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Earth-Science Reviews 141 (2015) 4555
Contents lists available at ScienceDirect
Earth-Science Reviews
j ourna l homepage: www.e lsev ie r .com/ locate /earsc i
revCoseismic slip on shallow dcollement megathrusts: implications
forseismic and tsunami hazardJudith Hubbard , Sylvain Barbot, Emma
M. Hill, Paul TapponnierEarth Observatory of Singapore, 50 Nanyang
Avenue, Nanyang Technological University, 639798, Singapore
Corresponding author. Tel.: +65 9855 2730, +65E-mail addresses:
[email protected] (J. H
http://dx.doi.org/10.1016/j.earscirev.2014.11.0030012-8252/ 2014
Elsevier B.V. All rights reserved.a b s t r a c ta r t i c l e i n
f oArticle history:Received 24 July 2014Accepted 8 November
2014Available online 15 November 2014
Keywords:EarthquakesDcollementSubduction zoneFold-and-thrust
beltSeismogenic zoneFaultsFor years,many studies of subduction
zones and on-land fold-and-thrust belts have assumed that the
frontal por-tions of accretionary prisms are tooweak to rupture
coseismically andmust therefore be fully creeping.We pres-ent a
series of examples, both on-land and offshore, demonstrating that
inmany cases, shallow dcollements arecapable of large, coseismic
slip events that rupture to the toes of the fault systems. Some of
these events are as-sociatedwith ruptures that initiate down-dip,
while others appear to be limited to the frontal, shallow portion
ofthe wedge.We suggest that this behavior is not limited to the
examples described here, but rather is common tomany (per-haps
most) accretionary wedges and fold-and-thrust belts around the
world. Indeed, there may be many otherexamples of similar
earthquakes, where existing data cannot constrain slip at the toe.
We do not characterizethe regions and events described here as
unusual, as they encompass a wide range of settings. This study
indi-cates that there is an urgent need to reevaluate seismic and
tsunami hazard in fold-and-thrust belts and subduc-tion zones
around the world, allowing for the possibility of shallow
dcollement rupture.
2014 Elsevier B.V. All rights reserved.Contents1. Introduction .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . 462. Evidence
for shallow slip in megathrust earthquakes . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
47
2.1. The Himalaya . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. 472.2. Bolivia . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . 482.3. Western Taiwan . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . 492.4. Santa Barbara Channel, Southern California . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . 492.5. Sumatra subduction zone, Mentawai segment . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . 502.6. Sumatra subduction zone, northern segment . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . 502.7. Java subduction zone . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . 502.8. Japan Trench . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . 502.9. Kuril Trench . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . 502.10. Japan, Nankai Trough . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . 502.11. Solomon Islands . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . 512.12. Middle America megathrust . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . 512.13. Alaskan Trench . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . 512.14. Peru . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . 512.15. Chile . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . 512.16. Summary . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . 51
3. How can weak decollements behave seismically? . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . 513.1. The state of stress within the wedge, and preferred
slip planes at geological time scales . . . . . . . . . . . . . . .
. . . . . . . . . . 523.2. Variations in strength through the
earthquake cycle . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . 53
4. Consequences of seismogenic decollements . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . 536592 7537.ubbard), [email protected] (S. Barbot),
[email protected] (E.M. Hill), [email protected] (P.
Tapponnier).
http://crossmark.crossref.org/dialog/?doi=10.1016/j.earscirev.2014.11.003&domain=pdfhttp://dx.doi.org/10.1016/j.earscirev.2014.11.003mailto:[email protected]:[email protected]:[email protected]:[email protected]://dx.doi.org/10.1016/j.earscirev.2014.11.003http://www.sciencedirect.com/science/journal/00128252www.elsevier.com/locate/earscirev
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46 J. Hubbard et al. / Earth-Science Reviews 141 (2015) 45555.
Discussion and conclusions . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . 53Acknowledgements . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . 53References . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . 531. Introduction
Many regions of plate convergence are underlain by
dcollementmegathrusts. They form the base of both accretionary
wedges and fold-and-thrust belts. These faults may extend laterally
for hundreds or thou-sands of kilometers, and downdip for tens to
hundreds of kilometers. Tra-ditionally, estimates of seismic hazard
have assumed that these faults slipaseismically, without radiating
significant seismic energy (e.g., Pachecoet al., 1993; Hyndman et
al., 1997; Oleskevich et al., 1999). However, inseveral recent
cases, shallow dcollements have been shown to slip inlarge,
discrete events (e.g., 2011 Mw 9.0 Tohoku-Oki earthquake,
Japan;2010 Mw 7.8 Mentawai earthquake, Indonesia; 1999 Mw 7.6
Chi-Chiearthquake, Taiwan). On land, dcollements frequently lie
along the bor-ders of large, populated basins, and therefore pose
an important seismichazard, threatening such large cities as Dhaka
(Bangladesh), Chengdu(China), Baghdad (Iraq), and Delhi (India).
Offshore, they form thelower boundaries of accretionary prisms in
subduction zones, and shouldbe considered in seismic and tsunami
hazard assessment in subductionzones around the world.
