Potential role of mantle-derived fluids in weakening the San Andreas Fault et al_JGR_2009.pdf ·...

Potential role of mantle-derived fluids in weakening the San Andreas

Fault

Patrick M. Fulton1,2 and Demian M. Saffer1

Received 5 September 2008; revised 2 March 2009; accepted 15 April 2009; published 31 July 2009.

[1] On the basis of both geomechanical and thermal data, the San Andreas Fault (SAF)has been interpreted to act as a weak plane within much stronger crust, allowing it toslip at very low shear stresses. One explanation for this weakness is that large fluidoverpressures exist locally within the fault zone. However, mechanisms for generating,maintaining, and localizing pressures within the fault are poorly quantified. Here weevaluate whether realistic sources of mantle-derived fluids, proposed on the basis of highmantle helium signatures near the SAF, can generate localized fluid pressures withinthe fault zone in a manner consistent with a wide range of observations along the fault andin the San Andreas Fault Observatory at Depth borehole. We first calculate a reasonableestimate of the magnitude and location of a mantle-derived flux of water into thecrust. This fluid flux results from dehydration of a serpentinized mantle wedge followingthe northward migration of the Mendocino Triple Junction and the transition fromsubduction to strike-slip tectonics. We then evaluate the potential effect of this water onfluid pressures within the crust using 2-D cross-sectional models of coupled fluid flowand heat transport. We show that in models with realistic permeability anisotropy,controlled by NE dipping faults and fractures within the country rock, large localized fluidpressures can be focused within a SAF acting as a hydrologic barrier. Our results illustratea simple and potentially plausible means of weakening the SAF in a manner generallyconsistent with available hydrologic, thermal, and mechanical constraints.

Citation: Fulton, P. M., and D. M. Saffer (2009), Potential role of mantle-derived fluids in weakening the San Andreas Fault,

J. Geophys. Res., 114, B07408, doi:10.1029/2008JB006087.

1. Introduction

[2] The San Andreas Fault (SAF), along with many otherplate boundary faults, has been interpreted to slip at shearstresses considerably lower than expected for typical rockand fault gouge friction coefficients, and assuming hydro-static pore fluid pressures [e.g., Lachenbruch and Sass,1980; Zoback et al., 1987; Hickman, 1991]. The lack of athermal anomaly from frictional heating near the SAF [e.g.,Lachenbruch and Sass, 1980; Sass et al., 1997; Williams etal., 2004; Fulton et al., 2004b], and maximum horizontalprincipal stress (sH) directions inferred to be at a very highangle (up to 85�) with the SAF [e.g., Townend and Zoback,2004; Hickman and Zoback, 2004] imply that the fault isweak in an absolute sense, and also considerably weakerthan the surrounding crust. One hypothesis for both theapparent absolute and relative fault weakness is that fluidpressures within the fault zone are markedly higher than inthe surrounding crust [Rice, 1992; Byerlee, 1990, 1992].The mechanisms for generating and maintaining such pres-

sures are poorly understood, although there has beenconsiderable speculation regarding the potential role ofmantle-derived fluids, primarily on the basis of geochemicaldata [e.g., Rice, 1992; Kennedy et al., 1997; Kharaka et al.,1999; Faulkner and Rutter, 2001; Wiersberg and Erzinger,2007]. Here, we use numerical models of fluid flow andheat transport to investigate the potential role of mantle-derived fluids in generating and sustaining large fluidpressures that are localized within the fault zone, and testingwhether such a mechanism is consistent with a range of datafrom the region and the San Andreas Fault Observatory atDepth (SAFOD).

1.1. Mantle Fluids and a Weak San Andreas Fault

[3] SAFOD is a borehole observatory consisting of a 3.2 kmdeep main borehole and a 2.2 km deep pilot hole locatednear the town of Parkfield in central California, a focusregion for this study (Figure 1). The main borehole isinclined and crosses the fault from the SW at �2.7 kmdepth (Figure 2). Numerous observations at SAFOD and inthe surrounding region provide important constraints onconceptual models of fault strength and crustal-scale hydro-geology (Figure 2, described in detail below). Notably,isotopic evidence from mud gases including high 3He/4Heratios and values of d13C from HCO3

� suggest the presenceof mantle-derived fluids in wells and springs primarily tothe NE of the SAF, as well as within the SAFOD borehole

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1Department of Geosciences, Pennsylvania State University, UniversityPark, Pennsylvania, USA.

2Now at College of Oceanic and Atmospheric Sciences, Oregon StateUniversity, Corvallis, Oregon, USA.

Copyright 2009 by the American Geophysical Union.0148-0227/09/2008JB006087$09.00

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as it crosses into the North American plate [Kennedy et al.,1997; Kharaka et al., 1999; Wiersberg and Erzinger, 2007].[4] Previous authors have suggested that a permeable

fault or damage zone surrounded by low-permeabilitycountry rock may act as a conduit for mantle-derived fluidsand allow for large fluid pressures to be localized within thefault zone [Rice, 1992; Kennedy et al., 1997]. Rice [1992]showed that with strongly anisotropic fault zone permeabil-ity, permeability in the surrounding crust considerably lowerthan in the fault, and a near-lithostatic pressure boundarydirectly beneath the fault at the base of the brittle crust, highfluid pressures develop locally within the fault duringtransient numerical simulations. Observations in and aroundSAFOD, however, suggest that this may not be an adequaterepresentation of the SAF, at least within the Parkfieldregion. SAFOD observations, including contrasts in bothfluid chemistry and fluid pressure across the fault zone,suggest that rather than acting as a permeable conduit, thefault appears to be considerably less permeable than thesurrounding crust and acts as a barrier for fluid flow [e.g.,Wiersberg and Erzinger, 2007] (Figure 2).[5] In addition, several observations suggest that the

source of mantle fluids is greatest NE of the fault zonerather than directly beneath it. In particular, helium isotoperatios from within the SAFOD main hole SWof the fault aretypical of meteoric and crustal-derived fluids, whereashigher values indicating a strong mantle component areobserved as the SAFOD borehole crosses the SAF into units

within the North American plate [Wiersberg and Erzinger,2007]. This mantle signature increases as the boreholecontinues to the NE, and additional well data 1.4 km NEof the fault zone near Parkfield reveal even greater valuesindicating that as much as 25% of the helium is of mantleorigin [Kennedy et al., 1997; Wiersberg and Erzinger,2007].[6] Within the Franciscan Complex to the NE of the SAF

(Figure 3), elevated pore pressures have been observed indeep boreholes, and a variety of geophysical surveysthrough the Parkfield region (seismic, gravity and magnetic,and magnetotelluric) have inferred a potentially overpres-sured zone capped by a serpentinite unit extending �12 kmfrom the fault to the NE at depths >�3 km (Figure 3)[Eberhart-Phillips and Michael, 1993; McPhee et al., 2004;Unsworth and Bedrosian, 2004]. Although no indirectindications of high fluid pressures within the fault zonewere observed during SAFOD drilling [Zoback and Hick-man, 2005], direct hydrologic measurements withinSAFOD have not yet been conducted (M. Zoback andS. Hickman, personal communication, 2009). It is alsoimportant to note that the borehole crosses the fault atrelatively shallow depth, such that inferences regardingfluid pressure at SAFOD may not fully represent processescontrolling fault strength throughout the upper �10–15 km.[7] The most promising source of mantle-derived fluids

comes from mantle wedge material, which was potentially

Figure 1. Map of California and Nevada. Quaternaryfaults are shown in gray, with the San Andreas Fault inblack (U.S. Geological Survey and California GeologicalSurvey and Nevada Bureau of Mines and Geology,Quaternary fault and fold database for the United States,2006, accessed 26 June 2008, available at http://earthquakes.usgs.gov/regional/qfaults/). The location of Parkfield,California, the San Andreas Fault Observatory at Depth(SAFOD), and the Mendocino Triple Junction (MTJ) arelabeled.

Figure 2. Schematic summary illustrating the SAFODboreholes and key hydrogeologic and geomechanicalobservations described in the text. Large-scale geologicunits are labeled for Salinian granite (Ksg) SW of the SAFand Great Valley Sequence (Kgv) to the NE.

B07408 FULTON AND SAFFER: MANTLE FLUIDS AND THE SAN ANDREAS FAULT

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hydrated and serpentinized during subduction prior to theformation of the SAF, and which would subsequentlydehydrate after the plate boundary along California transi-tioned from subduction to strike-slip tectonics (Figure 4)[Kirby et al., 2002]. The SAF was created �25 Ma when aportion of the East Pacific Rise spreading center subductedbeneath the North American plate creating two triplejunctions, the Mendocino Triple Junction (MTJ) to thenorth and the Rivera Triple Junction to the south [Atwaterand Stock, 1998]. The MTJ and the subducting slab havemigrated northward over time, and in its wake, subductionanalogous to that occurring in present-day Cascadia hasbeen replaced by the strike-slip tectonics of the SAF system[Atwater and Stock, 1998]. As the subducting slab isreplaced with upwelled asthenosphere (or remaining slabstalls beneath the overlying crust and thermally reequili-brates), serpentinized upper mantle, believed to remainbeneath the North American crust, is heated (Figure 4b)[ten Brink et al., 1999; Guzofski and Furlong, 2002; Kirbyet al., 2002; Brocher et al., 2003; Erkan and Blackwell,2008]. Serpentinite minerals are unique, and relevant to thisstudy, in that they contain up to 10–13 wt% water in theirmineral structure [Iwamori, 1998]. Kirby et al. [2002]combined a detailed thermal model with a kinetic reactionmodel of serpentinite formation and dehydration to illustratethat the temperature increase following the migration of theMendocino Triple Junction could drive continued progradedehydration of serpentinized mantle for tens of millions ofyears resulting in a large and sustained flux of water into theoverlying crust.[8] The presence of a substantial volume of highly

serpentinized mantle has been recognized through geophys-ical imaging beneath Cascadia [Bostock et al., 2002;Brocher et al., 2003; Blakely et al., 2005], as well as innumerous other subduction zones [e.g., Graeber and Asch,1999; Kamiya and Kobayashi, 2000; Kamimura et al.,2002; DeShon and Schwartz, 2004; Tibi et al., 2008; seealso Wada et al., 2008]. In Cascadia, on the basis of lowshear wave velocities, high Poisson’s ratio, and an invertedpolarity Moho along a trench-perpendicular transect, a zoneaveraging �50 km wide and �20 km thick within thecorner of the mantle wedge is interpreted to be at least50–60% serpentinized (Figure 4a) [Bostock et al., 2002].[9] Detailed models of the thermal response of the mantle

wedge to MTJ migration and subsequent serpentinite dehy-dration [Kirby et al., 2002; Guzofski and Furlong, 2002] are

also generally consistent with observations from a line ofvolcanic centers associated with the passage of the MTJ[e.g., Johnson and O’Neil, 1984; Whitlock et al., 2001,Whitlock, 2002; Furlong and Schwartz, 2004; Schmitt et al.,2006]. The petrology of these volcanics has been interpreted

