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A mechanism for sustained groundwater pressure changes
induced by distant earthquakes
Emily E. Brodsky,1 Evelyn Roeloffs,2 Douglas Woodcock,3 Ivan Gall,4
and Michael Manga5
Received 25 November 2002; revised 8 March 2003; accepted 3 April 2003; published 22 August 2003.
[1] Large, sustained well water level changes (>10 cm) in response to distant (more than
hundreds of kilometers) earthquakes have proven enigmatic for over 30 years. Here we
use high sampling rates at a well near Grants Pass, Oregon, to perform the first
simultaneous analysis of both the dynamic response of water level and sustained changes,
or steps. We observe a factor of 40 increase in the ratio of water level amplitude to seismic
wave ground velocity during a sudden coseismic step. On the basis of this observation
we propose a new model for coseismic pore pressure steps in which a temporary barrier
deposited by groundwater flow is entrained and removed by the more rapid flow induced
by the seismic waves. In hydrothermal areas, this mechanism could lead to 4  102 MPa
pressure changes and triggered seismicity. INDEX TERMS: 1829 Hydrology: Groundwater
hydrology; 7209 Seismology: Earthquake dynamics and mechanics; 7212 Seismology: Earthquake ground
motions and engineering; 7260 Seismology: Theory and modeling; 7294 Seismology: Instruments and
techniques; KEYWORDS: earthquakes, triggering, time-dependent hydrology, fractures
Citation: Brodsky, E. E., E. Roeloffs, D. Woodcock, I. Gall, and M. Manga, A mechanism for sustained groundwater pressure
changes induced by distant earthquakes, J. Geophys. Res., 108(B8), 2390, doi:10.1029/2002JB002321, 2003.
1. Introduction
[2] Earthquakes can produce sustained water level
changes in certain distant wells [Coble, 1965; Bower and
Heaton, 1978; Matsumoto, 1992; Roeloffs, 1998; King et
al., 1999] that are often orders of magnitude larger than can
be explained by static stress changes [Bower and Heaton,
1978]. Many researchers suggest that seismic waves interacting
with aquifers produce the sustained changes in pore
pressure, or steps, hundreds of kilometers from an earthquake
[Bower and Heaton, 1978; Roeloffs, 1998; King et
al., 1999]. The redistribution of pore pressure can generate
crustal deformation [Johnston et al., 1995] and perhaps
even trigger seismicity [Hill et al., 1993; Brodsky et al.,
2000; U.S. Geological Survey (USGS), 2000]. However, the
mechanism by which small cyclic stresses induce persistent
pore pressure changes has remained uncertain.
[3] Here we constrain the mechanism for coseismic steps
in a well near Grants Pass, Oregon, by using both high
sample rate water level data from the well and seismic data
from the broadband Berkeley Digital Seismic Network
station Yreka Blue Horn Mine (YBH) in Yreka, California
(Figure 1). The water level in a well penetrating a confined
aquifer is a manometer measuring the pore pressure at a
point. During 1993–2001, several seismic water level
oscillations and two coseismic steps were recorded digitally.
The 1 September 1994 Mw = 7.2 Petrolia, California
(epicentral distance  = 2.71), earthquake generated a
15 cm decrease in water level over 2.5 days, and the
30 September 1999 Mw = 7.4 Oaxaca, Mexico ( =
34.65) earthquake generated an immediate 11 cm decrease
in water level. We show that (1) the coseismic steps are
related to the passage of seismic waves, (2) the amplitude
of the water level oscillations relative to the seismic ground
velocity increased abruptly at the time of the step induced
by the Oaxaca earthquake, and (3) gradual water level
steps are consistent with pore pressure changes diffusing to
the well from within the aquifer. These observations
motivate a new model for distant water level changes.
Seismic waves remove a temporary barrier of sediment or
solid precipitate resulting in both an increase in the seismic
wave amplification and a persistent water level change. We
then test the model with a new observation during the
3 November 2002 Mw = 7.9 Alaska earthquake.
2. Observations
[4] The 91.4 m deep NVIP-3 well near Grants Pass,
Oregon, has been monitored continuously since 1984
[Woodcock and Roeloffs, 1996]. The well is drilled into a
fractured granodiorite confined aquifer and a float measures
the water level. The chart recorder installed in 1984 was
replaced in November 1993 with a digital data logger
recording at 1.7  103 or 1.1  103 Hz. If the water
level changed more than 0.6 mm, the sampling rate increased up to a maximum of 1 Hz. In October 1998 a pressure
transducer was added sampling at 1.7  103 Hz. Since
March 2001, 1 Hz data from both the float and the transducer
have been collected to verify that no instrumental delay is
introduced by the float. The well geometry and hydrological
properties are given in Table 1. The dynamic response
cannot be modeled for the eight earthquake-related water
level drops before 1994 [Woodcock and Roeloffs, 1996]
because the chart records lack sufficient resolution.
[5] Below we first discuss the hydrological and seismological
observations pertaining to the oscillatory response of
the well in the seismic frequency band (0.02–0.2 Hz). We
then present direct observations of steps in water level.
2.1. Oscillatory Well Response to Shaking
[6] During 1993–2001, several earthquakes produced
ground shaking on the order of mm s1 at the site and water well level oscillations with amplitudes 10 cm. These
large responses imply a large amplification in the wellaquifer
system (Figure 2). The water level displacement in
the well measures the head change in the aquifer induced by
the strain of the seismic waves. Hydraulic head h is defined
as h  p/rg  z, where p is the pore pressure, r is the density
of water, g is gravitational acceleration, and z is the elevation.
For waves in an elastic medium, strain is proportional to
particle velocity [e.g., Love, 1927, equation XIII.17]. Therefore
the amplification of the seismic waves in the well is
measured by the ratio c of the amplitude of the water level
oscillations to the particle velocity in the seismic waves. The
units of c are m/(m/s).
[7] The amplification factor c is computed by dividing
the observed well spectra from NVIP-3 by the seismically
observed vertical ground velocity spectra from YBH for the
records (Figure 3b). Amplitude corrections are applied to
the seismograms to account for differences in geometric
spreading and radiation pattern between YBH and the well.
These corrections are small (<15%) for all of the events
discussed in this paper. A 2002 seismic installation showed
that both YBH and NVIP-3 are hard rock sites and no site or soil correction is necessary. Regional earthquakes generate
relatively high-frequency signals (>0.05 Hz) while teleseismic
waves have lower frequencies. We use the regional
events to determine the high-frequency c and the teleseismic
events for the low-frequency c (Figure 3b). During
four regional earthquakes the average observed value of the
amplification factor c is 280 m/(m/s) at 0.06 Hz with a
standard deviation of 49 m/(m/s).
[8] The large value of c quantifies the large amplification
in the well and is primarily determined by the small specific
storage Ss. Specific storage is defined as the volume of fluid
a unit volume of aquifer releases under a unit decrease in
head. The definition can be shown to be equivalent to
Ss ¼ rgða þ fbÞ; ð1Þ
where a and b are the compressibility of the aquifer and
fluid, respectively, and f is the porosity [Freeze and Cherry,
1979]. Small Ss, indicative of low porosity, is expected for
crystalline rock. The small strains of the Earth tides (107)
which produce large amplitude signals (>6 cm) in the well
(Figures 1b, 1d, and 1f ) constrain Ss (Table 1). The hydraulic
conductivity and local structure also affect c. The flow
through the porous medium and the resonance of the water
column in the well both introduce a frequency dependence as
discussed below.
[9] We use a pumping test to constrain the geometry of the
aquifer near the well. Water was pumped from the well at a
rate of 20 GPM (20 GPM = 1.26  103 m3 s1) for 52 hours
(1 sample per 10 min or 1 sample per 15 min, respectively).
Figure 3 shows that the form of the drawdown curve is well
modeled by flow through a single, infinitesimally thin square
planar fracture embedded in an unbounded, homogeneous
and isotropic confined aquifer. Flux is constant over the
surface area of the fracture. The solution for the drawdown
curve in this geometry can be derived by integrating point
sources or using the Green’s function method of Gringarten
and Ramey [1973] to derive
h t ð Þ  h t ¼ 0 ð Þ ¼
1
Ss
Z t
0
QðtÞ
L2
1
2
ffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi
pKðt  tÞ=Ss
p
erf
L
4
ffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi
Kðt  tÞ=Ss
p
!2
dt; ð2Þ
where h(t) is the head in the well, Q(t) is the volumetric
withdrawal rate, K is the hydraulic conductivity, and L is the
fracture length. The relatively small-radius well intersects
and samples the fracture but has no effect on the modeled
flow. The shape of both the pumping test curve and the
seismic wave response are well-fit by a single fracture
model (Figure 3). A simpler, cylindrical flow model (Theis
solution) does not fit the data. The geometry is also
consistent with the drilling log which recorded a marked
increase in flow in the final 2 m of the well. Note that the
slope of 1/2 at early times on the log-log drawdown curve
precludes significant storage of water in the fracture and is
consistent with the small fracture aperture (1 mm) used in
the mechanism discussion below.
[10] The linearized response function of the well to
seismic waves is calculated following the method of Cooper
et al. [1965] with the geometry modified to be consistent
with the pumping test data in Figure 3. Cooper et al. [1965]
assumed cylindrical flow into the well; we assume strictly
linear flow into the fracture that intersects the well. The pore
pressure changes in the aquifer are assumed to be proportional
to the dilatational strain of the seismic waves. The
most significant dilatation is generated by the Rayleigh
wave, rather than the body waves; therefore we only
consider the Rayleigh waves as sources of pressure oscillations.
The flow directly into the well is assumed negligible
because the surface area of the fracture is much greater than
the well. The amplification c( f ) is
cð f Þ ¼ Aðn; cÞ 1 
4p2H f2
g
þ
p
2
r2
w
L2
ffiffiffiffiffiffiffiffi
p f
KSs
s
ð1 þ iÞ


