with a fully locked fault [e.g., Hyndman and
Wang, 1995; McCaffrey et al., 2013] and the lack of significant interseismic thrust seismicity suggest that there is currently little interseismic slip. The thrust is fully locked after
some postseismic transients. Also, the estimated coseismic
rupture displacement in past events approximately equals the
interseismic plate convergence slip deficit based on the magnitude of coseismic coastal subsidence [Leonard et al., 2004;
2010]. The one area of uncertainty is off central Oregon where
the inferred locked/rupture zone is well offshore such that the
expected and observed interseismic uplift rate and coseismic
subsidence are small. From these two types of data, it is therefore difficult to exclude some degree of interseismic aseismic
creep motion for the thrust in this region with the partially
seismic 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 are
theoretical problems of uncertain magnitude in the elastic
backslip model for the locked zone, especially postseismic
transients [e.g., Hu et al., 2004; Wang 2007; Kanda and
Simons, 2010; Hetland and Simons, 2010; Suito and
Freymueller, 2009; Wang et al., 2012]. The downdip extent
of fully seismic rupture is expected to be somewhat seaward
of the limit for the modeled interseismic locked zone because
some release of the locked zone strain buildup, especially the
deep part, is in postseismic slip [e.g., Wang, 2007; Hu and
Wang, 2012; Wang et al., 2012], so the limit from the geodetic data may be slightly landward of the actual coseismic
full rupture limit. Records of local deformation as a function
of time for a few subduction zones that have been adequately
monitored indicate that the long-term interseismic strain
buildup is approximately linear after a sometimes substantial
postseismic transient that may depend on the earthquake
magnitude [e.g., Wang, 2007]. However, this question
requires more rigorous analysis and calibration by recent great
earthquakes elsewhere. An example from the M ~ 8 1944 and
1946 events off southwest Japan is shown in Figure 6 that
indicates good correspondence of the preseismic pattern of
deformation with that for the coseismic rupture deformation.
The downdip rupture extent of defined by dislocation modeling of leveling data was very close to that for the locked
zone 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
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Figure 6. SW Japan 1944/1946 earthquakes, comparison of
coseismic vertical displacements and interseismic vertical
motion from repeated leveling measurement, and dislocation
models with a linear transition zone (modified from Hyndman
et al. [1995]). The coseismic vertical has been inverted and
scaled by the average earthquake cycle duration. The horizontal lines at the top show the three widths of the locked (thick
line), transition (thin line), and free slip zones for the three
vertical models shown, i.e., ±10 km from best fit model (blue
curve). The good agreement between the interseismic locked
zone and the coseismic rupture zone deformations indicates
that the geodetically determined locked/transition zones are
good estimators for the rupture/transition zone.
is quite small for at least these megathrust events. There also is
good correspondence of the modeled locked zone using geodetic data and the main downdip rupture zone for the 2010
M8.8 Chile event [e.g., Delouis et al., 2010; see also Ruegg
et al., 2009], the 2004 M9 Sumatra earthquake [e.g., Shearer
and Burgmann, 2010, and references therein], and the 2011
M9 Tohoku, Japan event [e.g., Loveless and Meade, 2011].
The interseismic locked zone estimated a few decades after
the last great earthquake is likely a reasonable estimator of
the rupture zone for Cascadia, as discussed for past great earthquakes off SW Japan.
[24] Another possible effect on the use of the locked/transition zone as a rupture estimate is that the zone of ETS slow
slip downward of the main rupture zone is locked or nearly
locked on a timescale averaging about 14 months [e.g.,
Rogers and Dragert, 2003]. Strain on this short-term locked
zone is released during a period of several weeks of slow slip.
However, if longer-term average geodetic data are used, the
effect of the slow slip zone events on the estimate of the main
locked zone and the estimate of the subsequent rupture zone
is small [e.g., Holtkamp and Brudzinski, 2010].
[25] Both forward models and inversions have been
applied to the backslip modeling with similar results. The
inversion results generally have a longer taper inland as a
consequence of the smoothing required in the method, unless
the 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 increasing
dip landward [e.g., Hyndman and Wang, 1995], and Flück
et al. [1997] developed a 3-D fault model that has been
updated by McCrory et al. [2006; 2012a] and used in most
subsequent modeling. A simple fully locked zone and linear
transition zone have been used in the initial models, but a
more appropriate exponential transition zone was proposed
by Wang et al. [2003] (Figure 4).
[26] Vertical geodetic data give a quite well-defined limit to
rupture because both coseismic and interseismic vertical displacements change from upward to downward approximately
at the limit of significant rupture on the fault, i.e., hinge line
(Figures 5 and 6). Most Cascadia vertical constraints have
come from repeated high-precision leveling and long-duration
tide gauges [e.g., Mitchell et al., 1994; Hyndman and Wang,
1995; Verdonck, 2005; Burgette et al., 2009]. GPS vertical
data have lower resolution than horizontal data and are just
beginning to have useful accuracy of a few mm/yr [e.g.,
Mazzotti et al., 2007; McCaffrey and Schmidt, 2012]. Highresolution gravity and borehole strainmeters also are providing
additional strain data for comparison with dislocation models
[e.g., Mazzotti et al., 2007; Henton et al., 2010]. An example
of vertical releveling data and models depicting the observed
and predicted pattern of coastal uplift are shown in Figure 7
[Hyndman and Wang, 1995; Burgette et al., 2009]. Note that
Figure 7. Examples of leveling data and corresponding dislocation models across the Cascadia margin (modified from
Hyndman and Wang [1995]). The rates of vertical motion
derived from leveling data are compared to those predicted
by the three different width dislocation models shown; the
labels give the model locked and transition zones widths in
kilometers and the average dip angle for the locked part of
the thrust. Detailed data for the Oregon margin have been
presented by Burgette et al. [2009].
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Figure 8. (a) Cascadia GPS vectors of current deformation
(modified from McCaffrey et al. [2007]). The motion rates
include the effect of elastic strain buildup on the subduction
thrust and fore-arc crustal motions. (b) The elastic dislocation
model from inversion of horizontal and vertical data. The
red dashed line is approximately 50% locking (~25 mm/yr)
(modified from McCaffrey et al. [2013]). A locked or nearly
fully locked thrust and transition zone is indicated for the
whole margin, very similar to the Flück et al. [1997] reference model shown as black solid and dashed lines.
the maximum uplift rates of about 1– 4 mm/yr at the coast from
geodetic data, multiplied by the average interevent period of
500–600 years, predict maximum coseismic subsidences of
0.5–2 m, in general agreement with those observed in the
coastal marsh data along the coast [e.g., Leonard et al.,
2010]. The patterns of the locked and transition zones for all
of the studies are very similar to those of the Flück et al.