Figure 3. Idealized cross section showing geologic structure of the San Andreas Fault and surroundingcrust in central California [after McLaughlin et al., 2000; Bailey et al., 1964; Irwin and Barnes, 1975;Eberhart-Phillips and Michael, 1993; McPhee et al., 2004; Bleibinhaus et al., 2007]. NE dippingmelange units within the Franciscan Complex (dark gray) and faults and fractures (annotated with arrowsindicating sense of motion) associated with both the accretion of the Franciscan Complex and present-daythrust faulting likely act as permeable pathways that may focus fluid pressure toward the SAF.

Figure 4. Schematic cross section illustrating the transi-tion from Cascadia-like subduction to strike-slip tectonicsfollowing the northward migration of the Mendocino TripleJunction. (a) Cartoon illustrating subduction prior to thecreation of the SAF (adapted with permission fromMacmillanPublishers Ltd. after Kirby [2000], copyright 2000), showingthe general location of serpentinized mantle wedge materialpredicted by thermal models and observed by geophysicalmethods [e.g., Bostock et al., 2002; Brocher et al., 2003;Wada et al., 2008]. (b) Idealized cross section followingMTJpassage; as the downgoing slab has migrated northward withthe MTJ, the space the slab once occupied is replaced byupwelling aesthenosphere which heats the overlying crustand mantle [e.g., Guzofski and Furlong, 2002]. This tem-perature rise dehydrates the serpentinized mantle [Kirby etal., 2002].


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to reflect ponding of small amounts of mantle-derived meltwithin the crust, which in turn drive crustal melting[Johnson and O’Neil, 1984]. The geochemistry of eruptedbasaltic lavas suggests that the mantle-derived melt issourced from a hydrated mantle wedge [Johnson andO’Neil, 1984; Whitlock et al., 2001, Whitlock, 2002; Fur-long and Schwartz, 2004; Schmitt et al., 2006]. Of relevanceto this study, these petrologic interpretations suggest that(1) hydrous mantle wedge material is present beneath the N.American plate following MTJ passage, and (2) that thismantle wedge material is significantly heated. In compari-son to the �20 km thick by �50 km wide zone of highlyserpentinized mantle material inferred in cross sectionbeneath Cascadia [Bostock et al., 2002], estimates basedon numerical modeling and petrologic constraints for theamount of mantle material that melts suggest a cross-sectional area �4–5 km thick and �10–20 km wide [Liuand Furlong, 1992; Schmitt et al., 2006]. This implies thatless than 10% of the inferred zone of serpentinized mantlemelts following MTJ passage; it is thus plausible that muchof the hydrated mantle wedge material would remain inplace following MTJ passage and gradually dehydrate[Kirby et al., 2002].[10] Here, we explore the potential role of mantle-derived

fluids in weakening the SAF by evaluating whether realisticfluid sources (in terms of both magnitude and location) fromthe dehydration of serpentinized mantle, together withrealistic permeability architecture for the fault and surround-ing crust, can generate large localized fluid pressures in amanner consistent with observations. This study differsfrom previous work examining the effect of mantle fluidson fluid pressures along the SAF system, in that (1) we use2-D hydrogeologic models to evaluate a realistic fluidsource rather than implementing a prescribed basal pressureboundary immediately below the fault zone; (2) we treat theSAF as a low-permeability barrier rather than as a conduit;and (3) we evaluate the possible role of NE dipping faultsand fractures within the surrounding country rock, whichmay act to focus pressures toward the fault zone, as a meansof localizing fluid pressures (Figure 3). Overall, this studyaddresses whether a prolonged, yet realistic mantle-derivedfluid flux can generate fluid pressures needed to explain aweak SAF when considering realistic 2-D permeabilitystructure; and whether a fault acting as a barrier and theexpected location of mantle fluid sources NE of the faultprecludes the role of mantle fluids in localizing pore fluidpressures within or along a near-vertical SAF.

1.2. Constraints on Fluid Pressures Necessaryto Explain Apparent Fault Weakness

[11] High angles between the SAF and the maximumhorizontal principal stress (y � 65�–85�) have beeninferred from earthquake focal mechanisms and borehole-based observations throughout California [e.g., Zoback etal., 1987; Mount and Suppe, 1987; Provost and Houston,2003; Townend and Zoback, 2004]. The largest values (y =�85�) occur in the San Francisco Peninsula region. In theParkfield area and throughout much of central and southernCalifornia, values of y = �68� are observed [Townend andZoback, 2004]. Within the SAFOD boreholes, similarlyhigh angles between the maximum horizontal principalstress and the fault plane have been determined from

borehole breakouts and drilling-induced tensile fractures(y � 69� in the SAFOD pilot hole) [Hickman and Zoback,2004], and from shear velocity anisotropy (y � 70� at adistance of 200 m from the fault) [Boness and Zoback,2006]. These high angles imply that the fault acts as a weakplane within a stronger country rock [Zoback et al., 1987;Mount and Suppe, 1987; Townend and Zoback, 2004;Hickman and Zoback, 2004]. The relative weakness canbe explained if there is material of low friction coefficientlocalized within the fault zone, the fluid pressure ratiolocally within the fault (lf) is very large, or a combinationof both (see Discussion). Fluid pressure ratio l is defined as

l ¼ P=rrgz ð1Þ

where P is the pore fluid pressure (Pa), rr is bulk rockdensity (kg m�3), g is gravitational acceleration (m s�2),and z is depth (m) (all variables in our analysis, and theirdefinitions and units, are summarized in Table 1). Together,rrgz defines the lithostatic (i.e., overburden) stress (Pa); forrr = 2600 kg m�3, hydrostatic fluid pressures correspond toa value of l = 0.38.[12] The fluid pressure ratio within the fault (lf) neces-

sary for failure along the SAF is a function of the shearstress (ty), the total normal stress on the fault plane (sy),and the fault zone friction coefficient (mf) (Figure 5):

lf ¼Pf

rrgz¼ sy � ty

mf

!=rrgz ð2Þ

where Pf is the fault zone pore fluid pressure.[13] Observations within the SAFOD pilot hole and

microearthquake focal mechanisms suggest that for the crustsurrounding the SAF, the overburden stress sv equals s3, theleast principal stress [Eberhart-Phillips and Michael, 1993;Hickman and Zoback, 2004]. Assuming the surroundingcrust is near critically stressed for failure [Townend andZoback, 2000], the shear stress (ty) and total normal stress(sy) on a near-vertical strike-slip fault with sv = s3 aredefined by

ty ¼ 1� lcð Þ2

rrgzM3

M1

�M3

M1

X þ X � 1

� �sin 2y ð3Þ

sy ¼ 1� lcð Þ M3

M1

sin2 y þ M3

M1

X �X þ 1

� �cos2 y

� �þlc

� �rrgz

ð4Þ

where lc is the fluid pressure ratio within the country rockand M1 and M3 are functions of the country rock frictioncoefficient, mc [Lachenbruch and McGarr, 1990],

M1 ¼ 1þ mcffiffiffiffiffiffiffiffiffiffiffiffiffiffi1þ m2

c

p !�1

ð5Þ

M3 ¼ 1� mcffiffiffiffiffiffiffiffiffiffiffiffiffiffi1þ m2

c

p !�1

ð6Þ


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and X ranging from 0 to 1 describes the magnitude of s2relative to s1 and s3 [Ramsay and Lisle, 2000]:

X ¼ s2 � s3

s1 � s3

ð7Þ

Analysis of borehole breakouts and drilling-induced tensilefractures within the SAFOD pilot hole show that X � 0.2 ata distance of 1.8 km SW of the surface trace of the SAF[Hickman and Zoback, 2004]. Laboratory experiments onsamples from the SAFOD pilot hole and surface outcropsreveal country rock friction coefficients of mc � 0.6 [Tembeet al., 2006; Carpenter et al., 2009]. Assuming the faultzone has a similar friction coefficient, such that mf = mc =0.6 (relaxation of this assumption is discussed in section 4)and values of X = 0.2, y = 85�, and lc = 0.38, the fluidpressure ratio in the fault zone required for slip on the SAF islf = 2.15 (equation (2)) (Figure 5). For a more conservative

value of y = 69�, the value determined from observationswithin the SAFOD pilot hole [Hickman and Zoback, 2004],the necessary fluid pressure ratio in the fault is lf = 1.59.Several studies have illustrated that total principal stresseswithin mature fault zones may be magnified to values greaterthan the surrounding country rock as a result of contrastingrheologic properties [Rice, 1992; Chery et al., 2004;Faulkner et al., 2006]. Such a stress state would allow forfluid pressures considerably greater than lithostatic to existlocally within the fault zone without causing hydrofracture[Rice, 1992; Chery et al., 2004; Faulkner et al., 2006].

