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   Metasomatic OIB

The isotopic signature in OIB mantle sources:
the metasomatic alternative

Sébastien Pilet1, Jean Hernandez1 & Paul J. Sylvester2

1University of Lausanne, Switzerland
2Memorial University of Newfoundland, Canada,,


Since the chemical variability of ocean island basalts (OIB) was described in the 1980s, it has become clear that the Earth's mantle is heterogeneous on small and large scales. In order to explain the isotopic variability observed in OIB, several distinct mantle "end-members" have been proposed (e.g. DM, HIMU, EM1, EM2, FOZO and C) [White, 1985; Zindler & Hart, 1986]. The term "end-member" (or "component") implies a mantle reservoir or reservoirs with specific isotopic composition, and implies two things:

  1. Fractionation of the radioactive parent/daughter ratio for different isotopic systems (e.g. U/Pb, Th/Pb, Sm/Nd, Rb/Sr, Re/Os etc.). Because of the extremely slow diffusion rate in the deep mantle [Hofmann & Hart, 1978], this fractionation can happen only at shallow depth, i.e., in the upper mantle or the crust.
  2. Isolation of the "reservoir" for a long period of time (1- 2 Ga). This is necessary in order for the different isotopic systems to evolve to the specific signatures observed in the "end members".

In addition, basalt generation from the mantle source implies:

  1. The source material must yield, by partial melting, the major- and trace-element chemical composition observed in OIB. This limits possible parent/daughter isotope-ratio fractionation in the source and, as a result, the possible isolation time interval required to generate the different isotopic signatures.
  2. Pure end-member compositions are practically never observed in oceanic islands, suggesting that several end-members must be present in the source in order to generate the isotopic composition range observed in the OIB of a single island.

1. Hypotheses for the origin of OIB

1.1 Ancient recycled subducted oceanic crust model

The proposal made more than a decade ago, that the variation of trace element and isotopic ratios in OIB is related to the recycling of ancient oceanic crust [Hofmann & White, 1982] and pelagic or terrigenous sediment assimilation [Weaver, 1991, Chauvel et al., 1992], is now widely accepted. This model is popular because:

  1. The recycling of oceanic crust through the mantle is a undeniable fact;
  2. The generation of arc volcanism from the subduction and dehydration of oceanic crust suggests major fractionation of trace elements in the residual subducted oceanic crust;
  3. Positive Nb and negative Pb anomalies are observed both in OIB and in MORB. This suggests a link in the generation of these two types of basalt;
  4. The addition of pelagic and terrigenous sediments to subducted oceanic crust could explain the chemical and isotopic differences observed between different end-members signatures.

Nevertheless, several chemical and physical arguments suggest that subducted oceanic crust is not the source of OIB [e.g. Sun & McDonough, 1989; Halliday et al., 1995. Niu & O'Hara [2003] state, provocatively:

"There is no obvious association between ancient recycled oceanic crust and OIB (mantle plumes) in terms of petrology, geochemistry, and mineral physics".

This claim is based on the following arguments:

  1. Melting of oceanic crust with basaltic or picritic composition cannot produce the high-magnesian lavas that are parental to most OIB. Primitive OIB melts (>15% MgO) are probably more magnesian than bulk oceanic crust (<13% MgO).
  2. Oceanic crusts subjected to subduction zone dehydration reactions must be depleted in water-soluble incompatible elements such as Ba, Rb, Cs, Th, U, K, Sr and Pb relative to water-insoluble incompatible elements such as Nb, Ta, Zr, Hf and Ti, etc. Residual crust with such trace-element systematics is unlikely to be the fertile source of OIB.
  3. Ancient oceanic crust is isotopically too depleted to produce the isotopic signatures of most OIB. OIB Sr-Nd-Hf isotopes preserve no signature that indicates previous subduction zone dehydration histories.
  4. Subducted oceanic crust at shallow lower-mantle conditions form mineral assemblages that are much (>2.3%) denser than the ambient peridotitic mantle. If subducted crust melts in the deep lower mantle, this melt, depending upon its composition, may have still greater (up to 15%) density than solid peridotitic mantle. Therefore ancient oceanic crust that has subducted into the deep lower mantle cannot return in bulk to the upper mantle in either a solid or molten state.

1.2 The problem of major element composition

In addition to the problems highlighted by Niu & O'Hara [2003], the question of the silica-deficient composition of OIB was pointed out many years ago [Green et al., 1967]. Many ocean island rocks are characterized by silica undersaturated compositions (low silica and high alkali contents), whereas the partial melts of oceanic-crust lithologies produce silica saturated liquids [Green et al., 1967]. Partial melting of a mixture of subducted crust and peridotite produced by mechanical desegregation of subducted eclogite [Kogiso et al., 1998], or by hybridization of peridotite by partial melts of eclogite [Yaxley & Green, 1998], generate alkali glasses. But these glasses contain too high Al2O3 and too low CaO to represent a potential source of OIB. Experimental data indicate that partial melting of peridotite in the presence of carbonate [Hirose, 1997] or silica-deficient pyroxenite [Hirschmann et al., 2003; Kogiso & Hirschmann, 2003; Keshava et al., 2004] produces liquids which are close to the nephelinite, basanite or alkali-basalt major element compositions observed in OIB. Conversion of silica-saturated MORB to silica-deficient pyroxenite is possible, but requires either hybridization with peridotite, metasomatism by silica-deficient fluid such as carbonatite, or removal of low-pressure partial melts or fluids in the presence of olivine or spinel [Hirschmann et al., 2003, Lundstrom et al., 2000]. However, the effect of these different processes on the trace element budget of pyroxenites is not known.