[1997] reference model, with successive improvements from
more accurate data and dislocation parameters. Burgette
et al. [2009] inferred an offset to a shallower locked depth over
the southern end of the coastal Siletz Terrane, but the effect
is small.
[27] Horizontal GPS data show a clear signal of the coastal
region being squeezed landward [e.g., McCaffrey et al.,
2013, Figure 8a]. The vectors also show the strong forearc crustal deformation that must be accounted for when
conducting megathrust dislocation modeling. Forward dislocation modeling that matches observations of landward
shortening has been presented by many authors with successive improvements from more extensive and accurate data
and 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 exponentially
landward to allow for the postseismic transients (Figure 4).
The recent models corrected for the effect of the Oregon
block fore-arc block motion, i.e., used the convergence
across the near-coastal margin rather than the ocean plate
relative to North America.
[28] Backslip inversions of the GPS and vertical data have
been presented by Verdonck [2005], Yoshioka et al. [2005],
and McCaffrey et al. [2007, 2013] which are generally in
agreement with the results of forward modeling. In the most
recent inversion analysis McCaffrey et al., [2013] found a
good fit of both horizontal and vertical data (Figure 8b).
They concluded that there is a somewhat wider transition
zone off central Oregon compared to southern Oregon and
northernmost California but the differences compared to previous models are quite small.
[29] From modeling both horizontal and vertical data, the
estimated full rupture (locked zone) is offshore everywhere
except near Cape Mendocino in Northern California, with
model linear transition zones extending onshore in northern
Washington and near the coast in Oregon and Vancouver
Island. The recent inversion model from McCaffrey et al.
[2013] is shown in Figure 8b. This 50% locked location is
very close to that for the Flück et al. [1997] reference model,
i.e., middle of the transition zone. It is not yet clear how much
of the small-scale variability from this type of inversion is
real or due to data uncertainty. Smooth inversion transition
zones have rupture downdip displacement of 2 m or ~10%
of full rupture near the landward limit of the linear forward
model transition zones, but this part of the models is very
poorly constrained.
3. Rupture Zone From Paleoseismic Coastal
Marsh Subsidence in Great Earthquakes
[30] The abrupt vertical deformation associated with
the most recent great earthquake in 1700 and earlier
events 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-buried
marsh surfaces are covered by sediments deposited after
successive coseismic subsidence events of up to 2 m (e.g.,
Figure 9a). Although the coseismic subsidence is approximately recovered through strain buildup in the interseismic
period, ongoing sea level rise results in progressive burial
of subsided marshes [e.g., Atwater, 1987]. On the coast of
Vancouver Island, there is ongoing land uplift faster than
sea level rise so only the 1700 buried marsh remains
[Clague and Bobrowsky, 1994]. The vertical distance
between the present marsh top and the 1700 marsh is not
the coseismic subsidence because of eustatic sea level rise,
the elastic strain buildup since that event, and tectonic vertical motions. However, a rough estimate can be obtained from
this vertical distance after correcting for these effects. Better
displacement estimates come from careful calibration of the
elevation range of a variety of coastal organisms, allowing
accurate estimation of the paleo-elevation differences
between the sediments just above and just below the top of
the buried marshes. The resolution is commonly ~0.5 m but
may be as good as ±0.3 m [e.g., Guilbault et al., 1996;
Peterson et al., 2012; Hawkes et al., 2011].
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Figure 9. (a) Example of a buried intertidal marsh off central Washington attributed to coseismic subsidence during the M9 earthquake of 1700 (photo by L. Leonard). A tsunami-driven sand sheet overlies the
buried marsh. (b) Comparison of coastal coseismic subsidence estimates for an area of the Oregon coast
plotted with distance orthogonal to the coastal trend, and corresponding dislocation models (location shown
in Figure 10), for the 1700 and pre-1700 events [after Leonard et al., 2010]. The model results are shown
for a 14 m full downdip rupture displacement and for the rupture that releases the total shortening elastic
strain for a 500 year return period, using the variable convergence rates along the margin.
[31] For individual events, there is some trade-off between
the amount of rupture displacement and the landward limit of
rupture, but the change from seaward coseismic uplift to
landward coseismic subsidence, i.e., hinge line, is approximately at the full displacement rupture limit (see Figures 5
and 6). Because coseismic coastal subsidence is observed
for most of the Cascadia margin, this hinge line and the
full coseismic rupture are concluded to be well offshore.
The only exception is for a small area just north of Cape
Mendocino where there has been coastal coseismic uplift
and full rupture displacement is inferred to extend to beneath
the coast. This hinge line constraint limits the uncertainty in
the landward limit from this trade-off. Where there is only a
very small coastal subsidence in central Oregon, there is an
additional uncertainty [e.g., Wang, 2012; Pollitz et al., 2010].
[32] The observed coastal subsidence for the 1700 event
and the average of previous events compared to the predictions of the backslip model from Leonard et al., [2010]
for the whole margin are shown in Figure 10a along with
the high-resolution coastal subsidence data from Hawkes
et al. [2011] for Oregon shown in Figure 10b. There is good
agreement of the observed coastal subsidence rupture limits
with the estimates from the geodetic data locked and transition zones, within the considerable downdip uncertainties
of 20–40 km. Wang [2012] modeled the coseismic subsidence pattern with rupture that varied along the margin to
get a better fit to a section of the coast with high-resolution
data. A slightly better model fit was found to the data but
the misfit differences compared to those for rupture scaled
directly to convergence rate along the margin are within
the uncertainties.
4. Downdip Temperatures: The Seismic-Aseismic
Transition
[33] Above a critical temperature, rocks deform by ductiletype mechanisms and seismic rupture is not expected,
as much discussed for continental transcurrent faults like
the San Andreas [e.g., Chen and Molnar, 1983; Wong and
Chapman, 1990; Ito, 1990]. The transition from brittle to
ductile behavior is closely related but not identical to the limit
of seismic instability [e.g., Scholz, 1998]. For seismic behavior, the critical change downdip is from velocity-weakening
to velocity-strengthening slip. With velocity weakening, the
slip surface weakens with ongoing displacement and the effective coefficient of friction decreases. Runaway earthquake
displacement can occur that releases much if not all of the
built-up elastic energy. For velocity strengthening, the slip
zone strengthens with ongoing displacement so slip speeds
are limited, a viscous-like behavior. For felsic crustal rock
compositions appropriate for the continental crust overlying
the thrust, for accretionary sedimentary prisms, and subduction channel sediments, the velocity-weakening to velocitystrengthening transition from laboratory data is about 350°C
[Tse and Rice, 1986; Blanpied et al., 1991; 1995; Bürgmann
and Dresen, 2008] which agrees reasonably well with the
estimated temperatures at the maximum depth of seismicity
in 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 higher
seismogenic temperature limit, but the thrust deformation is
expected to occur in the weaker overlying sediments or forearc crust. The exception of local areas where there is shearing
5536
HYNDMAN: CASCADIA GREAT EARTHQUAKE RUPTURE LIMIT
Figure 10. (a) Along the coast comparison of the magnitude of coastal marsh subsidence and predictions from
dislocation models for 200 and 500 years of strain buildup
at the convergence rates shown [after Leonard et al., 2010].
