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Two main processes are invoked for generating melts in subduction zones. One process involves melting in the mantle wedge due to the introduction of hydrous fluids released by the dehydration of the subducted slab (Gill, 1 98 1 ; Plank and Langmuir, 1 988; Davis and Bickle, 199 1 ; Hawkesworth et aI., 1991 ; Arculus, 1994; Pearce and Peate, 1 995). The second process is decompression melting due to comer flow and/or extension (Luhr, 1997; Sisson and Bronto, 1 998; Elkins-Tanton et aI., 2001). However, both these processes imply a "steady-state" situation irivolving active plate subduction (eg. Davies and Stevenson, 1 992), where, provided the subducted oceanic slab enters the asthenosphere at velocities > 6cmiyr, the mantle wedge will be hot enough to generate arc melts either through decompression, or hydrous fluxing. Importantly, in such models the subduction-

derived component can remain metastable, or cryptic in the form of vein-glasses, or as melt and/or fluid residing in the intergranular interfaces or as fluid inclusions in primary minerals (Kepezhinskas and Defant, 1 996).

In Section 2.3 it was noted that, with very few exceptions, shoshonitic magmatism rarely occurs in direct association with active or "steady-state" arc-settings, in contrast, shoshonitic magmatism typically occurs long after the cessation of active subduction and often in association with arc-rifting and/or postcollisional settings. In other words, a lack of active plate subduction would appear to be the rule rather than the exception regarding the tectonic and geodynamic association of shoshonitic magmatism. This observation has important implications, since the subduction­ derived metasomatic components (fluid and/or melt) are unlikely to remain in the mantle environment for extended periods (geological time scales) in the form of cryptic and metastable melts or fluids (Dawson, 1984). Accordingly, the near­ ubiquitous occurrence of shoshonitic magmatism, long after the cessation of subduction, ego Fiji (see Chapter 4), or in more extreme examples such as the Tibetan Plateau (eg. Turner et aI., 1996a, Section 2.3), requires that the metasomatic components reside in the non-convecting upper mantle in the form of hydrous K­ bearing and incompatible element-rich phases such as phlogopite, K-richterite, apatite and possibly carbonate (eg. Dawson, 1984; Vidal et aI., 1989, Foley 1992a, 1992b; McInnes and Cameron, 1994; Franz et aI., 2002). Hesse and Grove (2003) note tpat there are three main reasons to postulate a veined mantle rather than a more homogeneously enriched mantle:

1) the existence of veined mantle xenoliths (Irving, 1980; Vidal et aI., 1989; Obrien et aI., 1 99 1 ; Gregoire et aI., 200 1 ; Franz et aI., 2002);

2001) or as solitary waves (Nakayama and Mason, 1999) or by pervasive porous flow (Watson and Brenan, 1 987).

The general melting mechanism envisioned for the genesis of Fijian shoshonitic magmas (and arguably most shoshonitic magmas in general) follows from the vein­ plus-wall-rock melting mechanism proposed by Foley (1992b). The model of Foley (1992b) proposes that potassic melts originate from melting of a veined lithosphere. The veins are rich in clinopyroxene and mica

(±

apatite and carbonate), whereas the wall rocks consist of peridotite phases, predominantly olivine and orthopyroxene. The veins originate by solidification of low-degree melts (McKenzie, 1989), or from reaction between peridotite phases with melt and/or fluid, which in the case of shoshonitic magmatism, are ultimately derived from the subducted slab (McInnes and Cameron, 1994; Franz et aI., 2002). According to the model of Foley (1 992b), the melting event producing potassic (shoshonitic) magma begins in the metasomatic vein assemblage due to the concentration of hydrous phases and incompatible elements and then spreads to include the surrounding wall rocks by a combination of two mechanisms.