Unlike reverse faults, which form at dips of ~2060 according
toboth observation and theory, dcollements dip gently, at angles of
b110 (Davis et al., 1983). This is possible because these faults
take advan-tage of preexisting weaknesses in the rock, forming
along stratigraphichorizons with weak materials like salt or shale
(Suppe, 2007; Hubbardet al., 2010), in some cases with high pore
pressures (Behrmann et al.,1988; Bilotti and Shaw, 2005; Cubas et
al., 2013). The existence andlong-term deformation associated with
dcollements is understood notonly through observation (e.g., Ye et
al., 1997; Adam et al., 2004;Moore et al., 2009; Morley et al.,
2011), but also through laboratory(e.g., Malaveille, 2010;
Graveleau et al., 2012), computer (e.g., Strayeret al., 2001;
Burbidge and Braun, 2002), and theoretical modeling(Davis et al.,
1983; Dahlen et al., 1984; Dahlen, 1990). Large, activedcollements
are known to exist in regions both onshore (Himalayas;Taiwan;
Bolivia; Bangladesh; Sichuan, China;) and offshore (Sumatra,Java,
Japan, Peru, Cascadia, Antilles, Makran, Guatemala) (Davis et
al.,1983; Bilotti and Shaw, 2005; Hubbard et al., 2010; Morley et
al., 2011).
Many studies of seismic hazard have assumed that
becausedcollements are weak, they are unable to support large
stresses andstore sufficient elastic energy to produce hazardous
earthquakes, butmust rather be fully creeping, generating only
small and micro-earthquakes (e.g., Byrne et al., 1988; Hyndman et
al., 1997). Studies ofsubduction zones have generally observed an
updip limit to interplateseismicity that persists for several
decades (the seismic front, Byrneet al., 1988; Fig. 1),which has
led to the application of the term aseismicto the portion of the
dcollement underlying the accretionary prism. Al-though this term
is correctly applied in that we observe little seismicityin this
region, it has also been taken to mean that this part of the
wedgenever slips in association with moderate to large earthquakes
an as-sumption rather than an observation.
Experimental studies of clay and gouge materials show that
barerock surfaces and thin gouge layers exhibit potentially
unstablevelocity-weakening behavior, while slip within thick gouge
showsvelocity-strengthening behavior (Marone and Scholz, 1988).
Theupdip limit has been inferred to be associated with a zone of
thick,velocity-strengthening material in the accretionary prism
that resistsrapid rupture (Marone and Scholz, 1988). Thus, the
absence of observedmoderate to large earthquakes in this regionhas
been used to infer a slipbehavior that in turn has been used to
justify a stratigraphic model thatis consistent with creeping
behavior.As a consequence, until recently, manymodels of both
coseismic slipand interseismic coupling on subduction zones started
with the as-sumption that there was no coseismic slip at the tip of
the wedge, andthat the region updip of the seismic front was fully
creeping (Byrneet al., 1988; Hyndman et al., 1997; Chlieh et al.,
2007, 2008; LovelessandMeade, 2009). Inversions for coseismic slip
or interseismic couplingoften appear to support the idea that the
shallow region is creeping.However, these inversions often force
the near-trench area of the faultto creep during the interseismic
period or prevent it from slipping dur-ing earthquakes. Inferences
about the kinematic behavior of the toe ofthe prism are therefore
not usually directly supported by data, asdiscussed by Rhie et al.
(2007) and Loveless and Meade (2011).
We agree that dcollements appear to have long-term weak
behav-ior (Suppe, 2007), and that there is often a strong drop-off
in recordedseismicity updip of the seismic front. However, neither
of these observa-tions is a compelling reason to infer creeping
behavior. Seismic behavioris possible for a weak fault: for example
with a low effective confiningpressure of 50 MPa and a low
effective coefficient of friction of 0.1,there is enough frictional
resistance for a complete coseismic stressdrop of 5 MPa, a value
larger than that for typical inter-plate earth-quakes (Venkataraman
and Kanamori, 2004). The opposite can also betrue: a creeping fault
may not be weak, but rather be creeping at ahigher shear stress
than a stick-slip fault, averaged over the seismiccycle. And
indeed, contrary to the suggestion that low seismicity indi-cates
creeping behavior, current understanding of fault zones
rathersuggests the opposite: we often expect to see seismicity
where faultsare creeping, and little seismicity where they are
locked (e.g., Rubinet al., 1999; Barbot et al., 2013). Further,
these observations of limitedseismicity may be dependent on the
earthquake cycle: in Sumatra, theseismic front is clear prior to
the year 2000, but following the Mw 7.9earthquake in Southern
Sumatra in 2000 (and the great earthquakesin 2004 and 2005), many
earthquakes were recorded in the frontalpart of the system (Fig. 2:
Aceh, Nias/Simeulue, andMentawai 2010 seg-ments). In addition to
the temporal change in seismicity in Sumatra, wealso observe
tremendous spatial variability, with portions of themegathrust
exhibiting minimal seismicity everywhere, other showinga clear
seismic front, and yet others generating earthquakes everywhereup
to 300 km from the trench (Fig. 2). Thus, a single dcollement
mayhave extremely variable seismic patterns both spatially and
temporally.As described below, we can find evidence inmany
dcollement systemsfor large, episodic slip events at the tips of
wedge systems, often associ-ated with recorded earthquakes. This
indicates that current models offully creeping behavior on these
systems are flawed (Fig. 1).
Several recent earthquakes, including the 2010 Mw 7.8
Mentawaitsunami earthquake, Indonesia, and the 2011Mw 9.0
Tohoku-Oki earth-quake, Japan have prompted a reevaluation of the
assumption of creep-ing behavior near the trench (Lay and Bilek,
2007; McCaffrey, 2008;Avouac, 2011; Faulkner et al., 2011; Loveless
and Meade, 2011; Hillet al., 2012; Kozdon and Dunham, 2013) and the
development of newmodels where creep and coseismic slip occur at
different stages of theearthquake cycle (Noda and Lapusta, 2013).
Here, we present a reviewof additional data from these and other
dcollement systems aroundthe world, both on- and offshore, that
support the inference that theycan rupture in earthquakes
associated with large slip at their tips.
