Journal of Petrology | Volume 43 | Number 10 | Pages 1857-1883 | 2002
© Oxford University Press 2002
Plume-Associated Ultramafic Magmas of Phanerozoic Age
1DEPARTMENT OF GEOLOGICAL SCIENCES, RUTGERS UNIVERSITY, NEW BRUNSWICK, NJ 08903, USA
2DEPARTMENT OF EARTH SCIENCES, CARDIFF UNIVERSITY, PO BOX 914, CARDIFF CF10 3YE, UK
Received February 23, 2001; Revised typescript accepted April 1, 2002
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONA HYBRID FORWARD AND...MAGMAS IN PROJECTIONCALCULATED MAGMA COMPOSITIONS:...INVERSE MODELSGORGONA KOMATIITES: AN EXAMPLE...HAWAII (KILAUEA)TERTIARY PICRITES AND KOMATIITES...MID-OCEAN RIDGE BASALTSPLUME THERMAL AND PETROLOGICAL...CONCLUSIONSREFERENCES |
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A parameterization of experimental data in the 0·2–7·0 GPa pressure range constrains both forward models of potential primary magma compositions that exit the melting regime in the mantle and inverse models for computing the effects of olivine fractionation for any olivine-phyric lava suite. This is used to infer the MgO contents of primary magmas from Gorgona, Hawaii, Baffin Island and West Greenland. They typically contain 18–20% MgO for wide variations in assumed peridotite source compositions, but MgO can drop to 14–17% for Fe-enriched sources, and increase to 24–26% for fractional melts from Gorgona. Primary magmas with 18–20% MgO have potential temperatures of 1520–1570°C. For Gorgona picrites with 24% MgO, the potential temperature and initial melting pressure were about 1700°C and 8·0 GPa, respectively; melting was hot and deep, consistent with the plume model. There are important restrictions to magma mixing in mantle plumes. Primary magmas that exit the melting regime are both well-mixed aggregate fractional melts and isolated fractional melts. The latter can originate from a hot plume axis and be in equilibrium with olivines having mg-numbers of 93·0–93·6, but they have MgO contents and thermal characteristics that are difficult to constrain.
KEY WORDS: komatiite; picrite; basalt; MORB; olivine; mantle plumes; primary magmas; equilibrium melting; accumulated fractional melting
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONA HYBRID FORWARD AND...MAGMAS IN PROJECTIONCALCULATED MAGMA COMPOSITIONS:...INVERSE MODELSGORGONA KOMATIITES: AN EXAMPLE...HAWAII (KILAUEA)TERTIARY PICRITES AND KOMATIITES...MID-OCEAN RIDGE BASALTSPLUME THERMAL AND PETROLOGICAL...CONCLUSIONSREFERENCES |
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Mantle plumes are localized regions of hot mantle that are thought to offer ideal opportunities for ultramafic magmatism during the Archaean (Cawthorn, 1975; Fyfe, 1978; Jarvis & Campbell, 1983; Arndt, 1986). However, the importance of ultramafic magmatism is often diminished in models of Phanerozoic hotspot volcanism because basalt eruptives are more common, and the plume model is subject to debate (Anderson, 2000). Even where ultramafic lavas occur, they often have a large olivine phenocryst content and this complicates interpretations of the MgO content of the primary magma. Indeed, the interpretation that picrites can only form by the accumulation of olivine in basaltic liquids (Bowen, 1928) went unchallenged until petrographic evidence demonstrated that skeletal olivine can grow in situ in ultrabasic liquids (Drever, 1956; Drever & Johnston, 1957).
Discussions on primary magma compositions are often reduced to the following shortlist of frequently asked questions: How can the composition of a lava flow be used to derive the composition of its primary magma extracted from the melting regime in the mantle? How can an evaluation be made of the effects of olivine fractionation in crust and mantle lithosphere, and the possible sampling of wall-rock olivines from the melting regime? How can we best describe potentially complex fractional melting products of the mantle with equilibrium experimental data? Can the physics of the melt collection process be understood by identifying geochemical differences in the magmatic products of fractional and equilibrium melting? Can inferences be made concerning the extent of melting and crustal thickness from the major element geochemistry of ultramafic lavas? What is the role of heterogeneous mantle? How hot and deep must the melting regime be to produce an ultramafic magma, and can this be used to test the plume model?
