Bulletin of the Seismological Society of America, Vol. 104, No. 3, pp. 1100–1110, June 2014, doi: 10.1785/0120130222
Ⓔ
Shear-Wave Splitting Study of Crustal Anisotropy
in the New Madrid Seismic Zone
by P. Martin, P. Arroucau, and G. Vlahovic
Abstract
This study investigates crustal anisotropy in the New Madrid Seismic
Zone (NMSZ) by analyzing shear-wave splitting measurements from local earthquake
data. In addition to the waveforms provided by the Center for Earthquake Research and
Information (CERI) for over 3000 events, seismograms recorded by the portable array
for numerical data acquisition (PANDA) network were obtained for over 800 events.
Data reduction led to a final data set of 168 and 43 events from the CERI and PANDA
data, respectively. One-hundred and eighty-six pairs of measurements were produced
from the CERI data set by means of the automated shear-wave splitting measurement
program mfast and 49 from the PANDA data set. Two dominant directions, respectively
striking northeast–southwest and west-northwest–east-southeast, are identified and interpreted to be due to stress-aligned microcracks. The northeast–southwest polarization
direction is consistent with the maximum horizontal stress orientation of the region and
has previously been observed in the NMSZ, whereas the west-northwest–east-southeast
polarization direction has not. Path-normalized time delays range from 1 to 33 ms=km
for the CERI network data and 2 to 31 ms=km for the PANDA data. These results produce a range of estimated differential shear-wave anisotropy between 1% and 8%.
These values are higher than those previously determined in the region. The majority
of large path-normalized time delays (> 20 ms=km) are located along the Reelfoot fault
segment. These high values are believed to be indicative of high crack densities and high
pore fluid pressures, which agrees with previous results from local earthquake tomography and microseismic swarm analysis.
Online Material: Tables of stations, events, and associated shear-wave splitting
measurements.
Introduction
The New Madrid Seismic Zone (NMSZ) is an active
intraplate seismic zone centrally located in the United States,
spanning portions of western Tennessee, northeastern Arkansas, and southeastern Missouri (Fig. 1a). With over 5000
recorded earthquakes since 1974 (Center for Earthquake
Research and Information [CERI], New Madrid catalog), the
NMSZ has the highest level of seismicity in the United States
east of the Rocky Mountains (Hamilton and Johnston, 1990).
Although small magnitude (M < 4) seismic events are currently dominant in the region, the NMSZ is also the location
in which three of the largest known earthquakes that took
place in the contiguous United States (magnitude > 7), occurred during the winter of 1811–1812 (Johnston, 1996).
There are four major segments of seismicity in the
NMSZ (Fig. 1a). The two segments that trend northeast to
north-northeast appear as almost vertical, dextral strike-slip
faults (Vlahovic et al., 2000). The longest of these two
segments outlines the Axial fault (AF), whereas the shorter
segment in the north outlines portions of the Western Rift
Margin (WM). These two segments are connected via the
60 km central segment trending northwest, where the seismic
activity is located along planes dipping between 30° and
50° SW (Himes et al., 1998; Vlahovic et al., 2000). This
central segment is associated with the Reelfoot fault (RF),
in which the largest concentration of the seismic activity occurs in the region (Himes et al., 1998). The fourth segment of
seismicity trends west-northwest and is loosely spatially
associated with the Grand River Tectonic Zone (GRTZ).
Although there exists some seismicity near the eastern Rift
Margin faults (EM), there is not an appreciable amount for it
to be considered a dominant source.
The large-magnitude earthquakes that occurred in 1811–
1812 caused extensive damage in the region, and if similar
magnitude events were to occur today, the local population
and infrastructure would be at risk. The possibility of such
a catastrophic scenario has made the NMSZ the focus of many
1100
Shear-Wave Splitting Study of Crustal Anisotropy in the New Madrid Seismic Zone
1101
Figure 1.
(a) Map of faults and seismicity in the New Madrid Seismic Zone (NMSZ). Earthquake epicenters (white circles) in the NMSZ
from 1974 to present (provided by Center for Earthquake Research and Information [CERI] New Madrid catalog), with the large events of
1811–1812 shown as solid black circles (location from National Earthquake Information Center [NEIC] catalog of significant historical
events). Regional faults shown from Csontos and Van Arsdale (2008), including the Reelfoot fault (RF), Axial fault (AF), western Rift
Margin fault (WM), and the Grand River tectonic zone (GRTZ). Other faults include Osceola fault zone (OFZ), Bolivar-Mansfield tectonic
zone (BMTZ), central Missouri tectonic zone (CMTZ), and the eastern Rift Margin faults (EM). Rift margin faults (EM and WM) outline the
Reelfoot rift. (b) Simplified movements (thin arrows) along faults in NMSZ derived from Johnston (1996). Large arrows indicate direction of
maximum horizontal compressional stress within the NMSZ (Zoback and Zoback, 1991; Hurd and Zoback, 2012).
