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JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 113, B04415, doi:10.1029/2007JB004936, 2008
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Plate motion at the ridge-transform boundary of the
south Cleft segment of the Juan de Fuca Ridge from
GPS-Acoustic data
C. David Chadwell1 and Fred N. Spiess1,2
Received 10 January 2007; revised 3 November 2007; accepted 21 December 2007; published 30 April 2008.
[1] We measure the present-day plate velocity of the Juan de Fuca Ridge 25 km off-axis
to be 63.6 ± 3.6 mm/a at S67.2°E ± 7.9° degrees (1-s) relative to the Pacific plate (PA).
This velocity was derived from GPS-Acoustic (GPSA) measurements in 2000, 2001,
2002, and 2003 that observed the position of a seafloor array (44°430N,130°030W, 2900 m
depth) with a repeatability of ±4–6 mm. Three transient events at the Juan de Fuca Ridge
and Blanco Transform account for 10% of this motion in viscoelastic modeling,
suggesting that the observed GPSA-PA velocity is due primarily to steady state plate
dynamics. Subtracting the modeled transient motion gives a velocity of 57.3 ± 3.9 mm/a at
S72.9°E ± 12.1° degrees (1-s), which is consistent at the 95% confidence level with the
velocity calculated from the Wilson (1993) 0–0.78 Ma Euler pole. Therefore this site
is interpreted to be in a region of continuous, full-rate plate motion, a robust result of
this study which holds with and without correcting for transient motions. These results
provide direct geodetic evidence that spreading occurs predominantly within 25 km of the
axis at this intermediate spreading-rate ridge. Previously reported geodetic monitoring
across the 1-km-wide axial valley from 1994–1999 and 2000–2003 shows no significant
extension (Chadwell et al., 1999; Hildebrand et al., 1999; Chadwick and Stapp, 2002;
W. W. Chadwick, personal communication, 2006) and seismic monitoring shows no
activity. This suggests the crust between 0.5 and 25 km off-axis accommodates 26 mm
of aseismic deformation each year through some combination of near-axis fault motion
and elastic strain accumulation.
Citation: Chadwell, C. D., and F. N. Spiess (2008), Plate motion at the ridge-transform boundary of the south Cleft segment of the
Juan de Fuca Ridge from GPS-Acoustic data, J. Geophys. Res., 113, B04415, doi:10.1029/2007JB004936.
1. Introduction
[2] The Juan de Fuca plate provides an easily accessible
laboratory for studies of plate motion, creation, and subduction. This paper provides geodetic observations relevant
to understanding the behavior of crust newly formed at the
southern Cleft segment of the Juan de Fuca Ridge (JdFR,
Figure 1). This segment is a simple linear feature extending
about 50 km north from the Ridge intersection with the
Blanco Transform. Southern Cleft resembles typical intermediate rate spreading centers with a km-wide axial valley
and flanking ridges, and a long-term full-spreading rate of
approximately 52 mm/a as determined from geomagnetic
anomalies [Wilson, 1993]. The ridge crest has been an area
of study for over two decades [e.g., Kappel and Ryan, 1986;
Brett, 1987; Delaney et al., 1981; Embley et al., 1994;
Canales et al., 2005]. It has also been a site for new seafloor
geodetic tools [Morton et al., 1994; Chadwell et al., 1999;
Chadwick and Stapp, 2002].
1
Marine Physical Lab, Scripps Institution of Oceanography, University
of California, San Diego, La Jolla, California, USA.
2
Deceased 8 September 2006.
Copyright 2008 by the American Geophysical Union.
0148-0227/08/2007JB004936$09.00
[3] In the early 1990s, the U.S. Geological Survey
installed an acoustic ranging system to measure horizontal
deformation [Morton et al., 1994] across the 1-km-wide
axial valley at the south Cleft segment (44°40 0N,
130°200W). Chadwell et al. [1999] and Hildebrand et al.
[1999] used this system to measure horizontal motion from
1994– 1999. They measured the motion to be 3 ± 5 mm/a,
which implies no significant extension across the axial
valley floor. In July 2000, researchers from Oregon State
University established an array to include the valley floor
and partial valley wall 100 m south of the USGS array.
From 2000 – 2001, the displacement was 0 ± 20 mm
[Chadwick and Stapp, 2002]. Recently updated through
June 2003, preliminary analysis indicates no motion with
an at-most uncertainty of ±10 mm [W. W. Chadwick,
personal communication, 2006]. These results suggest
spreading occurs episodically within the axial valley walls.
[4] By contrast, far to the east of the Ridge, the JdF plate
is moving continuously at the average rate determined
geologically. This has been measured directly by one
globally referenced seafloor geodetic station at 48°100N,
127°100W [Spiess et al., 1998] that from 1994 to 1996
observed convergence between the JdF and North America
plates that agrees with the geologic rate within the measurement error. Convergence is also implied from contrac-
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Figure 1. The Juan de Fuca Ridge and Blanco Transform boundaries separating the Juan de Fuca and
Pacific plates. The GPSA site was measured with GPS and acoustics to determine its present-day velocity
both uncorrected (black) and corrected (red) for transient motions generated by boundary events. The
boundary events are the mid 1980s dike [Chadwick et al., 1991], the Blanco Transform event of 2 June
2000 (white star) [Dziak et al., 2003], and the 16 January 2003 event (white star) [D. Bohnenstiehl,
personal communication, 2006]. Thick black lines show rupture lengths of these events as modeled in
this paper. Acoustically monitored seismic activity is shown from 2 –4 June 2000 (solid gray circles) and
from 4 June 2000 through 4 May 2002 (solid black circles), the present end of available SOSUS
earthquake locations Fox et al. [1995]. Geologically predicted velocities of the Juan de Fuca plate relative
to the Pacific plate are plotted at the GPSA site for Wilson [1993] 0 – 0.78 Ma (blue) and 0 – 3.075 Ma
(green) Euler poles, showing agreement with the observed present-day motions. Also shown are the
USGS Tripods [Chadwell et al., 1999; Hildebrand et al., 1999] and OSU Extensometers [Chadwick and
Stapp, 2002; W. W. Chadwick, personal communication, 2006] at the JdFR; these detected no extension
across the axial valley floor. Inset shows general tectonic setting, location of shore GPS stations, and
coverage of detailed map. Bathymetry from the RIDGE Multibeam Synthesis Project.
