Magmatic processes in the Alaska subduction zone

by combined 3-D b value imaging and targeted

seismic tomography

T. van Stiphout,1 E. Kissling,2 S. Wiemer,1 and N. Ruppert3

Received 28 July 2008; revised 26 February 2009; accepted 5 August 2009; published 3 November 2009.

[1] We combine two complementary seismological methods to study the deep roots of arcvolcanism in the Alaskan subduction zone: targeted seismic body wave tomography andimaging of the relative size distribution (b value) of earthquakes. For the tomography weapply a staggered inversion scheme, starting with the minimum 1-D velocity model,progressing to a 3-D velocity model. Inversions are based on traveltime data from about5500 events (1971–2002) with approximately 100,000 P and 50,000 S high-qualityarrivals, allowing us to resolve the 3-D velocity field on a 20 � 20 km grid. For the bvalue imaging we have introduced a new 3-D sampling approach to map the b values onsurfaces in the Wadati-Benioff zone (WBZ). Our b value study is based on 12,474relocated events of magnitude of completeness �2.4 and depth >40 km. We observeanomalously high b values at depths around 100 km, appearing locally concentratedbeneath active subduction volcanoes. We believe that these zones are associated withdehydration of the slab and subsequent magma generation. The analysis of VP/VS ratioslikewise shows indications of the presence of fluids: significantly higher VP/VS ratiosrising in columns from the top of the WBZ at 100 km depth below active volcanic centers.On the basis of the joint interpretation, we propose that our images track fluids from theirgenesis at 100 km depth at the top of the subducting plate to the bottom of the crust belowvolcanoes.

Citation: van Stiphout, T., E. Kissling, S. Wiemer, and N. Ruppert (2009), Magmatic processes in the Alaska subduction zone by

combined 3-D b value imaging and targeted seismic tomography, J. Geophys. Res., 114, B11302, doi:10.1029/2008JB005958.

1. Introduction

[2] Kirby et al. [1996] and, more recently, van Keken[2003] summarized the current understanding of arc magmageneration that is based on a rather limited amount of in situobservables. The magmatic oceanic crust becomes fissuredalready near mid-ocean ridges by normal faulting. Ranero etal. [2003] document evidences for bending-related faultingat the trench that allows water to penetrate down to theoceanic mantle. Later during the subduction process, thesefissures filled with hydrothermal fluids and hydrous mineralsas well as the sediments on top are reactivated and dehy-drate progressively. Gill [1981], Syracuse and Abers [2006],and Tatsumi and Eggins [1995] found that independent fromthe type of subduction zone, the magmatic front, i.e., arcvolcanoes, lies 124 ± 38 km above the Wadati-Benioff zone(WBZ). This depth is presumed to be the source region ofarc volcanism. The process of slab dehydration is regardedas the major process of magma generation (e.g., see Ulmer[2001] for an overview). The flux of slab-released water

into the overlying mantle wedge lowers the solidus of theperidotite and furthers development of partial melts. Con-sequently, melts or aqueous fluids will penetrate upwardinto the crust of the overlying plate. Depending on mantlewedge dynamics, these melts may either flow vertically orbe deviated by mantle wedge flow [Wyss et al., 2001;Manea et al., 2005]. Direct geophysical evidence for slabdehydration at depth beneath arc volcanoes and its relationto melt in the mantle wedges and fluid conduits through theoverriding lithosphere, however, remain sparse.[3] To improve the understanding of generation and path

of arc magma, we apply targeted high-resolution body wavetomography and frequency-magnitude distribution analysis(FMD) of local earthquakes, also called b value study[Wiemer and Wyss, 2002], to illuminate in detail thestructure from slab top to arc volcanoes. By combiningthese two independent methods, we are able to constrainpossible physical processes leading to magma generation.Seismic tomography provides information mainly on thepath of magma but also partly on the source (we can resolveseismic velocities along the seismic ray paths that cover thepotential path of magma and partly the potential location ofits source in the slab), whereas b value imaging of the WBZfocuses on the source of the magma (b values revealinformation on physical properties at the potential magmasource). Central Alaska, with its high seismicity recorded

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1Swiss Seismological Service, ETH Zurich, Zurich, Switzerland.2Institute of Geophysics, ETH Zurich, Zurich, Switzerland.3Geophysical Institute, University of Alaska Fairbanks, Fairbanks,

Alaska, USA.

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over several decades by an extensive station network and, inparticular, with stations in the back arc, represents an ideallocation for such a joint analysis (Figure 1).[4] Several previous tomographic studies in central

Alaska provided a wealth of information on the subductionzone, though each had its specific focus and limitations.Kissling and Lahr [1991] defined lithosphere thickness ofthe two plates and resolved details of 20 km minimal extentwithin the subducting lithosphere and within the overridinglithosphere. Zhao et al. [1995] investigated the regional 3-DVP structure down to 200 km on a 30–60 km grid, andEberhart-Phillips et al. [2006] combined local earthquakeand active source data in their inversion of P and S wavedata to provide a detailed image mainly on the overridinglithosphere. In our tomography study, we combine regionaland local tomography of P and S body waves targeted atmagma source and path regions, where the term ‘‘targeted’’means that the data selection is optimized to achieve thehighest possible resolution in the area of interest, i.e., apossible source and path of magma.[5] So far, there have been only a few studies of detailed

b value mapping of microseismicity targeted to the sourceregion of arc volcanism. Wiemer and Benoit [1996] foundsignificant changes in b value at depths of 90–100 km onthe upper side of the WBZ in New Zealand, Cook Inlet, andthe Shumagin Island region. Wyss et al. [2001] found a highb value at 150 km depth in the subduction zone beneath theTohoku area (northern Japan). Schorlemmer et al. [2003]interpreted a high b value anomaly 90–200 km below theAeolian volcanic district as related with the origin of themagma. The study of Wyss et al. [2001] compares b valueanalysis of microseismicity with published results of seis-mic body wave tomography for the Tohoku region, suggest-ing that the b value anomaly at 150 km is connected to alow S wave velocity anomaly originating at that depth. Thiswould suggest that, in fact, magma genesis takes places atgreater depth, and the magma travels a curved path throughthe mantle wedge.[6] Following a brief summary of the tectonic framework,

in section 2, we discuss the seismic data available forAlaska and the selected data used in this study. In section 3we present the method and resolution of the b values studyand show the 2-D and 3-D b value structure of the WBZbased on relocalized earthquakes. Third, in section 4 wesummarize the method for targeted local earthquake tomog-raphy including calculation of the minimum 1-D velocitymodel and estimation of the resolution and present thetomographic results for VP and VS fields of the southernAlaska subduction zone. In sections 5 and 6, the interpre-tation of combined results of b value and tomography studyare discussed.

