Igneous rock associations 5. oceanic island volcanism II: mantle processes
John D. Greenough
Oceanic island basalts (OIBs) have been central to understanding evolution of the Earth and mantle because their isolated positions in ocean basins limit the potential for magma contamination by continental crust. Melting processes (e.g., percentage melting) affect OIB chemistry but isotopic and trace-element ratios provide information on mantle-source compositions. They indicate that OIB mantle sources represent mixtures between mid-ocean ridge basalt (MORB) mantle and four other mantle components: EM1 (enriched mantle 1), EM2, HIMU (High U/Pb = Hi [mu]) and FOZO (FOcal ZOne). Mass-balance and noble-gas arguments indicate that most of the mantle is depleted but He and Ne isotopes, and convergence of Sr-Nd-Pb isotopic arrays suggest that FOZO is a somewhat primitive (unmelted) component common to all oceanic basalt sources. The other components contain “materials” such as basaltic ocean floor (HIMU), pelagic sediments (EM1), oceanic plateaus (EM1), subcontinental lithosphere (EM1, EM2), terrigenous sediments or subducted continental crust (EM2), which have been recycled by subduction processes, and mixed back into the depleted mantle. How these components cycle through the mantle is debated but heterogeneities occur on all length-scales. One school argues that oceanic islands develop above mantle plume convection cells that deliver recycled components and FOZO (lower mantle?) for mixing with depleted upper mantle. Others contend that propagating cracks in the lithosphere create oceanic islands, that plumes do not exist, that the upper and lower mantle are isolated and depleted, and that MORB and OIB form from the same upper-mantle reservoir. Small-scale melting allows OIB to sample local, low-melting-point heterogeneities that are averaged-out by the large-scale melting that forms MORB. These radically different views of mantle structure and composition indicate that OIB will continue to be a focal point in studies of Earth’s evolution.
L’etude des basaltes d’iles oceaniques (BIOs, ou OIBs en anglais) s’est avere essentielle pour la comprehension de l’evolution de la Terre et de son manteau, et cela, de par l’isolement de ces iles dans les bassins oceaniques, ce qui limite les possibilites de contamination par des materiaux de la croute continentale. Les mecanismes de fusion (le pourcentage de fusion par ex.) delimitent la composition chimique des BIOs, mais les ratios isotopiques et des elements traces permettent d’obtenir des indications sur la composition des sources mantelliques. Ils indiquent que les sources mantelliques des BIOs sont des melanges de basaltes de dorsales oceaniques (BDOs ou MORBs en anglais) de quatre autres composantes du manteau, soit des EM1 (enriched mantle), EM2, HIMU (ratio eleve de U/Pb = Hi [mu]), et FOZO (FOcal ZOne). Les etudes des bilans massiques et des gaz nobles indiquent que la plus grande partie du manteau a subit un appauvrissement, mais les isotopes He et Ne, ainsi que la convergence des ensembles isotopiques Sr-Nd-Pb portent a penser que la composante FOZO serait de composition a peu pros primitive (n’aurait pas subit de fusion) qui serait commune toutes les sources de basaltes oceaniques. Les autres composantes renferment des “materiaux” issus de plancher oceanique basaltique (HIMU), de sediments pelagiques (EM1), de plateaux oceaniques (EM1), de lithosphere sous-continentale (EM1 et EM2), de sediments terrigenes ou de croutes continentales enfouies (EM2) et qui ont ete recycles par des meanismes de subduction et reinjecte dans les materiaux appauvris du manteau. La facon dont ces composantes sont recyclees dans le manteau fait l’objet de discussions serrees et on observe la presence d’heterogeneite a toute echelle. Une des ecoles de pensee soutient que les iles oceaniques se forment au-dessus de cellules de convection de panaches mantelliques qui apportent des composantes recyclees et de la FOZO (manteau inferieur?) et les melangent avec les couches superieures appauvries du manteau. D’autres croient plutot que ce sont des fissures de la croute qui permettent la formation des iles oceaniques, qu’il n’y pas de panaches, que les couches inferieures et superieures du manteau sont isolees et appauvries et que les BIO et les BDO sont formes a partirr des materiaux des meme couches superieures. Les BIO seraient le reflet de fusions d’heterogeneites locales a faibles temperatures de fusion, alors que les BDO seraient le resultat de fusions a grande echelle expliquant une composition correspondant a la moyenne de toutes les heterogeneites. L’existence de points de vue si radicalement opposes sur la structure et la composition du manteau demontrent que les BIOs seront encore l’objet d’&udes sur l’evolution de la Terre.
Oceanic islands represent a small proportion of the Earth’s surface yet their volcanic rocks are amongst the most studied; this is because they occur far from continental crust that might contaminate rising magma, and they are unaffected by orogenic, metamorphic and tectonic processes (Basaltic Volcanism Study Project = BVSP, 1981, p. 161). Thus, oceanic island basalts (OIBs) provide important chemical evidence regarding mantle composition, magma formation processes and magma evolution. Perhaps reflecting the availability and accuracy of major element, trace element and isotopic analyses, the emphasis of studies on OIB has shifted over 40 years from the effects of differentiation, to the impact of melting processes on magmas, to studying what magma chemistry can tell us about mantle-source compositions. Today, based mostly on isotopic data from OIB and mid-ocean ridge basalt (MORB), five types of mantle-source regions, or components, are recognized (Zindler and Hart, 1986; Hofmann, 2003). These appear to reflect variable melt-extraction and subduction-related recycling of materials back into the mantle over Earth history, although the specifics of how each “component” formed are debated.
Models for the distribution, scale and melting behaviour of these components are even more contentious and lead to very different hypotheses for chemical structure and convection patterns in the mantle. One school sees some amount of plume-related(?) exchange between primitive (relatively unmelted) lower mantle and “depleted” (previously melted) upper mantle (e.g., Allegre, 2002; Hofmann, 2003). The other sees the entire mantle as convecting and depleted; the lower mantle is separated from the upper mantle that is locally enriched in subduction-recycled components (Anderson, 1999; Hamilton, 2003). Mantle plumes may not exist (Anderson, 2000). Thus a healthy, and at present unresolved, debate has emerged that reflects fundamental views of how planet Earth works, and OIBs are front and centre in the discussion.
This paper looks at how OIB compositions (Table 1) reflect source region compositions and mantle melting processes (e.g., pressure effects, impact of fluids, melting percentage). Hypotheses for the origin of mantle heterogeneity (the mantle components), evidence for the survival of primitive mantie, and models for chemical structure in the Earth are then reviewed. This manuscript complements Greenough et al., (2005) that covers the mineralogy, petrology and differentiation of oceanic island basalts.
Controls on Major-Element Compositions
A recurring theme in petrology is whether a basalt has experienced differentiation or is “primary” and derived directly from the mantle. Although debated, primary magmas are thought to have Mg# values of about 0.72 as a result of equilibration with [Fo.sub.92] olivine that is found in most mantle lherzolite xenoliths (Roeder and Emslie, 1970). At least a few aphyric basalts on most oceanic islands have Mg# values around 0.72, which is consistent with this hypothesis. Olivine partitioning data indicate that primary basaltic magmas should have Ni/MgO ratios (ppm/wt.% oxide) between 22 and 50 with Ni = 200 to 1000 ppm (Basaltic Volcanism Study Project, 1981, p. 424). A review of the Ni, Cr, Co and Sc contents in approximately 200 “primary” basalts (Mg# values of 0.72) from French Polynesia, using regression analysis, gives concentration ranges of 350-550, 650-930, 70-80 and 20-40 ppm, respectively.
Figure 1a illustrates the pressure effects on the beginning-of-melting point in a synthetic ternary system modelling lherzolite. Points [I.sub.0], [I.sub.1], [I.sub.2][, [I.sub.3] show that magmas become increasingly silica-undersaturated, and Ne normative, with increasing pressure (1 atmosphere, 10, 20, and 30 kbar, respectively). Experiments on an anhydrous spinel lherzolite indicate that magmas change from Ne normative (alkaline) to Hy normative (tholeiitic) as melting percentages increase, and the Ne component increases as pressure increases (Fig. 1a). As Haase (1996) emphasized, the silica content of magmas is very sensitive to pressure. Melting experiments also show that C[O.sub.2] leads to undersaturated magmas, whereas [H.sub.2]0 produces more silica-rich magmas (Fig. 1b). These generalizations suggest that relative to tholeiites, oceanic island alkaline magmas reflect small percentages of melting of C[O.sub.2]-rich lherzolite at high pressures.
