EARTH AS AN EVOLVING PLANETARY SYSTEM Part 4 ppsx

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 The Lithosphere 127 89 90 91 92 93 Mg Number of Olivine 100 80 60 40 Modal Olivine (%) Archean Post-Archean Figure 4.9 Mg number 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). 0.1 1.0 10.0 Ba Post-Archean Lithosphere Archean Lithosphere Metasomatized Archean Lithosphere Depleted Mantle PM-Normalized Value Rb K TaTh Nb La Ce Sr P Hf Zr Eu Ti Yb Y 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. 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 128 The Mantle −180 −120 −60 0 60 120 180 −44 500 400 300 200 100 25 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). 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 The Lithosphere 129 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. 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). 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 The Lithosphere 131 1.0 2.0 3.0 4.00 20 40 60 80 Crustal Age (Ga) Heat Flow (mW/m 2 ) 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 (1993). magmatic event that left the Archean mantle lithosphere as a restite, Re was completely or largely extracted from the rock and Os was unaffected (Carlson et al., 1994; Pearson et al., 2002). When Re was extracted from the rock, the Os isotopic composition was “frozen” into the system; hence, by analyzing a mantle xenolith later brought to the surface, scientists can date the Re depletion event. Calculated Os isotopic ages of xenoliths from two kimberlite pipes in South Africa are shown in Figure 4.14, plotted at depths of origin inferred from thermobarometry. In the Premier and North Lesotho pipes, ages range from 3.3 to 2.2 Ga, similar to the ranges found in xenoliths from pipes in the Archean Siberian craton. Results support the diamond inclusion ages, indicating an Archean age for the thick Archean mantle keels. It is not yet clear whether the range in ages from a given pipe records the range in formation ages of the lithosphere or a series of metasomatic remo- bilization events of lithosphere that occurred about 3.0 Ga. The maximum isotopic ages obtained for mantle keels in the Siberia and South African cratons are similar to the oldest isotopic ages obtained from the overlying crust (Carlson et al., 1994; Pearson et al. 1995; 2002). This suggests that substantial portions of the mantle keels beneath the continents formed at the same time as the overlying crust and that they have remained firmly attached to the crust. The Low-Velocity Zone The LVZ in the upper mantle is characterized by low seismic-wave velocities, high seismic-energy attenuation, and high electric conductivity. The bottom of the LVZ, some- times called the Lehmann discontinuity, has been identified from the study of surface- wave and S-wave data in some continental areas (Gaherty and Jordan, 1995) (Fig. 4.1). This discontinuity, which occurs at depths from 180 to 220 km, appears to be thermally 132 The Mantle 2.8 3.1 3.3 2.2 2.3 2.6 3.3 2.9 2.7 200 km 40 0 Premier North Lesotho CRUST Figure 4.14 Idealized cross-section of the Archean lithosphere in South Africa constructed from mantle xenoliths from two kimberlite pipes. Ages are Re-depletion model ages (in Ga) from Carlson et al. (1994). controlled and to at least partly reflect a change from an anisotropic lithosphere to an isotropic asthenosphere. The LVZ plays a major role in plate tectonics, providing a rela- tively low-viscosity region upon which lithospheric plates can slide with little friction. Because of the dramatic drop in S-wave velocity and the increase in attenuation of seismic energy, it would appear that partial melting must contribute to LVZ production. The probable importance of incipient melting is attested to by the high surface-heat flow observed when the LVZ reaches shallow depths, such as beneath ocean ridges and in con- tinental rifts. Experimental results show that incipient melting in the LVZ requires a minor amount of water to depress silicate melting points (Wyllie, 1971). With only 0.05 to 0.10% water in the mantle, partial melting of garnet lherzolite occurs in the appropri- ate depth range for the LVZ, as shown by the geotherm–mantle solidus intersections in Figure 4.5. The source of water in the upper mantle may be from the breakdown of minor phases that contain water such as amphibole, mica, titanoclinohumite, or other hydrated silicates. The theory of elastic-wave velocities in two-phase materials indicates that only 1% melt is required to produce the lowest S-wave velocities measured in the LVZ (Anderson et al., 1971). If, however, melt fractions