2. Seismicity and Tectonics of Southern Alaska

[7] Southern Alaska is part of the ‘‘ring of fire,’’ acircum-Pacific zone of high seismic, volcanic, and tectonicactivity that is associated with plate boundaries. Here, thePacific plate, which has an age of about 47 Ma at the trench,is obliquely subducting with a velocity of about 50 km/Mabeneath the continental North American plate [DeMets etal., 1990; Syracuse and Abers, 2006]. Farther west alongthe Aleutian island arc, the oceanic Pacific plate subducts

under the oceanic North American plate. Active volcanismspreads from the western Aleutian toward Mount Spurr andMount Hayes to the east (Figure 1).[8] Tectonics and seismicity of southern Alaska are

dominated by processes of plate convergence (5–7 cm/yr).The Alaskan crust is a mosaic of a few oceanic and variouscontinental crustal fragments that have been accreted to thecontinental margin of North America during the past 200Ma [Ambos et al., 1995] superimposed by volcanic arcintrusions and deposits. The major faults in continentalAlaska include the strike-slip Denali, Kaltag, Tintina, andCastle Mountain faults (Figure 1). The crustal seismicity,however, is distributed over a much wider area with activeseismic zones and clusters located in virtually all parts of thestate. A major part of seismicity in Alaska is located in theWBZ of the subducting slab. At its northeast end, thesubducting plate reaches depths of approximately 200 kmas documented by the deepest earthquakes and earliertomographic studies [Kissling and Lahr, 1991]. In southernAlaska, earthquakes with magnitudes greater than 5 occur ata rate of about five events per year. Active volcanoes on themainland are mainly aligned above the 100 km deep WBZ.The target of this combined b value and tomographic studyis the WBZ and the mantle wedge below the arc volcanoes(Figures 1 and 2).

2.1. Local Earthquake Data from Southern Alaska

[9] The Alaska regional seismic network comprises near-ly 400 sites. The Alaska Earthquake Information Center(AEIC) serves as the center for integration of all seismicnetworks within Alaska, including stations of the AlaskaVolcano Observatory (AVO), the West Coast and AlaskaTsunami Warning Center (WC/ATWC), and few sites of theIncorporated Research Institute for Seismology (IRIS). Asof the end of the 1990s, the majority of the stations wereshort-period vertical instruments, with a handful of three-component short-period stations. Beginning in 1998, theAEIC, AVO, and WC/ATWC have been upgrading somesites to the broadband instrumentation with digital teleme-try. In 2006, about 50 broadband sites were operational inthe network.[10] In this study, we used 145 stations in central Alaska

(between 59�N and 64�N and 146�Wand 156�W, Figure 1).These stations provide good coverage of the study regionwith seismic rays between hypocenters and seismic stations.Stations without any readings were ignored.[11] To maximize resolution by joint interpretation of

targeted seismic tomography and b value study, we usethe same database for both methodologies. The data selec-tion criteria based on the requirements of the two methodsare explained in sections 2.2 and 2.3. We will demonstratein sections 3.2 and 4.2 that these data provide similar spatialresolution for b value anomalies and P and S wave velocity(VP and VS) anomalies in the WBZ at the depth of potentialmagma source regions. Additionally, the VP, VS, and VP/VS

fields can be mapped in the mantle wedge. Figure 2 outlinesthe regional expected resolution capabilities of b values andseismic tomography studies in subduction zones.

2.2. Seismic Data for b Value Study

[12] For b value studies, relative hypocenter locationaccuracy and consistency are crucial quality criteria. Of

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particular interest are systematic shifts in hypocenter loca-tions as a function of magnitude, which can lead to artifactsin b value distribution [Wiemer and Wyss, 2002], whichagain may bias the physical interpretation. Consistency ofmagnitude reporting and a good knowledge of the com-pleteness of reporting are also critical. Therefore, the hypo-centers used in the b value study are relocated using aminimum 1-D velocity model [Kissling et al., 1994] basedon traveltime data between 1971 and June 2002. We preferthe minimum 1-D velocity model with station delays forrelocation over the use of the calculated 3-D velocity modelbecause the former represents a global root-mean-squareminimum for all data, whereas the quality of the 3-Dvelocity model varies strongly and thus may locally biassome hypocenters. The calculation of the minimum 1-Dvelocity model is described in section 4.1. Studies of b value

require homogeneity of reporting with time. Such homoge-neity is only guaranteed for the AEIC catalog after May1989 [Ratchkovski et al., 2004]. Since our study focuses onb values in the slab, we use hypocenters with depths, z,larger than 40 km and thus selected a data set for b valuecalculation of 12,474 events (1989–2006).

2.3. Seismic Data for Tomographic Study

[13] For uniformly high spatial resolution in the targetvolume, an approximately homogeneous ray coverage isneeded [Kissling et al., 2001]. Generally, the smaller theazimuthal gap and the more observations (consistentlytimed with consistent error assessment) an earthquakerecord has, the higher the quality of this event for seismictomography. This demand on quality regards P and S wavetraveltimes alike, since the latter add important additional

Figure 1. Overview on study region of Cook Inlet, central Alaska. Inset is location of Cook Inlet inAlaska. Black triangles indicate stations of the AEIC network. Red triangles indicate named volcanoes.Hypocenters are indicated by gray points, and crustal faults are indicated by thin solid black lines. Blackcrosses indicate the horizontal grid used for seismic tomography. The two solid black lines indicatelocations of the cross sections (AA0 is 154.01�W, 60.88�N to 150.50�W, 60.37�N, and BB0 is 153.45�W,61.68�N to 149.81�W, 61.21�N) perpendicular to the strike of the subduction zone. The direction and thevelocity of the incoming plate are given in the lower right corner. The depth contour lines of the WBZ areindicated by the black dashed lines.

