plate motion at the ridge-transform boundary of the south cleft

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-axisto 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 mdepth) with a repeatability of ±4–6 mm. Three transient events at the Juan de Fuca Ridgeand Blanco Transform account for �10% of this motion in viscoelastic modeling,suggesting that the observed GPSA-PA velocity is due primarily to steady state platedynamics. Subtracting the modeled transient motion gives a velocity of 57.3 ± 3.9 mm/a atS72.9�E ± 12.1� degrees (1-s), which is consistent at the 95% confidence level with thevelocity calculated from the Wilson (1993) 0–0.78 Ma Euler pole. Therefore this siteis interpreted to be in a region of continuous, full-rate plate motion, a robust result ofthis study which holds with and without correcting for transient motions. These resultsprovide direct geodetic evidence that spreading occurs predominantly within 25 km of theaxis at this intermediate spreading-rate ridge. Previously reported geodetic monitoringacross the 1-km-wide axial valley from 1994–1999 and 2000–2003 shows no significantextension (Chadwell et al., 1999; Hildebrand et al., 1999; Chadwick and Stapp, 2002;W. W. Chadwick, personal communication, 2006) and seismic monitoring shows noactivity. This suggests the crust between 0.5 and 25 km off-axis accommodates �26 mmof aseismic deformation each year through some combination of near-axis fault motionand 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 accessiblelaboratory for studies of plate motion, creation, and sub-duction. This paper provides geodetic observations relevantto understanding the behavior of crust newly formed at thesouthern Cleft segment of the Juan de Fuca Ridge (JdFR,Figure 1). This segment is a simple linear feature extendingabout 50 km north from the Ridge intersection with theBlanco Transform. Southern Cleft resembles typical inter-mediate rate spreading centers with a km-wide axial valleyand flanking ridges, and a long-term full-spreading rate ofapproximately 52 mm/a as determined from geomagneticanomalies [Wilson, 1993]. The ridge crest has been an areaof 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 seafloorgeodetic tools [Morton et al., 1994; Chadwell et al., 1999;Chadwick and Stapp, 2002].

[3] In the early 1990s, the U.S. Geological Surveyinstalled an acoustic ranging system to measure horizontaldeformation [Morton et al., 1994] across the 1-km-wideaxial valley at the south Cleft segment (44�400N,130�200W). Chadwell et al. [1999] and Hildebrand et al.[1999] used this system to measure horizontal motion from1994–1999. They measured the motion to be �3 ± 5 mm/a,which implies no significant extension across the axialvalley floor. In July 2000, researchers from Oregon StateUniversity established an array to include the valley floorand 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 throughJune 2003, preliminary analysis indicates no motion withan at-most uncertainty of ±10 mm [W. W. Chadwick,personal communication, 2006]. These results suggestspreading 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 determinedgeologically. This has been measured directly by oneglobally referenced seafloor geodetic station at 48�100N,127�100W [Spiess et al., 1998] that from 1994 to 1996observed convergence between the JdF and North Americaplates that agrees with the geologic rate within the mea-surement 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 andPacific plates. The GPSA site was measured with GPS and acoustics to determine its present-day velocityboth uncorrected (black) and corrected (red) for transient motions generated by boundary events. Theboundary events are the mid 1980s dike [Chadwick et al., 1991], the Blanco Transform event of 2 June2000 (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 inthis paper. Acoustically monitored seismic activity is shown from 2–4 June 2000 (solid gray circles) andfrom 4 June 2000 through 4 May 2002 (solid black circles), the present end of available SOSUSearthquake locations Fox et al. [1995]. Geologically predicted velocities of the Juan de Fuca plate relativeto 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 theUSGS Tripods [Chadwell et al., 1999; Hildebrand et al., 1999] and OSU Extensometers [Chadwick andStapp, 2002; W. W. Chadwick, personal communication, 2006] at the JdFR; these detected no extensionacross the axial valley floor. Inset shows general tectonic setting, location of shore GPS stations, andcoverage of detailed map. Bathymetry from the RIDGE Multibeam Synthesis Project.

B04415 CHADWELL AND SPIESS: MOTION AT RIDGE-TRANSFORM BOUNDARY

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tion and uplift measured along the coast above the Cascadiasubduction 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, contin-

uous motion at divergent plate boundaries remains largelyunobserved with geodetic techniques. Land-based systemsare limited to the two exposures, Iceland and Afar, both ofwhich confirm that spreading at the axis is episodic [Stein etal., 1991]. Direct geodetic measurements in Northeast Ice-land of crustal response to a 1975–85 episode of seismicityin the Krafla volcanic system have been modeled [Foulgeret al., 1992; Heki et al., 1993] to infer crustal properties butin an environment quite different from typical sub-oceanicregions.[6] The developing combination of seafloor geodetic

techniques now makes the relevant oceanic observationspossible. As a step in this direction, in June 2000 weinstalled a system on the Ridge flank, 25 km to the eastof the Juan de Fuca Ridge, to monitor seafloor motion in aglobal frame using GPS and acoustic measurements [Spiesset al., 2000].[7] We use the GPSA-measured plate motion in an

attempt to find the transition between episodic, intermedi-ate, and continuous motion. Off-axis, spreading ridges arecharacterized kinematically, transitioning from episodic tointermediate to continuous motion (Figure 2). The episodicregion is where plate creation occurs intermittently withdike intrusions separated by long spans of no motion. Thecontinuous region is where the crust acts as a coherentlithospheric unit that moves at a constant velocity driven byplate-scale forces. Between these two is the intermediate

region where the interplay of magmatic and tectonic pro-cesses moves the new crust at varying rates as it coalescesinto a rigid plate. We attempt to find the transition bymeasuring the present-day velocity and comparing it to aprediction of full-rate motion after accounting for transientdisplacements from boundary slip events.