studies investigating possible mechanisms for seismicity in
the region. Some proposed mechanisms favor activation of
faults in the NMSZ by regional stresses, specifically the North
American stress field (Zoback and Zoback, 1991; Frankel
et al., 2012; Hurd and Zoback, 2012) (see Fig. 1b); faults that
are favorably oriented with respect to the northeast–southwest
regional stress field (Zoback and Zoback, 1980; Heidbach
et al., 2008) are activated by the buildup of stress over time,
without the need for local anomalous behavior. Recent interpretations of Global Positioning System (GPS) measurements
in the region (Frankel et al., 2012) are in agreement with this
scenario, in which the motion between GPS stations that
crossed the Reelfoot fault was found to be on the order of
0:37 mm=year. Motions of this magnitude agree with an
∼4 mm=year interseismic slip at 12–20 km depth, which if
constant through time could produce the necessary slip for
M 7.3 earthquake along the shallow Reelfoot fault with an
occurrence rate of 500 years. This would produce the necessary occurrence intervals of larger earthquakes in the NMSZ of
500 (300) years, as proposed by Tuttle et al. (2002). Other
interpretations of GPS measurements in the NMSZ only give
surface velocities along faults a magnitude of 0:2 mm=year
(Calais and Stein, 2009). If steady-state stress accumulation
was the only driving force of horizontal movement along
faults in the NMSZ, velocities of 0:2 mm=year would require
a minimum of 10,000 years to repeat a magnitude 7 earthquake and a minimum of 100,000 years to repeat a magnitude
8 (Calais and Stein, 2009). These slip rates would therefore
not satisfy the recurrence intervals of 500 (300) years that
were proposed by Tuttle et al. (2002). In this scenario, mech-
anisms for activation of faults in the NMSZ must include local
stress concentrations or anomalous behavior of the lithosphere. Such mechanisms have been proposed, including
flexure uplift leading to the unclamping of faults due to sediment removal (Calais et al., 2010) or creep at depth promoted
by the abundance of quartz, fluid overpressure conditions, and
the associated shear strain loading (Powell et al., 2010). This
study seeks to further investigate any indications of local
anomalous behavior in stress and fluid pressure within the
NMSZ, by studying the phenomenon known as shear-wave
splitting through the analysis of local earthquake data.
Anisotropy and Shear-Wave Splitting
Anisotropy is the term that characterizes directional
dependence within a given medium; this includes seismic
anisotropy, which is manifested by the dependence of velocity on propagation direction. Although there are various
sources of seismic anisotropy, the common cause in the
upper crust results from the opening and extension of fluidfilled cracks known as extensive-dilatancy anisotropy (EDA)
cracks (Crampin and Atkinson, 1985). When microcracks
are the cause of anisotropy, the fast component of the split
shear wave is polarized parallel to the preferential orientation
of the cracks, which in turn lay parallel to the direction of
maximum stress. This results in a measured fast polarization
direction (ϕ) that is therefore directly related to the direction
of maximum stress. The measurement of time delay of slow
shear wave (δt) in the case of microcracks is directly related
to the density, ε, and geometric properties of EDA cracks
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P. Martin, P. Arroucau, and G. Vlahovic
through the relationship ε Na3 =ν, with N representing
the number of cracks and a as the crack radius, within a volume, ν (Crampin, 1994). The information gathered through
shear-wave splitting measurements can then be used to estimate the percentage of differential shear-wave velocity
anisotropy (SWVA), defined by Crampin (1989) as
min
V max
S1 − V S2
;
max
V S1 x100
1
in which V S1 and V S2 are the velocities of the fast and slow
components of the split shear wave, respectively. For most
intact rocks, regardless of rock type, SWVA ranges between
1.5% and 4.5%, which is equivalent to crack densities on the
order of 0.015–0.045 (roughly 0:001 × SWVA) (Crampin, 1994).
By measuring shear-wave splitting in the upper crust of
the NMSZ, properties of stress on both the local and regional
scale can be resolved. The dominant fast polarization measurements provide information about the orientation of the
microcracks in the crust, which gives insight into the direction of maximum compressional stress in the area. Possible
deviations from expected alignment of fast polarization
directions and regional stress can be interpreted as due to
local stress perturbation or more complex anisotropic fabric.
Previous study of shear-wave splitting in the NMSZ by Rowlands et al. (1993) showed that the dominant stress direction
was parallel to the regional maximum compressional stress.
The authors also concluded that there was no evidence for
anomalous distribution of microcracks within the NMSZ and
therefore no anomalous local stress environment. Although
the results from Rowlands et al. (1993) support mechanisms
that favor regional stress as an activation source for seismicity
in the NMSZ, the data that the authors analyzed only spanned 3
years (October 1989–June 1991), a time period that may have
been too short to capture all the complexity of the stress field
in the NMSZ. This study attempts to provide a more complete
analysis of shear-wave splitting in the NMSZ, by combining
both the data set used in Rowlands et al. (1993) and a data set
spanning the time period of 2003–2011, with events being
reported by the CERI, Saint Louis University Earthquake
Center, and contributions from Southeastern U.S. Seismic
Networks (SEUSSN) bulletins (see Data and Resources).