1.3 An alternative model for the source of OIB

Silica deficient pyroxenite veins could be produced within continental or oceanic lithospheric mantle by percolation of magmatic liquids generated at low degrees of partial melt (low F) in upper-mantle rocks enriched in volatiles, alkali and incompatible trace elements at the base of the lithosphere [Sun & McDonough, 1989; Halliday et al., 1995; Niu et al., 1996; 1997; 2002; 2003]. As suggested by Niu & O'Hara [2003], this metasomatic process could occur as oceanic lithosphere cools and is transported away from the mid-ocean ridge. Subducted oceanic basal lithosphere enriched by metasomatic veins provides an alternative for the source of OIB. A similar hypothesis for the source of OIB, based on a comparison of OIB and MORB from the central Atlantic, was proposed by Halliday et al. [1995]. They postulated that the source region of OIB is enriched by small-degree partials melt soon after formation of the oceanic lithosphere at the periphery of plumes or at mid-ocean ridges. Metasomatism of the uppermost mantle would generate a "near-surface fractionated" (NSF) source with high U/Pb, Ce/Pb whilst maintaining Ce/U relatively unfractionated. The similarity of composition between basalt produced by modest extension in continental regions and OIB led McKenzie & O'Nions [1995] to propose that the source of OIB could be delaminated, metasomatised, continental lithosphere. As with models for metasomatism of ocean lithosphere, metasomatism of continental lithosphere implies the percolation of metasomatic melts rich in volatiles and incompatible elements into lithospheric mantle previously depleted by melt extraction [McKenzie & O'Nions,1995].

The problem with these various alternative models for the source of OIB is that they do not explain:

  1. Why there are systematic isotopic and trace element variations in the source of basalt from a single island, and
  2. How metasomatism can produce the apparent end-member mantle sources of OIB.

Moreover, the different mantle end-members (HIMU, EM1 and EM2), which are defined by their isotopic compositions, are also characterized by different trace-element ratios such Nb/Th, Ba/Nb, Th/La [Weaver, 1991; Chauvel et al., 1992]. In the recycled, ancient oceanic crust model [Hofmann & White, 1982], these elemental variations are related largely to the addition and assimilation of recycled sediment to the recycled oceanic crust [Weaver, 1991; Chauvel et al., 1992]. In contrast, no clear model for the correlation between trace-element and isotopic variations has been developed for different metasomatic alternative models.

The Cantal basalts from the Massif Central, France, have been interpreted as the product of metasomatism within the lithospheric mantle. This work provides new information that can constrain the formation of distinct compositions in lherzolitic OIB mantle sources.

2. A source for Cantal basalt involving veins plus enclosing lithospheric mantle

2.1 Trace element variations in the Cantal basalt source

The basalts erupted between 13 and 3 Ma ago in the Cantal alkali massif are similar to OIB in major and trace elements composition (Figures 1 & 2) [Wilson & Downes, 1991; Pilet et al., 2002]. They present a substantial composition range, from alkali basalts to basanites, which can be explained by a variation in the degree of partial melting of a common mantle source.

Figure 1: Total alkalis vs silica diagram for 13-9 Ma and 9-3 Ma Cantal basalts.

Figure 2: Primitive-mantle-normalised trace-element variation diagram for typical 9-3 Ma Cantal basalts. Normalisation values from McDonough [1995]. Grey field: OIB variation range. The variation in the incompatible trace element content in mafic Cantal rocks might be explained by a smaller degree of partial melting for basanites than for alkali basalt in a common source.

With respect to trace element compositions, these basalts show an unusual increase of Nb/Th and Ce/Pb ratios from 10.5 and 23 respectively during the first period of basalt eruption (13-9 Ma) to 18.7 and 44 during the last period (9-3 Ma) (Figure 3). The ranges in Nb/Th and Ce/Pb encompass practically all the variation observed in the major oceanic islands [Halliday et al., 1995; Jochum et al., 2001]. Moreover, Figure 3 indicates that the first erupted Cantal basalts have similar Nb/Th and Ce/Pb ratios as are observed in EM1, while the last are similar to HIMU.

Figure 3: Ce/Pb vs. Nb/Th diagram showing the variation between the first (13-9 Ma) and last erupted basalts (9-3 Ma) in the Cantal Massif, together with the composition of representative analyses from HIMU (Mangaia, Rurutu and St Helena islands), EM1 (Rarotonga, Tristan da Cunha, Gough and Pitcairn islands) and EM2 (Tahaa, Tutuila and Upola islands) mantle end-members. Data for these different oceanic islands were taken from the GEOROC database (

Variations in Ce/Pb and Nb/Th in mafic rocks are generally attributed to sediment assimilation in the mantle source [Weaver, 1991; Chauvel et al., 1992] or during passage through the crust. However, the basalts are homogeneous with respect to their Sr, Nd and Pb isotopic compositions (Figure 4) and do not display any correlation between Ce/Pb or Nb/Th and the Sr, Nd or Pb isotopic ratios [Pilet et al., 2002; 2004].

