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Spatial-temporal complexity of
continental intraplate seismicity: insights
from geodynamic modeling and implications for seismic hazard estimation
Qingsong Li, Lunar and Planetary Institute, USRA, Houston, TX 77058
Mian Liu, Dept. of Geological Sciences, University of Missouri, Columbia, MO 65211
Seth Stein, Dept. of Earth and Planetary Sciences, Northwestern University, Evanston, IL 60208
Abstract:
Continental intraplate seismicity occurs incomplex spatiotemporal patterns. The earthquakes are
often episodic, clustered, and migrating. Hence we observe both spatial clustering in seismic
zones and scattering across large plate interiors; temporal clustering followed by long periods of
quiescence; and migration of seismicity from one seismic zone to another. To seek insight into
these complexitiers, we have explored the spatiotemporal evolution of intraplate seismicity
using a 3D visco-elasto-plastic finite element model. The model simulates tectonic loading,
crustal failure in earthquakes, and coseismic and postseismic stress evolution. With pre-specified
perturbations of crustal strength, the model predicts various spatiotemporal patterns of seismicity
at different timescales: spatial clustering (in narrow belts) and scattering (across large regions)
overhundreds of years, connected seismic belts overthousands of years, and seismicity scattering
over the whole model region over tens of thousands of years. The orientation of seismic belts
coincides with the optimal failure directions predicted by the prescribed tectonic loading. The
model results also show temporal clustering of earthquakes in seismic zones. When weak zones
are included in the model, the predicted seismicity initiates within the weak zones but then
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extends far beyond them. If a fault zone is weakened following a large earthquake, repeated
large earthquakes can occur on the same fault zone even at the absence of strong tectonic
loading. These results provide useful insights into the complex spatial-temporal patterns of
intraplate seismicity, and call for caution when assessing intraplate earthquake hazard.
Assessment of earthquake hazard based on the limited historic record may be biased toward
overestimating the risks in regions of recent large earthquakes and underestimating the risks
where seismicity has been quiescent.
Introduction
The distribution of earthquakes in space and time within continental interiors is far more
complex than on plate boundaries. Earthquakes on plate boundaries result from strain produced
by steady relative plate motions. Hence plate boundaries remain the loci of large quasi-periodic
earthquakes for long intervals until the plate boundary geometry changes. As a result, the longterm pattern of seismicity inferred from geological fault studies is consistent with that observed
in large earthquakes e.g., [Cisternas, et al., 2005; Marco, et al., 1996; Weldon, et al., 2004].
Moreover, the direction and rate of geodetically observed strain accumulation is generally
consistent with the mechanisms and recurrence of large earthquakes [Stein and Freymueller,
2002].
The situation is quite different within continental plate interiors, where earthquakes are
clustered, episodic, and migrate (Figure 1). Paleoseismic data show that intraplate earthquakes
often occur in temporal clusters on faults that remain active for some time, and then have long
quiescent periods during which seismicity migrates to other faults [Crone, et al., 2003]. Thus,
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faults without historically recorded earthquakes, such as the Meers fault in Oklahoma [Crone
and Luza, 1990], appear to have had prehistoric large events [Clark and McCue, 2003].
The earthquakes in continental intraplate regions may also occur in temporal clusters. For
example, earthquakes in North China’s Weihe-Shanxi graben are temporally clustered before
1700 (Figure 2a). Similarly, he large earthquakes (M>6) in the North China Plain are clustered
between 1900 and 2000 (Figure 2b). Moreover, the locus of seismicity seems migrate between
faults and regions, such as from the Weihe-Shanxi graben eastward to the north China Plain in
the past 200 years (Figure 1b).
The limited observations of continental intraplate seismicity show complex spatial patterns:
spatial clustering of earthquakes in seismic zones and scattering in large regions (Figure 1). For
example, historic and instrumentally recorded earthquakes in the Central and Eastern U.S.