Methods
Shear-wave splitting measurements in this study were
made using the automated program mfast developed by
Savage et al. (2010). The program will not be discussed in
detail here, but readers are directed to the works of Silver and
Chan (1991), Teanby et al. (2004), and Savage et al. (2010).
The mfast program relies on the grid-search algorithm from
Silver and Chan (1991) to make the initial measurements of
(ϕ, δt). The grid search is used to find an inverse operator
(which is a function of ϕ and δt) that removes the effects of
splitting from the given seismic signal, which can be based
on the linearity of the corrected signals particle motion
(Silver and Chan, 1991). In order to quantify the linearity,
this automated method calculates the covariance matrix between the orthogonal components of the corrected signal.
The presence of anisotropy in a signal is represented by the
eigenvalues of this matrix. When no anisotropy is present,
there should be only one eigenvalue that is non-zero. In the
real world, however, where there is always noise, there will
be a second non-zero eigenvalue. Therefore, the lower the
value of this second eigenvalue, the more the effects of shearwave splitting have been removed from the signal. Thus, the
grid search is used to search all possible values of ϕ and δt in
a given measurement window, looking for the solution that
gives the lowest second eigenvalue. This is done for many
measurement windows, and the ϕ and δt measurements that
gave the lowest second eigenvalue are stored as the solution
for each individual window. The solutions from the grid
search are then used in conjunction with the cluster analysis
developed in Teanby et al. (2004) to determine the final
values and apply grading criteria. Any events that are not
rejected by the grading criteria are checked for null measurements. The event is considered null if the difference between
incoming and fast polarization falls below 20° or above 70°.
It is then assumed that either there is no anisotropy present or
that initial shear wave is polarized along the fast polarization
direction. Similarly, anomalously high time delay measurements can imply cycle skipping or noisy data (Evans et al.,
2006). Therefore any time delay measurements, that is 0.8
times a given maximum δt measurement are also considered
null, in which the max δt measurement in this study is 1.0 s
as suggested by Savage et al. (2010). An example of a completed measurement can be seen in Figure 2.
Two separate data sets were used in this study. The
larger data set of 3603 earthquakes spans the time period of
2003–2011, with events being reported by the CERI, Saint
Louis University Earthquake Center, and contributions from
SEUSSN bulletins. The second data set of 899 earthquakes
was collected by the portable array for numerical data acquisition (PANDA) network from October 1989 to June 1991
(Chiu et al., 1992), which was previously analyzed in the
study by Rowlands et al. (1993). For an event to provide an
accurate shear-wave splitting measurement it must fall within
the shear-wave window of at least one seismic station, in
which the shear-wave window is defined by all incidence angles less than a critical angle for a given event. This requirement results from distortion that occurs when a shear wave
arrives at the free-surface interface, such as the conversion of
S to P waves resultant from reflection or the creation of a P
wave that is inhomogeneous (Booth and Crampin, 1985).
Therefore, the ability to clearly read a shear wave on a
seismogram recorded at the surface is directly related to the
incident angle of the wave (Booth and Crampin, 1985). Uncertainty in velocity models as well as earthquake locations
have led to the use of the apparent incident angle (ia ) as
described by equation (2) (Rowlands et al., 1993).
Shear-Wave Splitting Study of Crustal Anisotropy in the New Madrid Seismic Zone
1103
Figure 2. Example of output from mfast program obtained in this study. Visual and numeric output of the mfast program: (a) seismic
event shown on all three components of the station DLAR, the event window shown in gray. (b) Horizontal components of seismogram
before (top two traces) and after correction (bottom two traces). The first two dotted lines represent the times in which measurement windows
can start, and the last two dotted lines represent the times in which measurement windows can end. The window used for this measurement is
shown in gray. (c) Plot of measurement results from 80 measurement windows, with the best measurement denoted by the X (note that it is
located within a stable region). (d) Plot of clusters from the cluster analysis, the X denotes the best cluster. (e) Waveforms before and after
correction for δt in both time domain and particle motion plot. Note the similar waveforms in the time domain and the linearity of the
corrected particle motion plot. (f) Contour plot of the grid search for all measurements, through the final measurement window. The minimum
value is denoted by the X. (g) Information about the station, event, and measurement window selection. (h) Results including the final
measurements for ϕ, which is given here in degrees from north, and δt. The color version of this figure is available only in the electronic
edition.
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P. Martin, P. Arroucau, and G. Vlahovic
Figure 3. Locations of all 185 seismic events (gray dots) used
in this study from CERI data set. Only seismic stations that had at
least one event fall within their shear-wave window were used in
this study shown here by black triangles.
ia tan−1
e
;
d
Figure 4.
