Figure 6.16 (a) Map of a hypothetical oceanic plate hit by a major impact that caused a crater (IC) with a diameter of 300 km. The outer limits of the area of 3*106 km2, in which the impact induced low basal restraint (LBR) are also indicated. It should be noted that the area of LBR is almost certainly extremely conservative. (b) Schematic section through the crater to the edge of the LBR.

the stresses in the plate and also to the boundary stresses in adjacent plates which, in turn, gives rise to changes of speed and direction of motion of the impacted plate.

Let us initially assume that our hypothetical target is mainly comprised of oceanic lithosphere which has a thickness of less than 100 km and, prior to impact, is underlain by a LVZ with an average thickness of perhaps 60 km. This LVZ zone normally restricts the rate of plate movement, because of the inherently high coefficient of viscosity of about 1019 Pa s. If, as the result of a major impact, a significant proportion of this area can be substantially, and extremely rapidly, heated (or 'unloaded'), then the viscosity of the LVZ within the 1000 km radius about GZ will be very materially reduced.

The degree of melting at a distance from GZ of 1000 km or more, will be mainly determined by the two forms of stress-pulse that are generated by the impact. As we saw in the previous chapter, when a large, fast-moving asteroid or comet makes contact with the Earth's surface, it sets in train a complex succession of events.

Let us consider the first phases of deformation, which relate to the passage of a high-intensity, fast-moving compressive stress-pulse which is initiated at the impact zone. We have seen that the magnitude of this transient stress degenerates away from the point of impact (GZ), and causes a series of effects that may range from vapourisation, melting, plastic deformation and, finally, fracturing (see Figure 5.30 and Figure 5.31). As we shall see, depending upon the size of the impacting body, transient cratering may reach down through the crust to the lower lithosphere or even into the asthenosphere. The spreading of the stress-wave generated in the Earth by the impacting body takes a brief, but finite, time. The compressive stress-wave propagates at a velocity of about 8-10 km s-1, so that the time taken to pass through the various zones of deformation, for a large event, will be about 2-3 minutes.

Adjacent to the crater area, significant melting will take place, and because the transient compressive stress-pulse will exceed its elastic limit, the rock mass will deform by plastic flow (which generates a considerable amount of heat). Hence, significant partial melting of the LVZ, in the areas beyond the crater, is likely to occur. Even a small percentage of induced melting in the LVZ will significantly reduce the viscosity of that zone. Moreover, the thickness of the LVZ beyond the limits of the crater is likely to increase as the result of transient pressure-induced melting. Beyond, but still close to the limit of the crater, where the magnitude of the transient stress-pulses are extremely high, the viscosity could be reduced by several orders of magnitude. The basal drag on this specific area of LVZ will be radically, and almost instantaneously, reduced by a comparable magnitude (see Figure 3.13 and relevant equations).

The inner zones shown in Figure 6.16b will be excavated to form a transient crater. However, the plastic and brittle response to the compressive stress-wave can extend to well beyond the crater. Hence, for large impacts, which give rise to, say, a diameter of 300 km, the zone of plastic deformation and fracturing can extend down to a depth well beyond the LVZ, as well as spreading out at the LVZ depth for at least 1000 km beyond GZ. One can infer that plastic deformation of the mantle results in the generation of heat. Consequently, if the plastic deformation extends down to and beyond the LVZ, which is at the PT conditions close to melting, then this deformation can cause further small amounts of melting to take place, thereby lowering the viscosity of the LVZ, at a distance of 1000 km from GZ, possibly by a factor of perhaps 2-3. The overall energy of this stress-pulse event is more than sufficient to supply, instantly, the latent heat to cause melting.

It is known that the development of shear planes in dry rock can cause melting of the fracture surfaces which, in relatively cold rock, give rise to the development of pseudotachylyte (Sibson, 1975). The development of such fractures in hot rocks generates a small but further, significant, additional amount of melt, which would further contribute to a reduction of the viscosity of the LVZ. We have seen in Chapter 5 that the stress-pulse caused by a large impacting body, even when it falls as low as 20 kb, is still able to generate heat in the target rock. Depending upon the velocity on impact, this 20 kb stress level may extend from GZ by as much as 10 or more times the 'diameter' of the impacting body.

We have also noted that the compressive stress-pulse may be followed by a weaker, tensile stress-pulse. Nevertheless, although weaker, this tensile pulse may have a considerable effect upon the mantle rock. For example, if the tensile pulse at any distance from GZ has (for a very brief period) an upward component of magnitude greater than 30 kb it is capable of reducing the effects of gravitational loading to zero, down to a depth of about 100 km. In this brief tensile-stress phase, relatively little deformation may take place. The temperature in the LVZ will be little affected. However, as the ambient gravitationally induced rock pressure (P) will be reduced, perhaps to zero, the original PT conditions in the LVZ (where the rock is already at or close to melting) will cause the tensile stress-pulse to further enhances the degree of melting in the LVZ. Moreover, this is not a situation which is quickly reversed. Heat must be conducted away from the LVZ before the original state can be approached. This will be a slow process.

The combined effects of the compressive and tensile pulses are therefore very large, and will result in a significant area of the LVZ experiencing a further, sudden and highly significant reduction in the coefficient of viscosity. Hence, over an area of at least 3,000,000 km2 this average reduction will completely change the balance of forces in the plate. We estimate that the average viscosity of the LVZ within the zone of 'enhanced LVZ' will be reduced by about 2 orders of magnitude.

In Chapter 3, we noted that the stresses involved in driving lithospheric plates was simply related to the thickness and viscosity of the LVZ (Figure 3.13). For a relatively fast-moving, oceanic plate, the stresses could move the lithosphere over the LVZ at about 7 cm a-1, or, 7 km in 1.0 Ma. If the viscosity of the LVZ is suddenly reduced by an average 2 orders of magnitude over an area of at least 3,000,000 km2, then, if the effects of inertia are ignored, we suggest that a speed of lithospheric motion of several hundred km per Ma could be attained. This would permit a relatively sudden increase in the speed of readjustment to plate motion in the regions surrounding the major impact, that would give rise to a change in direction and/or rate of plate motion, so that considerable readjustment, regarding direction of plate movement, could be largely completed in a period of between 20,000-50,000 a.

These arguments have been based on the assumption that the 'enhanced LVZ' occupies a very large proportion of the area of the plate. If the enhanced LVZ occupies only a fraction of the plate, the potential rate of change of motion will be correspondingly reduced. The boundary stresses on the plate will also influence the final rate and direction of movement of the plate. It is, of course, difficult to quantify these movement rates precisely. Fortunately, nature has presented an example which indicates that high plate velocities may be engendered by a major impact. We refer the reader to the track associated with the Paraña continental flood basalt, discussed earlier.

It was shown (Figure 6.8a) that in the period 139-135.1 Ma, S America moved northward at a rate of 7.6 km Ma-1. Between 135.1 and 134 Ma, the rate of motion was about 150 km Ma-1. These figures, we suggest, completely support the argument presented above.

We note that the Paraña event relates to movement of a continental area, where, because of the, commonly observed, variable thickness of lithosphere this is likely to provide considerably higher resistance to horizontal movement than that of an oceanic plate. Moreover, there is little evidence of an LVZ layer beneath continental lithospheres. However, where PT conditions below the continental lithosphere are appropriate, a major impact can induce a small percentage of melt such that a layer comparable with an oceanic LVZ may be initiated. Even so, it is likely that basal drag beneath a continent will usually be higher than that for an oceanic plate. Hence, we infer that early rates of adjustment of oceanic plates may be many tens of km per year.

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