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Figure 2. Plate motion at extensional (a) and strike-slip (b) boundaries occurs episodically near the
boundary, is continuous in the far-field, and is intermediate in between. At a ridge-transform boundary
(c) these regions overlap. Relative to the fault (d), no-motion is punctuated by rapid slip events in the
episodic region while in the far-field plate motion is steady at the half-rate. Within the intermediate
region, motion is steady state late into the post-event cycle, but can exceed the steady state rate following
a transient slip event.
tion and uplift measured along the coast above the Cascadia
subduction zone [Ando and Balazs, 1979; Savage et al.,
1981, 1991, 2000; Dragert et al., 1994; Mitchell et al.,
1994; Dragert and Hyndman, 1995; McCaffrey et al., 2000;
Murray and Lisowski, 2000; Miller et al., 2001; Svarc et al.,
2002].
[5] Globally, the transition from episodic to full, continuous motion at divergent plate boundaries remains largely
unobserved with geodetic techniques. Land-based systems
are limited to the two exposures, Iceland and Afar, both of
which confirm that spreading at the axis is episodic [Stein et
al., 1991]. Direct geodetic measurements in Northeast Iceland of crustal response to a 1975– 85 episode of seismicity
in the Krafla volcanic system have been modeled [Foulger
et al., 1992; Heki et al., 1993] to infer crustal properties but
in an environment quite different from typical sub-oceanic
regions.
[6] The developing combination of seafloor geodetic
techniques now makes the relevant oceanic observations
possible. As a step in this direction, in June 2000 we
installed a system on the Ridge flank, 25 km to the east
of the Juan de Fuca Ridge, to monitor seafloor motion in a
global frame using GPS and acoustic measurements [Spiess
et al., 2000].
[7] We use the GPSA-measured plate motion in an
attempt to find the transition between episodic, intermediate, and continuous motion. Off-axis, spreading ridges are
characterized kinematically, transitioning from episodic to
intermediate to continuous motion (Figure 2). The episodic
region is where plate creation occurs intermittently with
dike intrusions separated by long spans of no motion. The
continuous region is where the crust acts as a coherent
lithospheric unit that moves at a constant velocity driven by
plate-scale forces. Between these two is the intermediate
region where the interplay of magmatic and tectonic processes moves the new crust at varying rates as it coalesces
into a rigid plate. We attempt to find the transition by
measuring the present-day velocity and comparing it to a
prediction of full-rate motion after accounting for transient
displacements from boundary slip events.
2. GPS-Acoustic Measurements
[8] The GPS-Acoustic (GPSA) approach (Figure 3)
extends GPS positioning for crustal motion studies to the
seafloor. It combines GPS with acoustic ranging to measure
the position of seafloor transponders with centimeter-level
resolution in the same global reference frame as land-based
GPS sites [Spiess, 1985; Spiess et al., 1998]. The seafloor
array can be 100s of km from shore allowing geodetic
measurements of plate motion across the seafloor/continental interface or between widely separated seafloor points.
[9] GPS determines the precise location of a platform
(ship or buoy) on the sea surface, while underwater acoustic
ranging measures the distance to the seafloor array. Acoustic signals are needed because electromagnetic energy, on
which GPS is based, does not propagate significantly in
seawater. The basic underwater measurement is the time-offlight of an acoustic pulse from the ship to a seafloor unit
and back to the ship and the speed at which the acoustic
signal travels in seawater (sound speed). From these two
measurements the geometric range can be calculated. The
time-of-flight can be measured to ±3 microseconds (equivalent to 2 mm of range) using a variety of techniques [e.g.,
Spiess et al., 1997]. The main challenge is accommodating
changes in sound speed particularly in the upper ocean
where oceanographic forces drive variability that is significant in both space and time.
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Figure 3. The GPS-Acoustic approach determines the horizontal position of the seafloor array
~ with a
[horizontal components of (~
A)] by combining GPS positioning of shipboard antennas (D)
~
~
shipboard survey among antennas and hydrophone (C) with acoustic ranging (E) to seafloor transponders
whose relative positions (~
B) are known. Maintaining the ship near the array center assures that acoustic
velocity variations are primarily a function of depth and do not bias the horizontal components of ~
A.
[10] To date, there is no practical method to sample the
sound speed profile with sufficient temporal and spatial
resolution to account directly for changes in sound speed.
The horizontal stratification of sound speed, however, can
be exploited to mitigate its effect on positioning resolution
in the following manner. Three or four precision transponders are deployed on the seafloor to form an equilateral
triangle or square inscribed in a circle with the radius of the
nominal water depth (Figure 3). By maintaining the ship
near the center of the array (10 m), the vertical (launch)
angle from the shipboard transducer to each transponder can
be made equal, forcing the acoustic signals to spend the
same amount of time within each horizontal layer. As sound
speed changes in the upper ocean, all rays lengthen and
shorten equally. Because the transponders are evenly spaced
around the circumference of the inscribing circle with the
ship at the center, the coherent lengthening and shortening
of ranges is balanced in the horizontal. The upper ocean
sound speed variability will appear to move the seafloor
array vertically, but will not bias the horizontal position
estimate.
[11] To implement this approach we maintain the ship at
the array center and collect several tens-of-hours of continuous GPS and acoustic data. Traveltimes from the ship are
measured to seafloor transponders and back, and converted to
geometric range by ray-tracing through the mean sound
speed profile. To estimate the mean sound speed we repeatedly sample the ocean with a conductivity-temperature and
depth (CTD) device cycled from the surface to the seafloor.
These casts are averaged to provide the background profile
that includes the lower order components of the sound speed
field. These casts cannot provide the temporal and spatial
resolution to model sound speed on the scale of each acoustic
interrogation. The un-sampled sound speed variability is
mitigated by exploiting the horizontal stratification. With
4– 5 days of continuous data, the horizontal position of the
seafloor array can be determined with at least centimeterlevel repeatability in the global reference frame [Gagnon et
al., 2005].
[12] The GPSA approach relies on a ship (or buoy) to
provide the interface between the GPS and acoustic systems. Specifically, the shipboard configuration includes
three GPS antennas mounted on the ship to form a triangle
with as large fore-and-aft and athwart-ship dimensions as
are practical. Dual-frequency GPS carrier phase data are
sampled at 1 Hz at the ship and on shore to provide the
second-by-second positions of the shipboard GPS antennas.
A hydrophone is mounted within a hollow, vertical tube that
passes from the work decks through the bottom of the ship’s
hull. The hydrophone extends less than a meter below the
hull and is held rigidly in place against the sides of the tube.