Locations of all 49 seismic events (gray dots) used in
this study from the portable array for numerical data acquisition
(PANDA) data set. Only seismic stations that had at least one event
fall within their shear-wave window were used in this study shown
here by black triangles.
2
in which e is the distance to the epicenter, and d is the earthquake depth.
In this study, events from either data set must have their
apparent incident angle ia ≤ 35° at least at one station as suggested by Rowlands et al. (1993). Events for use from the
CERI data set were subject to further restrictions, such that
the event must have a standard horizontal error ≤ 2:5 km and
a standard error on depth ≤ 5:0 km; and each event must have
a magnitude of 2 or greater. Events for use from the PANDA
data set required the same location accuracy; however, most
events had a very low magnitude (M < 0:1), so no magnitude
threshold was applied. Because of the large difference in
magnitudes and to preserve the integrity of both data sets,
they were analyzed separately. There were 1160 events from
the CERI data set and 504 events from the PANDA data set
that met the quality criteria discussed above. There were 681
events from the CERI data and 329 from the PANDA data set,
eliminated either due to a low signal-to-noise ratio (SNR) or
an inability to distinguish the S-phase pick accurately. Thus,
479 pairs of (ϕ, δt) measurements were made from the remaining CERI events, 288 of which were defined as nulls
in the mfast program. Similarly, 175 measurements were
made with the remaining PANDA events, in which 111 of
these measurements were defined as nulls. Data reductions
continued with an additional loss of 30 events from the CERI
and 15 from PANDA due to the quality requirement that only
measurements with quality B (mfast criteria: SNR > 3, standard deviation of ϕ < 25) or greater be used. The final CERI
data set contained a total of 186 pairs of measurements, in
which the events used ranged in magnitude from 2.0 to 4.6
and depths from 3.8 to 23.7 km. The final PANDA data set
contained a total of 49 pairs of measurements, with magnitudes of 0.0 and depths ranging from 3.2 to 15.3 km. The
locations of these events are plotted in Figures 3 and 4, along
with the stations that had at least one event fall within their
shear-wave window. Ⓔ Tables S1 and S2, which include station, event, and measurement information for both the CERI
and PANDA data sets, are available in the electronic supplement to this article.
Results
The results from the shear-wave splitting analysis of the
CERI data set indicate two dominant fast polarization directions: northeast–southwest as well as west-northwest–eastsoutheast. These polarization directions can be seen plotted
in Figure 5a–h. The northeast–southwest fast polarization
direction has been identified before in the shear-wave splitting study of Rowlands et al. (1993). They attributed this
polarization to vertical EDA cracks aligned in the direction
of the proposed regional maximum horizontal stress (Zoback
and Zoback, 1991; Hurd and Zoback, 2012).
The northeast–southwest polarization direction is seen
strongly at stations MARM, WALK, and WYBT (Fig. 5a–h).
The west-northwest–east-southeast polarization direction is
seen strongly at stations CATM, NMDM, and LEPT
(Fig. 5a–h). The west-northwest–east-southeast fast polarization direction is found throughout the NMSZ (Fig. 5h) and has
not been identified by any previous shear-wave splitting studies in the region.
Shear-Wave Splitting Study of Crustal Anisotropy in the New Madrid Seismic Zone
1105
Figure 5. (a) Rectangles outline stations and earthquakes from CERI data set, grouped into zones 1, 2, and 3. (b) Enlargement of zone 1.
Equal area plots show fast polarization directions for a given event (shown as lines), plotted by azimuth and distance in respect to the station’s
location (triangle in center), segmented ring indicates shear-wave window of 35° and outer ring at 40°. Dashed lines connect equal area
projection plots with station locations. (c) Enlargement of zone 1. Radial histograms (rose diagrams) indicate fast polarization directions
measured at each station. The width of each bin is equivalent to 10°, north indicated at top. Each segmented ring within the rose diagram
represents a frequency increment of two events. The dashed line indicates station location. (d) Enlargement of zone 2. Equal area plots
indicate fast polarization directions. (e) Enlargement of zone 2. Radial histograms (rose diagrams) indicate fast polarization directions measured at each station. (f) Enlargement of zone 3. Equal area plots indicate fast polarization directions. (g) Enlargement of zone 3. Radial
histograms (rose diagrams) indicate fast polarization directions measured at each station. (h) Radial histogram (rose diagram) of all measured
fast polarization directions from the CERI data set. The width of each bin is equivalent to 10°, north indicated at top. Each segmented ring
within the rose diagram represents a frequency increment of five events. The most frequent (northeast–southwest) fast polarization direction
as well as the secondary west-northwest–east-southeast direction are clearly visible. Gray arrows indicate direction of maximum horizontal
compressional stress within the NMSZ (Zoback and Zoback, 1991; Hurd and Zoback, 2012).