1
; ð3Þ
where A(n, c) is the ratio of the dilatational strain to the
vertical ground velocity which is a function of the Poisson’s
ratio n and the seismic phase velocity c,  is the ratio of the
confined aquifer fluid pressure to changes in the dilatational
static strain and is inferred from the tidal response
[Rojstaczer and Agnew, 1989], f is frequency, H is the
water column height, and rw is the well bore radius. The
quantity A(n, c) is 0.46/c for a Rayleigh wave in a Poisson
solid [Ben-Menahem and Singh, 1981]. We use c =
3.7 km s1 as a representative value for the Rayleigh wave
phase velocity at the site. On the basis of records of the
2002 Denali earthquake, the actual frequency-dependent
phase velocity could possibly vary by as much as 3.3–
4.1 km s1 for 20–35 s Rayleigh waves at this site. The
variability in c introduces an error of as much as 10% in c.
More significantly, we expect some coupling between the
fluid pressure and the other seismic phases. In fractured
rock, shear stresses can produce oscillations in some
orientations and Love waves are occasionally observed in
this well. However, full prediction of the effects of all the
seismic phases requires that we know the full threedimensional
strain tensor at the site. Such an inversion
would require a local seismic array, which was not
emplaced at the time of the earthquakes studied here. We
show below that the overall spectral response and relative
changes can be satisfactorily explained with a simplified
model in which only the large dilatation of the Rayleigh
waves affects the hydrological system. The specific
parameters of the response is also consistent with
independent measurements from a pumping test. Therefore
we accept the simplified model of equation (3) as an
adequate representation of the coupling.
2.2. Water Level Steps
[11] The rapid step in water level during the Oaxaca
earthquake occurred during the seismic shaking (Figure 4).
The water level is recorded independently on both the float
and the pressure transducer. Therefore the step in the
record was not caused by instrumental error. Even given
a 10 min ambiguity in arrival time due to a station clock
error (Figure 4), the step must have occurred during the
passage of the seismic waves. We use the frequency
content of the dispersed Rayleigh wave to align the records
more precisely.
[12] The well record for the Oaxaca event shows only
small oscillations due to surface waves prior to the step (Figure 4). The float was clearly free to track the water level
as tides and random small fluctuations were recorded, yet
significant amplification of the Rayleigh wave train did not
occur until after the step. More specifically, before the step
the amplification factor c is less than 10 m/(m/s) at 0.06 Hz,
whereas after the step c at 0.06 Hz resumes a nearly normal
value of 380 ± 19 m/(m/s) (Figure 4b). Since neither the
tidal amplification nor the well geometry changes with time,
the change in c must be due to a local change in K or, more
plausibly, the geometry of the fracture. Equation (3) shows
that much larger changes in K than L are required to achieve
a significant effect. The reduction in c before the step by a
factor of nearly 40 requires at least a 75% reduction in
fracture length L if all other parameters in equation (3) are
constant (Figure 4). As will be detailed below, a simultaneous
change in c and a step can occur if the seismic waves
remove a temporary low permeability barrier in the fracture
near the well. Removing the barrier both returns the
amplification to normal and allows the well to drain. A
local blockage in the fracture has only small effects on the
amplitude of the aquifer tidal signal as the long-period tidal
stresses sample the average wall rock (matrix) properties.
[13] Drops in water level can be sudden, as in the case of
Oaxaca, or gradual, as in the case of Petrolia (Figure 1). The
different drop durations correspond to differing distances
from the well to localized sources of pore pressure change
[Roeloffs, 1998]. For times much greater than the duration
of the seismic wave train, we model the effect of the
pressure step on the well as the one-dimensional solution
to a diffusion equation in an unbounded, homogeneous
aquifer [Crank, 1975]
W ¼ W0 
p
rg
erfc d=
ffiffiffiffiffiffiffiffi
4Dt
 p
; ð4Þ
where W is the water level, W0 is the initial water level, p
is the amplitude of the pressure drop at the source, d is the
distance from the source of the pore pressure change to the
fracture, t is the time since the drop at the source, and D 
K/Ss is the hydraulic diffusivity [Freeze and Cherry, 1979].
On the basis of the pumping test and the seismic wave response, D in the granodiorite matrix is 0.2 m2 s1
(Table 1). A least squares fit of the Petrolia water level
record to equation (4) yields d = 70 m and p/rg = 22 cm
(Figure 5). The misfit at large times may be due to unrelated
seasonal trends or the flow encountering aquifer boundaries
not sampled in the 52 hour pumping test. Unfortunately,
there is no other monitored well sampling the same deep,
confined aquifer within a 1 km radius; therefore we cannot
triangulate to determine a more accurate location for the
pressure source.
[14] The process generating the step did not affect c in
the well during the Petrolia earthquake since d greatly
exceeds the pressure diffusion length scale for the seismic
waves (
ffiffiffiffiffiffi
Dt
p
= 2 m, where t is the dominant wave period of
21 s). For Oaxaca, the drop is a step function to within the
data resolution and the solution to the diffusion equation is
the trivial one, i.e., d = 0 and p/rg = 11 cm. Both steps
are consistent with a localized, instantaneous source of
pressure. The consistency suggests that the same mechanism
is active both near and far from the well, i.e., the steps
are not generated by the well bore itself.
3. Mechanism
[15] The static stress change at the well directly generated
by the 3850 km distant ( = 34.650) Mw = 7.4 Oaxaca
earthquake is less than 0.2 Pa. A static stress change of
103 Pa is required to explain the observed 11 cm water level
drop. Therefore we can eliminate static stress as the cause of
the drop and limit our investigations to a detailed study of
the dynamic stresses (seismic waves).
[16] On the basis of the observations we propose a new
model for water level changes in wells far from an earthquake