The back of the hydrophone is at the bottom of the open,
air-filled tube that extends up to the work deck from where
it remains visible. Corner cube reflectors are mounted below
each GPS antenna and above the back of the hydrophone. A
surveying instrument is placed overtop of the tube to
measure the distances and angles in two perpendicular
planes between all reflectors. These data give the offsets
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between the antennas and hydrophone. With these offsets,
GPS positions at the antennas can be transferred to the
hydrophone giving the global position of the hydrophone on
a second-by-second basis.
3. Site Selection
[13] Site selection was influenced by a variety of factors.
Although the Iceland data implied that the relaxation zone
might be 100 km wide [e.g., Foulger et al., 1992], this
seemed unlikely for faster spreading, thus warmer oceanic
crust. At the time of installation, there were reports that
pressure transients due to seismic activity had been observed downhole in northern Juan de Fuca CORK installations as much as 50 km from the ridge crest (Davis et al.,
2001, Earl Davis, personal communication, 2000). To be in
the transition zone or determine its outer limit, it would be
desirable to be high up on the ridge flank. As noted above,
the transponder array must be on the order of the water
depth in radius, which would mean a footprint of about 5 km
in this area. Since crustal deformation within the array
would present problems of interpretation, it would be
appropriate to site the installation a distance off axis such
that the footprint is small compared with the distance from
the axis. This consideration, combined with the nature of the
topography, led to a site 25 km from the ridge crest. The
along-strike location was chosen to minimize edge effects
from the Blanco Trough to the south, and Axial Seamount
to the north as well as to be related to geodetic installations
at the ridge crest [Morton et al., 1994; Chadwick and Stapp,
2002]. The result was selection of the site at 44°430N,
130°030W at a nominal depth of 2900 m (Figure 1).
4. Data and Results
[14] In August 2000, the four transponders comprising
the array were installed to form a square with sides of
approximately 4 km (Figure 3). A total of 98 h of simultaneous GPS and acoustic ranging data were collected from
the center of the array while CTD casts were conducted
concurrently. Return visits provided contemporaneous GPS,
acoustic, and CTD data for 87 and 83 h in May 2001 and
June 2002, respectively. In September 2003, one of the four
seafloor transponders had failed and was replaced with a
new transponder. The new transponder position was
referenced to the old one to maintain the continuity of the
time series.
[15] This was done by temporarily placing an additional
active transponder adjacent (1– 2 m) to the inactive and
replacement transponders. By moving the GPS-positioned
ship in a 2-km radius circle while simultaneously ranging
on the two active units their relative position was determined aligned with Earth-Centered-Earth-Fixed (ECEF)
frame. An optical survey device was then deployed to the
seafloor and maneuvered to within 1 – 2 m of the units to
measure geometric ranges among the three units. This was
repeated from several locations around the transponder
cluster to determine the position offset between the inactive
and replacement unit. This offset was then rotated into the
ECEF giving the location of the new transponder relative to
the old transponder in the global frame (see Gagnon and
Chadwell [2007] for details).
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[16] Then, the temporary transponder was recovered and
a total 14 h of GPS, acoustic, and CTD data were collected.
During each campaign, 1-Hz GPS data were collected at
three stations (CHZZ, TPW2, and NEWP) along the Oregon
coast (Figure 1, inset).
[17] The shipboard and shore GPS data were processed
with NASA Jet Propulsion Laboratory’s GIPSY OASIS-II
software [Webb and Zumberge, 1997] using analysis described by Spiess et al. [1998], Chadwell and Bock [2001].
In all years, second-by-second repeatability of the GPS
antenna positions is 10– 20 mm in the horizontal [Miura et
al., 2002] and 20– 30 mm in the vertical Chadwell and Bock
[2001]. In each year, the shipboard optical survey data were
reduced to connect the GPS antenna phase centers to
the acoustic hydrophone phase center with a precision of
2– 3 mm [Chadwell, 2003]. GPS antenna positions were
transferred to the hydrophone, providing 20– 30 mm level
positions of the shipboard hydrophone on a second-bysecond basis. Finally, these are combined with the traveltimes
and mean sound speed profiles to estimate the location of the
seafloor array in the International Terrestrial Reference
Frame 2000 (ITRF2000) [Altamimi et al., 2002] at the epoch
of each campaign (Figure 4).
[18] The 1-s uncertainties for each epoch are given in
Table 1 and shown as error bars in Figure 4. To calculate
the positional uncertainties, the formal error estimates
from GIPSY are multiplied by 3 to account for the
well-known underestimation of the formal error estimates
from GIPSY [Larson et al., 1997]. Then, the scaled GPS
position uncertainties are propagated to the hydrophone
through the transformation equation that includes the
uncertainties of the surveyed antenna-hydrophone offsets.
Next, the time series of hydrophone position uncertainties
is propagated with the traveltime uncertainties through a
least squares estimator of the array position uncertainty.
In this calculation sound speed was fixed to the average
of all CTD profiles collected during the cruise. The
departure of the instantaneous sound speed from the
mean causes scatter in the traveltime residuals [Spiess et
al., 1998]. This scatter increases the reduced Chi-square
to 5– 10. The propagated array position uncertainty is
scaled by this factor as is required by least squares
estimation theory.
[19] The east and north position 1-s uncertainties range
from ±4 to ±6 mm from 2000 to 2002, or about an order
of magnitude more than what might be expected from a
land-based site using continuous GPS tracking. Undoubtedly, there is more to learn about the error budget of
GPSA positioning; however, we omit discussions of more
sophisticated error components [e.g., Mao et al., 1997]
and instead limit our analysis to that more consistent with
the assessments of the first applications of GPS for
crustal deformation measurements [e.g., Davis et al.,
1989]. In 2003, the uncertainty increased to ±18 –22 mm
due to a shorter data span (14 versus 80 h, with ±10 mm) and
additional uncertainty (±17 – 20 mm) from registering a
replacement transponder.
4.1. GPSA Velocity Relative to ITRF2000
[20] The estimated east and north positions were weighted by the inverse square of their 1-s uncertainty and linear
fits made to estimate the velocity of the seafloor site in the
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Figure 4. East and north position estimates (filled circles) in the ITRF2000 frame estimated from the
GPS-Acoustic solutions. The positions are shown with their 1-s uncertainty in mm beneath the error bar.
Solid line depicts the weighted linear fits for the velocity (see text for discussion). Open circles show the
2003 position estimate which is based on only 14 h of data (±10 mm) and contains additional uncertainty
from replacing an inactive seafloor transponder (±17 – 20 mm). Vertical lines show epochs of the 2 June
2000 and 16 January 2003 earthquakes along the Blanco Transform. The epoch for the mid-1980s event
along north Cleft segment is not shown.