are interconnected by a network of tubes along grain boundaries, the amount of melting may exceed 5% (Marko, 1980). The downward termination of the LVZ appears to reflect the depth at which geotherms pass below the mantle solidus (Fig. 4.5). Also possibly contributing to the base of the LVZ is a rapid decrease in the amount of water available (perhaps free water enters high- pressure silicate phases at this depth). The width or even the existence of the LVZ depends on the steepness of the geotherms. With steep geotherms such as those charac- teristic of ocean ridges and continental rifts, the range of penetration of the mantle solidus is large; hence the LVZ should be relatively wide (lines A and B, Fig. 4.5). The gentle geotherms in continental platforms, which show a narrow range of intersection with the hydrated mantle solidus, produce a thin or poorly defined LVZ (line C, Fig. 4.5). Beneath Archean shields, geotherms do not intersect the mantle solidus; hence there is no LVZ (line D, Fig. 4.5). The Transition Zone The 410-km Discontinuity The transition zone is that part of the upper mantle in which two major seismic disconti- nuities occur: one at 410 km and the other at 660 km (Fig. 4.1). High-pressure experi- mental studies document the breakdown of Mg-rich olivine to a high-pressure phase known as wadsleyite (beta phase) around 14 GPa, which is equivalent to a 410-km depth in the Earth (Fig. 4.15). There is no change in chemical composition accompanying this phase change or other phase changes described in this section. Mantle olivine (Fo 90 ) transforms to wadsleyite at pressures less than 300 MPa at appropriate temperatures for the 410-km discontinuity (~1000° C) (Ita and Stixrude, 1992; Helffrich and Wood, 2001). This pressure range agrees with the less than 10-km width of the 410-km discontinuity deduced from seismic data (Vidale et al., 1995). In some places, the discontinuity is The Transition Zone 133 broader than normal (20–25 km), a feature that may reflect water incorporated into the wadsleyite crystal structure (van der Meijde et al., 2003). If olivine composes 40 to 60% of the rock, as it does in garnet lherzolite, the olivine–wadsleyite phase change may account for the approximately 6% increase in density observed at this discontinuity (Table 4.2). Measurements of elastic moduli of olivine at high pressures suggest that 40% olivine explains the velocity contrast better than 60% olivine (Duffy et al., 1995). Because garnet lherzolites typically have 50 to 60% olivine, modal olivine may decrease with depth in the upper mantle to meet this constraint. Experimental data indicate that wadsleyite should transform to a more densely packed spinel-structured phase (gamma phase) at equivalent burial depths of 500 to 550 km. This mineral, hereafter called spinel, has the same composition as Mg-rich olivine but the crystallographic structure of spinel. The small density change (~2%) asso- ciated with this transition, however, does not generally produce a resolvable seismic discontinuity. High-pressure experimental data also indicate that at depths from 350 to 450 km, both clino- and orthopyroxene are transformed into a garnet-structured mineral known as majorite garnet, involving a density increase of about 6% (Christensen, 1995). This transition has been petrographically observed as pyroxene exsolution laminae in garnet in mantle xenoliths derived from the Archean lithosphere at depths from 300 to 400 km (Haggerty and Sautter, 1990). It is probable that an increase in velocity gradient some- times observed from 350 km to the 410-km discontinuity is caused by these pyroxene transformations. At a slightly higher temperature, Ca-garnet begins to transform to Ca-perovskite (a mineral with Ca-garnet composition but perovskite structure). All of the preceding phase changes have positive slopes in P–T space; thus the reactions are exothermic (Table 4.2). 134 The Mantle 1000 1200 1400 1600 1800 2000 2200 2400 Temperature (°C) 30 25 20 15 10 5 Pressure (GPa) 200 400 600 800 Depth (km) Perovskite + Magnesiowustite Spinel Wadsleyite Olivine Mantle Solidus Mantle Adiabat 410-km Discontinuity 660-km Discontinuity Figure 4.15 Summary of phase relations for Mg 2 SiO 4 in the mantle from high- pressure and high- temperature experimental studies. The dashed line is the mantle adiabat. Modified from Christensen (1995). The 660-km Discontinuity One of the most important questions related to the style of mantle convection in the Earth is the nature of the 660-km discontinuity (Fig. 4.1). If descending slabs cannot readily penetrate this boundary or if the boundary