Figure 4: 207Pb/204Pb vs 206Pb/204Pb diagram for Cantal basalts together with a selection of oceanic basalts. Dark grey field: MORB isotopic variation; light grey fields: isotopic variations observed in different oceanic islands [Zindler & Hart, 1986].

This implies that:

  1. Variable sediment contamination of their mantle sources or crustal assimilation can be excluded by the observed Nb/Th and Ce/Pb variations, because a correlation between isotope and trace element variation would then be seen, which is not the case, and
  2. The variation in trace elements in the source reflects "short-term" mineralogical control of trace-element ratios in a vein-plus-enclosing-mantle source rather than "long term" extraction of continental crust from the mantle, sediment recycling and incomplete homogenization of ancient heterogeneities in the mantle.

This mineralogical modification of the mantle source may result from the metasomatic process "percolative fractional crystallization" (PclFC) [Harte et al., 1993], which produces veins in the upper mantle.

2.2 Mineralogical evidence for a PclFC mechanism in the Cantal basalt source

The PclFC mechanism results in clinopyroxene (CPX) xenocrysts in the Cantal basalts [Pilet et al., 2002; 2004]. These xenocrysts have a compositional range from Ti-augite with Mg# > 70 (Mg# = Mg/[Mg+Fe2+]) to green aegirine-augite with Mg# < 50 (hereafter GCPX) (Figure 5).

Figure 5: Mg# vs Zr diagram for Cantal CPX xenocrysts. Photographs present typical CPX xenocrysts from mafic (Mg# >70) to differentiated CPX compositions (Mg# <50).

The GCPX are interpreted as mantle xenocrysts because:

  1. CPX with compositions similar to Cantal GCPX is observed in mantle veins from La Palmas [Wulff-Perderson et al., 1999, Pilet et al., 2002];
  2. A fluid-inclusion study performed on GCPX from Hungarian basalts [Szabo & Bodnar, 1999] points to an upper-mantle origin;
  3. Sr, Nd, Pb and O isotopic compositions of GCPX in continental flood basalts from Yemen [Baker et al., 2000] indicate the absence of crustal contamination during GCPX crystallization and a direct relationship to an asthenosphere-derived melts;
  4. The observation of mantle xenoliths in Cantal rock where GCPX are observed excludes practically all possibility of mixing in the lower crust to explain the presence of the GCPX.
The trace element variations in the different types of CPX xenocrysts constrain the evolution of the mineralogical and chemical compositions of veins during the PclFC process (Figure 6). We observe, in Figure 6, first an increase in all trace elements with a decrease of Mg# in CPX xenocrysts (type 1 to type 2, Figure 6). This is explained by fractional crystallization and by the increase in the Kd Cpx/liq for the different elements in differentiated liquid [Dobosi & Jenner, 1999]. CPX with lower Mg# exhibit distinctive primitive mantle-normalized trace-element patterns. In the most differentiated CPX composition (type 3 in Figure 6), we observe a decrease in all trace elements except Pb and Th. The decrease in all high-field-strength elements (HFSEs) is ascribed to the fractionation and deposition of Nb-rich rutile during the PclFC process which generates the GCPX [Pilet et al., 2004]. For all other elements, these decreases are related to the formation of apatite which is frequently observed as inclusions within the xenocrysts. The increase of Pb relative to the other elements reflects differences in distribution coefficients (Kd) for Pb and La, Ce, Sr etc. in all phases that crystallized during the PclFC process (Kdmin/liq Pb < Kdmin/liq La, Ce, Sr etc. for CPX, olivine, amphibole, phlogopite etc. [Halliday et al., 1995].

Figure 6: a) Primitive-mantle-normalised trace-element compositions of CPX xenocrysts. CPX with Mg# > 70 (dark blue line); intermediate CPX composition (light blue line); differentiated CPX composition (red line). The decrease in Nb, Ta and REE from intermediate to differentiated CPX is the result of fractionation and deposition of Nb-rich oxides and apatite respectively in intermediate veins. b) Schematic metasomatised lithosphere. The evolution of CPX xenocrysts reflects the evolution of different vein compositions produced by the PclFC process (mafic, intermediate and differentiated veins). Increases in Nb and Nb/Th in intermediate veins are attributed to the presence of Nb-rich oxide (rutile). Decreases in Ce/Pb ratios in the differentiated veins reflect differences in distribution coefficients (Kds) for Ce and Pb in all phases that crystallized during the PclFC process (Kdmin/liq Ce > Kd min/liq Pb for CPX, olivine, amphibole, phlogopite, apatite etc. [Halliday et al., 1995]).

Following these observations, high- and low-Nb/Th and Ce/Pb ratios characterize the intermediate and differentiated veins respectively. Pilet et al. [2004] postulated that the variation in composition between the first and last erupted basalt is the result of evolution of vein compositions present at the depth of melting in the vein-plus-enclosing-lithospheric-mantle source (Figure 7). Basalts erupted during the first period of volcanism included differentiated veins in their source whereas later ones tapped intermediate veins.