(Figure 6). The permeability structure is dominated
by highly conductive fractures. From time to time, these
fractures become clogged with weathering products and low
permeability flocs of colloidal material. The fluid pressure
on the upgradient (or upstream) side of the low-permeability
barrier in steady state increases relative to the unclogged
state (Figure 6b). In addition, the effective fracture size is
smaller and the value of c measured by an intersecting
well is therefore reduced. When a seismic wave passes, it
induces rapid flow between the formation and fracture
which removes the barrier by loosening particles and
entraining them. Once the barrier is removed, water drains
from the well to produce a step in pressure as the permeability
structure returns to its normal state. If the barrier
forms at some distance from the fracture intersected by the
well, the pressure change diffuses gradually to the fracture,
but if the barrier is immediately adjacent to the fracture,
then the observed drop is very rapid.
[17] Water pumped from deep in the aquifer contains
4  107 micron-size aluminosilicate particles per liter.
These suspended weathering products could aggregate to
form the requisite blockages. Dense clay flocs have conductivities
comparable to the wall rock value 7  108 m s1
[Freeze and Cherry, 1979] and could thus effectively block
the fracture. For simplicity, we assume that the floc and wall
rock conductivities are identical and the water flows
through both according to Darcy’s law,
u ¼ K=ðfrgÞrp; ð5Þ
where u is the interstitial fluid velocity and f is the porosity.
The Darcy (volumetrically average) velocity is fu. In order
to deposit a 0.3 mm thick barriers of densely packed
colloids in two years, we require a Darcy velocity of 5 m d1.
In steady state, the head drop across this barrier is 25 cm
(2.5 kPa of pressure). When the barrier is removed, the
upgradient side returns to the initial pressure P0 by dropping
by half the head difference, or 13 cm.
[18] The fluid velocity at the edge of the flocs adjacent to
the wall is of the same order as that draining the matrix by
continuity. The seismic waves induce a flow rate Qclog that
is observed during Oaxaca to be 5  105 m3 s1 during the
Rayleigh waves before the step. If the fracture thickness w is
1 mm and the clogged fracture length L is at most 30 m as
inferred from the response (Figure 4), the fluid velocity is
Qclog/wL = 1 mm s1. Viscous stress on the particles is
approximately hu/R, where h is the viscosity of water
(103 Pa s) and R is the radius of the particle (1 mm). The
interstitial fluid velocity u is at least the Darcy velocity, fu,
and can be substantially greater if f is small. The resulting
viscous shear is 1 Pa which is sufficient to initiate motion
and disaggregate flocs [Kessler, 1993]. Low permeability
clay flocs can have very high porosity (>99%), therefore
when the flocs are disaggregated, the separated micronscale
clay particles no longer have a significant effect on the
permeability [Kessler, 1993]. It is possible that the small
amplitude solid strains of the shaking also contribute
directly to loosening the barriers, but there is no need to
invoke such a difficult mechanism as the observed induced
flow velocities are sufficient for entrainment.
[19] The gradual Petrolia drop suggests that seismic
waves can induce floc-entraining flow velocities outside
the immediate vicinity of the well bore. Seismic waves
induce pressure gradients between any zones with different
values of , i.e., different compressibilities and porosities.
The head difference between two geological units can be
80% of the head difference between one of the units and the
well bore. The pressure difference p12 between two units
with tidal amplifications 1 and 2 in response to a
dilatational strain q is (1  2/1)1q. If a well is drilled
into the 1 unit, the pressure difference pw between the
unit and the well is 1q. Therefore the ratio p12/pw is
(1  2/1). The range of  observed in nature is at least a
factor of 5 [Roeloffs, 1998]. In open fractures,  is 1. As
long as these contrasts are sustained over sharp boundaries,
such as a fracture wall, flow will be on the order of that
observed in the well. In a fracture that does not intersect a
well, there will be no water column resonance, i.e., the
second term of equation (3) will be negligible. The absence
of water column resonance far from the well reduces flow
velocities by only 40% relative to those observed here.
[20] The Oaxaca earthquake was followed 2 weeks later
by shaking with 20% greater vertical ground velocity from
the Hector Mine earthquake (Figure 2). No rainfall occurred
between the events and all other observables were indistinguishable,
yet the second event produced no step. The
above model predicts that a barrier could not reform within
2 weeks as the particle concentration is too low. Therefore
no step is expected with the second earthquake despite its
size.
[21] One observation not directly addressed by the model
is that the pressure changes at the well are always drops.
An abundant source of weathering material or disequilibrium
precipitation may exist downgradient of the well.
Alternatively, the fracture may have an easily blocked
constriction downgradient. In either case, the downgradient
location of the blockage would then favor drops at the
well.
[22] Other mechanisms that have been suggested for farfield
coseismic pore pressure changes include mobilization of
gas bubbles [Linde et al., 1994; Roeloffs, 1998; Sturtevant et
al., 1996], shaking-induced dilatancy [Bower and Heaton,
1978] and fracture of an impermeable fault [King et al.,
1999]. The high tidal amplification implies that a compressible
gas phase comprises <104% of the aquifer; therefore
bubbles cannot account for the observed water level drop.
The other two mechanisms fail to explain the observed drop
in effective permeability followed by the return to the original
value. Both dilatancy and fracture models predict that the
effective permeability, and hence c, should increase immediately
following the earthquake. We observe that the poststep
response of Oaxaca is of the same order as the long-term
average.
4. The 3 November 2002 Mw = 7.9 Denali
Earthquake
[23] While this manuscript was in preparation, another