Although many studies implicitly treat subduction zones and
conti-nental fold-and-thrust belts as different classes, we show
throughoutthis paper that they are structurally similar (Fig. 3)
and exhibit muchthe same behavior. This paper brings together
perspectives from struc-tural geology, geodesy, and earthquake
dynamics, both on-land and
-
Fig. 1. Schematic model of a subduction zone, showing the
locations of different types of earthquakes. (A) Redrawn after
Byrne et al. (1988). (B) Our model. We suggest that thedcollement
at the base of the accretionary prism is seismogenic and/or capable
of participating in large thrust events on the subduction
interface. We suggest that the faults risingfrom the shallow
dcollement are also capable of participating in earthquakes, either
in addition to or instead of rupture all the way to the toe.
47J. Hubbard et al. / Earth-Science Reviews 141 (2015)
4555offshore.We suggest that exposed fold-and-thrust belts can
provide im-portant constraints and observations that can be used to
understand thebehavior of subduction zones, where observations are
much morelimited.2. Evidence for shallow slip in megathrust
earthquakes
Below, we describe a set of regions with large dcollements
showingevidence for earthquakes that rupture to their tips in large
slip events(Fig. 3; Shaw and Suppe, 1994; Ye et al., 1997; Lav and
Avouac, 2000;Ranero et al., 2000; Yue et al., 2005; Singh et al.,
2008; Moore et al.,2009; Uba et al., 2009; Shulgin et al., 2011;
Singh et al., 2011; Hubbardet al, 2014). This evidence falls
primarily into four categories: (1) lockingoccurs along the
dcollement, with geodetic observations of strain accu-mulating
downdip; (2) slip propagating to the toe of the system
overtimescales of hundreds of years to tens of thousands of years,
implyingthe toe is active; (3) slip that occurs episodically (i.e.
not as interseismiccreep), and (4) slip associated with
earthquakes, as opposed to justslow slip events (e.g., Meade and
Loveless, 2009).2.1. The Himalaya
One of the most compelling cases for large earthquakes
ondcollements is in the Himalaya. This ongoing
continent-continentcollision is now primarily accommodated along a
2000-km-widesubhorizontal thrust that dips gently beneath the
lesser Himalaya.In the Nepal Himalaya, the dcollement extends about
90 kmdowndip, and then steepens beneath the high Himalaya on a
blindthrust ramp (Schelling and Arita, 1991; Jackson and Bilham,
1994;Pandey et al., 1995). Much like subduction zones, the Himalaya
ex-hibits a seismic front, with intense microseismicity and
frequentsmall earthquakes along the ramp, and limited seismicity
along thedcollement itself (Pandey et al., 1995).
Geodetic measurements in the Himalaya demonstrate that
thedcollement is currently locked from the surface to about 100
kmdowndip (Bilham et al., 1997; Larson et al., 1999; Jouanne et
al., 2004;Betinelli et al., 2006; Ader et al., 2012). Over the ten
thousand year time-scale, the intraplatemotion is accommodated
bymigrating slip onto thepresently locked zone, andmost of this
slip propagates to the toe, basedon uplifted terraces on the
frontal fault of the system (Lav and Avouac,
-
Inve
stig
ator
Frac
ture
Zon
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40 km 9
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Sumatra
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a) Aceh segment
b) Nias/Simeulue segment
c) Mentawai seismic gap
d) Pagai
e) Mentawai 2010 patch
f) Enggano segment
a
b
c
d
e
f
N
pre-2000
post-2000
Fig. 2. (Left)Mapof Sumatra showing seismicity,filtered for
thrust events (data extracted from theGlobal CMT catalog). Contours
for the subduction zone are shownas thin black lines. Dashedlines
show boundaries of regions for seismicity histograms (a-f; shown on
the right). Focal mechanisms for earthquakes M N =7.0 are shown.
(Right) Histograms of seismicity along the Su-matran megathrust
seismicity (red) prior to the 2000 Mw 7.9 earthquake in Southern
Sumatra and the later great earthquakes, and (black) from 20002014,
for six different portions of themegathrust. Seismicity is binned
into 10-km stripes down-dip. The megathrust demonstrates strong
variability in seismicity, both temporally and spatially (along
strike and down dip).
48 J. Hubbard et al. / Earth-Science Reviews 141 (2015)
45552000). This happens in discrete events with large amounts of
slip at thetoe that can be tied to historical large earthquakes
(Sapkota et al., 2013;Bollinger et al., 2014). Thus, in Nepal it
appears not only that thedcollement is slipping in large events
that reach the surface, but alsothat these events are radiating
significant seismic energy.2.2. Bolivia
The eastern side of the Central Andes forms a backarc
fold-and-thrust belt underlain by a shallowly dipping dcollement.
Shorteningacross the region is estimated at ~7-13 mm/yr from both
geological
-
Fig. 3. Schematic diagrams of continental fold-and-thrust belts
and subduction zones from around the world that show evidence of
shallow dcollement earthquakes, shown at a fixedscale. Most of the
regions discussed in the text are represented here.
49J. Hubbard et al. / Earth-Science Reviews 141 (2015) 4555and
geodetic data (Uba et al., 2009; Brooks et al., 2011b). Like in
theHimalaya, geodeticmeasurements suggest that the dcollement is
lockedfrom the toe to about 100 kmdowndip (Brooks et al., 2011b).
