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models use the same type of
regularization for poorly resolved regions, a strong correlation between model results does
not exclude the possibility of other patterns of heterogeneity (in shape and amplitude) that
are consistent with observations.
With these words of caution in mind, a comparison of large-scale tomographic models
allows for recognition of robust, model- and inversion-independent, features of the upper
mantle. In comparing different models, it is important to recognize reliably determined
features which supply important information on the physical properties and structure of the
upper mantle. Due to the limitations of the approach, only well-resolved structures should be
compared between different tomography models.
Body-wave seismic tomography: uncertainty and resolution
Resolution of a tomography model is determined by seismic wavelength, which in turn is
proportional to the seismic wave period and inversely proportional to frequency. Generally,
the higher the frequency, the higher the model resolution, and regional body-wave tomog-
raphy typically utilizes seismic waves with a frequency of 0.1–10 Hz (e.g. Sipkin and
Jordan, 1975; Inoue et al. 1990; Grand, 1994; Su et al., 1994; Masters et al. 1996;
Trampert and Woodhouse 1996). Recent global travel time P-wave tomography models
(e.g. Spakman et al. 1993; Zhou, 1996; Bijwaard et al., 1998) can image the seismic velocity
structure of particular regions of the upper mantle with lateral resolution as detailed as
that in regional tomographic studies (50–100 km) allowing us to distinguish even localized
anomalies such as slabs in the upper mantle of the present subduction zones (e.g. van der
Hilst et al., 1997; Grand, 2002) and lower mantle slabs associated with ancient subduction
zones (e.g. van der Voo et al., 1999b; Fukao et al., 2001). However, lateral resolution of
teleseismic body-wave tomography based on direct arrivals is significantly non-uniform
and is limited by ray path coverage which is good only in tectonically active regions and
areas with dense distribution of seismometers (direct arrivals are assumed to be sensitive
mainly to the upper mantle structure in a narrow zone beneath sources and receivers).
To overcome this problem, teleseismic body-waves multiply reflected from the surface
and turned in the upper mantle have been used in some regional studies (e.g. Grand and
Helmberger, 1985).
Most teleseismic phases arrive at the receiver along near-vertical ray paths, in particular
phases from the core which are widely used in body-wave tomography studies (Figs. 3.75,
3.76). The near-vertical propagation in the upper mantle puts strong limitations on vertical
resolution which at present does not exceed 50–100 km. One way to improve the vertical
resolution is to also use travel times of waves from local earthquakes observed at local
135 3.6 Teleseismic seismology
stations, as these waves tend to travel sub-horizontally in the upper mantle. However, this
approach is limited by local seismicity and also by the tendency of earthquakes to occur only
along distinct fault zones.
Another approach is used in body-wave tomographic models of S. Grand (e.g. Grand,
1994, 2002; Grand et al., 1997), where all velocity anomalies are put ad hoc into the upper
200 km. The approach is validated by the spectrum of mantle velocity anomalies as
revealed by both body-wave and surface-wave global whole mantle tomography: velocity
perturbations with amplitude exceeding 2% are restricted to the shallow mantle and
the core–mantle boundary, while the spectrum of middle-mantle anomalies is “white”
(e.g. Ritsema et al., 2004).
Computations of the spatial resolution of body-wave tomography based on several
traveltime phases have demonstrated the following (Vasco et al., 2003):
Fig. 3.75 Sketch showing ray paths in body-wave (a) and surface-wave (b) tomography. Gray shading – region sampled by
seismic waves. Due to near-vertical propagation of body-waves, they provide low vertical resolution. In contrast,
surface waves smear lateral velocity structure.
Fig. 3.76 Geometric ray paths for the shear wave phases (i.e. body-wave phases that propagate through the mantle with
shear wave speed). These phases (i) propagate through the lower mantle (S) or diffract along the core–mantle
boundary, (ii) reflect once (SS), twice (SSS), or three times (SSSS) off the Earth’s surface, (iii) reflect once (ScS), twice
(ScS2), or three times (ScS3) off the core–mantle boundary, or (iv) propagate as compressional waves through the
core (SKS and SKKS) (from Ritsema and and van Heijst, 2000b).
136 Seismic structure of the lithosphere
* The errors associated with different phases differ significantly and core phases can lead to
greater errors than first arriving P phases. As a result, inclusion of various phases into the
inversion does not guarantee an improvement in model resolution.
* Model parameter resolution is correlated with ray density which is controlled by distri-
bution of seismic events and seismic stations. This leads to highly variable resolution of
tomographic images.
* The best resolved cells lie beneath the southern parts of Eurasia and North America, and a
narrow zone of subduction zones encircling the Pacific Ocean (Fig. 3.77). Localized high
velocity anomalies interpreted as slabs often coincide with narrow zones of high reso-
lution in the mantle along event-station corridors and are a consequence of ray density–
model resolution correlation.
* The mantle structure is moderately resolved beneath mid-ocean ridges and in the northern
parts of Eurasia and North America. The worst resolved cells include the Pacific, Atlantic,
and Indian ocean basins, the Arctic, Siberia, and most of the southern hemisphere except
for Australia, southern Africa, and the Andes. In poorly constrained cells, the errors can
exceed several percent.
* Mantle anisotropy affects ray density–model resolution correlation, in particular beneath
the central Pacific: cells with a high ray density may still be poorly resolved.
* Existing crustal models allow for calculation of local velocity perturbations due to
heterogeneities in crustal thickness and average crustal velocities. Although such correc-
tions are of particular importance in surface-wave modeling (see below in this section),
teleseismic frequency-dependent site effects (as compared to frequency-independent
station corrections commonly used to correct for site effects) may have a significant
effect on teleseismic P-wave amplitudes with the major contribution of amplitude losses
in the sedimentary layer (Zhou et al., 2003).
* Depth “leakage” due to smoothing in tomographic inversions reduces vertical resolution
of the models (see also next section).
Fig. 3.77 Ray path sampling by compressional waves in the upper mantle. The logarithmic scale shows the number of
rays intersecting each 3°× 3° cell (unsampled blocks are shown in white) (from Vasco et al., 2003).
137 3.6 Teleseismic seismology
Surface-wave tomography: uncertainty and resolution
Dispersion, vertical resolution, and depth leakage
The advent of surface-wave tomography methods (Toksöz and Anderson, 1966; Cara, 1979;
Woodhouse and Dziewonski, 1984; Nolet, 1990; Kennett, 1995; Trampert and Woodhouse,
1995, 1996) provided some of the best constraints on the structure of the Earth’s upper
mantle (e.g. Zhang and Tanimoto, 1993; Zielhuis and Nolet, 1994; Laske andMasters, 1996;
Ekström et al., 1997; van der Lee and Nolet, 1997a; Shapiro and Ritzwoller, 2002; Ritsema
et al., 2004; Debayle et al., 2005; Panning and Romanowicz, 2006; Kustowski et al.,
2008a). Surface waves are commonly the strongest arrivals recorded at teleseismic distan-
ces. Seismic surface waves that propagate through the upper mantle include Rayleigh waves
and Love waves. The latter are essentially horizontally polarized shear waves (SH waves)
and their sensitivity