large earthquake generated a coseismic step at NVIP-3. The 3 November 2002 Denali earthquake (Figure 1) allowed us
to independently test the model already developed based on
the Oaxaca data.
[24] The Denali earthquake ( = 25.2) generated a
sudden drop in water level during the shaking (Figure 7).
The ambiguity in aligning the seismic and hydrologic
records is <5 s. Uncertainties in site effects were eliminated
since there was a Streickeisen STS-2 broadband seismometer
at the well site beginning in December 2001. Like
Oaxaca, the Denali record shows very little response to
shaking before the step and a normal response afterward.
However, unlike the previous case, the step occurred at the
beginning of the Rayleigh wave, therefore the responses
before and after cannot be directly compared as in Figure 4c.
For Denali, the water level is responding first to the S and
then to the Rayleigh wave. Shear phase coupling is observed
for other earthquakes and is probably due to an
anisotropic poroelastic response in the fractured rock
[Wang, 2000]. The Oaxaca earthquake appears to have been
very convenient in that it had a long, dispersed Rayleigh
wave which generated a constant excitation with a varying
response.
[25] Before the 12 cm step, the peak flow rate was 4 
105 m3 s1, i.e., nearly identical to the 5  105m3 s1
value of Qclog observed for Oaxaca. The entrainment
threshold for Denali is consistent with that proposed for
Oaxaca.
[26] The 4 mm s1 shaking of the Denali earthquake is
some of the strongest shaking recorded on site. The last
earthquake with a normal oscillatory record at the well prior
to the Denali event was the 23 June 2001 Mw = 8.4 Peru
earthquake that produced 17 cm oscillatory motion in the
well in response to 0.7 mm s1 shaking (Figure 8). The
observation is consistent with the predicted amplitude of
19 ± 2 cm calculated from equation (3) using the unblocked
fracture parameters in Table 1. In contrast, the 23 October
2002 Mw = 6.7 foreshock to the Denali event shook the
Oregon site with an amplitude of 0.04 mm s1. Although
this shaking is smaller than most earthquakes we studied,
according to equation (3) with the values in Table 1, the
water level should have oscillated with an amplitude of 1 cm.
The observed water level was less than the 0.6 mm trigger
for high sample rate recording. This order of magnitude
suppression of water level shaking is consistent with the barrier model. If the barrier formed over the 1.4 year interval
between the Peru and Denali earthquake and all other
parameters are the same as in the above model, we predict
a 9 cm head drop could occur during barrier removal. This
result from our order-of-magnitude model is reasonably near
the observed value of 12 cm.
5. Implications for Earthquakes
[27] We have constrained here a naturally occurring
process that suddenly redistributes pore pressure on fractures.
Faults, like fractures, form hydrological boundaries
that have contrasting porosity (storage) with the surrounding
rock and can accumulate sediment. Seismically induced
pore pressure steps can occur by the mechanism proposed
here in any hydrogeological systems that has (1) low matrix
specific storage, (2) fractures or faults, and (3) a source of
material for clogging. At least one other well-studied site of
coseismic steps is in a granitoid pluton with conditions 1–2
documented and condition 3 likely [King et al., 1999].
[28] Recent studies of regional-scale seismic triggering
suggest that seismic waves generate pore pressure
changes in geothermal areas that in turn generate seismicity
[Hill et al., 1993; Brodsky et al., 2000]. Geothermal
systems satisfy all three conditions above. In
particular, rapid precipitation of minerals is more common
in geothermal areas than in ordinary hydrogeologic environments
because of the large temperature and chemical
gradients [Lowell et al., 1993]. We speculate that high
incidences of triggered seismicity have been observed in
geothermal systems [Hill et al., 1993; Brodsky et al.,
2000; USGS, 2000] because of unstable permeability
structures generated by hot, circulating fluids. Shaking
by seismic waves loosens new precipitate and readjusts
the pore pressure on faults and fractures as is observed in
the Grants Pass well. The rapid redistribution of pore
pressure may promote earthquakes by quickly reducing
the effective stress on faults locally. Although the NVIP-3
well records only water level drops, the barrier removal
model requires pore pressure increases on the downgradient
end of the fracture or fault.
[29] In geothermal areas under typical conditions, precipitation
rates can be a factor of 15 greater than modeled for
Grants Pass [Lowell et al., 1993] resulting in barrier thicknesses
and step amplitudes proportionally larger. If all other
factors are equal to what we observe and model above, the
pressure change on a fault in a geothermal area will be 4 
102 MPa which is sufficient to trigger an earthquake
according to static stress studies [Hardebeck et al., 1998].
6. Conclusions
[30] The water well record from the Oaxaca earthquake
presented here is the first high sample rate recording of a
rapid far-field coseismic well step ever published to the best
of our knowledge. Such records are rare because (1) only a
small fraction of wells show far-field coseismic steps and
(2) hydrogeologists normally record water level at rates no
greater than 2  103 Hz. We analyzed this unique water
well record in conjunction with seismic data to show that during a rapid drop in water level, as the well drains, the
response c increases from the unusually low value of 9 ±
2 m/(m/s) to a more normal value of 380 ± 19 m/(m/s). We
interpret the change as the removal of a temporary blockage
in a fracture. Whether or not a step occurs depends on the
preexisting hydrogeology as well as the seismic input.
[31] Acknowledgments. This work was supported in part by a NSF
Earth Science Postdoctoral Fellowship and the Miller Institute for Basic
Research. We thank N. Beeler, D. Hill and S. Rojstaczer for reviews of an
early form of this manuscript. J. Barker gave insights into interpreting the
pumping test data, P. Hsieh advised on flow in fractured rocks and C. Chen
contributed SEM imaging of the particulates. Data from Yreka Blue Hill are
courtesy Berkeley Digital Seismic Network. PASSCAL instrument center
provided the STS-2 seismometer for the 2002 installation.