Figure 5. (a) Compressional velocity profile from McDonald et al. [1994] at the Cleft segment
combined with those from Cudrak and Clowes [1993] and Barclay and Wilcock [2004] from the
Endeavour segment of the JdFR. The shear wave velocity profile is also from Barclay and Wilcock
[2004]. (b) The bulk modulus (k) and shear modulus (m) of the crust calculated in 2-km-thick layers from
the profiles from the seafloor to the base of elastic layer. Below this is the visco-elastic half-space with a
constant shear modulus 50 GPa and a bulk modulus of 150 GPa.
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Table 1. Model Estimates of Displacement at GPSA Site From Ridge and Transform Events
Observed
Measurement Epoch
East, mm
2000.5914
2001.3915
2002.4490
2003.6849
East rate (mm/yr) from
weighted fit North (mm)
2000.5914
2001.3915
2002.4490
2003.6849
North rate (mm/yr) from
weighted fit
Correlation coefficient(r)
Reduced Chi-square(c2)
GPSA-ITRF2000
0.0 ± 4.2
11.7 ± 4.9
45.4 ± 5.8
42.4 ± 22.1
Corrected Solutionsa
Model Displacement
0
0
0
1
Ext. 85 SS 00
0.0
0.8
2.5
4.0
0.0
1.8
3.5
5.1
0
2
0
3
SS 00
SS 00
0.0
0.7
1.4
2.0
0.0
1.7
3.7
5.7
0
1
SS 03 GPSA-ITRF2000
0.0
0.0
0.0
b
5.9
GPSA-ITRF20002 GPSA-ITRF20003
0.0 ± 4.2
12.7 ± 4.9
46.4 ± 5.8
37.6 ± 22.1
0.0 ± 4.2
11.6 ± 4.9
44.3 ± 5.8
34.6 ± 22.1
0.0 ± 4.2
9.2 ± 4.9
39.2 ± 5.8
26.8 ± 22.1
22.0 ± 3.8
20.8 ± 4.0
18.2 ± 3.9
0.0 ± 3.7
12.0 ± 4.4
34.8 ± 5.3
62.4 ± 19.4
0.0 ± 3.7
2.1 ± 4.4
14.1 ± 5.3
32.0 ± 19.4
0.0 ± 3.7
5.2 ± 4.4
20.8 ± 5.3
42.0 ± 19.4
4.5 ± 3.3
19.6 ± 3.4
8.7 ± 3.6
12.2 ± 3.5
0.0369
1.03
0.0369
1.15
0.0369
1.31
0.0369
1.26
22.0 ± 3.6
0.0 ± 3.7
1.0 ± 4.4
7.4 ± 5.3
13.6 ± 19.4
0.0
0.5
1.3
2.2
0.0
12.5
26.1
38.5
0.0
2.6
5.4
8.0
0.0
5.7
12.1
17.9
0.0
0.0
0.0
b
8.1
a
GPSA-ITRF2000i = GPSA-ITRF20000 + Ext. 085 + SS003 + SS000i, where i = 1, 2, 3 corresponds to models of the 2 June 2000 event. GPSAITRF20000 is observed GPSA motion in ITRF2000 with no correction for transient deformation. Model SS0001 from Dziak et al. [2003]; length(l) = 44 km,
strike(st) = 128°, dip(di) = 73°, depth(de) = 12.6 km, slip(sl) = 60 cm. Model SS0002 from Dziak et al. [2003] with scaling law applied to M0; l = 22.9 km, st
= 128°, di = 73°, de = 12.6 km, sl = 19 cm. Model SS0003 from Dziak et al. [2003] with strike aligned with transform; l = 44 km, st =115°, di = 73°, de =
12.6 km, sl = 60 cm.
b
Contains both the co-seismic and post-seismic components.
ITRF2000 frame as 22.0 ± 3.6 mm/a east and 4.5 ± 3.3 mm/
a north with a 0.04 correlation coefficient (Table 1). For
the current solutions, at the 1-s uncertainty level (67%) all
east and north components lie along the linear fit within
their positional uncertainty.
[21] Spiess et al. [1998] reported the 1-s positional
uncertainty of ±39 mm east and ±8 mm north from 3
surveys conducted in 1994, 1995, and 1996 with 23.6,
32.0, and 33.5 h of GPSA data, respectively. At that time
in the mid-90s, a combination of self-generated ship noise
and poor acoustic signal recognition and processing, meant
that only about 10% of the interrogation pings resulted in
usable traveltimes. In addition, the ship used to collect the
data was manually steered and could only hold station to
within 150 m of the array center rather than the preferred
10-m level possible with a dynamically positioned ship.
Here, 99% of all interrogations result in a valid traveltime.
These factors along with better modeling of GPSA data and
increased time on station have improved the precision from
±39 mm in the mid-90s to ±4– 6 mm.
4.2. GPSA Velocity Relative to Pacific Plate
[22] Beavan et al. [2002] measured the motion of the
interior of the PA plate with present-day geodetic techniques. This is not possible for the Juan de Fuca plate because
it has no exposed landmass for conventional geodetic
stations from which a present-day Euler pole can be
calculated. The present-day motion of the Juan de FucaPacific plates must be inferred from the geomagnetic
anomaly pattern until such time when at least three GPSA
sites are established within the interior of the JdF plate.
[23] For most of the Earth’s major tectonic plates,
present-day Euler poles have been determined from
space-based geodesy through tracking the motion of the
plates over a few years, primarily with GPS [e.g., Altamimi
et al., 2002; Sella et al., 2002]. Independently, present-day
Euler poles have been determined by averaging over the
finite rotation from the present back to geomagnetic
anomaly 2A (3.075 Ma) [e.g., DeMets et al., 1994]. Rather
consistently for most plates, the space-based velocities
measured at the stable interior of the plate and the
geomagnetic-based velocities agree within their respective
measurement uncertainties.
[24] It is, therefore, quite natural to propose the anomaly
2A pattern to describe the present-day relative motion of
the JdF-PA plates, just as the NUVEL-1 and NUVEL-1A
JdF-PA Euler poles are constructed [DeMets et al., 1990,
1994]. However, Wilson [1993] cautions that there may be
applications where it is important to recognize changes in
the pole since 3.075 Ma. Such pole changes can result from
reorientations of the JdF plate in response to interactions
with the larger PA and North America (NA) plates. More
recently investigators have adopted the total reconstruction
pole from 0 – 0.78 Ma of Wilson [1993] to calculate the
mean rate and instantaneous velocity of JdF-PA motion. We
also adopt this to define the rigid-body motion of the JdFPA plates and to compare with the observed GPSA-PA
velocity, though we note the velocities implied by these two
poles do not differ significantly for the purposes of this
study (Figure 1).