2. GPS-Acoustic Measurements

[8] The GPS-Acoustic (GPSA) approach (Figure 3)extends GPS positioning for crustal motion studies to theseafloor. It combines GPS with acoustic ranging to measurethe position of seafloor transponders with centimeter-levelresolution in the same global reference frame as land-basedGPS sites [Spiess, 1985; Spiess et al., 1998]. The seafloorarray can be 100s of km from shore allowing geodeticmeasurements of plate motion across the seafloor/continen-tal 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 acousticranging measures the distance to the seafloor array. Acous-tic signals are needed because electromagnetic energy, onwhich GPS is based, does not propagate significantly inseawater. The basic underwater measurement is the time-of-flight of an acoustic pulse from the ship to a seafloor unitand back to the ship and the speed at which the acousticsignal travels in seawater (sound speed). From these twomeasurements the geometric range can be calculated. Thetime-of-flight can be measured to ±3 microseconds (equiv-alent to �2 mm of range) using a variety of techniques [e.g.,Spiess et al., 1997]. The main challenge is accommodatingchanges in sound speed particularly in the upper oceanwhere oceanographic forces drive variability that is signif-icant in both space and time.

Figure 2. Plate motion at extensional (a) and strike-slip (b) boundaries occurs episodically near theboundary, 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 theepisodic region while in the far-field plate motion is steady at the half-rate. Within the intermediateregion, motion is steady state late into the post-event cycle, but can exceed the steady state rate followinga transient slip event.


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[10] To date, there is no practical method to sample thesound speed profile with sufficient temporal and spatialresolution to account directly for changes in sound speed.The horizontal stratification of sound speed, however, canbe exploited to mitigate its effect on positioning resolutionin the following manner. Three or four precision trans-ponders are deployed on the seafloor to form an equilateraltriangle or square inscribed in a circle with the radius of thenominal water depth (Figure 3). By maintaining the shipnear the center of the array (�10 m), the vertical (launch)angle from the shipboard transducer to each transponder canbe made equal, forcing the acoustic signals to spend thesame amount of time within each horizontal layer. As soundspeed changes in the upper ocean, all rays lengthen andshorten equally. Because the transponders are evenly spacedaround the circumference of the inscribing circle with theship at the center, the coherent lengthening and shorteningof ranges is balanced in the horizontal. The upper oceansound speed variability will appear to move the seafloorarray vertically, but will not bias the horizontal positionestimate.[11] To implement this approach we maintain the ship at

the array center and collect several tens-of-hours of contin-uous GPS and acoustic data. Traveltimes from the ship aremeasured to seafloor transponders and back, and converted togeometric range by ray-tracing through the mean soundspeed profile. To estimate the mean sound speed we repeat-edly sample the ocean with a conductivity-temperature anddepth (CTD) device cycled from the surface to the seafloor.

These casts are averaged to provide the background profilethat includes the lower order components of the sound speedfield. These casts cannot provide the temporal and spatialresolution to model sound speed on the scale of each acousticinterrogation. The un-sampled sound speed variability ismitigated by exploiting the horizontal stratification. With4–5 days of continuous data, the horizontal position of theseafloor array can be determined with at least centimeter-level repeatability in the global reference frame [Gagnon etal., 2005].[12] The GPSA approach relies on a ship (or buoy) to

provide the interface between the GPS and acoustic sys-tems. Specifically, the shipboard configuration includesthree GPS antennas mounted on the ship to form a trianglewith as large fore-and-aft and athwart-ship dimensions asare practical. Dual-frequency GPS carrier phase data aresampled at 1 Hz at the ship and on shore to provide thesecond-by-second positions of the shipboard GPS antennas.A hydrophone is mounted within a hollow, vertical tube thatpasses from the work decks through the bottom of the ship’shull. The hydrophone extends less than a meter below thehull 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 whereit remains visible. Corner cube reflectors are mounted beloweach GPS antenna and above the back of the hydrophone. Asurveying instrument is placed overtop of the tube tomeasure the distances and angles in two perpendicularplanes between all reflectors. These data give the offsets

Figure 3. The GPS-Acoustic approach determines the horizontal position of the seafloor array[horizontal components of (~A)] by combining GPS positioning of shipboard antennas (~D) with ashipboard survey among antennas and hydrophone (~C) with acoustic ranging (~E) to seafloor transponderswhose relative positions (~B) are known. Maintaining the ship near the array center assures that acousticvelocity variations are primarily a function of depth and do not bias the horizontal components of ~A.


