6.5       Identification of mantle components

 

Since the study of Hart et al. (1986), major efforts have been devoted to identifying the proposed mantle components in geological terms, and explaining how they have interacted to generate OIB sources. To a large extent the debate has been polarised between those who invoke metasomatic enrichment models (e.g. Hart et al., 1986) and those who invoke crustal recycling models (e.g. Weaver, 1991) to explain the enriched components. Some of the arguments will be briefly examined for the different end-members.

 

 

6.5.1    HIMU

 

Many authors have proposed that HIMU represents subducted oceanic crust (e.g. Chase, 1981; Palacz and Saunders, 1986; Staudigel et al., 1991; Chauvel et al., 1992; Hauri et al., 1993). The great advantage of this model is that it attributes HIMU to a known major subducted component.

 

            The U/Pb ratios of normal MORB are not high enough to explain the composition of HIMU, but various means have been proposed to raise the U/Pb ratio of subducted oceanic crust (section 6.3.2). An alternative site for possible U/Pb enrichment in the subducted slab is the sub-oceanic lithosphere. For example, Halliday et al. (1990; 1992) argued that shallow Pb)Pb isotope arrays in the Cameroon Line volcanics and other Atlantic islands were best explained by recent strong U/Pb enrichment of the oceanic lithosphere (section 7.3.2). This process cannot directly explain the much steeper correlation between 206Pb and 207Pb in HIMU islands; however, after storage for ca. 1 Byr, and mixing with less radiogenic Pb from other parts of the subducted slab, this represents an additional mechanism to generate the HIMU component.

 

            This type of model has been adapted by several workers (e.g. Thirlwall, 1997) to suggest that OIB signatures with radiogenic Pb that are less extreme than the HIMU end-member could be attributed to ‘young HIMU mantle’, rather than to mixing with the specific HIMU reservoir seen at St Helena and Mangaia. However, it is worth noting a possible corollary of this model, by which the HIMU end-member itself was created as a special case within this general model. Perhaps HIMU was formed by subduction of very U-enriched oceanic crust about 2 Byr ago, reflecting the sudden release of uranium to the oceans in response to changing atmospheric conditions (section 6.3.3).

 

            A rather different concept of the origin of some mantle plumes was suggested by Class et al. (1993; 1996), who proposed that plumes such as Ninetyeast–Kerguelen and Tristan–Walvis could undergo in situ growth of radiogenic Pb. This model was based on an observed correlation of Pb isotope ratio with age along the Ninetyeast Ridge, increasing from an unradiogenic composition similar to the Rajmahal traps of eastern India (ca. 120 Myr ago) to a composition at the radiogenic Pb end of the Heard Island array (at the present day). This trend is shown by the bold line in Fig. 6.36, and has been dubbed the ‘evolving plume’ model. Class et al. envisaged that U/Pb enrichment occurred when sub-continental lithosphere was delaminated and recycled back into the convecting mantle, where it would reside for less than 1 Byr at a mantle boundary layer. New plumes would then be generated from this material, either at the 760 km discontinuity or the core–mantle boundary.

 

            Helium isotope data for Heard Island have indeed provided evidence for a plume signature at this island (section 11.1.5). However, isotope data for lavas of different ages in the Kerguelen islands appear to define an evolution line in the opposite direction to the ‘evolving plume’ trend (Fig. 6.36). This led Frey and Weiss (1995; 1996) to argue that the Evolving Plume model could not explain the Kerguelen data, and could only partially explain the Ninetyeast Ridge data. Since that time a considerable amount of additional work has been done on the Kerguelen Plateau, as well as other features near the SE Indian Ridge. The picture that has arisen is very complex, with evidence for contamination by upper crust, lower crust, or lithospheric mantle in some localities, as well as evidence for a heterogeneous plume composition (e.g. Neal et al., 2002; Mattielli et al., 2002). Hence the overall consensus seems to be that the evolving plume model is too simple to explain all the observed variability. Possibly, the unradiogenic Pb signatures in older Ninetyeast basalts were due to lithospheric contamination, whereas several different plume components with varying radiogenic Pb signatures are now involved.

