Earthquake focal depths, effective elastic thickness,
and the strength of the continental lithosphere
A. Maggi
J.A. Jackson
D. McKenzie
K. Priestley
University of Cambridge, Department of Earth Sciences, Bullard Laboratories, Madingley Road, Cambridge CB3 OEZ, UK
ABSTRACT
Almost all earthquakes on the continents are confined within a crustal layer that varies in
thickness (Ts ) from about 10 to 40 km, and are not in the mantle. Variations in Ts correlate with
variations in the effective elastic thickness (Te ), both of them having similar values, although Te
is usually the smaller of the two. These observations suggest that the lower crust, at least in some
places, is stronger than the mantle beneath the Moho, contrary to most models of continental
rheology. Thus the strength of the continental lithosphere is likely to be contained within the
seismogenic layer, variations in the thickness of this strong layer determining the heights of the
mountain ranges it can support. The aseismic nature of the continental mantle and the lower
crustal seismicity beneath some shields are probably related to their water contents.
Keywords: focal depth, elastic thickness, seismogenic crust, continental lithosphere.
INTRODUCTION
The two most accessible indicators of lithospheric strength are the focal depth distribution
of earthquakes and the association of gravity
anomalies with topography. Most earthquakes in
the continental lithosphere occur in the upper
crust, the lower crust being relatively, or even
completely, aseismic (Chen and Molnar, 1983;
Chen, 1988). While earthquakes certainly occur
in the oceanic mantle lithosphere (Wiens and
Stein, 1983), earthquakes in the continental
mantle lithosphere are known to be rare, but were
thought to exist and to represent an important
strength contrast between a weak lower crust
and relatively strong upper mantle (Chen and
Molnar, 1983; Chen, 1988; Wong and Chapman,
1990). In a study of continental gravity and
topography, McKenzie and Fairhead (1997) concluded that the effective elastic thickness (Te ) on
the continents is usually close to the thickness of
the seismogenic crust Ts allowing the simple
interpretation that the strength of the lithosphere
resides in that layer, although they were unable
to estimate the depth to the top of any elastic
layer. In this paper, we argue that (1) contrary to
popular belief, the upper mantle beneath the
continents has relatively little strength; (2) the
effective strength of the continental lithosphere
is in the seismogenic crust; and (3) some modifications to our ideas of lithosphere rheology and
what influences it are required.
FOCAL DEPTH DISTRIBUTIONS ON
THE CONTINENTS
In a separate study (Maggi et al., 2000) we reviewed the seismological evidence for earthquake depth distributions on the continents. We
Geology; June 2000; v. 28; no. 6; p. 495–498; 5 figures.
found no convincing evidence that seismicity
occurs anywhere in the continental mantle except
for a single, small (ML 3.8), isolated event at
90 km depth reported in northern Utah (Zandt
and Richins, 1979). In several parts of Africa,
Asia, and North America seismicity continues
throughout the crust to Moho depths (Fig. 1). In
East Africa, the Tien Shan, and the Indian shield
some of the deepest earthquakes in the crust occur
so close to the Moho that uncertainties in their
depth and crustal thickness allow them to be in
the uppermost mantle. However, because other
earthquakes in the same regions indicate that the
lower crust is seismically active, we suspect that
all the events occurred above the Moho. Earthquakes in the mantle beneath continents do occur
Figure 1. Histograms of earthquake focal depths determined by modeling of long-period teleseismic P (primary) and SH (secondary horizontal) seismograms (solid bars).White bar in North
India (G) is depth determined from short period depth phases in Shillong Plateau by Chen and
Molnar (1990). White bars in Tibet (C) are subcrustal earthquakes, but not necessarily in mantle
of continental origin. Approximate Moho depths are indicated by dashed lines. Focal depth and
Moho data are from various sources, including Nelson et al. (1987), Molnar and Lyon-Caen
(1989), Foster and Jackson (1998), Mangino et al. (1999), and Maggi et al. (2000). Focal depths
based on arrival times recorded at local seismic networks have also found seismicity throughout crust in several parts of North America (e.g., Wong and Chapman, 1990).