The best fit is for approximately a 500 year great earthquake
return period indicating that most of the interseismic plate
convergence is taken up as elastic strain. The ellipse shows
the location of the data in Figure 9. (b) High-resolution
1700 coseismic subsidence estimates for Oregon compared
to elastic dislocation models (modified from Hawkes et al.
[2011]). The shaded vertical bar has lower accuracy. The
gray band is the mean estimate of coastal subsidence from
Leonard et al. [2010].
of mafic seamounts, etc., having a higher temperature limit,
is discussed below. Also, as discussed below, aseismic
serpentinite/talc may be important downdip of the fore-arc
mantle corner.
[34] At the landward downdip end of great earthquake rupture, there must be a transition zone within which rupture
cannot initiate but into which slip can continue downdip if
rupture is initiated updip. This zone represents “conditional
stability” [e.g., Scholz, 2002]. The downdip limit of the
transition zone is expected to be where the instantaneous
shear stress for slip increases rapidly downward, for example, from the effect of increasing temperature. Laboratory
studies suggest this limit is at about 450°C for crustal rocks
[e.g., Blanpied et al., 1991; 1995]. Thermal model estimates
of downdip temperatures are then required to define the location of 350°C and 450°C on the thrust.
[35] Recent subduction thermal models use finite element
modeling, a landward steepening thrust profile, accurate
incoming plate thermal regimes that include the effect of
overlying insulating sediments, and the effect of radioactive
heat 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 and
450°C isotherms on the thrust are shown in Figure 2. The
locations are within the uncertainties of the limits of the reference model locked and transition zones from geodetic data.
It has been argued that the effect of frictional heating on
the thrust in the depth range of the seismogenic zone is
small and usually may be neglected [Wang et al., 1995;
Wang et al., 1995; Wang and He, 1999] but the magnitude
of heating remains a source of uncertainty [see McKenna
and Blackwell, 2002]. Currie et al. [2004] included convection in the back arc in the models but this made no significant
difference to temperatures in the region of the seismogenic
zone. Gutscher and Peacock [2003] proposed a flat slab profile for the Cascadia subducting slab and estimated about
30 km farther landward positions of the 350 and 450°C temperatures 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 flattening used in their Cascadia models.
[36] Cooling of the crust beneath the margin by hydrothermal circulation has been proposed by Cozzens and Spinelli
[2012] such that the critical thermal limits could be as much
as 30–55 km landward relative to results from simulations
without fluid flow. Heat may be transferred from beneath
the accretionary sedimentary prism to the adjacent deep-sea
basin through fluid flow in the subducting crust, decreasing
the heat flow beneath the continental slope and increasing
the heat flow from the adjacent deep sea floor. Off SW
Japan where the incoming sediment section is much thinner,
the effect of hydrothermal circulation on the thermal data is
clear [Spinelli and Wang, 2008], with heat flow seaward of
the deformation front much higher than for a conductive
model. However, off Cascadia where the sediment section
offshore is very thick and there may be little crustal hydrothermal circulation, there is no indication of high heat flow
seaward of the deformation front; the heat flow corresponds
to that expected for a cooling plate model after correction
for the sedimentation effect. Landward of the deformation
front, the best model fit to the heat flow data across the
Vancouver Island margin is for little or no hydrothermal circulation cooling effect (Figure 11). On this profile, the heat
flow has been corrected for the thermal effect of accretionary
prism 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 hydrothermal circulation would also agree with the thermal data off
Washington, Oregon, and California if a correction was made
for the sediment thickening reduction effect. However,
Cozzens and Spinelli [2012] also argued that the effect may
be important based on the depth to the basalt-eclogite transition indicated in receiver function structure data, which
emphasizes the uncertainties in the thermal model predicted
thrust temperatures.
5. Change in Thrust Seismic Ref lection Character
From Thin Sharp Seismic to Thick Aseismic Ductile
Shear Zone
[37] From the thermal analysis summarized above, it is
expected that there is a thermally controlled downdip change
from brittle-seismic to ductile-aseismic behavior at about
350°C–450°C. This transition roughly corresponds to the
transition as seen in field geology studies from cold brittle
type behavior and thin fault zones, to broad mylonitic, and
then ductile shear zones [e.g., Sibson, 1977, 1983; Cole,
et al., 2007]. The thin shear represents strain weakening
5537
HYNDMAN: CASCADIA GREAT EARTHQUAKE RUPTURE LIMIT
Figure 11. (a) Heat flow data across the northern Cascadia margin compared to the thermal model prediction of Hyndman and Wang [1995]; the dashed red line shows the prism thickening effect that depresses
the 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 the
map of Figure 2, compared to the locked and transition zones from modeling geodetic data. If ocean crustal
hydrothermal cooling is important [Cozzens and Spinelli, 2012], these temperature limits may be several
tens of kilometers landward.
whereas the thick shear likely represents strain strengthening.
In the latter, the fault zone continually moves outward from
the initial shear zone which becomes stronger with deformation [e.g., Kirby, 1985] and thickens into weaker as yet
undeformed rocks. The transition from thin brittle to thick
ductile behavior is expected to be at a slightly higher temperature 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 change
from a thin strong detachment reflector to a reflective
5–10 km thick band in marine and land multichannel seismic
reflection data approximately corresponds to the downdip
limit of seismic rupture for M > 8 earthquakes at a number
of subduction zones, e.g., Alaska, Chile, and SW Japan.
They concluded that the thin to thick reflection transition
marks the brittle to ductile transition. They mapped this transition along the margin of southern Vancouver Island (reflection
zone 2 on map of Figure 12) and found it generally lies just
seaward of the coast and corresponds reasonably well to the
modeled transition zone downdip limit from other constraints,
i.e., approximately 450°C. Despite considerable variation in
the reflection character of the subduction thrust on a number
of seismic lines, there is a consistent downdip change from a
thin reflection zone offshore (reflection zone 1) to a broad
reflection band (E layer; reflection zone 3) some 5 to 7 km
thick, near the west coast of Vancouver Island (Figure 12).
The thin reflection is less than 2 km thick, but the true boundary is probably thinner, the image being thickened by the
reflection signature. The transition is a few tens of kilometers
wide that corresponds in position very approximately to
the downdip limit of the geodetic and thermal transition
(450°C) zones (Figure 12).
