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Physics of the Earth and Planetary Interiors 186 (2011) 49–58
Contents lists available at ScienceDirect
Physics of the Earth and Planetary Interiors
journal homepage: www.elsevier.com/locate/pepi
Control of high oceanic features and subduction channel on earthquake ruptures
along the Chile–Peru subduction zone
Eduardo Contreras-Reyes a,∗ , Daniel Carrizo a,b
a
b
Departamento de Geofísica, Facultad de Ciencias Físicas y Matemáticas, Universidad de Chile, Santiago, Chile
Advanced Mining Technology Center AMTC, Facultad de Ciencias Físicas y Matemáticas, Universidad de Chile, Santiago, Chile
a r t i c l e
i n f o
Article history:
Received 8 September 2010
Received in revised form 8 March 2011
Accepted 8 March 2011
Edited by: George Helffrich.
Keywords:
Bathymetry
Asperity
Barrier
Subduction channel
Earthquake rupture
1960 and 2010 megathrust earthquakes
a b s t r a c t
We discuss the earthquake rupture behavior along the Chile–Peru subduction zone in terms of the buoyancy of the subducting high oceanic features (HOF’s), and the effect of the interplay between HOF and
subduction channel thickness on the degree of interplate coupling. We show a strong relation between
subduction of HOF’s and earthquake rupture segments along the Chile–Peru margin, elucidating how
these subducting features play a key role in seismic segmentation. Within this context, the extra increase
of normal stress at the subduction interface is strongly controlled by the buoyancy of HOF’s which is
likely caused by crustal thickening and mantle serpentinization beneath hotspot ridges and fracture
zones, respectively. Buoyancy of HOF’s provide an increase in normal stress estimated to be as high as
10–50 MPa. This significant increase of normal stress will enhance seismic coupling across the subduction interface and hence will affect the seismicity. In particular, several large earthquakes (Mw ≥ 7.5) have
occurred in regions characterized by subduction of HOF’s including fracture zones (e.g., Nazca, Challenger
and Mocha), hotspot ridges (e.g., Nazca, Iquique, and Juan Fernández) and the active Nazca-Antarctic
spreading center. For instance, the giant 1960 earthquake (Mw = 9.5) is coincident with the linear projections of the Mocha Fracture Zone and the buoyant Chile Rise, while the active seismic gap of north
Chile spatially correlates with the subduction of the Iquique Ridge. Further comparison of rupture characteristics of large underthrusting earthquakes and the locations of subducting features provide evidence
that HOF’s control earthquake rupture acting as both asperities and barriers. This dual behavior can be
partially controlled by the subduction channel thickness. A thick subduction channel smooths the degree
of coupling caused by the subducted HOF which allows lateral earthquake rupture propagation. This may
explain why the 1960 rupture propagates through six major fracture zones, and ceased near the Mocha
Fracture Zone in the north and at the Chile Rise in the south (regions characterized by a thin subduction
channel). In addition, the thin subduction channel (north of the Juan Fernández Ridge) reflects a heterogeneous frictional behavior of the subduction interface which appears to be mainly controlled by the
subduction of HOF’s.
© 2011 Elsevier B.V. All rights reserved.
1. Introduction
Along the Chile–Peru Trench the oceanic Nazca plate subducts
beneath the South American plate at a current rate of ∼6.5 cm/year
(Khazaradze and Klotz, 2003). The Nazca plate hosts a large population of high oceanic features (HOF’s) including hotspot tracks
and oceanic fracture zones (FZ’s). Perhaps the most striking feature of the oceanic plate is the Chile Rise, an active spreading
center that marks the boundary between the oceanic Nazca and
Antarctic plates and is currently subducting at ∼46◦ S beneath
the South American plate. Bathymetric data show four main
oceanic hotspot tracks: Nazca, Iquique, Copiapó and Juan Fer-
∗ Corresponding author. Tel.: +56 2 9784296; fax: +56 2 6968686.
E-mail address: [email protected] (E. Contreras-Reyes).
0031-9201/$ – see front matter © 2011 Elsevier B.V. All rights reserved.
doi:10.1016/j.pepi.2011.03.002
nández Ridges (Fig. 1). The Nazca Ridge (NR) was formed at
the Easter Island hotspot on the Pacific-Farallon/Nazca spreading center (Pilger, 1984), while the Juan Fernández Ridge (JFR) is
an off-ridge hotspot formed at the JFR hotspot onto 27 Myr old
oceanic crust (Yáñez et al., 2001). The oceanic Nazca plate is further
segmented by several oceanic fracture zones formed at the PacificNazca and Antarctic-Nazca spreading centers (Tebbens et al.,
1997).