[25] Next, we transform the GPSA velocity at the 25-km
site from ITRF2000 to a velocity relative to the Pacific
plate. At this site, the Pacific-ITRF2000 velocity is 36.6 ±
0.2 mm/a east and +29.1 ± 0.6 mm/a north with a correlation of 0.34 Beavan et al. [2002]. The GPSA-PA vector
for the observed motion is compared to the velocity calculated from the 0– 0.78 Ma Euler pole of Wilson [1993] in
Figure 1. Upon inspection of Figure 1 the 95% error ellipses
of GPSA-PA and that from the Wilson [1993] 0 – 0.78 Ma
Euler pole do not overlap, indicating the possible importance of modeling transient motions.
5. Transient Motions
[26] Large transient motions that persist for years have
been observed by direct geodetic measurements in northern
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Figure 6
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Iceland with the 1975– 1981 Krafla rifting episode between
the Eurasian and North American plates [Foulger et al.,
1992; Heki et al., 1993; Hofton and Foulger, 1996a, 1996b;
Pollitz, 1996]. Here, a dike intrusion of width 2 m
occurred at the Rift. Nearly 10 years following the event,
velocities were still more than twice the full-spreading rate
within 50 km of the Rift. Similarly, co-seismic slip followed
by accelerated motion exceeding 2 – 3 times steady state has
been observed with major strike-slip earthquakes along and
near the San Andreas Fault in 1906 [Thatcher, 1974], 1989
[Bürgmann et al., 1997], 1992 [Pollitz et al., 2000], and
1999 [Pollitz et al., 2001].
[27] One model of these phenomena begins with crust
composed of an elastic layer overlying a viscoelastic
halfspace. A slip event within the elastic layer rapidly
releases accumulated elastic strain displacing the elastic
layer. This causes co-seismic motion. The newly displaced
upper layer transfers stress across the elastic/viscoelastic
boundary, which more slowly diffuses to relieve the
boundary stress. The relaxation of the viscoelastic halfspace further displaces the surface of the elastic layer
resulting in post-seismic motion [Elsasser, 1969; Bott
and Dean, 1973].
[28] Therefore slip events at either the JdF Ridge or the
Blanco Transform may contribute to motion observed at the
GPSA site. We must calculate the motion induced by elastic
and viscoelastic response of the JdF plate to recorded slip
events at the Ridge and Transform. We begin by reviewing
the boundary events.
5.1. Cleft Extensional Events
[29] Between 1983 and 1987, a volcanic event occurred
along a 30 km-section of the northern Cleft segment of the
Juan de Fuca Ridge [Chadwick et al., 1991; Embley et al.,
1991]. Evidence of an event was first detected in 1986 by
observance of a hydrothermal plume in the water column
over the north Cleft [Baker et al., 1989]. The exact dates are
unknown because the event(s) occurred prior to scientific
monitoring of the U.S. Navy’s SOund SUrveillance System
(SOSUS) and were not recorded by seismic arrays on land,
which have a detection threshold of mb > 4.0 in this region.
Comparisons between bathymetry collected from ship and
subsequent deep-towed platforms constrain the eruption
dates between 1983 and 1987. A quantitative differencing
of the bathymetric data revealed an eruptive volume of
0.05 km3 [Fox et al., 1992]. Using geologic and hydrothermal observations, Embley and Chadwick [1994] have proposed an intrusion/eruption episode that likely involved two
separate events separated by at least 7 months. They suggest
two lateral dikes that in total extended from 44°530N to
45°100N.
[30] There is no evidence for recent extensional events in
the southern Cleft segment. Normark et al. [1983] mapped
the lava plain covering the axial valley floor at south Cleft
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and estimate the age to be less than a few hundred years.
Geodetic monitoring did not detect extension across the
axial valley from 1994 – 1999 [Chadwell et al., 1999;
Hildebrand et al., 1999] and from July 2000 through June
2003 [Chadwick and Stapp, 2002; W. W. Chadwick, personal
communication, 2006]. SOSUS monitoring since 1994 has
detected no significant activity in the Cleft region.
5.2. Blanco Transform Strike-Slip Events
[31] On 2 June 2000 at 1113Z, a Mw 6.2 main shock
occurred at the east side of a foreshock cluster that initiated
16 h earlier centered at 44°210N to 130°150W. During the
following 33 h aftershocks extended from the main shock
location along a 44-km-length section of the Transform
Dziak et al. [2003]. The National Earthquake Information
Center (NEIC) derived moment tensor suggested an event
that was 12 km deep, mostly right-lateral in motion, striking
128°, and dipping 73°.
[32] On 16 January 2003 at 0053Z, a Mw 6.3 main shock
occurred at 44°170N to 129°010W. The NEIC estimated a
right lateral motion with 17 km depth, strike of 127°, and
dip of 83°. D. Bohnenstiehl, personal communication
[2006] characterized this event with a 9 km depth based
upon a 600 °C isotherm from a simple thermal model,
30 km rupture length, and 0.25 m slip.
5.3. Modeling and Results
[33] We use the crustal structure and fault models of the
events in an elastic [Okada, 1992; Gomberg and Ellis,
1994] and a viscoelastic model [Pollitz, 1992] to calculate
the response of the JdF plate to the transient events at the
boundaries.
[34] We construct a model of the crust in the vicinity of
the south Cleft segment from a composite of sources. The
elastic properties of the crust are calculated from the
compressional (Vp) and shear wave (Vs) velocities and
crustal density (r). We use the compressional velocity
profile from McDonald et al. [1994] at the Cleft segment
combined with those from Cudrak and Clowes [1993] and
Barclay and Wilcock [2004] from the Endeavour segment of
the JdFR. The shear wave velocity profile is also from
Barclay and Wilcock [2004]. Stevenson et al. [1994] used
seafloor and sea-surface gravity data to model the density
structure along the southern Juan de Fuca Ridge. They
estimate the average density to be 2630 ± 50 km m3 in the
upper 2 km of ocean crust. We adopted a density profile that
increases with depth with values of 2700 km m3 at the
seafloor increasing to 3100 km m3 12 km depth. The bulk
modulus (k) and shear modulus (m) of the crust are
calculated in 2-km-thick layers from these profiles from
the seafloor to the base of elastic layer (Figure 5). Below
this is the visco-elastic half-space with a constant shear
modulus 50 GPa, a bulk modulus of 150 GPa and a
viscosity of 3.0 1018 GPaS.