Previous discussions of ultramafic rocks have been aggravated by taxonomical confusion over the distinction between picrites and komatiites. For example, we show in this paper that Cretaceous ‘komatiite’ lava flows from Gorgona Island and Tertiary ‘picrites’ from West Greenland crystallized from primary magmas with nearly identical major element compositions and T–P conditions of melting. The difference, therefore, is purely textural in that the ‘komatiite’ flows are spinifex-textured whereas the picrites are generally not. The qualification that komatiites be spinifex-textured (Arndt & Nisbet, 1982) has been dropped with the new IUGS chemical classification (Le Bas, 2000). Nevertheless, discussion continues over the merits of the new classification (Kerr & Arndt, 2001; Le Bas, 2001). As we are concerned mostly with the origin of primary magmas that fit the definitions of both picrite and komatiite, we will use the term ‘ultramafic’ magma to refer to all magmas with more than 18 wt % MgO.
In previous papers it was shown that Gorgona primary magmas are similar in composition to experimentally constrained anhydrous liquids in equilibrium with harzburgite [L + Ol + Opx] at 3–4 GPa (Herzberg & O’Hara, 1998; Herzberg & Zhang, 1998). Experimental liquid compositions and natural lavas were examined in simplified four-component CMAS space and projected into coordinates defined by olivine, pyroxenes, plagioclase, spinel or garnet. This method of data representation, developed by O’Hara (1968a), has the advantage of permitting a simple visual inspection to be made of the full range of analogue magma compositions that can be extracted from the mantle. However, all FeO is combined with MgO in projections, and information is lost in this procedure. Furthermore, the representation of lava compositions in projection is inadequate for distinguishing equilibrium melting from fractional melting processes.
We attempt to provide answers to the above shortlist of frequently asked questions by reporting a new method for constraining the composition of a primary magma for an olivine-phyric lava series. The method involves interfacing projections with diagrams that show absolute weight percentages of FeO and MgO for both equilibrium and fractional melting of various assumed peridotite sources. It differs from the computation of a parental magma using olivine compositions and FeO–MgO exchange coefficients (e.g. Albarède, 1992; Langmuir et al., 1992; Nisbet et al., 1993; Hauri, 1996; Larsen & Pedersen, 2000; Thompson & Gibson, 2000). Parental magmas may not be the same as primary magmas that exit the melting regime owing to modification by olivine crystallization in crustal magma chambers or mantle conduit walls (Larsen & Pedersen, 2000). Our method sees through olivine fractionation effects, and it anticipates potential problems that can arise from the use of olivine phenocrysts or wall-rock xenocrysts with complex fractional melting histories. Indeed, we demonstrate that olivine compositions can provide poor constraints on primary magma MgO content. Model primary magma compositions are used to evaluate the temperatures, pressures and petrological structures of the melting regimes for picrites and komatiites of Phanerozoic age. We show that melting is characteristically hot and deep, consistent with the plume model.
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A HYBRID FORWARD AND INVERSE MODEL FOR CALCULATING PRIMARY MAGMA COMPOSITION |
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TOPABSTRACTINTRODUCTIONA HYBRID FORWARD AND...MAGMAS IN PROJECTIONCALCULATED MAGMA COMPOSITIONS:...INVERSE MODELSGORGONA KOMATIITES: AN EXAMPLE...HAWAII (KILAUEA)TERTIARY PICRITES AND KOMATIITES...MID-OCEAN RIDGE BASALTSPLUME THERMAL AND PETROLOGICAL...CONCLUSIONSREFERENCES |
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The forward component identifies potential primary magma compositions as functions of melt fraction for an assumed peridotite source and melting mechanism. This is done by a parameterization of a large experimental database over a wide pressure range using diagrams of CMAS projections and FeO–MgO. Peridotite compositions that we examine are listed in Table 1. The melting mechanisms that we explore are equilibrium, perfect fractional, and accumulated perfect fractional melting. The inverse component selects a derivative liquid in an erupted lava suite for which olivine fractionation is computed, and this provides an array of possible primary magma compositions. We compare the arrays of potential primary magmas in both forward and inverse models, and make use of model melt fractions in the forward component to seek a unique primary magma solution. Model solutions provide primary magma composition, melt fraction, residuum mineralogy, and temperature of eruption and melt collection in the melting regime. Computed primary magma compositions depend on assumptions concerning peridotite source composition and melting mechanism, and we evaluate these effects in the forward model component.