2. Models of Coupled Fluid Flow and HeatTransport

[14] To assess whether realistic fluxes of mantle fluidhave the potential to generate large fluid pressures and

Table 1. Definition of Variables

Variable Parameter Units

%ss Equation (9); relates excess pressure from transientresults to steady state values

dimensionless

g Gravitational acceleration m s�2 (Lt�2)k Permeability m2 (L2)kc Country rock permeability m2 (L2)kcmax Country rock permeability in direction

of maximum permeabilitym2 (L2)

kdepth14 Permeability-depth relation with kcmax definedby equation (8).

N/A

kdepth15 Permeability-depth relation with kcmax 1 order of magnitudeless than defined by equation (8).

N/A

kf Fault zone permeability m2 (L2)kfmax Fault zone permeability in direction of maximum

permeabilitym2 (L2)

kcmin Country rock permeability in directionminimum permeability

m2 (L2)

M1 Equation (5); function of mc dimensionlessM3 Equation (6); function of mc dimensionlessP Pore pressure Pa (ML�1t�2)Pc Pore pressure in country rock Pa (ML�1t�2)P* Pore pressure above hydrostatic Pa (ML�1t�2)Pss* Pore pressure above hydrostatic at steady state Pa (ML�1t�2)

Pf Pore pressure in fault zone Pa (ML�1t�2)Phydro Hydrostatic fluid pressure Pa (ML�1t�2)X Equation (7); parameter describing s2

relative to s1 and s3dimensionless

q Heat flow mW m�2 (Mt�3)v Surface discharge cm a�1 (Lt�1)z Depth (units of meters, unless

otherwise noted)m (L)

l Equation (1); Fluid pressure ratio dimensionlesslc Country rock fluid pressure ratio dimensionlesslf Fault zone fluid pressure ratio dimensionlessm Friction coefficient dimensionlessmc Country rock friction coefficient dimensionlessmf Fault zone friction coefficient dimensionlessrr Bulk rock density kg m�3 (M L�3)s Normal stress Pa (ML�1t�2)s1 Maximum principal stress Pa (ML�1t�2)s2 Intermediate principal stress Pa (ML�1t�2)s3 Minimum principal stress Pa (ML�1t�2)sH Maximum horizontal principal stress Pa (ML�1t�2)sh Minimum horizontal principal stress Pa (ML�1t�2)sHf Maximum horizontal principal stress

within fault zonePa (ML�1t�2)

shf Minimum horizontal principal stress within fault zone Pa (ML�1t�2)sv Vertical principal stress (i.e., overburden stress) Pa (ML�1t�2)sy Fault zone total normal stress Pa (ML�1t�2)t Shear stress Pa (ML�1t�2)ty Shear stress on fault zone oriented at angle y to s1 Pa (ML�1t�2)y Angle between sH and fault zone degrees


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localize them onto a near-vertical fault, we incorporate amantle source of fluids in 2-D numerical models of coupledfluid flow and heat transport [e.g., Hanson, 1992, 1995,1997; Fulton et al., 2009]. We also evaluate whether thisscenario is consistent with heat flow and near-surfacehydrologic observations. Our model domain consists of a150 km long � 12 km thick 2-D cross section representativeof the seismogenic crust perpendicular to the SAF withinthe Parkfield region near SAFOD (Figures 6a– 8a; note thatFigures 6a–8a highlight only the 50 km to each side of thefault). The fault zone is located 50 km from the SW modelboundary. We prescribe a boundary condition of atmosphericpressure and temperature at the top of the model domain,and define the sides of the model domain as no-flowboundaries for fluid flow and heat transport. We assign aheat flux of 64 mW/m2 at the base of the model, and setradiogenic heat production at a constant value of 1.2 mW/m3

throughout the model domain, typical of measured valuesfor rocks within the area [Lachenbruch and Sass, 1980;Williams et al., 2004]. We prescribe thermal conductivity of2.5 W m�1 K�1 for rock and 0.6 W m�1 K�1 for water[Voss, 1984; Sass et al., 1997]. Together, the heat sourcesand thermal conductivities result in a near-surface heat flowof 78 mW m�2 (equivalent to the observed regionalaverage), and a temperature of �350�C at 12 km depthalong the model base, corresponding to the base of theseismogenic crust [Lachenbruch and Sass, 1980; Sass et al.,1997; Williams et al., 2004; Fulton et al., 2004b]. In thesemodels, we do not include any frictional heat production onthe fault.

2.1. Flux of Mantle-Derived Fluids

[15] On the basis of the inferred volume of serpentinizedmantle wedge beneath Cascadia and the age of the SAFsystem, we calculate a realistic estimate of the potentialdehydration fluid source. Assuming a 20 km thick wedgeof mantle material that is 50% serpentinized [Bostock etal., 2002; Kirby et al., 2002] and which fully dehydrates

over 25 Ma, we estimate an average water flux of 3.4 �10�9 kg s�1 m�2 of basal area. This is comparable to themaximum pulse of water (averaged over a 0.5 Ma interval)estimated from lower crustal dehydration but is sustainedfor 25 Ma [Fulton et al., 2004a, 2005, 2009]. Even largerfluid fluxes may be possible if dehydration of the serpenti-nized mantle occurs over a shorter time, but this wouldlikely allow any overpressures to dissipate within a fewhundreds of thousands of year after the sources are dimin-ished, as is the case for crustal dehydration [e.g., Fulton etal., 2009]. By choosing 25 Ma as the duration for dehydra-tion, our calculations incorporate the maximum sustainedfluid flux from dehydrating serpentinized mantle that wouldaffect locations along the entire strike of the SAF. Our studyfocuses solely on mantle-derived water, although additionalfluids such as CO2 may play a complementary role ingenerating overpressures [e.g., Bredehoeft and Ingebritsen,1990; Kharaka et al., 1999; Faulkner and Rutter, 2001].However, our estimated mass flux of mantle-derived wateris >12 times the estimated fluxes of mantle-derived CO2

inferred from geochemical observations near Parkfield[Kennedy et al., 1997; Kharaka et al., 1999; Faulknerand Rutter, 2001].[16] On the basis of observations that suggest the greatest

mantle fluid source is NE of the fault zone (see section 1.1),and the spatial extent of the zone of serpentinized mantlewedge material inferred beneath Cascadia (Figure 4)[Bostock et al., 2002], we prescribe the dehydration drivenfluid flux for 50 km along the model base, from the faultzone to the NE (Figures 6a–8a). The location of these fluidsources to the NE of the fault is also consistent with theexpected location of serpentinized mantle relative to theestimated location of the paleotrench �160 km SW ofthe SAF at Parkfield (Figure 4b) [Kirby et al., 2002]. Thefluids entering the model base are prescribed a temperatureof 350�C. We assign densities of 2600 kg m�3 for rock[Boness and Zoback, 2004] and 1000 kg m�3 for water at25�C, which changes by �0.375 kg m�3 �C�1 [Voss, 1984].

Figure 5. Mohr-Coulomb diagram illustrating the role of localized fluid pressures in explaining therelative weakness of the SAF [after Rice, 1992] scaled for mc = mf = 0.6, lc = 0.38, y = 85� and assumingcohesion is equal to zero. As mentioned in the text, differences in rheologic properties may allow totalprincipal stresses within the fault zone to be substantially higher than in the country rock [Rice, 1992;Faulkner et al., 2006]. The star represents the stress state on the SAF inferred from the orientation of thefault (y) with respect to themaximumhorizontal principal stress in the surrounding country rock (sH). SolidMohr circles represent stress states on vertical planes in the country rock (on the left) and within the fault zone(on the right). Failure envelopes (mc = mf = 0.6) are shown for a hydrostatic fluid pressure ratio in the countryrock lc = 0.38 and the fluid pressure ratio in the fault zone necessary for failure on the SAF (lf = 2.15).


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[17] We solve the equations for 2-D coupled fluid flowand heat transport using the finite element code SUTRA[Voss, 1984]. We run steady state simulations in order todetermine the maximum fluid pressures that could begenerated. We also run transient simulations for a subsetof our models, in order to illustrate the evolution of fluidpressure. In these transient simulations, we start with ahydrostatic initial condition to illustrate the potential forfluid pressure generation solely from mantle dehydration.These transient simulations incorporate values of rock andwater compressibility of 1 � 10�9 and 4 � 10�10 Pa�1

[Neuzil, 1986; Ge and Garven, 1992; Voss, 1984] and heatcapacities of 1000 and 4182 J kg�1K�1, for rock and waterrespectively [Taylor et al., 1982; Mossop and Segall, 1997;Voss, 1984]. Because total principal stresses within the faultzone may be magnified [e.g., Rice, 1992; Chery et al., 2004;Faulkner et al., 2006], and because our focus is to test thebasic hypothesis that mantle fluids have the potential tocontribute in localized overpressure development, our mod-els do not include a feedback on permeability and pressurethat may result from possible hydrofracture.