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between the antennas and hydrophone. With these offsets,GPS positions at the antennas can be transferred to thehydrophone giving the global position of the hydrophone ona 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 zonemight be 100 km wide [e.g., Foulger et al., 1992], thisseemed unlikely for faster spreading, thus warmer oceaniccrust. At the time of installation, there were reports thatpressure transients due to seismic activity had been ob-served downhole in northern Juan de Fuca CORK installa-tions as much as 50 km from the ridge crest (Davis et al.,2001, Earl Davis, personal communication, 2000). To be inthe transition zone or determine its outer limit, it would bedesirable to be high up on the ridge flank. As noted above,the transponder array must be on the order of the waterdepth in radius, which would mean a footprint of about 5 kmin this area. Since crustal deformation within the arraywould present problems of interpretation, it would beappropriate to site the installation a distance off axis suchthat the footprint is small compared with the distance fromthe axis. This consideration, combined with the nature of thetopography, led to a site 25 km from the ridge crest. Thealong-strike location was chosen to minimize edge effectsfrom the Blanco Trough to the south, and Axial Seamountto the north as well as to be related to geodetic installationsat 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 comprisingthe array were installed to form a square with sides ofapproximately 4 km (Figure 3). A total of 98 h of simul-taneous GPS and acoustic ranging data were collected fromthe center of the array while CTD casts were conductedconcurrently. Return visits provided contemporaneous GPS,acoustic, and CTD data for 87 and 83 h in May 2001 andJune 2002, respectively. In September 2003, one of the fourseafloor transponders had failed and was replaced with anew transponder. The new transponder position wasreferenced to the old one to maintain the continuity of thetime series.[15] This was done by temporarily placing an additional

active transponder adjacent (1–2 m) to the inactive andreplacement transponders. By moving the GPS-positionedship in a �2-km radius circle while simultaneously rangingon the two active units their relative position was deter-mined aligned with Earth-Centered-Earth-Fixed (ECEF)frame. An optical survey device was then deployed to theseafloor and maneuvered to within 1–2 m of the units tomeasure geometric ranges among the three units. This wasrepeated from several locations around the transpondercluster to determine the position offset between the inactiveand replacement unit. This offset was then rotated into theECEF giving the location of the new transponder relative tothe old transponder in the global frame (see Gagnon andChadwell [2007] for details).

[16] Then, the temporary transponder was recovered anda total 14 h of GPS, acoustic, and CTD data were collected.During each campaign, 1-Hz GPS data were collected atthree stations (CHZZ, TPW2, and NEWP) along the Oregoncoast (Figure 1, inset).[17] The shipboard and shore GPS data were processed

with NASA Jet Propulsion Laboratory’s GIPSY OASIS-IIsoftware [Webb and Zumberge, 1997] using analysis de-scribed by Spiess et al. [1998], Chadwell and Bock [2001].In all years, second-by-second repeatability of the GPSantenna positions is 10–20 mm in the horizontal [Miura etal., 2002] and 20–30 mm in the vertical Chadwell and Bock[2001]. In each year, the shipboard optical survey data werereduced to connect the GPS antenna phase centers tothe acoustic hydrophone phase center with a precision of2–3 mm [Chadwell, 2003]. GPS antenna positions weretransferred to the hydrophone, providing 20–30 mm levelpositions of the shipboard hydrophone on a second-by-second basis. Finally, these are combined with the traveltimesand mean sound speed profiles to estimate the location of theseafloor array in the International Terrestrial ReferenceFrame 2000 (ITRF2000) [Altamimi et al., 2002] at the epochof 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 calculatethe positional uncertainties, the formal error estimatesfrom GIPSY are multiplied by 3 to account for thewell-known underestimation of the formal error estimatesfrom GIPSY [Larson et al., 1997]. Then, the scaled GPSposition uncertainties are propagated to the hydrophonethrough the transformation equation that includes theuncertainties of the surveyed antenna-hydrophone offsets.Next, the time series of hydrophone position uncertaintiesis propagated with the traveltime uncertainties through aleast squares estimator of the array position uncertainty.In this calculation sound speed was fixed to the averageof all CTD profiles collected during the cruise. Thedeparture of the instantaneous sound speed from themean causes scatter in the traveltime residuals [Spiess etal., 1998]. This scatter increases the reduced Chi-squareto 5–10. The propagated array position uncertainty isscaled by this factor as is required by least squaresestimation theory.[19] The east and north position 1-s uncertainties range

from ±4 to ±6 mm from 2000 to 2002, or about an orderof magnitude more than what might be expected from aland-based site using continuous GPS tracking. Undoubt-edly, there is more to learn about the error budget ofGPSA positioning; however, we omit discussions of moresophisticated error components [e.g., Mao et al., 1997]and instead limit our analysis to that more consistent withthe assessments of the first applications of GPS forcrustal deformation measurements [e.g., Davis et al.,1989]. In 2003, the uncertainty increased to ±18–22 mmdue to a shorter data span (14 versus 80 h, with ±10 mm) andadditional uncertainty (±17–20 mm) from registering areplacement transponder.

4.1. GPSA Velocity Relative to ITRF2000

[20] The estimated east and north positions were weight-ed by the inverse square of their 1-s uncertainty and linearfits 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 theGPS-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 the2003 position estimate which is based on only 14 h of data (±10 mm) and contains additional uncertaintyfrom replacing an inactive seafloor transponder (±17–20 mm). Vertical lines show epochs of the 2 June2000 and 16 January 2003 earthquakes along the Blanco Transform. The epoch for the mid-1980s eventalong north Cleft segment is not shown.

Figure 5. (a) Compressional velocity profile from McDonald et al. [1994] at the Cleft segmentcombined with those from Cudrak and Clowes [1993] and Barclay and Wilcock [2004] from theEndeavour 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 fromthe profiles from the seafloor to the base of elastic layer. Below this is the visco-elastic half-space with aconstant shear modulus 50 GPa and a bulk modulus of 150 GPa.