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MAGMAS IN PROJECTION |
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TOPABSTRACTINTRODUCTIONA HYBRID FORWARD AND...MAGMAS IN PROJECTIONCALCULATED MAGMA COMPOSITIONS:...INVERSE MODELSGORGONA KOMATIITES: AN EXAMPLE...HAWAII (KILAUEA)TERTIARY PICRITES AND KOMATIITES...MID-OCEAN RIDGE BASALTSPLUME THERMAL AND PETROLOGICAL...CONCLUSIONSREFERENCES |
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The compositions of magmas formed by equilibrium melting of mantle peridotite are shown in projection in Fig. 1a and b. These projections are based on experiments in the system CaO–MgO–Al2O3–SiO2 summarized by Herzberg & O’Hara (1998). Other components such as TiO2, Cr2O3, FeO, MnO, Na2O, K2O and NiO can expand and contract liquidus crystallization fields to varying degrees. The replacement of MgO by FeO does little to change the projected locations of the cotectics and invariant points, but the effects of TiO2, Na2O and K2O are more substantial. Cotectic shifts have been evaluated in the projection code by optimizing the match to the phase equilibria in the system CaO–MgO–Al2O3–SiO2 (Herzberg & O’Hara, 1998). A description of the projection code is given in Electronic Appendix 1, which may be downloaded from the Journal of Petrology Web site at http://www.petrology.oupjournals.org. A CMAS description of mantle melting reduces the system SiO2–TiO2–Al2O3–Cr2O3–FeO–MnO–MgO–CaO–Na2O–K2O–NiO to four ‘equivalent’ components in CaO–MgO–Al2O3–SiO2. The experimental database used in the construction of these projections was listed by Herzberg & O’Hara (1998), and is here upgraded with new data at pressures from 1 to 2·7 GPa (Falloon et al., 1999a, 1999b; Gudfinnsson & Presnall, 2000), previously published data that we had overlooked (Kushiro, 1996) and the experimental results of Walter (1998). A preliminary draft of Walter’s initial results was made available to the first author, and these were encorporated into the projections of Herzberg & O’Hara (1998). The final set of experimental results that were published by Walter (1998) differed somewhat in that only superior data were included, those being data for which an optimum mass balance was demonstrated. Data of Kinzler (1997) have not been used because totals in many reported liquid compositions are low (i.e. 97–99%), and high pressures arise as an artefact of low SiO2.
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The new projections in Fig. 1 are very similar to those reported by Herzberg & O’Hara (1998), differing mainly in a more precise location of the harzburgite cotectics [L + Ol + Opx] at 3 and 4 GPa. In most cases the revised harzburgite cotectics at 2–6 GPa describe the experimental data to within ±0·5 GPa (Herzberg & O’Hara, 1998). Experimental picritic liquid compositions reported by Longhi (1995) exhibit model pressures in projection that are typically 0·5 GPa higher than his experimental nominal pressures, a difference that points to uncertainties associated with pressure calibrations (Herzberg et al., 2000; Hirschmann, 2000). The 1 GPa cotectics for lherzolite [L + Ol + Opx + Cpx ± Spinel ± Chromite) and harzburgite [L + Ol + Opx ± Chromite) shown in Fig. 1a describe data for a wide range of compositions to within ±0·2 GPa (Hirose & Kushiro, 1993; Baker & Stolper, 1994; Baker et al., 1995; Kushiro, 1996; Hirschmann et al., 1998; Falloon et al., 1999a).
The phase relations shown in Fig. 1a and b apply to equilibrium melting of a fertile and a depleted mantle peridotite (Table 1). The fertile peridotite is a xenolith from Kettle River, British Columbia (KR-4003; Xue et al., 1990; Walter, 1998), and the depleted composition is a sample of abyssal peridotite (Baker & Beckett, 1999). Oxides in KR-4003 sum to 99·18 (Xue et al., 1990), and we suspect that this arises from low MgO (Table 1); liquids analysed by Walter (1998) in experiments on this sample containing only liquid + olivine are better described with a bulk composition having 38·12% MgO rather than 37·3% MgO (Table 1). Liquids in low-variance assemblages such as spinel-lherzolite [L + Ol + Opx + Cpx + Sp] project along common cotectics. Differences can arise, however, for liquids in equilibrium with olivine [L + Ol] and harzburgite [L + Ol + Opx] near the melting out of Opx owing to the differences in bulk composition (Fig. 1).