2.2. Permeability Architecture

[18] In our models, we consider a series of differentpermeability architectures for the crust. For the permeabilityof the country rock (kc), we assign maximum values

following a permeability-depth relation derived specificallyfor continental crust undergoing prograde metamorphism(termed kdepth14) [Manning and Ingebritsen, 1999]:

log kcmaxð Þ ¼ �14� 3:2 log zð Þ ð8Þ

where z is depth in kilometers. We also consider countryrock permeability 1 order of magnitude lower than thatdefined by equation (8) (kdepth15), which is also consistentwith the range of constraints used to construct this relation[Manning and Ingebritsen, 1999]. Porosity ranging from10�1 to �10�3 is defined by assuming a Kozeny-Carmanrelationship between permeability and porosity and followsa similar decrease with depth [Carman, 1956; Manning andIngebritsen, 1999]. Details regarding fault zone perme-ability and the degree and direction of permeabilityanisotropy for each suite of models are described insections 2.2.1–2.2.3.2.2.1. Fault Barrier[19] Observations in and around the SAFOD borehole

have been interpreted to indicate that the SAF acts as abarrier to fluid flow within more permeable country rock.Well data from both sides of the fault within the Parkfieldregion reveal a fluid pressure contrast across the fault, withsubhydrostatic pressures to the SW (lc 0.38) and

Figure 6. Model domain and results for scenarios with maximum country rock permeability orientedhorizontally, showing the 50 km to each side of the fault, although the model domain extends 50 kmfarther to the NE. (a) Schematic showing the orientation of country rock permeability and a 500 m widelow-permeability fault zone representing the SAF. A flux of mantle-derived water is prescribed along themodel base from the fault zone to 50 km NE (blue arrows). (b) Simulated steady state fluid pressure ratio(l) for country rock permeability defined by kdepth15 and an anisotropy ratio of 100:1. (c) Depth-averagedvalues of l for simulations with crustal permeability defined by kdepth14 (red) and kdepth15 (blue).(d) Simulated heat flow (q) taken at 100–200 m model depth for both simulations (kdepth14 = red; kdepth15 =blue); there is very little difference between the two. (e) Surface discharge (v) from both simulations; againthere is very little difference between the two.


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�12 MPa of overpressure at 1.5 km depth, 1.4 km to the NEof the fault (lc = 0.42) [Johnson and McEvilly, 1995;Zoback and Hickman, 2005]. Mud gas chemistry observa-tions during drilling of the SAFOD borehole include sharpcontrasts in the mantle helium signature and other fluidchemistry across the fault [Wiersberg and Erzinger, 2007;Erzinger and Wiersberg, 2007], as well as reduced gas fluxwithin the fault zone [Erzinger and Wiersberg, 2007].Geophysical logs, cuttings, and core observations are con-sistent with this conceptual model, as they reveal a clay-richprimary fault plane formed within a <650 m thick interval ofshale and siltstone [e.g., Solum et al., 2006] (Figure 2).[20] To simulate a fault acting as a barrier, we implement

a �500 m wide low-permeability zone in our models, withthe fault permeability (kfmax) 3 orders of magnitude lowerthan that of the country rock (kcmax). This corresponds to thedifference between permeability curves calibrated specifi-cally for brittle continental crust [Manning and Ingebritsen,1999] and laboratory data for clay-rich fault gouge thatdocument permeabilities of <10�21 to �10�19 m2 under�0–200 MPa of effective pressure [Faulkner, 2004]. As aninitial model scenario, we specify that the direction ofmaximum permeability in the country rock (kcmax) ishorizontal. For all scenarios, we specify the direction ofmaximum permeability within the fault zone (kfmax) to bevertical. The ratio of anisotropy for both the fault andcountry rock is prescribed a value of 100:1.2.2.2. Fault Barrier and NE Dipping Country RockAnisotropy[21] For a wide range of 2-D hydrogeologic and fault

zone structures, including models similar to our initial faultbarrier model described above, the ability to localize fluidpressure from a regional fluid source within a near-verticalfault zone has proved difficult [e.g., Fulton et al., 2009]. On

the basis of previous studies, which show that dippingpermeable strata can focus overpressures toward their updiptermini by a hydraulic condition known as the centroideffect [e.g., Flemings et al., 2002], we evaluate the role ofdipping hydrogeologic structures within the country rock asa potential means of localizing fluid pressures onto the faultzone. Rather than prescribing discrete permeable pathways,we assume the direction of maximum permeability in thecountry rock (kcmax) dips 30� to the NE. This correspondsto the direction of paleosubduction recorded within theFranciscan Complex and to the orientation of several largethrust faults and geologic structures within the FranciscanComplex and other geologic units NE of the fault zone(Figure 3) [e.g., Bailey et al., 1964; Wakabayashi, 1999;Dickinson, 2002; Guzofski et al., 2007]. Additionally, atthis orientation faults and fractures in the country rock arelikely to be critically stressed for failure (assuming mc �0.6 and with s1 nearly fault-normal), and as observed inother deep boreholes, these critically stressed faults andfractures may act as the most permeable pathways forfluid flow [Barton et al., 1995; Townend and Zoback,2000]. As in our other model simulations, we prescribethe ratio of anisotropy for both the fault and country rockas 100:1. We evaluate the sensitivity of our model resultsto the magnitude of anisotropy by conducting a subsetof simulations in which we maintain the value of themaximum permeability and allow the anisotropy ratio forboth the fault and country rock to vary between 5:1 and500:1.2.2.3. SAFOD-Specific NE Dipping Anisotropy and aSerpentinite Seal[22] In addition to model scenarios with NE dipping

permeability anisotropy in the country rock, which maygenerally represent much of the San Andreas Fault system

Figure 7. Model domain and results for simulations with maximum country rock permeability dipping30� to the NE. Descriptions are the same as in Figure 6.


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along its strike, we also evaluate a scenario specific to theParkfield region, where additional large-scale hydrogeo-logic structures have been inferred. In particular, in thisarea a subhorizontal serpentinite unit is thought to extend�10–15 km from the fault zone to the NE at a depth of�2–3 km [e.g., McPhee et al., 2004]. This unit is associ-ated with the Coast Range Ophiolite and has been mappedor inferred along much of the SAF system in northern andcentral California [Irwin and Barnes, 1975; Irwin, 1990].On the basis of geophysical observations and the distribu-tion of surface springs, the serpentinite has been interpretedto act as a low-permeability seal that promotes overpressuredevelopment beneath it [Irwin and Barnes, 1975; Eberhart-Phillips and Michael, 1993]. For a set of Parkfield-specificscenarios, we model this structure as a low-permeabilityzone (10�20 m2) [Tenthorey and Cox, 2003] from 3 to 4 kmdepth, extending from the fault zone 10 km to the NE(Figure 8).

3. Results

3.1. Pore Pressures

[23] Results of model simulations with a fault acting as abarrier and with maximum permeability anisotropy in thecountry rock oriented horizontally show that moderatesteady state fluid overpressures can be generated by arealistic mantle fluid source (Figure 6). For our low-perme-ability scenario (kdepth15), depth-averaged values of l reach0.61, with pressures along much of the base of the model tothe NE of the fault reaching values of lc � 1.10 (Figures 6band 6c). The zone of elevated fluid pressure is generallylimited to depths >5 km (Figure 6b); excess pore pressure

within the fault zone (Pf � Phydro) is �1 MPa at 2.7 kmdepth and 17.2 MPa at 6 km depth. These pore pressures areconsiderably larger than those generated by crustal dehy-dration; depth-averaged pore pressure ratios from models ofcrustal dehydration peak at values of l 0.46 for a similarpermeability architecture, and are short-lived owing to thelimited time over which sources are active [Fulton et al.,2009]. For our higher permeability scenario (kdepth14),depth-averaged overpressure is considerably smaller, withthe largest values of l � 0.40 (Figure 6c). Although thefluid pressure ratios generated in the low-permeabilityscenario are much larger than hydrostatic, they are not highenough to explain a weak SAF, and they are distributed overa broad zone NE of the fault and above the fluid sourceregion, rather then localized within the fault zone as neededto explain relative fault weakness (Figures 6b and 6c).[24] In model simulations with NE dipping permeable

pathways truncated by a low-permeability SAF acting as abarrier, we find that large fluid pressures can be focusedtoward the fault (i.e., localized) (Figure 7). Fluid pressureratios within the fault reach values considerably larger than forthe case in which maximum permeability in the country rockis horizontal (Figures 6b and 6c). For the low-permeabilityscenario (kdepth15), depth-averaged values of l are 0.66–0.71within 1 km NE of the fault plane (Figure 7). Along the edgeof the fault zone toward the model base, pore pressure ratiosreach values of lf � 2.0. Notably, the highest pore pressureratios occur within the 500 m wide low-permeability faultzone or straddle its NE edge (Figure 9a). Values of l to theNE of the fault diminish to near hydrostatic values (l < 0.4)over a distance less than �13 km from the fault (Figures 7b,7c, and 9). To the SW of the fault, there is an even sharper

Figure 8. Model domain and results for simulations representative of the Parkfield region aroundSAFOD, with maximum country rock permeability dipping 30� to the NE and a low-permeabilityserpentinite unit acting as a seal, extending from the fault zone 10 km NE at 3–4 km depth. Descriptionsare the same as in Figures 6–7.


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contrast in pore pressure, with near-hydrostatic values im-mediately on the SW side of the low-permeability fault zone(l < 0.4). The simulated excess pore pressure within the faultzone is 6 MPa at 2.7 km depth and 45 MPa at 6 km depth.Results of simulations for our high-permeability scenario(kdepth14) are similar in character, but with lower fluidpressure ratios; the maximum depth-averaged pore pressureratio for this scenario is l = 0.41 and occurs within the faultzone (compared with 0.71 for kdepth15; Figure 7c).[25] For the suite of models designed to represent the

hydrogeologic structure around Parkfield, California, andthe SAFOD site, the low-permeability serpentinite unit bothenhances the generation of elevated pore pressure, andlocalizes it within the fault zone, at all depths below�3 km (Figure 8). For the low-permeability case (kdepth15),depth-averaged overpressures peak at lf = �1.16 within thefault zone. Simulated pore pressures within the fault zonenear the model base are lf � 2.0–2.5, and l = �1.6 beneaththe serpentinite seal at �4 km depth within the region 500 mNE of the fault (Figures 8b and 8c). Excess pressure is8 MPa at 2.7 km depth and 160 MPa at 6 km depth(Figure 9b). For the high-permeability scenario (kdepth14),

depth-averaged values of lf are �0.46, maximum fluidpressure ratios near the base of the fault are lf = �0.58,and l = �0.50–0.51 beneath the serpentinite seal at 4 kmdepth, within the fault zone and extending to 2.5 km NE(Figures 8b and 8c).