Most of the seafloor features hosted by the oceanic Nazca plate
are in the throes of being subducted (Herron et al., 1981; Yáñez
et al., 2001; Rosenbaum et al., 2005). Subduction of a HOF may
modify the geodynamics and tectonic setting of the outer forearc
region, and disrupt and erode material from the overriding plate.
In addition, subduction of a HOF is expected to influence dramatically the degree of coupling across the subduction interface and
may affect seismicity, in particular the size and frequency of large
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Fig. 1. (A) Main oceanic bathymetric features and most significant earthquakes (red elipses) from the 19th to 21st centuries are shown. Bathymetry is taken from Sandwell
and Smith (1997) and main oceanic fracture zones are taken from Tebbens et al. (1997). (B) Earthquake ruptures (Mw > 7.0) were compiled from Beck et al. (1998), Comte and
Pardo (1991), Cisternas et al. (2005) and Bilek (2010). (For interpretation of the references to color in this sentence, the reader is referred to the web version of the article.)
earthquakes (Kelleher and McCann, 1976; Cloos, 1992; Scholz and
Small, 1997).
The slip distribution at the subduction interface is highly heterogeneous in terms of the strength or frictional characteristics of
the material in the fault zone. These stress heterogeneities have
been recognized as asperities (Kanamori, 1994) and barriers (Das
and Aki, 1977) to earthquake rupture which control the seismic
moment release and rupture area. A seismic asperity is an area with
locally increased friction and exhibits little aseismic slip during the
interseismic period relative to the surrounding regions. Once the
shear yield stress o (critical shear stress required for failure) along
these heterogeneities is reached by the accumulated interseismic
shear stress 1 , the asperity concentrates the coseismic moment
release and slip during the earthquake. The regions with little or no
slip during rupture propagation are generally called barriers and
control the size of the earthquake rupture area (ERA) (Aki, 1979;
Kanamori and McNally, 1982).
Once the earthquake initiates ( 1 ≥ o ), the rupture front propagates through regions where the dynamic stress associated with
the rupture process is larger than o . If the rupture front encounters an obstacle where the dynamic stress is lower than o , the fault
motion slows down and eventually stops (barrier zone). In other
words, the interplay between the amount of dynamic stress and o
will define whether a region behaves as an asperity with large slip
or behaves as a barrier with no slip. Therefore, it is fundamental
investigate the shear yield stress distribution along the subduction
interface. A heterogeneous distribution of o might be useful to predict the location of local barriers and asperities before occurrence of
an earthquake. o varies over time and depends on frictional properties, fluid pore-pressure, and state stress anomaly. For example,
subduction of a rigid seamount will result in an increase of normal
stress n and hence o preventing eventual rupture (strong coupling). Likewise, buoyancy of HOF’s will enhance seismic coupling
and anomalous increase of o . In contrast, an increase of fluid pore
pressure will reduce o facilitating rupture propagation.
Since HOF’s dramatically modify the geodynamics of the outer
forearc and interplate boundary conditions, they are probably the
most obvious candidates for asperities and/or barriers. The frictional behavior of HOF’s is clearly influenced by the increase of
normal stress at the subduction interface caused by the excess
buoyancy of these features (Scholz and Small, 1997). Furthermore,
Scholz and Small (1997) claim that the “geometric” normal stress
effect due to the need of the overriding plate to accommodate the
shape of the HOF is a significant additional source of increasing normal stress at the subduction interface. Thus, the anomalous strong
coupling associated with HOF’s (buoyancy and geometry) might
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E. Contreras-Reyes, D. Carrizo / Physics of the Earth and Planetary Interiors 186 (2011) 49–58
play an important role on seismicity. This appears to be the case
for South America, where many large earthquakes occur in regions
where these large Nazca plate features enter the trench.
We study the spatial relationship between large underthrusting earthquakes (Mw ≥ 7.5) and bathymetric heterogeneities on the
subducting Nazca oceanic plate along the Chile–Peru subduction
zone. Such heterogeneities are manifested by subduction of bathymetric anomalies (HOF’s), which are likely to affect the seismic
coupling across the subduction interface due to their buoyancy.
For instance, hotspot seamounts usually host thick oceanic crust
and magmatic material that may have erupted, ponded beneath
or intruded the base of the oceanic crust (magmatic underplating)
(Koppers and Watts, 2010). On the other hand, oceanic fracture
zones are characterized by a high degree of hydration in both the
crust and mantle, and they commonly comprise bodies of serpentinites deep into the oceanic upper mantle (Contreras-Reyes et al.,
2008). Thus, buoyancy is linked directly to structural or chemical
alteration of the crust or mantle beneath the oceanic feature that
stores rocks less dense than its surroundings. We use published
seismic velocity models across the Nazca Ridge (NR), JFR and Mocha
FZ in order to estimate the anomalous normal stress associated with
the buoyancy of these HOF’s. In addition, several HOF’s are identified from bathymetric data and compared with published coseismic
slip models of large underthrusting earthquakes in order to explore
if subducting HOF correlate with zones of large slip (asperities)
and/or little slip (barriers).