Figure 7: Schematic model for (A) 13-9 Ma and (B) 9-3 Ma basalt genesis. (a) Basaltic melts are produced by partial melting in the low-velocity zone observed below the European lithosphere [Granet et al., 1995]. The low velocity zone is thought to be rich in alkali, trace elements and volatiles [e.g. Niu & O’Hara, 2003]. (b) Rising melts percolate along channels or pervasively through asthenospheric and then lithospheric peridotite. They undergo a PFC process, precipitating amphibole, clinopyroxene, oxide, and apatite in veins. Three main zones with different compositions are distinguished in veins: (1) mafic pyroxenite veins; (2) differentiated veins with clinopyroxene, amphibole, Nb- and Ta-rich oxide (such as rutile), ilmenite and apatite; and (3) veins extremely enriched in Th, U, and LREEs relative to Ta and Nb, which are depleted as a result of rutile fractionation. These latter veins are composed of GCPX, apatite, amphibole, ilmenite, and residual melts. Geochemical evolution of the differentiated GCPX-bearing veins is similar to that proposed by Bedini et al. [1997] to explain evolution during the PclFC process. (c) Melt percolation and vein deposition are associated with migration of a thermal front across lithospheric mantle [Lenoir et al., 2001]. Partial melting of veins plus enclosing depleted lithospheric mantle forms Cantal basaltic liquids. Because 13-9 Ma and 9-3 Ma basalts have similar major element compositions, their depth of mantle melting must have been similar [Albarède, 1992]. As PFC proceeded, vein compositions at the depth of melting evolved from type 3 during the early basalt activity to type 2 during the late basalt activity.

2.3 Influence of the enclosing lithosphere over the metasomatic process

The mineralogical and chemical constraints described above indicate that continental basalts with compositions similar to OIB could be generated by melting veins plus enclosing continental lithospheric mantle. But does this model apply to metasomatic vein formation within the oceanic lithosphere? Several lines of evidence suggest that it does:

  1. The model suggests that the trace-element composition of basalts is mainly controlled by the composition of veins in the source [Pilet et al., 2004]. This is because of the difference in the degree of trace-element enrichment in veins compared to the surrounding mantle. For example, concentrations of incompatible trace elements are 40-80 times higher in vein minerals than in lithospheric spinel peridotite minerals (e.g., GCPX have 12-40 ppm of La whereas clinopyroxene from depleted peridotite xenoliths from the Massif Central contain only 0.3-0.6 ppm of La [Zangana et al., 1999]). Thus the nature of the enclosing lithosphere does not play an essential role in the generation of OIB trace-element compositions;
  2. Cantal lithosphere has a depleted composition, which has been attributed to the extraction of basaltic liquid from an initial near-chondritic mantle source [Zangana et al., 1999; Lenoir et al., 2000; Downes et al., 2003]. Thus, before Neogene volcanic activity, Cantal lithospheric mantle was similar to typical unmetasomatised oceanic lithosphere;
  3. GCPX xenocrysts in alkali basalt are described from both oceanic and continental settings [Pilet et al., 2002]. This suggests that the PclFC process is global.

Furthermore, the PclFC process observed in Cantal basalt sources can be linked with the low-F, volatile- and incompatible-element-enriched, metasomatic agent that seems to affect the oceanic basal lithosphere [Halliday et al., 1995; Niu et al., 1996; 1997; 2002; 2003]. The similarities in the two processes are:

  1. Initial liquids derived from the PclFC mechanism must be produced at low F, i.e., the melt fractions must be sufficiently large to segregate from the peridotitic source but not so large that they are transported out of the upper mantle into the crust.
  2. The initial PclFC model liquid must have high-H2O and CO2 contents, as demonstrated by amphibole inclusions in CPX xenocrysts, carbonate in association with GCPX in xenoliths from Australian basalt [Wass, 1979] and CO2 fluid inclusions in GCPX from Hungarian basalts [Szabo & Bodnar, 1999].
  3. In both the continental and oceanic environments, the initial metasomatic model melt is generated in the low velocity zone at the base of the lithosphere [Pilet et al., 2004, Hoernle et al., 1995, Granet et al., 1995, Niu et al., 2002; 2003].

3. Evolution of metasomatised heterogeneous lithosphere

3.1 Trace element variation within metasomatised heterogeneous lithosphere

The compositional differences between Cantal basalts erupted during the two periods of volcanism provide an unusual opportunity to characterize compositional variations introduced by metasomatic processes in lithospheric mantle. As indicated in Figure 6b, which is a schematic model of metasomatized lithosphere, we distinguish upper and lower parts of this hypothetical lithosphere. The composition of the upper part is constrained by the composition of first-erupted Cantal basalt while the lower part is related to the composition of last-erupted basalt. As shown in Figure 8, there are striking differences between the two parts.

Figure 8: Trace element variations observed between the first (13-9 Ma) and the last (9-3 Ma) emitted basalts in the Cantal Massif. Diagram normalised to the average composition of the 9-3 Ma basalts, measured at 9 ppm of Th, which accounts for differences in degrees of partial melting. Light blue lines: minimum and maximum normalised value of 9-3 Ma basalt measured at 9 ppm of Th. Red line: average and confidence interval of 13-9 Ma basalt. The range of ratios measured in 13-9 Ma (30.3 - 42.8) and 9-3 Ma (38.8 - 53.0) basalts. Cantal basalts may be used to calculate ratios of 24 and 31.2 in their respective sources.