In addition,Brooks et al. (2011a) have found geological evidence
for a surface-rupturing event with at least 7 m of slip at the
range-front fault at thetip of the system.2.3. Western Taiwan
The 1999 Mw 7.6 Chi-Chi earthquake represents a clear example
inwhich a shallow, bedding-parallel dcollement ruptured with large
sur-face slip (3-10 m) (Lee and Ma, 1999; Johnson et al., 2001; Yue
et al.,2005). Although the rupture involved a complex geometry and
faultproperties (Ma et al., 2000, 2003), it is clear that the
earthquake producedsignificant coseismic slip (and afterslip) on
the bedding-parallelChelungpu-Sanyi thrust system, which extends as
a shallow dcollement(5-8 km deep) beneath western Taiwan for tens
of kilometers (Yue et al.,2005; Rousset et al., 2012).
2.4. Santa Barbara Channel, Southern California
The southwesternmargin of the Transverse Ranges extends
offshoreinto the Santa Barbara Channel along a set of
bedding-paralleldcollements at 36 km depth, imaged by seismic
reflection data(Shaw and Suppe, 1994). Syntectonic sediments
deposited on thesestructures produce distinctive growth triangles
that record active slipin the Quaternary accommodating ~45% of the
geodetically measuredshortening rate across the channel. North of
the toe of the dcollement,the Ventura-Pitas Point fault splays
upward, producing the Ventura Av-enue anticline (Hubbard et al.,
2014); recent terrace measurementsdemonstrate that this fault
produces large, episodic uplift events(Rockwell, 2011). These large
uplift events would require rupture ofnot only the Ventura-Pitas
Point fault, but also the dcollement and
-
50 J. Hubbard et al. / Earth-Science Reviews 141 (2015)
4555other fault systems along strike (Hubbard et al., 2014). We
thereforeconclude that shortening is at least in part accommodated
by large, ep-isodic events on the dcollement system.
2.5. Sumatra subduction zone, Mentawai segment
The 2010 Mw 7.8 Mentawai earthquake produced a tsunami thatwas
much larger than expected based on the seismic magnitude, lead-ing
to its classification as a tsunami earthquake (Kanamori, 1972).Hill
et al. (2012) used GPS data and a tsunami field survey to
demon-strate that the earthquake must have been associated with
high faultslip at shallow depths close to the oceanic trench, with
most of theslip at depths shallower than 6 km. Yue et al. (2014)
used a joint inver-sion of high-rate GPS and teleseismic data to
confirm this result, placingeven higher slip (up to 23 m) in the
shallowest section (above 5 km).Such high slip at the tip of an
accretionary prism must be the result ofslip along the dcollement
at its base. This event provides a possible ex-ample of a large
earthquake that initiated updip of the seismic front,indicating
that the dcollement here may be not only seismic, butseismogenic.
Alternatively, it is possible that the rupture initiateddeeper, but
that the deeper slip was low in magnitude compared tothe shallower
slip, and is therefore not well constrained. Coralmicroatolls in
this region also record a slightly deeper, but still shallowslip
event in ~A.D. 1314, suggesting that the 2010 earthquake was notan
isolated event (Philibosian et al., 2012).
2.6. Sumatra subduction zone, northern segment
The greatMw9.2 Sumatra earthquake of 26December 2004 propagat-ed
~1300 km along strike and produced a devastating tsunami. Singhet
al. (2008) imaged large, active thrust faults at the front of the
accretion-arywedgewith seismic reflection data. They also note the
presence of af-tershocks with thrust mechanisms consistent with the
imaged faults.These observations suggest that the
2004earthquakemayhave propagat-ed updip to the tip of the
accretionary prism. This is compatiblewith geo-detic data (Rhie et
al., 2007), and supported by backprojections ofradiated
energy,which indicate significant energy release near the trenchfor
hundreds of kilometers along strike (Ishii et al., 2005). However,
theresolution of the slip distribution based on seismic energy
release andgeodetic deformation is poor, and several alternative
slip models fit thedata reasonably well with primarily deeper slip
patches (Ammon et al.,2005; Chlieh et al., 2007; Pietrzak et al.,
2007; Shearer and Brgmann,2010). This is a common problem for
offshore earthquakes far fromland. If significant slip near the
trench did occur, it could have producedlarge uplift of the
seafloor and contributed to the height of the tsunami.
Nearly a century prior to the 2004 Sumatran earthquake, a ~ M
7.8earthquake occurred near the southern extent of the 2004 event
in1907 (Kanamori et al., 2010). This earthquake produced an
extensivetsunami that affected nearly 950 km of the Sumatran coast.
Detailed in-vestigations of historical seismograms has led to the
conclusion that thisearthquake probably initiated on the subduction
interface, and propa-gated up-dip into the shallow sediments,
causing the large tsunami(Kanamori et al., 2010).
2.7. Java subduction zone
The 17 July 2006 Mw 7.8 Java earthquake produced a tsunami
withlocal runup heights over 20 m and average tsunami heights of 5
to7 m along 200 m of coastline (Ammon et al., 2006; Fritz et al.,
2007).Like the 2010 Mentawai earthquake, this qualifies as a
tsunami earth-quake, due to the size of the tsunami relative to the
earthquakemagni-tude (Kanamori, 1972). Also like the Mentawai
earthquake, a rupturetowards the toe of the accretionarywedge could
produce significant up-lift, and generate a tsunami of larger size
than would be predicted by adeeper rupture. This is consistentwith
the observation of low energy re-lease and slow rupture, features
that are inferred to be associated withrupture within the
low-rigidity materials of an accretionary prism(Newman and Okal,
1998; Polet and Kanamori, 2000). Thus, we inferthat the Java
portion of the Sunda subduction zone, like Sumatra tothe north, may
be capable of large ruptures in the frontal part of thewedge.