[FIGURE 1 OMITTED]
Metasomatism is the process whereby migration of C[O.sub.2]- and [H.sub.2]0-rich fluids (or melts), through the mantle, leads to local incompatible element enrichment prior to magma formation. Alkaline OIBs commonly contain too much Rb for their observed [sup.87]Sr/[sup.86]Sr ratios. This suggests that metasomatic fluids or melts carry large ion lithophile elements to their sources a short time prior to melt extraction. Metasomatism may also occur long before magma genesis. Many experiments suggest that ancient subduction-related metasomatism affected OIB mantle sources long before island magmatism. Clearly, metasomatism coincident with magmatism is important, but it cannot yield the large variations in daughter radiogenic isotopes observed among islands because they require millions or billions of years to develop from parent isotopes having long half-lives. However, Kamber and Collerson (1999) proposed that OIB Pb isotope variations partly reflect metasomatism by melts derived from the lower mantle only 150 million years ago. Halliday et al. (1995) argued that melt metasomatism during lithosphere formation at the ocean ridges produces high U/Pb ratios (high [mu] = high “mu”) that with aging ([~10.sub.8] years) can explain the high [sup.206]Pb/[sup.204]Pb ratios of HIMU oceanic islands (see Isotopes section below).
Percentage of Melting
The percentage of melting affects trace-element concentrations as illustrated by chondrite-normalized rare-earth-element (REE) diagrams (Fig. 2; Table 1). Highly alkaline magmas (e.g., Aitutaki; low % melting) show steep slopes and tholeiites low slopes (e.g., Iceland; high % melting). La (largest REE cation) is substantially more incompatible in mantle minerals than Lu (smallest REE). Nearly all La enters the first drops of magma yielding high concentrations in alkaline magmas. Concentrations fall rapidly as the percentage of melting increases. Lu is less incompatible, so magma concentrations are low and change slowly, causing La/Lu ratios to decrease as melting increases. Ratios reflecting relative percentages of melting are calculated from elements [greater than or equal to] 10 elements apart in the Sun and McDonough (1989) incompatibility list (Table 2, notes). Thus, Nb/Y (23 elements apart) is an excellent chemical separator of alkaline (ratios > 1) and tholeiitic (< 1) magmas and provides a continuous measure of "alkalinity". The correlation in Figure 3 shows that as the depth/pressure of melting increases (reflected by SI[O.sub.2]; Haase, 1996), the percentage of melting (monitored by Nb/Y) decreases. The depth and pressure of melting correlate with the age (Haase, 1996) and thickness (Watson and McKenzie, 1991) of the lithosphere. Thus, strongly alkaline magmas form where the lithosphere is old, cold and thick, and this causes melting to occur at the greatest depths, where the percentage of melting is apparently low.
[FIGURES 2-3 OMITTED]
The percentage of melting correlates with, and appears responsible for, excesses in the U-series ratios ([sup.230]Th/[sup.238]U), ([sup.226]Ra/[sup.230]Th), and ([sup.231]Pa/[sup.235]U) (brackets = activity ratios), which increase between tholeiites (high % melting) and alkali basalts (Bourdon and Sims, 2003; see brief introduction to U-series dating in Oceanic Island Volcanism I: Mineralogy and Petrology, Greenough et al., 2005). Apparently, magma is enriched in more-incompatible daughter isotopes ([sup.230]Th, [sup.226]Ra, [sup.231pa) during melting and the process of melt formation and extraction occurs rapidly enough that daughter excesses are not erased (equilibrium is not re-established). Plume models for Hawaiian volcanism indicate that excesses are negatively correlated with mantle upwelling velocity; both upwelling velocity and melting percentage (faster = higher % melting) decrease away from the axis of the plume.
Residual Minor Phases
The phases that are left behind after magma extraction from the mantle can be inferred from trace-element data. For example, most “undifferentiated” OIBs exhibit low chondrite-normalized heavy REE (Yb, Lu) concentrations (6-9 times chondrites) that change little between alkaline (e.g., Kauai) and tholeiitic (Kilauea) magmas (Fig. 2). A persistent phase creates a bulk crystal/liquid partitioning coefficient for these elements close to 1. Garnet shows high partition coefficients for the heavy REE (e.g., Rollinson, 1993; p. 108). Apparently most OIBs form at pressures great enough for garnet stability, and the mineral forms a residual phase over a wide range of melting percentages. Excesses in the activity of ([sup.230]Th) over (238]U) are commonly ascribed to the presence of residual garnet in oceanic island sources (Bourdon and Sims, 2003).
Highly alkaline rocks (La > 50 ppm) show lower element/La ratios for Rb, K, Sr, Th, Zr, Nb and Ta indicating residual minor phases in their sources (Sun and McDonough, 1989). In single-locality suites of primitive (Mg# = 0.67-0.75) basalts, rocks with the highest Sr show intermediate Rb concentrations, those with highest Rb have intermediate Sr, and the lowest Rb and Sr occur together (Greenough, 1988). This may reflect retention of Rb in phlogopite at low percentages of melting (highest Sr rocks). As phlogopite melts, magma Rb concentrations rise, and they peak just as it is entirely consumed. With more melting, both Rb and Sr decrease because of dilution. A similar “arrow-shaped” pattern for Ti versus Sr suggests a residual titanate mineral or Ti-phlogopite at low-percentages of melting. Highly alkaline magmas also have somewhat low Nb, Ta, Zr and Hf (relative to Ba, La) supporting residual futile or ilmenite although the Sr-Ti relationship favours phlogopite (Greenough, 1988). Halliday et al. (1995) argued that OIB sources with high U/Pb ratios (HIMU) result from migration of small percentage melts that equilibrated with minor amphibole, sulfide and phlogopite. In summary, minor minerals probably play a major role in controlling the chemistry of some OIB source region types and their alkaline magmas.
ISOTOPES AND SOURCE REGION COMPOSITIONS
Lithophile and Siderophile Isotopes
Since the pioneering work of Gast, Tilton and Hedge in the 1960s, isotopic ratios have been the cornerstone of mantle composition studies. This is because, unlike element concentrations, these ratios are generally thought to be unaffected by melting percentages (but see Os discussion below). Variations in [sup.87]Sr/[sup.86]Sr, [sup.143]Nd/[sup.144]Nd, [sup.206]Pb/[sup.204]Pb, [sup.207]Pb/[sup.204]Pb, [sup.208]Pb/[sup.204]Pb, [sup.187]Os/[sup.188]Os and [sup.176]Hf/[sup.177]Hf reflect long-lived fractionation of parent isotopes ([sup.87]Rb/[sup.86]Sr, [sup.147]Sm/[sup.144]Nd, [sup.238]U/[sup.204]Pb, [sup.235]U/[sup.204]Pb, [sup.232]Th/[sup.204]Pb, [sup.187]Re/[sup.188]Os, [sup.176]Lu/[sup.177]Hf, respectively) from fore-listed daughter isotopes. For example, Rb is more incompatible in mantle minerals than Sr during melting. Magma removal to form oceanic crust results in mantle residue that is depleted in Rb relative to Sr, and the crust is Rb-enriched. Because [sup.87]Rb decays to [sup.87]Sr with time, the crust develops a high [sup.87]Sr/[sup.86]Sr ratio relative to unmelted mantle, but this ratio in the residue increases slowly (is low today) because of being depleted in [sup.87]Rb (see Faure, 2001).
Isotopes are used to identify four end-member types of mantle formed by previous melt extraction and recycling of materials (e.g., crust) by subduction (Zindler and Hart, 1986; Hofmann, 2003). The component (type) sampled by MORB has low incompatible element concentrations, low [sup.87]Sr/[sup.86]Sr, and high [sup.143]Nd/[sup.144]Nd, reflecting previous (ancient) melt extraction (Fig. 4). This depleted MORB mantle (DMM) is not well represented by OIB (see Iceland debate; Hanan et al., 2000; Fitton et al., 2003), but less-extreme depleted mantle (DM) underlies some islands (e.g., Easter). High Pb isotopic ratios in OIB (e.g., St. Helena, Fig. 4) reflect high mantle U/Pb and Th/Pb ratios in the HIMU component (= high [mu] = high [sup.238]U/[sup.204]Pb). The enriched mantle (EM) components have high incompatible element concentrations relative to hypothetical primitive (unmelted) mantle. The EM1 islands show the lowest [sup.143]Nd/[sup.144]Nd and [sup.176]Hf/[sup.177]Hf ratios and high [sup.87]Sr/[sup.86]Sr (e.g., Gough, Fig. 4) and EM2 islands have the highest [sup.87]Sr/[sup.86]Sr ratios (e.g., Tumila). Most OIBs have intermediate compositions explained by convective mixing between components (Zindier and Hart, 1986). Three-dimensional graphs of isotopic ratios (e.g., Pb, Sr, Nd) show that islands plot close to a mixing plane, determined by regression analysis (Fig. 5; Zindler and Hart, 1986). These simple isotopic relationships are easier to explain if the components formed at particular times by specific processes. Thus they may reflect major events in Earth’s evolution.
[FIGURES 4-5 OMITTED]
The Nd-Sr isotopic array (Fig. 4a) might be explained by mixing between depleted MORB mantle (depleted to form continental crust) and primordial (unmelted) mantle but EM1 and EM2 islands plot off the array and Pb isotope variations are unaccounted for (e.g., HIMU; Fig. 5a). Thus this simple model advocating unidirectional movement of incompatible elements into the crust is not tenable.