These differences are particularly important for Ba, Nb, Ta, and Pb, yielding an increase in U/Pb and Th/Pb. Smaller variations observed for Sm and Nd are associated with a decrease in Sm/Nd for early basalts relative to late ones. There is also probably a decrease in Rb/Sr from early to late basalts but this is less safe since Rb was mobile during alteration of the 13-9 Ma Cantal basalts. The decrease in Rb/Sr is inferred on the basis of (1) a decrease in Ba/Sr between the two periods of basalt emission (Ba is generally correlated with Rb in basaltic rocks [Hofmann & White, 1983]) and (2) a decrease in Sr and an increase in Rb content in the most differentiated veins. The µ ratio average is estimated to be 31.2 in the source of the 9-3 Ma basalt (with extreme values up to 37.1) and 24.0 in the source of  the 13-9 Ma basalt (with extreme values up to 23.1). Analogous estimates for 147Sm/144Nd indicate 147Sm/144Nd values of 0.144 and 0.153 for the first and last erupted basalt respectively.

3.2 Temporal evolution of the metasomatised heterogeneous lithosphere

Metasomatised lithosphere could be entrained in asthenosphere flow during subduction of the oceanic lithosphere or delamination of the continental lithosphere. It would then become a long-term heterogeneity producing isotopically enriched mantle. We can model the temporal evolution of isotopic systems in the enriched mantle.

The subduction process generates a metamorphic transformation of vein mineralogy and probably hybridization within the enclosing mantle. However, because of the very low diffusion rate at high pressure [Hofmann & Hart, 1978; Kogiso et al., 2004], chemical heterogeneity will be preserved within the subducted mantle lithosphere. Nevertheless, during these processes, the enriched component will be greatly diluted in a manner similar to that proposed for enrichment by subduction of oceanic crust [Chauvel et al., 1992], erosion/delamination of sub-continental lithosphere [McKenzie & O'Nions, 1995] or model NSF mantle recycling [Halliday et al., 1995]. It is extremely difficult to estimate the dilution factor, but the average of the enriched component calculated previously and the depleted mantle ( µ DM = 10 and 147Sm/144Nd DM = 0.21 [Rehkämper & Hofmann, 1997]) may be used for model evolution calculations.

Because metasomatised lithosphere would be heterogeneous with variable U/Pb, Th/Pb and Sm/Nd ratios, different Pb and Nd isotopic ratios will be produced in the upper and lower parts of the lithosphere. The µ ratio of the basal lithosphere would, after 1.77 Ga, have a 206Pb/204Pb isotopic ratio similar to that of the sources of HIMU basalts (Figure 9a). The absolute value of this isotopic ratio is highly dependant of the estimate of the µ ratio in the source, isolation time, dilution factor and choice of initial isotopic composition which is taken here to be similar to depleted mantle at 1.77 Ga or 1 Ga. Conversely, the difference in µ calculated between the lower and upper lithosphere (and in the corresponding last and first emitted Cantal basalt) is better constrained and varies linearly with the dilution factor or with the method used to estimate the µ value of the sources. The different values of µ produce 206Pb/204Pb variations of ~1.2 over 1.77 Ga (Figure 9a) using the average µ values, and ~2.2 using the extreme µ values for the first (16.55, average of the value calculated from Cantal basalts and µ value of DM) and last (23.55) emitted basalts. For an isolation time of 1 Ga, the 206Pb/204Pb variation range calculated is somewhat smaller than, though similar in scale to, that seen in individual oceanic islands (~0.6 for average µ values and ~1.9 for extreme µ values) (Figure 9a).

Figure 9: a) 207Pb/204Pb vs 206Pb/204Pb and b) εNd (today) vs 206Pb/204Pb diagrams for oceanic basalts. εNd (t) = (143Nd/144Nd (t,sample)/ 143Nd/144Nd (t, CHUR) -1) x 104 with 143Nd/144Nd (today, CHUR)= 0.512638. Red and blue points: present-day Pb and Nd isotopic ratios generated after 1.77 Ga and 1 Ga respectively from the lower part (full point, μ ratio: 20.6, Sm/Nd: 0.289, see text for explanation) and the upper part (empty point, μ ratio: 17.0, Sm/Nd: 0.282) of metasomatised lithosphere. Green triangles: isotopic compositions of unmetasomatised lithosphere. Arrows indicate the effect of evolution of the metasomatic agent within lithosphere over the isotopic composition after 1 Ga (blue) and 1.77 Ga (red). Dark grey field: MORB isotopic variation; light grey fields: isotopic variations observed in different oceanic islands [Zindler & Hart, 1986]. Orange, green and blue fields: Mangaia (Cook Islands) Tahaa (Society Islands) and Rarotonga (Cook Islands) isotopic compositions respectively. Data from GEOROC data base ( Orange and green arrows: 206Pb/204Pb variation range in olivine-hosted melt inclusion from Mangaia and Tahaa basalt respectively analysed by Saal et al. [1998]. Direction of arrows reflect a decrease in Cr content in CPX crystallized in melt inclusion [Saal et al., 1998]. Model calculation: unmetasomatised lithosphere is formed at 1.77 Ga or at 1 Ga with initial Pb isotopic ratios calculated using present-day 206Pb/204Pb and 207Pb/204Pb of 18 and 15.46 respectively and with μ ratio fixed to 10 [Rehkämper& Hofmann, 1997]. The Nd isotopic composition of unmetasomatised lithosphere at 1 or 1.77 Ga is calculated using a present day 143Nd/144Nd value of 0.5132 and 147Sm/144Nd ratio of 0.21. [Rehkämper & Hofmann, 1997] .