Recordings

  • makale ( recorded by David_B ), American (Midwest)

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    A mechanism for sustained groundwater pressure changes induced by distant earthquakes
    Emily E. Brodsky, Evelyn Roeloffs, Douglas Woodcock, Ivan Gall, Michael Manga

    First published: 22 August 2003

    ABSTRACT

    Large, sustained well water level changes (>10 cm) in response to distant (more than hundreds of kilometers) earthquakes have proven enigmatic for over 30 years. Here we use high sampling rates at a well near Grants Pass, Oregon, to perform the first simultaneous analysis of both the dynamic response of water level and sustained changes, or steps. We observe a factor of 40 increase in the ratio of water level amplitude to seismic wave ground velocity during a sudden coseismic step. On the basis of this observation we propose a new model for coseismic pore pressure steps in which a temporary barrier deposited by groundwater flow is entrained and removed by the more rapid flow induced by the seismic waves. In hydrothermal areas, this mechanism could lead to 4 × 10^(−2) MPa pressure changes and triggered seismicity.

  • makale ( recorded by David_B ), American (Midwest)

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    Introduction

    Earthquakes can produce sustained water level changes in certain distant wells that are often orders of magnitude larger than can be explain
    ed by static stress changes. Many researchers suggest that seismic waves interacting with aquifers produce the sustained changes in pore pressure, or steps, hundreds of kilometers from an earthquake. The redistribution of pore pressure can generate crustal deformation and perhaps even trigger seismicity. However, the mechanism by which small cyclic stresses induce persistent pore pressure changes has remained uncertain.

    Here we constrain the mechanism for coseismic steps in a well near Grants Pass
    , Oregon, by using both high sample rate water level data from the well and seismic data from the broadband Berkeley Digital Seismic Network station Yreka Blue Horn Mine (YBH) in Yreka, California (Figure 1). The water level in a well penetrating a confined aquifer is a manometer measuring the pore pressure at a point. During 1993–2001, several seismic water level oscillations and two coseismic steps were recorded digitally. The 1 September 1994 Mw = 7.2 Petrolia, California (epicentral distance Δ = 2.71°), earthquake generated a 15 cm decrease in water level over 2.5 days, and the 30 September 1999 Mw = 7.4 Oaxaca, Mexico (Δ = 34.65°) earthquake generated an immediate 11 cm decrease in water level. We show that (1) the coseismic steps are related to the passage of seismic waves, (2) the amplitude of the water level oscillations relative to the seismic ground velocity increased abruptly at the time of the step induced by the Oaxaca earthquake, and (3) gradual water level steps are consistent with pore pressure changes diffusing to the well from within the aquifer. These observations motivate a new model for distant water level changes. Seismic waves remove a temporary barrier of sediment or solid precipitate resulting in both an increase in the seismic wave amplification and a persistent water level change. We then test the model with a new observation during the 3 November 2002 Mw = 7.9 Alaska earthquake.

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    Observations

    The 91.4 m deep NVIP-3 well near Grants Pass, Oregon, has been monitored continuously since 1984. The well is drilled into a fractured granodiorite confined aquifer and a float measures the water level. The chart recorder installed in 1984 was replaced in November 1993 with a digital data logger recording at 1.7 × 10^(−3) or 1.1 × 10^(−3) Hz. If the water level changed more than 0.6 mm, the sampling rate increased up to a maximum of 1 Hz. In October 1998 a pressure transducer was added sampling at 1.7 × 10^(−3) Hz. Since March 2001, 1 Hz data from both the float and the transducer have been collected to verify that no instrumental delay is introduced by the float. The well geometry and hydrological properties are given in Table 1. The dynamic response cannot be modeled for the eight earthquake-related water level drops before 1994 because the chart records lack sufficient resolution.

    Below we first discuss the hydrological and seismological observations pertaining to the oscillatory response of the well in the seismic frequency band (0.02–0.2 Hz). We then present direct observations of steps in water level.

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    Oscillatory Well Response to Shaking

    During 1993–2001, several earthquakes produced ground shaking on the order of mm s^(−1) at the site and water well level oscillations with amplitudes ∼10 cm. These large responses imply a large amplification in the well-aquifer system (Figure 2). The water level displacement in the well measures the head change in the aquifer induced by the strain of the seismic waves. Hydraulic head h is defined as h ≡ p/ρg − z, where p is the pore pressure, ρ is the density of water, g is gravitational acceleration, and z is the elevation. For waves in an elastic medium, strain is proportional to particle velocity. Therefore the amplification of the seismic waves in the well is measured by the ratio χ of the amplitude of the water level oscillations to the particle velocity in the seismic waves. The units of χ are m/(m/s).

    The amplification factor χ is computed by dividing the observed well spectra from NVIP-3 by the seismically observed vertical ground velocity spectra from YBH for the records (Figure 3b). Amplitude corrections are applied to the seismograms to account for differences in geometric spreading and radiation pattern between YBH and the well. These corrections are small (<15%) for all of the events discussed in this paper. A 2002 seismic installation showed that both YBH and NVIP-3 are hard rock sites and no site or soil correction is necessary. Regional earthquakes generate relatively high-frequency signals (>0.05 Hz) while teleseismic waves have lower frequencies. We use the regional events to determine the high-frequency χ and the teleseismic events for the low-frequency χ (Figure 3b). During four regional earthquakes the average observed value of the amplification factor χ is 280 m/(m/s) at 0.06 Hz with a standard deviation of 49 m/(m/s).