2.8. Japan Trench
The 2011 Mw 9.0 Tohoku-Oki earthquake ruptured the plate
inter-face between the Pacific and Okhotsk plates, east of the
island of Hon-shu. Tremendous slip vertically displaced the sea
bottom by up to50 m, as measured by differential bathymetry, GPS
measurements, un-derwater acoustic sounding, and sea-bottom
pressure records of thetsunami waves, demonstrating that rupture
reached close to the Japantrench (Avouac, 2011; Fujiwara et al.,
2011; Sato et al., 2011). Indeed,seismic reflection profiles pre-
and post-earthquake directly imageslip at the trench axis during
the earthquake (Kodaira et al., 2012),and demonstrate that the
dcollement underlying this prism slippedin this event, with large
coseismic slip reaching the toe of the wedge.
Shouldwe have known prior to the Tohoku-Oki earthquake that
thissubduction zone was capable of trench ruptures? Indeed, we
shouldhave. The Mw 8.5 1896 Sanriku earthquake ruptured a nearby
portionof the trench and generated a tsunami with a maximum run-up
of25 m, despite only weak shaking being felt onshore (Tanioka
andSatake, 1996). Tsunami modeling suggests that slip was
concentratednear the trench (Tanioka and Satake, 1996). Repeated
greater-than-expected tsunamis with weak shaking support the
hypothesis of multi-ple toe ruptures.
2.9. Kuril Trench
In addition to the two earthquakes in the southern part of the
Japan-Kuril trench, two tsunami earthquakes occurred further north:
a Ms 7.2event in 1963, and a Ms 7.0 event in 1975 (Fukao, 1979;
Pelayo andWiens, 1992). The Oct. 20 1963 earthquake occurred as an
aftershockof a larger, more down-dip earthquake on Oct. 13, and has
beeninterpreted as having a shallow origin (~5 km depth, Fukao,
1979;~9 km depth, Pelayo and Wiens, 1992). The June 10, 1979
tsunamiearthquake occurred about 300 km southwest of the 1963
earthquake;body wave inversion provides a best-fit source depth of
5 km, which isconsistent with the location of the epicenter
relative to the trench(Pelayo and Wiens, 1992). Both of these
events exhibited anomalouslyslow rupture speeds, suggesting that
they occurred within the weaksediments of the accretionary prism
(Pelayo and Wiens, 1992).
In 2006, a Mw 8.3 event ruptured just northeast of the 1963
earth-quake region (Ammon et al., 2008; Lay et al., 2009). Although
the tsuna-mi was relatively limited in this event (mostly b1.2 m),
this is in partbecause of the lack of local tide gauge recordings.
A finite sourcemodel for the rupture suggests that at shallow
depths (b25 km), a250 km long segment of the subduction zone
slipped an average of4.3-6.5 m (Lay et al., 2009). This calculation
constrains slip at the trenchto zero. It seems likely that this
rupture extended to the trench.
2.10. Japan, Nankai Trough
High-resolution seismic reflection profiles across the Nankai
Troughimage both a series of shallowly dipping thrust faults in the
frontal partof the wedge and a large thrust to the hinterland
termed a megasplaythatmergeswith the subduction interface at ~89
kmdepth (Park et al.,2002; Moore et al., 2009). Cores were drilled
through both the frontalthrust at the toe of the accretionary prism
and across the megasplay,and analyzed for vitrinite reflectance
geothermometry (Sakaguchiet al., 2011). It was determined that both
fault zones underwent local-ized temperatures of more than 380 C,
implying that frictional heatingoccurred, likely due to coseismic
slip. This evidence strongly suggeststhat the frontal dcollement
slipped coseismically at least once.
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51J. Hubbard et al. / Earth-Science Reviews 141 (2015)
4555Although it is not possible to link this recorded heating event
to a partic-ular earthquake, both the Mw 8.1 1944 Tonankai and the
Mw 8.1 1946Nankaido earthquakes are potential candidates, as they
producedground shaking and damaging tsunamis in this region (Kato
andAndo, 1997).2.11. Solomon Islands
The 2007 Mw 8.1 Solomon Islands earthquake produced a large
tsu-nami, with up to 12m of runup (Chen et al., 2009; Furlong et
al., 2009).Inversion of geodetic data and surveys of uplifted coral
reefs and othercoastal features suggests that there was up to 30 m
of slip at the trench(Chen et al., 2009). A few years later, the
2010 Mw 7.1 Solomon Islandstsunami earthquake producedwidespread
coseismic subsidencewithin20 kmof the San Cristobal trench (Newman
et al., 2011). Analysis of thenear-trench deformation, tsunami
run-up, open-ocean wave heightdata, and seismic records indicates
that the earthquake occurred onthe shallow part of the low-angle
dcollement in this region. The best-fit slip model suggests that
the maximum slip was close to the trenchand reached over 7 m
(Newman et al., 2011).2.12. Middle America megathrust
The September 2 1992 M 7.6 Nicaragua earthquake generated up
to10mof tsunami runup, but only exhibitedweak ground shaking.
Exten-sive studies of this earthquake and its associated tsunami
demonstratethat it ruptured a 40-50 kmwide fault plane extending
from the trenchto depths of ~10 km (Satake, 1994; Lay and Bilek,
2007). Finite-fault in-version based on teleseismic P wave
inversion shows slip down to~20 km depth, but with large amounts
within the upper 10 km (Yeet al., 2013). In 2012, another
earthquake ruptured the MiddleAmerica megathrust, this time a Mw
7.3 in El Salvador. Both this eventand the 1992 earthquake exhibit
low radiated energy for their seismicmoment and slow rupture
velocities, characteristics of tsunami earth-quakes (Kanamori and
Kikuchi, 1993; Ye et al., 2013). Although the2012 earthquake did
not produce a tsunami due to its relatively lowmagnitude,
finite-fault inversion based on teleseismic P-wave
inversionindicates that most of the slip was within the upper 15
km. The inver-sion model of Ye et al. (2013) imposes zero slip at
the trench, butexhibits a steep slip gradient in the upper 5
km.