The combination of µ and Sm/Nd heterogeneities in Cantal basalt sources produces variations in 143Nd/144Nd – 206Pb/204Pb that simulate mixing between HIMU and EM components reported for OIB worldwide (Figure 9b). The inferred high Rb/Sr in the upper part of the hypothesised lithosphere relative to the lower part would produce significant 87Sr/86Sr variations with time in the derived basalts. This estimate of 87Sr/86Sr variation is in agreement with the apparent mixing trends between HIMU and EM components inferred from 206Pb/204Pb, 207Pb/204Pb and 143Nd/144Nd variations calculated from the upper and lower parts of the heterogeneous lithosphere.

Isotopic trends of Nd-Sr-Pb in oceanic basalts are correlated with trace-element ratios that are used to infer elemental signatures of the hypothetical EM and HIMU end-member mantle components [Weaver, 1991; Chauvel et al., 1992] (Figure 3, 10a, 10b). As indicated in Figure 10, the variations in trace element ratios from the first to the last period of Cantal basalt emission define a trend that mimics mixing between HIMU and an EM component.

Figure 10: a) Ba/Nb vs La/Nb and b) Th/Nb vs Th/La diagrams for Cantal basalts; 13-9 Ma (open red circles) and 9-3 Ma (filled red circles) from the upper and lower parts of metasomatised heterogeneous lithosphere. HIMU (triangles), EM1 (filled squares) and EM2 (open squares) compositions are from Weaver [1991].

These data demonstrate that fractionation of trace-element ratios (U/Pb, Th/Pb, Rb/Sr, Sm/Nd, Nb/La etc.), necessary to generate EM and HIMU isotopic characteristics over time, can be simulated by the evolution of a metasomatic agent within the lithosphere. These observations are compelling evidence that variable isotopic and trace-element-ratio signatures in the mantle sources of single oceanic islands can be produced by metasomatised and subducted oceanic lithosphere.

4. Interpretation of end-member compositions

4.1 Example from HIMU – EM variation in Polynesia basalts

Mangaia (Cook Islands) and Tahaa (Society Islands) basaltic rocks are considered to be derived from the partial melting of nearly "pure" HIMU and EM2 mantle components respectively [Hart et al., 1992]. Trace-element variations observed between basalts from these two islands (Figure 11a) are remarkably similar to the variations observed between the two periods of Cantal basalt emission (Figure 8). So, by analogy with the model proposed for Cantal basalt formation, the Mangaia and Tahaa mantle sources may have been produced during a metasomatic event within the oceanic lithosphere. In this hypothesis, Mangaia and Tahaa source compositions correspond to a heterogeneous lithosphere containing mafic-to-intermediate and differentiated veins respectively.

Data from olivine-hosted melt inclusions analyzed in both Mangaia and Tahaa basalts [Saal et al., 1998] strongly support this idea. Melt inclusions from Mangaia alone have 206Pb/204Pb, 207Pb/206 and 208Pb/206Pb variations which cover much of the HIMU to EM2 lead isotope composition range (Figure 9a). The lead isotope range of Mangaia melt inclusions is slightly larger than that deduced for the Cantal basalts from the variation of µ after an isolation time of 1.77 Ga. However, there is no reason to suggest that the lithospheric sources of Cantal basalt were the most affected by the postulated metasomatism. Tahaa melt inclusions are similar to Mangaia melt inclusions, but show a smaller Pb isotopic range corresponding to EM2 isotopic compositions.

The progressive variation from HIMU to EM2 isotopic compositions in melt inclusions is associated with an increase in Pb content and a decrease of Cr in CPX crystallized within the inclusions [Saal et al., 1998]. These relationships indicate that a differentiation process was involved in the generation of the mantle reservoirs prior to long-term evolution of the Pb isotopic system [Saal et al., 1998]. These variations are in perfect agreement with the hypothesis proposed here. The differentiation process which generates the different vein composition within the lithosphere leads to a decrease of U/Pb, Th/Pb and Cr and an increase of Pb from the bottom to the top of the oceanic lithosphere prior to subduction (Figures 6 and 8). The continuous variation of U/Pb and Th/Pb from the bottom to the top of the lithosphere will generate, after sufficient isolation time, a decrease in 206Pb /204Pb associated with an increase of 207Pb/206 and 208Pb/206Pb, which are correlated to a decrease of Cr and an increase of Pb.