A plot of [sup.206]Pb/[sup.204]Pb versus [sup.204]Pb/[sup.204]Pb sheds light on mantle evolution (Fig. 4c). The ratio of [sup.235]U/[sup.238]U has changed throughout Earth history because the two isotopes have different decay rates, but at any one time the ratio was everywhere the same. This allows the calculation of model ages. Any mantle fractionation process that creates variable U/Pb ratios and a single set of [sup.206]Pb/[sup.204]Pb and [sup.207]Pb/[sup.204]Pb ratios will, with time, evolve sets of Pb isotopic ratios that define a linear array dating the disturbance (see Faure, 2001). The Geochron dates the Earth (Fig.4c). Most OIBs and MORBs plot to the right of the Geochron (Fig. 4c) and this is referred to as the Pb paradox. It implies that U was added or Pb removed from the mantle sources of all oceanic basalts in the ancient past. MORB basalts are mildly enriched in U (and Th), relative to Pb during mantle melting and the residue should have a low U/Pb ratio. However, MORBs have high [sup.206]Pb/[sup.204]Pb and [sup.207]Pb/[sup.204]Pb ratios, requiring a long-term high U/Pb ratio. To explain the Pb paradox, it has been suggested that fluids efficiently and permanently delivered Pb from oceanic lithosphere to the continental crust during ancient subduction events, whereas U was more effectively recycled back into the mantle (Hofmann, 1997; 2003).
Given the importance of subduction, it is reasonable to expect that OIB magma compositions might hold evidence for recycling of oceanic lithosphere. HIMU basalts are commonly ascribed to melting of mantle containing recycled oceanic crust (Zindler and Hart, 1986; Hofmann, 1997). Mixing between this reservoir and depleted mantle can explain first-order correlations between Pb isotopes (Fig. 4c, 4d). If so, the slope on the line (model Pb age) has no age significance. However, HIMU sources with the highest [sup.206]Pb/[sup.204]Pb ratios could be 2 billion years old (Hofmann, 1997). Kamber and Collerson (1999) suggested that mixing between depleted MORB mantle and melts from EM1-type mantle can produce HIMU-type sources. HIMU Pb isotopic ratios are produced after only 150 Ma. Lithium isotopes ([sup.7]Li/[sup.6]Li reported as [delta][sup.7]Li), appear to be primarily fractionated by low- (surface) and medium-temperature (crust and lithosphere) processes. There is a significant range in OIB [delta][sup.7]Li ratios and this supports widespread distribution of recycled materials in the mantle (see review, Elliott et al., 2004). However, HIMU sources appear enriched in the heavy [sup.7]Li isotope. This is the opposite of what is predicted for mantle with basaltic ocean floor that has been “processed” during subduction. Thus, Li may place new constraints on HIMU formation and is potentially more useful than other stable isotopes for monitoring recycling processes.
EM1 and EM2 cannot be explained by mixing between depleted mantle and HIMU components (Fig. 4a, 4b; Hofmann, 1997). The Sr, Nd, Hf and Pb isotopes of EM1-sourced OIBs resemble lower continental crust but Pb/Zr ratios are low. A “primitive” mantle explanation for the enriched nature of EM1 is inconsistent with low [sup.3]He/[sup.4]He ratios that imply previous and extensive outgassing and melting. Gasparini et al. (2000) proposed that EM1 basalts form by melting subducted (recycled) oceanic plateaux containing enriched basalts produced by an ancient plume head. Alternatively, EM1 sources may bear an imprint from subducted pelagic sediment subduction (Weaver, 1991) or represent subcontinental lithospheric mantle scraped from beneath continents (e.g., Milner and le Roex, 1996). Enriched mantle sources are more common in a globe-encircling band possibly related to the detachment of African and South American subcontinental lithosphere during the assembly and breakup of Gondwana. EM1 may be delaminated Archean subcontinental lithosphere, given that the most extreme isotopic examples come from some Cenozoic, cratonic mafic magmas (Greenough and Kyser, 2003). Diamonds are most commonly associated with Archean subcontinental lithosphere but nanodiamonds were recently reported from mantle xenoliths in EM1like Hawaiian basalts (Wirth and Rocholl, 2003). Tatsumi (2000) showed that EM1 Sr-Nd-Pb isotopic compositions are consistent with melting Archean pyroxenite from the subcontinental lithosphere, and the [sup.206]Pb/[sup.204]Pb versus [sup.207]Pb/[sup.204]Pb model age from EM1 sources supports an Archean age (Fig. 4c).
The high [sup.207]Pb/[sup.204]Pb and [sup.87]Sr/[sup.86]Sr of EM2 sources (Figs.4 and 5) suggest that they contain recycled (subducted) terrigenous sediments or pelagic sediments having a continental signature (Weaver, 1991). Elevated oxygen isotopic ratios ([delta] [sup.18]0), reflecting near-surface, low-temperature, fractionation processes, support the subducted sediment model (Eiler et al., 1997). EM2 may also represent delaminated, post-Archean, subcontinental lithosphere (Greenough and Kyser, 2003).
The [sup.187]Re – [sup.187]Os system provides insights into the origin and distribution of mantle components. Compared to lithophile isotopic systems (e.g., Rb-Sr) that behave incompatibly during mantle melting, Os is compatible and Re is incompatible; their chalcophyic/siderophilic geochemistry suggests they are mostly stored in the core (Hauri, 2002). Basalts are enriched in Re over Os, and they develop high [sup.187]Os/[sup.188]Os ratios over time, which should identify recycled ocean crust in the mantle. Indeed, Os isotopes suggest that most OIB sources contain > 10% recycled basalt and HIMU sources are even higher (i.e., high [sup.187]Os/[sup.188][Os]; Fig. 4e) (Hauri, 2002).
The geochemistry of Os is not well understood. Mantle minerals holding Os include sulphides, chromite and metal alloys; their distribution and melting behaviour are not well known. Silicate phases hosting Re may melt before Os-rich phases, resulting in magmas with high [sup.187][Os]/[sup.188]Os ratios (from Re decay) out of equilibrium with the bulk source composition; thus magmas may not reflect average mantle Os compositions (Becker, 2000; Hofmann, 2003). Walker et al. (1999) cautioned against using Os to estimate the percentage of recycled ocean floor in OIB sources. Although examples of correlations between Os and lithophile isotopes exist (Shiano et al., 2001), they are “disconnected” in most OIBs. The EM1 basalts, which show both high and low Os ratios, illustrate this well (Fig. 4e). An alternative to disequilibrium melting is that Os isotopes reflect the proportion of peridotite to recycled oceanic crust, whereas other isotopes are controlled by the amount of recycled sediment (Eisele et al., 2002). Possibly OIB Os isotopes, along with platinum-group-element concentrations are controlled by mantle-outer core (high [sup.187]Os/[sup.188]Os? ratios) interaction, or the distribution of late-bombardment meteorite material (high Pt, Os etc.) in the mantle (Fryer and Greenough, 1992; Walker et al., 1999). There is much yet to learn about and from Os isotopes.
Hafnium and Nd isotopes support the presence of recycled ocean floor in the mantle. In a plot of [sup.143]Nd/[sup.144]Nd versus [sup.176]Hf/[sup.177]Hf, a regression line passed through depleted mantle, OIB and continental crust does not pass through bulk silicate Earth. This suggests there is a “hidden” mantle reservoir with prolonged, high Sm/Nd and low Lu/Hf (time-integrated high [sup.143]Nd/[sup.144]Nd and low [sup.176]Hf/[sup.177]Hf) similar to that predicted for extreme HIMU sources (Bizzarro et al., 2002; Pearson and Nowell, 2002). Modern MORB contains the requisite trace-element ratios for such a hidden reservoir. The Sm/Nd ratio is high reflecting previous melt extraction whereas the incompatibility of Hf, relative to Lu during melting, results in low Lu/Hf ratios. Thus, there may be a “hidden”, extreme, HIMU-like mantle reservoir containing basaltic crust that represents 1 to 15% of the silicate Earth.
Noble Gas Isotopes
Noble gas isotopes (e.g., [sup.3]He/[sup.4]He, [sup.20]Ne/[sup.22]Ne, [sup.21]Ne/[sup.22]Ne, [sup.40]Ar/[sup.36]Ar, [sup.129]Xe/[sup.130]Xe) are central to some heated debates on the composition and structure of the mantle (Hilton and Porcelli, 2003; McDougall and Honda, 1998). Some isotopes are radiogenic ([sup.4]He and [sup.21]Ne from U and Th decay, [sup.40]Ar from [sup.40]K, [sup.129]Xe from [sup.129]I, and exceptionally high ratios in mantle reservoirs (e.g., MORB [sup.40]Ar/[sup.36]Ar > 20,000 compared to 295.5 in atmosphere) reflect major enrichment in the radiogenic, relative to the primordial (unradioactive and unradiogenic) isotopes. Apparently the Earth underwent massive degassing so that mantle gases are now dominated by the radiogenic isotopes. High [sup.129]Xe/[sup.130]Xe ratios in MORBs compared to the atmosphere indicate that the catastrophic degassing occurred during the first 100 Ma of Earth history (McDougall and Honda, 1998) because [sup.129]Xe is produced by the radioactive decay of short-lived [sup.129]I (1/2 life = 17 Ma). Lower atmospheric [sup.129]Xe/[sup.130]Xe ratios indicate that the degassing occurred before most [sup.129]I had decayed.
