Slab low-velocity layer in the eastern Aleutian subduction zone
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Geophys. J. Int. (1997) 130,640-648
Slab low-velocity layer in the eastern Aleutian subduction zone
George Helffrich' and Geoffrey A. Abers2
' Geology Department. University of Bristol, Wills Memoriul Building, Queens Road, Bristol BS8 1RJ, U K . E-mail: george@geology.bristol.ac.uk
'Department of Geology, 120 Lindley Hall, University of Kansas, Lawrence, KS 66045, USA. E-mail: g-abers@ukans.edu
Accepted 1997 April 22. Received 1997 April 21; in original form 1996 August 28
SUMMARY
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Local earthquakes in the vicinity of the Alaskan Peninsula's Shumagin Islands often
produce arrivals between the main P and S arrivals not predicted by standard traveltime
tables. Based on traveltime and polarization, these anomalous arrivals appear to be
from P-to-S conversions at the surface of the subducted Pacific Plate beneath the
recording stations. The P-to-S conversion occurs at the top of a low-velocity layer
which extends to at least 150 km depth and is 8 f2 per cent slower than the overlying
mantle. The slab is -7 per cent faster than the mantle. The low-velocity layer contains
the foci of the earthquakes in the upper plane of the double seismic zone and confines
PS ray paths to lie within it. These observations indicate that layered structures persist
to positions well past the surface location of the volcanic front. Reactions forming
high-pressure minerals do not yield slab-like velocities until beyond the point that
subduction zone magma genesis occurs. If the subducted oceanic crust forms the layer,
it is subducted essentially intact.
Key words: body waves, subduction, phase transitions.
Fischer, Creager & Jordan 1991). Closer in, this generalization
INTRODUCTION
changes. Seismic networks deployed above subduction zones
Viewed from the perspective of the distant seismic observatory, record a variety of anomalous phases between the P and the
slabs of subducted lithosphere are high-velocity regions (Fig. 1). S arrivals that involve interactions with the slab, mostly
Seismic waves emanating from subduction zones generally reporting regions of low velocities relative to those expected
arrive early with a dependence on the azimuth from the source for a slab. Mitronovas & Isacks (1971) were the first to show
to the receiver, consistent with the effects of material that is examples of arrivals of this type, inferring from them slab/
cooler, and therefore faster, than the ambient mantle (Davies mantle velocity contrasts of 6-7 per cent. Suyehiro & Sacks
& McKenzie 1969; Sleep 1973; Creager & Jordan 1984, 1986; (1979) and Huppert & Frohlich ( 1981) analysed traveltimes to
Japanese and Fijian networks, respectively, for information on
the velocity structure within the slab, finding a slab interior
faster than its surface. Reyners & Coles (1982) documented P
and S arrivals between the direct P and S in Alaska. A low-
velocity layer on the subducted Philippine Sea plate causes
Homogeneous
unusual P arrival patterns west of Tokyo, Japan (Fukao, Hori
fast slab
(distant view) & Ukawa 1983; Hori et at. 1985). In Vanuatu, Chiu, Isacks &
Cardwell (1985) identified P multipathing in the slab as well
as P-to-S conversions at the slab/mantle interface, and found
in the observations evidence of low velocity layering in the
slab. Matsuzawa et al. (1986, 1987) also observed P-to-S
conversions arriving between P and S in Japan, similarly
finding a low-velocity layer at the upper slab surface. Body-
wave dispersion analysis, showing the earliest arriving energy
Figure 1. Two views of the subducted slab, from the near and far to be low frequency, suggests that a thin low-velocity layer is
seismic perspective. Slabs generate bulk velocity contrasts of roughly present in the slab inland from Cook Inlet in Alaska (Abers
5 per cent when viewed teleseismically. Stations nearer the slab see & Sarker 1996).
more structure in it: faster or slower layering at its upper surface. This New Zealand is an exception. Here, the uniformly early
layering is the subject of the present work. arrival times from Tonga, Kermadec and North Island
640 0 1997 RASAlaskan slab low-velocity layer 641
earthquakes require high shallow-slab wavespeeds (Galea 1992; Table 1. Shumagin velocity
van der Hilst & Snieder 1996). The earliest arriving energy is model.
high frequency, an observation that corroborates traveltime- Depth V, v,
based inferences and suggests that a high-velocity layer exists km km s-' km s-'
in ;he slab, stretching from southern Tonga to New Zealand
-1 6.2 1 3.31
(Ansell & Gubbins 1986; Gubbins & Snieder 1991; van der
10. 6.56 3.75
Hilst & Snieder 1996).