(CEUS) appear to be concentrated in several zones, notably the New Madrid, Charleston, and St.
Lawrence seismic zones (Figure 1a). In addition, scattered seismicity occurs outside these zones
(Figure 1a). Some seismic zones, such as New Madrid, occur along the failed rift zones, but
other seismic zones do not [Schulte and Mooney, 2005]. Conversely, some prominent rifts, such
as the Mid-continent rift in central North America, have little seismicity. As a result, the relation
between continental intraplate seismicity and geologic structure remains uncertain and the
subject of active investigation, as summarized by papers in Stein and Mazzotti [2007].
MAYBE IN FIG 1a USE THE VERSION WITH RIFTS MAPPED
The complex variation of seismicity in space and time poses a major challenge for seismic
hazard estimation. It seems likely that inferring seismic hazards from historical seismicity can
overestimate the hazard where recent earthquakes have occurred and underestimate hazard
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elsewhere [e.g., Swafford and Stein, 2007; van Lanen and Mooney, 2007]. To reduce this
difficulty, researchers are developing hazard maps that also reflect geologic structure, yielding
maps with more diffuse hazards [Halchuk and Adams, 1999; Tόth, et al., 2006].
Numerous models have attempted to explain the complexities of intraplate seismicity.
Generally, these models involve temporal variations of the physical properties of faults and the
stress acting on them. Crone et al. [2003] suggest that temporal clustering, in which faults "turn
on" to generate a series of large earthquakes and then "turn off” for along time, may reflect the
evolution of pore fluid pressure in the fault zone [Sibson, 1992]. In this model, low-permeability
seals form around the fault zone as stress accumulates, raising the pore pressure until an
earthquake happens. Faults that weaken during earthquakes and heal during interseismic phases
could also contribute the complexity of seismicity [Lyakhovsky, et al., 2001]. The migration of
seismicity between faults may result from stress transfer, in which stress changes due to
earthquakes affect the location of future events [Chery, et al., 2001; Mueller, et al., 2004; Rundle,
et al., 2006; Stein, 1999].
Finite element model
We explored the spatiotemporal complexity of intraplate earthquakes with a 3D viscoelasto-plastic finite element model [Li and Liu, 2006; Li and Liu, 2007]. The model simulates
tectonic loading, crustal failuring, and the associated coseismic and postseismic stress evolution.
Tectonic loading is simulated by applying velocity boundary conditions at the model boundaries.
Crustal failure in an earthquake is simulated by abruptly decreasing strength in an element by a
given amount when stress in the element reaches the failure criterion. This process releases stress
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at the element and causes deformation in the surrounding model domain. Postseismic stress
evolution following the failure event is simulated by viscous relaxation of the lower crust and
the upper mantle. The coseismic and postseismic stress changes in the elements that failed may
increase (or decrease) stress in the neighboring regions, thus promoting (or
inhibiting)
earthquake occurrence there.
The model dimension is 2000 km by2000 km and×60 km thick (Figure 3). A 20km thick
elasto-plastic upper layer simulates the upper crust and a 40km thick visco-elastic lower layer
simulates the lower crust and upper mantle. The eastern and western boundaries are compressed
uniformly at 1 mm/yr, while the northern and southern boundaries are extended at the same rate.
The bottom boundary is free-slip. Young’s modulus is 8.75×1010 Pa and Poison’s ratio is 0.25
for both the crust and the mantle. The viscosity for the lower layer is 1×1020 Pa s. The strength of
the upper layer is described by a cohesive strength of 50 MPa and a coefficient of internal
friction of value of 0.4.
Figure 4 shows an example of how the predicted stress and earthquakes evolve. . Figure 4a
shows the stress at the beginning of a 1000-year time period, andFigure 4b shows the stress at
the end of the period. The black dots in Figure 4a show the earthquakes that occurred during the
period. The earthquakes occur within regions of high stresses. After the earthquakes, stress in
these regions drops below the crustal strength. The earthquakes also cause stress changes in the
surrounding regions, whichare much smaller than the stress drops during the earthquakes.