3.2. Effect on Heat Flow and Surface Discharge

[26] Although large fluid pressures are localized towardthe fault zone in models that include dipping permeabilityanisotropy in the country rock, the associated fluid flow hasvery little effect on heat flow or on surface fluid discharge(Figures 6d–8d and 6e–8e). Simulated heat flow from thesemodel scenarios is marked by an increase of <10 mW m�2

near the fault (Figures 6e–8e). This spatial variability inheat flow is consistent with the range of observed heat flowscatter around the fault and is considerably less than the�40 mW m�2 near-fault anomaly expected from frictionalheating [e.g., Lachenbruch and Sass, 1980; Fulton andSaffer, 2009].[27] Similarly, simulated surface discharge exhibits only a

small increase near the fault. In simulations with NE dippingcountry rock anisotropy, the peak discharge is <1.0 cm a�1.This small rate of discharge is not sufficient to generate anoticeable seep or spring at the surface (Figure 7e), and isconsistent with the general lack of springs along the surfacetrace of the fault [Waring, 1965; Lachenbruch and Sass,1980; Kharaka et al., 1999]. For both simulated heat flowand surface discharge, there is very little difference betweenour high- and low-permeability scenarios (kdepth14 andkdepth15) (Figures 6e–8e).

3.3. Sensitivity to Degree of Anisotropy

[28] With larger permeability anisotropy in the countryrock, the effect of NE dipping permeable pathways ingenerating localized excess fluid pressures along the faultis more pronounced (Figure 10). For our low-permeabilityscenario (kdepth15), an anisotropy ratio of 10:1 generates apeak depth-averaged fluid pressure ratio at the fault of lf =0.43, whereas anisotropy ratios of 100:1 and 500:1 generatedepth-averaged lf values as high as 0.71 and 1.12, respec-tively. Although increased anisotropy results in larger fluidpressures on the fault, the associated fluid flow in all caseshas a small effect on near-surface heat flow (<10 mW m�2)(Figure 10a). Similarly, varying the permeability anisotropyvalue results in only slight differences in surface discharge(not shown; similar to Figure 7e); simulated surface dis-charge peaks near the fault at <1.4 cm a�1 for an anisotropyratio of 500:1 and at <1.0 cm a�1 for anisotropy of 100:1.As mentioned above, these small, simulated dischargevalues are consistent with the general lack of observedsurface springs along the fault trace.

3.4. Transient Results

[29] To investigate the evolution of fluid overpressures,we ran transient simulations for a subset of our models. Forour model scenario incorporating low crustal permeability(kdepth15), a fault barrier, and NE dipping maximum countryrock permeability (the model scenario which best representsthe known geology in the Parkfield and SAFOD region),large localized fluid pressures (similar to those illustrated inthe steady state results of Figure 7) develop within 2–3 Ma(Figure 11a). We report overpressures normalized to the

Figure 9. Detailed view of simulated fluid pressure near thefault zone from low-permeability scenarios for (a) dippinganisotropy case (e.g., Figure 7) and (b) Parkfield-specificcase (e.g., Figure 8). The lateral extent of the low-permeability fault zone is marked by vertical dotted lines.Labels for each profile describe the depth at which the resultshave been extracted and the peak excess pressure.


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steady state value, defined by the excess pore pressure (P �Phydro) at a given time divided by excess pressure at steadystate:

%ss ¼ P*=P*ss100 ð9Þ

By �2 Ma, depth-averaged pressures within the fault zonereach 50% of the steady state values (Figure 11b); thissuggests that a realistic mantle source would strongly affectpore pressures in the vicinity of the fault zone shortly afterthe passage of the MTJ and the establishment of the faultzone as a barrier to fluid flow. Fluid pressures within thefault zone continue to rise quickly, reaching >90% of the

steady state value by �10 Ma, and are maintained at theselevels until 25 Ma, the end of our simulation and thecessation of our fluid source, based on the assumptionsdescribed above (section 2.1).

4. Discussion

[30] Using a realistic estimate of mantle-derived fluid fluxand incorporating hydrogeologic structures consistent withobservations, our results illustrate the potential role ofmantle-derived fluids in causing both the absolute andrelative weakness of the SAF. Our results are unique in that(1) they provide a plausible mechanism for the generation ofsustained excess fluid pressures that are localized along anear-vertical fault zone, without the need to invoke a basalpressure boundary directly beneath the fault, complex faultpermeability behaviors, or transient processes such as shearcompaction or thermal pressurization [e.g., Byerlee, 1990;

Figure 10. Sensitivity of model results to the magnitudeof anisotropy in country rock permeability. Results areshown for models with NE dipping maximum permeabilityanisotropy. Maximum permeability is held constant for eachsimulation and defined by the lower fit to the relation ofManning and Ingebritsen [1999] (kdepth15), but theanisotropy ratio is varied. (a) Depth-averaged values of lfor anisotropy ratios of 5:1, 10:1, 50:1, 100:1, and 500:1. (b)Simulated heat flow for the same models as Figure 10a.There is very little difference between simulations in termsof simulated heat flow (Figure 10b) and fluid discharge (notshown).

Figure 11. Transient simulation results for the modelscenario shown in Figure 7, using the low-permeabilityrelation for the country rock (kdepth15). (a) Depth-averagedvalues of l as a function of distance from the fault, shownfor several time steps and for the steady state result.(b) Evolution of depth-averaged excess pore pressure withinthe fault zone from an initial hydrostatic condition,expressed as a percentage of the steady state value.


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Rice, 1992; Sleep and Blanpied, 1992; Andrews, 2002] and(2) they are consistent with observations from SAFOD andthe Parkfield region that suggest the SAF acts as a barrier tofluid flow [Zoback and Hickman, 2005; Solum et al., 2006;Wiersberg and Erzinger, 2007], with elevated fluid pres-sures observed in wells immediately NE of the SAF[Johnson and McEvilly, 1995; Zoback and Hickman, 2005],and with regional overpressures beneath a serpentinite unitextending NE from the SAF as inferred from high Vp/Vs

ratios [Eberhart-Phillips and Michael, 1993].[31] Although direct pore pressure measurements have

not yet been conducted in the SAFOD boreholes, measure-ments during drilling showed no evidence for significantlyelevated pore pressure within the fault zone [Zoback andHickman, 2005; M. Zoback and S. Hickman, personalcommunication, 2009]. However, our results demonstratethat fluid pressure ratios at the depth of the SAFODborehole may not fully reflect the fluid pressure state withinthe fault zone over much of its depth extent. For example,our results show that at �2.7 km (the depth at which theSAFOD main borehole crosses the fault zone), values ofboth fluid pressure ratio and excess fluid pressure within thefault zone should be considerably smaller than those at even6 km depth (excess pore pressures of 8 MPa versus160 MPa; Figure 9b). This difference depends on the depthof the borehole, and to a lesser extent on whether the modelincludes a low-permeability seal within the country rock NEof the fault zone (Figures 9a and 9b).[32] In a general sense, our results show that for realistic

2-D permeability architectures, mantle-derived fluid sourcescan generate large excess pore pressures. Additionally, ourresults illustrate that strongly anisotropic country rockpermeability resulting from NE dipping faults and fracturescan focus fluid pressures along a low-permeability faultzone, yet the associated fluid flow should have very littleeffect on heat flow and surface fluid discharge, causing<10 mW m�2 increase in heat flow and <1.0 cm a�1 ofsurface discharge near the fault. Fluid pressures along andimmediately to the NE of the fault may be quite large (lf >1.0 and lc > 0.7), and diminish to near-hydrostatic valuesby �10–20 km from the fault (Figures 8 and 9). In order forrealistic mantle fluid sources to generate excess fluidpressures localized along the SAF sufficient to explain itsapparent weakness, crustal permeability values near thelower bound of those defined by Manning and Ingebritsen[1999] are required (Figure 8). However, it is important tonote that additional processes may act in combination withmantle sources to drive elevated pore pressures. Pressuresgenerated by other regional geologic forcings such as tec-tonic compaction [Berry, 1973], crustal dehydration [Fultonet al., 2009], or CO2 release [Faulkner and Rutter, 2001]would also be focused toward the fault zone along dippingpermeable pathways and add to the effects of mantle dehy-dration sources.[33] Additionally, localized fluid pressures and low fric-

tion coefficient material within the fault zone may act incombination to account for geomechanical observations thatimply a weak fault in a much stronger crust. However, asillustrated in Figure 12 (derived from equations (2)–(4)), forall fault zone friction coefficients mf 0.3, fluid pressures inthe fault zone would need to be lithostatic or greater (lf 1)in order to explain an active fault plane oriented at 69� to s1