20. 6.92 3.89
That layered structures in a subducted slab may exist is not
sunrising given that what is subducted, oceanic lithosphere, 30. 7.32 4.11
is already compositionally, mineralogically and elastically 40. 7.66 4.38
layered (Fox & Stroup 1981). Persisting as it does to depths 75. 8.03 4.57
in excess of 150 km, the layering argues strongly against crustal 145. 8.38 4.79
delamination or underplating during subduction (Helffrich
1996), and it is worthwhile to establish whether slab layers
are common or rare. What the layering implies about the
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mineralogic evolution of the slab is also of interest due to the
Shurnagin network
57
56
55
54
53
-164 -163 -162 -161 -160 -159 -158 -157 -156 - 5
Swath width: 60km
( 57.06, -162.34) az 150 SAS 150 cross section ( 53.60, -158.81)
0
50
y 100
x
v
f
," 150
0
200
250 I I
1
I
V.F.
I I I
-200 -1 00 0 100 200
Distance (km)
Figure 2. (a) Map of the Shumagin Islands area of the Alaskan Peninsula. Crosses denote locations of three-component seismic stations. Volcano
positions are indicated by triangles. The trench position is evident by contours of bathymetry at 200, 3000 and 6000 m depth. Points are relocated
seismicity in the vicinity (Abers 1994); circles indicate events used in this study. The line is the cross-section through station SAS approximately
perpendicular to trench. (b) Cross-section of seismicity along the line in (a). The line marked V.F. indicates position of the volcanic front. The
seismicity clearly defines upper and lower seismic planes. Only the hypocentres 30 km to either side of the projection line are shown, but up to
600 km along-strike projection swaths define a double seismic zone.
0 1997 RAS, G J I 130, 640-648642 G. Helffrich and G. A. Abers
Evtnk1987.299 Sta:sos DiskO.6 k 4 0 . 7 Eaz:221.2 54.916N -161.140E 80.4km
2'- tangent
0.
n -2-
0
x c
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Scclrdr
usw Yla
Figure 3. Local earthquake recorded at station SAS. The three traces record the radial, tangential and vertical components of motion. Vertical
lines indicate IASP91 (Kennett & Engdahl 1991) arrival times for P and S. Secondary P arrivals occur at -22 s on the vertical component, as
well as at -28 s (just before the S arrival). The prominent S arrival at -25 s is X .
Table 2. Polarization of X arrivals recorded at SGB.