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P. Martin, P. Arroucau, and G. Vlahovic
Figure 6. (a) Enlargement of zone 1 (see Fig. 5a). Time delays for a given event are shown as scaled circles inside the equal area
projection, plotted by azimuth and distance in respect to the station location (triangle in center), segmented ring indicates shear-wave window
of 35° and outer ring at 40°. The dashed line indicates station location. (b) Enlargement of zone 2. Time delays for a given event are shown as
scaled circles inside the equal area projection, plotted by azimuth and distance in respect to the station’s location (triangle in center), segmented ring indicates shear-wave window of 35° and outer ring at 40°. The dashed line indicates station location. (c) Enlargement of zone 3.
Time delays for a given event are shown as scaled circles inside the equal area projection, plotted by azimuth and distance in respect to the
station’s location.
Our results indicate a wide range of path-normalized
time delays. The path-normalized time delays are calculated
by dividing each time delay by the length of its associated
travel path δt=l, which results in values ranging from 1 to
33 ms=km. As noted by Rowlands et al. (1993), actual pathnormalized time delays could be even larger if the thick top
layers of sediments in the Mississippi embayment were considered isotropic, which would reduce the effective path
length in the anisotropic medium, thus increasing the pathnormalized time delays. Stations GRAT, LEPT, and RDGT
surround a cluster of the largest time delays in the section of
southern Reelfoot. Detailed results for these stations can be
seen in Figure 6a–c.
It is important to note that path-normalized time delays
larger than 20 ms=km in this study have an associated event
depth range between 5 and 10 km (Fig. 7). The short-effective
path length due to shallow event depths in the southern Reel-
foot region coupled with the large time delays measured there
suggest a high density of EDA cracks and a higher percent of
SWVA, possibly due to the presence of high pore fluid pressure. There is also a possibility that this behavior results from
the confinement of EDA anisotropy to above 10 km depth.
This is consistent with the results from the Bisrat et al.
(2012) study of swarm activity in the southern Reelfoot region
between 1995 and 2008. Their study identified the Ridgley
swarm area as the most active swarm region in the NMSZ
and attributed it to changes in pore fluid pressure within highly
fractured basement rock. The locations of large time delays in
this study are also consistent with prominent low-velocity
anomalies from Powell et al. (2010). In their study, Powell
et al. (2010) describe the velocity anomalies that occur in
the northern portion of the Reelfoot fault as caused by variations in rock composition. However, they attribute the velocity
anomalies located in the central as well as the southern
Shear-Wave Splitting Study of Crustal Anisotropy in the New Madrid Seismic Zone
Figure 7. Path-normalized time delays plotted against event
depth for the CERI and PANDA data sets.
Reelfoot fault section to fluid-saturated, highly fractured
crust. In the central Reelfoot fault segment where low V P values are found, the low V S velocity anomalies are even more
pronounced. This suggests that the resultant high V P =V S ratio
is consistent with high pore pressure and fluid-filled, smallaspect ratio cracks. In the southern Reelfoot fault segment, the
low V P velocity anomalies are of smaller magnitude than the
low V S velocity anomalies. Therefore, Powell et al. (2010)
concluded that the resultant low V P =V S value was due to
the existence of high aspect ratio, fluid-filled cracks.
Results from the PANDA data set are similar to those from
the CERI data, including northeast–southwest dominant fast
polarization direction and a second, less dominant west-northwest–east-southeast polarization direction (see Fig. 8a,b).
Path-normalized time delays range from 2 to 31 ms=km, with
a cluster of large delays concentrated in the southern Reelfoot
fault segment, analogous to the results from the CERI data
analysis (Fig. 9). The larger path-normalized time delays from
the PANDA data are not located as far south as those from the
CERI analysis but are still within the area where the presence
of high pore fluid pressure is indicated by other studies
(Powell et al., 2010; Bisrat et al., 2012).
Discussion
The majority of fast polarization directions found in this
study were northeast–southwest oriented, which is consistent
with the dominant fast polarization direction identified by
Rowlands et al. (1993) in their shear-wave splitting study
and is coincident to North American stress field (see Zoback
and Zoback, 1980; Heidbach et al., 2008). Both of these
results are attributed to EDA cracks polarized in the direction
of maximum horizontal compressional stress within the
NMSZ (Zoback and Zoback, 1991; Hurd and Zoback, 2012).