The experimental data of O’Hara et al. (1971) and Walter (1998) are consistent with a 2·7 GPa pressure for the transformation of spinel to garnet peridotite (Herzberg et al., 2000). This pressure defines an important break in the pressure-induced track of liquids formed by initial melting of mantle peridotite (Fig. 1a). At higher pressures, liquids in equilibrium with garnet peridotite become substantially enriched in MgO and depleted in Al2O3 (see also Herzberg, 1992). At pressures between 3·0 and 4·0 GPa, initial melts of fertile peridotite KR-4003 are in equilibrium with Opx (L + Ol + Opx + Cpx + Gt; Walter, 1998). At pressures between 4·0 and 4·5 GPa, initial melts of KR-4003 are produced from the Opx-free assemblage [L + Ol + Cpx + Gt] owing to the stabilization of subcalcic clinopyroxene (Walter, 1998). When Opx becomes unstable, liquids on the solidus must have lower SiO2 than those in equilibrium with Opx because of the reaction Ol + Cpx + Gt = L + Opx (e.g. Herzberg & Zhang, 1998; Walter, 1998).
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CALCULATED MAGMA COMPOSITIONS: BALANCE OF MgO AND FeO IN MAGMAS, SOURCES AND RESIDUES |
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TOPABSTRACTINTRODUCTIONA HYBRID FORWARD AND...MAGMAS IN PROJECTIONCALCULATED MAGMA COMPOSITIONS:...INVERSE MODELSGORGONA KOMATIITES: AN EXAMPLE...HAWAII (KILAUEA)TERTIARY PICRITES AND KOMATIITES...MID-OCEAN RIDGE BASALTSPLUME THERMAL AND PETROLOGICAL...CONCLUSIONSREFERENCES |
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Partition and exchange coefficients
The procedure of Beattie et al. (1991) and Jones (1995) has been adopted for computing the compositions of olivine and orthopyroxene in equilibrium with a liquid of a known composition. The liquid composition in weight percent oxides is first recalculated to mole percent oxides, and the compositions of olivine and orthopyroxene are computed using the molecular percent partition coefficients:
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We have added to and re-evaluated the constants Ai and Bi of Beattie et al. (1991) and Jones (1995) for TiO2, Al2O3, Cr2O3, Fe2O3, MnO, CaO, Na2O and K2O, using experimental data described in Electronic Appendix 2 (available from the Journal of Petrology Web site at http://www.petrology.oupjournals.org). Results are listed in Tables 2 and 3. For the partitioning of Fe2O3, DFe2O3Ol/L = 0, and DFe2O3Opx/L = 0·3; this partition coefficient for Opx/Liquid yields Fe3+/
Fe = 0·06 as observed in orthopyroxenes from a wide range of peridotites (Canil et al., 1994). For the partitioning of FeO, AFeOOl/L and AFeOOpx/L are equivalent to the exchange coefficients
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Roeder & Emslie (1970) determined the exchange coefficient for olivine to be 0·30 ± 0·03 for basaltic liquids with about 8% MgO at 1 atm, a value that is commonly used in petrological modelling. This is observed again in Fig. 2 with our database, but it can also be seen that KDOl/LFeO/MgO increases to 0·35–0·36 for liquids with MgO > 20%. Clearly, the values of AFeO in equations (3) and (4) are not constant. Experimental data shown in Fig. 2 have been regressed with the equations
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The uncertainty in calculating KDOl/LFeO/MgO from equations (8) and (9) is ±0·019 at the 1
level, compared with an uncertainty of ±0·011 (1) that arises strictly from errors in the analysis of olivine and glass using the electron microprobe. Uncertainties arising from Fe2O3 should be small because experimental samples were encapsulated in either carbon or metallic iron, as discussed in Electronic Appendix 2. The difference between ±0·019 and ±0·011 is fairly small, and most of it is likely to arise from temperature and pressure effects (Takahashi & Kushiro, 1983; Ulmer, 1989; Sobolev & Danyushevsky, 1994). Indeed, we observe positive correlations of KDOl/L with temperature and pressure. However, regressions of KDOl/L vs pressure and temperature yield 1 uncertainties of ±0·024 (correlation coefficient R = 0·46) and ±0·021 (R = 0·60), respectively, larger than standard deviations for KDOl/L vs liquid MgO content (±0·019; R = 0·69). Values of KDOl/L estimated from equations (8) and (9) are significantly lower than those of Ulmer (1989). For example, both experiments reported by Walter (1998) and our regression in Fig. 2 yield KDOl/L = 0·34–0·36 for liquids equilibrated at 3·0–7·0 GPa, whereas Ulmer (1989) predicted 0·38–0·51. We have not attempted to resolve the independent effects of temperature, pressure and composition on KDOl/L for several reasons. First, it is difficult to do so with confidence because the MgO content of the liquid is highest in experiments conducted at elevated temperatures and pressures. Second, our method of calculating KDOl/L from liquid composition has the advantage of considerable simplicity because it does not require an independent way of evaluating pressure.