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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). Forthe current solutions, at the 1-s uncertainty level (67%) alleast and north components lie along the linear fit withintheir positional uncertainty.[21] Spiess et al. [1998] reported the 1-s positional

uncertainty of ±39 mm east and ±8 mm north from 3surveys conducted in 1994, 1995, and 1996 with 23.6,32.0, and 33.5 h of GPSA data, respectively. At that timein the mid-90s, a combination of self-generated ship noiseand poor acoustic signal recognition and processing, meantthat only about 10% of the interrogation pings resulted inusable traveltimes. In addition, the ship used to collect thedata was manually steered and could only hold station towithin �150 m of the array center rather than the preferred10-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 andincreased 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 theinterior of the PA plate with present-day geodetic techni-ques. This is not possible for the Juan de Fuca plate becauseit has no exposed landmass for conventional geodeticstations from which a present-day Euler pole can becalculated. The present-day motion of the Juan de Fuca-Pacific plates must be inferred from the geomagneticanomaly pattern until such time when at least three GPSAsites 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 fromspace-based geodesy through tracking the motion of theplates over a few years, primarily with GPS [e.g., Altamimiet al., 2002; Sella et al., 2002]. Independently, present-dayEuler poles have been determined by averaging over thefinite rotation from the present back to geomagnetic

anomaly 2A (3.075 Ma) [e.g., DeMets et al., 1994]. Ratherconsistently for most plates, the space-based velocitiesmeasured at the stable interior of the plate and thegeomagnetic-based velocities agree within their respectivemeasurement uncertainties.[24] It is, therefore, quite natural to propose the anomaly

2A pattern to describe the present-day relative motion ofthe JdF-PA plates, just as the NUVEL-1 and NUVEL-1AJdF-PA Euler poles are constructed [DeMets et al., 1990,1994]. However, Wilson [1993] cautions that there may beapplications where it is important to recognize changes inthe pole since 3.075 Ma. Such pole changes can result fromreorientations of the JdF plate in response to interactionswith the larger PA and North America (NA) plates. Morerecently investigators have adopted the total reconstructionpole from 0–0.78 Ma of Wilson [1993] to calculate themean rate and instantaneous velocity of JdF-PA motion. Wealso adopt this to define the rigid-body motion of the JdF-PA plates and to compare with the observed GPSA-PAvelocity, though we note the velocities implied by these twopoles do not differ significantly for the purposes of thisstudy (Figure 1).[25] Next, we transform the GPSA velocity at the 25-km

site from ITRF2000 to a velocity relative to the Pacificplate. 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 correla-tion of �0.34 Beavan et al. [2002]. The GPSA-PA vectorfor the observed motion is compared to the velocity calcu-lated from the 0–0.78 Ma Euler pole of Wilson [1993] inFigure 1. Upon inspection of Figure 1 the 95% error ellipsesof GPSA-PA and that from the Wilson [1993] 0–0.78 MaEuler pole do not overlap, indicating the possible impor-tance of modeling transient motions.

5. Transient Motions

[26] Large transient motions that persist for years havebeen observed by direct geodetic measurements in northern

Table 1. Model Estimates of Displacement at GPSA Site From Ridge and Transform Events

Measurement Epoch

Observed Model Displacement Corrected Solutionsa

GPSA-ITRF20000 Ext.085 SS0001 SS0002 SS0003 SS003 GPSA-ITRF20001 GPSA-ITRF20002 GPSA-ITRF20003

East, mm2000.5914 0.0 ± 4.2 0.0 0.0 0.0 0.0 0.0 0.0 ± 4.2 0.0 ± 4.2 0.0 ± 4.22001.3915 11.7 ± 4.9 �0.8 1.8 0.7 �1.7 0.0 12.7 ± 4.9 11.6 ± 4.9 9.2 ± 4.92002.4490 45.4 ± 5.8 �2.5 3.5 1.4 �3.7 0.0 46.4 ± 5.8 44.3 ± 5.8 39.2 ± 5.82003.6849 42.4 ± 22.1 �4.0 5.1 2.0 �5.7 b�5.9 37.6 ± 22.1 34.6 ± 22.1 26.8 ± 22.1

East rate (mm/yr) fromweighted fit North (mm)

22.0 ± 3.6 22.0 ± 3.8 20.8 ± 4.0 18.2 ± 3.9

2000.5914 0.0 ± 3.7 0.0 0.0 0.0 0.0 0.0 0.0 ± 3.7 0.0 ± 3.7 0.0 ± 3.72001.3915 �1.0 ± 4.4 0.5 12.5 2.6 5.7 0.0 12.0 ± 4.4 2.1 ± 4.4 5.2 ± 4.42002.4490 7.4 ± 5.3 1.3 26.1 5.4 12.1 0.0 34.8 ± 5.3 14.1 ± 5.3 20.8 ± 5.32003.6849 13.6 ± 19.4 2.2 38.5 8.0 17.9 b8.1 62.4 ± 19.4 32.0 ± 19.4 42.0 ± 19.4

North rate (mm/yr) fromweighted fit

4.5 ± 3.3 19.6 ± 3.4 8.7 ± 3.6 12.2 ± 3.5

Correlation coefficient(r) �0.0369 �0.0369 �0.0369 �0.0369Reduced Chi-square(c2) 1.03 1.15 1.31 1.26

aGPSA-ITRF2000i = GPSA-ITRF20000 + Ext. 085 + SS003 + SS000i, where i = 1, 2, 3 corresponds to models of the 2 June 2000 event. GPSA-ITRF20000 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 toM0; 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.

bContains both the co-seismic and post-seismic components.