Model results
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effect of various factors on the predicted spatiotemporal patterns of
intraplate seismicity, we start with a simple model with horizontally homogeneous crust and
mantle, and then incrementally add additional factors.
Case o1: timescale-dependent spatiotemporal patterns of seismicity
In this case , an initial randomly pre-specified perturbation of the crustal strength is
applied to an otherwise laterally homogeneous crust and mantle (Figure 3). The resulting
seismicity shows both spatial clustering (in spots and belts) and scattering (in large regions) over
a period of 300 years (Figure 5a). Over a longer period, 3000 years, the earthquake clusters
connect to form a network of seismic zones (Figure 5b). In an even longer time, 30,000 years,
seismicity covers the entire model domain (Figure 5c). This result is expected because the model
is horizontally homogeneous. Hence seismicity is uniform over the long term, but is concentrated
in seismic zones during shorter intervals.
The model results also show temporal clustering of earthquakes in seismic zones. Figure 6
shows the predicted earthquake sequences for four 100 km2 regions. Temporal clustering of
earthquakes occurs in region B between 10,000 and 20,000 years, and in region C around 5000
and 22,000 years. The temporal clustering is caused by coseismic and postseismic stress
triggering. Without coseismic and postseismic stress interaction between regions, we would
expect periodic earthquake occurrence given that the model has constant loading and constant
stress release in earthquakes.
The orientations of the predicted seismic belts, NW-SE or NE-SW, coincide with the
optimal failure directions resulting from the boundary conditions of east-west compression and
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north-south extension. This loading produces a pure shear stress state in the model, with the
maximum shear stress oriented NW-SE and NE-SW.
Case 2: effects of weak zones.
In the second case the model does not have a prespecified perturbation of crustal strength.
Instead, it has three weak zones whose yield strength is 5 MPa lower than the rest of the model.
Two of the zones are parallel and perpendicular to the third (Figure 7a):. The initial stress is set
to zero, and the model is constantly loaded from at the boundaries.Initially , earthquakes occur
only within the weak zones. After ~80,000 years, stress in the surrounding regions begins to
reach the strength and cause crustal failure. After another 100,000 years, earthquake occurrence
enters a steady state, in which the seismicity during one time period is similar to another,
provided that the period is long enough (>60,000 years).
IS FIGURE 7 STARTING AT TIME ZERO OR 100,000 YEARS (STEADY STATE)?
The predicted seismicity ondifferent time scales is shown in Figure 7. On a short time
scale (600 years), earthquakes occur both inside and outside the weak zones (Figure 7a). The
seismicity shows some spatial clustering with no clear correlation with the weak zones. Over a
longer period (6000 years), the seismicity forms belts extending from the weak zone (Figure 7b).
This pattern remains the same over longer periods (Figure 7c). Hence on the long time scales,
seismicity is not be confined to the weak zones. Moreover, as in the previous case, the short-term
seismicity pattern does not fully represent that over longer terms/
.
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Case 3: effects of fault weakening.
Paleoearthquake studies have found evidence of repeated ruptures on individual intraplate
faults and fault zones in short periods, followed by long periods of quiescence [Crone, et al.,
2003; Tuttle, et al., 2002]. The short time between the repeated earthquakes on these intraplate
faults, such as the ~500 years found for the recent large events in the New Madrid Zone [Tuttle,
et al., 2002]) is at odds with the low rates of deformation and far field loading in plate interiors
[Calais, et al., 2005; Newman, et al., 1999]. Li et al. [2005] have shown that a large intraplate
earthquake can reduce stress in the fault zone to levels below that requires for rupture for
thousands of years. Thus, large intraplate earthquakes repeating in a short time interval would
require either local loading to build up the stress, or local weakening to permit repeated failure at
lowered stress levels.