The fraction of sites available for Fe2+ and Mg is 0·667 and 0·5 for olivine and orthopyroxene, respectively. These are filled with Mn, Ca, Na, K and Ni; Ti is partitioned exclusively into the tetrahedral sites; Al and Cr are distributed equally between tetrahedral and octrahedral sites to maintain charge balance. The remaining octahedral sites are filled with Fe2+ and Mg, calculated by rearranging equations (6) and (7) to
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Effect of H2O on exchange coefficients
Ulmer’s (1989) FeO–MgO exchange coefficients for experimentally equilibrated olivine–liquid pairs were for starting material compositions that contained 1·98–3·28 wt % H2O. These hydrous data can be reproduced to within ±0·023 (1) using equations (8) and (9) for anhydrous data, indicating that H2O has little effect on KDOl/LFeO/MgO.
Liquids in equilibrium with olivine and harzburgite
The compositions of liquids in equilibrium with olivine [L + Ol] and harzburgite [L + Ol + Opx] were computed by mass balance solutions to the equilibrium melting equation
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The bulk distribution coefficient is the concentration of the oxide component in the multiphase residue divided by the concentration in the liquid, and takes the form
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Coexisting liquid and crystalline compositions were solved with equation (13) using an iterative procedure that incrementally varies F from 1·0 to a value that is regulated by the stability of clinopyroxene. Liquid compositions in equilibrium with olivine (L + Ol) are solved for XOpx = 0, and liquid compositions in equilibrium with harzburgite (L + Ol + Opx) are solved for XOpx = 0 to a value that is also regulated by the stability of clinopyroxene (L + Ol + Opx + Cpx). Clinopyroxene stability is constrained experimentally in Fig. 1, into which computed liquid compositions are projected. Harzburgite cotectics in Fig. 1 at 1–6 GPa constrain the pressure for projected liquid compositions. For example, in the case of equilibrium melting of fertile mantle peridotite KR-4003 (Table 1), Cpx melts out when F reaches 0·23 at 1·0 GPa, similar to that for other fertile compositions (Langmuir et al., 1992; Baker & Stolper, 1994). The effect of pressure is to increase F at which Cpx remains stable, and this is shown in more detail below.
The contents of FeO and MgO in liquids equilibrated with olivine and harzburgite residua at 1 GPa have been calculated for peridotite compositions PHN1611 (Kushiro, 1996) and MM3 (Hirschmann et al., 1998; Falloon et al., 1999a), and the results are shown in Fig. 3. The calculated compositions agree well with the experimentally observed liquids, and are consistently within the uncertainties associated with both the location of the harzburgite cotectic (Fig. 1a; ±0·2 GPa) and KDOl/L (Fig. 2; ±0·019 at the 1 level). Peridotite 1611 is considerably more FeO-rich than MM3 (Table 1), and this propagates to the partial melt products in a predictable way. Indeed, it is possible to independently compute the liquid compositions of PHN1611 from those for MM3. This can be done by the approximations
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, data of 
, data of The approximations arise because D values for PHN1611 are not the same as D values for MM3 at a given melt fraction F. Nevertheless, the compositions of liquids formed by equilibrium melting of PHN1611 calculated by these different methods agree very well with the experimentally observed compositions (Fig. 3). We therefore have independent ways of evaluating the FeO and MgO contents of liquids formed by partial melting of a peridotite composition for which no experimental data are available.
Liquids on the solidus
We can constrain the compositions of liquids that form immediately on the solidus where F 
0 by combining equations (8) and (11):
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A forward model for equilibrium melting of a fertile peridotite source (FeO 
8·0%)
A computation has been made of contents of SiO2, TiO2, Al2O3, Cr2O3, FeO, MnO, MgO, CaO, Na2O, K2O and NiO for magmas produced by equilibrium melting of a fertile peridotite source with 8·0% FeO (Table 1; KR-4003) at pressures that range from 1 to 6 GPa. Results for FeO and MgO are shown in Fig. 4a and representative model liquid compositions are listed in Table A1 of Electronic Appendix 3 on the Journal of Petrology Web site at http://www.petrology.oupjournals.org. Liquids were computed for olivine [L + Ol] and harzburgite [L + Ol + Opx] assemblages by the method discussed above, and for liquids immediately on the solidus. Uncertainties arising from 1 in KDOl/LFeO/MgO are typically ±1·4% MgO and ±0·52% FeO, and propagate to an uncertainty of ±1·5 GPa where isobaric cotectics are compressed in FeO–MgO space (Fig. 4a).