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constraints to the earthquake location problems, e.g., bydetermining focal depths [Gomberg et al., 1990]. Therefore,selection of consistent high-quality P and S arrivals isequally important for hypocenter locations and high-preci-sion seismic tomography [Kissling, 1988].[14] We benefit from a consistent high-quality data set

provided by the AEIC. The data set on which our tomog-raphy study is based contains 83,546 events betweenOctober 1971 and June 2002, with about 7 million P andS observations of all Observation Quality Classes (OQC).AEIC classifies observations into 5 OQC, 0 being highestand 4 denoting lowest. To guarantee the requested qualityfor our study, we only use phase readings of OQC 0 and 1,accepting thereby the associated errors of ±0.1 s and ±0.5 s,respectively. In the case of poorly to fairly well resolvedcells, the maximum velocity uncertainty is about 4%, andfor well-resolved cells it is less than 2%. With localearthquake tomography we seek solutions to the coupledhypocenter-velocity problem. The quality of tomographic

images also depends on location quality. By including onlyevents with an azimuthal gap smaller than 160� and at leasteight P and eight S observations, we ensure high data quality,thereby further reducing the data set for tomography to about5526 events with 98,875 P and 53,006 S observations.

3. The b Values of Local Earthquakes in theSubducting Slab of Southern Alaska

3.1. Method

[15] The most critical parameter when estimating b valuesis the magnitude of completeness, MC [Wiemer and Wyss,2002]. We follow the approach of Woessner and Wiemer[2005] and calculate MC with the maximum curvaturemethod [Wiemer and Wyss, 2000] in combination withbootstrapping. We further add a correction term of 0.2, alsotaking into account the fact that changes in detectionthreshold and homogeneity in magnitude reporting havebeen documented for the AEIC data set [Zuniga and

Figure 2. Sketch of a subduction zone with volcanic arc. It shows the resolution potential of b valuestudy (dark gray shaded area) and seismic tomography (light gray shaded area), including hypocenter(black dots) and ray paths (gray line) to the stations (black triangles). The geometry of the WBZ iscalculated by applying a vertical cylinder of 25 km radius (see text). The rectangles show the typicalresolution for 3-D tomography. The horizontal cylinder and sphere indicate typical sampling volumes forb value.

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Wiemer, 1999; Ratchkovski et al., 2004]. This yields aMC of2.4, and therefore, we cut the catalog at this value. Wecomputed maximum likelihood b values according to Utsu[1965] and Aki [1965], and uncertainties in b are based onthe formulation of Shi and Bolt [1982].[16] Previous mapping of b values in subduction zones

has primarily used cross sections perpendicular to the strikeof the subduction zone [Wiemer and Benoit, 1996; Wyss etal., 2001]. Events within a certain width (generally 80–150km) are projected onto this cross section, and b values aredetermined for circular volumes with either constant radii orconstant sample size (see Wiemer and Wyss [2002] for adiscussion of the merits of either approach). We want tocalculate the 3-D distribution of the b value on differentsurfaces of the WBZ. The previous approach is welljustified given that the processes we are interested in areassumed to be rather a function of depth and position in theslab than a function along strike. This can potentially biasour analysis of the 3-D structure of the Alaskan subductionzone. Also, it is not readily comparable with the 3-Dtomographic approach simultaneously pursued in this study.To overcome these limitations and to study for the first timethe 3-D (or lateral) variability in b in the WBZ, we devisefor this study a new mapping approach.[17] To perform a 3-D mapping of b values projected onto

the subducting slab in central Alaska, we initially use therelocated seismicity and remove crustal events (z < 40 km)from the earthquake catalog to define the slab geometry.Next we compute the depth distribution of hypocenterswithin vertical columns with radii of 25 km, centered atnodes spaced 10 � 10 km. Each column must contain atleast 50 earthquakes to define the slab geometry. Note thatthe sampling radii overlap, providing an intended smooth-ing effect of slab contours. The percentiles (e.g., the 10th or90th of the depth distribution define the shape of the slabsurface in a Cartesian coordinate system. The samplingapproach is sketched in Figure 2.[18] To map the b values, we use two different ways of

sampling at a given node of a selected depth contour (e.g,we use the 10% and 90% depth contour for the top and thebottom of the slab): (1) one way is to use spherical volumesof constant radii or constant number; this allows for ageometrically unbiased sampling; and (2) another way isto use cylindrical volumes; at each node a cylindricalvolume sample with constant number or constant radiussampling, whereas the axes of the cylinders are orientedalong the strike of the surface in the WBZ.[19] Additionally, we analyzed a series of cross sections

placed perpendicular to the strike. The cross sections arelying in the direction of the maximum dip. To compute bvalues, we project the seismicity within 50 km on each sideof the cross-sectional plane and compute b values for the100 nearest hypocenters to each node on a grid spaced 2 �2 km. We only plot events within a maximal cylinder radiusof 15 km. The locations of these cross sections are depictedin Figure 1.[20] The stability of the solution in the cross sections was

tested by (1) rotating the orientation of the cross section byup to 10�, (2) varying MC, and (3) varying the width of thecylinder. The stability of our new sampling approach wastested by varying MC and using different sampling volumes

(cylindrical or spherical volumes) of varying size. Theb values reported do not critically depend on our choice ofMC estimation, time period selection, or 3-D samplingapproach, e.g., cylindrical or spherical sampling. Withcylindrical sampling volumes having an axis along strike,these anomalies tend to smear along strike. These linea-ments can be seen northward up to 61.0�N (Figure 3).

3.2. Resolution

[21] The lateral resolution of the b value mapping on theslab is defined by the number of events within verticalcylinders with radii of 25 km. These cylinders are used todetermine the depths of different surfaces in the slab. Thereare no b values determined if there are less than 50 eventswithin a vertical cylinder.[22] For b value calculation on the surface of the WBZ,

we use 100 events within a horizontal cylinder with max-imum radius of 25 km. The typical horizontal cylinderradius in our target region is between 5 and 10 km(Figure 3d). The maximum radius of a spherical volumeis 30 km and typically between 15 and 20 km in the targetregion. Thus, the resolution depends on the density of thelocal seismicity, and the b values are representative for thevolume. In our target region and on the WBZ plane, we mayresolve b value anomalies of about 10% (see section 3.3)with an extension of about 40 km.[23] Figure 3 shows the resolvable regions in the WBZ.