Fig. 6.36. Plot of Sr–Pb isotope signatures for volcanic units of the Kerguelen plateau ( " ) and Heard Island ( ! ) relative to an ‘evolving plume’ model based on Ninetyeast Ridge data. Numbers indicate approximate ages of volcanism in Myr. After Class et al. (1996).

 

 

6.5.2    EM II

 

The case for EMII as subducted continental material is almost universally agreed, since this end-member is squarely located on mixing lines between depleted mantle and marine sediments. This model was further strengthened by evidence from peridotite xenoliths in Samoan lavas (Hauri et al., 1993). Trace element data for these xenoliths point to an origin from carbonate-rich melts within the Samoan plume, and the isotopic compositions of the xenoliths are therefore taken as indicative of the EMII mantle component. These xenoliths extend the EMII array directly into the field of marine sediments (Fig. 6.37) and thus provide a compelling case for this material as the source of the EMII component. Similar xenoliths from Tubuai also support the concept of a discrete HIMU component, as previously observed in lavas from Mangaia, Tubuai, and the nearby Macdonald seamount chain.

Fig. 6.37.Plot of  Sr versus Pb isotope data for cpx grains ( ! ) and glass inclusions ( * ) in peridotite xenoliths from Savaii (Samoa) and Tubuai (Austral Islands), indicating affinity with the EMII and HIMU mantle end-members. After Hauri et al. (1993).

 

            Further evidence for sediment recycling into the EMII source comes from oxygen isotope measurements, which have consistently revealed elevated signatures in these plumes relative to MORB values. Other enriched mantle reservoirs have also appeared in the past to display oxygen isotope variations outside the range of MORB values, based on the analysis of whole-rock basalts or basaltic glasses (see review by Harmon and Hoefs, 1995). However, more recent analysis of olivine phenocrysts from a variety of plume sources showed a much more restricted range (Eiler et al., 1997), suggesting that most of the earlier variations were due to shallow contamination effects, either at the magmatic stage by oceanic crust, or under sub-solidus conditions after eruption. In contrast, phenocryst analyses from EMII plumes continue to show elevated oxygen isotope ratios correlated with 87Sr/86Sr, and are therefore considered to be a strong indicator of sediment recycling.

 

 

6.5.3    EM I

 

In their early synthesis on the nature of enriched mantle sources, Hart et al. (1986) argued that both HIMU and EMI were derived from recycled sub-continental lithosphere. More recent work has supported the argument that these two components are very closely related (e.g. Chauvel et al., 1992), but with the increasing acceptance of oceanic crust/lithosphere as the origin of HIMU, a sub-continental origin for EMI became more problematical. Dickin (1995) attempted to resurrect this model by advocating the juxtaposition of oceanic and continental lithosphere in the plume source by subduction erosion. However, since that time there has been increasing evidence to support the suggestion of Weaver (1991) that subducted pelagic sediment is an important component in the EMI source (in contrast to EMII which is attributed to recycled terrigenous sediment).

 

            Some of this evidence comes from a re-evaluation of Hf isotope data, which were earlier believed to militate against the recycling of variable types of sediment into the plume source (section 9.2.5). Other evidence comes from stable isotope analysis. However, this has had a checkered history.

 

            Woodhead et al. (1993) re-initiated interest in this problem when they found significantly elevated oxygen isotope signatures in submarine glasses from the Pitcairn seamounts, which display the most extreme EMI signatures for several isotope systems. Unfortunately there are many ways in which oxygen isotope ratios can be perturbed by sea floor processes to yield spurious signals. Nevertheless, Woodhead et al. argued that none of these processes could explain their data, which they attributed to isotopic variations in the plume source itself. Hence, their preferred explanation for this effect was the recycling of marine sediment into the EMI source by subduction.