The second dominant fast polarization direction found in this
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study is west-northwest–east-southeast, which was not identified by Rowlands et al. (1993). The widespread distribution
of this secondary fast polarization direction throughout the
NMSZ suggests that it is not due to a local change in maximum horizontal stress. Fast polarization directions offset up
to 90° from the direction of the crack strike are possible in
cases of shallow-dipping or intersecting fracture systems
(Rial et al., 2005). Thus, geometry of upper crustal
anisotropy in NMSZ may be more complex than vertical EDA
cracks. Similarly, the distribution of the secondary westnorthwest–east-southeast fast polarization gives no evidence
of an association with any local igneous intrusions or plutons
in the region (Hildenbrand, 1985; Hildenbrand and Hendricks, 1995), which suggests that there is no local buildup
of stresses that occurs around their boundaries. For example,
the polarization directions measured closest to the large
Bloomifield pluton (located north-northeast of the Reelfoot
fault; Hildenbrand and Hendricks, 1995) exhibit similar directions to polarizations measured along the AF where more
felsic intrusions are believed to exist (Johnston and Schweig,
1996). The existence of a secondary dominant fast polarization direction, significantly offset from the direction of
regional maximum horizontal compressive stress, has been
observed before in other fault zones. For example, Zhang
et al. (2007) observed fast polarization directions recorded
at stations near the San Andreas fault that were parallel or
subparallel to the strike of the fault (northwest–southeast)
and significantly offset from north-northeast–southsouthwest fast polarizations that were suggested to be due
to the direction of regional maximum horizontal compressive
stress. The polarization directions oriented along the strike of
the fault were concluded to be due to the shear fabric in the
fault region. Another possible mechanism for two dominant
fast polarization directions can be a change of anisotropy
with depth (Winterstein and Meadows, 1991). Results of
our study show no indication of a relationship between depth
and fast polarization directions (Fig. 10).
Path-normalized time delays found in this study were
assumed to be caused by EDA cracks and ranged from 1 to
33 ms=km. The measured time delays in this study are assumed to be the result of EDA cracks within the shallow
crust. This assumption allows the path-normalized time delays combined with the average shear-wave velocity along
the travel path, to be related to crack densities in the region.
This relationship can be shown with a (unit-less) apparent
crack density, calculated by
δt
;
ε υS
l
3
in which ε is the apparent crack density, υS is the shear velocity in the uncracked medium, and δt=l is the path-normalized time delay (O’Connell and Budiansky, 1974; Hudson,
1981). The value for υS in this study was estimated using a
weighted average of the S velocities for all layers in the
velocity model of Chiu et al. (1992) along the straight ray
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P. Martin, P. Arroucau, and G. Vlahovic
Figure 8.
(a) Enlargement of PANDA station locations. Equal area plots indicate fast polarization directions for a given event (shown as
lines), plotted by azimuth and distance in respect to the station’s location (triangle in center). Segmented ring indicates shear-wave window of
35° and outer ring at 40°. The dashed line indicates station location. (b) Enlargement of PANDA station locations. Radial histograms (rose
diagrams) indicate fast polarization directions measured at each station. The width of each bin is equivalent to 10°, north indicated at top.
Each segmented ring within the rose diagram represents a frequency increment of two events. The dashed line indicates station location.
Figure 10.
Fast polarization directions plotted against depth for
CERI data set.
Figure 9.
Enlargement of PANDA station locations. Time delays
for a given event are shown as scaled circles inside the equal area
projection, plotted by azimuth and distance in respect to the station’s location (triangle in center), segmented ring indicates
shear-wave window of 35° and outer ring at 40°. The dashed line
indicates station location.
path for each event. The resulting average velocities are considered a conservative estimate, however, as the low S velocity (0:6 km=s) in the upper layer of sediments was included
in the calculation. A hundredth of the percentage of SWVA
(see equation 1) is considered to be the approximate value of
the crack density (Crampin, 1994), allowing the estimate of
SWVA through; SWVA εx100. This in turn gives estimated
differential shear-wave anisotropy between 1% and 8%.
These values are higher than those from the previous shearwave splitting study by Rowlands et al. (1993), in which
path-normalized time delays ranged from 1.4 to
14:6 ms=km, with anisotropy on the order of 4%. Large time
delays found in this study were concentrated in the southern
Reelfoot segment. The large time delays could be explained
by the existence of elevated pore fluid pressure and are consistent with the results of Bisrat et al. (2012), in which swarm
activity in this area was attributed to changes in pore fluid
pressure within highly fractured basement rock. Similarly,
Shear-Wave Splitting Study of Crustal Anisotropy in the New Madrid Seismic Zone
Powell et al. (2010) attribute the P- and S-velocity anomalies
that they identified in the same location of the southern Reelfoot fault section to a fluid-saturated volume, with a saturating medium of meteoric water or gas.
Hurd and Zoback (2012) concluded that the existence of
elevated pore fluid pressure, or local changes in stress, was
not needed to explain the faulting within the NMSZ. Instead
their study showed that faults in the NMSZ are favorably oriented for slip in the east-northeast–west-southwest compressive stress field. If this is the case, then the existence of high
pore fluid pressure in the Reelfoot region as indicated by this
study, Bisrat et al. (2012), and Powell et al. (2010), would
only help facilitate slip in the region.