represents a compositional change, two-layer mantle convection is favored, with the 660-km discontinuity representing the base of the upper layer. Large increases in both seismic-wave velocity (5–7%) and density (5%) occur at this boundary. High-frequency seismic waves reflected at the boundary suggest that it has a width of only about 5 km but has up to 20 km of relief over hundreds to thou- sands of kilometers (Wood, 1995). The Transition Zone 135 Table 4.2 Summary of Mantle Mineral Assemblages for Average Garnet Lherzolite from High-Pressure Studies Mineral Slope of Assemblage Density Reaction Depth (km) (minerals in vol. %) Contrast (%) (MPa/°C) <410 Olivine 58 Opx 11 Cpx 18 Garnet 13 350–450 Opx-cpx 6 +1.5 → Majorite garnet 410-km discontinuity 410 Olivine (α phase) 6 +5.5 → Wadsleyite (β phase) 410–550 Wadsleyite 58 Majorite garnet 30 Cpx 9 Opx 3 500–550 Wadsleyite 2 +3.0 → Spinel (γ phase) 550–660 Spinel 58 Majorite Garnet 37 Ca-perovskite 5 Ca-garnet → Ca-perovskite 660-km 660 Spinel → discontinuity Perovskite 5 –2.5 to –2.8 + magnesiowustite 650–680 Majorite garnet +1.5 to +2.5 → perovskite 680–2900 Perovskite 77 Magnesiowustite 15 Ca-perovskite 8 Silica (?) Data from Ita and Stixrude (1992), Christensen (1995), Mambole and Fleitout (2002), and Hirose (2002). As with the 410-km discontinuity, it appears that a phase change in Mg 2 SiO 4 is responsible for the 660-km discontinuity (Christensen, 1995; Helffrich and Wood, 2001). High-pressure experimental results indicate that spinel transforms to a mixture of per- ovskite and magnesiowustite at a pressure of about 23 GPa and can account for both the seismic velocity and the density increases at this boundary if the rock contains 50 to 60% spinel: (Mg,Fe) 2 SiO 4 → (Mg,Fe)SiO 3 + (Mg,Fe)O spinel → perovskite + magnesiowustite Mg-perovskite and magnesiowustite are extremely high-density minerals and appear to comprise most of the lower mantle. High-pressure experimental studies show that small amounts of water may be carried as deep as the 660-km discontinuity in hydrous phases stable to 23 GPa (Ohtani et al., 1995). Unlike the shallower phase transitions, the spinel–perovskite transition has a negative slope in P–T space (–2.5 to –2.8 MPa/°C) (Fig. 4.15; Table 4.2); thus the reaction is endothermic and may impede slabs from sinking into the deep mantle or impede plumes from rising into the upper mantle. The latent heat associated with phase transitions in descending slabs and rising plumes can deflect phase transitions to shallower depths for exothermic (positive P–T slope) reactions and to greater depths for endothermic (nega- tive P–T slope) reactions (Liu, 1994). For a descending slab in an exothermic case, such as the olivine–wadsleyite transition, the elevated region of the denser phase exerts a strong downward pull on the slab or an upward pull on a plume, helping drive convection. In contrast, for an endothermic reaction, such as the spinel–perovskite transition, the low- density phase is depressed, enhancing a slab’s buoyancy and resisting further sinking of the slab. This same reaction may retard a rising plume. Around the same depth (650–680 km), majorite garnet transforms to perovskite (Table 4.2), but unlike the spinel transition, the garnet transition is gradual and does not produce a seismic discontinuity. This transition is sensitive to temperature and the Al content of the system. Unlike the spinel–perovskite reaction, the garnet–perovskite reac- tion has a positive slope (1.5 to 2.5 MPa/°C) (Hirose, 2002; Mambole and Fleitout, 2002). Computer models by Davies (1995) suggest that stiff slabs can penetrate the bound- ary more readily than plume heads and that plume tails are the least able to penetrate it. Some investigators have suggested that slabs may locally accumulate at the 660-km discontinuity, culminating in occasional “avalanches” of slabs into the lower mantle. The fate of the oceanic crust in descending plates may be different from that of the sub- oceanic lithosphere because of their different compositions. Irifune and Ringwood (1993) suggested that the 660-km discontinuity may be a density “filter,” which causes the crust to separate from the mantle in descending slabs. At depths less than 720 km, basaltic crust has a greater density than surrounding mantle, which could cause separation of the two components when they intersect the discontinuity. At a depth of about 700 km, however, the crust becomes less dense than surrounding mantle because of the majorite–perovskite 136 The Mantle [...]... Modern Oceanic Basalts Ε Nd (T) 2.7-Ga Mafic Rocks +4 0 4 0.05 0.10 0.15 0.20 147 Sm/114Nd 0.25 0.30 Figure 4. 