[39] This interpretation across the Vancouver Island margin 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 interpretation that the reflective band is the subducting oceanic
crust or the upper crust. In contrast, from detailed seismic
structure analysis, Calvert et al. [2011] concluded that the
E-layer dipping reflective band is a shear zone above the
subducting plate as in the original interpretation of Clowes
et al. [1987]. This disagreement is not yet resolved. Further
calibration of this method using the change in reflection character downdip is needed for subduction zones that have had
recent great earthquakes with a well-constrained downdip
rupture limits. Initial results from a seismic structure survey
to examine the downdip change in reflection character for
the Alaska subduction zone were reported by Shillington
et al. [2011].
6. Fore-Arc Mantle Corner: Aseismic Serpentinite
and Talc on Thrust
[40] The fore-arc mantle corner may limit downdip rupture
because of the presence of aseismic serpentinite and talc. The
fore-arc mantle is commonly significantly serpentinized by
rising fluids driven upward from the underlying dehydrating
subducting crust [e.g., Hyndman and Peacock, 2003, and references therein] (Figure 13). The structure of serpentinite is
expected to make it weak where it is in contact with the
underlying fault zone, a conclusion supported by the arguments for no significant frictional heating on the thrust, at
least landward of the fore-arc mantle corner [e.g. Wada
et al., 2007]. There is some uncertainty as to the weakness
5538
HYNDMAN: CASCADIA GREAT EARTHQUAKE RUPTURE LIMIT
Figure 12. Change in megathrust multichannel seismic reflection character from a thin inferred seismic
signature (Reflection band 1), to a thick inferred aseismic shear zone (Reflection band 3; modified after
Nedimovic et al. [2003]), compared to Flück et al. [1997] reference locked and transition zones from
modeling geodetic data, i.e., approximately 350°C and 450°C. The reflection transition (Reflection band
2) occurs at slightly cooler than 450°C and seaward of the ETS zone.
of serpentinite [e.g., Escartın et al., 2001; Moore et al., 1997;
Moore and Rymer, 2007; Reynard, 2013], but another factor
may ensure that the thrust beneath the fore-arc mantle is usually aseismic. Talc is generated by the silica-saturated fluids
from the underlying dehydrating oceanic crust and underthrust sediments that react with the overlying fore-arc mantle
serpentinite [e.g., Peacock and Hyndman, 1999]. Talc is a
very weak mineral and fault zones containing talc are
very likely aseismic, as suggested for the San Andreas Fault
[Moore and Rymer, 2007]. Brocher et al. [2003], Preston
et al. [2003], and Ramachandran and Hyndman [2011] using
wide-angle seismic structure and seismic tomography studies
showed that there is a substantial reduction in seismic velocity
and 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 produces a significant magnetic anomaly high.
[41] Tichelaar and Ruff [1993], Hyndman et al. [1997],
and Oleskevich et al. [1999] concluded that the maximum
depth of great earthquake rupture, commonly defined by
aftershocks, often corresponds to the intersection with the
fore-arc mantle corner. Unfortunately, the location of the
fore-arc mantle corner has not yet been well defined on most
margins. Ruff and Tichelaar [1996] concluded that for the
South American subduction zone, the downdip limit of great
earthquake rupture approximately coincides with the fore-arc
mantle corner which in turn is approximately at the coast for
much of that margin. Heuret et al. [2011] compiled the maximum depth of intermediate magnitude thrust events for a
number of M5.5–7.0 thrust earthquakes in a global study
and concluded that the average maximum event depth for
subduction zones where ocean plates descend beneath continental lithosphere was about 40 km. This is somewhat deeper
than the average of 34 ± 11 km for the compiled depth estimates of fore-arc Moho depth by Wada and Wang [2009].
However, this is not a large difference since the Moho intersection depth for most subduction zones is very poorly
defined. For example, the Moho may not exhibit a crust-mantle
downward 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 allow
examination of this proposed downdip rupture limit for
M ~ 9 events, Sumatra in 2004, Chile in 2008, and NE
Japan in 2011. For the great earthquake of 2004 off
Sumatra, the rupture extent was to a depth of about 30 km
[e.g., Chlieh et al., 2007, and references therein] in agreement with the locked zone from geodetic and paleogeodetic
data [e.g., Chlieh et al., 2008]. The depth of ~30 km is within
the common range of 30–50 km for continental fore-arc mantle corners, but Dessa et al. [2009] and Simoes et al. [2004]
have concluded from seismic structure data that the fore-arc
Moho depth is very shallow, 21–25 km, on this margin such
that the locked/rupture zone extends to beneath the fore-arc
mantle. If this unusually shallow Moho depth does apply to
this fore-arc mantle corner, the corner does not constrain
downdip rupture on this margin. However, the inference of
an exceptionally shallow fore-arc mantle leaves considerable
uncertainty. Also, the depth and the critical temperature for
seismic behavior are similar [Hippchen and Hyndman,
2008], so the downdip limit may be thermally controlled
rather 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] from
geodetic data concluded a downdip coseismic rupture limit
near the depth where the subducting Nazca plate intersects
with the continental Moho of the South America plate at
40–50 km [Haberland et al., 2006]. This depth is in the range
of maximum rupture depths estimates for a compilation of
earlier large thrust earthquakes on this margin by Tichelaar
and Ruff [1991]. This depth also agrees with the limit of the
locked zone inferred from preseismic GPS data [Ruegg
et al., 2009; Moreno et al., 2010]. Haberland et al. [2006]
found the fore-arc Moho at a depth of about 50 km in good
agreement with the fore-arc mantle corner being the downdip
limit for rupture in this great earthquake.
[44] For the third recent great earthquake, Tohoku off NE
Japan in 2011, the main large displacement rupture appears
5539
HYNDMAN: CASCADIA GREAT EARTHQUAKE RUPTURE LIMIT
Figure 13. Cascadia section showing the inferred aseismic
serpentinite and talc on the subduction thrust that may limit
seismic rupture downdip (modified from Peacock and
Hyndman [1999]; bru = brucite; serp = serpentinite). The
ETS zone approximately corresponds to the fore-arc mantle
corner, whereas the estimated Cascadia seismogenic zone is
significantly shallower.
to terminate just seaward of the coast where the thrust depth
is 30–40 km deep [e.g., Simons et al., 2011; Loveless and
Meade, 2011], a common depth for the fore-arc mantle corner. Aftershocks extend to a depth of 40–45 km just seaward
of the coast. However, seismic structure data reported by
Hino et al. [2000] and Takahashi et al. [2004] just to the
north suggest that the fore-arc Moho is at an unusually
shallow depth of about 20 km. If the latter is correct, the
fore-arc mantle corner does not limit downdip rupture. Seno
[2005] also found the maximum depth of other thrust seismicity inferred to be on the subduction thrust to extend to over
50 km depth on this margin, much deeper than the estimated
fore-arc Moho. However, he found the maximum thrust earthquake depth to be only ~30 km to the north of the 2011 rupture
area and to the south off SW Japan, in general agreement
with the fore-arc Moho depth. As for Sumatra, rupture off
NE Japan appears to extend deeper than the fore-arc mantle
corner, but in both cases, this conclusion depends on the
fore-arc Moho being exceptionally shallow at 20–25 km. If
the fore-arc Moho corner is actually at the more common
35–50 km, there would be agreement with this rupture limit.
