EARTH AS AN EVOLVING PLANETARY SYSTEM Part 4

The Lithosphere 127 Figure 4.9 Mg number 93 92 Archean 91 90 Post-Archean 89 versus modal olivine in post-Archean mantle xeno-liths and ophiolite ultra-mafics and in Archean lithosphere xenoliths from South Africa. Mg number = Mg/Mg + Fe molecular ratio. Modified from Boyd (1989). 100 80 60 40 Modal Olivine (%) the Earth just after planetary accretion (Sun and McDonough, 1989). Compared with primitive mantle, incompatible element distributions in the mantle lithosphere and depleted mantle are distinct (Table 4.1). The striking depletion in the most incompatible elements (Rb, Ba, Th, etc.) in depleted mantle (Fig. 4.10), as represented by ophiolite ultramafics and mantle compositions calculated from ocean-ridge basalts, reflects the removal of basaltic liquids enriched in these elements, perhaps early in the Earth’s his-tory (Hofmann, 1988). In striking contrast to depleted mantle, lithosphere mantle shows prominent enrichment in the most incompatible elements. Spinel lherzolites from post-Archean lithosphere show a positive Nb-Ta anomaly, suggesting that they represent plume material plastered on the bottom of the lithosphere (McDonough, 1990). Archean garnet lherzolites commonly show textural and mineralogical evidence for metasomatism (i.e., modal metasomatism), such as veinlets of amphibole, micas, and other secondary minerals (Waters and Erlank, 1988). The high content of the most incompatible elements in modally metasomatized xenoliths (Fig. 4.10) probably records metasomatic additions Post-Archean Lithosphere Archean Lithosphere Metasomatized Archean Lithosphere 10.0 Depleted Mantle 1.0 Figure 4.10 Primitive mantle (PM), normalized, incompatible element distri-butions in subcontinental lithosphere and depleted mantle. Primitive mantle values from Sun and McDonough (1989); data from Nixon et al. (1981), Hawkesworth et al. (1990), Wood (1979), and miscella-neous sources. 0.1 Rb Ba Th K Ta Nb La Ce Sr P Hf Zr Eu Ti Yb Y 128 The Mantle of these elements to the lithosphere. In some cases, such as in the xenoliths from Kimberley in South Africa (Hawkesworth et al., 1990), these additions occurred after the Archean. Thickness of Continental Lithosphere The continental lithosphere varies considerably in thickness depending on its age and mechanism of formation. S-wave tomographic studies of the upper mantle have been most definitive in estimating the thickness of continental lithosphere (Grand, 1987; Polet and Anderson, 1995). Most post-Archean lithosphere is 100 to 200 km thick, and litho-sphere beneath Archean shields is commonly more than 300 km thick. Rheological models suggest thicknesses in these same ranges (Ranalli, 1991). In an S-wave tomo-graphic cross-section around the globe, high-velocity roots underlie Archean crust, as in northern Canada, central and southern Africa, and Antarctica (Fig. 4.11). The base of the lithosphere in these and other areas overlain by Archean crust may nearly reach the 410-km discontinuity. Under Proterozoic shields, however, lithospheric thicknesses rarely exceed 200 km. Consideration of elongation directions of Archean cratons relative to directions of modern plate motions suggests that the thick Archean lithosphere does not aid or hinder plate motions (Stoddard and Abbott, 1996). As expected, hotspots (plumes), such as Hawaii and Iceland, are associated with slow velocities between 50 and 200 km deep (Fig. 4.11). Thermal and geochemical modeling has shown that the lithosphere can be thinned by as much as 50 km by extension over mantle plumes (White and McKenzie, 1995). Isotopic and geochemical data from mantle xenoliths indicate that the mantle litho-sphere beneath Archean shields formed during the Archean and that it is chemically distinct from post-Archean lithosphere. Because of the buoyant nature of the depleted Figure 4.11 S-wave velocity distribution in the upper mantle along a great circle passing through Hawaii and Iceland. Darker shades indicate faster veloc-ities. The map shows the location of the great circle and major hotspots (black dots). From Zhang and Tanimoto (1993). 25 100 200 300 400 50−180 −120 −60 0 −4 60 120 180 4 The Lithosphere 129 Archean lithosphere, it tends to ride high compared with adjacent Proterozoic lithosphere, as shown by the extensive platform sediment cover on Proterozoic cratons compared with Archean cratons (Hoffman, 1990). The thick roots of Archean lithosphere often survive later tectonic events and thermal events, such as continental collisions and supercontinent rifting. However, mantle plumes or extensive later reactivation can remove the thick lith-osphere keels, as for instance is the case with the Archean Wyoming province in North America and the north China craton. Seismic Anisotropy P-wave velocity measurements at various orientations to rock fabric show that differ-ences in mineral alignment can produce significant anisotropy (Ave’Lallemant and Carter, 1970; Kumazawa et al., 1971). Vp differences of more than 15% occur in some ultramafic samples and are related primarily to the orientation of olivine grains. Seismic-wave anisotropy in the mantle lithosphere beneath ocean basins may be produced by recrystallization of olivine and pyroxene accompanying seafloor spreading with the [100] axes of olivine and [001] axes of orthopyroxene oriented normal to ridge axes (the higher velocity direction) (Estey and Douglas, 1986). Supporting evidence for alignment of these minerals comes from studies of ophiolites and upper mantle xenoliths, and flow patterns in the oceanic upper mantle can be studied by structural mapping of olivine ori-entations (Nicholas, 1986). The mechanism of mineral alignment requires upper mantle shear flow, which aligns minerals through dislocation glide. The crystallographic glide systems