In addition, the amount of subducted sediment may also play
an important role on megathrust seismicity since the subduction channel might smooth the subduction interface resulting in
a homogenous coupled region. It can be expected that a large
homogenous coupled region facilitates long lateral rupture propagation (Ruff, 1989; Scherwath et al., 2009; Contreras-Reyes et al.,
2010) and hence the earthquake size. Therefore, a thin subduction
channel above a subducting HOF may reflect the presence of a seismic barrier while a thick subduction channel would tend to reduce
the effect of the anomalous coupled region caused by the subducting HOF. In order to test this hypothesis for some earthquakes
recorded along the Chilean subduction we use a compilation of
geodetic data and seismic data constraining the subduction channel
thickness and we discuss the interplay between subduction channel thickness and the dual asperity-barrier behavior. Identification
of both subduction channel thickness variation and magnitude of
buoyancy forces associated to HOF’s would provide a very useful
context for future seismological works.
2. Relation between high oceanic features and earthquake
rupture areas along the Chile–Peru subduction zone
The Chile–Peru subduction zone is an extremely active convergent plate margin producing large earthquakes (Mw > 8) about
every 10 years and capable of generating at least one tsunamigenic
mega-thrust earthquake per century. This is documented in historical accounts and instrumentally well-recorded events that date
back to the 14th century (Darwin, 1851; Beck et al., 1998; Bilek,
2010). Fig. 1A shows the location of the rupture areas of the largest
events documented over the 19th to 21st centuries. These ERA’s
are spatially linked to the locations of the main oceanic incoming
features of the Nazca plate. One of the most important observations
is that at least one rupture limit for each event coincides approximately with a subducting HOF, and about half of the ERA’s are
bounded by HOF’s. In the following we discuss some examples:
(a) The southern Peru–northern Chile region (15.5–23◦ S): this
region is characterized by two active seismic gaps that mark the
zones of the (Mw = 8.7) 1868 southern Peru and the (Mw = 8.9)
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1877 northern Chile mega-thrust earthquakes (Comte and
Pardo, 1991). The southern Peru 1868 rupture segment broke
previously in 1604 and 1784 in a similar, approximately 450 km
long rupture. The northwest part of that segment broke again
in an Mw = 8.4 earthquake in 2001, filling at least partially the
south Peru gap (Ruegg et al., 2001). This event propagated for
∼70 km before encountering a 6000 km2 fault area that acted as
a barrier. Approximately 30 s after the start of rupture, this barrier began to break when the main rupture front was ∼200 km
from the epicenter (Robinson et al., 2006). In this case, the subducted feature acted as both a barrier and an asperity during
a single event and was associated with the oceanic Nazca FZ
(Robinson et al., 2006). Further to the south, the northern Chile
seismic gap (18.5–23◦ S) is roughly coincident with the subduction of the Iquique Ridge (IR). Furthermore, the 2007 Tocopilla
earthquake (Mw = 7.7) spatially correlates with the southern
part of the IR, while the 1995 Antofagasta event (Mw = 8.0) is
bounded by the IR and Tal–Tal ridge in the north and south,
respectively (Figs. 1A and 2).
(b) Central Chile (31–37.5◦ S): the most recent large Chilean
earthquake occurred on February 27th, 2010 (Mw = 8.8) and
propagated northward and southward achieving a final rupture
length of about 450 km (34–38◦ S; Lay et al., 2010). In the same
region, 176 years ago in 1835 a large earthquake with an estimated magnitude of Mw ∼ 8.5 was reported by Darwin (1851).
North of the rupture area of the 2010 earthquake, two megathrust earthquakes occurred: the 1906 (Mw ∼ 8.4; Beck et al.,
1998) and 1985 (Ms = 7.8; Barrientos, 1995) events. The southern limit of the 2010 earthquake is coincident with the Mocha
FZ, while the northern limits of the 1906 and 1985 earthquakes
are coincident with the subduction of JFR (Fig. 1). Similarly, the
southern limit of the 1943 (Ms = 7.9) earthquake correlates with
the JFR.