Figure 6. Total model calculated displacements from 2 August 2000 through 5 September 2003 of the Juan de Fuca plate
due to transient events at the boundaries. (a) Viscoelastic displacement from dike at north Cleft segment in 1985,
(b) Viscoelastic displacement from 2 June 2000 earthquake using aftershock distribution for rupture length, (c) using
rupture length from scaling law applied to moment magnitude, (d) using orientation aligned with strike of transform. Elastic
(e) and viscoelastic (f) response to the 16 January 2003 event. Gray arrows show total displacement at grid points, black
arrow at GPSA site. Insets show total east and north displacement at the GPSA site for each model.
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Figure 7. East and north position estimates (solid symbols) in the ITRF2000 frame estimated by
applying the calculated transient motions to the GPS-Acoustic positions. Again, the values are shown
with their 1-s uncertainty error bars along with weighted linear fits (solid lines) for the velocity (Table 1).
Open symbols show the 2003 position estimate which is based on 14 h of data (±10 mm) and contains
additional uncertainty (±17 – 20 mm) from replacing an inactive seafloor transponder. Vertical lines show
epochs of the 2 June 2000 and 16 January 2003 earthquakes along the Blanco Transform. The epoch for
the mid-1980s event along north Cleft segment is not shown.
[35] Constraints upon the input parameters controlling the
elastic and viscoelastic models are weak. The elastic layer
depth can reasonably range from 6 km, as suggested by
seismic imaging [Canales et al., 2005], to 12 km implied by
the depth of the strike-slip events. We calculate displacements using 6 km and 12 km elastic layer thicknesses and
show results from the thicker model.
[36] For the model calculations of the 1985 extensional
event, we assumed a single lateral dike intrusion that
extended the entire 30 km, that the dike was just below
the surface, and ruptured the entire layer 2C with a width of
2 m. The total displacement since 1985 is approximately
50 mm east and 15 mm south (Figure 6a, inset); however,
over the span of the survey the displacement is less than
5 mm (Ext085 in Figure 6a and Table 1). The deformation
transient from the 1980s intrusion has decayed down to near
background levels.
[37] The 2 June 2000 event, though perhaps a smaller
equivalent moment magnitude than the 1980s event, occurred just two months prior to the establishment of the
GPSA site and may have induced a significant motion.
Dziak et al. [2003] interpreted the aftershock distribution to
be the rupture length, and applied a scaling law [Scholz,
1982] to give a slip of 60 cm (SS0001 in Figure 6b and
Table 1). We note that calculation of the moment magnitude
with 44 km length, 12 km depth, and 60 cm slip exceeds the
seismically observed moment, though this is not significant
for the application by Dziak et al. [2003]. A possible
interpretation is that the main shock is the seismically
observed component, while the aftershock pattern during
the following 33 h delineates the likely extent of aseismic
slip, recognizing that transform faults can have large aseismic components, i.e., low seismic coupling [Boettcher and
Jordan, 2004].
[38] To gauge the sensitivity of the model to this interpretation, we calculate deformation for two additional slip
models. First, rather than using the aftershock cluster as the
length, we take the seismically observed moment magnitude
and apply the scaling law of Scholz [1982] to get a rupture
length of 22.9 km and a slip of 19 cm (SS0002 in Figure 6c
and Table 1). Second, we note that though the NEIC
estimated the event with a strike of 128°, the distribution
of the aftershocks allows the interpretation that the strike of
the slip be aligned with the bathymetric trace of the
Transform at 115°. Here we again use the aftershocks for
the rupture length, apply the scaling law for the slip
magnitude, but take the strike to be 115° (SS0003 in
Figure 6d and Table 1).
[39] As expected, the orientation of the displacement is
controlled by the strike of the slip, its magnitude, and the
rupture length. Both models (SS0001 and SS0002) with strike
of 128° give southerly displacements with a small westerly
component (Figures 6b and 6c; and Table 1), though the
total displacements are 38.8 and 8.2 mm for rupture lengths
and slips of 44 km, 60 cm and 23 km, 19 cm, respectively.
With the strike of the slip aligned with the general orientation of the Blanco Transform (115°), the motion (SS0003) is
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Figure 8. Velocity vectors relative to the PA plate along with their 95% confidence ellipses for the
Wilson [1993] 0 –0.78 Ma Euler pole (gray) and the GPSA observations (black) calculated for: (a) with
no deformation correction, (b) with correction for motion implied by a slip model from Dziak et al.
[2003] for 2 June 2000 event, (c) with slip model from Dziak et al. [2003] with scaling law applied to M0,
(d) with slip model from Dziak et al. [2003] with strike aligned with the Transform.
south-southeast with a total displacement of 18.8 mm
(Figure 6d and Table 1).
[40] Finally, the 16 January. 2003 event was modeled
with a 9 km depth, 30 km length, and 0.25 m slip. This
event has both co-seismic and post-seismic components
(SS003) that are generally southeasterly with a total displacement of 10 mm (Figures 6e, 6f, and Table 1).
[41] Constraints upon the input parameters controlling the
viscoelastic computations are weak. For example, the elastic
layer depth was chosen at 12 km, but could reasonably vary
from 6 – 12 km depth. Likewise, the slip event dimensions
(depth, rupture length, and fault slip) are not strongly
determined. To study these effects, solutions were repeated
with a 6 km elastic layer depth. Shallower elastic layer
depths tended to reduce the calculated relaxation, but do not
change any conclusions presented later. Therefore we note
that though our choice of fault models and crustal properties
is not unique, nor exhaustive, they are an appropriate
representation of the perturbation probable from transient
events given the detail of our geodetic data.
[42] Four solutions are constructed to probe the influence
of transient motions upon the observed GPSA motion. The
first solution is simply the observed GPSA motion without
correcting for transient motion (GPSA-ITRF20000). The
first corrected solution (GPSA-ITRF20001) subtracts from
GPSA-ITRF20000 the displacements from the 1985 extensional event (Ext.085), the 16 January 2003 strike slip
event (SS003) and model 1 of the 2 June 2000 strike slip
event (SS0001). The next solution (GPSA-ITRF20002)
subtracts Ext.085, SS003, and SS0002. The third corrected
solution (GPSA-ITRF20003) subtracts Ext.085, SS003, and
SS0003. GPSA-ITRF20000 was given in Figure 4 and
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Table 2. Velocity (mm/a) of GPSA Site Relative to the Pacific, and Juan de Fuca Plates
GPSA-PAa = GPSA-ITRF2000a - PA-ITRF2000b
GPSA-PA
e
n
r
0
58.6 ± 3.6
24.6 ± 3.4
0.0392
GPSA-PA
1
58.7 ± 3.8
9.5 ± 3.4
0.0401
GPSA-PA
2
57.5 ± 4.0
20.4 ± 3.6
0.0396
GPSA-JdFa = GPSA-PAa - JdF-PAc
GPSA-PA
3
GPSA-JdF
54.8 ± 3.9
16.9 ± 3.6
0.0396
0
8.5 ± 3.7
9.7 ± 3.8
+0.0386
GPSA-JdF1
GPSA-JdF2
GPSA-JdF3
8.6 ± 3.9
5.4 ± 3.8
+0.0283
7.4 ± 4.1
5.5 ± 3.9
+0.0283
4.7 ± 4.0
2.0 ± 3.9
+0.0300
a
With 0, 1, 2, 3 representing solutions given in Table 1.