Kanamori and Kikuchi (1993) suggest that the shallow rupture
in1992 may have been made possible by the lack of a
sedimentarywedge offshore, resulting in different frictional
properties than inother wedges. Heesemann et al. (2009) propose
that seamounts onthe subducting Cocos plate might alter the stress
state and promote dy-namic slip on a normally aseismic dcollement.
However, our discussionhere demonstrates that such ruptures have
occurred inmany other sub-duction zones without these
characteristics.2.13. Alaskan Trench
The 1964Mw 9.2 Great Alaskan earthquake ruptured the eastern
seg-ment of the Alaska-Aleutian Trench. Joint inversion of tsunami
and geo-detic data indicates that a large area of high slip
concentrated over PrinceWilliam Sound extended to the trench, with
an average of ~11 m of slipin the region above 15 km depth (Johnson
et al., 1996). However,the resolution of this slip model is
limited. Slip certainly extendedas far out as Middleton Island,
which lies about 8 km above themegathrust: uplift in the earthquake
produced the most recent ofsix beach terraces visible on the island
(Savage et al., 2014). Wheth-er the slip continued to the trench or
instead was diverted onto thesplay fault that is presumably
responsible for the growth of this is-land (or both) is not
known.2.14. Peru
The Mw 7.6 Peru tsunami earthquake of Nov. 20, 1960, had a
muchlonger rupture duration than would have been otherwise expected
foran earthquake of this magnitude (Pelayo and Wiens, 1990),
suggestingthat it occurred within the low-rigidity accretionary
prism. Body wave-form inversion indicates that it ruptured a fault
plane dipping ~6 closeto the trench axis (Pelayo and Wiens, 1992).
This earthquake occurredin a portion of the Peru trench that
previously had beenmostly aseismic,leading to the suggestion that
itmight not be capable of generating largeearthquakes (Nishenko,
1991). However, the 1960 event demonstratedthat shallow,
tsunamigenic earthquakes can occur in this region, andmay occur in
other, apparently aseismic regions as well.
The 1960 earthquake was followed by two other events with
evi-dence of shallow slip on the Peruvian subduction zone: a Mw 7.4
tsuna-mi earthquake in February 1996, which generated a large
tsunami thanexpected from its surfacemagnitude and exhibited a long
rupture dura-tion (Heinrich et al., 1998), and a Mw 8.5 earthquake
in June 2001, forwhich joint inversion of seismic and geodetic data
suggest significantslip above 15 km depth, reaching close to the
trench (Pritchard et al.,2007).
2.15. Chile
The great 1960 Mw 9.5 Chile earthquake was the largest
earthquakeever recorded. Inversions for the slip distribution from
ground deforma-tion using a 3D fault model suggest that there was
at least 1 m of slip atthe trench for a distance of 200 km, with 50
km of that distanceexperiencing over 5 m of slip (Barrientos and
Ward, 1990). In addition,they note that substantial undetected
offshore slip is likely, given thelarge tsunami generated by the
earthquake. Moreno et al. (2009) usedthe same geodetic dataset
combined with a 3D fault model to estimatethat over 200 km of the
trench along strike slipped N10 m, with over50 km of that distance
experiencing N20 m of slip. A joint inversion ofgeodetic and
tsunami data for the 1960 earthquake estimates 13-21 mof slip near
the trench in the southern part of the source (Fujii andSatake,
2013). These studies demonstrate the importance of
consideringtsunami data in slip inversions, since geodetic data
onshore cannot ef-fectively constrain slip occurring at the
trench.
2.16. Summary
The fifteen regional examples that we describe include
continent-continent, continent-ocean, and ocean-ocean convergent
settings, andeven oblique portions of transform margins; subduction
zones withold and young downgoing plates, with and without sediment
deposi-tion at the trench. They include fold-and-thrust belts that
extend for1000s of km and b100 km along strike, and systems that
deform pri-marily through forethrusts, backthrusts, and
combinations of the two.They include very large and moderate-sized
earthquake ruptures, andruptures that initiate at deep and shallow
levels. Thus, it is not possibleto point to one common feature that
makes rupture possible on theshallow dcollement, implying that
these faults may always have thepotential to be seismic and/or
seismogenic.
3. How can weak decollements behave seismically?
Given that there is evidence for large, episodic slip at the
tips ofdcollements, wemust reevaluate the assumption that
dcollementsare too weak to support stresses and store large strain
energy. How-ever, critical taper wedge mechanics implies that these
are weak fea-tures, significantly weaker than the surrounding rocks
(Suppe,2007). Reconciling these apparently conflicting views
requires usto examine our assumptions about fault strength,
earthquake dy-namics and rock properties.
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52 J. Hubbard et al. / Earth-Science Reviews 141 (2015) 45553.1.
The state of stress within the wedge, and preferred slip planes
atgeological time scales
We know that dcollements are weak because they form and slip
atunfavorable angles. However, we could turn this around and say
insteadthat these faults are oriented at unfavorable angles, and
therefore the dif-ferential stress in the ambient rocksmust be high
in order for these faultsto slip. In fact, because we see thrust
faults rising from dcollements, weknow that sometimes it is easier
to break a more steeply dipping thrustfault and slip on it than to
slip on the dcollement surface. At othertimes, however, it must be
easier to slip on the dcollement; we knowthis because we see slip
at the toes.