A number of mechanisms for local loading and/or fault weakening have been proposed
[Kenner and Segall, 2000; Pollitz, et al., 2001], although there is no direct evidence for the
proposed weakening [McKenna et al., 2007].
We explored the effect of fault weakening
[Lyakhovsky, et al., 2001; Sibson, 1992]. In this case we explicitly incorporated a weak fault
zone (zone A in Figure 7a), whose strength is 5 MPa lower than the rest of the model domain .
We considered here three different scenarios of fault weakening following a large fault rupture.
These assume different histories of weakening but do not specified its physical mechanism.
IT’S HARD IN FIGS 8 & 9 TO SEE WHICH IS WHICH. I SUGGEST USING
CONSISTENT NOTATION (A,B.C), WORDING, AND LABELS
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In the first scenario, the fault zone is weakened abruptly by 0.5 MPa immediately after the
first rupture, which had a stress drop of 1 Mpa. Stress on the fault then gradually rises due to
the far-field loading(Figure 8), and reaches the new lower fault strength and causes a second
fault rupture (Figure 8) in 1800 years. Because no further fault weakening occurs, it then takes
another 3000 years for the fault zone to rupture again (Figure 9a).
In the second scenario, the strength of the fault zone weakens continuously at a rate of 1
MPa per 1000 years after the first rupture. The stress in the weak zone recovers and is then
released (Figure 8) in a series of repeated earthquakes (Figure 9b). The time interval between
successive events is hundreds of years.
In the third scenario, the strength of the fault zone gradually weakens in the first 2000 years
following the first rupture, and then remains constant . The stress evolution in the first 2000
years is the same as that in the second scenario. After the weakening stops, it takes longer time
for the stress to recover (Figure 8). Thus the predicted earthquake series is identical to that in the
second scenario during the first 2000 years. After that, the recurrence interval becomes longer
but regular (Figure 9c).
Discussion
The strength perturbation in case 1 was randomly generated and so is irregular (Figure 10).
The magnitude of the perturbation ranges from -0.5 to 0.5 MPa. This perturbation simulates the
presence of weak zones in the ancient continental crust. A model in which uniform loading was
applied to a crust of uniform strength would predict uniform failure of the whole upper crust
instead of crustal failure at different locations and times.
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Because the strength perturbation is randomly distributed, the average strength in a large
region of the model does not differ significantly from that of other regions. Consequently, the
strength perturbation has only a limited effect on the distribution of seismicity. Figure 11 shows
the ratioof the earthquakes that occured in weak regions (negative strength perturbation) to the
total number of earthquakes in the entire model. During 500-year time windows The fractions
fluctuate around 0.5. In 30,000 years, 53.3% earthquakes occurred in weak regions, while 46.7%
occurred in strong regions (positive strength perturbation).
We might expect a correlation between the orientation of tectonic stress and that of seismic
zone. . In the simplest case, the maximum compressive stress would bisect conjugate trends of
seismicity. s However, this does not seem the case in the continents does not appear The seismic
belts in North China form conjugate NNE and SEE pairs (Figure 1b), whereas the direction of
maximum compressive tectonic stress is around N55E [Yuan, et al., 1999]. The seismic belts in
CEUS are oriented NE
CAN YOU SEE AN SE TREND?,
(Figure 1a), close to the average maximum compressive direction of tectonic stress of
around N60E [Zoback and Zoback, 1989],. Thus the orientation of seismic zones may instead
be controlled by ancient fault zones, which formed in a different stress field from the present
ones [Johnston and Schweig, 1996].