495
sisting of all or some of the crust, and that the
mantle part of the continental lithosphere is
relatively weak. The alternative is to argue that,
because seismicity is an indicator of frictional
instability rather than strength, the continental
mantle could still be strong despite being aseismic.
The problem then is that, because Te < Ts , it is
also necessary to argue that the seismically active
layer has long-term weakness while the aseismic
part has long-term strength, otherwise Te would
be greater than Ts . We cannot rule this possibility
out, but it seems to us improbable and unnecessarily complicated. We prefer the simpler explanation that the long-term strength resides in the
seismogenic layer.
A striking correlation exists between the
heights of the mountainous regions and the
strengths of the forelands that support them
(Fig. 5), the 4–5 km elevation of the Karakoram,
western Tibet, and Tien Shan being supported by
forelands with Te of ~35 km, while the 1–2 km
elevation of the Iranian plateau is supported by
forelands with Te of ~10–15 km. Higher mountains require greater support, with estimates of
the buoyancy force needed to support Tibet being
about 5 × 1012 N m–1 (e.g., Molnar and LyonCaen, 1988), equivalent to average deviatoric
stresses of ~120 MPa if contained within the
40 km elastic layer. From the profiles in Figure 4
we estimate that similar values of deviatoric stress
are required to maintain the gravity anomalies
associated with flexure. The conclusion must be
that relatively intact shields, like northern India,
are particularly strong. We suspect that the Persian
Gulf, which is part of the margin of the Arabian
shield that was stretched in the Mesozoic (Koop
and Stoneley, 1982), is currently too weak to support mountains higher than the Zagros. In this
way we expect the strength of the bounding foreland to be a major control on the height of the
adjacent mountains.
WHAT CONTROLS LITHOSPHERE
STRENGTH?
In the oceans the maximum depth of intraplate
earthquakes and the effective elastic thickness
both correlate with the age of the lithosphere,
suggesting that the main control on strength is
temperature (Watts et al., 1980; Chen and Molnar,
1983; Wiens and Stein, 1983). The concentration
of seismicity in the continental crust in all areas,
and its restriction to depths shallower than 20 km
in many, together with some correlation of maximum focal depth with surface heat flow, suggests
that temperature is also the main control on the
continents (e.g., Brace and Byerlee, 1970; Sibson,
1982; Chen and Molnar, 1983; Chen, 1988). It
is more problematic to explain why, given that
some old (possibly cold) continental shields are
seismically active throughout the crust, the
mantle beneath them is not seismically active as
well, because it clearly has that capability in the
oceans. A possible explanation relies on the rheological effects of water contained in hydrous
minerals, where even small concentrations of a
few tens of parts per million can greatly decrease
strength (e.g., Mackwell et al., 1998). Many
nodules from the continental mantle have been
metasomatized by small amounts of a melt that is
very enriched in incompatible elements such as
hydrogen (Harte et al., 1993). Such metasomatism is likely to be more widespread within the
continental lithosphere than the oceanic lithosphere, simply because the mean age of the continents is about 20 times that of the oceans. In
many regions the lower continental crust is
hydrous, weak, and aseismic, but in others it may
be dry because it has been partially melted at
temperatures of ~700 °C to generate granitic
magma and leave behind dry mafic granulites.
Because the solidus temperature of hydrous basic
rock is lower than that of hydrous peridotite
(~1000 °C), water can be preferentially removed
from the lower crust. The percolation of a melt
through a matrix is determined by its connectivity, which is controlled by the dihedral angle
(Θ) between solid and liquid (Waff and Bulau,
1979). If Θ > 60° melt cannot percolate at small
melt fractions. The mantle is dominantly olivine,
with a dihedral angle of ~45°. It is therefore permeable to melt. In contrast, olivine is generally
not present in crustal granulites, and therefore the
permeability will be controlled by the dihedral
angle of pyroxene. Toramaru and Fujii (1986)
estimated that the dihedral angle between orthopyroxene and melt is 76°, and that for clinopyroxene and melt is 98°. Both values are therefore large enough to prevent percolation, and
hence rehydration of the lower crust by upper
mantle melts.