Conclusion
In this work, we performed and analyzed shear-wave
splitting measurements from a total of 235 seismic events that
occurred within the NMSZ, recorded by the PANDA network
(Chiu et al., 1992) from the years 1989–1991 and by the CERI
from the years 2003 to 2011. The anisotropy detected by the
split waves is assumed to be due to EDA cracks in the shallow
crust. The results have shown two dominant fast polarization
directions throughout the NMSZ. The most frequent
northeast–southwest fast polarization direction is consistent
with EDA cracks oriented in the direction of maximum horizontal compressional stress within the NMSZ. The source of
the secondary west-northwest–east-southeast fast polarization direction may be, among others, shallow-dipping EDA
cracks, two intersecting fracture systems, or pervasive shear
fabric within the NMSZ. The path-normalized time delay measurements exhibit a wide range, from 1 to 33 ms=km, with a
cluster of large time delays occurring in the lower portion of
the southern Reelfoot section. The large time delays are
believed to be due to elevated pore fluid pressure within
volumes of highly fractured basement rock. These findings
agree with previous results from local earthquake tomography
(Powell et al., 2010) and microseismic swarm analysis (Bisrat
et al., 2012) and present a possible mechanism for the facilitation of slip in the region. There are many things that could be
done in future work to improve our understanding of the
NMSZ through shear-wave splitting from local earthquakes.
This includes increasing the density of stations in the region,
as well as possible tomographic inversions using time delays.
Such an inversion could aid in understanding the precise
location where the splitting occurred, which is a limitation
of the study presented here.
Data and Resources
A portion of the seismic data in this study was reported
by Center for Earthquake Research and Information (CERI),
Saint Louis University Earthquake Center, and contributions
from Southeastern U.S. Seismic Networks (SEUSSN) bulletins, at http://www.ceri.memphis.edu/seismic/catalogs/cat
_nm.html (last accessed January 2012). A portion of the seis-
1109
mic data used in this study was collected by the portable array for numerical data acquisition (PANDA) network from
October 1989–June 1991 (Chiu et al., 1992).
Acknowledgments
We would like to acknowledge the Center for Earthquake Research and
Information as well as Mitch Withers for making data available and Christine
Powell for sharing her insight into the New Madrid Seismic Zone. We also
acknowledge Jer-Ming Chiu for making the portable array for numerical
data acquisition data available. We would like to thank Associate Editor Anton Dainty and the two anonymous reviewers whose suggestions made this a
better paper. This study was in part supported by the National Science Foundation under Grant Number HRD-0833184 and National Aeronautics and
Space Administration under Award NNX09AV07A. Any opinions, findings,
and conclusions or recommendations expressed in this presentation are those
of the authors and do not necessarily reflect the views of the funding
agencies.
References
Bisrat, S., H. R. DeShon, and C. Rowe (2012). Microseismic swarm
activity in the New Madrid Seismic Zone, Bull. Seismol. Soc. Am.
102, 1167–1178.
Booth, D. C., and S. Crampin (1985). Shear-wave polarizations on a curved
wavefront at an isotropic free surface, Geophys. J. Int. 83, 31–45.
Calais, E., and S. Stein (2009). Time-variable deformation in the
New Madrid Seismic Zone, Science 323, 1442.
Calais, E., A. M. Freed, R. Van Arsdale, and S. Stein (2010). Triggering of
New Madrid seismicity by late-Pleistocene erosion, Nature 466,
608–611.
Chiu, J. M., A. C. Johnston, and Y. T. Yang (1992). Imaging the active faults
of the central NMSZ using PANDA array data, Seismol. Res. Lett. 63,
375–393.
Crampin, S. (1989). Suggestions for a consistent terminology for seismic
anisotropy, Geophys. Prospect. 37, 753–770.
Crampin, S. (1994). The fracture criticality of crustal rocks, Geophys. J. Int.
118, 428–438.
Crampin, S., and B. K. Atkinson (1985). Microcracks in the earth’s crust,
First Break 3, 16–20.
Csontos, R., and R. Van Arsdale (2008). New Madrid seismic zone fault
geometry, Geosphere 4, no. 5, 802–813.
Evans, M., J.-M. Kendall,, and R. Willeman (2006). Automated SKS splitting and upper-mantle anisotropy beneath Canadian seismic stations,
Geophys. J. Int. 165, no. 3, 931–942.
Frankel, A., R. Smalley, and J. Paul (2012). Significant motions between
GPS sites in the New Madrid region: Implications for seismic hazard,
Bull. Seismol. Soc. Am. 102, no. 2, 479–489.
Hamilton, R. M., and A. C. Johnston (1990). Tecumseh’s prophecy—
Preparing for the next New Madrid earthquake, U.S. Geol. Surv.
Circular 1066, 30 pp.
Heidbach, O., M. Tingay, A. Barth, J. Reinecker, D. Kurfeß, and B. Müller
(2008). The world stress map database release 2008, http://dc‑app3‑14
.gfz‑potsdam.de/ (last accessed December 2013).