In our models both the tectonic stress and the orientation of preexisting structures may
affect the orientations of seismicity. In case 1, the orientation of seismicity is controlled by
tectonic stress, because the strength perturbation is small and random. In case 2, the orientations
of the weak zones coincide to the optimal failure directions of the stress field. Thus the
seismicity belts extending from the weak zones are oriented in the same directions as the weak
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zone and optimal failure directions. I AGREE: THIS IS BEST LEFT OUT UNLESS MORE
ANALYSIS IS DONE
The regional temporal clustering in the model (Figure 6) is caused by coseismic and
postseismic stress triggering. Such stress triggering may have caused a series of large
earthquakes to occur on nearby faults within a short period, as in the 1811-1812 NMSZ
earthquakes [Mueller et al., 2004] and the 2005-2007 Sumatra earthquakes [McCloskey et al.,
2005]. However, our model cases 1 and 2, in which constant fault properties and tectonic
loading are assumed, can not explain repeated earthquakes on individual faults in short periods,
such as the approximately 500 year recurrence at New Madrid,. Such case3, we have repeated
earthquakes on an individual fault arise in case 3 due to the assumed fault weakening (Figure
11).
Although there is no direct evidence for
fault weakening, several possible causes have been proposed. Sibson [1992] suggested fault
weakening through change of water pore pressure on faults. Lyakhovsky et al. [2001] suggested
changes of fault rheology due to a balance between weakening due to earthquakes and
interseismic healing. Fault weakening and healing may help to explain the paleoseismic data
interpreted as showing that large New Madrid earthquakes occur a few hundred years apart in
clusters, and then do not occur again for a long time [Holbrook, et al., 2006; Schweig and Ellis,
1994] Thus the geodetic observations in the New Madrid seismic zone showing little or none of
the interseismic motion expected before a future large earthquake suggest that the recent cluster
of large earthquakes may be coming to an end [Calais, et al., 2005; Newman, et al., 1999],
McKenna et al. 2007].
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Conclusions
We have explored the spatiotemporal patterns of intraplate earthquakes with a 3D finite
element model. The model includes far field boundary loading, crustal failure, coseismic and
postseismic stress evolution, and fault weakening. Our main results include the following:
1) Even with a horizontally homogeneous crust and mantle, the model predicts complex
spatio-temporal seismicity patterns that vary with timescales. Over hundreds of years, the
seismicity may appear both clustered in small belts and scattered over a larger region. Over
thousands of years, the seismic clusters (belts) connect to form a network of seismic belts. Over
tens of thousands of years, scattered seismicity covers the whole model region.
2) Coseismic and posterseismic triggering can cause temporal clustering of seismicity at
individual small regions. If weak zones are incorporated in the model, the seismicity initiatesin,
but is not confined to, these weak zones.
3) Fault weakening can lead to repeated earthquakes on intraplate fault zones. The predicted
temporal patterns of these repeated earthquakes varies with the weakening history .
Thus this simple model can replicate some of the spatio-temporal complexity of migrating,
episodic, and clustered intraplate earthquakes. Thus the model is generally consistent with the
observations that the seismicity pattern observed from short-term seismic records may not reflect
the long term patterns intraplate seismicity. Hence assessment of earthquake hazard may be
biased toward overestimating the risks in regions of recent large earthquakes and
underestimating the risks where seismicity has been quiescent.
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References
Calais, E., et al. (2005), Tectonic strain in plate interiors?, Nature, 438, E9-E10.
Chery, J., et al. (2001), A physical basis for time clustering of large earthquakes, Bull. Seismol.
Soc. Am., 91, 1685-1693.
Cisternas, M., et al. (2005), Predecessors of the giant 1960 Chile earthquake, Nature, 437, 404407.
Clark, D., and K. McCue (2003), Australian paleoseismology: towards a better basis for seismic
hazard estimation, Annals of Geophysics, 46, 1087-1105.
Crone, A. J., et al. (2003), Paleoseismicity of two historically quiescent faults in Australia:
Implications for fault behavior in stable continental regions, Bull. Seismol. Soc. Am., 93, 19131934.
Crone, A. J., and K. V. Luza (1990), Style and timing of Holocene surface faulting on the Meers
fault, southwestern Oklahoma, GSA Bull., 102, 1-17.