[45] Lay et al. [2012] in a summary of global great earthquake behavior found that large earthquake displacements
most commonly occur to a depth of 35 km. Ruptures of smaller
isolated patches on the thrust extend deeper to ~55 km, generating only small magnitude earthquakes. Their results suggest
that some smaller thrust events may extend to greater depths
than the main displacements in megathrust events. Caution
therefore should be used in employing the maximum depth
of smaller thrust events for a limit to large displacement
rupture and to megathrust earthquake hazard.
[46] Another complicating factor that may result in thrust
earthquakes deeper than the fore-arc mantle corner is the possibility that some thrust earthquake ruptures occur within
subducted sediments, subducted seamounts, or other subduction channel material, isolated from overlying fore-arc
mantle serpentinite and talc. More detailed structural studies
are needed in the region of the Moho intersection where the
downdip rupture limit for both low- and high-frequency
energy has been well determined for recent great earthquakes.
[47] The thickness of the Cascadia fore-arc crust is quite
well determined in several areas by a number of seismic studies, including receiver function analyses, seismic tomography,
and wide-angle seismic structure experiments. However, the
dip angle of the thrust profile is not so well determined and
there remains some debate, especially for the Vancouver
Island margin, as to what represents the thrust in seismic structure data [e.g., Hansen et al., 2012]. McCrory et al. [2006,
2012a] provided a review of the thrust constraints. An initial
summary of constraints to the location of the fore-arc mantle
corner has been provided by McCrory et al. [2012b]. The
accuracy is sufficient to show that the fore-arc mantle corner
is generally well landward of the seismogenic limit from most
of the other constraints (e.g., Figure 14a). The differences in
various thrust profile estimates are not significant to a depth
of about 30 km [e.g., Wech et al., 2009, their Figure 4]
and so are not very important for the seismogenic zone as defined by the geodetic, thermal, and coastal marsh constraints.
However, should the seismogenic zone extend deeper, at
greater depths, there are significant differences in various estimates of the thrust dip profile. A constraint to the location of
the fore-arc mantle corner may come from the updip limit of
ETS tremor (discussed later). Wech et al. [2009] suggested
that the well-resolved, sharp updip edge of ETS tremor epicenters reflects a change in plate interface coupling properties.
This boundary in coupling may be associated with the fore-arc
mantle corner (e.g., Figures 13 and 14a), with ETS occurrence
being 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 recent
discovery, with the most detailed studies in Cascadia and SW
Japan [e.g., Rogers and Dragert, 2003; Obara et al., 2004;
Schwartz and Rokosky, 2007; Peacock, 2009; Gomberg
et al., 2010; Beroza and Ide, 2011]. A number of authors
have suggested that ETS may provide a great earthquake
downdip rupture limit. The slow slip associated with ETS
tremor likely occurs mainly on or just above the subduction
thrust, and for Cascadia accommodates much if not all of
the plate convergence in the ETS zone [e.g., Dragert and
Wang, 2011]. Therefore, there is little or no elastic strain
buildup in the ETS zone to be released in great events. If this
last conclusion is correct, the slow slip zone represents a
downdip limit for significant rupture displacement. The largest uncertainty as to whether most if not all of the plate convergence is so accommodated lies in the amount of slip in
small events between the main ETS events [e.g., Wang
et al., 2008]. The physical processes involved in ETS remain
uncertain, but the slow slip behavior itself indicates that the
fault zone has ductile characteristics at that depth.
[49] It cannot be excluded that there is some small elastic
strain buildup in the ETS zone and that great earthquake rupture extends into the ETS zone with very small displacement.
It is recognized that even a small displacement slip in this
region could substantially influence the resulting seismic
shaking at major cities in Cascadia. If slow slip represents
5540
HYNDMAN: CASCADIA GREAT EARTHQUAKE RUPTURE LIMIT
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 comparison. 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 estimates
of the downdip limit of significant great earthquake rupture. (b) Distribution of Cascadia tremor
2009–2013 (data from Wech et al. [2009] and Wech [2010], http://www.pnsn.org/tremor), compared
to Flück et al. [1997] reference model locked and transition zones. (c) Illustration that the sum of
the coseismic downdip transition and the updip transition for the slow slip does not equal the total
plate convergence rate over the great earthquake cycle. There must be an intervening region that does
not move coseismically or in slow slip events.
conditionally stable frictional behavior, is it possible that the
slow slip event zone could be pushed to seismic rupture by a
large earthquake on the seismogenic portion of the interface.
However, the intervening zone between ETS and the main
locked/rupture zone estimated from most constraints also
would need to rupture coseismically.
[50] The horizontal positions of ETS tremor are quite
well determined [e.g., Kao et al., 2009; Wech et al., 2009;
Gomberg et al., 2010; Dragert and Wang, 2011; Ghosh
et al., 2012] (e.g., Figure 14b) although the depths are somewhat uncertain. The location of the slow slip has been estimated through dislocation modeling of geodetic data and its
location is also quite well determined but not quite as well
as the tremor [e.g., Dragert and Wang, 2011; Wech et al.,
2009; McCaffrey, 2009; Schmidt and Gao, 2010; Bartlow
et al., 2011]. There is some indication that the slip is slightly
seaward of the tremor but this is within the uncertainties
(example in Figure 14a). Hawthorne and Rubin [2013] found
a short timescale correlation between slow slip and tremor in
Cascadia that indicates that they are closely related.
[51] Audet et al. [2010] and Ghosh et al. [2012] conclude
that the top of the oceanic plate is at a depth of 30–40 km
beneath the tremor based on various estimates of plate depth.