MORB sources show low, uniform [sup.3]He/[sup.4]He ratios whereas OIB sources are variable. The highest ratios occur in some Hawaiian and Icelandic basalts (Hilton et al., 1999). High [sup.3]He (relative to [sup.4]He) implies formation from less-degassed or undegassed primordial mantle. Supporting this, these basalts have FOZO isotopic compositions (Hilton et al., 1999). Arrays can be drawn, in Sr-Nd-Pb isotopic space, from the mantle components to a common point called FOZO for focal zone (Hart et al., 1992; Figs.4a, 4b and 5a). FOZO is regarded by some as a fifth mantle component of possibly primordial origin and common to (involved in mixing to form) all OIB.
An alternative explanation for high [sup.3]He/[sup.4]He ratios in specific OIB samples is that small-scale melting leads to localized sampling of highly depleted (low U+Th) sources that produced little [sup.4]He–the opposite of the primordial mantle hypothesis (Anderson, 1999; 2001; Meibom et al., 2003)! In this model, low but uniform [sup.3]He/[sup.4]He in MORB reflects a source having localized domains of depleted and enriched mantle, the latter containing recycled materials (e.g., ancient subducted crust) with high (U+Th)/He (high radiogenic [sup.4]He) and low [sup.3]He/[sup.4]He ratios. The large-scale melting that yields MORB homogenizes the isotopic signature from these domains but it is dominated by radiogenic He from enriched mantle. Slow-spreading ridges that have reduced melt production, and more localized melting, should have more variable [sup.3]He/[sup.4]He ratios (Anderson, 2001) but Georgen et al. (2003) present evidence that the slow-spreading Southwest Indian Ridge shows uniform [sup.3]He/[sup.4]He ratios. Another issue is that the Iceland samples with the highest [sup.3]He/[sup.4]He ratios on Earth have elevated [sup.87]Sr/[sup.86]Sr and lower [sup.143]Nd/[sup.144]Nd ratios relative to other Iceland basalts indicating they are not coming from a highly depleted source (Hilton et al., 1999).
Ne isotopes provide support for relatively undegassed domains in the mantle. Plots of [sup.20]Ne/[sup.22]Ne versus [sup.21]Ne/[sup.22]Ne show that basalts from the various localities define linear trends (e.g., MORB, Iceland and Loihi, Hawaii, Fig. 6; Hilton and Porcelli, 2003). Rare samples from each place have high [sup.20]Ne/[sup.22]Ne ratios that approach solar values. Because [sup.20]Ne and [sup.22]Ne are primordial (nonradiogenic or radioactive), it appears the mantle’s original Ne composition was “solar”. Each location has different maximum [sup.21]Ne/[sup.22]Ne ratios reflecting variable impact of nucleogenic [sup.21]Ne on mantle Ne isotopic ratios. The linear trends (Fig. 6) result from mixing between ‘LAir” and mantle Ne (McDougal and Honda, 1998) and most basalts are contaminated by atmospheric Ne derived from seawater or air. The Iceland and Loihi samples with the highest [sup.20]Ne/[sup.22]Ne ratios also have near-solar [sup.21]Ne/[sup.22]Ne ratios (e.g., Trieloff et al., 2000; 2003). Their primordial mantle Ne concentrations are apparently so high (un-degassed?) that nucleogenic production of [sup.21]Ne has not been able to substantially modify the [sup.21]Ne/[sup.22]Ne ratios from the solar values (Dixon et al., 2000; Moreira et al., 2001).
[FIGURE 6 OMITTED]
Most of the mantle may be substantially degassed/depleted (Anderson, 1999; Davies, 1999). Modeling by Coltice and Ricard (2002) suggests that the Loihi source contains < 3% primitive mantle, but these small amounts, perhaps as partially degassed peridotite, create the "primordial" noble gas signature. Consistent with this, Hanyu et al. (2001) report primitive Ne and Ar from Reunion, but lithophile isotopes (e.g., [sup.87]Sr/[sup.86]Sr) are unlike Iceland or Hawaii. There may be several types of relatively undegassed mantle sources with different evolutionary histories.
TRACE-ELEMENT RATIOS AND MANTLE COMPOSITIONS
Isotopes are used to identify the mantle components but they do not provide a clear chemical picture of what these components are, or how they are formed. Trace elements help but concentrations in flows reflect melting processes as well as source compositions. This problem is ameliorated by using the ratios of similarly incompatible elements. Two elements with similar incompatibility have comparable bulk distribution coefficients during melting. Their ratios in magmas are not overly affected by the percentage of melting and approach source ratios. Ratio differences between rocks will reflect variations in mantle composition, providing the elements used in the ratios are separated by [less than or equal to] 10 elements in the Sun and McDonough (1989) incompatibility list (Table 2, notes; Greenough et al., unpublished data). The pioneering work of Allegre et al. (1995) showed that the same mantle components identified using isotopes are delimited in multidimensional trace-element ratio space.
Figure 7 and Table 2 show selected, similarly incompatible-element-ratios that distinguish mantle components. All ratios have the slightly more incompatible element in the numerator. Thus, all normal – MORB ratios are lower than in OIB implying a highly depleted source for the former (Table 2). Even the average OIB depleted mantle (DM) has not experienced as much former melt extraction as MORB sources (Table 2) although Figure 6 illustrates their similarity.
[FIGURE 7 OMITTED]
HIMU shows high U/X, Nb/X, Ta/X and light REE/X ratios but low Rb/X, K/X, Pb/X and (commonly) Ba/X ratios where X represents various elements having similar but slightly lower incompatibility than the numerator element. This indicates HIMU is enriched in U, Nb, Ta and the light REE but depleted in K, Rb, Ba and Pb. It relates the low [sup.87]Sr/[sup.86]Sr ratio of HIMU to ancient Rb depletion and high [sup.206]Pb/[sup.204]Pb to U enrichment and/or Pb depletion. The data pattern supports the idea that HIMU sources contain ancient, recycled basaltic ocean crust (Zindler and Hart, 1986; Hofmann, 1997).
Processing in ancient subduction zones removed elements soluble in water-rich metasomatic fluids and concentrated insoluble high-field-strength elements (e.g., U, Nb, Ta; Weaver, 1991) by stabilizing oxide phases in an oxidizing environment. Possibly Nb is not enriched in HIMU sources (Niu and O’Hara, 2003) but this is not supported by Table 2, and other reviews. Kamber and Collerson (2000) proposed that correlated Zr and Nb concentrations in oceanic basalts indicate mixing between MORB mantle and sources (components) variably affected by metasomatic partial melts. Similarly, Halliday et al. (1995) argued that high U/Pb, low K/U and moderate Ba/Ce support HIMU formation at ocean ridges by metasomatism of small percentage melts that equilibrated with phlogopite, amphibole, and sulfides.
The EM2 sources show the highest Rb/X and lowest Sr/X ratios in the ocean basins (Table 2; X defined previous paragraph) which explains why they have the highest [sup.87]Sr/[sup.86]Sr ratios. Moderate enrichment in other large ion lithophile elements (Ba, K) and strong Nb depletion are consistent with several percent of subducted terrigenous sediment in the source (Weaver, 1991).
Many highly incompatible elements are more enriched in EM1 than EM2. The EM1 ratios involving Ba, Th, and K (in numerator) are the highest, and U, Nb and Ta are the lowest in the ocean basins, and comparing OIB only, HIMU is the reverse of EM1 (Table 2). For elements more compatible than Pb, (below Pb/P, Table 2), EM1 and EM2 form opposite extremes. The EM1 has the highest St/X, P/X, Zr/X, Hf/X, middle REE (Nd, Sm, Eu)/X ratios in OIB whereas EM2 ratios are mostly the lowest. The EM1 characteristics, particularly high Ba/X values, have been ascribed to ancient pelagic sediment (Weaver, 1991). Dostal et al. (1998) modeled EM1 by adding pelagic sediment to a HIMU-type source. Predicted large-ion lithophile-element concentrations were too high indicating that, in nature, these elements are partially lost from sediment during subduction; EM1 and EM2 may contain delaminated subcontinental lithospheric mantle. Continental basalts that melted Archean subcontinental lithosphere show the same pattern of high- or low-element ratios (relative to non-EM1 mantle components) as exhibited by EM1 (Table 2), but ratios are more extreme than in EM1 (Greenough et al., unpublished data).