Halchuk, S., and J. Adams (1999), Crossing the border: assessing the differences between new
Canadian and American seismic hazard maps, Proceedings Eighth Canadian Conference on
Earthquake Engineering, 77-82.
Holbrook, J., et al. (2006), Stratigraphic evidence for millennial-scale temporal clustering of
earthquakes on a continental-interior fault: Holocene Mississippi River floodplain deposits,
New Madrid seismic zone, USA, Tectonophysics, 420, 431-454.
Johnston, A. C., and E. S. Schweig (1996), The enigma of the New Madrid earthquakes of 18111812, Annu. Rev. Earth Planet. Sci., 24, 339-384.
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Kenner, S. J., and P. Segall (2000), A mechanical model for intraplate earthquakes; application
to the New Madrid seismic zone, Science, 289, 2329-2332.
Li, Q., and M. Liu (2006), Geometrical impact of the San Andreas Fault on stress and seismicity
in California, Geophys. Res. Lett., 33, L08302, doi:08310.01029/02005GL025661.
Li, Q., and M. Liu (2007), Initiation of the San Jacinto Fault and its interaction with the San
Andreas Fault: Insights from geodynamic modeling, Pure Appl. Geophys., 164, 1937-1945.
Li, Q., et al. (2005), Stress Evolution Following the 1811-1812 Large Earthquakes in the New
Madrid Seismic Zone, Geophys. Res. Lett, 32, doi:10.1029/2004GL022133.
Lyakhovsky, V., et al. (2001), Earthquake cycle, fault zones, and seismicity patterns in a
rheologically layered lithosphere, J. Geophys. Res., 106, 4103-4120.
Marco, S., et al. (1996), Long term earthquake clustering: a 50,000 year paleoseismic record in
the Dead Sea Graben, J. Geopghys. Res., 101, 6179-6192.
Mueller, K., et al. (2004), Analysing the 1811-1812 New Madrid earthquakes with recent
instrumentally recorded aftershocks, Nature, 429, 284-288.
Newman, A., et al. (1999), Slow deformation and lower seismic hazard at the New Madrid
Seismic Zone, Science, 284, 619-621.
Pollitz, F. F., et al. (2001), Sinking mafic body in a reactivated lower crust; a mechanism for
stress concentration at the New Madrid seismic zone, Bull. Seismol. Soc. Am., 91, 1882-1897.
Rundle, J. B., et al. (2006), Stress transfer in earthquakes, hazard estimation and ensemble
forecasting: inferences from numerical simulations, Tectonophysics, 413, 109-125.
Schulte, S. M., and W. D. Mooney (2005), An updated global earthquake catalogue for stable
continental regions: reassessing the correlation with ancient rifts, Geophys. J. Int., 161, 707-721.
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Schweig, E. S., and M. A. Ellis (1994), Reconciling short recurrence intervals with minor
deformation in the New Madrid seismic zone, Science, 264, 1308-1311.
Sibson, R. H. (1992), Implications of fault-valve behaviour for rupture nucleation and recurrence,
Tectonophysics, 211, 283-293.
Stein, R. S. (1999), The role of stress transfer in earthquake occurrence, Nature, 402, 605-609.
Stein, S., and J. T. Freymueller (Eds.) (2002), Plate boundary zones: Concepts and approaches,
425 pp., AGU, Washington, D.C.
Tuttle, M. P., et al. (2002), The earthquake potential of the New Madrid seismic zone, Bull.
Seismol. Soc. Am., 92, 2080-2089.
Tόth, L., et al. (2006), Seismic hazard in the Pannonian region, in The Adria Microplate: GPS
Geodesy, Tectonics and Hazards, edited by N. Pinter, et al., pp. 369-384, Springer Netherlands.
Weldon, R., et al. (2004), Wrightwood and the earthquake cycle: What a long recurrence record
tells us about how faults work, GSA Today, 14, 4-10.
Yuan, J., et al. (1999), Inversion analysis of present tectonic stress field in North China region
based on GPS data, Geological Science and Technology Information, 18, 99-103.