In their review, Gomberg and the Cascadia 2007 and
Beyond Working Group [2010] also concluded that the depth
range for Cascadia slow slip is generally 30–40 km. The forearc Moho is approximately at 30 km so the tremor appears to
be across and landward of the fore-arc mantle corner but the
depth uncertainties for the fore-arc Moho are substantial. In a
more conclusive analysis of slow slip from continuous GPS
data, Holtkamp and Brudzinski [2010] found that the ETS
zone is likely located where the subduction thrust underlies
the fore-arc mantle, such that the fore-arc mantle corner is
the updip limit of most of the ETS. Peacock [2009] also
concluded that the tremor for SW Japan and Cascadia is just
below or at the fore-arc mantle corner. A physical composition or structure constraint is supported by the conclusion
of Peacock [2009] that tremor occurs at different depths for
different subduction zones. He estimated the temperature
on the thrust at the depth of the Cascadia tremor to be about
5541
HYNDMAN: CASCADIA GREAT EARTHQUAKE RUPTURE LIMIT
Figure 15. Gravity anomalies along the Cascadia margin that may delineate the great earthquake rupture
extent downdip and possible along-strike segmentation defined by sedimentary basins A–E [after Wells
et al., 2003], compared to Flück et al. [1997] reference locked.
510°C, well above the proposed maximum of about 350°C
for seismic behavior. The slow slip is therefore 30–50 km
landward of the great earthquake rupture zone from the
thermal constraint and most of the other constraints for the
downdip limit of rupture.
[52] A possible explanation for the ETS tremor being
associated with the fore-arc mantle corner is that the fluids
responsible are channeled upward and seaward beneath the
fore-arc mantle to rise above the corner. Very low Poisson’s
ratio in the crust above this corner has been observed,
suggested by Ramachandran and Hyndman [2011] to result
from silica precipitation from the rising fluids. Katayama
et al. [2012] also proposed an impermeable barrier at the
fore-arc Moho that channels fluids seaward to rise into the
crust above the corner producing tremor.
[53] If the conclusion that the slow slip of ETS occurs well
landward of the great earthquake rupture zone is correct
(Figure 14), it means that there is an intervening section of
the thrust where long-term plate convergence is accommodated neither in great earthquake rupture nor in slow slip
events, but rather in some form of long-term aseismic creep
[e.g., Holtkamp and Brudzinski, 2010]. It should be noted
that Cascadia has an unusually shallow downdip great earthquake rupture limit, probably because there is a thermal
rupture limit for this very hot subduction zone. The more common cool subduction zones often have a much deeper limit
that may be at the fore-arc mantle corner and may not have a
gap between the great earthquake rupture and ETS slow slip.
8. Geological Associations of Rupture
With Offshore Basins
[54] There is a common correlation between offshore forearc basins, defined mainly by gravity, and the regions of seismic rupture in past great events globally [Wells et al., 2003;
Song and Simons, 2003]. In a related observation, Ruff and
Tichelaar [1996] found that the landward limit of great earthquake rupture is often beneath the coast. The coast often
defines the landward limit of the basins. This correlation
needs to be examined for recent great events such as the
2004 great earthquake of Sumatra, the 2008 great earthquake
of 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 largest
seismic moment release for the majority of great earthquakes
studied 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 with
increases or maxima in the trench parallel gravity field. A
physical mechanism has not yet been identified; however,
the deep-sea terraces and basins may evolve not only by
growth of the outer arc high but also by interseismic
subsidence that is not recovered by uplift during great
earthquakes. There could be a link between subsidence,
subduction erosion, and seismogenesis. Kilometers of late
Cenozoic subsidence and crustal thinning above some of
the source zones elsewhere are indicated by seismic profiling
and drilling. They are thought to be caused by basal subduction erosion [e.g., Scholl and von Huene, 2007] although
large amounts of erosion are most commonly argued for sediment deficient subduction zones with high convergence
rates [e.g., Stern, 2011] and may not apply to Cascadia.
[56] The source zone for the Cascadia 1700 A.D. earthquake contains five semicontinuous basin-centered gravity
lows (Figure 15) that may indicate asperities at depth [e.g.,
Wells et al., 2003]. The gravity gradient marking the inferred
downdip limit to large coseismic slip lies offshore, except in
northwestern Washington, where the low extends landward
beneath the coast. Figure 15 shows the good correlation between the locked and transitions zones defined geodetically
and thermally and the gravity-defined basins for Cascadia.
The basin landward extents are also in general agreement
with the estimated landward limit of rupture in 1700 and
earlier great events based on coastal coseismic subsidence
[Leonard et al., 2004; 2010]. On other margins, seismic slip
has been concluded to be less beneath along-strike structural
highs [e.g., Wells et al., 2003] so such highs may delineate
rupture variability along the margin with some aseismic slip.
9. Downdip Limit of Small Subduction
Thrust Earthquakes
[57] The downdip limit of small to intermediate size thrust
earthquakes on the subduction thrust provides an important
seismogenic zone constraint for many other subduction
zones [e.g., Tichelaar and Ruff, 1993]. These thrust events
tend to concentrate near the downdip limit of rupture for great
earthquakes elsewhere [e.g., Tichelaar and Ruff, 1993; Ruff
and Tichelaar, 1996].
[58] However, on the Cascadia margin, only a very few
small such events have been detected and in a limited area
[Tréhu et al., 2008; Williams et al., 2011] (Figure 16), but
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HYNDMAN: CASCADIA GREAT EARTHQUAKE RUPTURE LIMIT
displacements in megathrust events and therefore that the
maximum depth of smaller thrust events may not define the
limit to large megathrust rupture displacement.
10. Landward Limit of Earthquakes
on the Nootka Transform Fault
Figure 16. Small thrust earthquakes that may define the
downdip limit of seismic behavior on the subduction thrust
off Oregon [after Williams et al., 2011]. They occurred where
thermal model temperatures on the subduction thrust interface are 400–450°C, so they may represent rupture of mafic
seamounts that have a higher seismogenic temperature limit
than most of the subduction thrust. The downdip limits to
the locked (black solid line) and transition (dashed line)
zones are from the Flück et al. [1997] reference model.
they are at the estimated depth of the subduction thrust and
some have been shown to have rupture mechanisms at the
dip angle of the thrust. They may be related to local stress
concentrations associated with features such as subducted
seamounts [Tréhu et al., 2012]. Some are accurately located
with ocean bottom seismographs and they may provide an
estimate of the downdip seismogenic limit. The estimated
source depths are 9–11 km for a M4.9 event and 15–17 km
for a M4.8 event. Other small earthquakes are in the range
from 10 km to 18 km. These events are within the location
uncertainties of the subduction thrust and probably represent
seismogenic motion on the Cascadia megathrust. The deeper
events are within the transition zone for our reference model
[Flück et al., 1997] but well updip of where episodic tremor
and slip (ETS) has been documented and well updip of the forearc mantle corner. Williams et al. [2011] concluded that the
events occurred at the downdip edge of the locked/transition
zone which may be some 30 km farther landward than the
Flück et al. model (Figure 16) and suggested that the transition
zone extends approximately to the coast. However, an alternative explanation is that these events may have occurred on
sheared-off subducted seamounts [Tréhu et al., 2012] or local
small blocks of crustal material in the subduction channel with
a mafic composition. Such material could have a higher limiting temperature than the normal more felsic sediments and
fore-arc crustal rocks on the main thrust contact [e.g., Chen
and Molnar, 1983; He et al. 2007] and explain why these
thrust earthquakes occur where the temperature is higher than
the concluded common seismogenic limit of about 350°C.