The first mechanism involves solid solution melting. Minerals that are likely to be abundant in the vein assemblage such as CrlAI spinel, FIOH mica, amphibole and apatite, form extensive solid solutions (Foley, 1992b). The breakdown of these phases Jakes place over a temperature range between the solidus of the vein assemblage and the elimination of more refractory end-members. Foley (1 992b) suggested that this process spans the temperature gap between the solidi of vein and wall rock, so that a melt component from the wall rock is added to that from the vein before elimination of all vein minerals. Phlogopite notably forms the most effective of these solid solution reactions, from OH-phlogopite at low temperature to more refractory F-phlogopite at higher temperature (Otoyoshi and Ensen, 2001). T4e second mechanism involves the dissolution of wall rock minerals, whereby the initial vein melt fraction infiltrates the surrounding wall rock due to the dominance of surface energy minimisation on melt flow at the intergranular scale (Waff and Bulau, 1982). Following infiltration, dissolution of wall rock minerals occurs at temperatures lower than their respective melting temperatures, thus imparting a 306

refractory wall rock component to the melt composition. Dissolution of olivine and/or orthopyroxene occurs preferentially, since these minerals are' furthest from equilibrium with the strongly alkaline, vein-derived melt (Foley, 1 992b). The alkaline (shoshonitic) magma composition is thus a hybrid of vein and wall rock components.

The relationship of shoshonitic magmatism in Fiji, with regard to regional geodynamics will be discussed in the following section. However, there are a number of lines of evidence to suggest that the melting of the non-convective metasomatised (veined) sub-Fijian mantle occurred in response to conductive heating associated with vigorous convection of the underlying asthenosphere. In this sense the model proposed here for Fijian shoshonitic magmatism is notably similar to that suggested by Turner et aI. (1996a) to account for shoshonitic magmatism on Tibetan Plateau and Turner et aI. (1996b) regarding the generation of flood-basalts from conductive heating during plume-lithosphere interaction.

The melting of the metasomatic vein assemblages in non-convecting sub-Fijian mantle is inferred to be a progressive "lithospheric-unroofing" type of process. In this context, conductive heating at the base of the metasomatised mantle leads to melting of hydrous phases in the vein assemblage (eg. Foley, 1 992b). The vein-melt fraction then migrates upward re-equilibrating with the upper mantle dissolving perid�tite phases (olivine and orthopyroxene) and diluting its H20 content (Foley, 1992b; Grove et aI., 2002; Hesse and Grove 2003) until it reaches its solidus in the cooler regions in the upper mantle where it freezes. Subsequent heat input remobilises the hybrid (vein

+

wall rock) melt and the melt continues to migrate upward, dissolving peridotite phases before freezing again at higher level. According to Hesse and Grove (2003), such a melt-migration-re-equilibration

review), noted that their experimentally determined liquidus phase relations (olivine

+

orthopyroxene) are not consistent with the presence of non-peri doti tic veins at the depth of last equilibration. They suggested that their results imply that the shoshonitic component resided in the melt, at the point, or moment of last equilibration prior to melt segregation. The phases present at the point of multiple saturation (olivine

+

orthopyroxene) are therefore an indication of the phases present at the depth of last equilibration rather than the residue of the original melting process at greater depth. Accordingly, the experimental results of Hesse and Grove (2003) should be interpreted in terms of a batch melt that records the pressure of last equilibration and an apparent melt fraction at that depth. The major element oxides (Si02, MgO, FeO, CaO and Ah03), derived from the wall rock component contain information regarding the bulk composition at that pressure of last equilibration. In contrast, the alkali and incompatible trace element content of the melt, dominated by the (now diluted) original vein component provides information concerning the original metasomatised source (Foley, 1 992b; Hesse and Grove, 2003). The enriched vein component likely originates from melting of an

modally metasomatised (ie. phlogopite

±

apatite

±

carbonate-bearing) vein assemblage at depth (>2GPa, ie. >60km) (Wendlandt and Eggler, 1980; McInnes and Cameron, 1994). As the melt is progressively remobilised and migrates to shallower levels, the alkali-enriched vein component is likely to become increasingly metastable (ie. cryptic and residing in the melt) due to the punctuated nature �f the melt-migration re-equilibration process (eg. Hesse and Grove, 2003; Spiegelman and Keleman, 2002).

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