Few trace-element ratio constraints can be placed on the primordial mantle–deep mantle–FOZO component common to all oceanic basalts. Allegre et al. (1995) showed that Hawaii (representing FOZO) is distinctive, in multi-trace element ratio space, relative to MORB, EM1, EM2 and HIMU. Baksi (2001) argued that FOZO can be identified using Nb/Y and Zr/Y ratios.
CHEMICAL STRUCTURE OF THE MANTLE
Variations in OIB and MORB geochemistry lead to models for chemical structure in the mantle. Most mantle heterogeneity appears related to melt extraction to form continental and oceanic crust, and recycling of oceanic and continental lithosphere back into the mantle. Compositional variability exists at all scales in the mantle (Zindler and Hart, 1986). Xenoliths document heterogeneity at the mineral scale. Individual flows on volcanoes indicate heterogeneity at the km-scale in the source region. High [sup.187]Os/[sup.188]Os ratios in some Hawaiian flows are consistent with low-temperature mantle melting of pyroxenite “blobs” or veins (Lassiter et al., 2000). Melting experiments on peridotitebasalt-peridotite “sandwiches” confirm that low-temperature rocks (basalt as eclogite) melt first (Takahashi and Nakajima, 2002). The volume of stratigraphic units (Makapuu stage of Koolau, Hawaii) ascribed to eclogite melting suggests the presence of [10.sup.3] [km.sup.3] blocks in the underlying mantle. Similarly, early-formed seamounts along island chains show more extreme trace element and isotopic compositions than islands indicating that components (e.g., EM1) have lower melting temperatures, melt first, and form distinct, localized, rock masses in the mantle (Devey et al., 2003). Niu et al. (2002) came to similar conclusions on the scale and melting behaviour of mantle domains underlying seamounts near the East Pacific Rise. At larger scales, individual islands, island chains and ocean basins (e.g., Indian Ocean) exhibit distinct signatures (Allegre, 2002). Pb isotope maps show anomalous mantle between 30[degrees] and 40[degrees] S latitude encircling much of the planet (Dupal anomaly; Zindler and Hart, 1986).
Opinions differ on how chemical heterogeneities are cycled through the mantle (e.g., Allegre, 2002; Hofmann, 2003; Van Keken et al., 2003). Mass balance calculations originally indicated that melting of primitive upper mantle produced the continental crust. The residue forms present-day depleted MORB mantle. The percentage of depleted mantle (30%; more recent estimates higher) resembled the amount of mantle above the 670 km seismic discontinuity suggesting a depleted upper mantle, a primitive lower mantle, and little or no exchange between the two (Hofmann, 2003; Bennett; 2003). The discovery of high [sup.3]He/[sup.4]He ratios in some OIBs (and solar Ne isotopes) was attributed to small amounts of primitive, noble-gas-rich lower mantle being entrained by mantle plumes rising from the 670 km boundary. Plumes were also seen as the carriers of recycled material (components) seen in OIB geochemistry (Hofmann, 1997). When seismic tomography showed subducted oceanic lithosphere penetrating the lower mantle, and plumes rising from the core-mantle boundary, an isolated lower mantle seemed impossible (Van Keken et al., 2003). Nevertheless, others argued that whole-mantle convection has been episodic or is a recent phenomenon (Allegre, 2002; Hofmann, 2003).
Some models suggest that primitive/un-degassed mantle does not exist (e.g., Davies, 1999; Anderson, 1999; Coltice and Ricard, 2002; Hamilton, 2003). None of the mantle components, including FOZO (which is actually “depleted”) with high [sup.3]He/[sup.4]He ratios and solar Ne, have lithophlle isotopic compositions requisite of primitive mantle (Bennett, 2003). Most arguments for large amounts of un-degassed primitive mantle account for the [sup.40]Ar budget by estimating Earth’s [sup.40]K content from K/U ratios (Davies, 1999; Allegre, 2002). If the K content has been overestimated, primitive lower-mantle storing [sup.40]Ar is not required.
There is an “alternative” model for the Earth (e.g., Anderson, 1999, 2000; Hamilton, 2002, 2003). The entire mantle has been processed, melted and depleted, much of it during the Earth’s earliest history (see Bennett, 2003). It is layered and there is no mass transfer across the 670 km boundary (seismic tomography results are wrong!). Mantle plumes do not exist and hot-spot trails are seen as the result of propagating cracks in the lithosphere. Ocean basin heat flow implies that most radioactivity occurs in the upper mantle, which has been progressively enriched in radioactive K and U by recycling of crustal materials back into the mantle. Both MORB and OIB are derived from the same upper mantle source containing recycled materials. MORB magmas are more homogenous because of larger percentages and volumes of melting. The high [sup.3]He/[sup.4] He ratios of a few OIBs reflect melting of small pockets of highly depleted mantle (low U + Th) (Anderson, 2001). Central to testing this model is whether differences in sampling scale (MORB large, OIB small) and melting process (OIB preferentially sample low-melting point domains) can explain the major compositional differences (Allegre et al., 1995; Allegre, 2002) between OIB and MORB or whether their sources are truly different. Non-conclusive numerical modeling of heterogeneity development in the mantle suggests that melting process may account for the differences (Kellogg et al., 2002).
FUTURE RESEARCH DIRECTIONS
Many hypotheses for the origin of the mantle components show that there is much left to learn about their composition and significance. Possibly new isotopic systems such as Li, or further work with trace-element ratios will help clarify these issues. How components (heterogeneities) cycle through the mantle is also problematic. Whether they reflect plume-entrained material brought from the lower mantle, the 670 km discontinuity, or simply the scale and process of melting requires further work. Information is needed on 1) what the components are, 2) how they melt, 3) the length-scales of heterogeneities, and 4) whether seismic tomography results really confirm the existence of complete mantle convection. Clearly, OIBs will continue to play an important role in understanding the evolution of the mantle and the Earth.
R. Corney prepared diagrams. SMU provided office support for LMM and JDG. JDG and JD acknowledge NSERC operating grants. The comments of reviewers (G. Jenner, B. Murphy and D. Piper) and the editor (G. Pe-Piper) led to major improvements in the manuscript.
Supplementary material, including references and notes on whole-rock data, can be viewed at the following location: http://www.gac.ca/JOURNALS/geocan.html.
Allegre, C.J., 2002, The evolution of mantle mixing: Philosophical Transactions of the Royal Society of London A, v. 360, p. 2411-2431.
Allegre, C.J., Schiano, P. and Lewin, E., 1995, Differences between oceanic basalts by multitrace element ratio topology: Earth and Planetary Science Letters, v. 129, p. 1-12.
Anderson, D.L., 1999, A theory of the Earth; Hutton and Humpty Dumpty and Holmes, in Craig, G.Y. and Hull, J.H., eds., James Hutton; present and future: Geological Society Special Publications, v. 150, p. 13-35.
Anderson, D.L., 2000, The thermal state of the upper mantle; no role for mantle plumes: Geophysical Research Letters, v. 27, p. 3623-3626.
Anderson, D.L., 2001, A statistical test of the two reservoir model for helium isotopes: Earth and Planetary Science Letters, v. 193, p. 77-82.
Baksi, A.K., 2001, Search for a deep-mantle component in mafic lavas using a Nb-Y-Zr plot: Canadian Journal of Earth Sciences, v. 38, p. 813-824.
Basaltic Volcanism Study Project, 1981, Basaltic Volcanism on the Terrestrial Planets: Pergamon Press Inc., New York, 1286 pp.
Becker, H., 2000, Re-Os fractionation in eclogites and blueschists and the implications for recycling of oceanic crust into the mantle: Earth and Planetary Science Letters, v. 177, p. 287-300.
Bennett, V.C., 2003, Compositional evolution of the mantle, in Carlson, R.W., ed., The Mantle and Core: Holland, H.D. and Turekian, K.K., eds., Volume 2, Treatise on Geochemistry: Elsevier-Pergamon, Oxford, p. 493-519.
Bizzarro, M., Simonettie, A., Stevenson, R.K. and David, J., 2002, Hf isotope evidence for a hidden mantle reservoir: Geology; v. 30, p. 771-774.
Bourdon, B. and Sims, K.W.W., 2003, U-series constraints on intraplate basaltic magmatism: Reviews in Mineralogy and Geochemistry, v. 52, p. 215-254.
Coltice, N. and Ricard, Y. 2002. On the origin of noble gases in mantle plumes: Philosophical Transactions of the Royal Society of London, A, v. 360, p. 2633-2648.
Davies, G.F., 1999, Geophysically constrained mantle mass flows and the [sup.40]Ar budget: a degassed lower mantle?: Earth and Planetary Science Letters, v. 166, p. 149-162.
Devey, C.W., Lackschewitz, K.S., Mertz, D.E, Bourdon, B., Cheminee, J.-L., Dubois, J., Guivel, C., Hekinian, R. and Stoffers, P., 2003, Giving birth to hotspot volcanoes: distribution and composition of young seamounts from the seafloor near Tahiti and Pitcairn Islands: Geology, v. 31, p. 395-398.