Back- Polarization Back- Polarization Back- Polarization Back- Polarization
azimuth, O angle, O azimuth, O angle, O azimuth, ' angle, O azimuth, O angle, O
-144 52.2 32. 247.2 -136 34.0 -106 274.9
46. 236.6 -98 173.7 -53 330.9 93. 268.1
-24 152.8 - 104 307.6 -144 320.2 3. 252.2
-141 33 1.2 -5 345.2 69. 186.8 - 142 287.5
158. 2.5 -153 35.6 -164 219.9 -137 179.1
179. 2.7 59. 143.6 -106 292.8 49. 318.6
-148 235.5 -146 56.2 -162 197.0 -161 24.3
89. 279.0 51. 2 16.0 175. 331.1 -90 314.1
-148 208.8 -153 313.1 -156 4.7 - 157 313.4
-52 344.8 -107 267.5 -86 226.2 -163 352.4
90. 286.7 -139 49.7 -143 155.2 -82 101.1
101. 334.3 -150 289.1 -145 31.6 -154 149.8
-141 44.4 -128 329.7 -141 275.2 -132 255.1
- 108 262.0 -139 131.3 -123 270.9 -127 195.7
-136 257.4 -1 18 287.8 -148 3 19.6 53. 178.1
-139 274.2 - 140 149.4 - 141 273.0 - 127 249.0
-100 317.4 -131 271.5 -135 269.0 -131 268.8
-99 294.8 -135 2 10.I -139 264.0 -106 293.6
-134 169.9 -123 258.9 -106 72.5 -134 277.0
-120 28 1.O -121 274.1 -129 291.7 70. 234.0
-101 293.5 -1 14 184.9 - I 14 288.9 55. 261.5
-127 273.2 -113 249.3 -1 14 265.4 -126 354.0
- 126 272.8 -1 16 116.7 -125 275.4 - 125 273.5
-126 141.9 -95 307.7 - 103 269.9 -122 292.5
-99 240.5 -111 51.7 -96 262.8 -1 16 282.8
-121 124.0 -105 265.8 51. 182.5
0 1997 RAS, GJI 130, 640-648Alaskan slab low-velocity layer 643
SGB X Dolarization analvsis postulated link between intermediate-depth earthquakes and
90 I phase transformations in the slab (Kirby 1995). Here we
-
v
0
80
70
60
Teench
n o r m a l0
0
examine arrivals recorded by a local seismic network in the
Shumagin Islands of the Alaskan Peninsula for evidence of
layered structures in the slab. The observations indicate that
C 50 a z i m wt h a low-velocity layer, which is about 8 per cent slower than
.-0
./-
+
40 mantle wavespeeds, persists to more than 150 km depth in a
0 30 c 0 slab that is about 7 per cent faster than the mantle. Thus the
.-
N
0
-
L 20
10 -P-*
0
- Alaskan subduction zone joins the more numerous ranks of
0 0 those having slow layers. Moreover, if the nature of the layering
Q 0 is mineralogical, a slow mineralogy must persist past the depth
of subduction-zone magma genesis.
OBSERVATIONS
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Our data are from a Lamont-Doherty operated local seismic
network in the Alaskan Peninsula (Fig. 2 ) which ran stably
between 1982 and 1991 (Abers 1992). The network's digital
-90
-80 0l l l l ~ l l t l l l data streams are sampled at 100Hz. Most of the nominally
-180-150-120-90-60-30 O 30 6 0 90 120 150 180 1 Hz free-period HS-10 instruments are vertical component,
B a c k - a z i m u t h (") but five of the stations are equipped with horizontals. Since
our interest is in identifying the types and traveltimes of
Figure 4. Plot of the deviation between the back-azimuth from station arrivals between P and S, these latter stations provide the data
SGB to the event and the horizontal polarization of the X arrival, for the study. The events studied are drawn from an earthquake
which is an S wave. The X polarization is obtained by an eigenvalue
analysis of horizontal particle motion and accounts for any splitting
catalogue of relocated seismicity within the network (Abers
of the S wave (Silver & Chan 1991). If a P-to-S conversion occurs at 1994) (Fig. 2 ) . The relocation jointly solved for velocity
a dipping interface, the polarization will rotate to the down-dip structure and earthquake locations. The starting velocity model
direction. Thus the deviation between the back-azimuth and the for the relocation is a 1-D model developed for locating
polarization direction should reveal the dip direction. All the initial earthquakes in the.network (Table 1).