[59] As noted above, Lay et al. [2012] in a summary of
global great earthquake behavior found that large earthquake
displacements most commonly occur to a depth of 35 km but
that the rupture of smaller isolated patches often extends
deeper to ~55 km. Their results suggest that some smaller
thrust events may extend to greater depths than the main
[60] The left-lateral Nootka transform fault zone lies off
central Vancouver Island approximately orthogonal to the
margin. The landward downdip extent of the strike-slip seismicity on this fault zone (Figure 17) may provide a limit to
the downdip extent of seismic behavior on the overlying
subduction thrust. The strike-slip motion is between the main
Juan de Fuca plate to the south and Explorer plate complex to
the north [Hyndman et al., 1979]. The fault zone is concluded
to be the result of quite recent ~3 Ma breakup of the Juan
de Fuca plate [e.g., Riddihough, 1984; Braunmiller and
Nábělek, 2002]. There is a continuous change in the orientation of this fault zone, resulting in a broad irregular shear
zone. The epicenter uncertainties for the Nootka fault zone
earthquakes are significant because most seismograph stations are landward. However, a few events have been located
with ocean bottom seismographs [Hyndman and Rogers,
1981] and the location bias near the coast is small. Strong
seismicity on the Nootka fault zone extends approximately
to the coast well landward of the deformation front. Like
the small thrust earthquakes off Oregon, the Nootka fault
earthquakes extend approximately to the landward limit of
the 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 limited landward, in a similar manner to the postulated thermal
limit for continental strike-slip faults and for the subduction
thrust. The seismicity on the subducted Nootka transform
fault zone is slightly deeper than the megathrust and therefore
may be slightly hotter. The result is a landward seismogenic
Figure 17. Seismicity on the Nootka transform fault
extending off Vancouver Island that may provide a landward
limit of seismic behavior (from Geological Survey of Canada
data file), compared to the Flück et al. [1997] reference
locked and transition zones. These earthquakes extend to
where thermal model temperatures are about 450°C. This
higher than normal seismogenic temperature limit may be
explained by rupture of mafic ocean crust material that has
a higher temperature seismogenic limit than most of the
subduction thrust.
5543
HYNDMAN: CASCADIA GREAT EARTHQUAKE RUPTURE LIMIT
limit that is slightly seaward. However, the transform seismicity is argued to be only in the upper few kilometers of
the oceanic crust [Hyndman and Rogers, 1981; Willoughby
and Hyndman, 2005] so not much deeper and not much hotter than the subduction thrust. However, this fault cuts mafic
oceanic crustal rocks that likely have a higher maximum
seismogenic temperature than for the more felsic composition of the rocks and sediments in contact with the subduction
thrust, so there may be a higher maximum seismogenic temperature and therefore the landward limit may be farther landward. For example, Chen and Molnar [1986] estimated the
temperature at the base of the seismogenic zone for the fracture zones cutting oceanic lithosphere to be 600–800°C and
He et al. [2007] from laboratory data found that seismic
behavior could extend to over 510°C for gabbroic oceanic
lower crust rocks. This limiting temperature is in general
agreement with the landward limit of Nootka seismicity zone
(Figure 17) at about 450°C and similar to that inferred for the
small thrust earthquakes off Oregon discussed earlier. This
constraint provides evidence that significant thrust rupture
at least does not extend far beneath the Vancouver Island
coast. From this constraint, the maximum megathrust rupture
extent may be seaward but is unlikely to be landward.
[62] To the south of the Nootka fault zone, there is incoming plate seismicity that extends several tens of kilometers
landward of the coast (Figure 17). This seismicity does not
continue to the south off Washington and Oregon. It could
result from extensional faulting generated by plate bending
and steepening landward or due to dehydration fluid pressures [e.g., Abers et al., 2009]. Like the Nooka seismicity
and the small events off Oregon, these earthquakes extend
to where the estimate temperature is about 450°C which is
approximately that expected for mafic ocean crustal rocks.
Like the Nootka fault constraint, the maximum megathrust
rupture extent may be seaward but is unlikely to be landward
of this subducting plate seismicity.
11.
Discussion
[63] The following is a summary of constraint results:
[64] 1. The geodetic constraints for the current locked and
transition zones have excellent data and the highest resolution,
about ±10 km differences among the various models. The
greatest uncertainty is from the degree to which the locked/
transition zones represent the full rupture and the tapering
to zero rupture displacement. Recent great earthquakes elsewhere indicate that this is a good approximation. The seismic
rupture limit may be slightly seaward of the locked/transition
zones because of the postseismic transients that are not
included in most locked zone models.
[65] 2. Coastal coseismic subsidence provides the least
ambiguous constraint to the rupture zone but has quite low
resolution of about ±30 km downdip [see Leonard et al.,
2010], and there is some ambiguity trade-off of location of
rupture and amount of rupture displacement, especially off
Oregon where the subsidence is small. There also is the question of how much of the displacement occurs in postseismic
deformation. The coastal subsidence constraint agrees well
with the geodetic constraint.
[66] 3. The thermal constraint has significant uncertainties
in 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 hydrothermal
cooling in the incoming crust is significant, the critical temperatures may be a few tens of kilometers farther landward.
[67] 4. The transition in thrust reflection character from
thin seismic to thick aseismic has quite high resolution [see
Nedimovic et al., 2003] but has a large uncertainty because
of the lack of a good calibration where there has been a recent
megathrust earthquake. From laboratory and field data, ductile
behavior in crustal rocks is expected at roughly 450°C. The
reflection transition agrees well with the transition zone
thermal limit and the geodetic constraint transition zone limit.
[68] 5. The fore-arc mantle corner or Moho limit has
considerable uncertainty, about 20 km [e.g., McCrory et al.,
2012], but is about 50 km landward of the main rupture indicated by the geodetic models. It roughly agrees with the ETS
limit. This seismic rupture limit has been concluded for a
number of cold subduction zones elsewhere but there have
been several interpretations of thrust earthquakes extending
deeper. This constraint is only a maximum depth, and for
Cascadia the rupture limit appears to be thermally controlled
at 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 is
where the thrust is at a depth of about 30 km, which corresponds quite well with the fore-arc mantle corner, although
the latter is not yet well constrained in some areas of the
margin. The ETS updip limit is 30–50 km landward of the
geodetic locked/transition zone limit. This constraint is only
a maximum depth and allows a shallower rupture limit. For
Cascadia, the rupture limit appears to be thermally controlled
at a shallower depth.