Between 59�N and 60.75�N we have good resolution at the100 km depth contour line, northward; between 61.0�N and63.0�N, the resolution is too poor for results to be inter-preted. North of 63.0�N, where the dip direction of the slabis changing toward north, seismicity allows resolution ofhigh b values along the 100 km depth contour line. The lackof b values at about 62.3�N and 150.3�W is caused bysparse seismicity in this area, i.e., not enough events in thevertical columns to determine the depth levels of the slab.

3.3. Results

[24] First, we have analyzed b values interactively on twocross sections perpendicular to the strike of the slab (AA0

and BB0, in Figures 1, 4a, and 4b). The cross sections AA0

and BB0 are separated by 100 km. The upper limit of thesubducting lithosphere is defined by the WBZ, and weassume a 60 km thick oceanic lithosphere (dashed lines)according to the results of Kissling and Lahr [1991]. Thewhite areas in the b value cross section are regions wherethe mapping radius exceeds 15 km. Cross section AA0

(Figure 4a) shows a strong high b value anomaly (red)between 100 and 120 km depth, located below Redoubtvolcano. Cross section BB0 (Figure 4b), located north of thelast active volcano of the eastern edge of the Aleutian,Mount Spurr, however, does not show a high b valueanomaly at 100 km depth. To document the significanceof local anomalies, for two selected nodes (marked ascircles) we obtain b values of b = 1.47 ± 0.14 and b =0.98 ± 0.1. Following Utsu [1992], these two samples aresignificantly different with a probability of over 99%.[25] Results presented in Figures 3 and 4 document a

strong lateral lineament of significantly increased b valuesat a depth of 90–100 km that coincides with the overlyingvolcanoes and suggests a depth-related mechanism. Because

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Figure 3. A b value map view of the WBZ. The b values on the top layer of the WBZ are based onvolume sampling that is (a) cylindrical and (b) spherical. The b values on the bottom layer of the WBZare based on (c) cylindrical volume sampling. (d) The resolution (cylinder radius) of Figure 3a. The inletin Figure 3a indicates the sampling on top and bottom of the WBZ. The b value anomaly is marked with ablack polygon. The coastline of central Alaska is shown in gray, the depth contour lines of the WBZ aredashed and indicated with the associated depth, and the volcanoes are indicated with red triangles. Notethat only regions where both the depths of the slab as well as the b values are resolved are plotted.

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Figure 4. For the northern (AA0) and southern (BB0) cross section (Figure 1), (a and b) b value, (c andd) absolute VP, and (e and f) VP/VS ratio. The solid black line in cross section AA0 indicates theapproximate extension of the b value anomaly in Figure 4a. The location along the cross section isindicated by a red triangle for volcanoes within ±50 km from the cross section and by a black triangle forstations within ±100 km from the cross section. Black dots represent hypocenters within ±50 km from thecross section (relocalized with the minimum 1-D velocity model for b value in Figures 4a–4b and the 3-Dvelocity model for seismic tomography in Figures 4c–4f). The dashed lines indicate the upper and lowerboundaries of the oceanic lithosphere. The b values are calculated on the basis of cylindrical samplingvolumes perpendicular to the cross section, centered at the nodes, and with a cylinder length of 100 km.The grid nodes are equally spaced 2 � 2 km. Number of earthquakes per sample, N, is 100. The b valueof samples indicated by circles are shown as insets in Figures 4a and 4b. For absolute VP and VP/VS

ratio, the VP and VP/VS structures for the cross sections are calculated as described in the text. Whiteregions are unresolved.

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of sparse seismicity in some regions, we cannot map thisdepth layer for the whole area. However, additional crosssections show that the anomaly is present over a longdistance up to 62.25�N (see auxiliary material FiguresS1–S4).1 Active subduction volcanism does stop northwardwith Mount Hayes at the same latitude north. Another lesspronounced lineament of high b value exists along a SW-NE axis at depths of 60–80 km and does not coincide withvolcanoes.[26] Assuming a resolution of about 10–20 km diameter

of the cylinders on a WBZ thickness of 20 km allowscomparison of the top and bottom layer (10th and 90thpercentile) because they represent two different populations.Comparing the resulting b value maps at the top and at thebottom of the WBZ (Figure 3), we observe that prominentlineaments on the top level are not present deeper within theslab. Generally, the distribution of b values at the bottom ofthe WBZ is much more homogeneous and of lower valuewhen compared to the top level.

4. Targeted Seismic Tomography

4.1. Method

[27] Performing an inversion of seismic traveltimes toobtain the 3-D velocity field requires an appropriate initialreference model [Kissling et al., 1994]. The so-calledminimum 1-D velocity model is not only best suited as aninitial 1-D velocity model for the 3-D inversion but, incombination with station delays, accounts for near-surfaceeffects. It also allows regionally consistent highest-qualityhypocenter locations [Kissling, 1988]. For our study, aminimum 1-D model is calculated as the solution to thecoupled hypocenter-velocity model problem by joint inver-sion of traveltime data. Our local earthquake data set hasbeen selected to obtain uniform and high resolution in thetarget region, i.e., the location of b value anomalies in theslab and the region directly above (Figure 2). Thereby, Pand S arrivals are weighted in the same way since they havethe same OQCs (see section 2.3). To study the stability ofthe inversion and the dependence of the results on dataselection, we divide the data set into two sets and determinefor each a minimum 1-D velocity model with station

corrections. The differences between the two models aresmall, and the station corrections in both cases coincidewith local geology. Hence, we use the average of the twomodels as our final minimum 1-D velocity model (Table 1and Figure 5). In the next step, all 5526 events are relocatedusing the minimum 1-D velocity model and includingstation corrections. This set of hypocenters and the param-eterized minimum 1-D velocity model act as initial hypo-centers and as reference models for the 3-D inversion.[28] We use SIMULPS [Eberhart-Phillips, 1990;