polarizations are plotted without regard to quality, and a line is fit We examined waveforms from those of the 1448 relocated
using a robust fitting technique (Menke 1989). The line crosses zero events located deeper than 40km for secondary P and S
(no deviation between back-azimuth and initial polarization) at an arrivals in the window between the direct arrivals, classifying
azimuth of about -3O", which is perpendicular to the trench azimuth them as to their polarization. By omitting events shallower
and the volcanic front azimuth in this area. Thus the polarization than 40 km we emphasize paths leading from hypocentres in
trend has the expected characteristics for a P-to-S conversion at the
the subducted plate, maximizing their interaction with the
slab/mantle interface. Data are in Table 2. Scatter is due to most
events being located along strike (Fig. l), resulting in an effectively plate. 172 arrivals are S and 65 are P. Fig. 3 is an example of
zero interface dip and no polarization rotation. one of the records, which shows P and S arrivals from a local
event. Between the main phase arrivals, at about 26 s, is an
X arrivals at SAS X arrivals at SGB X arrivals at SQF
I I I I I I I I I I I I I I I I I I I I I
0 I5
'0
A
A A
0
A f
5
A
A
A
A
I I I I I I I I I I I I I I 0 I I I I I I I
5 10 1520253035 5 10 1520253035
S-P (sec) S-P (sec) S-P (sec)
Figure 5. Plot of picked X arrivals for the three stations where it is common. The estimated picking error is about the size of the symbol, 0.3 s.
Trends in the arrivals are highlighted by different symbols: solid where they correspond to a high-velocity arrival branch (smaller X - P times) and
open for low-velocity arrivals (larger X - P times). Triangles designate arrivals not clearly associated with either of these groups.
0 1997 RAS, GJI 130, 640-648644 G. Helflrich and G. A. Abers
arrival polarized dominantly as S. This phase X arrives before
MODELLING A N D RESULTS
S, so must have either travelled in material faster than did
direct S or accrued some of its traveltime as P. The latter The character of the flat trend, whose arrivals are S with nearly
possibility is unlikely because a regional tomographic study constant X - P times, suggests that the phase is generated
(Abers 1994) does not reveal extensive regions of high S near the station. As a working hypothesis we assume that X
velocities, which are required by this interpretation. Moreover, is a PS conversion at the slab surface near the station and
X is polarized nearly radially in the slab down-dip/trench trace rays through a 2-D model [using SEIS83 by Cerveny &
normal direction, implying an origin as a P-to-S conversion Psencik (1983)] with a 100 km thick slab 5 per cent faster
(Fig. 4). Three of the stations, SQF, SAS and SGB, reliably than the ambient mantle (Zhao, Christensen & Pulpan 1995)
record X. The individual picks for each station are best which dips at 30". The shallow seismicity cross-section (Fig. 2 )
visualized in terms of their differential traveltimes (Fig. 5 ) . supplies the dip, as well as the station location above the slab.
S - P roughly represents range (depth as well as epicentral From deep events the P-to-S conversion at the slab interface
distance), and X - P is the time lapse from the P arrival. What is essentially fixed, but it moves beneath the station to shallower
is seen in the X - P times in Fig. 5, most clearly for SAS, are levels as the source approaches the trench (Fig. 7a). This leads
two distinct traveltime trends. These emanate from an S - P to a nearly constant P S - P time at large range, but a trend
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time of about 10 s, one with increasing X - P and the other to smaller times at close range, with a sharply curved branch
with nearly the same X - P time. Henceforth the trend with forming that corresponds to up-dip source locations (Fig. 7b).
increasing X - P times will be called the 'increasing trend' If instead of a fast slab we hypothesize a slow one and trace
and the other the 'flat trend'. The sources comprising the two rays through it, the ray geometry of the P-to-S conversion
trends show a depth segregation in the slab (Fig. 6), but does not change much from that of a fast slab, but the
overlap in map view. traveltimes do. PS - P times lengthen with range because the
SCB
0 I I I I I
50
-
Y
E
100
v
so*
fn
0
150 0
.* 0
200
**
250 I I I I L
I00 -50 0 50 100 -50 0 50 1O(
o 1 I I I 1
20 - -
-c
E
40
60
-
- A A
-
-
A
fn A
0 00 - -
A
A
100 - -
120 - I I I I
-
Figure 6. Distribution of hypocentres associated with sources identified in Fig. 5, using the same symbol codings: solid symbols denote the arrivals
in the flat trend, and open symbols indicate events in the increasing trend. Events not clearly associated with any group are shown with triangles.