Dixon, E.T., Honda, M., McDougall, I., Campbell, I.H. and Sigurdsson, I., 2000, Preservation of near-solar neon isotopic ratios in Icelandic basalts: Earth and Planetary Science Letters, v. 180, p. 309-324.
Dostal, J., Cousens, B. and Dupuy, C., 1998, The incompatible element characteristics of an ancient subducted sedimentary component in ocean island basalts from French Polynesia: Journal of Petrology, v. 39, p. 937-952.
Eggler, D.H. and Holloway, J.R., 1977, Partial melting of peridotite in the presence of [H.sub.2]O and C[O.sub.2]: principles and review: Bulletin Oregon, Department of Geology and Mineral Industries, Issue 96, p. 1536.
Eiler, J.M., Farley, K.A., Valley, J.W., Hauri, E.H., Craig, H., Hart, S.R. and Stolper E.M., 1997, Oxygen isotope variations in ocean island basalt phenocrysts: Geochimica et Cosmochimica Acta, v. 61, p. 2281-2293.
Eisele, J., Sharma, M., Galer, S.J.G., Blichert Toft, J., Devey, C.W. and Hofmann, A.W., 2002, The role of sediment recycling in EM-1 inferred from Os, Pb, Hf, Nd, Sr isotope and trace element systematics of the Pitcairn Hotspot: Earth and Planetary Science Letters, v. 196, p. 197-212.
Elliott, T., Jeffcoate, A. and Bouman, C., 2004, The terrestrial Li isotope cycle: light-weight constraints on mantle convection: Earth and Planetary Science Letters, v. 220, p. 231-245.
Faure, G., 2001, Origin of Igneous Rocks; The isotopic Evidence: Springer, Berlin, 496 p.
Fitton, J.G., Saunders, A.D., Kempton, P.D. and Hardarson, B.S., 2003, Does depleted mantle form an intrinsic part of the Iceland plume?: Geochemistry, Geophysics, Geosystems, v. 4, Paper 2002GC000424.
Fryer, B.J. and Greenough, J.D., 1992, Evidence for mantle heterogeneity from platinum-group-element abundances in Indian Ocean basalts: Canadian Journal of Earth Sciences, v. 29, p. 2329-2340.
Gasperini, D., Blichert-Toff, J., Bosch, D., Del Moro, A., Macera, P., Telouk, P., Albarede, F., 2000, Evidence from Sardinian basalt geochemistry for recycling of plume heads into the Earth’s mantle: Nature, v. 408, p. 701-704.
Georgen, J.E., Kurz, M.D., Dick, H.J.B. and Lin, J., 2003, Low [sup.3]He/[sup.4]He ratios in basalt glasses from the western Southwest Indian Ridge (10[degrees]-24[degrees]E): Earth and Planetary Science Letters, v. 206, p. 509-528.
Greenough, J. D., 1988, Minor Phases in the Earth’s Mantle: Evidence from trace and minor element patterns in primitive alkaline magmas: Chemical Geology; v. 69, p. 177-192.
Greenough, J.D., Dostal, J. and Mallory-Greenough, L.M., 2005. Oceanic Island Volcanism I: Mineralogy and Petrology: Geoscience Canada, v. 32, p. 29-45.
Greenough, J.D. and Kyser, T.K., 2003, Contrasting Archean and Proterozoic lithospheric mantle: isotopic evidence from the Shonkin Sag sill (Montana): Contributions to Mineralogy and Petrology, v. 145, p. 169-181.
Haase, K.M., 1996, The relationship between the age of the lithosphere and the composition of oceanic magmas: constraints on partial melting, mantle sources and the thermal structure of the plates: Earth and Planetary Science Letters, v. 144, p. 75-92.
Halliday, A.N., Lee, D.-C., Tommasini, S., Davies, G.R., Paslick, C.R., Fitton, J.G. and James, D.E., 1995, Incompatible trace elements in OIB and MORB and source enrichment in the sub-oceanic mantle: Earth and Planetary Science Letters, v. 133, p. 379-395.
Hamilton, W.B., 2002, The closed upper-mantle circulation of plate tectonics: American Geophysical Union Geodynamics Series, v. 30, p. 359-410.
Hamilton, W.B., 2003, An alternative Earth: GSA Today, v. 13, p. 4-12.
Hanan, B.B., Blichert-Toft, J., Kingsley, R. and Schilling, J.-G., 2000, Depleted Iceland mantle plume geochemical signature: artifact of multicomponent mixing? Geochemistry, Geophysics, Geosystems, v. 1, Paper 1999GC000009.
Hanyu, T., Dunai, T.J., Davies, G.R., Kaneoka, I., Nohda, S. and Uto, K., 2001, Noble gas study of the Reunion hotspot: evidence for distinct less-degassed mantle sources: Earth and Planetary Science Letters, v. 193, p. 83-98.
Hart, S.R., Hauri, E.H., Oschmann, J.A. and Whitehead, J.A., 1992, Mantle plumes and entrainment–isotopic evidence: Science, v. 256, p. 517-520.
Hauri, E.H., 2002, Osmium isotopes and mantle convection: Philosophical Transactions of the Royal Society of London, A, v. 360, p. 2371-2382.
Hilton, D.R., Gronvold, K., Macpherson, C.G. and Castillo, RR., 1999, Extreme [sup.3]He/[sup.4]He ratios in Northwest Iceland; constraining the common component in mantle plumes: Earth and Planetary Science Letters, v. 173, p. 53-60.
Hilton, D.R. and Porcelli, D., 2003, Noble gases as mantle tracers, in Carlson, R.W., ed., The Mantle and Core: Holland, H.D. and Turekian, K.K., eds., Volume 2, Treatise on Geochemistry: Elsevier-Pergamon, Oxford, p. 277-318.
Hirose, K. and Kushiro, I., 1993, Partial melting of dry peridotites at high pressures: determination of composition of melts segregated from peridotite using aggregates of diamond: Earth and Planetary Science Letters, v. 114, p. 477-489.
Hofmann, A.W., 1997, Mantle geochemistry: the message from oceanic volcanism: Nature, v. 385, p. 219-229.
Hofmann, A.W., 2003, Sampling mantle heterogeneity through oceanic basalts: isotopes and trace elements, in Carlson, R.W., ed., The Mantle and Core: Holland, H.D. and Turekian, K.K., eds., Volume 2, Treatise on Geochemistry: Elsevier-Pergamon, Oxford, p. 61-101.
Kamber, B.S. and Collerson, K.D., 1999, Origin of ocean island basalts: a new model based on lead and helium isotope systematics: Journal of Geophysical Research, v. 104, p. 25479-25491.
Kamber, B.S. and Collerson, K.D., 2000, Zr/Nb systematics of ocean island basalts reassessed–the case for binary mixing: Journal of Petrology, v. 41, p. 1007-1021.
Kellogg, J.B., Jacobsen, S.B. and O’Connell, R.J., 2002, Modeling the distribution of isotopic ratios in geochemical reservoirs: Earth and Planetary Science Letters, v. 204, p. 183-202.
Kushiro, I., 1968, Compositions of magmas formed by partial zone melting of the Earth’s upper mantle: Journal of Geophysical Research, v. 73, p. 619-634.
Lassiter, J.C., Hauri, E.H., Reiners, P.W. and Garcia, M.O., 2000, Generation of Hawaiian post-erosional lavas by melting of a mixed lherzolite/pyroxenite source: Earth and Planetary Science Letters, v. 178, p. 269-284.
McDougall, I. and Honda, M., 1998, Primordial solar noble-gas component in the Earth: consequences for the origin and evolution of the Earth and its atmosphere, in Jackson, I., ed., The Earth’s Mantle: Composition, Structure and Evolution, Cambridge University Press, Cambridge, U.K., p. 159-187.
Meibom, A., Anderson, D.L., Sleep, N.H., Frei, R., Chamberlain, C.P., Hren, M.T. and Wooden, J.L., 2003, Are high [sup.3]He/[sup.4]He ratios in oceanic basalts an indicator of deep-mantle plume components?: Earth and Planetary Science Letters, v. 208, p. 197-204.
Milner, S.C. and le Roex, A.R, 1996, Isotope characteristics of the Okenyenya igneous complex, northwestern Namibia; constraints on the composition of the early Tristan Plume and the origin of the EM 1 mantle component: Earth and Planetary Science Letters, v. 141, p. 277-291.
Moreira, M., Breddam, K., Curtice, J. and Kurz, M.D., 2001, Solar neon in Icelandic mantle: new evidence for an undegassed lower mantle: Earth and Planetary Science Letters, v. 185, p. 15-23.
Niu, Y., Regelous, M., Wendt, I.J., Batiza, R. and O’Hara, MJ., 2002, Geochemistry of near-EPR seamounts; importance of source vs. process and the origin of enriched mantle component: Earth and Planetary Science Letters, v. 199, p. 327-345.