Thurber, 1983] to invert simultaneously for the 3-D velocityfield and hypocenter locations. The inversion code has beenextended by Haslinger and Kissling [2001] for 3-D rayshooting. The inversion code solves the nonlinear, coupledhypocenter-velocity model problem by a linearized, iterativeapproach. After each iteration step, hypocenters are recom-puted using the updated 3-D velocity field. Thus, weapplied a staggered inversion scheme, starting with theaforementioned minimum 1-D velocity model, progressingto the simultaneous inversion for the 3-D velocity field andhypocenter locations.[29] The velocity fields for P and S waves are defined by

a horizontally equidistant inversion grid of 20 � 20 km(Figure 1). In depth, the 15 vertical layers located between�5 and 200 km depths are spaced at 10, 15, 20, and 40 km

Table 1. VP and VS of the Minimum 1-D Velocity Model

Depth (km) VP (km/s) VS (km/s)

�5 5.40 3.650 5.70 3.677 6.70 3.8317 7.08 3.9627 7.43 4.1737 7.70 4.3347 7.92 4.4560 8.08 4.6175 8.25 4.8290 8.37 4.89105 8.46 4.95120 8.50 4.97140 8.53 4.98160 8.54 4.99200 8.55 5.00

1Auxiliary materials are available in the HTML. doi:10.1029/2008JB005958.

Figure 5. VP and VS of the minimum 1-D velocity modelthat was used for relocating the earthquakes and acted asinitial velocity model for seismic tomography.

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Figure

6.

(a)Differenttarget

layersbetween37and105km

depth

andmodel

Athat

was

usedas

input.Theredbody

showstheE-W

dippingslab

inthenorthwithnegativeanomaly(10%)withrespectto

theminim

um

1-D

velocity

model.In

thesouth,ahorizontalbar

isplacedabovetheslab

inthemantlewedgeto

show

theresolutionthere.

Model

Ahas

anomaliesin

VP,whereasVP/V

Swas

keptconstant.(b)ResultingVPfieldformodel

A.Theresidual

diagonal

elem

ents

(RDE)contourlineforPisindicated

byasolidgrayline.(c)ResultingVP/V

SfieldformodelA.TheRDEcontourlineforS

isindicated

byasolidgrayline.At90and105km

depth,thezoneofhighbvalues

isindicated

withasolidblack

linein

the

SW.Black

dotsindicatethegridnodes

oftheinversiongrid.

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(Table 1). The cell dimensions in the target region at 100 kmdepth are 20 � 20 � 15(z) km.

4.2. Resolution

[30] We first evaluate the resolution of cells by commonlyemployed resolution estimators: the number of hits, HIT;derivative weighted sums, DWS; residual diagonal ele-ments, RDE [Husen et al., 2000]; and spreading value,SPRD [Toomey and Foulger, 1989; Michelini and McEvilly,1991]. In a second step, quantitative resolution estimates forthe selected region are obtained by synthetic data testing[Haslinger et al., 1999]. Synthetic traveltimes are computedby using a 3-D synthetic model (Figures 6a and S5a) and thestation and hypocenter locations of the real data set. Forcalculation of synthetic traveltimes we use an approachbased on finite difference approximation according toPodvin and Lecomte [1991] and Husen and Kissling[2001]. Then, Gaussian noise was added to the traveltimesaccording to the OQC to mimic normal observation errors.Synthetic data tests are carried out by performing simulta-neous inversions of the synthetic data using the same 3-Dmodel grid as for real data. Appropriate damping is chosenon the basis of synthetic data testing with a characteristicmodel [Haslinger et al., 1999; Husen et al., 2000].[31] To quantify resolution capabilities of our data set

within the target region, we designed two specific syntheticmodels A and B. Model A (Figure 6a) has anomalies in VP,whereas VP/VS was kept constant. Model B (Figure S5a) hasanomalies in VS, whereas VP/VS was varying to keep VP

constant. The resulting velocity fields after inversion for VP

and VP/VS are shown in Figures 6b, 6c, S5b and S5c. Formodel A, within the 0.2 RDE contour line (see below), theamplitude recovery is better for VP, but the shape of theanomalies is well resolved for both VP and VS. We also notesome leakage and overswinging artifacts mainly at shal-lower depths. Recovering of the geometry anomalies withlarge amplitudes in VS only (model B) is good, despite somevertical leakage (e.g., at 60 km depth from below and at 105km from above). We also note only negligible leakage fromVS anomalies into the VP field.[32] On the basis of these tests, we define our well-

resolved region to be within the RDE for VP/VS of 0.2(Figures 6c and S5c) and the fairly well resolved regionwithin the RDE for P of 0.2 (Figures 6b and S5b). Forinterpretation, we only use well- and fairly well resolvedregions. Results in poorly and nonresolved regions outsideof RDE for P of 0.2 should be neglected. In general, wehave a good horizontal resolution, whereas the verticalresolution suffers from vertical leakage due to locallyoverrepresented subvertical rays. We do not have goodresolution in the back-arc region between 59.75�N and62.0�N and with increasing depth (Figures 6 and S5). Thetarget volume of high b value between 90 and 100 kmdepths (Figure 3) is at the border of, but still within, thewell-resolved region. Toward the north, the resolutiondecreases, e.g., below Mount Spurr and Mount Hayes,where it is fairly well resolved at 75 km and unresolvedat 90 and 105 km depth. In the target region we have a cellsize of 20 � 20 � 15 km. To prevent misinterpretation, weavoid analysis of single-cell anomalies. Thus, our resolutionestimates confirm that we can resolve structures of 30–40 km length with an amplitude of ±2% accuracy.

[33] Figures 4c–4f show the P velocities and VP/VS ratiosderived from the 3-D tomographic inversion projected ontothe cross sections AA0 and BB0. White areas represent thepoorly resolved and unresolved regions. For better compar-ison, the tomographic results are averaged over a similarvolume as the cylindrical sampling for b values. Hence, thevelocity values shown in the cross sections are distanceweighted averages over values at the cross section itself(100% weight), and at 25 km (50% weight), and 50 km(25% weight) distances on either side of the profile. Whilethe top part of the slab at 100 km depth is well resolved, weare not able to resolve the lower boundary of the slab atthese depths and infer this boundary from results by Kis-sling and Lahr [1991].