    The large value of χ quantifies the large amplification in the well and is primarily determined by the small specific storage Ss. Specific storage is defined as the volume of fluid a unit volume of aquifer releases under a unit decrease in head. The definition can be shown to be equivalent to

    Ss = ρg ( α + φ β),

    where α and β are the compressibility of the aquifer and fluid, respectively, and ϕ is the porosity
    . Small Ss, indicative of low porosity, is expected for crystalline rock. The small strains of the Earth tides (10^(−7)) which produce large amplitude signals (>6 cm) in the well (Figures 1b, 1d, and 1f) constrain Ss (Table 1). The hydraulic conductivity and local structure also affect χ. The flow through the porous medium and the resonance of the water column in the well both introduce a frequency dependence as discussed below.

    We use a pumping test to constrain the geometry of the aquifer near the well. Water was pumped from the well at a rate of 20 GPM (20 GPM = 1.26 × 10^−3 m^3 s^−1) for 52 hours (1 sample per 10 min or 1 sample per 15 min, respectively). Figure 3 shows that the form of the drawdown curve is well modeled by flow through a single, infinitesimally thin square planar fracture embedded in an unbounded, homogeneous and isotropic confined aquifer. Flux is constant over the surface area of the fracture. The solution for the drawdown curve in this geometry can be derived by integrating point sources or using the Green's function method of Gringarten and Ramey to derive

    h(t) - h(t=0) = 1/Ss ∫ _(0)^(t) Q(τ)/L^(2) * 1/(2 √(π K(t - τ)/Ss)) * (erf( L / ( 4 √ (K(t - τ)/Ss)))^2 dτ

    where h(t) is the head in the well
    , Q(t) is the volumetric withdrawal rate, K is the hydraulic conductivity, and L is the fracture length. The relatively small-radius well intersects and samples the fracture but has no effect on the modeled flow. The shape of both the pumping test curve and the seismic wave response are well-fit by a single fracture model (Figure 3). A simpler, cylindrical flow model (Theis solution) does not fit the data. The geometry is also consistent with the drilling log which recorded a marked increase in flow in the final 2 m of the well. Note that the slope of ∼1/2 at early times on the log-log drawdown curve precludes significant storage of water in the fracture and is consistent with the small fracture aperture (1 mm) used in the mechanism discussion below.

    The linearized response function of the well to seismic waves is calculated following the method of Cooper et al. with the geometry modified to be consistent with the pumping test data in Figure 3. Cooper et al. assumed cylindrical flow into the well; we assume strictly linear flow into the fracture that intersects the well. The pore pressure changes in the aquifer are assumed to be proportional to the dilatational strain of the seismic waves. The most significant dilatation is generated by the Rayleigh wave, rather than the body waves; therefore we only consider the Rayleigh waves as sources of pressure oscillations. The flow directly into the well is assumed negligible because the surface area of the fracture is much greater than the well. The amplification χ(f) is

    χ(f) = A (ν, c) Γ | 1 - (4 π^(2) H f^(2))/g + π/2 r_(w)^(2)/L^(2) √ ( (π f)/(KSs)) (1+i)|^(-1)

    where A(ν, c) is the ratio of the dilatational strain to the vertical ground velocity which is a function of the Poisson's ratio ν and the seismic phase velocity c, Γ is the ratio of the confined aquifer fluid pressure to changes in the dilatational static strain and is inferred from the tidal response, f is frequency, H is the water column height, and r_(w) is the well bore radius. The quantity A(ν, c) is 0.46/c for a Rayleigh wave in a Poisson solid. We use c = 3.7 km s^(−1) as a representative value for the Rayleigh wave phase velocity at the site. On the basis of records of the 2002 Denali earthquake, the actual frequency-dependent phase velocity could possibly vary by as much as 3.3–4.1 km s^(−1) for 20–35 s Rayleigh waves at this site. The variability in c introduces an error of as much as 10% in χ. More significantly, we expect some coupling between the fluid pressure and the other seismic phases. In fractured rock, shear stresses can produce oscillations in some orientations and Love waves are occasionally observed in this well. However, full prediction of the effects of all the seismic phases requires that we know the full three-dimensional strain tensor at the site. Such an inversion would require a local seismic array, which was not emplaced at the time of the earthquakes studied here. We show below that the overall spectral response and relative changes can be satisfactorily explained with a simplified model in which only the large dilatation of the Rayleigh waves affects the hydrological system. The specific parameters of the response is also consistent with independent measurements from a pumping test. Therefore we accept the simplified model of equation (3) as an adequate representation of the coupling.

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    Water Level Steps

    The rapid step in water level during the Oaxaca earthquake occurred during the seismic shaking (Figure 4). The water level is recorded independently on both the float and the pressure transducer. Therefore the step in the record was not caus
    ed by instrumental error. Even given a 10 min ambiguity in arrival time due to a station clock error (Figure 4), the step must have occurred during the passage of the seismic waves. We use the frequency content of the dispersed Rayleigh wave to align the records more precisely.

    The well record for the Oaxaca event shows only small oscillations due to surface waves prior to the step (Figure 4). The float was clearly free to track the water level as tides and random small fluctuations were recorded, yet significant amplification of the Rayleigh wave train did not occur until after the step. More specifically, before the step the amplification factor χ is less than 10 m/(m/s) at 0.06 Hz, whereas after the step χ at 0.06 Hz resumes a nearly normal value of 380 ± 19 m/(m/s) (Figure 4b). Since neither the tidal amplification nor the well geometry changes with time, the change in χ must be due to a local change in K or, more plausibly, the geometry of the fracture. Equation (3) shows that much larger changes in K than L are required to achieve a significant effect. The reduction in χ before the step by a factor of nearly 40 requires at least a 75% reduction in fracture length L if all other parameters in equation (3) are constant (Figure 4). As will be detailed below, a simultaneous change in χ and a step can occur if the seismic waves remove a temporary low permeability barrier in the fracture near the well. Removing the barrier both returns the amplification to normal and allows the well to drain. A local blockage in the fracture has only small effects on the amplitude of the aquifer tidal signal as the long-period tidal stresses sample the average wall rock (matrix) properties.