From Beavan et al. [2002] with e = 36.6 ± 0.2 mm/a, n = 29.1 ± 0.6 mm/a, r = 0.3453 for PA-ITRF2000.
c
From Wilson [1993] 0 – 0.78 Ma Euler pole with e = 50.1 ± 1.0 mm/a, n = 14.9 ± 1.6 mm/a, r = +0.6397 for JdF-PA.
b
Table 1, and repeated in Figure 7. GPSA-ITRF20001 – 3
positions, corrected for the transient effects, and weighted
linear fits estimated for each solution are given in Table 1
and Figure 7.
5.4. Transient Component of GPSA Velocity
[43] To determine the kinematic regime of the GPSA site
we compare the observed GPSA motion with and without
transient motion components to the average motion of the
rigid JdF-Pacific (PA) plates.
[44] The GPSA-PA vectors for observed motion with and
without transient motions removed are computed in Table 2
and shown in Figure 8. Upon inspection of Figure 8, two
deductions are immediately clear: (1) The GPSA-PA1 – 3
motions are in general agreement with JdF-PA motion
predicted by the Wilson [1993] 0 – 0.78 Ma Euler pole. (2)
The motions induced by the transient events at the boundaries, while perhaps significant, are modest, not exceeding
10% of the total motion at the GPSA site.
[45] Support for the first deduction is the overlap of the
95% error ellipses of the GPSA-PA1 – 3 velocity with that
calculated from the Wilson [1993] 0– 0.78 Ma Euler pole
(Figure 8). Also, the right-hand side of Table 2 gives the
GPSA-JdF motion, i.e., the motion of the site as measured
by GPS-Acoustics relative to the JdF plate itself as predicted by the geomagnetic anomaly at 0.78 Ma. If the site
was measured by GPSA to be moving as predicted by the
Wilson [1993] 0 – 0.78 Ma Euler pole, then the discrepancy
would be zero. The GPSA-JdF1 – 3 velocities are small and
Figure 9. (a) Location of episodic, intermediate, and continuous zones supported by seafloor geodetic
data. Plate velocities relative to the axis are plotted across strike of JdF Ridge at location of geodetic
experiments. Shown are the velocities for the USGS array Chadwell et al. [1999], the OSU extensometers
[Chadwick and Stapp, 2002; W. W. Chadwick, personal communication, 2006], the 0 – 0.78 Ma Euler
pole of Wilson [1993], and the GPSA-PA velocities without (GPSA-PA0) and with (GPSA-PA3)
correction for transient motions. No-motion condition observed by the on-axis geodetic experiments
defines the extent of episodic region (dark gray shading). Full 1/2 rate at GPSA site indicates continuous
zone (dark gray shading). Additional, future geodetic monitoring in the region from 0.5 to 25 km off-axis
could further constrain the width of each zone. (b) Shows spreading boundaries implied by geophysical
observations. Black line shows bathymetric profile across-strike. On-axis light gray region is approximate
limit of episodic activity from Carbotte et al. [2006] and Karson et al. [2002]. Curved-dashed line is
approximate seafloor profile. Inflection of the curvature from concave down to convex occurs between
5 – 15 km off-axis. Beyond begins continuous motion of the rigid plate (light gray area) [Buck et al.,
2005].
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only 1 or 2 times their uncertainty, a statistical indicator the
discrepancy is insignificant and the two estimates agree.
However, the uncorrected velocity, GPSA-JdF0, is 2 – 3
times its uncertainty and the 95% error ellipses of GPSAPA0 and that from the Wilson [1993] 0 – 0.78 Ma Euler pole
do not overlap, indicating the importance of modeling
transient motion.
[46] The second deduction is supported by the small
difference (10%) between GPSA-PA0 without correction
and GPSA-PA1 – 3 with corrections. This leads to the interpretation that the motion observed at the GPSA site is due
primarily to steady state conditions and not to the three
modeled transient effects. This is an important point given
results from Iceland and San Andreas have shown transient
motions can be 2 to 3 times the full-rate plate velocity.
[47] Examining GPSA-PA1 – 3 vectors shows the influence
of slip, strike, and rupture length (Figure 8). GPSA-PA1 and
GPSA-PA2 model the 2 June 2000 event as SS0001 and
SS0002, respectively, and show the difference between two
rupture lengths at the same strike. Because the aftershock
distribution clearly indicates a change in stress conditions
on the fault, the total slip may be under-represented by
SS0002, which limits application of the scaling law to the
observed seismic moment magnitude. Therefore the GPSAPA2 velocity may be incomplete. SS0001 represents a more
complete rupture length because it reflects the actual aftershock pattern, but the strike is not well constrained by the
seismic solution, forcing GPSA-PA1 to have a larger north
component. GPSA-PA3 best agrees with the velocity calculated from the Wilson [1993] 0 –0.78 Ma Euler pole. GPSAPA3 modeled the 2 June 2000 event defining the rupture
length by the aftershock pattern and the slip strike to be
parallel to the Transform, which is a long-term, persistent
plane of weakness.
6. Mid-Ocean Ridge Dynamics
[48] We next examine the implications of these results to
dynamics at mid-ocean ridges. We assume the GPSA-PA
velocity represents near steady state conditions as we have
shown the calculated transient effects are only 10% of the
full-rate velocity. Figure 9a shows a plot of the velocities
measured in the vicinity of the site including no significant
extension across the the 1-km-wide axial valley from 1994 –
1999 and 2000– 2003 [Chadwell et al., 1999; Hildebrand et
al., 1999; Chadwick and Stapp, 2002; W. W. Chadwick,
personal communication, 2006]. As discussed previously,
no motion under steady state conditions indicates that the
site lies within the episodic zone. We note that the total
extent of the on-axis arrays is 1 km thus the episodic zone
extends to at least 0.5 km to either side of the axis. The
continuous zone is defined by the velocity from the 0 – 0.78
Ma Euler pole Wilson [1993] and is plotted relative to the
axis (i.e., half-rate). To show the sensitivity of our data we
plot the half-rate velocity for both GPSA-PA0 and GPSAPA3 that bracket our results. Figure 9 shows that the GPSA
site is moving at the half-rate, constraining the continuous
region to begin at no more than 25 km off-axis.