4.3. Results

[34] In the deeper part of the slab, the oceanic lithosphereis documented by VP over 8.0 km/s (Figures 4c and 4d). VP

values between 6.5 and 8.0 km/s characterize the seismi-cally most active zone of the slab, the WBZ. These valuesare typical for the oceanic crust and uppermost mantlelithosphere. The continental crust of the overriding plateis characterized by VP velocities below 7 km/s and reachesdepths of 40 km, consistent with previous studies that mapthe Moho at around 40 km [Kissling and Lahr, 1991; Zhaoet al., 1995; Eberhart-Phillips et al., 2006]. The mantlewedge above the zone where the WBZ reaches 80 to 100km depth is characterized by VP of 8 km/s or higher(Figure 4c). However, a small patch of lower VP can beobserved directly above the location of the b value anomaly.These general structures are not as easily identified in theVP/VS ratios (Figures 4e and 4f) with the exception of thelatter local anomaly in the mantle wedge. In profile AA0wesee a clear increase in the VP/VS field at and above the zoneof b values at 100 km depths (black solid triangle). Aplume-like anomaly of high VP/VS ratio rises from the highb value anomaly at the WBZ to the continental crustbeneath Mount Redoubt. In profile BB0, we see a ratherundisturbed mantle wedge below Mount Spurr and MountHayes with no indication of high VP/VS columns ascendingfrom the top of the WBZ; however, our resolution in thisregion is more limited. We note that the high VP/VS ratio‘‘plume’’ seen in profile AA0 is caused by reduction in VS,since VP does not change significantly in that region.[35] Figure 7 shows tomographic results plotted on the

surface of the WBZ as defined by seismicity distributiondocumenting a lineament of low VP/VS ratio along the 80 kmdepth contour line and a high VP/VS ratio lineament at the100 km contour line. Tomographic results are only shownfor regions where the VP/VS ratio is well resolved. Follow-ing the VP/VS ratio in the slab along dip from the trenchdownward, we observe (1) a high VP/VS ratio above 60 kmdepth that can be associated with the contact zone of theoceanic crust with the overriding plate, (2) a relativedecrease in VP/VS at about 80 km depth, and (3) againincreased VP/VS values at 100 km depth contour line. Withrespect to the 100 km depth contour line, Figure 8 gives abetter overview over the tomographic results regarding theVP/VS ratio. It shows the VP/VS ratio on different depthslayers. Depth layers 75, 90, and 105 km suggest SW-NElineaments of VP/VS ratios. At the 105 km depth a relativelyhigh VP/VS ratio of about 1.8 is located just below Mount

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Iliamna and Mount Redoubt and consists of approximately10 bounded grid nodes. In general, VP is gradually increas-ing along dip with depths. At depths where we observe anincrease in the VP/VS ratio (approximately at the 100 kmdepth contour line), VP is still gradually increasing withdepth. This means that changes in VP/VS at these depths aredue to changes in VS. The bottom layer of the WBZ cannotbe resolved with seismic tomography. Therefore, interpre-tation of the relative change in the velocity field between thetop and bottom layer of the WBZ is not possible, and thisdiscussion must rely explicitly on b value results (Figure 3).

5. Discussion and Interpretation of CombinedResults

[36] By introducing a new mapping approach for 3-Db value distribution in the WBZ and using relocatedhypocenters, we found clear signals for depth-related mech-anisms along the 100 km depth contour line (Figures 3a and3b). This enables us not only to reproduce results ofprevious studies [Wiemer and Benoit, 1996; Wyss et al.,2001; Schorlemmer et al., 2003] but also to provide newinsights into the extension, location, and size of theseanomalies. Figures 3a and 3b indicate that the anomaliesare rather spotwise distributed along the 100 km depthcontour line (it is best seen if spherical volume samplingis applied). Unfortunately, we can only resolve the b valuestructure over 150 km along the 100 km depth contour linein southern Alaska (Figure 3). The sudden changes ofelevated b values along dip indicate a noncontinuousmechanism with depth and coincide with the location of

the volcanoes. We discuss possible mechanisms for thisoccurrence further below. The thickness of the WBZ isabout 20 km, and below this depth, brittle deformation isabsent. Because we observe significantly different b valuepatterns between the top and the bottom layer in the WBZ(Figures 3a and 3c), we suggest that mainly the uppermostlayer of the WBZ, i.e., mainly the oceanic crust, is fissuredand filled with metasediments and hydrous minerals [Hackeret al., 2003].[37] We have high confidence in our tomographic results

since we observe a subduction zone structure although wedid not include a priori information. As opposed to theapproach by Zhao et al. [1992], who used predefined plateboundary discontinuity, the shape of the slab in our study isentirely based on data since we employed a 1-D initialvelocity model. Thus, we have verified the specific low-and high-VP zone that indicates the WBZ (Figure 4) as byKissling and Lahr [1991] without including a prioriinformation.[38] With almost uniform dip direction and subduction

velocity (see section 1 and Figure 1) over this 200 km broadpart of slab, temperature and pressure conditions alongstrike can be assumed to be likewise uniform. The obser-vation of lineament-like along-strike anomalies in the VP/VS

ratio (Figure 7) at the upper surface (top layer) of the WBZsupports the findings using b values. This gives confidencefor the significance of the detected b value anomaly withrespect to location. If we assume a slab that is only affectedby increasing temperature and pressure, we expect velocities(VP as well as VS) within each layer of the oceanic crust tochange continuously with depth (thus the slab is penetrating

Figure 7. Map view of (a) VP and (b) VP/VS ratio at the slab surface. The slab surface is defined by thedepth distribution of the earthquakes within vertical cylinders on nodes which are equally spaced on a10 � 10 km grid and are identical to Figure 3. The coastline of central Alaska is shown in gray, the depthcontour lines are dashed and indicated with the associated depth, and the volcanoes are indicated with redtriangles.

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deeper into the mantle). The VP field (Figure 7) at the uppersurface of the WBZ shows this continuous increase withdepth, whereas VP/VS does not show this clear linear trend.Thus, we can conclude that the changes in VP/VS are causedby changes in VS and that these changes in VS are discrete.[39] We observed a lateral lineament of significantly

increased b values at 90–100 km depth at the top surface,coinciding with the overlying volcanoes (Figure 3). Theseb value anomalies correlate with the boundary between thelow VP/VS ratio lineament along strike at 100 km depth andthe high VP/VS ratios farther down the slab (Figures 7 and 8).We propose that these VS and b value anomalies have thesame origin. The high b value anomalies at 60–80 km inthe SW of our study regions (Figure 3) do not coincide withthe location of the volcanoes and do not correlate clearlywith tomography results. Next, we discuss possible candi-date mechanisms such as mechanical stress, inheritedmechanical structural damage, material heterogeneity, or

dehydration of minerals from the oceanic crust and relatedprocesses in the mantle wedge.