    Drops in water level can be sudden, as in the case of Oaxaca, or gradual, as in the case of Petrolia (Figure 1). The different drop durations correspond to differing distances from the well to localized sources of pore pressure change. For times much greater than the duration of the seismic wave train, we model the effect of the pressure step on the well as the one-dimensional solution to a diffusion equation in an unbounded, homogeneous aquifer

    W = W_(0) - ∆ p/(ρg) erfc (d/√(4Dt)

    where W is the water level, W0 is the initial water level, Δp is the amplitude of the pressure drop at the source, d is the distance from the source of the pore pressure change to the fracture, t is the time since the drop at the source, and DK/Ss is the hydraulic diffusivity. On the basis of the pumping test and the seismic wave response, D in the granodiorite matrix is 0.2 m2 s1 (Table 1). A least squares fit of the Petrolia water level record to equation (4) yields d = 70 m and Δp/ρg = 22 cm (Figure 5). The misfit at large times may be due to unrelated seasonal trends or the flow encountering aquifer boundaries not sampled in the 52 hour pumping test. Unfortunately, there is no other monitored well sampling the same deep, confined aquifer within a 1 km radius; therefore we cannot triangulate to determine a more accurate location for the pressure source.

    The process generating the step did not affect χ in the well during the Petrolia earthquake since d greatly exceeds the pressure diffusion length scale for the seismic waves (√(
    4Dt) = 2 m, where τ is the dominant wave period of 21 s). For Oaxaca, the drop is a step function to within the data resolution and the solution to the diffusion equation is the trivial one, i.e., d = 0 and Δp/ρg = 11 cm. Both steps are consistent with a localized, instantaneous source of pressure. The consistency suggests that the same mechanism is active both near and far from the well, i.e., the steps are not generated by the well bore itself.

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    Mechanism

    The static stress change at the well directly generated by the 3850 km distant (Δ = 34.650) Mw = 7.4 Oaxaca earthquake is less than 0.2 Pa. A static stress change of 10^3 Pa is required to explain the observed 11 cm water level drop. Therefore we can eliminate static stress as the cause of the drop and limit our investigations to a detailed study of the dynamic stresses (seismic waves).

    On the basis of the observations we propose a new model for water level changes in wells far from an earthquake (Figure 6). The permeability structure is dominated by highly conductive fractures. From time to time, these fractures become clogged with weathering products and low permeability flocs of colloidal material. The fluid pressure on the upgradient (or upstream) side of the low-permeability barrier in steady state increases relative to the unclogged state (Figure 6b). In addition, the effective fracture size is smaller and the value of χ measured by an intersecting well is therefore reduced. When a seismic wave passes, it induces rapid flow between the formation and fracture which removes the barrier by loosening particles and entraining them. Once the barrier is removed, water drains from the well to produce a step in pressure as the permeability structure returns to its normal state. If the barrier forms at some distance from the fracture intersected by the well, the pressure change diffuses gradually to the fracture, but if the barrier is immediately adjacent to the fracture, then the observed drop is very rapid.

    Water pumped from deep in the aquifer contains 4 × 10^7 micron-size aluminosilicate particles per liter. These suspended weathering products could aggregate to form the requisite blockages. Dense clay flocs have conductivities comparable to the wall rock value 7 × 10^−8 m s^−1 and could thus effectively block the fracture. For simplicity, we assume that the floc and wall rock conductivities are identical and the water flows through both according to Darcy's law,

    u = - K / (ϕ ρ g) ∇ p

    where u is the interstitial fluid velocity and ϕ is the porosity. The Darcy (volumetrically average) velocity is ϕu. In order to deposit a 0.3 mm thick barriers of densely packed colloids in two years, we require a Darcy velocity of 5 m d1. In steady state, the head drop across this barrier is 25 cm (2.5 kPa of pressure). When the barrier is removed, the upgradient side returns to the initial pressure P0 by dropping by half the head difference, or 13 cm.

    The fluid velocity at the edge of the flocs adjacent to the wall is of the same order as that draining the matrix by continuity. The seismic waves induce a flow rate Qclog that is observed during Oaxaca to be 5 × 10^−5 m^3 s^−1 during the Rayleigh waves before the step. If the fracture thickness w is 1 mm and the clogged fracture length L is at most 30 m as inferred from the response (Figure 4), the fluid velocity is Qclog/wL = 1 mm s^−1. Viscous stress on the particles is approximately −ηu/R, where η is the viscosity of water (10^−3 Pa s) and R is the radius of the particle (1 μm). The interstitial fluid velocity u is at least the Darcy velocity, ϕu, and can be substantially greater if ϕ is small. The resulting viscous shear is 1 Pa which is sufficient to initiate motion and disaggregate flocs. Low permeability clay flocs can have very high porosity (>99%), therefore when the flocs are disaggregated, the separated micron-scale clay particles no longer have a significant effect on the permeability. It is possible that the small amplitude solid strains of the shaking also contribute directly to loosening the barriers, but there is no need to invoke such a difficult mechanism as the observed induced flow velocities are sufficient for entrainment.

    The gradual Petrolia drop suggests that seismic waves can induce floc-entraining flow velocities outside the immediate vicinity of the well bore. Seismic waves induce pressure gradients between any zones with different values of Γ, i.e., different compressibilities and porosities. The head difference between two geological units can be 80% of the head difference between one of the units and the well bore. The pressure difference Δp_(12) between two units with tidal amplifications Γ1 and Γ2 in response to a dilatational strain θ is (1 − Γ2/Γ1)Γ1 θ. If a well is drilled into the Γ1 unit, the pressure difference Δpw between the unit and the well is Γ1θ. Therefore the ratio Δp12/Δpw is (1 − Γ2/Γ1). The range of Γ observed in nature is at least a factor of 5. In open fractures, Γ is 1. As long as these contrasts are sustained over sharp boundaries, such as a fracture wall, flow will be on the order of that observed in the well. In a fracture that does not intersect a well, there will be no water column resonance, i.e., the second term of equation (3) will be negligible. The absence of water column resonance far from the well reduces flow velocities by only 40% relative to those observed here.

    The Oaxaca earthquake was followed
    2 weeks later by shaking with 20% greater vertical ground velocity from the Hector Mine earthquake (Figure 2). No rainfall occurred between the events and all other observables were indistinguishable, yet the second event produced no step. The above model predicts that a barrier could not reform within 2 weeks as the particle concentration is too low. Therefore no step is expected with the second earthquake despite its size.

    One observation not directly addressed by the model is that the pressure changes at the well are always drops. An abundant source of weathering material or disequilibrium precipitation may exist downgradient of the well. Alternatively, the fracture may have an easily blocked constriction downgradient. In either case, the downgradient location of the blockage would then favor drops at the well.