(c) The 1960 mega-thrust event (37.5–46◦ S): the largest instrumentally recorded earthquake, which occurred in 1960 with
a magnitude of Mw = 9.5, ruptured over a length of almost
1000 km from ∼37.5◦ S to 46◦ S near the Chile Triple Junction
(CTJ) of the Nazca, Antarctic and South American plates (Moreno
et al., 2009). The Chile Rise is currently subducting at the CTJ,
which is a region characterized by a shallow trench associated
with the buoyancy of this active spreading ridge (Herron et al.,
1981). The 1960 earthquake ruptured across six major fractures
zones and it stopped near the Mocha FZ to the north and at
the CTJ to the south. In this case, the fractures zones involved
did not act as barriers, except for the Mocha FZ and Chile Rise
which are the northern and southern earthquake rupture limits,
respectively.
Other examples showing spatial correlations between ERA locations and incoming HOF’s can be inferred from Fig. 1A. In addition,
such spatial correlations also appear to apply for the historical seismic segmentation of the last 500 years (see Fig. 1B). Considering
the historical ERA estimates, not every limit of an ERA coincides
with the subduction of an HOF, but at least one limit of each ERA
is coincident with a prominent HOF as is shown in Fig. 1A, and
half of the events are bounded at both limits by HOF’s. Thus, the
large earthquakes of the Chile–Peru margin appear to be highly
segmented by bathymetric features. Obviously, the uncertainty of
earthquake rupture areas are large for historical reports and the
correlation between HOF and ERA position is far from perfect for
old events (Fig. 1). Nevertheless, events instrumentally recorded
during the 20th to 21st centuries present a reasonably good correlation between HOF and ERA position (Fig. 1). Additionally, geodetical
data provide further constraints for the size of ERA’s and fault slip
distribution as we discuss in the following section.
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Fig. 2. (A) Southern Peru-northern Chile seismic gap. Ellipses represent the approximate extent of rupture zones of the 1868 (Mw = 8.7) and 1877 (Mw = 8.9) events after
Comte and Pardo (1991). The rupture area of the 2001 Arequipa earthquake (Mw = 8.4) is almost coincident with the middle part of the 1868 rupture area. In northern Chile,
no great event had been recorded since 1877. The Mw = 8.1 Antofagasta earthquake of 1995 July 30 ruptured an 180 km long segment located south of the 1877 rupture zone.
The 2007 Tocopilla earthquake Mw = 7.7 is almost coincident with the southern part of the 1877 rupture zone. For those earthquakes with known distributed slip we use the
outermost contour to represent the rupture area (for the 1995 earthquake from Chlieh et al. (2004), for the 2001 Arequipa earthquake from Pritchard et al. (2007) and for the
2007 earthquake Béjar-Pizarro et al. (2010). Starts denote epicenter locations. (B) Fault slip model after Chlieh et al. (2004) and Béjar-Pizarro et al. (2010) for the 1995 and
2007 events, respectively. Slip distribution with 3 m contour interval of the 2001 Arequipa earthquake based on the coseismic fault slip model of Pritchard et al. (2007). Gray
and salmon regions represent the swell associated to the Iquique and Tal–Tal Ridges, respectively. Red dots denote the largest seamounts identified in A. (For interpretation
of the references to color in this sentence, the reader is referred to the web version of the article.)
3. Geodetical observations
Geodetical data provide further and crucial evidence for seismic coupling at the subduction interface, allowing identification
of regions with high coseismic slip (asperities) or no slip (barriers). In this section, we show a compilation of fault slip models for
(1) the active seismic gap of the southern Peru-north Chile region
(15.5–23◦ S), (2) Central Chile (31–38◦ S), and (3) the region of the
largest ever recorded 1960 earthquake (∼38–46◦ S).
3.1. The active seismic gap of south Peru-north Chile
The region between the Ilo Peninsula (15.5◦ S, South Peru) and
the Mejillones Peninsula (23.5◦ S, North Chile) represents a major
seismic gap that has not experienced a significant megathrust
earthquake since the 1868 earthquake (Mw = 8.7) and the 1877
earthquake (Mw = 8.9) (Comte and Pardo, 1991, Fig. 2A). The 1995
Antofagasta (Mw = 8.1) and the 2001 Arequipa (Mw = 8.4) earthquakes appear to have provided an extra load at both extremities
of the remaining ∼500 km length unruptured segment. After the
1877 event and before the 2007 Tocopilla earthquake, a few events
(Mw < 7.5) have been reported in the region (Comte and Pardo,
1991; Tichelaar and Ruff, 1991) but they were not large enough
to release a significant part of the 9 m of slip deficit accumulated in
the gap in the last 130 years (Béjar-Pizarro et al., 2010).
Fig. 2B shows the slip distribution with 3 m contour interval of
the 2001 Arequipa earthquake (Mw = 8.4) based on the coseismic
slip model of Pritchard et al. (2007). This event only partially filled
the rupture zone of the 1868 earthquake (Fig. 2A). The rupture front
propagates southward and encounters the Nazca FZ at 70 km from
the hypocenter. Although, this HOF apparently acted as barrier, the
rupture continued around this barrier, which remained unbroken
for ∼30 s and then began to break when the main rupture front was
∼200 km from the epicenter (Robinson et al., 2006).