How is it possible that both of these options are favorable at
differenttimes? We can use a Mohr diagram, which describes the
normal andshear stresses acting on a plane of any orientation, and
comparesthem to failure criteria, to try to understand this problem
(Fig. 4). Wepropose a setting where minor changes in stress or rock
propertiescan cause slip on either the dcollement or on a thrust
fault to be pref-erable at a given time.Fig. 4.Mohr circle analysis
of failure in a critical taperwedge. (A) Basic concept of aMohr
circle.at an orientation that forms an anglewith respect to the
primary stress orientation1. The sizeThe black line represents the
failure criterion of the material, with slope = the angle of
interhesion (here considered cohesionless). If the circle
intersects the failure criterion, thematerial wAlaskanwedge at the
toe of thewedge andwithin the interior. Herewe consider the
potential foangles of internal friction for the ramp and the
detachment, respectively. A value of r=0.64wafor rupture). d can
have a value as low as 0.035 without triggering slip on the
dcollement atshaded blue area on theMohr circles represents this
range of possible d values. Note that theseremain the same. At the
toe of the wedge, 1 is horizontal, and the ramp reaches failure,
even ttowards the toe, rotating the planes represented on theMohr
circle clockwise. This causes the dfor both the unrotated and
rotated 1 orientation, to better illustrate the effect of the
rotation.However, we do not see randomvariations in slip on the
dcollementvs. thrusting. Instead, we typically observe in-sequence
break-forwardpropagation: the thrust faults towards the toe are
more active, andthose in the hinterland have been abandoned
(Morley, 1988). Thus, rup-tures today are typically choosing to
continue along the dcollement allthe way to the toe rather than to
break upward along preexisting thrustfaults, even though they
previously chose those branches. This pattern ofbreak-forward
propagation can be understood by the fact that the prin-cipal
stresses are rotatedwithin thewedge due to the topographic
gradi-ent (Dahlen et al., 1984; Savage et al., 1985; Cubas et al.,
2008;Mary et al.,2013). The maximum stress direction dips gently in
the direction of thewedge toe (see 1 orientation on the right side
of Fig. 4). This increasesthe angle between the maximum stress
orientation, 1, and thedcollement. There is a lateral transition in
the amount of stress rotation,with 1 horizontal just beyond the toe
of the wedge to tilted in the inte-rior of the wedge. A larger
angle between the 1 direction and thedcollement dip is equivalent
to a smaller in the Mohr diagram(Fig. 4a). This makes slip on the
dcollement more favorable. Thus, weexpect ruptures to follow the
dcollement until they reach a pointEvery point on the circle
represents the normal () and shear () stresses acting on a planeand
location of the circle is defined by the largest (1) and smallest
(3) principal stresses.
nal friction, and the y-axis intercept for the failure criterion
(here zero) represents the co-ill fail on the plane represented by
the point of intersection. (B)Mohr circle analysis of ther failure
of the dcollement and a ramp rising from the dcollement. r and d
represent thes chosen tomatch the observed dips of the fault ramps
(assumed to be the ideal orientationthe toe, and must have a value
below 0.187 to allow slip in the interior of the wedge.
Thevalueswould change in the presence of pore pressures or
cohesion, but the principlewouldhough the rampmaterial has a higher
coefficient of friction. In the interior, 1 dips slightlycollement,
rather than the ramp, to reach failure. TheMohr circle on the right
shows lines
-
53J. Hubbard et al. / Earth-Science Reviews 141 (2015)
4555(usually near the toe)where the stressfield is less tilted, and
at that pointto break upward along a thrust fault. In other words,
the slip on theramps builds a topographic gradient that causes
those ramps to shutoff, and the dcollement to propagate
forward.
3.2. Variations in strength through the earthquake cycle
There are additional mechanisms that could allow stress to
accumu-late on dcollements. In particular, the interseismic and
coseismicstrengths of the dcollement may differ. In this case, the
long-term re-cord visible as the total deformation of the systemmay
reflect theweak-er of the two. Low-temperaturemechanical processes
ofweakening andhealing (abrasion, pulverization, and fusion) are
not likely to be able todramatically affect effective friction
(Biegel et al., 1992; Wang andScholz, 1994). However, at high slip
velocities, frictional heating mayallow strong-weakening phenomena,
such as melt-welts (Brown andFialko, 2012) and other
flash-weakening effects (Rice, 2006) and porepressurization through
thermal expansion (Ghabezloo and Sulem,2009; Ferri et al., 2010) or
mineral decomposition (Han et al., 2007).Strong-weakening
mechanisms are activated under specific conditions,so we can expect
different frictional resistance at subseismic and seis-mic slip
speeds. For example, reaching sufficiently high temperaturefor
melting or for substantial expansion of pore fluids requires
largeslip (Rice, 2006), so these strong-weakening effects are
expected tooccur mostly during large earthquakes. But the
complexity of fault rhe-ology may also affect the interseismic
period. For example, slow slipevents occurred near the hypocenter
of the 2011 Mw 9.0 Tohoku-Okiearthquake and the 2014 Mw 8.1
Iquique, Chile earthquake with signif-icant overlap with the
seismic rupture (Ito et al., 2013; Kato andNakagawa, 2014; Ruiz et
al., 2014), and models of fault slip evolutionover many earthquake
cycles suggest the possibility that the same seg-ment may
experience creep, slow slip, and seismic ruptures during dif-ferent
periods (Noda and Lapusta, 2013; Noda and Hori, 2014).
Thisindicates that the seismogenic potential of faults may not be
fullyassessed during a short period of observation.
4. Consequences of seismogenic decollements
Dcollements are some of the largest faults on Earth.