The arrivals in the increasing trend are segregated to the upper seismic plane, whereas the flat-trend arrivals are in the lower seismic plane.
0 1997 RAS, GJI 130, 640-648Alaskan slab low-velocity layer 645
0
PS ray paths and secondarily on the high-velocity slab speed. A 10 km thick
layer shifts the asymptote by about 1 s on account of the delay
50
(a) accrued by PS traversing it as P from lower-plane sources
J;
(Fig. 8b). This suggests that the layer is thin, closer to 5 km as
~
100 opposed to 10 km. There is a mild trade-off with the speed of
~ the high-velocity plate, but a 10 km layer fits the data only
J: with a 15 per cent faster plate, which is faster than the 3-5
1- 150
a.
L&J per cent increases in shallow plate speeds relative to overlying
0
200 mantle reported elsewhere (Hori et al. 1985; Matsuzawa et al.
1986). Thus we opt for a 5 km layer and a 7 ± 2 per cent faster
250 slab. This is a compromise between results from SGB and SAS.
150 250 350 450 550
DISTANCE IN KM
DISCUSSION
Calculated PS-P times
(b) These results extend an earlier study of the Alaskan subduction
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15 "' 1sh b dip 1 zone ~ 400 km along-strike to the east of the Shumagins by
oo ut-of-plane Abers & Sarker ( 1996). Body-wave dispersion characteristics
00 0 suggest the existence of a 2-6 km thick layer with 2.5-5 per
oo o rrors
~
10 oo cent lower velocities than the mantle and a 5 per cent faster
-!"
0..
I
Vl
0.. I
•~·········
• Fast sla
slab. In the Shumagins, we find a slower layer speed, a similar
slab speed, and a comparable layer thickness. The thickness is
5 roughly equivalent to other reported values; 6 km by Hori
0 Slow slab et al. (1985); 5 km by Matsuzawa et al. (1986); 6-15 km by
Gubbins & Snieder (1991); and 8 km by van der Hilst &
0 L-~~--~-L--L-~~~ Snieder ( 1996). Oceanic crust is generally 7.1 ± 0.8 km thick
0 5 10 15 20 25 30 35
S-P {sec) (White et al. 1992; Mutter & Mutter 1993), and specifically
6.5-7 km thick in the Gulf of Alaska environs (Moore et al.
Figure 7. (a) PS ray paths for sources near the top of a slab pro- 1991; Bracher et al. 1994), with ~0.8 km sediment overburden
portionately faster (or slower) than the mantle above it. Each cross
(Winterer 1989). This suggests that the low-velocity layer
designates a source. The rays travel up the slab as P and convert to S
at the slab/mantle interface. For the deepest source, a direct P path is
comprises all of the extrusive (layer 2) and most or all of the
also shown. (b) If the slab is faster than the mantle, the S leg contributes intrusive (layer 3) package if it indeed represents the oceanic
little total traveltime to PS, yielding the trend shown with solid crust. The observed layer thickness, 5 km, is nearly the thick-
symbols. A slower slab, because the long P leg is also delayed, yields ness of layer 3, based on it typically being two-thirds of the
a positively sloping traveltime curve (open symbols). Note also the crust thickness (Spudich & Orcutt 1980), so if there is tectonic
sharply curved reverse branch caused by up-dip sources. Brackets on erosion or underplating of the sediments and extrusives, what
the right of the plot indicate ±so uncertainty in the plate dip, and of would remain is a layer approximately as thick as we infer.