Figure 2. Map showing areas where effective elastic thickness (Te ) has been calculated either
from stacked gravity profiles in foreland basins (open regions) or from admittance between
topography and gravity in frequency domain (cross-hachures), using techniques in McKenzie
and Fairhead (1997). Te values are given alongside in kilometers, with round brackets indicating
values that are less well constrained. Dotted lines show location of profiles in Figure. 5.
CONCLUSIONS
For almost 20 years the popular view of continental strength profiles has consisted of a weak
lower crust sandwiched between relatively strong
to depths of ~100 km in regions where subduction is active now or has been in the late Tertiary,
such as the Atlas, Karakoram, or southernmost
Tibet, but these earthquakes could be in oceanic
rather than continental lithosphere. If the view of
continental seismicity presented by Maggi et al.
(2000) is correct, then seismicity in the continental mantle is unimportant, and we examine
the consequences in the following.
EFFECTIVE ELASTIC THICKNESS ON
THE CONTINENTS
Figure 2 shows a summary of elastic thickness
determinations in the Alpine-Himalayan region
and central Asia, obtained using either the admittance between free-air gravity and topography or
from directly modeling flexural signals in the
free-air gravity profiles across foreland basins.
The methods were described fully in McKenzie
and Fairhead (1997). Note particularly the high
values of Te for the northern Tien Shan and Himalayan forelands (31 km and 37 km) compared
with the more typical values of 5–15 km elsewhere. These high values are evident in visual
inspection of the gravity maps (Fig. 3) as well as
in the profiles and their misfit curves (Fig. 4).
Most remarkably, in both the northern Tien Shan
and northern India earthquakes occur to depths of
40–50 km in the continental crust, whereas in all
the places with smaller Te earthquakes are restricted to the upper crust. There is an obvious
correlation between Te and the seismogenic
thickness (Ts ), with Te usually being the smaller
of the two (Figs. 1 and 2).
STRENGTH OF THE CONTINENTAL
LITHOSPHERE AND THE HEIGHT OF
MOUNTAIN BELTS
The simplest interpretation of the seismicity
and gravity data is that the strength of the continental lithosphere is in its seismogenic part, con-
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GEOLOGY, June 2000
Figure 3. Maps of free-air gravity showing foreland flexures associated with (A) Kopet Dag and (B) Dzungaria (northern Tien Shan), at same scale.
layers in the upper crust and mantle. We now
believe this view to be incorrect. Earthquake focal
depths and gravity anomalies instead suggest that
the strength of the continents resides in the seismogenic layer within the crust, and that the continental mantle lithosphere is relatively weak. Spatial
variations in continental strength are probably related to temperature structure and the presence or
absence of small amounts of water.
REFERENCES CITED
Brace, W.F., and Byerlee, J.D., 1970, California earthquakes: Why only shallow focus?: Science,
v. 168, p. 1573–1576.
Chen, W.-P., 1988, A brief update on the focal depths of
intracontinental earthquakes and their correlations with heat flow and tectonic age: Seismological Research Letters, v. 59, p. 263–272.
Chen, W.-P., and Molnar, P., 1983, Focal depths of
intracontinental and intraplate earthquakes and
their implications for the thermal and mechanical
properties of the lithosphere: Journal of Geophysical Research, v. 88, p. 4183–4214.
Chen, W.-P., and Molnar, P., 1990, Source parameters of
earthquakes and intraplate deformation beneath
the Shillong Plateau and northern Indoburman
ranges: Journal of Geophysical Research, v. 95,
p. 12,527–12,552.
Foster, A.N., and Jackson, J.A., 1998, Source parameters of large African earthquakes: Implications for
crustal rheology and regional kinematics: Geophysical Journal International, v. 134, p. 422–448.
Harte, B., Hunter, R.H., and Kinny, P.D., 1993, Melt
geometry, movement and crystallization, in relation to mantle dikes, veins and metasomatism:
Royal Society of London Philosophical Transactions, ser. A, v. 342, p. 1–21.