 

            Eiler et al. (1995) tested these findings by analysing the oxygen isotope compositions of olivine and plagioclase phenocrysts from Pitcairn Island. All olivines had * 18O values close to 5.2, whereas plagioclase had * 18O values near 6.1, consistent with mass fractionation effects at magmatic temperatures. Therefore, it was concluded that the variations seen by Woodhead et al. are almost certainly due to local contamination of the glasses, either before or after solidification. Local contamination effects are also believed to explain oxygen isotope data from Iceland that lie well below the range of MORB values, although less extreme oxygen signatures could originate from a plume source (Harmon and Hoefs, 1995).

 

            In contrast to these negative results, more recent analysis of olivine phenocrysts from Hawaiian lavas has provided the first strong evidence in support of sediment contamination of the EMI source (Eiler et al., 1996). In this study, oxygen isotope ratios were correlated with lithophile isotope tracers such as Nd, in addition to osmium and helium, consistent with the involvement of three end-members in the Hawaiian plume (Fig. 6.38). The Loihi component, with an oxygen signature similar to MORB, has helium and osmium signatures indicative of a core component (section 8.3.5). The Kea component, with a depleted * 18O signature relative to MORB, is attributed to melting of recycled oceanic lithosphere. Finally, the Koolau component has an enriched * 18O signature, indicative of a component of recycled sediment, combined with unradiogenic Pb and Nd signatures characteristic of the EMI source. Additional evidence for this model comes from Hf isotope data (section 9.2.5).

 

            This strong evidence for a pelagic sediment signature in the EMI source is problematical for the alternative origin of EMI in the sub-continental lithosphere. However, the two models may be reconciled by proposing that the lithospheric mantle wedge above subduction zones takes on the isotopic signature of pelagic sediment by enrichment with magmatic and metasomatic fluids derived from pelagic sediments in the subducting slab (section 6.6.2).

Fig. 6.38. Plots of  , Nd and He isotope data (R/RA) against * 18O for olivine phenocrysts in lavas from selected Hawaiian volcanos: ( <> ) = Koolau–Lanai; ( o ) = Mauna Kea; ( ! ) = Loihi. Error bars represent suites where different tracers were determined on different samples, so the symbol shows the mean and standard deviation of the suite. Modified after Eiler et al. (1996).

 

            A final note on EMI concerns basalts dredged from the Afanasy–Nikitin rise (at the southern end of the non-seismic 85o East Ridge) which may extend EMI to even more enriched compositions. Two samples analysed by Mahoney et al. (1996) were colinear with an extension of the Pitcairn seamount array in Nd–Pb, Sr–Pb and Pb–Pb isotope space. In addition, nine analyses quoted from a Russian source (Sushchevskaya et al., 1996) define an array which is colinear with the data of Mahoney et al. (and Pitcairn data) in Nd–Pb and Sr–Pb isotope space. The Russian data are not colinear on a Pb–Pb isochron diagram (not shown), but this could be due to analytical fractionation effects. Mahoney et al. ruled out in situ lithospheric mantle as a source for the A–N data, because at the time of volcanism the A–N rise was situated on young oceanic crust, far from any continent. However, the data may not represent a lower mantle source composition because they are quite distinct from the Crozet hot spot, to which the 85o E ridge and A–N rise are attributed. Hence, Mahoney et al. speculated that the A–N component may represent lithospheric mantle material entrained for a brief time into the Crozet plume.