The 1995 Antofagasta earthquake (Mw = 8.1) ruptured a 180 kmlong segment just south of the main 1877 gap. Geodetic slip
inversions give a consistent picture of the coseismic slip distribution during the earthquake with a single asperity slipping 5–10 m
(Chlieh et al., 2004; Fig. 2B). The earthquake rupture area is limited by the southern part of the Iquique Ridge in the north and the
Tal–Tal ridge in the south (Fig. 2B). Thus, these features appear to
act as barrier for rupture propagation.
The 2007 Tocopilla event (Mw = 7.7) ruptured the deeper part of
the seismogenic interface (30–50 km) and did not reach the surface,
indicating that the shallow part of the seismogenic interface (from
the trench to 30 km depth) remains unbroken with the exception of
the southern edge of the rupture (at ∼23◦ S), where coseismic slip
seems to have propagated up to ∼25 km depth (Béjar-Pizarro et al.,
2010). The fault slip model shows slip concentrated in two main
asperities with maximum values lower than 2 m (Béjar-Pizarro
et al., 2010). The rupture area of this event is roughly coincident
with southern region of the IR which may have arrested the updip
rupture propagation (Fig. 2B).
3.2. Central Chile (31–37.5◦ S)
The rupture area of the most recent megathrust earthquake; the
2010 event (Mw = 8.8), is characterized by two regions of high seismic slip with maximum slip of 10–15 m (Lay et al., 2010; Moreno
et al., 2010). Between these asperities, the rupture bridged a zone
that was creeping interseismically with consistently low coseismic
slip. The bilateral rupture propagated towards the south, where one
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Fig. 3. (A) Earthquake segmentation along the south central Chile margin is indicated by ellipses that enclose the approximate rupture areas of historic earthquakes (Beck
et al., 1998; Ruegg et al., 2009). Starts denote epicenter location. (B) Fault slip distribution of the 2010 Maule earthquake obtained by inversion of teleseismic P waves, SH
waves, and resulting R1 source–time functions after Lay et al. (2010). Slip distribution at 0.5 m contour interval of the 1985 earthquake based on the dislocation model of
Mendoza et al. (1994).
asperity failed, and bridged a previously creeping zone to the north,
where a second asperity failed (Moreno et al., 2010).
The southern limit of the rupture area is fairly coincident with
the subduction of the Mocha FZ (Fig. 3A), and hence this HOF might
have acted as seismic barrier. In the northen limit of the 2010
earthquake, the rupture apparently stops some 100 km south of
the subduction JFR according to the teleseismic fault slip model of
Lay et al. (2010) (Fig. 3B). Moreno et al. (2010) proposed that prestress was lowered in the northern periphery of the 2010 ERA by
previous earthquakes such as 1906 and 1985 events resulting in a
low-stress barrier. Nevertheless, the JFR appears to be the northern
barrier for both the 1906 and 1985 events, and further north the
southern barrier of the 1943 earthquake (Ms = 7.9) (Fig. 3). Thus,
the JFR plays a crucial role on seismic segmentation in south central Chile. In fact, at ∼72.5◦ W/32.5◦ S, magnetic anomalies reveal
the subduction of a large seamount belonging to the Juan Fernández hotspot track (Yáñez et al., 2001). The size of this seamount is
about 40 km in diameter and 4 km high, and therefore equivalent to
the size of the O’Higgins guyot which host a wide crustal root and a
magmatic underplating body beneath the Moho (Kopp et al., 2004,
see Section 4). Thus, the subduction of the JFR might considerably
affect seismicity in this segment of the Chilean margin.
3.3. The 1960 megathrust earthquake (Mw = 9.5)
Fig. 4 shows the fault slip model inverted by Moreno et al. (2009)
using the geodetic data compiled by Plafker and Savage (1970). The
northern limit of the 1960 ERA is coincident with both the Peninsula
de Arauco and the subducting Mocha FZ, while in the south fault
slip ceased a few kilometres north of the CTJ. The most constrained
region of the model is north of Isla de Chiloé, where most geodetical
data were acquired (Moreno, personal communication). Here, the
two main slip patches are bounded by fractures zones, and these
HOF’s present a reduction of fault slip or increase of coupling.