5.1. Mechanical Stress

[40] Mechanical stress due to bending is unlikely to beconcentrated in such a small area. In addition, if this wouldbe the case, we would see such anomalies at the lowersurface as well (Figure 3c). Furthermore, we would expect asignificant change in dip angle in comparison with the areaabove and below. Finally, why should mechanical stressonly affect VS but not VP?

5.2. Inherited Mechanical Structural Damage

[41] Near-surface layers have experienced strongermechanical damage than bottom layers due to shock coolingat the mid-ocean ridge (MOR) and locally varying mechan-ical loading. The fissures in the oceanic mafic crust anduppermost lithospheric mantle are filled with hydrothermal

Figure 8. VP/VS ratio on four depth layers in the Alaska subduction zone between 60 and 105 km depth.The depth and the underlying VP/VS ratio of the starting model (minimum 1-D velocity model) are givenfor each plane. The VP/VS ratio is only plotted at nodes with sufficient resolution. The nodes of theinversion grid are indicated with black dots, and volcanoes are indicated with red triangles. The locationand extent of the b value anomaly are indicated in the depth levels of 90 and 105 km by black lines. In the60 km depth layer, the station locations are indicated with black triangles.

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fluids, with hydrous minerals, or, at the top, with sediments.Subduction processes most easily reactivate existing fis-sures, though because of bending, new cracks and fissuresare created as well [Ranero et al., 2003]. With increasingfracturing, seismic velocities (VP and VS) decrease [Boaduand Long, 1996] and b values become higher [Wang et al.,2000]. Circulating fluids in such a fracture system mightaffect shear wave velocity more strongly than P wavevelocity.

5.3. Material Heterogeneity

[42] Geologically, the oceanic crust is primarily layeredhorizontally and is laterally rather homogeneous on the10 km scale. Sediments in existing fissures and on top ofthe basalt, as well as variation in thickness and internalstructure due to the spreading history, may be the main localheterogeneities in the upper surface of the oceanic crust.The thickness of the sediment layer near the trench is a fewhundreds of meters [Ranero et al., 2003]. It is unlikely thatthere has been a time when the production of oceanic crustat the MOR has changed in such a way that a lineament withdifferent material would be parallel to the MOR. The factthat the Alaska subduction zone in central Alaska is obliquemakes such a MOR-parallel lineament as the main reasonfor the observed anomalies even more unlikely. Anotherstructure that could cause material heterogeneity with themafic oceanic crust would be a former seamount that hasbeen subducted. As there are no obvious seamount struc-tures visible in the tomographic images [Husen et al., 2003]and since similar b value anomalies have been observed inother subduction zones with similar settings [Wiemer andBenoit, 1996; Wyss et al., 2001], material heterogeneitiesseem a rather unlikely candidate to be the cause for ourobservations.

5.4. Dehydration

[43] The process of dehydration is, in the presence ofenough hydrous mineral mass, largely controlled by tem-perature, T, pressure, p, and effective stress. Our observa-tions of b values and seismic tomography clearly favor aprocess of dehydration of minerals as the main cause forvarious reasons: (1) The b lineament follows strongly thedepth contour lines, and thus, it can be assumed that thesame temperature and pressure regime produces theseanomalies (under constant plate motion). (2) The b anoma-lies can only be found in the upper surface where preexist-ing fissures filled with hydrothermal fluids, hydrousminerals and/or sediments do exist. (3) With advancingsubduction, T and p are increasing; therefore, oceanic crustalbasalt will dehydrate and be transformed to eclogite [Hacker,2008], and fluids will be released to the hanging rock. Thisrelease of fluids/water due to increased p-T conditionsresults in a decrease in shear wave velocity. Our observa-tions (Figures 4c and 4e) indicate such an increase in VP/VS

ratio that results from the VS field. Raleigh and Paterson[1965] and Green and Houston [1995] showed that anincrease in pore pressure will result from dehydration,sufficiently reducing the effective normal stress to bringthe system toward brittle failure state, thus producing earth-quakes. They further pointed out that dehydration of ser-pentine (the main constituent of near-Moho oceanic mantle)results in low-energy acoustic emission events within the

hydrous material. Translating this to the subduction zonemay indicate a high b value.[44] Obviously, there are different phases that dehydrate

progressively during the subduction of the slab [Schmidtand Poli, 1998]. Assuming a range of minerals dehydratingwith increasing temperature and pressure during the courseof subduction, we would expect a continuous high b valueand VP/VS ratio at the top of the slab, which we do notobserve. We observe locally high b values because dehy-dration causes more earthquakes with small magnitudeswhich consequently increases the ratio of small to bigmagnitude earthquakes (b value) in this zone. The highVP/VS ratio is explained by the water that is released bydehydration and therefore reduces VS. However, we observea high b value at 60–80 km depth in the SW of our studyregion that may be an indication for earlier dehydratingprocesses [Manea et al., 2004], but these signals are not asclear as the one at 90–100 km depth. Hence, in agreementwith petrological models, dehydration in general and pos-sibly dehydration of serpentine in particular seem to be themost likely candidates to fit our observations.[45] Above, we discussed different possible mechanisms

that take place in the WBZ and that coincide with ourobservations. Thereby, we have elaborated that dehydrationof minerals is a major player. Dehydration reactions gener-ate aqueous fluids and melts. These are physically charac-terized by lower velocities, especially dominant for VS. Thewater released by dehydration lowers the solidus of theperidotite located in the mantle wedge above it. Ringwood[1974] introduced the general idea that dehydration ofminerals at subducting oceanic crust could trigger partialmelting in the mantle wedge and play a crucial role ingeneration of arc magma. The combined results are shownin a 3-D representation in Figure 9. We project the b valueonto the top surface of the WBZ and also show the 1.78isosurface of the VP/VS ratio as a magenta body. The VP/VS

ratio isosurface show two prominent features. High VP/VS

ratio signature characterizes the crust, and from there, adominant vertical column of relatively high VP/VS ratios(1.78) draws down to the surface of the WBZ. There is ashift between the high b values and the high VP/VS ratio.Although the resolution does not allow interpretation downto the slab in the north below Mount Spurr, there is no onsetof a similar vertical structure of high VP/VS ratio as observedin the south. The isosurface of VP/VS ratio of 1.78 (shown inmagenta in Figure 9) and the prominent high VP/VS ratio inthe horizontal layers through the subduction zone (Figure 8)support this model with a zone of higher VP/VS ratio thatascends from the surface of the slab, at about 100 km depthup to crustal levels. This can be observed below thesouthern part of Cook Inlet, i.e., below Mount Redoubtand Mount Iliamna. Unfortunately, because of poor crossfiring, we are unable to resolve the volume between the slaband the continental crust farther to the north. Stations in theback-arc region would help enormously to improve theresolution above the slab.[46] Assuming the columns of relatively high VP/VS ratio