[70] 7. The center of gravity lows beneath the slope and
shelf gives a moderate resolution constraint of about ±30 km.
There is agreement with the maximum rupture for some but
not all great earthquakes elsewhere, but the reason is not yet
clear. There is general agreement with the geodetic, coastal
paleoseismic, and thermal constraints.
[71] 8. The small thrust earthquakes off Oregon that are
interpreted to be on the subduction thrust are located with
an accuracy of a few kilometers, near the downdip limit of
the 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 as
suggested by Tréhu et al. [2012] or other local mafic material
in this area, with a higher temperature limit of about 450°C.
[72] 9. The earthquakes on the Nootka transform fault zone
are located with an accuracy of a few kilometers. They
extend to about the coast of Vancouver Island, i.e., near the
landward limit of the transition zone of the geodetic models.
The temperature at this location is too hot for seismic initiation in felsic composition rocks. However, the fault cuts
mafic oceanic crust that likely has a higher temperature limit
than the normal thrust seismic temperature limit.
[73] Three of the downdip limits to great earthquake rupture
agree within the uncertainties with reasonable resolution
of ±10 to ±20 km, i.e., geodetic, coastal paleoseismic, and
thermal, and are close to our reference model (e.g., Figures 2
and 8–10). The other limits allow the reference model limit.
For major rupture, i.e., ~10 m or 50% of the maximum
displacement, the uncertainty is about ±20 km, with a small
possibility that the constraints are systematically incorrectly
interpreted. The limit for the landward tapering 1–2 m rupture
5544
HYNDMAN: CASCADIA GREAT EARTHQUAKE RUPTURE LIMIT
that can still provide strong shaking is less well estimated but appears to be 25–50 km farther landward. This latter limit extends
some distance inland of the coast for northern Washington.
[74] The small thrust earthquakes off Oregon and the earthquakes on the Nootka fault are interpreted to be subject to a
higher temperature limit, ~450°C, than the subduction thrust,
~350°C, because they are interpreted to rupture more mafic
rocks, i.e., seamounts and ocean crust, respectively. The
change in thrust seismic reflection character from a thin fault
to a broad shear zone also occurs at about 450°C as expected
from laboratory and field data. The remaining constraints are
consistent with thermally limited rupture mainly offshore for
this especially hot subduction zone.
[75] There are two constraints that provide a fairly strict
maximum rupture depth if they are correctly interpreted,
albeit with significant uncertainty in position. These are the
aseismic fore-arc mantle corner and the updip limit of ETS
slow slip that accommodates most if not all plate convergence. The two constraints agree within the large uncertainties. They are a little deeper than the transition zone
downdip thermal limit of ~450°C. Both allow a shallower
rupture limit and are likely deeper than the actual maximum
depth for significant rupture displacement for this hot
subduction zone.
[76] A surprising result from the analyses of the different
downdip rupture limits is that there appears to be a gap of
50–100 km between the downdip limit of significant rupture
displacement in great earthquakes and the region of slow slip
of ETS. The analysis of slow slip from continuous GPS data
by Holtkamp and Brudzinski [2010] supports this conclusion., i.e., a gap between the great earthquake interseismic
locked zone and the deeper inter-ETS event locked zone that
slips in ETS slow slip events. It is not yet clear how and when
this intermediate depth zone moves. Does it move continuously with slow creep with plate convergence accommodated
by episodic rupture displacement above and slow slip events
below? As an alternative, are there smaller very frequent
slow slip events that are too small to be observed geodetically
and do not result in tremor? This intermediate depth zone
result suggests that the ETS zone does not provide a downdip
limit of great earthquake coseismic rupture. Rupture is limited to a shallower depth, although more work is needed to
confirm this conclusion. It is noted that for at least some cool
subduction zones, the downdip great earthquake rupture limit
may coincide with updip limit of ETS slow slip, since both
limits may be at the fore-arc mantle corner.
[77] The mechanisms responsible for the downdip limit
of seismogenesis on subduction thrust faults globally are
still not well understood. The thermal limit of subduction
thrust rupture appears to apply to hot subduction zones like
Cascadia, SW Japan, and Mexico. However, there are only
a few subduction zones that are sufficiently hot for the critical
temperatures to be reached at a shallow depth to test this
hypothesis. The downdip limit of serpentinite/talc at the
fore-arc mantle corner agrees with data from a number of
subduction zones but is questioned by a number of great
earthquakes, including the Sumatra and NE Japan Tohoku
M9 earthquakes. Both these earthquakes have been concluded to have some rupture deeper than the fore-arc mantle
corner. However, in both cases, this corner is inferred to be
exceptionally shallow, 20–25 km, and it is recognized that
the depth of the fore-arc Moho tends to be especially difficult
to determine. The location of the fore-arc mantle corner is an
important need for future study. There also is the question of
smaller thrust earthquakes that occur at depths greater than
the main rupture displacements. They may represent rupture
of seamounts or other mafic crustal material with a higher
seismogenic temperature than the main thrust.
12. Seismic Ground Motion From Landward
Limits 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 and
Macias, 2009]. The landward limit of rupture is one of the
most important control of shaking inland for these models.
The lack of significant historical earthquakes in Cascadia
means it is very difficult to test existing ground motion
attenuation models. There are several significant problems
in converting the available landward limit constraints for
Cascadia to ground motion:
[79] 1. The ground motion attenuation with distance relations so far presented used poorly defined distances from
the earthquake rupture. Usually, they have used the landward
rupture extent for great earthquakes elsewhere as defined by
aftershocks or by seismic rupture models which have a
large uncertainty.
[80] 2. The nature of rupture displacement landward taper,
i.e., the transition zone, may be very important, since, as
discussed earlier, 1–2 m of displacement can produce very
damaging ground motions. This landward taper is not incorporated in most ground motion models. Most great earthquake rupture inversions and models have a transition zone
with decreasing rupture displacement landward, corresponding roughly with the estimated tapering of displacement in
observed recent great events. For Cascadia with approximately 20 m average full rupture, the 50% or 10 m rupture
limit can be estimated with reasonable assurance to within
about ±20 km, but the position of the taper to ~2 m (~10%)
is less accurately defined 25–50 km farther landward.
[81] 3. Within the great earthquake rupture zone, there may
be a frequency dependence of seismic energy that varies with
downdip distance. More damaging high frequencies may be
generated near the landward limit of rupture, even though
the rupture displacement may be small there.
[82] Acknowledgments. The author would like to thank scientists at the
Pacific Geoscience Centre, Geological Survey of Canada for many useful discussions, and a number of authors whose work is discussed in this article for
clarifications and additional information.
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