25 Distribution of late Archean and modern basalts on an ΕNd (T) versus 147 Sm/ 144 Nd graph T (age) = 0 Ma for modern basalts and 2.7 Ga for Archean rocks Modified from Blichert-Toft and Albarede (19 94) 1 54 The Mantle convecting mantle to lose their geochemical and isotopic signatures Many of these... mantle The Mantle 16 DM 0.5132 MORB ICELAND EASTER CANARIES ASCENSION GUADALUPE 12 PM ST HELENA 0 TAHITI BOUVET 4 WALVIS R 0.51 24 4 Nd HIMU ε SAMOA 0.5128 AZORES I 8 HA W AI Figure 4. 22 Nd and Pb isotope distributions in MORB and oceanic island basalts Modified from Zindler and Hart (1986) DM, depleted mantle; EM1 and EM2, enriched mantle components; HIMU, high U/Pb ratio; PM, primitive mantle 143 Nd/ 144 Nd... involving extraction of basaltic magmas Depleted mantle is known to underlie ocean ridges and probably extends beneath ocean basins, although it is not the source of oceanic-island magmas The depleted isotopic character (low 87Sr/86Sr, 206Pb/204Pb, and high 143 Nd/ 144 Nd) and low LIL element contents of NMORB require the existence in the Earth of a widespread depleted mantle reservoir Rare gas isotopic compositions... dominant force for continued subduction Mantle Plumes Introduction A mantle plume is a buoyant mass of material in the mantle that, because of its buoyancy, rises The existence of mantle plumes in the Earth was first suggested by Wilson (1963) as an explanation of oceanic-island chains, such as the Hawaiian-Emperor chain, which change progressively in age along the chain Wilson proposed that as a lithospheric... the end members has to be Paleoproterozoic Identifying Mantle Components Summary At least four and perhaps as many as six isotopic end members may exist in the mantle from results available from oceanic basalts (Hart, 1988; Hart et al., 1992; Helffrich and Mantle Geochemical Components 147 Wood, 2001) (Figs 4. 21 and 4. 22) These are depleted mantle, the source of normal midocean-ridge basalts (NMORB);... whereas midocean-ridge basalts (MORB) generally has R/RA values of 7 to 9 (Hanan and Graham, 1996) In some Pacific MORB, R/RA values increase with increasing 206Pb/204Pb ratios, whereas in the Atlantic, the opposite trend is observed Also, in the Atlantic and Indian Oceans, helium isotope ratios tend to be higher in basalts with EM1 characteristics These high ratios, which apply to both MORB and OIB,... The Mantle Figure 4. 24 Threedimensional plot of Sr, Nd, and Pb isotopic ratios from intraplate ocean basalts The tetrahedron is defined by the mantle end members DM, depleted mantle; EM1 and EM2, enriched mantle components; FOZO, focal zone; HIMU, high U/Pb ratio; NMORB, normal midocean-ridge basalts Modified from Hauri et al (19 94) EM1 EM2 Samoa 87/86 Sr Pitcairn Society Islands 143 / 144 Nd 206/2 04 Pb... Sm/Nd ratios provide a means to compare the Archean mantle with the modern mantle by analyzing for these quantities Mantle Geochemical Components 153 in greenstone basalts The mantle array on an ΕNd- 147 Sm/ 144 Nd plot varies significantly from the modern arrays (Fig 4. 25) For a similar overall range in Sm/Nd ratio, the Archean array varies by only 1 epsilon (Ε) unit compared with at least 8Ε units in modern... ratios and moderate 206Pb/204Pb ratios Both have low 143 Nd/ 144 Nd ratios Among the numerous candidates proposed for EM1 are old oceanic mantle lithosphere Mantle Geochemical Components (±sediments) that has been recycled back into the mantle (Hart et al., 1992; Hauri et al., 19 94) and metasomatized lower mantle (Kamber and Collerson, 1999) Also, studies of Hf isotope distributions in Hawaiian basalts... Dupal Anomaly With some important exceptions, most oceanic islands showing enriched mantle components occur in the Southern Hemisphere (Hart, 1988) Also, NMORB in the South Atlantic and Indian Oceans exhibit definite contributions from enriched mantle sources, unlike NMORB from other ocean-ridge systems This band of mantle enrichment in the Southern Hemisphere is known as the Dupal anomaly (Fig 4. 23) . core as it cools. 144 The Mantle Deccan Rajmahal 65 57 49 82 77 58 38 35 8 2 Reunion 62 43 27 Kerguelen Figure 4. 19 Computer- generated hotspot tracks for the Indian Ocean and South Atlantic basins,. high-pressure phase known as wadsleyite (beta phase) around 14 GPa, which is equivalent to a 41 0-km depth in the Earth (Fig. 4. 15). There is no change in chemical composition accompanying this phase change. buoyant mass of material in the mantle that, because of its buoy- ancy, rises. The existence of mantle plumes in the Earth was first suggested by Wilson (1963) as an explanation of oceanic-island