(y = 69�), regardless of the friction coefficient (mc) or porepressure ratio (lc) in the surrounding crust. To account foran orientation of y = 85�, as is inferred for the SanFrancisco Peninsula [Townend and Zoback, 2004], supra-lithostatic pressures are needed for all values of mf 0.07.For comparison, the lowest friction coefficient measured todate on SAFOD samples is m = 0.4, for cuttings from a fine-grained shear zone [Tembe et al., 2006]; measurements oncuttings from phases 1 and 2 of SAFOD drilling and fromrock outcrop samples within the Parkfield area documentvalues of mc > 0.55–0.6 [Tembe et al., 2006; Carpenter etal., 2009]. Weak minerals that may be present in the SAF,such as serpentinite or talc [Moore and Rymer, 2007],generally exhibit values of m = �0.1–0.2 when tested aspure monomineralic standards [Reinen et al., 1992; Mooreet al., 1997; Moore and Lockner, 2008; Carpenter et al.,2009]. Even with weak materials such as these in the faultzone, localized fluid pressures of lf = �1 would still beneeded to account for the steepest inferred orientations (y)between sH and the SAF (Figure 12).[34] In our simulations, the highest pore pressures occur

within the fault zone and to the NE (Figures 8 and 9).Although pore pressures in excess of the lithostatic pressureare possible if total principal stresses within the fault zoneare magnified owing to contrasts in rheologic properties[e.g., Rice, 1992; Faulkner et al., 2006], in a subset of ourmodels the simulation of similarly high pore pressures to theNE of the fault may be unrealistic, because total principalstresses would not be magnified in this region. Although ourmodels do not include a feedback on pressure and perme-ability from hydrofracture or fluid-induced failure, ourresults would not be affected substantially for two mainreasons. First, at stress states relevant over most of the depthextent of our models (>�3 km), supralithostatic fluid pres-sures would cause shear failure on preexisting faults strikingparallel to the SAF [Sibson, 1998], which would increasepermeability anisotropy and focusing of pressure onto thefault zone. Second, if rock failure or hydraulic fracturingoccur, this would effectively buffer the pore pressure in thecountry rock at or near a value of lc = �1, thus resultingin even greater localization of the highest pressures only onthe fault itself where magnified total principal stresses mayprevent hydrofracture.[35] Results from our transient model simulations illus-

trate that the focusing mechanisms described here couldhave an effect on fault zone fluid pressures soon after thecreation of the SAF (Figure 11). Simulated fluid pressureswithin the fault zone reach 50% of steady state valueswithin 2 Ma after the passage of the MTJ and the estab-lishment of the SAF as a hydrologic barrier and reach 80%of steady state values within 5 Ma. Assuming the MTJmigrates northward at a rate �40 km Ma�1, the mainte-nance of large localized fluid pressures for 25 Ma impliesthat mantle-derived water and the focusing mechanismsdescribed above could affect fault mechanical strength overa region encompassing nearly the entire length of the SAF,unlike the short-lived effects expected for crustal dehydra-tion [Fulton et al., 2009]. The southward increase in porepressures implied by our transient results is also consistentwith progressively increasing values of y (from 55� in thenorthern California immediately south of the MTJ, to 62� inthe San Francisco Bay area, to 74� just north of Parkfield)


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inferred from earthquake focal mechanisms in the crustsurrounding the SAF, which suggest progressive weakeningof the fault with age [Provost and Houston, 2003].[36] Our models show that NE dipping faults and fractures

may allow fluid pressures from a broad fluid source regionto the NE of the fault to be focused toward the fault zonethrough hydraulic transfer (i.e., the centroid effect). Thisshould result in a 20–30 km wide zone of slightly increasedheat flow (<10 mW/m2) with a peak near the fault (Figures7d, 8d, and 10b). Although a detailed evaluation of Hetransport is beyond the scope of this study, because thechemical diffusivity of He is at least 2 orders of magnitudelower than thermal diffusivity [e.g., Castro et al., 2005], weanticipate that transport of mantle helium would be stronglyaffected by advection along NE dipping permeable path-ways. Thus, we speculate that a mantle He signal would beubiquitous above and SW of the source region. This isconsistent with the observed sharp increase in the 3He/4Heratio across the SAF, with a mantle-derived component ashigh as 25% within 1.4 km NE of the SAF near Parkfield[Kennedy et al., 1997; Kharaka et al., 1999; Wiersberg andErzinger, 2007].

5. Conclusions

[37] Our models illustrate a plausible mechanism forlocalizing fluid pressures along the SAF, and in a mannerconsistent with inferred mechanical weakness along much

of its length. We illustrate the potential for a large source ofwater from the mantle and quantify the contribution of thissource to the generation of elevated fluid pressures withinthe seismogenic crust, and along the SAF in particular. Ourresults also draw attention to a potential mechanism forlocalizing fluid pressures along the fault. The SAF, or anyother major fault zone that acts as a flow barrier andtruncates dipping permeable pathways within the surround-ing country rock, could allow for the focusing of fluidpressure along the fault zone. Although our study focuseson a near-vertical SAF, similar localization of fluid pres-sures may also occur on dipping faults; for example withinaccretionary wedges where relative fault weakness hasalso been inferred [e.g., Byrne and Fisher, 1990], a low-permeability decollement or large-offset thrust splays maytruncate dipping strata that could transfer pressure fromgreater depths.[38] Our analysis also suggests that even with reasonably

low friction coefficient material, large localized fluid over-pressures may still be necessary to account for the steepinferred orientation of themaximum principal stress directionwith the fault (Figure 12). Overall, our results illustrate thatrealistic fluid sources and simple permeability architectureshave the potential to generate large, localized, and sustainedfluid pressures that could account for both the absolute andrelative weakness of the SAF and possibly other major plateboundary fault zones, and in amanner consistent with a broadrange of geological, geochemical, geomechanical and ther-

Figure 12. Fluid pressure ratio within the fault (lf) necessary for slip on the SAF, as a function ofassumed friction coefficient on the fault (mf), under a range of assumptions about pore pressure (lc) andfriction coefficient (mc) in the surrounding crust (from equations (2)–(4)). We assume that s3 is thevertical principal stress, X = 0.2 and mc = 0.6 (solid lines), mc = 0.4 (dashed lines), or mc = 0.2 (dottedlines). Red and blue lines represent solutions for y = 69� and y = 85�, respectively. Figures 12a, 12b, and12c show solutions for lc = 0.38, 0.6, and 0.8, respectively.


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mal data [e.g., Eberhart-Phillips and Michael, 1993; Fultonet al., 2004b; McPhee et al., 2004; Townend and Zoback,2004; Williams et al., 2004; Hickman and Zoback, 2004;Zoback and Hickman, 2005;Wiersberg and Erzinger, 2007].

[39] Acknowledgments. This project was supported by NSF grantsEAR 0545702 and EAR 0125189 to D.M.S. We thank B. Bekins andP. Flemings for helpful discussions in the early stages of this work andN. Sleep, C. Collettini, and an Associate Editor for constructive commentsand reviews.

ReferencesAndrews, D. J. (2002), A fault constitutive relation accounting for thermalpressurization of pore fluid, J. Geophys. Res., 107(B12), 2363,doi:10.1029/2002JB001942.

Atwater, T., and J. M. Stock (1998), Pacific-North America plate tectonicsof the Neogene southwestern United States: An update, Int. Geol. Rev.,40, 375–402, doi:10.1080/00206819809465216.

Bailey, E. H., W. P. Irwin, and D. L. Jones (1964), Franciscan and relatedrocks, and their significance in the geology of western California, Bull.Calif. Div. Mines Geol., 183, 177 pp.

Barton, C. A., M. D. Zoback, and D. Moos (1995), Fluid flow alongpotentially active faults in crystalline rock, Geology, 23, 683–686,doi:10.1130/0091-7613(1995)023<0683:FFAPAF>2.3.CO;2.

Berry, F. (1973), High fluid potentials in California Coast Ranges and theirtectonic significance, Am. Assoc. Pet. Geol. Bull., 57, 1219–1249.

Blakely, R. J., T. M. Brocher, and R. E. Wells (2005), Subduction-zonemagnetic anomalies and implications for hydrated forearc mantle,Geology,33, 445–448, doi:10.1130/G21447.1.

Bleibinhaus, F., J. A. Hole, T. Ryberg, and G. S. Fuis (2007), Structure ofthe California Coast Ranges and San Andreas Fault at SAFOD fromseismic waveform inversion and reflection imaging, J. Geophys. Res.,112, B06315, doi:10.1029/2006JB004611.

Boness, N. L., and M. D. Zoback (2004), Stress-induced seismic velocityanisotropy and physical properties in the SAFOD Pilot Hole in Parkfield,CA, Geophys. Res. Lett., 31, L15S17, doi:10.1029/2003GL019020.

Boness, N. L., and M. D. Zoback (2006), A multiscale study of themechanisms controlling shear velocity anisotropy in the San AndreasFault Observatory at Depth, Geophysics, 71, F131–F146, doi:10.1190/1.2231107.

Bostock, M. G., R. D. Hyndman, S. Rondenay, and S. M. Peacock (2002),An inverted continental Moho and serpentinization of the forearc mantle,Nature, 417, 536–538, doi:10.1038/417536a.

Bredehoeft, J. D., and S. E. Ingebritsen (1990), Degassing of carbondioxide as a possible source of high pore pressures in the crust, in TheRole of Fluids in Crustal Processes, edited by J. D. Bredehoeft and D. L.Norton, pp. 158–164, Natl. Acad. Press, Washington, D. C.

Brocher, T. M., T. Parsons, A. M. Trehu, C. M. Snelson, and M. A. Fisher(2003), Seismic evidence for widespread serpentinized forearc uppermantle along the Cascadia margin, Geology, 31 , 267 – 270,doi:10.1130/0091-7613(2003)031<0267:SEFWSF>2.0.CO;2.

Byerlee, J. (1990), Friction, overpressure, and fault normal compression,Geophys. Res. Lett., 17(12), 2109–2112, doi:10.1029/GL017i012p02109.