Hildenbrand, T. G. (1985). Rift structure of the northern Mississippi embayment from the analysis of gravity and magnetic data, J. Geophys. Res.
90, 12607–12622.
Hildenbrand, T. G., and J. D. Hendricks (1995). Geophysical setting of the
Reelfoot rift and relations between rift structures and the New Madrid
seismic zone, U.S. Geol. Surv. Profess. Pap. 1538-E, 30 pp.
Himes, L., W. Stauder, and R. B. Hermann (1998). Indication of active faults
in the New Madrid seismic zone from precise location of hypocenters,
Seismol. Res. Lett. 59, 123–131.
Hudson, J. A. (1981). Wave speeds and attenuation of elastic waves in
material containing cracks, Geophys. J. Int. 64, 133–150.
1110
Hurd, O., and M. D. Zoback (2012). Regional stress orientations and slip
compatibility of earthquake focal planes in the New Madrid seismic
zone, Seismol. Res. Lett. 83, 110–122.
Johnston, A. C. (1996). Seismic moment assessment of earthquakes in stable
continental regions—III. New Madrid 1811–1812, Charleston 1886
and Lisbon 1755, Geophys. J. Int. 126, 314–344.
Johnston, A. C., and E. S. Schweig (1996). The Enigma of the New Madrid
earthquakes of 1811–18121, Annu. Rev. Earth. Planet. Sci. 24, 339–
384.
O’Connell, R. J., and B. Budiansky (1974). Seismic velocities in dry and
saturated cracked solids, J. Geophys. Res. 79, 5412–5426.
Powell, C. A., M. M. Withers, H. R. DeShon, and M. M. Dunn (2010).
Intrusions and anomalous V P =V S ratios associated with the
New Madrid Seismic Zone, J. Geophys. Res. 115, no. B08311, doi:
10.1029/2009JB007107.
Rial, J., M. Elkibbi, and M. Yang (2005). Shear-wave splitting as a tool for
the characterization of geothermal fractured reservoirs: Lessons
learned, Geothermics 34, 365–385.
Rowlands, H. J., D. C. Booth, and J. M. Chiu (1993). Shear-wave
splitting from microearthquakes in the New Madrid Seismic Zone,
Can. J. Explor. Geophys. 29, 352–362.
Savage, M. K., A. Wessel, N. A. Teanby, and A. W. Hurst (2010). Automatic
measurement of shear wave splitting and applications to time varying
anisotropy at Mount Ruapehu volcano, New Zealand, J. Geophys. Res.
115, no. B12321, doi: 10.1029/2010JB007722.
Silver, P. G., and W. W. Chan (1991). Shear wave splitting and
subcontinental mantle deformation, J. Geophys. Res. 96, 16429–
16454.
Teanby, N., J.-M. Kendall, and M. van der Baan (2004). Automation of
shear-wave splitting measurements using cluster analysis, Bull. Seismol. Soc. Am. 94, 453–463.
Tuttle, M. P., E. S. Schweig, J. D. Sims, R. H. Lafferty, L. W. Wolf, and
M. L. Haynes (2002). The earthquake potential of the New Madrid
Seismic Zone, Bull. Seismol. Soc. Am. 92, 2080–2089.
Vlahovic, G., C. A. Powell, and J.-M. Chiu (2000). Three-dimensional
wave velocity structure in the New Madrid seismic zone, J. Geophys.
Res. 105, 7999–8011.
P. Martin, P. Arroucau, and G. Vlahovic
Winterstein, D., and M. Meadows (1991). Changes in shear-wave
polarization azimuth with depth in Cymric and Railroad Gap oil fields,
Geophysics 56, no. 9, 1349–1364.
Zhang, H., Y. Liu, C. Thurber, and S. Roecker (2007). Three-dimensional
shear-wave splitting tomography in the Parkfield, California, region,
Geophys. Res. Lett. 34, L24308, doi: 10.1029/2007GL031951.
Zoback, M. L., and M. D. Zoback (1980). State of stress of the conterminous
United States, J. Geophys. Res. 85, 6113–6156.
Zoback, M. D., and M. L. Zoback (1991). Tectonic stress field of North
America and relative plate motions, Neotectonics N. Am. 1, 339–366.
EES Department
Lehigh University
1 West Packer Avenue
Bethlehem, Pennsylvania 18015
(P.M.)
Instituto Dom Luiz
Centro de Geofísica
Faculdade de Ciências da Universidade de Lisboa
Campo Grande, Ed. C8
1749-016 Lisboa, Portugal
[email protected]
(P.A.)
North Carolina Central University
Department of Environmental Earth and Geospatial Sciences
2202 Mary M. Townes Science Complex
Durham, North Carolina 27707
(G.V.)
Manuscript received 15 August 2013;
Published Online 13 May 2014