have a threshold temperature necessary for recrystallization of about 900° C, which yields a thermally defined lithosphere depth similar to that deduced from seismic data (~100 km). Creep actively maintains mineral alignment below this boundary in the LVZ, and it is preserved in a fossil state in the overlying lithosphere. The subcontinental lithosphere also exhibits seismic anisotropy of S-waves parallel to the surface of the Earth. This is evidenced by S-wave splitting, where the incident wave is polarized into two orthogonal directions traveling at different velocities (Silver and Chan, 1991). As with the oceanic lithosphere, this seismic anisotropy appears to be caused by the strain-induced preferred orientation of anisotropic crystals such as olivine. Seismic and thermal modeling indicate that the continental anisotropy occurs within the lithosphere at depths from 150 to 400 km. The major problem in the subcontinental lithosphere has been to determine how and when such alignment occurred in tectonically stable cratons. Was it produced during assembly of the craton in the Precambrian, or is it a recent feature caused by deformation of the base of the lithosphere as it moves about? In most continental sites, the azimuth of the fast S-wave has been closely aligned with the direction of absolute plate motion for the last 100 My (Silver and Chan, 1991; Vinnik et al., 1995) (Fig. 4.12). This coincidence suggests that the anisotropy is not a Precambrian feature but results from resistive drag along the base of the lithosphere. Supporting this interpreta-tion, seismic anisotropy does not correlate with single terranes in Precambrian crustal provinces. These provinces were assembled in the Precambrian by terrane collisions, and if anisotropy was acquired at this time, the azimuths should vary from terrane to terrane 130 The Mantle Figure 4.12 Fast S-wave velocity directions in the subcontinental lithosphere compared with motion directions of modern plates (bold arrows). Modified from Silver and Chan (1991). according to their preassembly deformational histories. Instead, the seismic anisotropies show a uniform direction across cratons, aligned parallel to modern plate motions (Fig. 4.12). Thermal Structure of Precambrian Continental Lithosphere It has long been known that heat flow from Archean cratons is less than that from Proterozoic cratons (Fig. 4.13). Two explanations for this relationship have been pro-posed (Nyblade and Pollack, 1993; Jaupart and Mareschal, 1999): 1. There is greater heat production in Proterozoic crust than in Archean crust. 2. A thick lithospheric root beneath the Archean lithosphere is depleted in radiogenic elements. The Proterozoic upper continental crust appears to be enriched in K, U, and Th, whose isotopes produce most of the heat in the Earth, relative to its Archean counterpart. Using estimates of the concentration of these elements in the crust (Condie, 1993), only part of the difference in heat flow from Proterozoic and Archean lithosphere can be explained. The Lithosphere 131 80 60 40 200 1.0 2.0 Figure 4.13 Heat flow versus age of Precambrian lithosphere. The width of each box shows the age range, and the height is one standard deviation of the mean heat flow. From Nyblade and Pollack 3.0 4.0 (1993). Crustal Age (Ga) Thus, it would appear that the thick root beneath the Archean cratons must also be depleted in radiogenic elements and contribute to the difference in heat flows. Age of Subcontinental Lithosphere It is important in terms of crust–mantle evolution to know whether the thick lithospheric roots beneath Archean cratons formed in the Archean in association with the overlying crust or whether they were added later by underplating. Although in theory scientists can use mantle xenoliths to isotopically date the lithosphere, because later deformation and metasomatism may reset isotopic clocks, ages obtained from xenoliths are generally too young. Some xenoliths give the isotopic age of eruption of the host magma. What scientists really need to determine the original age of the subcontinental lithosphere are minerals that did not recrystallize during later events or an isotopic system that was not affected by later events. At this point, diamonds and Os isotopes enter the picture. Diamonds, which form at depths greater than 150 km, are resistant to recrystallization at lithosphere temperatures. Sometimes they trap silicate phases as they grow, shielding these minerals from later recrystallization (Richardson, 1990). Pyroxene and garnet inclusions in dia-monds, which range from about 50 to 300 microns in size, have been successfully dated by the Sm-Nd isotopic method and appear to record the age of the original ultramafic rock. Often more than one age is recorded by diamond inclusions from the same kim-berlite pipe, as with the Premier pipe in South Africa. Diamonds in this pipe with garnet-opx inclusions have Nd and Sr mineral isochron ages older than 3 Gy, suggesting that the mantle lithosphere formed in the early Archean when the overlying crust formed (Richardson et al., 1993). Those diamonds with garnet-cpx-opx inclusions record an age of 1.93 Ga, and those with garnet-cpx (eclogitic) record an age of 1.15 Ga, only slightly older than kimberlite emplacement. The younger ages clearly indicate multiple events in the South African lithosphere. Diamond inclusion ages from lithosphere xenoliths from Archean cratons in South Africa, Siberia, and Western Australia indicate that the litho-sphere in these regions is also Archean. The Re-Os isotopic system differs from the Sm-Nd and Rb-Sr systems in that Re is incompatible in the mantle but Os is compatible. In contrast, in most other isotopic systems, both parent and daughter elements are incompatible. Hence, during the early ... - tailieumienphi.vn