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Figure 6


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Iceland with the 1975–1981 Krafla rifting episode betweenthe 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 moccurred at the Rift. Nearly 10 years following the event,velocities were still more than twice the full-spreading ratewithin 50 km of the Rift. Similarly, co-seismic slip followedby accelerated motion exceeding 2–3 times steady state hasbeen observed with major strike-slip earthquakes along andnear the San Andreas Fault in 1906 [Thatcher, 1974], 1989[Burgmann et al., 1997], 1992 [Pollitz et al., 2000], and1999 [Pollitz et al., 2001].[27] One model of these phenomena begins with crust

composed of an elastic layer overlying a viscoelastichalfspace. A slip event within the elastic layer rapidlyreleases accumulated elastic strain displacing the elasticlayer. This causes co-seismic motion. The newly displacedupper layer transfers stress across the elastic/viscoelasticboundary, which more slowly diffuses to relieve theboundary stress. The relaxation of the viscoelastic half-space further displaces the surface of the elastic layerresulting in post-seismic motion [Elsasser, 1969; Bottand Dean, 1973].[28] Therefore slip events at either the JdF Ridge or the

Blanco Transform may contribute to motion observed at theGPSA site. We must calculate the motion induced by elasticand viscoelastic response of the JdF plate to recorded slipevents at the Ridge and Transform. We begin by reviewingthe boundary events.

5.1. Cleft Extensional Events

[29] Between 1983 and 1987, a volcanic event occurredalong a �30 km-section of the northern Cleft segment of theJuan de Fuca Ridge [Chadwick et al., 1991; Embley et al.,1991]. Evidence of an event was first detected in 1986 byobservance of a hydrothermal plume in the water columnover the north Cleft [Baker et al., 1989]. The exact dates areunknown because the event(s) occurred prior to scientificmonitoring 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 andsubsequent deep-towed platforms constrain the eruptiondates between 1983 and 1987. A quantitative differencingof the bathymetric data revealed an eruptive volume of0.05 km3 [Fox et al., 1992]. Using geologic and hydrother-mal observations, Embley and Chadwick [1994] have pro-posed an intrusion/eruption episode that likely involved twoseparate events separated by at least 7 months. They suggesttwo lateral dikes that in total extended from 44�530N to45�100N.[30] There is no evidence for recent extensional events in

the southern Cleft segment. Normark et al. [1983] mappedthe lava plain covering the axial valley floor at south Cleft

and estimate the age to be less than a few hundred years.Geodetic monitoring did not detect extension across theaxial valley from 1994–1999 [Chadwell et al., 1999;Hildebrand et al., 1999] and from July 2000 through June2003 [Chadwick and Stapp, 2002;W.W. Chadwick, personalcommunication, 2006]. SOSUS monitoring since 1994 hasdetected 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 shockoccurred at the east side of a foreshock cluster that initiated16 h earlier centered at �44�210N to 130�150W. During thefollowing 33 h aftershocks extended from the main shocklocation along a 44-km-length section of the TransformDziak et al. [2003]. The National Earthquake InformationCenter (NEIC) derived moment tensor suggested an eventthat was 12 km deep, mostly right-lateral in motion, striking128�, 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 aright lateral motion with 17 km depth, strike of 127�, anddip of 83�. D. Bohnenstiehl, personal communication[2006] characterized this event with a 9 km depth basedupon 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 theevents in an elastic [Okada, 1992; Gomberg and Ellis,1994] and a viscoelastic model [Pollitz, 1992] to calculatethe response of the JdF plate to the transient events at theboundaries.[34] We construct a model of the crust in the vicinity of

the south Cleft segment from a composite of sources. Theelastic properties of the crust are calculated from thecompressional (Vp) and shear wave (Vs) velocities andcrustal density (r). We use the compressional velocityprofile from McDonald et al. [1994] at the Cleft segmentcombined with those from Cudrak and Clowes [1993] andBarclay and Wilcock [2004] from the Endeavour segment ofthe JdFR. The shear wave velocity profile is also fromBarclay and Wilcock [2004]. Stevenson et al. [1994] usedseafloor and sea-surface gravity data to model the densitystructure along the southern Juan de Fuca Ridge. Theyestimate the average density to be 2630 ± 50 km m�3 in theupper 2 km of ocean crust. We adopted a density profile thatincreases with depth with values of 2700 km m�3 at theseafloor increasing to 3100 km m�3 12 km depth. The bulkmodulus (k) and shear modulus (m) of the crust arecalculated in 2-km-thick layers from these profiles fromthe seafloor to the base of elastic layer (Figure 5). Belowthis is the visco-elastic half-space with a constant shearmodulus 50 GPa, a bulk modulus of 150 GPa and aviscosity of 3.0 � 1018 GPaS.

Figure 6. Total model calculated displacements from 2 August 2000 through 5 September 2003 of the Juan de Fuca platedue 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) usingrupture 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, blackarrow at GPSA site. Insets show total east and north displacement at the GPSA site for each model.