4. Seafloor morphology and buoyant zones of the oceanic
Nazca plate
Individual features on the subducting Nazca oceanic plate are
large, reaching elevations of up to 5 km above the surrounding
regional seafloor depth and widths up to 70 km (Fig. 1). The buoyancy of these features is a key factor controlling interplate seismic
coupling and depends on the local mode of isostatic compensation. Fig. 5 shows the seismic structure of the NR (Hampel et al.,
2004), easternmost portion of the JFR (Kopp et al., 2004), and
the Mocha FZ (Contreras-Reyes et al., 2008) derived from wideangle seismic transects. The NR comprises of an anomalously thick
crust (∼14 km), twice the thickness of the typical oceanic Nazca
crust, and characterized by underplating (Hampel et al., 2004)
(Fig. 5A). Both the thick crust and the underplated subcrustal material represent anomalous bodies with density similar to gabbro, but
lower than the surrounding mantle peridotite. The easternmost
portion of the JFR corresponds to the O’Higgins seamount group
(3–4 km high) which shows a moderately thick crust according to
the seismic model of Kopp et al. (2004) (Fig. 5B). However, the
major buoyancy source is owed to reduced mantle velocities and
hence densities. That is, low upper mantle velocities in the range
of 7.5 ≤ Vp ≤ 8.0 km/s are commonly attributed to underplating of
melt bodies beneath the crust which are usually centered at the volcanic edifice, and an asymmetric distribution is rather uncommon.
Fig. 4B shows that the reduced mantle velocities extend further
trenchward suggesting mantle serpentinization in the outer rise
region (Kopp et al., 2004). Thus, the serpentinized mantle also represents an extra buoyant force since the bulk density of peridotite
decreases with the degree of hydration. Regardless if low mantle
seismic velocities and densities are caused by mantle serpentinization or magmatic underplating, both processes are an important
source for buoyancy of the oceanic lithosphere.
Contreras-Reyes et al. (2008), reported both low crustal and
upper mantle velocities beneath the Mocha FZ which were interpreted in terms of an intensely fractured and hydrated upper
oceanic lithosphere. The maximum depth of mantle serpentinization is up to 6–8 km within the uppermost mantle, providing an
important buoyancy source (Fig. 5C). Thus, buoyancy due to crustal
and upper mantle hydration may contribute an important mechanism for increased coupling at the subduction interface.
Table 1 and Fig. 5 summarize the resulting anomalous normal
stress, n , associated with the buoyancy of the NR, JFR and Mocha
FZ. n ranges between 10 and 50 MPa providing a significant
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Fig. 4. (A) Rupture area of the mega-thrust 1960 earthquake. Start denotes epicenter location. (B) Fault slip model after Moreno et al. (2009) and main oceanic bathymetric
features identified. Note the strong spatial correlation between HOF’s and low slip values, which suggests that HOF’s enhance strong coupling and hence reduced fault slip.
increase of coupling at the plate interface compared to the apparent stress drop of about 30 MPa reported as an average for Chilean
interplate earthquakes (Leyton et al., 2009). Our n estimates are
based only on the buoyancy of HOF’s and represent a minimum of
the actual anomalous coupling. Actually, n is likely to be higher
due to the additional normal stress associated with the accommodation of a subducting high relief topographic feature (topographic
effect). Scholz and Small (1997) roughly estimated that a typical
seamount 4 km high and 60 km wide results in an extra normal
stress of 100 MPa. Thus, the resulting n is expected to be higher
than estimates shown in Table 1 and Fig. 5 for large seamounts such
as the NR and JFR.
5. Discussion
5.1. Role of subducting HOF’s on seismic coupling
The spatial correlation between ERA locations and incoming
HOF’s indicates that seamounts and fracture zones play a key role
in the dynamics of the earthquake rupture along the Chile–Peru
subduction zone. For instance, part of the southern Peru seismic
gap and the entire northern Chile seismic gap is coincident with
the subduction of the IR, which likely controls at least in part the
degree of coupling in the seismogenic zone. The IR is composed by
a number of seamounts (1.0–2.0 km high) and it is surrounded by
a smooth and broad region of shallow seafloor (swell). This swell
is 200–300 km wide and more than 500 m in elevation (Fig. 2). The
gray region representing the IR’s swell shown in Fig. 2B is based
on the shallow topography of this HOF inferred from the bathymetric data. The extension and distribution of the swell is not clear
due to presence of the prominent outer rise topography. McNutt
and Bonneville (2000) proposed that the swell is mainly a consequence of the buoyancy of magmatic underplating beneath the
Moho or crustal thickening, as imaged by seismic refraction experiment. Unfortunately, the deep structure of the IR has not yet been
imaged by seismic refraction data and a high resolution image of the
IR deep structure is unknown. Nevertheless, the 3D density model
of Tassara et al. (2006) shows that the IR hosts an anomalously thick
crust ∼15 km thick providing an important source of buoyancy. A
smooth and uplifted topography (like the IR swell) might provide a
Table 1
The anomalous normal stress ␴n is calculated as n = gH + gH, where the n (buoyancy force) depends on the thickness of the corresponding anomalous
crustal thickness (Hc ), thickness of the underplated magmatic material beneath the crust and/or thickness of the serpentinized mantle (Hm ). n also depends on the
mantle–crust density contrast (mc = 530 kg/m3 ), “normal” mantle–serpentinized mantle density contrast (m = 230 kg/m3 ) and/or “normal” crust-altered crust density
contrast (c = 50 kg/m3 ). g = 9.81 m/s2 is the acceleration due to gravity. Hm and Hc are based on the information shown in Fig. 5.