Byerlee, J. (1992), The change in orientation of subsidiary shears near faultscontaining pore fluid under high pressure, Tectonophysics, 211, 295–303,doi:10.1016/0040-1951(92)90066-F.

Byrne, T., and D. Fisher (1990), Evidence for a weak and overpressureddecollement beneath sediment-dominated accretionary prisms, J. Geophys.Res., 95(B6), 9081–9097, doi:10.1029/JB095iB06p09081.

Carman, P. C. (1956), Flow of Gases Through Porous Media, Academic,New York.

Carpenter, B. M., C. Marone, and D. M. Saffer (2009), Frictional behaviorof materials in the 3D SAFOD volume, Geophys. Res. Lett., 36, L05302,doi:10.1029/2008GL036660.

Castro, M. C., D. Patriarche, and P. Goblet (2005), 2-D numerical simula-tions of groundwater flow, heat transfer and 4He transport—Implicationsfor the He terrestrial budget and the mantle helium-heat imbalance, EarthPlanet. Sci. Lett., 237, 893–910, doi:10.1016/j.epsl.2005.06.037.

Chery, J., M. D. Zoback, and S. Hickman (2004), A mechanical model ofthe San Andreas Fault and SAFOD Pilot Hole stress measurements,Geophys. Res. Lett., 31, L15S13, doi:10.1029/2004GL019521.

DeShon, H. R., and S. Y. Schwartz (2004), Evidence for serpentinization ofthe forearc mantle wedge along the Nicoya Peninsula, Costa Rica,Geophys. Res. Lett., 31, L21611, doi:10.1029/2004GL021179.

Dickinson, W. R. (2002), Reappraisal of hypothetical Franciscan thrustwedging at Coalinga: Implications for tectonic relations along the GreatValley flank of the California Coast Ranges, Tectonics, 21(5), 1039,doi:10.1029/2001TC001315.

Eberhart-Phillips, D., and A. J. Michael (1993), Three-dimensional velocitystructure, seismicity, and fault structure in the Parkfield region, centralCalifornia, J. Geophys. Res., 98(B9), 15,737–15,758, doi:10.1029/93JB01029.

Erkan, K., and D. D. Blackwell (2008), A thermal test of the post-subductiontectonic evolution along the California transform margin, Geophys. Res.Lett., 35, L07309, doi:10.1029/2008GL033479.

Erzinger, J., and T. Wiersberg (2007), Characterization of gas from seismo-genic depths of the San Andreas Fault at SAFOD, Eos Trans. AGU,88(52), Fall Meet. Suppl., Abstract H11B-0486.

Faulkner, D. R. (2004), A model for the variation in permeability of clay-bearing fault gouge with depth in the brittle crust, Geophys. Res. Lett., 31,L19611, doi:10.1029/2004GL020736.

Faulkner, D. R., and E.H. Rutter (2001), Can themaintenance of overpressuredfluids in large strike-slip fault zones explain their apparent weakness?,Geology, 29 , 503 – 506, doi :10.1130/0091-7613(2001)029<0503:CTMOOF>2.0.CO;2.

Faulkner, D. R., T. M. Mitchell, D. Healy, and M. J. Heap (2006), Slip on‘weak’ faults by the rotation of regional stress in the fracture damagezone, Nature, 444, 922–925, doi:10.1038/nature05353.

Flemings, P. B., B. B. Stump, T. Finkbeiner, and M. Zoback (2002), FlowFocusing in Overpressured Sandstones: Theory, Observations, andApplications, Am. J. Sci., 302, 827–855, doi:10.2475/ajs.302.10.827.

Fulton, P. M., and D. M. Saffer (2009), The effect of thermal refraction onheat flow scatter near the San Andreas Fault, Parkfield, California,J. Geophys. Res., 114, B06408, doi:10.1029/2008JB005796.

Fulton, P. M., D. M. Saffer, and B. A. Bekins (2004a), Fluid overpressures onthe San Andreas Fault following the passage of the Mendocino TripleJunction,Eos Trans. AGU, 85(47), Fall Meet. Suppl., Abstract T41F-1285.

Fulton, P. M., D. M. Saffer, R. N. Harris, and B. A. Bekins (2004b), Re-evaluation of heat flow data near Parkfield, CA: Evidence for a weak SanAndreas Fault, Geophys. Res. Lett., 31, L15S15, doi:10.1029/2003GL019378.

Fulton, P. M., D. M. Saffer, and B. A. Bekins (2005), Crustal dehydrationand overpressure development on the San Andreas Fault, Eos Trans.AGU, 86(52), Fall Meet. Suppl., Abstract T43C-02.

Fulton, P. M., D. M. Saffer, and B. A. Bekins (2009), A critical evaluation ofcrustal dehydration as the cause of an overpressured and weak SanAndreasFault, Earth Planet. Sci. Lett., doi:10.1016/j.epsl.2009.05.009, in press.

Furlong, K. P., and S. Y. Schwartz (2004), Influence of the MendocinoTriple junction on the tectonics of coastal California, Annu. Rev. EarthPlanet. Sci., 32, 403–433, doi:10.1146/annurev.earth.32.101802.120252.

Ge, S., and G. Garven (1992), Hydromechanical modeling of tectonicallydriven groundwater flow with application to the Arkoma Foreland Basin,J. Geophys. Res., 97, 9119–9144, doi:10.1029/92JB00677.

Graeber, F. M., and G. Asch (1999), Three-dimensional models of P wavevelocity and P-to-S velocity ratio in the southern central Andes by simul-taneous inversion of local earthquake data, J. Geophys. Res., 104(B9),20,237–20,256, doi:10.1029/1999JB900037.

Guzofski, C. A., and K. P. Furlong (2002), Migration of the Mendocinotriple junction and ephemeral crustal deformation: Implications for Cali-fornia Coast range heat flow, Geophys. Res. Lett., 29(1), 1012,doi:10.1029/2001GL013614.

Guzofski, C. A., J. H. Shaw, G. Lin, and P. M. Shearer (2007), Seismicallyactive wedge structure beneath the Coalinga anticline, San Joaquin basin,California, J. Geophys. Res., 112, B03S05, doi:10.1029/2006JB004465.

Hanson, R. B. (1992), Effects of fluid production on fluid flow duringregional and contact metamorphism, J. Metamorph. Geol., 10, 87–97,doi:10.1111/j.1525-1314.1992.tb00073.x.

Hanson, R. B. (1995), The hydrodynamics of contact metamorphism,Geol. Soc. Am. Bull., 107, 595–611, doi:10.1130/0016-7606(1995)107<0595:THOCM>2.3.CO;2.

Hanson, R. B. (1997), Hydrodynamics of regional metamorphism due to con-tinental collision, Econ. Geol., 92, 880–891, doi:10.2113/gsecongeo.92.7-8.880.

Hickman, S. (1991), Stress in the lithosphere and the strength of activefaults, Rev. Geophys., 29, 759–775.

Hickman, S., and M. Zoback (2004), Stress orientations and magnitudes inthe SAFOD pilot hole, Geophys. Res. Lett., 31, L15S12, doi:10.1029/2004GL020043.

Irwin, W. P. (1990), Geology and plate tectonic development, in The SanAndreas Fault System, California, edited by R. E. Wallace, U.S. Geol.Surv. Prof. Pap., 1515, 61–80.

Irwin, W. P., and I. Barnes (1975), Effect of geologic structure and meta-morphic fluids on seismic behavior of the San Andreas Fault system incentral and northern California, Geology, 3, 713–716, doi:10.1130/0091-7613(1975)3<713:EOGSAM>2.0.CO;2.

Iwamori, H. (1998), Transportation of H2O and melting in subductionzones, Earth Planet. Sci. Lett., 160, 65 – 80, doi:10.1016/S0012-821X(98)00080-6.


14 of 15

B07408

Johnson, C. M., and J. R. O’Neil (1984), Triple junction magmatism: Ageochemical study of Neogene volcanic rocks in western California,Earth Planet. Sci. Lett., 71, 241 – 262, doi:10.1016/0012-821X(84)90090-6.

Johnson, P., and T. McEvilly (1995), Parkfield seismicity: Fluid-driven?,J. Geophys. Res., 100(B7), 12,937–12,950, doi:10.1029/95JB00474.

Kamimura, A., J. Kasahara,M. Shinohara, R. Hino, H. Shiobar, G. Fujie, andT. Kanazawa (2002), Crustal structure study at the Izu-Bonin subductionzone around 31�N: Implications of serpentinized materials along the sub-duction plate boundary, Phys. Earth Planet. Inter., 132, 105 – 129,doi:10.1016/S0031-9201(02)00047-X.

Kamiya, S., and Y. Kobayashi (2000), Seismological evidence for theexistence of serpentinized wedge mantle, Geophys. Res. Lett., 27(6),819–822, doi:10.1029/1999GL011080.

Kennedy, B. M., Y. K. Kharaka, W. C. Evans, A. Ellwood, D. J. DePaolo,J. Thordsen, G. Ambats, and R. H. Mariner (1997), Mantle fluids in theSan Andreas fault system, Science, 278, 1278–1281, doi:10.1126/science.278.5341.1278.

Kharaka, Y. K., J. J. Thordsen, W. C. Evans, and B. M. Kennedy (1999),Geochemistry and hydromechanical interactions of fluids associated withthe San Andreas Fault system, California, in Faults and Subsurface FluidFlow in the Shallow Crust, Geophys. Monogr. Ser., vol. 113, edited byW. C. Haneberg et al., pp. 129–148, AGU, Washington, D. C.

Kirby, S. H. (2000), Taking the temperature of slabs, Nature, 403, 31–34,doi:10.1038/47382.