Niu, Y. and O’Hara, MJ. 2003. Origin of ocean island basalt: a new perspective from petrology, geochemistry, and mineral physics considerations: Journal of Geophysical Research, v. 108.
Pearson, D.G. and Nowell, G.M., 2002, The continental lithospheric mantle: characteristics and significance as a mantle reservoir: Philosophical Transactions of the Royal Society of London, A, v. 360, p. 2383-2410.
Roeder, P.L. and Emslie, R.F., 1970, Olivine-liquid equilibrium: Contributions to Mineralogy and Petrology, v. 29, p. 275-289.
Rollinson, H.G., 1993, Using Geochemical Data: Evaluation, Presentation, Interpretation, Longman Group, Harlow, UK, 352 p.
Schiano, P., Burton, K.W., Dupre, B., Birck, J.L., Guille, G. and Allegre, C.J., 2001. Correlated Os-Pb-Nd-Sr isotopes in the Austral-Cook Chain basalts; the nature of mantle components in plume sources: Earth and Planetary Science Letters, v. 186, p. 527-537.
Sun, S.S and McDonough, WE, 1989, Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes, in Saunders A.D. and Norry M.J., eds., Magmatism in the Ocean Basins: Geological Society Special Publication 42, p. 313-345.
Takahashi, E. and Nakajima, K., 2002, Melting process in the Hawaiian plume: an experimental study, in Takahashi, E., Lipman, P.W., Garcia, M.O., Naka, J. and Aramaki, S., eds., Hawaiian Volcanoes, Deep Water Perspectives: Geophysical Monograph 128, American Geophysical Union, Washington DC, p. 403-418.
Tatsumi, Y., 2000, Continental crust formation by crustal delamination in subduction zones and complementary accumulation of the enriched mantle I component in the mantle: Geochemistry, Geophysics Geosystems, v. 1, paper 2000GC000094.
Trieloff, M., Kunz, J., Clague, D.A., Harrison, D. and Allegre, C.J., 2000, The nature of pristine noble gases in mantle plumes: Science, v 288, p. 1036-1038.
Trieloff, M., Falter, M. and Jessberger, E.K. 2003. The distribution of mantle and atmospheric argon in oceanic basalt glasses: Geochimica et Cosmochimica Acta, v. 67, p. 1229-1245.
Van Keken, RE., Ballentine, C.J. and Hauri, E.H., 2003, Convective mixing in the Earth’s mantle, in Carlson, R.W., ed., The Mantle and Core: Holland, H.D. and Turekian, K.K., eds., Volume 2, Treatise on Geochemistry: Elsevier-Pergamon, Oxford, p. 471-491.
Walker, R.J., Storey, M., Kerr, A.C., Tarney, J. and Arndt, N.T. 1999. Implications of [sup.187]Os isotopic heterogeneities in a mantle plume; evidence from Gorgona Island and Curacao: Geochimica et Cosmochimica Acta, v 63, p. 713-728.
Watson, S. and McKenzie, D., 1991, Melt generation by plumes; a study of Hawaiian volcanism: Journal of Petrology, v. 32, p. 501537.
Weaver, B.L., 1991, The origin of ocean island basalt end-member compositions: trace element and isotopic constraints: Earth and Planetary Science Letters, v. 104, p. 381-397.
Wirth, R. and Rocholl, A., 2003, Nanocrystalline diamond from the Earth’s mantle underneath Hawaii: Earth and Planetary Science Letters, v 211, p. 357-369.
Zindler, A. and Hart, S.R., 1986, Chemical geodynamics: Annual Review of Earth and Planetary Sciences, v. 14, p. 493-571.
Accepted as received 1 Dec 2004
John D. Greenough (1), Jaroslav Dostal (2), and Leanne M. Mallory-Greenough (3)
(1) Department of Earth and Environmental Sciences, University of British Columbia-Okanagan, 3333 College Way, Kelowna, BC, VIV 1V7, Canada. E-mail: John.Greenough@ubc.ca Tel: 250 762 5445 ext. 7520,” Fax: 250 470 6005
(2) Department of Geology, Saint Mary’s University, Halifax, NS, B3H 3C3, Canada. E-mail: email@example.com
(3) Department of Geology, University of Toronto, 22 Russell St., Toronto, ON, M5S 3B1, Canada. E-mail: firstname.lastname@example.org
Table 1: Average whole-rock major element, trace element and isotopic
data for selected oceanic islands.
Island Gough Sao Miquel Easter Is.
Series Alkaline Alkaline Alkaline
Type EM1 EM2 DM
Si[O.sub.2] 49.53 47.42 48.67
Ti[O.sub.2] 3.1 3.45 3.71[Al.sub.2][O.sub.3] 14.27 13.82 16.04
FeO 10.4 10.87 10.77
CaO 8.21 10.09 9.79
MgO 8.02 8.74 6.34
MnO 0.15 0.17 0.17[K.sub.2]O 2.46 2 0.81[Na.sub.2]O 3.22 2.9 3.27[P.sub.2][O.sub.5] 0.65 0.55 0.43
Mg# 0.58 0.61 0.53
Cs 0.23 0.46 0.088
Rb 48 46 19
Sr 769 670 406
Ba 726 584 214
La 66.4 50.3 28.6
Ce 96.7 105 57.3
Pr 10.8 13.1 5.3
Nd 51.8 47.7 35.6
Sm 9.85 9.34 8.03
Eu 3.38 3.12 2.33
Gd 8.68 6.95
Tb 1.01 1.21 1.11
Dy 6.32 6.34
Er 2.88 3.31
Yb 2.35 3.19
Lu 0.333 0.465
Y 28.9 30.9 32.8
Zr 309 290 247
Hf 8.21 7.99 6.58
Nb 50.5 60.1 34.7
Ta 3.56 4.32 2.63
Th 5.42 5.23 3.02
U 1.33 1.37 0.95
Ga 21.3 22
V 185 264 248
Sc 20.6 22.4 25.9
Cr 217 357 143
Co 46.8 43.7 34.3
Ni 194 179 76
Cu 33.5 61.2 52.5
Pb 4.3 1.53
Zn 108 115 101
Sn 3.09 3.46
Sb 0.217 0.29
Re 90.2 392.3
Os 30.6 63.4[sup.143]Nd/[sup.144]Nd 0.512574 0.512803 0.5129694[sup.87]Sr/[sup.86]Sr 0.70532 0.7044491 0.7031304[sup.206]Pb/[sup.204]Pb 18.4235 19.74831 19.512556[sup.207]Pb/[sup.204]Pb 15.608 15.686621 15.611[[sup.208]Pb/[sup.204]Pb 38.892 39.84669 39.197[sup.187]OS/[sup.188]OS 0.140293[sup.176]Hf/[[sup.177]Hf
Island Tristan Isabela Tutuila
Series S. Alkaline Tholeiitic Alkaline
Type EM1 DM EM2
Si[O.sub.2] 46.31 49.11 46.84
Ti[O.sub.2] 3.31 3.34 3.94[Al.sub.2][O.sub.3] 15.57 14 12.68
FeO 11.38 12.88 12.49
CaO 10.21 10.43 9.42
MgO 6.54 5.94 9.95
MnO 0.18 0.2 0.18[K.sub.2]O 2.21 0.58 1.09[Na.sub.2]O 3.57 3.14 2.91[P.sub.2][O.sub.5] 0.73 0.38 0.51
Mg# 0.52 0.47 0.61
Cs 0.87 4.8
Rb 54 11 36
Sr 1059 317 526
Ba 666 120 221
La 56.7 25.6 24.1
Ce 124.4 55.9 56.8
Pr 14.3 6.3
Nd 65 33.2 39.6
Sm 10.7 8.53 9.57
Eu 3.17 2.58 2.76
Gd 7.74 8.79 9.51
Tb 0.98 1.46 1.15
Dy 5.83 7.61 4.6
Ho 1.02 1.52
Er 2.82 4
Tm 0.44 0.549
Yb 2.1 3.89 2.12
Lu 0.288 0.579 0.25
Y 27.2 39 32
Zr 281 328 234
Hf 7.73 7.17 5.77
Nb 72.2 28.7 33.5
Ta 4.54 2.16 2.19
Th 6.97 2.76 3.34
U 1.72 0.78 0.66
V 282 332 280
Sc 21.6 27.6 29
Cr 133 121 364
Co 49.3 48.3 42
