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 (T s
) from about 10 to 40 km, and are not in the mantle. Variations in T s correlate with variations in the effective elastic thickness (T e
), both of them having similar values, although T is usually the smaller of the two. These observations suggest that the lower crust, at least in some e 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 (T e
) on the continents is usually close to the thickness of the seismogenic crust T s 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.
found no convincing evidence that seismicity occurs anywhere in the continental mantle except for a single, small (M
L
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
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
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).
Geology; June 2000; v. 28; no. 6; p. 495–498; 5 figures.
495

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 T e 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 T e earthquakes are restricted to the upper crust. There is an obvious correlation between T e thickness (T s
), with T e and the seismogenic 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, consisting 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 T e
< T s
, it is also necessary to argue that the seismically active layer has long-term weakness while the aseismic part has long-term strength, otherwise T e would be greater than T s
. 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 T e of ~35 km, while the 1–2 km elevation of the Iranian plateau is supported by forelands with T e of ~10–15 km. Higher mountains require greater support, with estimates of the buoyancy force needed to support Tibet being about 5 × 10 12 N m –1 (e.g., Molnar and Lyon-
Caen, 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 ( T e
) 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).
T e 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
496 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.
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Figure 4. Stacked gravity profiles across (A) Kopet Dag and (B) Dzungaria foreland basins, at same scale. Solid line is average of stacked profiles g m
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
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GEOLOGY, June 2000 497

Figure 5. Cross sections through Iranian plateau and through Karakoram, western Tibet, and Tien
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Manuscript received December 6, 1999
Revised manuscript received March 20, 2000
Manuscript accepted March 28, 2000
498 Printed in USA GEOLOGY, June 2000