 

 

6.5.4.   Kinematic models for mantle recycling

 

If HIMU, EMII and EMI are attributed to recycling of oceanic crust, continental sediment, and mantle wedge plus pelagic sediment, respectively, a simple plate tectonic model can explain recycling of these components into the deep mantle in two conjugate pairs: EMI)HIMU and EMII)HIMU. This is based on the two different tectonic settings of subduction zones recognised by Uyeda (1982). The Chilean-type setting (Fig. 6.39a) is characterised by a compressional stress regime across the arc)trench gap. This causes tectonic erosion of the underside of the arc lithosphere, which may then give rise to a composite sheet in the downgoing slab consisting of oceanic crust overlain by lithospheric mantle. If the mantle wedge carries an EMI signature introduced from subducted pelagic sediment, the combination with subducted oceanic crust can generate the conjugate pair HIMU)EMI in the OIB source.

 

            In contrast to the Chilean type, the Mariana-type setting (Fig. 6.39b) is characterised by a tensional stress regime across the arc)trench gap. This causes subsidence of the trench bottom so that ocean-floor sediments are efficiently subducted, but tectonic erosion of sub-arc lithospheric mantle does not occur. The results of this process can be seen in the Lesser Antilles arc (section 6.7). The north end of the arc (with a low sediment supply) subducts barren oceanic crust, which can form a pure HIMU component after storage in the lower mantle. In contrast, the south end of the arc (with a large sediment supply) subducts a composite sheet of oceanic crust and continental sediment, which can generate the conjugate pair HIMU)EMII in the lower mantle. Isotopic evidence for these alternative processes in subduction zones will be examined in section 6.7.

Fig. 6.39. Schematic illustrations of two different tectonic styles at subduction zones which may generate conjugate pairs of enriched mantle signatures: (a) HIMU)EMI; and (b) HIMU)EMII. Modified after Uyeda (1982).

 

            The subduction of oceanic crust, together with marine sediment or eroded sub-continental lithosphere, may give rise to large-scale isotopic structure in the mantle. For example, Hart (1984) argued that recycling into the asthenosphere was responsible for generating a Pb and Sr isotope anomaly of global scale which he observed to form a small circle of approximately constant latitude encircling the southern hemisphere. He named it the ‘Dupal’ anomaly because its characteristic signature was first described in Indian Ocean volcanics by Dupre and Allegre (1983).

 

            Hart quantified the Dupal anomaly in terms of its deviation from the typical Pb isotope signatures of MORB and OIB in the northern hemisphere, which form a series of coherent Pb/Pb arrays (Fig. 6.21). Hence, he defined the terms ) 207/204 and ) 208/204 as per mil deviations of Pb isotope ratio from a ‘Northern Hemisphere Reference Line’. Not only are the southern tropics characterised by the Dupal anomaly with a positive ) 208/204 value, but a HIMU source with negative ) 208/204 is also seen in the same area.

 

            Hart (1988) and Castillo (1988) argued that the configuration of these anomalies was an indicator of the convective structure of the deep mantle. Staudigel et al. (1991) further suggested that large-scale regional isotope signatures such as the Dupal anomaly could be explained by ‘focussed subduction’ from a group of destructive plate margins, such as are presently seen in SE Asia.

 

 

6.5.5    Depleted OIB sources

 

When Zindler and Hart (1986) integrated enriched mantle signatures into a coherent model, and proposed the three well-known end-members discussed above, they also recognised that many OIBs were intermediate in composition between the end-members. However, rather than suggest that all such plumes represent complex mixing between all three end-members, they suggested the existence of a large reservoir of intermediate composition, which they called PREvalent MAntle (PREMA). Zindler and Hart also noticed that the group of islands exhibiting this signature included Hawaii and Iceland, with enriched helium isotope signatures, but with lithophile isotope signatures depleted relative to Bulk Earth. As a result of these observations, Zindler and Hart suggested two possible alternative origins for PREMA. It could either be a result of mixing of all the other mantle sources, or it could be a kind of depleted mantle formed early in Earth history, before two-layered mantle convection established the existence of the MORB reservoir.

