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been subject neither to the effects of oceanic break-up nor to intensive rifting episodes
with related crustal thinning as in the North Sea, the crust is 30–35 km thick. In island arcs
(e.g. Japan) the crustal thickness is about 20–30 km and has a pronounced 3D structure.
A low velocity mantle with Pn velocity of 7.5–7.8 km/s suggests high mantle temperatures.
Oceanic crust
General patterns
Compared to the continental crust, the oldest oceanic crust is Jurassic (~180Ma). Classically
the oceanic crust is divided into three layers (Fig. 3.32). Layer 1 is the sedimentary layer
with a highly variable thickness. Layer 2 has P-wave velocities of 4.5–5.6 km/s, is
c. 1.5–2.0 km thick, and is made of extrusive volcanic rocks (pillow basalts at shallow
depths which grade downwards into sheeted dikes). Layer 3 (typical thickness 4.5–5.0 km)
with P-wave velocities of 6.5–7.0 km/s has gabbro composition similar to the composition
of ophiolite complexes. The top of Layer 3 is transitional with interfingering of sheeted dikes
into the lower part of Layer 2. Further down, the upper part of Layer 3 consists of isotropic
gabbro, underlain by layered gabbro and harzburgite. In mid-ocean ridges, where the new
oceanic crust is formed, the base of the crust also marks the base of the lithosphere. The
thickness of the oceanic crust (commonly defined as the thickness of Layers 2 and 3) is fairly
uniform globally (5.6–7.1 km according to different authors) and does not show any age
dependence (Tanimoto, 1995).
Anomalous oceanic crust
Some oceanic regions have crustal thickness significantly thicker than global observations.
These regions are of a particular interest because, as a rule, bathymetry there does not follow
the square root dependence on ocean floor age (see Chapter 4). This fact, together with the
presence of anomalously thick oceanic crust, is attributed to mantle thermal anomalies with
a high degree of melt generation. In ocean plateaux, crustal thickness reaches 20 km (and
may be as thick as 35 km beneath the Ontong–Java plateau); the origin of this thick crust is
commonly ascribed to a large amount of melt generated by a mantle plume. However, the
seismic structure of many other ocean plateaux is still poorly known and some of them (such
as the Falkland Plateau, Lord Howe Rise, the Kerguelen Ridge, the Seychelles Ridge, and
95 3.3 Major seismic discontinuities in the lithosphere
the Arctic Ridge) may contain fragments of the continental crust (Mooney et al., 1998 and
references therein).
A thick (15–25 km) crust beneath several aseismic ocean ridges such as theWalvis and the
Ninety-East Ridge in the Indian Ocean (Detrick and Watts, 1979) has been also explained
by the presence of large volumes of melt associated with mantle upwellings (e.g. White
et al., 1992). Similarly, anomalously thick oceanic crust (30–35 km) underlies the Faeroe–
Iceland–Greenland ridge that extends from Greenland to the British Isles across the
Mid-Atlantic Ridge (Bott and Gunnarsson, 1980; White et al., 2008). The presence of a
thick oceanic crust can be interpreted in favor of the presence of a mantle plume with
anomalously high melting temperature at Iceland, while the east–west symmetric crustal
structure of the Faeroe–Iceland–Greenland ridge around Iceland is used as an argument
against the plume origin of Iceland unless the plume location has been semi-stationary with
respect to the plate boundary (Lundin and Dore, 2005).
Crustal thickness in Iceland is a topic of hot debate: two competing models, “thin crust”
and “thick crust”, have been proposed for Iceland (see review by Foulger et al., 2005).
Numerous seismic studies indicate that the oceanic Layer 3 (down to a depth of 10–20 km)
has Vp velocity of 6.5–7.0 km/s. Below this depth, in Layer 4 (which may extend down to
60 km depth), the P-wave velocity gradually increases to 7.0–7.6 km/s. The principal differ-
ence between the two crustal models is the petrologic interpretation of Layer 4: in the “thin
crust” model it is interpreted as anomalous peridotite mantle with c. 2% of melt, while in
the “thick crust” model the same layer is interpreted as gabbroic “lower crust” with lenses
of melt.
It should be noted that some oceanic regions have anomalously thin crust. They include
regions with slow spreading rates (< 2 cm/y), non-volcanic rifted margins that have earlier
undergone extreme extension, and fracture zones where Layer 3 is very thin or absent.
3.3.2 Seismic discontinuities in the upper mantle
LVZ, G-discontinuity and the base of seismic lithosphere
Seismic lithosphere (or the lid) is defined as the seismic high-velocity region on the top of
the mantle. It generally overlies a low-velocity zone (LVZ) that was first recognized by Beno
Gutenberg (1959). Thus, the top of the LVZ defines the bottom of the lid (Anderson, 1995).
Since the LVZ is located between the thermal boundary layer (above) and the nearly
adiabatic mantle (below), in the oceans the top of the LVZ and the top of the asthenosphere
(a layer with lowered viscosity) are equivalent and are marked by the seismically sharp,
Gutenberg (G-), discontinuity that is characterized by an abrupt seismic velocity decrease of
roughly 9% (Bagley and Revenaugh, 2008).
Starting from the early works, a decrease in seismic velocities in the LVZ was attributed
solely to partial melting (e.g. Lambert and Wyllie, 1970). The mechanism behind this
melting is still the subject of discussion; volatile enrichment of the mantle is a possible
cause. Alternatively, a sharp decrease of water solubility in mantle minerals at the depths
that correspond to the LVZ may cause excess water to form a hydrous silicate melt
(Fig. 3.42) (Mierdel et al., 2007). If the LVZ below the seismic lithosphere results from
96 Seismic structure of the lithosphere
partial melt, the lithospheric base should be a rheological boundary and it may be marked
by a change in mantle anisotropy pattern from frozen-in anisotropy in the lithosphere to
anisotropy in the asthenosphere due to mantle flow.
Recent studies favor explanations other than partial melting, such as high-temperature
relaxation (Anderson, 1980), a contrast in volatile content from water-depleted lithosphere
to water-rich asthenosphere (Karato and Jung, 1998), and grain size variations without
requiring the presence of melts (Faul and Jackson, 2007). These mechanisms suggest that
the base of a seismic lithosphere should be a diffuse boundary which extends over a certain
depth interval. Seismic studies in the western and central Pacific based on multiple ScS
reverberations put an upper limit of 30 km on the transition interval where most of the
velocity decrease from the lid to the LVZ takes place (Bagley and Revenaugh, 2008), while a
recent study based on converted P-to-S and S-to-P waves suggests that the velocity drop at
the discontinuity occurs over depths of 11 km or less (Rychert and Shearer, 2007). The width
of this transitional interval is more easily explained by the termination of the zone of partial
melting rather than by a change in grain size. Furthermore, grain size variations with depth
are insufficient to explain the seismic LVZ in the absence of water and/or melt and can only
be applied to wet mantle.
The low-velocity zone, as its name indicates, is associated with negative velocity gradient:
velocity in the upper mantle across the upper boundary of the LVZ decreases with increasing
Temperature (°C)
600
continental geotherm
0
20
40
60
Pr
es
su
re
 (k
ba
r)
80
100
upper mantle adiabat
olivine
pure
enstatite
continental oceanic
100
200
D
epth (km)
300
400
60% olivine +
40% AI-saturated enstatite
AI-saturated
enstatite
LVZ
LVZ
120
140
1000
oceanic geotherm
1400 0 500 1000 1500 2000 2500 0 500 1000 1500 2000 2500 3000
Water solubility (ppm H2O) Water solubility (ppm H2O)