out-of-plane sources (along-strike arrivals in a 3-D rather than a 2-D Seismic waves emitted by earthquakes in the low-velocity
geometry). These are of little significance to the overall trend. layer must remain in it-otherwise the rigorous segregation of
upper- and lower-plate sources into low- and high-velocity
P path, confined to the slab, is slower than direct P which paths would not be observed. Since the mantle wedge is
leaves the slab close to the source on account of Snell's Law generally aseismic (Abers 1992; Zhao, Hasegawa & Horiuchi
acting at the slow slab's surface. The slow slab thus accounts 1992), and since the layer in which the upper-plane earthquakes
for the increasing trend whereas a fast slab is responsible for occur is roughly the oceanic crust thickness, we believe that
the fiat trend. The PS conversion at the slab/mantle interface the layer is the top of the slab, and not the hydrous layer
appears to explain both branches of X. above the slab as Tatsumi, Ito & Go to ( 1994; Ito & Tatsumi
While it seems paradoxical that a slab is both fast and slow, 1995) envisage.
some insight may be gained from the distribution of hypo- A further insight into slab velocity gradients also derives
centres that generate the fast and the slow arrivals. Their from the observed spatial separation of flat- and increasing-
distribution for SAS PS observations (Fig. 6) clearly restricts trend arrivals. Snell's Jaw forbids a near-source PS conversion
the slow arrivals (the increasing trend in Fig. 7) to the upper of a vertical ray, but downward propagating rays conceivably
surface of the slab and the faster arrivals (fiat trend) to its could leave the low-velocity layer, bottom in the slab, and
interior. A similar separation, while not as rigorous as at SAS, ascend through the layer to the station. Since they generally
is suggested by the hypocentres for SGB. These results suggest do not do this, this probably means that velocity gradients in
that a low-velocity layer exists at the upper-slab surface. the slab below the low-velocity layer are too small to turn
By tracing rays through a variety of models, some constraints downward propagating rays upwards at the short ranges found
on the low-velocity layer speed and thickness may be sought. in the Shumagins.
Slab wave speeds in these models are relative to the surround- Two effects may conspire to prevent observation of PS
ing mantle speeds given in Table 1. The arrivals forming the conversions at the other stations in the network (Fig. 1). At
increasing trend are sensitive to the low-velocity layer speed BLH, the deepest into the backarc, PS and S arrive nearly
and are fit best if the layer is 8 ± 2 per cent slower (Fig. Sa). simultaneously for S- P times less than ~ 21 s due to either
In contrast, the asymptotic PS- P time in the fiat trend the long upgoing updip S leg or long downdip P legs (Fig. 7a).
depends primarily on the thickness of the low-velocity layer Since few earthquakes are more distant than this, observable
© 1997 RAS, GJI 130, 640-648646 G. Helfrich and G. A . Abers
SGB SAS
I I I I I
15 0
a dip
00
iIIt 0
O
& geometry
errors
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I l l I I t
0 5 10 15 20 25 30 35 0 5 10 15 20 25 30 35
S-P (sec) S-P (sec)
0
(b)
50
5 100
z
I
i- 150
a
w
D
200
250
150 250 350 450 550
DISTANCE IN KM
Figure 8. (a) Calculated and observed PS arrival times, assuming a slab/mantle interface PS conversion. The preferred slab model has a 5 km
thick layer 8 per cent slower than the overlying mantle on top of an interior 7 per cent faster than the overlying mantle (a 21 per cent velocity
increase at the bottom of the low-velocity layer). Brackets on calculated times are + 2 per cent variation in layer velocity (low-velocity trend) and
joint -+2 per cent variation of both slab and layer velocity (high-velocity group). Lower brackets indicate uncertainty due to plate dip and 3-D
geometry (Fig. 7). (b) PS and P ray paths from a lower-plane source, illustrating the effect of the low-velocity layer thickness on PS - P times. Its
thickness only affects flat-trend times (Fig. 7b), displacing them to larger PS - P values. For a given thickness, the layer adds a constant delay to
PS because the conversion point is essentially fixed (Fig. 7a). The P delay depends on source depth due to the changing P geometry. An increase
in PS - P time is the net result.