Kirby, S., K. Wang, and T. Brocher (2002), A possible deep, long-termsource for water in the Northern San Andreas Fault system: A ghost ofCascadia subduction past?, Eos Trans. AGU, 83(47), Fall Meet. Suppl.,Abstract S22B-1038.

Lachenbruch, A. H., and A. McGarr (1990), Stress and heat flow, The SanAndreas Fault System, California, edited by R. E. Wallace, U.S. Geol.Surv. Prof. Pap., 1515, 261–277.

Lachenbruch, A. H., and J. H. Sass (1980), Heat flow and energetics of theSan Andreas fault zone, J. Geophys. Res., 85, 6185–6222, doi:10.1029/JB085iB11p06185.

Liu, M., and K. Furlong (1992), Cenozoic volcanism in the CaliforniaCoast Ranges: Numerical solutions, J. Geophys. Res., 97(B4), 4941–4951, doi:10.1029/92JB00193.

Manning, C. E., and S. E. Ingebritsen (1999), Permeability of thecontinental crust: The implications of geothermal data and metamorphicsystems, Rev. Geophys., 37, 127–150, doi:10.1029/1998RG900002.

McLaughlin, R. J., et al. (2000), Geology of Cape Mendocino, Eureka,Garberville, and southwestern part of the Hayfork 30 � 60 minute quad-rangles and adjacent offshore area, northern California, U.S. Geol. Surv.Misc. Field Stud., MF-2336.

McPhee, D. K., R. C. Jachens, and C. M. Wentworth (2004), Crustalstructure across the San Andreas Fault at the SAFOD site from potentialfield and geologic studies, Geophys. Res. Lett., 31, L12S03, doi:10.1029/2003GL019363.

Moore, D. E., and D. A. Lockner (2008), Talc friction in the temperaturerange 25�–400�C: Relevance for fault-zone weakening, Tectonophysics,449, 120–132, doi:10.1016/j.tecto.2007.11.039.

Moore, D. E., and M. J. Rymer (2007), Talc-bearing serpentinite and thecreeping section of the San Andreas Fault, Nature, 448, 795–797,doi:10.1038/nature06064.

Moore, D., D. Lockner, M. Shengli, R. Summers, and J. Byerlee (1997),Strengths of serpentinite gouges at elevated temperatures, J. Geophys.Res., 102(B7), 14,787–14,801, doi:10.1029/97JB00995.

Mossop, A., and P. Segall (1997), Subsidence at The Geysers GeothermalField, N. California from a comparison of GPS and leveling surveys,Geophys. Res. Lett., 24(14), 1839–1842, doi:10.1029/97GL51792.

Mount, V. S., and J. Suppe (1987), State of stress near the San AndreasFault: Implications for wrench tectonics, Geology, 15, 1143–1146,doi:10.1130/0091-7613(1987)15<1143:SOSNTS>2.0.CO;2.

Neuzil, C. E. (1986), Groundwater flow in low-permeability environments,Water Resour. Res., 22, 1163–1195, doi:10.1029/WR022i008p01163.

Provost, A.-S., and H. Houston (2003), Stress orientations in northern andcentral California: Evidence for the evolution of frictional strength alongthe San Andreas plate boundary system, J. Geophys. Res., 108(B3), 2175,doi:10.1029/2001JB001123.

Ramsay, J. G., and R. J. Lisle (2000), The Techniques of Modern StructuralGeology, vol. 3, Applications of Continuum Mechanics in StructuralGeology, Academic, London.

Reinen, L. A., T. E. Tullis, and J. D. Weeks (1992), Two-mechanism modelfor frictional sliding of serpentinite, Geophys. Res. Lett., 19(15), 1535–1538, doi:10.1029/92GL01388.

Rice, J. R. (1992), Fault stress states, pore pressure distributions, and theweakness of the San Andreas Fault, in Fault Mechanics and TransportProperties of Rocks, edited by B. Evans and T.-F. Wong, pp. 475–503,Academic, San Diego, Calif.

Sass, J. H., et al. (1997), Thermal regime of the San Andreas Fault nearParkfield, California, J. Geophys. Res., 102(B12), 27,575 – 27,585,doi:10.1029/JB102iB12p27575.

Schmitt, A. K., R. Romer, and J. Stimac (2006), Geochemistry of volcanicrocks from the Geysers geothermal reservoir, Californian Coast Ranges,Lithos, 87, 80–103, doi:10.1016/j.lithos.2005.05.005.

Sibson, R. H. (1998), Brittle failure mode plots for compressional andextensional tectonic regimes, J. Struct. Geol., 20, 655 – 660,doi:10.1016/S0191-8141(98)00116-3.

Sleep, N. H., and M. L. Blanpied (1992), Creep, compaction and the weakrheology of major faults, Nature, 359, 687–692, doi:10.1038/359687a0.

Solum, J. G., S. H. Hickman, D. A. Lockner, D. E. Moore, B. A. van derPluijm, A. M. Schleicher, and J. P. Evans (2006), Mineralogical charac-terization of protolith and fault rocks from the SAFOD Main Hole,Geophys. Res. Lett., 33, L21314, doi:10.1029/2006GL027285.

Taylor, R. E., R. L. Shoemaker, and H. Groot (1982), ThermophysicalProperties of Selected Rocks: A Report to U.S.G.S., TPRL 271, 32 pp.,Thermophys. Prop. Res. Lab., Purdue Univ., Indiana.

Tembe, S., D. A. Lockner, J. G. Solum, C. A. Morrow, T. Wong, and D. E.Moore (2006), Frictional strength of cuttings and core from SAFODdrillhole phases 1 and 2, Geophys. Res. Lett., 33, L23307, doi:10.1029/2006GL027626.

ten Brink, U. S., N. Shimizu, and P. C. Molzer (1999), Plate deformation atdepth under northern California: Slab gap or stretched slab?, Tectonics,18, 1084–1098, doi:10.1029/1999TC900050.

Tenthorey, E., and S. F. Cox (2003), Reaction-enhanced permeability dur-ing serpentinite dehydration, Geology, 31, 921–924, doi:10.1130/G19724.1.

Tibi, R., D. A. Wiens, and X. Yuan (2008), Seismic evidence for wide-spread serpentinized forearc mantle along the Mariana convergence mar-gin, Geophys. Res. Lett., 35, L13303, doi:10.1029/2008GL034163.

Townend, J., and M. D. Zoback (2000), How faulting keeps the cruststrong, Geology, 28, 399 – 402, doi:10.1130/0091-7613(2000)28<399:HFKTCS>2.0.CO;2.

Townend, J., and M. D. Zoback (2004), Regional tectonic stress near theSan Andreas Fault in central and southern California, Geophys. Res. Lett.,31, L15S11, doi:10.1029/2003GL018918.

Unsworth, M., and P. A. Bedrosian (2004), Electrical resistivity structure atthe SAFOD site from magnetotelluric exploration, Geophys. Res. Lett.,31, L12S05, doi:10.1029/2003GL019405.

Voss, C. I. (1984), A finite-element simulation model for saturated-unsaturatedfluid density-dependent groundwater flow with energy transport or chemi-cally reactive single-species solute transport, U.S. Geol. Surv. Water Resour.Invest., 84-4369.

Wada, I., K. Wang, J. He, and R. D. Hyndman (2008), Weakening of thesubduction interface and its effects on surface heat flow, slab dehydration,and mantle wedge serpentinization, J. Geophys. Res., 113, B04402,doi:10.1029/2007JB005190.

Wakabayashi, J. (1999), Subduction and the rock record: Concepts developedin the Franciscan Complex, California, in Classic Concepts in CordilleranGeology: A View From California, edited by E. M. Moores, D. Sloan, andD. L. Stout, Spec. Pap. Geol. Soc. Am., 338, 123–135.

Waring, G. A. (1965), Thermal springs of the United States and othercountries of the world, U.S. Geol. Surv. Prof., 492.

Whitlock, J. S. (2002), Evidence of a mantle wedge source for slab windowvolcanism in the northern California Coast Ranges,Master’s thesis, 79 pp.,Pa. State Univ., Univ. Park.

Whitlock, J. S., K. P. Furlong, C. E. Lesher, and T. Furman (2001), TheJuan de Fuca Slab-window and Coast Range volcanics, California: Cor-relation between subducted slab age and mantle wedge geochemistry, EosTrans. AGU, 82(47), Fall Meet. Suppl., Abstract T32-03.

Wiersberg, T., and J. Erzinger (2007), A helium isotope cross-section studythrough the San Andreas Fault at seismogenic depths, Geochem.Geophys. Geosyst., 8, Q01002, doi:10.1029/2006GC001388.

Williams, C. F., F. V. Grubb, and S. P. Galanis Jr. (2004), Heat flow in theSAFOD pilot hole and implications for the strength of the San AndreasFault, Geophys. Res. Lett., 31, L15S14, doi:10.1029/2003GL019352.

Zoback, M. D., and S. H. Hickman (2005), Preliminary observations ofstress and fluid pressure in and near the San Andreas Fault at depth in theSAFOD boreholes, Eos Trans. AGU, 86(52), Fall Meet. Suppl., AbstractT21A-0438.

Zoback, M. D., M. L. Zoback, J. P. Eaton, V. S. Mount, and J. Suppe(1987), New evidence on the state of stress of the San Andreas Fault,Science, 238, 1105–1111, doi:10.1126/science.238.4830.1105.

��P. M. Fulton, College of Oceanic and Atmospheric Sciences, Oregon

State University, 104 COAS Building, Corvallis, OR 97331, USA.([email protected])D. M. Saffer, Department of Geosciences, Pennsylvania State University,

310 Deike Building, University Park, PA 16802-0000, USA.


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