Oceanic feature
Hc (km)
NR
JFR
Mocha FZ
7.4
2.0
0.0
mc (kg/m3 )
530
530
–
Hm (km)
m (kg/m3 )
c (kg/m3 )
n (MPa)
3.0
3.0
8.0
230
230
230
0
0
50
50
10
20
Fig. 5. (upper) Structure and interpretation of the (A) Nazca Ridge, (B) easternmost portion of the Juan Fernández Ridge, and (C) Mocha FZ based on the 2D seismic velocity model of Hampel et al. (2004), Kopp et al. (2004), and
Contreras-Reyes et al. (2008), respectively. Map locations of seismic profiles are shown in Fig. 1A. (below) Direct comparison of these HOF’s structure with typical Nazca oceanic crust (6.5 km thick). The anomalous normal stress
n (buoyancy force) depends on the thickness of the corresponding anomalous crustal thickness (Hc ) and on thickness of the underplated magmatic material beneath the crust and/or thickness of the serpentinized mantle
(Hm ). n also depends on the mantle–crust density contrast (mc = 530 kg/m3 ), “normal” mantle–serpentinized mantle density contrast (m = 230 kg/m3 ) and “normal” crust-altered crust density contrast (c = 50 kg/m3 ).
See further details in Table 1. Density values are taken from the density model of Tassara et al. (2006).
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Fig. 6. Cartoons showing the subduction interface with a thin (a) and thick (b) subduction channel cover. A thin subduction channel is typical for the northern Chile and
Peruvian margins (see Contreras-Reyes et al. (2010) for a review), whereas a thick subduction channel thicker than 1.0 km correlates with the rupture area of the 1960
earthquake (Scherwath et al., 2009; Contreras-Reyes et al., 2010). A thick subduction channel is expected to smooth the plate interface contact and favor earthquake rupture
propagation. FAP denotes frontal accretionary prism.
high degree of coupling homogeneously distributed. Furthermore,
the IR subduction is highly oblique beneath South America, effectively spreading the anomalously coupled zone over a wider region.
Similarly, the obliqueness of other subducted HOF’s can extend the
anomalously coupled region along strike (e.g., Nazca FZ, Challenger
FZ, Iquique FZ and Mocha FZ). For large HOF’s such as the IR, oblique
subduction can explain the large size and recurrence interval of
great earthquakes.
The impact of subducting seamounts on the degree of coupling
has been broadly documented (e.g., Scholz and Small, 1997; Cloos
and Shreve, 1996) in contrast to the effect of fracture zones on seismic coupling which has been given less attention. Even though the
importance of oceanic fracture zones on seismic segmentation has
been proposed by some authors (e.g., Ruff and Miller, 1994), the
mechanism responsable for such segmentation has not been disscused much. We proposed here that the deep serpentinized bodies
within the uppermost mantle hosted by oceanic FZs affect dramatically the buoyancy and hence the degree of coupling (Fig. 5).
Consequently, their impact on seismicity is of similar importance
to hotspot seamounts. For instance, the subduction of the Nazca FZ,
Iquique FZ, Challenger FZ and Mocha FZ appears to strongly correlate with the northern and southern limits of several ERA’s (Fig. 1A).
In most cases these FZ’s coincide with the ERA’s borders and hence
act as seismic barriers.
5.2. The interplay between HOF and subduction channel
thickness and its impact on rupture propagation
An additional and fundamental element affecting the coupling
degree along the subduction interface is the subduction channel
thickness (e.g., Ruff, 1989). For example, a rough subducting HOF
covered by a thin subduction channel translates into highly heterogeneous strength distribution, and between regions of high relief
would behave aseismic due to its weak coupling, which will result
in small earthquake generation (Ruff, 1989). This process is illustrated in Fig. 6A, which shows the typical structure of the northern
erosional Chile margin characterized by a thin subduction channel. A thin subduction channel would result in a potential asperity
and large earthquake nucleation if the dynamic stress is larger than
the yield shear stress, otherwise it will behave as a seismic barrier.