This differentiation process can explain the different correlations and variations observed in Mangaia and Tahaa melt inclusions and whole rocks. This suggests that HIMU and EM2 end-member isotopic compositions correspond to the extreme trace-element fractionation generated by metasomatic processes within the lithosphere, as illustrated at Cantal. A similar metasomatic origin for the EM2 signature has been proposed previously by Workman et al. [2004] based on a study of the Samoan volcanic chain. The metasomatic interpretation of the EM2 component is also in agreement with the hypothesis of Eiler et al. [1997] to explain the origin of high δO18 observed in some EM2 rocks. Following these authors, the high δO18 values in EM2 could be linked either to addition of ~2-6% of terrigenous sediment in the mantle source or with a metasomatic process within mantle peridotite [Workman et al., 2004; Eiler et al., 1997].

EM1 isotopic compositions differ from EM2 by having lower 206Pb/204Pb, 87Sr/86Sr and 143Nd/144Nd ratios. However, the variations in trace element ratios such as Nb/Th, La/Th, Ce/Pb and Ba/Nb observed in these two end-members mainly overlap (Figure 3, 10a, 10b) [Weaver, 1991; Chauvel et al., 1992]. Rarotonga, located amongst the Cook Islands in Polynesia, and also Pitcairn, trends towards the EM1 end-member [Chauvel et al., 1992]. Comparisons between Rarotonga, Tahaa and Mangaia chemical compositions allow us to constrain the primary variations between the EM1, EM2 and HIMU end-members. In Figures 11a and 11b, the trace element variations between Mangaia-Tahaa and Mangaia-Rarotonga can be seen to have similarities.

Figure 11: a) Trace element variations observed between the Mangaia (Cook Islands) [Kogiso et al., 1997; Dupuy et al., 1989] and Tahaa (Society Islands) [White & Duncan, 1996]. b) Trace element variations observed between the Mangaia (Cook Islands) and Rarotonga (Cook Islands) [Zindler & Hart, 1986; Nakamura & Tatsumoto, 1988; Hauri & Hart, 1993; Dostal et al., 1998; Thompson et al., 2001]. Diagrams normalised to the average composition of Mangaia basalts, measured at 5.5 ppm Th. Black lines: minimum and maximum normalised value of Mangaia basalt measured at 5.5 ppm of Th. Grey line: average and confidence interval of Tahaa and Rarotonga basalt.


Rarotonga has lower 206Pb/204Pb and 87Sr/86Sr than Tahaa but 143Nd/144Nd is similar for both islands (Figures 9a, 9b). This implies a slightly lower U/Pb and Rb/Sr in the source of Rarotonga basalts than in Tahaa basalts. These variations are observed in U vs Pb and Rb vs Sr diagrams (Figures 12a, 12b) where Rarotonga basalts exhibit lower U/Pb and Rb/Sr ratios than Tahaa basalts. Following the previous interpretation of Tahaa (EM2) compositions, we suggest that the source of Rarotonga was generated by a metasomatic process similar to that which we invoke to explain the Tahaa signature, that is to say, the final stages of metasomatic melt evolution within the lithosphere. Small differences during melt differentiation could explain the small variation between Tahaa and Rarotonga sources. For example, fractionation of a small amount of phlogopite instead of amphibole during the process which generated the Rarotonga source could have decreased the Rb/Sr ratio and increased the Pb content of the differentiated liquid that formed veins in the lithosphere (KdPhlogopite/liquid is higher that Kdamphibole/liquid for Rb and lower for Sr and Pb [Halliday et al., 1995]).

Figure 12: a) U-Pb and b) Sr-Rb diagrams from Mangaia (Cook Island), Tahaa (Society Island) and Rarotonga (Cook Island) basalts which represent HIMU, EM1 and EM2 mantle end-members respectively. Data for these different oceanic islands are from GEOROC data base ( Basalt analyses are considered to be altered based on the Rb-Ba correlation defined by Hofmann & White [1983], where the two elements were analysed in the same sample or when the Rb/Sr ratio seems unrealistic.

Rarotonga basalts do not present the low 143Nd/144Nd observed in Pitcairn basalts which are considered to be the melt products of a near "pure" EM1 source. A hypothesis for the decrease of Sm/Nd in the source of Pitcairn basalts is the fractionation of garnet during the metasomatic process which generates the differentiated veins (KdGarnet/liquidSm > KdGarnet/liquidNd). Phlogopite and garnet represent high pressure phases in alkaline rocks. Thus, we propose that metasomatic formation of the EM1 isotopic composition, rather than EM2, is linked to the thickness of the lithosphere – thick for the formation of EM1 (continental lithosphere?) and thin for EM2 (oceanic lithosphere?). However, more studies are necessary to investigate further these last suggestions.