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[35] Constraints upon the input parameters controlling theelastic and viscoelastic models are weak. The elastic layerdepth can reasonably range from 6 km, as suggested byseismic imaging [Canales et al., 2005], to 12 km implied bythe depth of the strike-slip events. We calculate displace-ments using 6 km and 12 km elastic layer thicknesses andshow results from the thicker model.[36] For the model calculations of the 1985 extensional

event, we assumed a single lateral dike intrusion thatextended the entire 30 km, that the dike was just belowthe surface, and ruptured the entire layer 2C with a width of2 m. The total displacement since 1985 is approximately50 mm east and 15 mm south (Figure 6a, inset); however,over the span of the survey the displacement is less than5 mm (Ext085 in Figure 6a and Table 1). The deformationtransient from the 1980s intrusion has decayed down to nearbackground levels.[37] The 2 June 2000 event, though perhaps a smaller

equivalent moment magnitude than the 1980s event, oc-curred just two months prior to the establishment of theGPSA site and may have induced a significant motion.Dziak et al. [2003] interpreted the aftershock distribution tobe the rupture length, and applied a scaling law [Scholz,1982] to give a slip of 60 cm (SS0001 in Figure 6b andTable 1). We note that calculation of the moment magnitudewith 44 km length, 12 km depth, and 60 cm slip exceeds theseismically observed moment, though this is not significantfor the application by Dziak et al. [2003]. A possibleinterpretation is that the main shock is the seismically

observed component, while the aftershock pattern duringthe following 33 h delineates the likely extent of aseismicslip, recognizing that transform faults can have large aseis-mic components, i.e., low seismic coupling [Boettcher andJordan, 2004].[38] To gauge the sensitivity of the model to this inter-

pretation, we calculate deformation for two additional slipmodels. First, rather than using the aftershock cluster as thelength, we take the seismically observed moment magnitudeand apply the scaling law of Scholz [1982] to get a rupturelength of 22.9 km and a slip of 19 cm (SS0002 in Figure 6cand Table 1). Second, we note that though the NEICestimated the event with a strike of 128�, the distributionof the aftershocks allows the interpretation that the strike ofthe slip be aligned with the bathymetric trace of theTransform at 115�. Here we again use the aftershocks forthe rupture length, apply the scaling law for the slipmagnitude, but take the strike to be 115� (SS0003 inFigure 6d and Table 1).[39] As expected, the orientation of the displacement is

controlled by the strike of the slip, its magnitude, and therupture length. Both models (SS0001 and SS0002) with strikeof 128� give southerly displacements with a small westerlycomponent (Figures 6b and 6c; and Table 1), though thetotal displacements are 38.8 and 8.2 mm for rupture lengthsand slips of 44 km, 60 cm and 23 km, 19 cm, respectively.With the strike of the slip aligned with the general orienta-tion of the Blanco Transform (115�), the motion (SS0003) is

Figure 7. East and north position estimates (solid symbols) in the ITRF2000 frame estimated byapplying the calculated transient motions to the GPS-Acoustic positions. Again, the values are shownwith 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 containsadditional uncertainty (±17–20 mm) from replacing an inactive seafloor transponder. Vertical lines showepochs of the 2 June 2000 and 16 January 2003 earthquakes along the Blanco Transform. The epoch forthe mid-1980s event along north Cleft segment is not shown.


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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. Thisevent has both co-seismic and post-seismic components(SS003) that are generally southeasterly with a total dis-placement of 10 mm (Figures 6e, 6f, and Table 1).[41] Constraints upon the input parameters controlling the

viscoelastic computations are weak. For example, the elasticlayer depth was chosen at 12 km, but could reasonably varyfrom 6–12 km depth. Likewise, the slip event dimensions(depth, rupture length, and fault slip) are not stronglydetermined. To study these effects, solutions were repeatedwith a 6 km elastic layer depth. Shallower elastic layerdepths tended to reduce the calculated relaxation, but do notchange any conclusions presented later. Therefore we note

that though our choice of fault models and crustal propertiesis not unique, nor exhaustive, they are an appropriaterepresentation of the perturbation probable from transientevents given the detail of our geodetic data.[42] Four solutions are constructed to probe the influence

of transient motions upon the observed GPSA motion. Thefirst solution is simply the observed GPSA motion withoutcorrecting for transient motion (GPSA-ITRF20000). Thefirst corrected solution (GPSA-ITRF20001) subtracts fromGPSA-ITRF20000 the displacements from the 1985 exten-sional event (Ext.085), the 16 January 2003 strike slipevent (SS003) and model 1 of the 2 June 2000 strike slipevent (SS0001). The next solution (GPSA-ITRF20002)subtracts Ext.085, SS003, and SS0002. The third correctedsolution (GPSA-ITRF20003) subtracts Ext.085, SS003, andSS0003. GPSA-ITRF20000 was given in Figure 4 and

Figure 8. Velocity vectors relative to the PA plate along with their 95% confidence ellipses for theWilson [1993] 0–0.78 Ma Euler pole (gray) and the GPSA observations (black) calculated for: (a) withno 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 toM0,(d) with slip model from Dziak et al. [2003] with strike aligned with the Transform.


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Table 1, and repeated in Figure 7. GPSA-ITRF20001–3

positions, corrected for the transient effects, and weightedlinear fits estimated for each solution are given in Table 1and Figure 7.

