Fate of the subducted slab: Ringwood 1991 Model


One of the problems of plate tectonics is the fate of the subducting slab. This can be traced, from seismic evidence, to descend to about 650 km; but the evidence is somewhat conflicting regarding the extent to which it penetrates the dense 650 km discontinuity. (See references by Jordan and Hilst). Because the phase changes with depth are now known in some detail, both for ultramafic mantle material and for subducted basaltic ocean crust, it is possible to calculate their modal compositions with depth. For instance, the modal composition of pyrolite with depth is shown in Fig. 10:

 

Fig. 11 shows the same calculations for basaltic ocean crust. Note that the plate which is subducting is not uniform mantle pyrolite but, because of melting at the ridge axis, it has segregated into a basaltic ocean crust (ca 5 km thick), residual harzburgite (from which the basalts were extracted) underlain by ordinary pyrolite. Knowing the mineral proportions and the densities of the minerals in each of the main rock types, undepleted pyrolite, depleted harzburgite, and basaltic ocean crust, it is then possible to calculate the density changes in each of these rock types with depth. The thermally equilibrated densities for these three rock types with depth are shown in Fig. 12.

Fig. 12. Densities (g/cc) of thermally equilibrated basaltic ocean crust, subducted harzburgite lithosphere compared with undepleted pyrolite mantle to depths of 800 km. Note that the ocean crust is mostly more dense and the harzburgite is less dense than pyrolite down to 650km depth, but then their positions are reversed.

 

Fig. 13. Density differences between basalt - pyrolite and harzburgite - pyrolite as the subducted ocean crust sinks. The basaltic slab becomes less dense than mantle pyrolite in the depth range 650 - 750 km.

 

The important point is that the subducted plate sinks because the basaltic component of the slab (now eclogite) is ca. 0.2 - 0.1 g/cc more dense than the enclosing host pyrolite to depths of 650 km, and exerts a strong 'slab-pull' force at subduction zones. The harzburgite part of the plate may also be slightly more dense initially because it is cold, but is inherently less dense once it has thermally equilibrated with the surrounding mantle pyrolite. However, because of the phase changes in pyrolite at the 670 km discontinuity, the basaltic crust suddenly becomes 0.2 g/cc less dense than the pyrolite in the depth range 650-750 km, whereas the harzburgite component of the slab becomes very slightly more dense. These effects are very clearly shown in Fig. 13. Ringwood (1991) argues that these changes then have the effect of trapping subducted basaltic ocean crust at the 670 km discontinuity, as shown in Fig. 14.

Fig. 14. The effect of density differences is that basaltic ocean crust becomes trapped at the 670km discontinuity.

 

Ringwood has suggested that the slab piles up at the base of the upper mantle, as shown in Fig. 14. By the end of the Archaean (2500 my ago) he envisages that the mantle structure around the 650km discontinuity would be as shown in Fig. 15. This layer is source for diamond-bearing kimberlite magmas according to Ringwood et al. (1992).

 

Fig. 15. Likely mantle structure at the end of the Archaean as a result of subducted mafic ocean crust piling up at the 650 km discontinuity (after Ringwood).
Fig. 17. Comparison between oceanic and continental lithosphere.

 

Assuming constant spreading rates (present day) it can be calculated that, throughout Earth history, the amount of ocean crust which may have accumulated at the 650 km discontinuity would be at least 100 km thick. However because the harzburgite is inherently less dense and potentially more buoyant than the surrounding mantle, then when it heats up it may begin to ascend as blobs or diapirs, as shown in Fig. 16 (below).

 

Fig. 16. Models of mantle differentiation involving storage of ocean crust at the 650 km discontinuity.

 

This has interesting consequences as a mechanism to generate mantle plumes and 'hotspots'. Most plumes need to be generated at a discontinuity, either the 650 km one or at the core-mantle boundary. The mantle model (Fig. 16) shows that these plumes rise and penetrate the lithosphere to become the source of hotspot ocean islands. If these rising diapirs cannot penetrate the lithosphere, they may just add to the base of the lithsophere, and melts may penetrate it and metasomatise and chemically alter it. The ocean lithosphere is young (almost all less than 200 my) whereas the sub-continental lithosphere is older, cooler, thicker and more complex, as shown in both Figs. 16 & 17.

Because of the different thermal regimes, and the influences of plumes, it is likely that there have been differences in the make-up of the lithosphere during Earth history. Fig. 18 (below) shows the likely structure of the modern mature Phanerozoic ocean lithosphere (left), which is regarded as being less depleted with increasing depth. Ocean plateaus (centre) have a very thick ocean crust, with (implicitly) a much more depleted harzburgitic mantle underlying it. In the Archaean (right) one suggestion is that the high mantle temperatures would have led to very high degrees of melting, to produce high-Mg komatiitic lavas, and leaving an extremely depleted pure-olivine dunitic residue. This oceanic structure is much more like that of modern oceanic plateaus, so was there plate tectonics in the Archaean or plateau tectonics (see later PlateLect-F)?