4.2 Variation of high field strength elements in HIMU and non-HIMU ocean island basalt

In the subducted oceanic crust hypothesis for the source of OIB, enrichments in Ba, Rb, Th, U and La relative to Nb and Ta in the EM1 and EM2 end-members, compared to the HIMU source (Figures 10a and 10b), is ascribed to recycling of pelagic (EM1) and terrigenous (EM2) sediments into the mantle [Weaver, 1991; Chauvel et al., 1992]. Following Kogiso et al. [1997], this hypothesis is problematic for HIMU and other OIBs in southern Polynesia. The major element characteristics of those OIB magmas do not support significant involvement of basaltic components in non-HIMU sources; and the Nb enrichment of HIMU basalts relative to other incompatible elements is difficult to explain simply by the addition of a crustal component to the source of HIMU rocks [Kogiso et al., 1997]. These inconsistencies are solved by the metasomatic hypothesis presented here, where the high Nb/Th in HIMU is explained by the presence of a small amount of Ti-rich oxide in metasomatic veins. Ti-oxides have large distribution coefficients (Kd) for HFSEs. Thus, low La/Nb, Ba/Nb, Th/Nb and high Hf/Lu in the HIMU source is explained by precipitation of Ti-oxide during modal metasomatism, preferentially enriching Nb and Hf relative to other incompatible elements in the lithospheric source. Hf and Nd isotope ratios in non-HIMU OIB exhibit a strong positive correlation (Figure 13), indicating a control by magmatic processes [Salter & White, 1998; Niu & O'Hara, 2003]. HIMU basalts differ from other basalts falling on the main OIB array in that they have a low 176Hf/177Hf value for a given 143Nd/144Nd, reflecting a larger Hf/Lu ratio in their source. Even if recycled subducted oceanic crust can duplicate the Hf-Nd isotopic composition observed in HIMU basalts, mixing the subducted residuum with sediments in the mantle does not produce the main 176Hf/177Hf - 143Nd/144Nd array for non-HIMU OIBs [Salter & White, 1998]. Enrichment in Hf relative to Lu in the HIMU source by modal metasomatism before the evolution of the isotopic system explains the Hf-Nd isotopic variation in all OIB sources more simply (Figure 13).

Figure 13: εHf vs εNd diagram for oceanic basalts. εHf = (176Hf/177Hf(t,sample)/176Hf/177Hf(t,CH) -1) x104 with 176Hf/177HfCH,today =0.282772 [Blichert-Toft & Albaréde, 1997]. Filled grey circles: non-HIMU-OIB compositions; open circles: HIMU-OIB compositions; data from Salters & White [1998]. Open squares: evolution of a hypothetical OIB mantle source after 2 Ga from an initial Bulk Silicate Earth isotopic composition with 176Lu/177Hf and 147Sm/144Nd of 0.0381 and 0.2115 respectively. Filled squares: model mantle as in open squares except with the addition of 0.01% of rutile. Hf concentration in rutile = 340 ppm [Kalfoun et al., 2002]. Arrow shows the effect on εHf (decrease of ~3.0) in non-HIMU basalts by the addition of 0.01% of rutile. This decrease is equivalent to the observed difference in εHf between HIMU and non-HIMU basalts.

5. Summary

5.1 General

  1. Variations in the chemical compositions of Cantal basalts are interpreted as the result of a metasomatic agent within the lithosphere. The Cantal example demonstrates that basalts characterized by OIB chemical compositions could be generated by partial melting of metasomatic veins plus enclosing lithospheric mantle. Such a heterogeneous source may form within either the continental or oceanic lithosphere.
  2. The evolution of the metasomatic agent could generate trace element variations (U/Pb, Th/Pb, Rb/Sr, Sm/Nd, Nb/La etc.) in different parts of the lithosphere. Isolation of this metasomatic lithosphere would generate isotopic and trace element ratio variations that are similar to the variations observed in oceanic islands. These observations are compelling evidence that variable isotopic and trace element ratio signatures in the mantle sources of single oceanic islands can be produced by subduction of metasomatised oceanic lithosphere or delamination of metasomatised continental lithosphere.
  3. The similarity between the trace element variation observed in Cantal basalt and the variation observe between Mangaia (HIMU) and Tahaa (EM2) - Rarotonga (EM1) islands suggest that model HIMU and EM end-member isotopic and chemical compositions could correspond to the extreme trace-element fractionation generated by metasomatic process within the lithosphere.

5.2 Implications for the existence of plumes

  1. The new interpretation of OIB sources proposed here requires re-evaluation of the processes that control the chemical evolution of Earth's mantle reservoirs and their chemical and isotopic indicators.
  2. The metasomatic hypothesis is entirely consistent with the isotopic heterogeneity observed in the source of oceanic basalts and experimental data which indicate that the most plausible source of OIB material are pyroxenites [Hirschmann et al., 2003; Kogiso et al., 2003]. This model is particularly interesting for the formation of high-silica undersaturated rocks as nephelinite or basanite where the common interpretation as recycled oceanic crust clearly fails to explain their major element composition.
  3. Recycled metasomatised lherzolite may be the major mantle source sampled by OIBs; recycled oceanic crust and sediment may be less common in OIB sources than is usually assumed.
  4. The addition of volatiles, in particular of CO2, to lherzolite significantly decreases the solidus temperature. The metasomatic model for the origin of OIB suggests that OIB could be produce at a lower temperature than is commonly assumed in the recycled oceanic crust model.


We are grateful for long discussions with Mike O'Hara and Yaoling Niu. We also thank D. Fontignie, A. Potrel, M. Poujol, J. Kramers, B. Villemant, M. Tubrett and L. Hewa for help with TIMS, INAA, MC-ICP-MS and laser ablation ICP-MS analyses.


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last updated 20th February, 2005