[49] Given no motion within the axial valley floor and
continuous motion 25 km off-axis there is an intermediate
region between 0.5 and 25 km that accommodates 26 mm
deformation each year. This deformation must occur
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aseismically because no seismic activity above magnitude
1.8 was recorded by SOSUS during the survey span
anywhere between the Ridge and survey site [Fox et al.,
1995].
[50] We compare our results with suggested models of
plate creation at intermediate spreading rate ridges and the
JdFR specifically. Carbotte et al. [2006] suggest the width
affected by episodic magmatic processes is approximately
equal to the depth of the axial magma chamber that they
measure to be 2.0 km deep. In their model, the axial
volcanic ridge that lies within 2 km of the axis is potentially
still active and would be affected by a subsequent dike
injection. Beyond 2 km, the graben faults are no longer
affected by magmatic intrusions at the ridge [see also
Canales et al., 2005]. Karson et al. [2002] suggest that
extensive vertical subsidence must occur (likely episodically)
within the axial valley out to 1.5 km to accommodate
features observed in the exposed wall of the Western Blanco
Transform Fault. We note the extent of the seafloor
geodetic arrays on the axial valley floor is 0.5 km to
either side of the axial Cleft. Both Carbotte et al. [2006]
and Karson et al. [2002] would extend the region of episodic
activity to 1.5 – 2 km off-axis. The geodetic measurements
showing no motion are consistent with this interpretation
(Figure 9b).
[51] Moving farther off-axis beyond 2 km, the crust cools
increasing its density and pushing deeper the isotherm
defining the boundary between the lithosphere and asthenosphere. This thickens the elastic layer that at an unknown
distance off-axis is sufficient to act as a stress guide. In
addition, faulting and possible off-axis magmatic intrusion
forming the abyssal hills can be seen to begin as close as 5 km
off-axis. At some unknown distance off-axis these faults
become inactive and no longer accommodate motion. The
site 25 km off-axis is at the full-half rate and places some
constraint on these two processes. It suggests that the elastic
layer is sufficiently thick to be a stress guide and behave as
part of the rigid JdF plate and that to the east of the GPSA
site, i.e., toward the plate interior that the abyssal hill
faulting has predominately ceased. This is generally consistent with the Buck et al. [2005] unbending model that
predicts fault motion ceases where the curvature of the
plate flattens which occurs 20 km off-axis at this site
(Figure 9b).
7. Conclusions
[52] GPS-Acoustic positions of a seafloor array over 4
annual campaigns are consistent with a linear trend. Approximately 80 h of continuous GPS-Acoustic data allows a
precision of ±4– 6 mm, sufficient to address a number of
seafloor tectonic questions. Also, replacing a seafloor transponder is possible with at-most 10 –20 mm uncertainty,
demonstrating that seafloor geodetic arrays can be maintained for long-term (>5 years) time series measurements.
[53] The observed GPSA-PA velocity is consistent, at the
95% confidence level, with the velocities calculated from
the Wilson [1993] 0 – 0.78 (and 0 – 3.075) Ma Euler pole(s)
once transient motions are removed. Transient events at the
plate boundaries account for 10% of the total GPSA-PA
motion according to elastic/viscoelastic models. This suggests the GPS-PA velocity is due primarily to steady state
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plate dynamics. Assuming the geologically derived rate
represents full-rate JdF-PA spreading, the site 25 km east
of the Ridge is interpreted to be in a region of continuous
plate motion. This is a robust result of this study and holds
for GPSA-PA velocities both uncorrected and corrected for
transient motions. These are the first geodetic results to
directly constrain the width of active spreading to lie within
25 km of the axis at an intermediate spreading-rate ridge.
[54] The region between 0.5 and 25 km off-axis likely
accommodates 26 mm of aseismic deformation each year,
given the results of this study, a lack of observed seismicity,
and previously reported geodetic monitoring showing no
significant extension across the 1-km-wide axial valley.
This raises several questions of how the strain is accumulated. The geomorphology of the abyssal hills suggests the
faults accommodate deformation is some manner. We speculate that as the upper mantle creeps steadily from the ridge
and the elastic layer thickens, the upper crust accumulates
strain which is accommodated by motion along the faults.
The frequency of the fault motion is unknown. If this
motion occurred during the span of our measurements then
it was aseismic to the threshold of the SOSUS system,
which have a detection threshold of mb > 1.8 in this region
[Fox et al., 1995]. Alternatively, the faults may not have
displaced and instead strain accumulated in the near-axis
elastic crust. Eventually, this elastic strain will be released
and accommodated by fault motion perhaps generating a
seismically recorded event. The partitioning between fault
motion and elastic strain accumulation might be discerned
with additional near-axis geodetic studies.
[55] Acknowledgments. Herb Dragert, M. Meghan Miller, Andrew
Miner, Dan Johnson, Chris Goldfinger, Jason Chaytor, Chris Fox, Jonathan
Klay, Bill Chadwick and UNAVCO helped with collection of the 1-Hz GPS
data at shore stations. John Hildebrand aided with site selection and joined
the 2000 cruise. Captain and crew of the R/V Roger Revelle provided at-sea
support. Richard Zimmerman, Dave Jabson, Dennis Rimington and Dave
Price provided engineering support. We thank Katie Phillips, Katie Gagnon,
and Neil Kussat for at-sea help, discussions about data reduction, and
comments on this manuscript. We also thank Kelin Wang for comments on
a early version of this manuscript. Bob Dziak and Del Bohnenstiehl
provided data and insight on the Blanco Earthquakes. Bill Chadwick made
available recent extensometer results. Douglas Wilson and John Beaven
provided insightful reviews that much improved the paper. Bathymetry
from the RIDGE Multibeam Synthesis Project. Map was produced using
the Generic Mapping Tools package [Wessel and Smith, 1998]. This work
was supported by NSF grant OCE-9907247 from the Marine Geology and
Geophysics Program. Fred N. Spiess died on 8 September 2006 while this
paper was in final preparation.
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C. D. Chadwell and F. N. Spiess, Marine Physical Lab, Scripps
Institution of Oceanography, University of California, San Diego, 9500
Gilman Drive, La Jolla, CA 92093-0205, USA. ([email protected])
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