Ni 72 52 229
Cu 43.7 82.9
Pb 4.48 1.62 3.2
Zn 100 120
Sn 3.43 1.94
Re 455.4 209
Os 38.8[sup.143]Nd/[sup.144]Nd 0.5125603 0.5129837 0.5127315[sup.87]Sr/[sup.86]Sr 0.7049259 0.7030976 0.7058075[sup.206]Pb/[sup.204]Pb 18.610522 19.138839 19.095529[sup.207]Pb/[sup.204]Pb 15.547087 15.573037 15.608588[[sup.208]Pb/[sup.204]Pb 39.066913 38.763392 39.290706[sup.187]OS/[sup.188]OS[sup.176]Hf/[[sup.177]Hf 0.2830689
Island St. Helena Iceland Tahiti
Series Alkaline Tholeiitic Alkaline
Type HIMU DM EM2
Si[O.sub.2] 46.55 48.85 45.33
Ti[O.sub.2] 3.09 1.4 3.62[Al.sub.2][O.sub.3] 14.98 15.05 13.21
FeO 12.15 10.59 12.35
CaO 10.6 12.19 10.94
MgO 7.73 9.31 9.89
MnO 0.19 0.18 0.18[K.sub.2]O 1.11 0.22 1.24[Na.sub.2]O 3.05 2.02 2.6[P.sub.2][O.sub.5] 0.56 0.17 0.63
Mg# 0.54 0.63 0.59
Cs 0.15 0.15
Rb 27 6 36
Sr 596 180 662
Ba 338 62 430
La 50.7 7.6 43.3
Ce 102 18.9 99.9
Pr 10.5 1.6 12.2
Nd 43.3 12.1 48.8
Sm 8.61 3.25 9.88
Eu 2.68 1.16 3.22
Gd 7.78 2.96 8.99
Tb 1.2 0.65 1.33
Dy 5.85 3.24 6.82
Ho 1.28 0.69 1.19
Er 2.82 2.01 2.82
Tm 0.4 2.41 0.381
Yb 2.73 2.18 2.14
Lu 0.406 0.327 0.295
Y 32.6 24 31.5
Zr 245 92 305
Hf 5.62 3.83 7.64
Nb 64.9 9.9 47.9
Ta 3.48 1.19 3.54
Th 3.72 1.03 5.44
U 1.21 0.36 1.49
V 229 266 287
Sc 30 38.7 22.9
Cr 299 366 397
Co 65.1 51.8 57.3
Ni 136 157 209
Cu 54.0 122.6 79.1
Pb 2.48 1.61 3.62
Zn 109 89 115
NIo 4.44 0.88
Sn 2.39 3.08
Sb 0.247 0.152
Re 146.8 479
Os 17.9 122[sup.143]Nd/[sup.144]Nd 0.51288 0.5130465 0.5128357[sup.87]Sr/[sup.86]Sr 0.7029127 0.703183 0.7043635[sup.206]Pb/[sup.204]Pb 20.690347 18.690463 19.063913[sup.207]Pb/[sup.204]Pb 15.767774 15.476551 15.557[[sup.208]Pb/[sup.204]Pb 39.717754 38.043333 38.676696[sup.187]OS/[sup.188]OS 0.132117[sup.176]Hf/[[sup.177]Hf 0.282888 0.2832239
Island Aitutaki Mangaia Tubuai
Series S. Alkaline Alkaline S. Alkaline
Type EM1 HIMU HIMU
Si[O.sub.2] 43.01 45.19 44.04
Ti[O.sub.2] 2.53 2.97 2.93[Al.sub.2][O.sub.3] 12.01 13.29 13.32
FeO 12.03 12.99 13.48
CaO 12.11 12.2 11.68
MgO 11.57 8.94 9.16
MnO 0.2 0.21 0.24[K.sub.2]O 1.52 0.88 1.04[Na.sub.2]O 4.02 2.87 3.49[P.sub.2][O.sub.5] 1.01 0.46 0.63
Mg# 0.65 0.55 0.56
Li 13.7 8.8 6.6
Cs 0.29 0.27
Rb 43 18 28
Sr 1148 508 699
Ba 843 278 355
La 71.8 38.9 62.8
Ce 131.8 81.3 129.2
Pr 9.3 14.3
Nd 59.9 38.5 55
Sm 11.41 8.09 9.57
Eu 3.52 2.52 2.99
Gd 7.07 8.25
Tb 1.44 1.03 1.02
Dy 5.44 6.28
Ho 1.03 0.98
Er 2.6 2.82
Tm 0.331 0.322
Yb 1.9 2.06 2.09
Lu 0.256 0.298 0.293
Y 31 28.5 29.3
Zr 284 223 271
Hf 4.73 4.98 7.14
Nb 71.5 51.1 78.8
Ta 4.53 3.4 5.54
Th 12.97 4.17 8.02
U 2.59 1.12 2.06
Ga 20.4 18.2
V 300 263
Sc 31.3 26.2
Cr 173 464 361
Co 61.9 53.6
Ni 133 169 164
Cu 116.7 102.3
Pb 7.9 2.4 4
Zn 120 111
Re 51 325 398
Os 27.4 274 351[sup.143]Nd/[sup.144]Nd 0.5127642 0.5128875 0.512896[sup.87]Sr/[sup.86]Sr 0.7047786 0.7028407 0.7028247[sup.206]Pb/[sup.204]Pb 19.018375 21.658163 21.07552[sup.207]Pb/[sup.204]Pb 15.71375 15.888395 15.75992[[sup.208]Pb/[sup.204]Pb 38.951 40.540674 40.3344[sup.187]OS/[sup.188]OS 0.137 0.1768[sup.176]Hf/[[sup.177]Hf 0.2829436
Notes: Island = Island or area where data came from. Series = whether
rocks are tholeiitic, alkaline or strongly alkaline (S. Alkaline)
based on total alkalis versus silica diagram. Type = mantle source type
where EM1 = enriched mantle 1, EM2 = enriched mantle 2; HIMU =
High [micro] = High U/Pb; DM = depleted mantle. Columns represent
averages. Major element oxides in wt. % recalculated to 100% volatile
free with total Fe as FeO. Mg# = Mg/(Mg+0.9Fe) atomic. Trace element
concentrations in ppm except Re and Os in ppt and Au, Pd, Pt, Ru,
and Ir in ppb.
Radiogenic isotopes (Nd, Sr, Ph, Os, Hf) represent initial ratios
(corrected for age if necessary).
Averages for more islands, the number of samples averaged, standard
deviations, data sources and data reduction procedures appear with
the supplementary materials.
Table 2: Ratios of similarly incompatible elements in end-member
oceanic island basalts.
HIMU EM2 EM1 DM MORB
Rb/Ba1 0.074 0.104 0.077 0.084 0.089
Ba/Nb4 5 8 10.1 8.1 2.7
Th/Ce6 0.05 0.052 0.066 0.062 0.016
U/K3 0.000175 0.000111 0.000094 0.000133 0.000078
U/Pb6 0.49 0.34 0.34 0.35 0.16
Nb/K2 0.0077 0.004 0.0034 0.0049 0.0039
Nb/Sr8 0.107 0.069 0.07 0.071 0.026
Ta/K1 0.0005 0.00032 0.00024 0.00043 0.00022
K/La1 168 277 356 259 240
Ce/Pb1 36 29 23 24 25
Ce/P5 0.043 0.039 0.033 0.029 0.015
Pb/P4 0.0012 0.0015 0.0015 0.0015 0.0006
Sr/Zr6 2.4 2.2 2.8 2.2 1.2
P/Zr5 9.7 8.8 10.6 8.6 6.9
Nd/Ti8 0.0025 0.0021 0.0027 0.0015 0.001
Zr/Ti5 0.014 0.012 0.015 0.012 0.01
Hf/Ti4 0.00033 0.00031 0.00036 0.00033 0.00027
Ti/Y5 598 680 638 489 271
Y/Yb4 13 15 16 10 9
Notes: Bold = highest ratio in OIB; underlined italics = lowest ratio
in OIB. Trace element ratios were calculated from average element
concentrations (ppm of cations) for each oceanic island listed below.
These ratios were then averaged. All ratios have the more incompatible
element in the numerator and the number following the ratio (e.g.
Ta/K1) shows how far apart the elements are in the Sun and McDonough
(1989) incompatibility list (most incompatible Cs, Tl, Rb, Ba, W, Th,
U, Nb, Ta, K, La, Ce, Pb, Pr, Mo, Sr, P, Nd, F, Sm, Zr, Hf, Eu, Sn, Sb,
Ti, Gd, Tb, Dy, Li, Y, Ho, Er. Tm, Yb, Lu; least incompatible).
Ratios are organized according to the incompatibility of numerator
cations (e.g. Rb most incompatible, Ba somewhat less incompatible,
etc.). HIMU = 3 islands (St. Helena, Mangaia (Austral-Cook), Tubuai
(Austral-Cook)); EM2 = 5 islands (Sao Miguel (Azores), Tutuila
(Samoa), Upolu (Samoa), Tahiti (Society), Eiao (Marquesas)); EM1 = 5
islands (Gough, Kergulen, Pitcairn, Tristan da Cunha, Aitutaki
(Austral-Cook)); DM = 4 islands (Easter Island, Floreana (Galapagos),
Isabela (Galapagos), Iceland); MORB = Ratios from aver-age N-MORB,
in Sun and McDonough (1989).
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