    Other mechanisms that have been suggested for far-field coseismic pore pressure changes include mobilization of gas bubbles, shaking-induced dilatancy and fracture of an impermeable fault. The high tidal amplification implies that a compressible gas phase comprises <10^−4% of the aquifer; therefore bubbles cannot account for the observed water level drop. The other two mechanisms fail to explain the observed drop in effective permeability followed by the return to the original value. Both dilatancy and fracture models predict that the effective permeability, and hence χ, should increase immediately following the earthquake. We observe that the poststep response of Oaxaca is of the same order as the long-term average.

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    The 3 November 2002 Mw = 7.9 Denali Earthquake

    While this manuscript was in preparation, another large earthquake generated a coseismic step at NVIP-3. The 3 November 2002 Denali earthquake (Figure 1) allowed us to independently test the model already developed based on the Oaxaca data.

    The Denali earthquake (Δ = 25.2°) generated a sudden drop in water level during the shaking (Figure 7). The ambiguity in aligning the seismic and hydrologic records is <5 s. Uncertainties in site effects were eliminated since there was a Streickeisen STS-2 broadband seismometer at the well site beginning in December 2001. Like Oaxaca, the Denali record shows very little response to shaking before the step and a normal response afterward. However, unlike the previous case, the step occurred at the beginning of the Rayleigh wave, therefore the responses before and after cannot be directly compared as in Figure 4c. For Denali, the water level is responding first to the S and then to the Rayleigh wave. Shear phase coupling is observed for other earthquakes and is probably due to an anisotropic poroelastic response in the fractured rock. The Oaxaca earthquake appears to have been very convenient in that it had a long, dispersed Rayleigh wave which generated a constant excitation with a varying response.

    Before the 12 cm step, the peak flow rate was 4 × 10^−5 m^−3 s^−1, i.e., nearly identical to the 5 × 10^−5m^3 s^−1
    value of Qclog observed for Oaxaca. The entrainment threshold for Denali is consistent with that proposed for Oaxaca.

    The 4 mm s^−1 shaking of the
    Denali earthquake is some of the strongest shaking recorded on site. The last earthquake with a normal oscillatory record at the well prior to the Denali event was the 23 June 2001 Mw = 8.4 Peru earthquake that produced 17 cm oscillatory motion in the well in response to 0.7 mm s^−1 shaking (Figure 8). The observation is consistent with the predicted amplitude of 19 ± 2 cm calculated from equation (3) using the unblocked fracture parameters in Table 1. In contrast, the 23 October 2002 Mw = 6.7 foreshock to the Denali event shook the Oregon site with an amplitude of 0.04 mm s^−1. Although this shaking is smaller than most earthquakes we studied, according to equation (3) with the values in Table 1, the water level should have oscillated with an amplitude of 1 cm. The observed water level was less than the 0.6 mm trigger for high sample rate recording. This order of magnitude suppression of water level shaking is consistent with the barrier model. If the barrier formed over the 1.4 year interval between the Peru and Denali earthquake and all other parameters are the same as in the above model, we predict a 9 cm head drop could occur during barrier removal. This result from our order-of-magnitude model is reasonably near the observed value of 12 cm.

    Implications for Earthquakes

    We have constrained here a naturally occurring process that suddenly redistributes pore pressure on fractures. Faults, like fractures, form hydrological boundaries that have contrasting porosity (storage) with the surrounding rock and can accumulate sediment. Seismically induced pore pressure steps can occur by the mechanism proposed here in any hydrogeological systems that has (1) low matrix specific storage, (2) fractures or faults, and (3) a source of material for clogging. At least one other well-studied site of coseismic steps is in a granitoid pluton with conditions 1–2 documented and condition 3 likely.

    Recent studies of regional-scale seismic triggering suggest that seismic waves generate pore pressure changes in geothermal areas that in turn generate seismicity. Geothermal systems satisfy all three conditions above. In particular, rapid precipitation of minerals is more common in geothermal areas than in ordinary hydrogeologic environments because of the large temperature and chemical gradients. We speculate that high incidences of triggered seismicity have been observed in geothermal systems because of unstable permeability structures generated by hot, circulating fluids. Shaking by seismic waves loosens new precipitate and readjusts the pore pressure on faults and fractures as is observed in the Grants Pass well. The rapid redistribution of pore pressure may promote earthquakes by quickly reducing the effective stress on faults locally. Although the NVIP-3 well records only water level drops, the barrier removal model requires pore pressure increases on the downgradient end of the fracture or fault.

    In geothermal areas under typical conditions, precipitation rates can be a factor of 15 greater than modeled for Grants Pass resulting in barrier thicknesses and step amplitudes proportionally larger. If all other factors are equal to what we observe and model above, the pressure change on a fault in a geothermal area will be 4 × 10^−2 MPa which is sufficient to trigger an earthquake according to static stress studies.

    Conclusions

    The water well record from the Oaxaca earthquake presented here is the first high sample rate recording of a rapid far-field coseismic well step ever published to the best of our knowledge. Such records are rare because (1) only a small fraction of wells show far-field coseismic steps and (2) hydrogeologists normally record water level at rates no greater than 2 × 103 Hz. We analyzed this unique water well record in conjunction with seismic data to show that during a rapid drop in water level, as the well drains, the response χ increases from the unusually low value of 9 ± 2 m/(m/s) to a more normal value of 380 ± 19 m/(m/s). We interpret the change as the removal of a temporary blockage in a fracture. Whether or not a step occurs depends on the preexisting hydrogeology as well as the seismic input.

Comments

PrinceJerry
March 7, 2015

The equations need to be corrected before an audio can be made. Otherwise it will not make sense.

manthonyd1618
March 15, 2015

Agreed, much of this is incomprehensible.

timothy72
July 18, 2015

I agree with the previous two commenters. Might I also suggest breaking this up into smaller chunks of text. I can almost guarantee that ten requests of ~450 words will be fulfilled much more quickly than one request of ~4500 words.

David_B
Sept. 17, 2015

I agree that much of this is currently incomprehensible. Since the journal article information is at least comprehensible, it was not difficult to find the article in more readable format:

http://onlinelibrary.wiley.com/doi/10.1029/2002JB002321/full

Also, you do not have to record the entire article at once. You are free to submit audio as long or as short as you desire. I've seen some people post over 10 audio responses to a single audio request. And each audio response counts towards credits.

Hope this helps!

David_B
Sept. 19, 2015

Nevermind, it looks like you no longer receive credit for each submission. Only one credit per submission.

David_B
Sept. 19, 2015

*only one credit total for each person that records something in response to a request regardless of the number of audio submitted.

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