PS conversion conditions are rare-85 per cent of the recorded invoked to explain basalt persistence past the equilibrium
distant than -
events are excluded. Similarly, PVV requires events more
17 s, excluding 70 per cent. Thus the eligible
population is small. The other effect that may impair obser-
reaction depth, -20 km (Wood 1987). Low-velocity material
extends beyond the depth of magma generation beneath the
volcanic arc at -100km depth, however. At this depth,
vation is the slab/mantle interface conditions at the points dehydration reactions are believed to transfer volatiles to the
where the P-to-S conversion occurs, which are nearly constant mantle wedge, which promotes melting of the mantle (Tatsumi
for distant events. Any heterogeneity in interface properties, if 1986; Davies & Stevenson 1992). Volatiles also lower kinetic
arising by chance at this point, would inhibit observation. barriers to reaction (Fyfe & Verhoogen 1958). Thus meta-
Abers’s ( 1994) tomographic study indicates some variability stability in basalt/gabbro would be difficult to maintain to the
above the seismicity, suggesting interface inhomogeneity. Note, depth where low velocities are observed.
however, that interface heterogeneity does not affect the travel- Alternatively, the low-velocity layer may represent hydrous
time trends from which the layer properties are derived. The metabasalt facies, which evolve by dehydration to high-
long up- or downdip paths in the low-velocity layer or high- pressure assemblages (Peacock 1993). At 700°C and 20kb
velocity slab govern the trends, not the interface properties at (65 km depth), Pawley & Holloway (1993) showed hydrous
the conversion point (Fig. 7). phases to be stable in metabasalts and Poli & Schmidt (1995)
Hypocentres associated with the low-velocity traveltime reported lawsonite stable to >60 kb. Down to depths of
branch extend down to 180 km (Fig. 6), suggesting an extent 100 km, calculated metabasalt velocities (lawsonite blueschist)
somewhat deeper than the 100-150 km found by Abers & remain 7 per cent slower than the mantle (Helffrich 1996),
Sarker (1996). Since both the thickness and the common low- with supporting exrsrimental evidence for an alternative
velocity character suggest that it represents the subducted hydrous mineralogy (Tatsumi et al. 1994; Ito & Tatsumi 1995).
oceanic crust, a conversion from basalt (or the intrusive While the mineralogy of this low-velocity layer is as yet
equivalent gabbro) to eclogite is expected (Ahrens & Schubert unknown, its depth extent may suggest pressure regimes for
1975; Gubbins et al. 1994). Typically, kinetic hinderance is future exploratory work in experimental petrology, as well as
0 1997 RAS, G J I 130, 640-648Alaskan slab low-velocity layer 641
a role in the overall to the slab/mantle velocity contrast a t Fukao, Y., Hori, S. & Ukawa, M., 1983. A seismological constraint
intermediate subduction-zone depths. More work is warranted on the depth of basalt-eclogite transition in a suhducting oceanic
to find how deep low-velocity layering persists. crust, Nature, 303, 413-415.
Fyfe, W.S. & Verhoogen, J., 1958. Kinetics of metamorphic reactions,
in Metamorphic Reactions and Metamorphic Facies, pp. 53-104, eds
Fyfe, W.S., Turner, F.J. & Verhoogen, J., Geol. SOC.Am., New
ACKNOWLEDGMENTS
York, NY.
This unusually high-quality dataset w a s collected over a 10 Galea, P., 1992. Observations of very high P-velocities in the subducted
year period by the seismology g r o u p a t Lamont-Doherty: slab, New Zealand, and their relation with the slab geometry,
K. Jacob, J. Davies, E. Hauksson, J. Taber, D. J o h n s o n and Geophys. J. Int., 110, 238-250.
others. We salute their successful, long-term network main- Guhhins, D. & Snieder, R., 1991. Dispersion of P waves in sub-
tenance efforts. GA acknowledges support from USGS a n d ducted lithosphere: evidence for an eclogite layer, J. geophys. Res.,
96, 6321-6333.
DOE grants 1408-0001-A0616 and DE-FG02-84ER13221F
Guhbins, D., Barnicoat, A. & Cann, J., 1994. Seismological constraints
and their predecessors at L a m o n t for operation and analysis
on the gahbro-eclogite transition in subducted oceanic crust, Earth
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