Zoback, M. L., and M. D. Zoback (1989), Tectonic stress field of the continental United States,
in Geophysical framework of the continental United States, edited by L. C. Pakiser and W. D.
Mooney, pp. 523-539, Geological Society of America Memoir 172, Boulder, Colorado.
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REFS TO CONSIDER ADDING
Stein, S., and S. Mazzotti, Editors, Continental intraplate earthquakes: science, hazard,
and policy issues, Special Paper 425, Geological Society of America, Boulder CO, 2007.
Swafford, L. and S. Stein, Limitations of the short earthquake record
for seismicity and seismic hazard studies, in Continental intraplate earthquakes: science, hazard,
and policy issues, Special Paper 425, S. Stein, and S. Mazzotti, Editors,
Geological Society of America, Boulder CO, 49-58, 2007.
McKenna, J., S. Stein, and C. Stein, Is the New Madrid seismic Zone hotter
and weaker than its surroundings? In Continental intraplate earthquakes: science, hazard,
and policy issues, Special Paper 425, S. Stein, and S. Mazzotti, Editors,
Geological Society of America, Boulder CO, 167-175, 2007.
Van Lanen, X. and W.D. Mooney, Integrated geologic and geophysical studies of North
American continental intraplate seismicity. in Continental intraplate earthquakes: science, hazard,
and policy issues, Special Paper 425, S. Stein, and S. Mazzotti, Editors,
Geological Society of America, Boulder CO, 101-112, 2007.
McCloskey, J., S. S. Nalbant, and S. Steacy, Indonesian earthquake: Earthquake risk from coseismic stress, Nature, 434, 291, 2005.
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Figure 1. Intracontinental earthquakes in the Central and Eastern U.S. (CEUS) (a), North China
(b), and Australia (c). Data for the CEUS and Australia are from the NEIC catalog. Earthquake
data of North China are from China Seismological Bureau. Green dots in CEUS and Australia
represent earthquakes before 1973, while red dots represent earthquakes after 1973. Green dots
in North China represent earthquakes before 1900, while red dots represent earthquakes after
1900.
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Figure 2. Historic records of intraplate earthquakes (M>6) in Weihe-Shanxi Graben
“Graben”in text
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(a) and North China Plain (b).
Figure 3. Finite element model for Scenario 1. The model is loaded by compression at the east
and west boundaries and extension at the north and south boundaries. The blue layer represents
the brittle upper crust, while the pink layer represents the ductile lower crust and upper mantle. A,
B, C, D are four small regions used for showing seismicity clustering in figure 6.
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Figure 4. Evolution of stress and seismicity over 1000 years in the model. (a) Initial stress level
and the locations of the resulting earthquakes (dots). (b) Stress level after 1000 years. The color
scale indicates how close the stress is from the crustal strength.
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Figure 5. Seismicity in Case1 that occurred over time scales of 300 years (a), 3000 years (b),
and 30,000 years (c). The colors of dots represent the time within each of the three intervals.
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Figure 6. Temporal distribution of earthquakes s in regions A, B, C, and D (Figure 3).
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Figure 7. Seismicity in Case 2 that occurred over time scales of 600 years (a), 6000 years (b),
and 60,000 years (c). The colors of dots represent the time within each of the three intervals.
The grey belts in (a) are weak zones.
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Figure 8. Stress evolution in the weak zone by incorporating three different weakening scenarios.
SUGGEST LABELING ABC IN KEY
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Figure 9. Predicted crustal rupture sequences for abrupt weakening (a), continuousweakening (b),
and weakening that stops after 2000 years (c).
CHANGE FIGURE LABEL FROM “GRADUAL” TO “CONTINUOUS”
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Figure 10. Initial random crustal strength perturbation used in case 1.
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Figure 11. The fraction of earthquakes that occurred in elements of negative strength
perturbation in successive 500-year time windows .