5.4. Transient Component of GPSA Velocity

[43] To determine the kinematic regime of the GPSA sitewe compare the observed GPSA motion with and withouttransient motion components to the average motion of therigid JdF-Pacific (PA) plates.[44] The GPSA-PA vectors for observed motion with and

without transient motions removed are computed in Table 2and shown in Figure 8. Upon inspection of Figure 8, twodeductions 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 bound-aries, while perhaps significant, are modest, not exceeding10% 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 thatcalculated from the Wilson [1993] 0–0.78 Ma Euler pole(Figure 8). Also, the right-hand side of Table 2 gives theGPSA-JdF motion, i.e., the motion of the site as measuredby GPS-Acoustics relative to the JdF plate itself as pre-dicted by the geomagnetic anomaly at 0.78 Ma. If the sitewas measured by GPSA to be moving as predicted by theWilson [1993] 0–0.78 Ma Euler pole, then the discrepancywould be zero. The GPSA-JdF1–3 velocities are small and

Table 2. Velocity (mm/a) of GPSA Site Relative to the Pacific, and Juan de Fuca Plates

GPSA-PAa = GPSA-ITRF2000a - PA-ITRF2000b GPSA-JdFa = GPSA-PAa - JdF-PAc

GPSA-PA0 GPSA-PA1 GPSA-PA2 GPSA-PA3 GPSA-JdF0 GPSA-JdF1 GPSA-JdF2 GPSA-JdF3

e 58.6 ± 3.6 58.7 ± 3.8 57.5 ± 4.0 54.8 ± 3.9 8.5 ± 3.7 8.6 ± 3.9 7.4 ± 4.1 4.7 ± 4.0n �24.6 ± 3.4 �9.5 ± 3.4 �20.4 ± 3.6 �16.9 ± 3.6 �9.7 ± 3.8 5.4 ± 3.8 �5.5 ± 3.9 �2.0 ± 3.9r �0.0392 �0.0401 �0.0396 �0.0396 +0.0386 +0.0283 +0.0283 +0.0300

aWith 0, 1, 2, 3 representing solutions given in Table 1.bFrom 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.cFrom 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.

Figure 9. (a) Location of episodic, intermediate, and continuous zones supported by seafloor geodeticdata. Plate velocities relative to the axis are plotted across strike of JdF Ridge at location of geodeticexperiments. 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 Eulerpole 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 experimentsdefines the extent of episodic region (dark gray shading). Full 1/2 rate at GPSA site indicates continuouszone (dark gray shading). Additional, future geodetic monitoring in the region from 0.5 to 25 km off-axiscould further constrain the width of each zone. (b) Shows spreading boundaries implied by geophysicalobservations. Black line shows bathymetric profile across-strike. On-axis light gray region is approximatelimit of episodic activity from Carbotte et al. [2006] and Karson et al. [2002]. Curved-dashed line isapproximate seafloor profile. Inflection of the curvature from concave down to convex occurs between5–15 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 thediscrepancy is insignificant and the two estimates agree.However, the uncorrected velocity, GPSA-JdF0, is 2–3times its uncertainty and the 95% error ellipses of GPSA-PA0 and that from the Wilson [1993] 0–0.78 Ma Euler poledo not overlap, indicating the importance of modelingtransient motion.[46] The second deduction is supported by the small

difference (�10%) between GPSA-PA0 without correctionand GPSA-PA1–3 with corrections. This leads to the inter-pretation that the motion observed at the GPSA site is dueprimarily to steady state conditions and not to the threemodeled transient effects. This is an important point givenresults from Iceland and San Andreas have shown transientmotions 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 andGPSA-PA2 model the 2 June 2000 event as SS0001 andSS0002, respectively, and show the difference between tworupture lengths at the same strike. Because the aftershockdistribution clearly indicates a change in stress conditionson the fault, the total slip may be under-represented bySS0002, which limits application of the scaling law to theobserved seismic moment magnitude. Therefore the GPSA-PA2 velocity may be incomplete. SS0001 represents a morecomplete rupture length because it reflects the actual after-shock pattern, but the strike is not well constrained by theseismic solution, forcing GPSA-PA1 to have a larger northcomponent. GPSA-PA3 best agrees with the velocity calcu-lated from the Wilson [1993] 0–0.78 Ma Euler pole. GPSA-PA3 modeled the 2 June 2000 event defining the rupturelength by the aftershock pattern and the slip strike to beparallel to the Transform, which is a long-term, persistentplane of weakness.

6. Mid-Ocean Ridge Dynamics

[48] We next examine the implications of these results todynamics at mid-ocean ridges. We assume the GPSA-PAvelocity represents near steady state conditions as we haveshown the calculated transient effects are only �10% of thefull-rate velocity. Figure 9a shows a plot of the velocitiesmeasured in the vicinity of the site including no significantextension across the the 1-km-wide axial valley from 1994–1999 and 2000–2003 [Chadwell et al., 1999; Hildebrand etal., 1999; Chadwick and Stapp, 2002; W. W. Chadwick,personal communication, 2006]. As discussed previously,no motion under steady state conditions indicates that thesite lies within the episodic zone. We note that the totalextent of the on-axis arrays is �1 km thus the episodic zoneextends to at least �0.5 km to either side of the axis. Thecontinuous zone is defined by the velocity from the 0–0.78Ma Euler pole Wilson [1993] and is plotted relative to theaxis (i.e., half-rate). To show the sensitivity of our data weplot the half-rate velocity for both GPSA-PA0 and GPSA-PA3 that bracket our results. Figure 9 shows that the GPSAsite is moving at the half-rate, constraining the continuousregion 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 intermediateregion between 0.5 and 25 km that accommodates 26 mmdeformation each year. This deformation must occur