However,existing scaling relationships between fault area and
earthquakemagni-tude (Wells and Coppersmith, 1994; Hanks and Bakun,
2008) may beinappropriate for dcollements, as they are structurally
different fromtypical dip-slip or strike-slip faults. In addition,
weak dcollementsand accretionarywedges exhibit much lower rigidity
thanmore consol-idated rock (Kanamori, 1972). Because earthquake
moment is definedas the product of the rigidity, slip, and area of
a fault, lower rigiditywill result in lower earthquake magnitudes
for a given amount of slipand area (Fukao, 1979).
Such an alteration to the typical slipmagnitude relationship
wouldreduce expected seismicmoments in regionswhere hazard is
estimatedfrom slip, uplift amounts, or GPS shortening. Initially,
one would expectthis alteration to reduce seismic hazard. However,
because the slipwould be expected to occur directly beneath large
regions at shallowdepths, the resulting ground shaking levelsmight
be as ormore hazard-ous than that predicted from higher seismic
moments on more distantfaults.
In contrast, if seismic hazard is estimated from earthquake
moment,we may see earthquakes with higher slip than expected based
on seis-mic records. This would help explain tsunami earthquakes,
as evenallowing slip to propagate to the tips of accretionary
prisms is not al-ways sufficient to explain resulting tsunami
heights (e.g., Hill et al.,2012).
This discrepancy between earthquake magnitude and slip can
alsobe partially solved by structural analysis. Seismic reflection
imaging ofaccretionary prisms demonstrates that dcollement tips
frequently ter-minate into more steeply dipping thrust faults that
break towards thesurface (Fig. 3). If slip propagates to the tip of
the dcollement, we canexpect that it will then extend upward onto
such a thrust fault and pro-duce a band of high uplift at the toe
of the wedge, resulting in largertsunamis.
5. Discussion and conclusions
For many years, studies of subduction zones have assumed that
thefrontal portions of accretionary prisms were too weak to
rupturecoseismically (Byrne et al., 1988; Hyndman et al., 1997).
Hazard studiesof fold-and-thrust belts on land have often suffered
from the same as-sumption, based on the assumption that the shallow
dcollement istoo weak to rupture in large earthquakes. We present a
series of exam-ples, both on-land and offshore, demonstrating that
inmany cases, shal-low dcollements are capable of producing large,
coseismic slip eventsthat rupture to the toes of the systems. Some
of these events are associ-ated with ruptures that initiate
down-dip of the seismic front, whileothers are limited to the
frontal, shallow portion of the wedge
We suggest that this behavior is not limited to the examples
de-scribed here, but rather is common to many (perhaps most)
accretion-ary wedges and fold-and-thrust belts. Although many
earthquakes insubduction zones have been interpreted to have no
slip at the tip ofthe accretionary prism, this interpretation is
typically driven by modelassumptions, rather than the data. In
addition to the examples providedhere, there may bemany other
examples of similar earthquakes, whereknown slip downdip obscures
the need for slip at the toe (hidden tsu-nami earthquakes).
We do not characterize the regions and events described here as
un-usual, as they encompass a wide range of settings. However, we
recog-nize that as far as reliably documented events go, these
events are stillfairly limited.Whether this fact is due to a lack
of data or a lack of eventsis not clear, and accurately assessing
this distinctionmay require the ac-quisition of new forms of data,
including pre- and post- bathymetry sur-veys and seafloor geodesy
(as has been used to observe deformation inthe 2011 Tohoku-Oki
earthquake). The fraction of slip at the trench thatoccurs in large
seismogenic events is unknown, and likely varies be-tween
dcollements, and thus assessing the recurrence rates of
suchearthquakes will in many cases be difficult, especially for
subductionzones, where paleoseismic evidence is lacking. As a
result, it would besafest to assume that trench-rupturing events
can occur until we haveenough evidence to show that a particular
fault system is behaving dif-ferently. Many attempts to assess
seismic hazard for particular regionshave relied on extrapolating
the likelihood of large earthquakes fromthe occurrence of small
ones, resulting in unexpected large earth-quakes with striking
underestimations (by two to three orders of mag-nitude) of expected
fatalities (Kossobokov and Nekrasova, 2012; Wysset al., 2012). This
study indicates that there is an urgent need for studiesthat
evaluate seismic and tsunami hazard in fold-and-thrust belts
andsubduction zones around theworld and allow for the possibility
of shal-low dcollement rupture.
Acknowledgements
We thank our two reviewers, Jeff Freymueller and Thorne Lay,
fortheir constructive comments. This research was supported by the
Na-tional Research Foundation of Singapore under the NRF
Fellowshipscheme (National Research Fellow Award No.
NRF-NRFF2013-06) andby the EOS, the National Research Foundation of
Singapore and theSingapore Ministry of Education under the Research
Centres of Excel-lence initiative. This is EOS paper number 76.
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Coseismic slip on shallow dcollement megathrusts: implications
for seismic and tsunami hazard1. Introduction2. Evidence for
shallow slip in megathrust earthquakes2.1. The Himalaya2.2.
Bolivia2.3. Western Taiwan2.4. Santa Barbara Channel, Southern
California2.5. Sumatra subduction zone, Mentawai segment2.6.
Sumatra subduction zone, northern segment2.7. Java subduction
zone2.8. Japan Trench2.9. Kuril Trench2.10. Japan, Nankai
Trough2.11. Solomon Islands2.12. Middle America megathrust2.13.
Alaskan Trench2.14. Peru2.15. Chile2.16. Summary
3. How can weak decollements behave seismically?3.1. The state
of stress within the wedge, and preferred slip planes at geological
time scales3.2. Variations in strength through the earthquake
cycle
4. Consequences of seismogenic
decollementsAcknowledgementsReferences