GEOLOGY, June 2000
Figure 4. Stacked gravity profiles across (A) Kopet Dag and (B) Dzungaria foreland basins, at
same scale. Solid line is average of stacked profiles gm , and gray band shows ±1 σ range. Dotted
lines are modeled gravity gc matched by bending elastic plate of thickness Te . C: Misfit
(
)
2
1 N 
g m – gc σ as function of Te for both regions. In both maps and profiles
∑

N n = 1 
greater wavelength of Dzungaria flexural signal is clear, requiring Te of ~31 km, compared with
only ~12 km for Kopet Dag.
H =
497
Figure 5. Cross sections through Iranian plateau and through Karakoram, western Tibet, and Tien
Shan along lines in Figure 2.Topography is exaggerated relative to depth below sea level. Crosshachured regions show approximate thickness of elastic layer. Gradational change in Moho
depth between mountains and foreland is schematic, but values under forelands are constrained
by receiver function data reported in Mangino et al. (1999) and Maggi et al. (2000).
Koop, W.J., and Stoneley, R., 1982, Subsidence history
of the Middle East Zagros basin, Permian to recent: Royal Society of London Philosophical
Transactions, ser. A, v. 305, p. 149–168.
Mackwell, S.J., Zimmerman, M.E., and Kohlstedt, D.L.,
1998, High-temperature deformation of dry diabase with application to tectonics on Venus: Journal of Geophysical Research, v. 103, p. 975–984.
Maggi, A., Jackson, J.A., Priestley, K., and Baker, C.,
2000, A re-assessment of focal depth distribution
in southern Iran, the Tien Shan and northern India:
Do earthquakes really occur in the continental
mantle?: Geophysical Journal International, v. 143
(in press).
498
Mangino, W., Priestley, K., and Ebel, J., 1999, The
receiver structure beneath the China Digital
Seismograph Network stations: Seismological
Society of America Bulletin, v. 89, p. 1053–1076.
McKenzie, D., and Fairhead, D., 1997, Estimates of the
effective elastic thickness of the continental lithosphere from Bouger and free air gravity anomalies: Journal of Geophysical Research, v. 102,
p. 27,523–27,552.
Printed in USA
Molnar, P., and Lyon-Caen, H., 1988, Some physical
aspects of the support, structure and evolution of
mountain belts, in Clark, S.P., Jr., ed., Processes
in continental and lithospheric deformation: Geological Society of America Special Paper 218,
p. 179–207.
Molnar, P., and Lyon-Caen, H., 1989, Fault plane solutions of earthquakes and active tectonics of the
Tibetan plateau and its margins: Geophysical
Journal International, v. 99, p. 123–153.
Nelson, M.R., McCaffrey, R., and Molnar, R., 1987,
Source parameters for 11 earthquakes in the Tien
Shan, Central Asia, determined by P and SH
waveform inversion: Journal of Geophysical Research, v. 92, p. 12,629–12,648.
Sibson, R.H., 1982, Fault zone models, heat flow and
depth distribution of earthquakes in the continental crust of the United States: Seismological
Society of America Bulletin, v. 72, p. 151–163.
Toramaru, A., and Fujii, N., 1986, Connectivity of melt
phase in a partially molten peridotite: Journal of
Geophysical Research, v. 91, p. 9239–9252.
Waff, H.F., and Bulau, J.R., 1979, Equilibrium fluid
distribution in an ultramafic partial melt under
hydrostatic stress conditions: Journal of Geophysical Research, v. 84, p. 6109–6114.
Watts, A.B., Bodine, J.H., and Ribe, N., 1980, Observations of flexure and the geological evolution of the
Pacific Ocean basin: Nature, v. 283, p. 532–537.
Wiens, D.A., and Stein, S., 1983, Age dependence of
intraplate seismicity and implications for lithospheric evolution: Journal of Geophysical Research, v. 88, p. 6455–6468.
Wong, I.G., and Chapman, D.S., 1990, Deep intraplate
earthquakes in the western United States and their
relationship to lithospheric temperatures: Seismological Society of America Bulletin, v. 80,
p. 589–599.
Zandt, G., and Richins, W.D., 1979, An upper mantle
earthquake beneath the middle Rocky Mountains
in NE Utah [abs.]: Earthquake Notes, v. 50,
p. 69–70.
Manuscript received December 6, 1999
Revised manuscript received March 20, 2000
Manuscript accepted March 28, 2000
GEOLOGY, June 2000