Certain volcanic hotspots at the Earth's surface appear to be essentially fixed with respect to the Earth's deep interior and to provide an absolute reference frame for plate motions for the past 40 Ma (Section 5.5). The fixed nature of hotspots such as Hawaii, first suggested by Wilson (1963), led Morgan (1971) to propose that they are located over plumes of mantle material upwell-ing from the lower mantle or even the core-mantle boundary. The plume hypothesis has been, and continues to be, controversial because it has proven difficult to provide unequivocal evidence of such plumes (Foulger & Natland, 2003). There is now, however, a growing body of evidence, from both modeling and observational data, that some hotspots at the Earth's surface may be fed by narrow plumes of high temperature, low viscosity material rising from essentially fixed points on the core-mantle boundary. There is also theoretical and empirical evidence that the source material for other hotspots may be derived from much shallower depths, within the mantle transition zone or the uppermost part of the lower mantle, or even from immediately beneath the lithosphere; the latter being a passive response to various forms of litho-spheric break-up (Anderson, 2000). The suggestion that there are three types of hotspot, in terms of their depth of origin, has been deduced by Courtillot et al. (2003), mostly by a consideration of the roles of the three potential thermal boundary layers in the mantle. However, there is substantial support for this from the results of seismic tomography (Montelli et al., 2004a, 2004b). Hotspots that are underlain by low seismic velocities in the upper mantle only appear to be limited in number. Examples in Fig. 5.7 are Bowie, Cobb, Galapagos, East Australia and, surprisingly perhaps, Iceland, although it is underlain by a very large upper mantle anomaly (Montelli et al., 2004b). Iceland is anomalous also in terms of geochemical indicators of deep mantle origin, notably the ratios of 3He/4He and 186Os/187Os in the lavas (Foulger & Pearson, 2001; Brandon, 2002). Montelli et al. (2004b) describe the tomographic anomaly beneath Yellowstone as being virtually nonexistent.
Laboratory experiments indicate that the peaks that develop on the ULVZ in the D" layer, where it is sufficiently dense, form where ridges between embay-ments in the surface of the ULVZ meet at an elevated point or arête (Jellinek & Manga, 2004). As a result the upwelling of the thermal boundary layer that produces these peaks is focused into a narrow, cylindrical conduit. The temperature difference between this plume and the surrounding mantle is probably 200-300°C, implying more than two orders of magnitude reduction in the viscosity across the boundary layer between them. If partial melt is present in the upwell-ing thermal boundary layer this would also lower the viscosity. Partial melt entrained from the ULVZ may be required to explain the osmium isotopic ratios in certain hotspot lavas that are thought to indicate that the source of the osmium is the outer core (Brandon et al., 1998).
Numerical and analogue models of these hot, low viscosity plumes, originating in the deep mantle, suggest that plume shape and mobility are controlled by the magnitude of the viscosity contrast with the surrounding mantle (Kellogg & King, 1997; Lowman et al., 2004; Lin & van Keken, 2006). As the contrast increases, the plume conduit becomes narrower and its head becomes broad and mushroom-shaped as hot material is able to move upward more efficiently (Fig. 12.14). This model, with a mushroom-shaped plume head and a long, thin tail extending to the depth of origin, has achieved widespread application. Nevertheless, numerical models also predict a great variety of plume shapes and sizes in cases where density contrasts due to chemical variations in the lowermost mantle are incorporated into models of plume formation (Section 12.9) (Farnetani & Samuel, 2005; Lin & van Keken, 2006).
In general, the model of a narrow, mushroom-shaped plume fits well with the initial expression of some hotspots in terms of continental flood basalts or oceanic plateaux, reflecting the arrival of the plume head beneath a thinned lithosphere, and the subsequent trace of the hotspot, in the form of a volcanic ridge or line of volcanoes, produced by the tail. Courtillot et al. (2003) suggest that these hotspots be termed primary hotspots (Section 5.5). They also suggested that the lifespan of primary hotspots might be approximately 140 Ma. Those initiated within the past 100 Ma, such as Afar and Reunion, are still active; those that are 100 to 140 Ma old, that is Louisville and Tristan, might be failing, and those that formed more than 140 Ma ago, such as Karoo and Siberia, have no active trace. Theoretical arguments predict that such large plume heads and long-lived tails must originate in a thermal boundary layer at great depth, presumably layer D" at the
4 Ma 43 Ma 83 Ma 100 Ma 121 Ma 176 Ma
78 Ma 94 Ma 98 Ma 106 Ma 114 Ma 137 Ma
0 Temperature (°C) 1846
Figure 12.14 Sequences from numerical models, scaled approximately to the mantle, in which a plume grows from a thermal boundary layer. In (a) the viscosity is a function of temperature, and in (b) the viscosity also increases by a factor of 20 at 700km depth. In (b) the plume slows and thickens through the 700km discontinuity but then narrows and speeds up in the low viscosity upper layer (from Davies, 1999. Copyright © Cambridge University Press, reproduced with permission).
core-mantle boundary. It has been estimated that a major plume may be fed for 100 Ma from a volume of layer D" only tens of kilometers thick and 500-1000 km in diameter.
If major upwellings, such as those beneath southern Africa and the south Pacific, reach the base of the transition zone they may well form a thermal boundary layer at this depth from which secondary plumes may originate (Brunet & Yuen, 2000; Courtillot et al., 2003). These would be relatively short-lived and without initial flood basalts but may well account for the hotspots on the south Pacific superswell such as the Society and Cook-Austral islands, Samoa, Pitcairn, and Caroline (Fig. 5.7) (Adam & Bonneville, 2005). By contrast, there are no plumes within the southern
African superswell although, as in the Pacific, there are several potential deep mantle, or primary, plumes around it (Figs 5.7, 12.11c). This contrast is also reflected in the marked difference in the seismic velocity anomalies in the upper mantle beneath the two areas (Plate 12.2b,d between pp. 244 and 245). The differing characteristics of the African and Pacific super-swells may arise from the fact that the south Pacific upwelling is the remnant of the Cretaceous superplume in this area (Section 5.7). The uplift of southern Africa was also initiated in the mid-Cretaceous, suggesting that major mantle upwellings or superplumes may have a life cycle analogous to, and perhaps related to, the life cycle of the assembly and break-up of supercontinents.
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