On the other hand, a thick subduction channel (Fig. 6B) will allow
a smoother strength distribution along the plate interface resulting into highly homogeneous strength distribution (Ruff, 1989).
It can be expected that a thick subduction channel would reduce
coupling above topographic heterogeneities and it would increase
plate coupling between high topographies (Fig. 6B). Thus, these two
processes together will favor a homogeneous plate contact and consequently earthquake rupture propagation and generation of giant
earthquakes.
Recently, seismic studies have reported large variations of
the subduction channel thickness along the Chile–Peru margin
(Scherwath et al., 2009; Contreras-Reyes et al., 2010). North of the
JFR, the erosional northern Chile and Peruvian margin are characterized by a thin subduction channel, typically thinner than 300 m
(Reichert et al., 2002). On the other hand, the subduction channel
south of the JFR and north of the Mocha FZ is relatively thin (≤1.0 km
thick) compared to the subduction channel south of the Mocha FZ
and north of the CTJ (>1.5 km) (Contreras-Reyes et al., 2010). Thus,
we would expect that south of the Mocha FZ, the thick subduction
channel has a strong influence on earthquake rupture propagation.
In contrast, north of the Mocha FZ where the subduction channel
becomes thinner we would expect that the HOF’s mainly control
the size of the earthquake. Although the limits of several earthquakes do not perfectly correlate with HOF subduction, this may be
explained by local anomalous regions of thick subduction channels
that would tend to homogenize the subduction interface facilitating
rupture propagation.
The Mocha FZ is a remarkable oceanic feature controlling the
size of both the 1960 and 2010 earthquakes. In fact, the highly coupled region where the Mocha FZ subducts appears to have arrested
the southward propagation of the recent Maule 2010 earthquake
(Fig. 3). On the other hand, the Mocha FZ stopped in the north the
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rupture propagation of the megathrust 1960 event (Fig. 4). This
earthquake then propagated all the way south to just north of the
CTJ. The thick subduction channel along the rupture zone may have
provided enough smoothness for a long rupture propagation across
the fracture zones involved. The thin subduction channels north of
the Mocha FZ and near the CTJ were too thin to reduce plate coupling, and rupturing ceased before crossing the Mocha FZ and the
Chile Rise in the north and south, respectively (Scherwath et al.,
2009). In addition, the fault-slip model shows that the slip patches
(Moreno et al., 2009) are approximately bounded by the subduction of fracture zones, which suggests that these HOF’s tend to act
as barriers, but both the large amount of energy carried by the rupture front and the thick subduction channel have allowed further
rupture propagation (Fig. 4).
A giant earthquake can also be generated in a region with a thin
subduction channel, if the topography of the subducting feature is
rather smooth and wide enough as for instance an oceanic plateau.
This case could be the active seismic gap of north Chile, where the
IR subducts and no major earthquake (Mw > 8.5) has occurred since
more than 120 years. The smooth topography and deep crustal root
of the IR (Tassara et al., 2006) favors efficiently the plate contact,
seismic coupling and elastic energy accumulation during the interseismic period. When the accumulated shear stress becomes larger
than critical stress required for failure, a large earthquake may generate with a rupture area of similar dimensions to the IR.
6. Summary
We propose that two important factors affect the earthquake
rupture behavior: (1) the buoyancy of the subducting HOF, and (2)
the effect of the interplay between HOF and subduction channel
thickness on the degree of coupling. We show a strong relation
between subduction of oceanic features and earthquake rupture
segments along the Chile–Peru margin, elucidating how these subducting features play a key role in seismic segmentation. Within
this context, the extra increase of normal stress at the subduction
interface is strongly controlled by the buoyancy of HOF’s. These
features generate anomalously coupled regions acting as barriers
and/or asperities. This dual behavior depends mainly on the interplay between the magnitude of the yield shear stress in the vicinity
of the HOF and the energy carried by the rupture front. A thick subduction channel smooths the degree of coupling and hence reduce
the yield shear stress at the subduction interface allowing lateral
earthquake rupture propagation. Future work should focus on more
seismic imaging HOF’s and subduction channels in order to better
understand their internal structures to provide better constraints
on rupture earthquake dynamics modelling and seismic hazard
assessment.
Acknowledgments
We thank Javier Ruiz and Sergio Ruiz for fruitful discussions.
Eduardo Contreras-Reyes acknowledges the support of the Chilean
National Science Foundation (FONDECYT) project # 11090009. We
thank Marta Béjar-Pizarro and Marcos Moreno for providing the
digital coseismic slip models. We also thank M.A. Gutscher, an
anonymous reviewer, and the Editor George Helffrich for comments on the manuscript. We appreciate comments on the English
provided by Dave Rahn.
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