(1.78) to be caused by partial melting of mantle materialperidotite, this will affect the crust above. The relative highVP/VS ratio can be traced up to the crust where we find aratio of 1.85 just below Mount Iliamna and Mount Spurr.We interpret these anomalies, which have an extent of about

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Figure 9. Source and path of the magma in the Alaska subduction zone from two different view angles:(a) azimuth is �115� and (b) azimuth is �60�. The b value is calculated and plotted on top of thesubducting slab. The 1.78 isosurface of VP/VS ratio is show by a magenta body. The geography andvolcanoes are shown. The vertical lines indicate the location of the volcanoes down to a depth of 100 km.An animated version of Figure 9 is available in the auxiliary material (Animation S1).

Figure 10. Model of source and path of the magma in the Alaska subduction zone based onobservations. The rather homogenous oceanic crust is carrying hydrothermal fluids, hydrous minerals,and sediments in preexisting cracks and fissures. During the descent of the plate, the oceanic plate iscontinuously dehydrating under increasing p and T conditions. The majority of the fluids are released at100 km depth, where we observe higher b values. The released fluids will partially melt the mantle,whereupon melts and fluids rise to the continental crust, where they initiate arc volcanism.

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10 � 40 km, as zones with partial melting that feed the arcvolcanoes.[47] In contrast to our findings, the results of Wyss et al.

[2001] for the VP/VS structure in the subduction zone ofnortheastern Japan show a bent upward flowing column.Also, with 140 km depth they map the high b valueanomaly deeper than in Alaska. Several reasons couldcontribute to such differences, though in principle, theobservations remain the same.[48] Anisotropy plays an important role in the mantle

wedge [Zhao, 2001] and is strongly connected to recrystal-lization of minerals because of partial melting, stress direc-tions, temperature, and pressure. Nakajima and Hasegawa[2004] indicate that the velocity field in the mantle wedge ofthe Japan subduction zone shows anisotropic behavior withhigh velocities parallel to the trench. According to thesestudies, distinguishing between the origin of the measuredanisotropy, i.e., whether the signal is coming from the crust,the mantle wedge, or the slab, seems to be tricky. Never-theless, the contribution due to the mantle wedge seems tobe dominant [Wiemer et al., 1999]. Since the seismicvelocities derived from tomography do represent an averageof all rays (ignoring its orientation) for a grid node,anisotropy values of a few percent do not affect our resultssignificantly.

6. Conclusions

[49] We have investigated the subduction zone of Alaskaapplying two methods, FMD and seismic tomography, andusing the same database of the AEIC to ensure consistency.In summary, our results give evidence on the location,source mechanism, and path of magma that feeds centralAlaska arc volcanism (Figure 10).[50] The b value pattern between the upper and lower

surface of the WBZ is changing completely. This is clearevidence for the occurrence of different mechanisms. The bvalues on the upper surface are likely associated withdehydration processes since they correlate with anomaliesin the S wave velocity field. In the lower region of theWBZ, the resolution in seismic tomography is poor, andtherefore, our interpretation relies entirely on the b valuestudy. The lower b values at the lower surface of the WBZsuggest that the process of dehydration is less dominantthere.[51] The discussion of the different possible causes leads

us to the conclusion that the observed b value and VP/VS

lineaments parallel to strike of the slab at its upper surfacein 100 km depth most probably are associated with dehy-dration. The significance of the observed b value anomalieswas tested by following the approach of Utsu [1992], andaccording to Wiemer and Benoit [1996], Db > 0.3 aresignificant on compatible data.[52] Above the slab we rely entirely on tomography

results since b values cannot be estimated in almost aseis-mogenic zones. We observe a high VP/VS ratio that dragsdown from the back-arc crust toward the regions of anom-alously high b values. This column is just at the edge of theresolution. Because of the size and consistency of theanomaly and on the basis of resolution ability in this volumetested by characteristic tests, we can exclude artifacts.

[53] Major constraints in this study were provided byseismic tomography of simultaneous inversion of P and Swave data. Though the AEIC did provide an unusuallyhigh-quality data set, the volume of sufficiently high reso-lution is limited to the fore-arc region and to the southernparts of our study region. Our results clearly document thatonly a few more stations in the back-arc region wouldprovide much improved ray distribution of the entiresubduction zone of central Alaska. Another limitation inthe resolution is due to the unusually large error of thesecond best class in relation to the best class. It seems likelythat the second quality class for S observation could besubdivided thus reducing significantly uncertainties in seis-mic tomography.

[54] Acknowledgments. This project benefited from the thoughtfuland constructive reviews by the two anonymous reviewers and S. Husen.Many figures in this work are made using GMT [Wessel and Smith, 1991].This study was part of the Project 20021108096 funded by the SwissNational Science Foundation (SNF).

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�����������������������E. Kissling, Institute of Geophysics, ETH Zurich, Sonneggstrasse 5,

NO H 53.2, CH-8092 Zurich, Switzerland. (kiss@tomo.ifg.ethz.ch)N. Ruppert, Geophysical Institute, University of Alaska Fairbanks,

P.O. Box 757320, Fairbanks, AK 99775-7320, USA. (natasha@gi.alaska.edu)T. van Stiphout and S. Wiemer, Swiss Seismological Service, ETH

Zurich, Sonneggstrasse 5, NO H 61, CH-8092 Zurich, Switzerland.(vanstiphout@sed.ethz.ch; wiemer@sed.ethz.ch)

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