 

            The increased amount of OIB data that became available over subsequent years led Hart et al. (1992) to argue that the most common feature in the isotope data from plume sources was a tendency to form linear arrays, which appeared fan out from a ‘Focus Zone’ at the base of the mantle tetrahedron towards a variety of enriched mantle end-members. Hence Hart et al. named this common component FOZO. The proposed composition of FOZO was on the edge of the tetrahedron between the depleted mantle (DMM) and HIMU. However, it was clearly distinct from DMM. Therefore, Hart et al. proposed that it was a lower mantle component that had been entrained around enriched mantle plumes rising from the core–mantle boundary.

 

            Subsequent work showed that the location of FOZO as originally proposed was not satisfactory, since several island arrays (e.g. the Macdonald Seamounts) trended from HIMU or DMM towards the middle of the mantle tetrahedron. Therefore, Hauri et al. (1994) revised the concept of FOZO to a somewhat less depleted signature bearing a very strong resemblance to PREMA (Fig. 6.40). Again, this component was also identified with elevated levels of 3He. In addition, fluid dynamic modeling by Hauri et al. suggested that the extent of lower mantle  entrainment into plumes was very variable (between 5% and 90%).

Fig. 6.40. View of the mantle tetrahedron, showing OIB arrays that converge on the revised composition of FOZO from many different directions. After Hauri et al. (1994).       

 

            Hauri et al. (1994) examined two alternative models for mixing between FOZO and enriched mantle in plumes. Since these plumes were argued to originate at the core–mantle boundary layer (CMBL), it was argued that one of the two end-members must be located at this point. Therefore, plumes either originate as enriched mantle at the CMBL and entrain FOZO as a kind of sheath, or they must originate as FOZO at the CMBL and entrain enriched blobs of lower mantle.

 

            These and other alternative models were recently reviewed by van Keken et al. (2002), who re-examined the correlation between helium and lithophile isotope systems. They plotted 3He data against a composite index of Sr and Pb isotope ratios (Fig. 6.41). Several ocean islands exhibit 3He enrichment, but the degree of correlation between helium and lithophile isotope tracers is generally quite weak (except for Samoa, where helium variations are attributed to shallow mixing processes; see section 11.1). Therefore, the evidence in Fig. 6.41 suggests that the 3He-enriched component is largely decoupled from and independent of the lithophile isotope signature of FOZO/PREMA.

 

            The simplest solution to the problem of juxtaposing these sources would be that elevated 186Os and 3He signatures come from the core and then become one of a suite of distinct components at the CMBL. These components then all undergo variable mixing with FOZO/PREMA during their ascent and entrainment in the mantle convection system, as originally proposed. This would imply that FOZO/PREMA is dispersed in the lower mantle between blobs of enriched material. In this case it can most easily be produced by recycling of oceanic lithospheric mantle, in contrast to the enriched components, which represent recycled crust or sub-continental lithosphere.

Fig. 6.41. Plot of helium against a composite index of Sr and Pb isotope data.

 

            Detailed examination of individual hot-spots may offer the best hope of determining the origins of ‘non enriched’ lower mantle sources such as FOZO, and the ways in which these sources are juxtaposed with enriched sources in rising plumes. For these purposes, the Iceland plume has been of particular interest.

 

            Thirlwall et al. (1994) showed that Icelandic basalts and North Atlantic MORB display sub-parallel but distinct Pb isotope arrays. Hence, they argued that isotopic heterogeneity in the Iceland plume could not be caused by mixing with MORB. They speculated that the Iceland plume array might represent a ‘young HIMU’ type source. On the other hand, Kerr et al. (1995) suggested that the Iceland plume was formed from recycled ‘lower oceanic lithosphere’. This would generate a more depleted signature than recycled oceanic crust (the proposed source of HIMU). A component similar to this was proposed by Hanan and Graham (1996) as a ‘common’ mantle component, ‘C’. This resembles FOZO/PREMA in some respects, since it has effectively the same Pb isotope ratio. However, the component C was envisaged as dispersed within the MORB reservoir. It could therefore represent FOZO/PREMA material in a ‘marble cake’ mantle.

 

 

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