aseismically because no seismic activity above magnitude1.8 was recorded by SOSUS during the survey spananywhere 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 theJdFR specifically. Carbotte et al. [2006] suggest the widthaffected by episodic magmatic processes is approximatelyequal to the depth of the axial magma chamber that theymeasure to be 2.0 km deep. In their model, the axialvolcanic ridge that lies within 2 km of the axis is potentiallystill active and would be affected by a subsequent dikeinjection. Beyond 2 km, the graben faults are no longeraffected by magmatic intrusions at the ridge [see alsoCanales et al., 2005]. Karson et al. [2002] suggest thatextensive vertical subsidencemust occur (likely episodically)within the axial valley out to �1.5 km to accommodatefeatures observed in the exposed wall of the Western BlancoTransform Fault. We note the extent of the seafloorgeodetic arrays on the axial valley floor is �0.5 km toeither side of the axial Cleft. Both Carbotte et al. [2006]and Karson et al. [2002] would extend the region of episodicactivity to 1.5–2 km off-axis. The geodetic measurementsshowing 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 isothermdefining the boundary between the lithosphere and astheno-sphere. This thickens the elastic layer that at an unknowndistance off-axis is sufficient to act as a stress guide. Inaddition, faulting and possible off-axis magmatic intrusionforming the abyssal hills can be seen to begin as close as 5 kmoff-axis. At some unknown distance off-axis these faultsbecome inactive and no longer accommodate motion. Thesite 25 km off-axis is at the full-half rate and places someconstraint on these two processes. It suggests that the elasticlayer is sufficiently thick to be a stress guide and behave aspart of the rigid JdF plate and that to the east of the GPSAsite, i.e., toward the plate interior that the abyssal hillfaulting has predominately ceased. This is generally consis-tent with the Buck et al. [2005] unbending model thatpredicts fault motion ceases where the curvature of theplate flattens which occurs �20 km off-axis at this site(Figure 9b).

7. Conclusions

[52] GPS-Acoustic positions of a seafloor array over 4annual campaigns are consistent with a linear trend. Ap-proximately 80 h of continuous GPS-Acoustic data allows aprecision of ±4–6 mm, sufficient to address a number ofseafloor tectonic questions. Also, replacing a seafloor tran-sponder is possible with at-most 10–20 mm uncertainty,demonstrating that seafloor geodetic arrays can be main-tained 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 fromthe Wilson [1993] 0–0.78 (and 0–3.075) Ma Euler pole(s)once transient motions are removed. Transient events at theplate boundaries account for �10% of the total GPSA-PAmotion according to elastic/viscoelastic models. This sug-gests the GPS-PA velocity is due primarily to steady state


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plate dynamics. Assuming the geologically derived raterepresents full-rate JdF-PA spreading, the site 25 km eastof the Ridge is interpreted to be in a region of continuousplate motion. This is a robust result of this study and holdsfor GPSA-PA velocities both uncorrected and corrected fortransient motions. These are the first geodetic results todirectly constrain the width of active spreading to lie within25 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 nosignificant extension across the 1-km-wide axial valley.This raises several questions of how the strain is accumu-lated. The geomorphology of the abyssal hills suggests thefaults accommodate deformation is some manner. We spec-ulate that as the upper mantle creeps steadily from the ridgeand the elastic layer thickens, the upper crust accumulatesstrain which is accommodated by motion along the faults.The frequency of the fault motion is unknown. If thismotion occurred during the span of our measurements thenit 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 havedisplaced and instead strain accumulated in the near-axiselastic crust. Eventually, this elastic strain will be releasedand accommodated by fault motion perhaps generating aseismically recorded event. The partitioning between faultmotion and elastic strain accumulation might be discernedwith additional near-axis geodetic studies.

[55] Acknowledgments. Herb Dragert, M. Meghan Miller, AndrewMiner, Dan Johnson, Chris Goldfinger, Jason Chaytor, Chris Fox, JonathanKlay, Bill Chadwick and UNAVCO helped with collection of the 1-Hz GPSdata at shore stations. John Hildebrand aided with site selection and joinedthe 2000 cruise. Captain and crew of the R/V Roger Revelle provided at-seasupport. Richard Zimmerman, Dave Jabson, Dennis Rimington and DavePrice provided engineering support. We thank Katie Phillips, Katie Gagnon,and Neil Kussat for at-sea help, discussions about data reduction, andcomments on this manuscript. We also thank Kelin Wang for comments ona early version of this manuscript. Bob Dziak and Del Bohnenstiehlprovided data and insight on the Blanco Earthquakes. Bill Chadwick madeavailable recent extensometer results. Douglas Wilson and John Beavenprovided insightful reviews that much improved the paper. Bathymetryfrom the RIDGE Multibeam Synthesis Project. Map was produced usingthe Generic Mapping Tools package [Wessel and Smith, 1998]. This workwas supported by NSF grant OCE-9907247 from the Marine Geology andGeophysics Program. Fred N. Spiess died on 8 September 2006 while thispaper 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, 9500Gilman Drive, La Jolla, CA 92093-0205, USA. ([email protected])


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plate motion at the ridge-transform boundary of the south cleft

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