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Liquids immediately on the anhydrous solidus were computed from equation (18) and the composition of subsolidus olivine, which has an mg-number of 89·5 (Walter, 1998; olivine analysis 60·02). The array of permissible solutions for FeO and MgO in liquids in equilibrium with this composition of olivine at the solidus is accurately defined (Fig. 4a), but difficulties arise in calibrating this array for pressure. For example, in the absence of any experimental data, the FeO and MgO contents of the first drop of liquid on the solidus at 4 GPa could be located anywhere on the solidus array.
Ironically, liquids on the solidus at 4–7 GPa may be better known than those at 1–3 GPa. The data of Walter (1998) show that FeO and MgO are negatively correlated and linear for [L + Ol + Cpx + Gt] at 6 and 7 GPa and for F ranging from 0·11 to 0·41. These isobaric FeO–MgO arrays are nearly orthogonal to the solidus array, which is intersected at well-defined points (Fig. 4a). Liquids on the solidus at 1 and 3 GPa were constrained by factoring liquid compositions for MM3 and PHN1611 according to bulk composition, as discussed above (Fig. 3). At 1 GPa the FeO and MgO contents of liquids at F in the 0–0·20 range nearly parallel those on the solidus FeO–MgO array (Figs 3 and 4a). For a basalt with 12% MgO, a ±1 uncertainty in KDOl/LFeO/MgO propagates to FeO contents at 1 GPa that cannot be statistically distinguished from those at 2 GPa. It is therefore understandable that disagreement exists over the FeO and MgO contents of the near-solidus liquid at 1 GPa for peridotite MM3 (Baker & Stolper, 1994; Baker et al., 1995; Hirschmann et al., 1998; Falloon et al., 1999a). The positions of the model isobaric cotectics between the solidus and harzburgite saturation shown in Fig. 4a usually differ from experimental determinations by no more than ±0·5 GPa. However, this uncertainty can increase to ±1·5 GPa where the isobaric cotectics are compressed in FeO and MgO space, similar to those for [L + Ol + Opx]. In most cases, the model FeO–MgO isobaric cotectics agree with experimental cotectics to within ±1 GPa in the 1–7 GPa range and therefore pressure information cannot be accurately obtained for natural lava compositions. We believe that a ±1 GPa uncertainty is a problem that is common to all petrological models that operate within the ±0·019 bound of uncertainty in KDOl/LFeO/MgO.
A forward model for equilibrium melting of an Fe-rich fertile peridotite source (FeO 9·0%)
We now examine the consequences of partial melting of a peridotite source that is enriched in FeO. The average of 590 samples of peridotite from off-craton occurrences contains 8·14% and this varies by ±0·90% FeO at the 1 level (Herzberg, 1993). Most peridotites from ocean basins and xenoliths display wide variations in MgO content at 8·0% FeO, and they display no relationship between FeO and indicators of fertility or depletion (e.g. TiO2, Al2O3, Na2O; Herzberg, 1993; Baker & Beckett, 1999; Griffin et al., 1999; excluding cratonic mantle). Calculations have been performed on a synthetic composition formed from Kettle River peridotite KR-4003 by a simple replacement of MgO with FeO. This Fe-rich source contains 9·04% FeO (Table 1), which is higher than the average FeO content by 1 (i.e. 8·14 + 0·90), and it is identically located in projection with its precursor, KR-4003.
Shown in Fig. 4b are the compositions of liquids formed by equilibrium melting of this Fe-rich fertile source with residual harzburgite and olivine. These liquids are elevated in FeO compared with normal fertile peridotite with about 8% FeO, a result that is intuitively obvious. Indeed, liquids are elevated in FeO by about 0·9% at most pressures, an amount that corresponds to enrichment of normal mantle with 0·9% FeO. We do not provide pressure information for compositions on the solidus or between the solidus and harzburgite-saturated assemblages because we have no experimental data for pressures in the 2–6 GPa range.
A forward model for equilibrium melting of a depleted peridotite source (FeO 8·0%)
The depleted peridotite sample chosen for this forward model is an abyssal peridotite from Baker & Beckett (1999). Inspection of Table 1 shows that it has nearly the same composition as the average of 590 mantle peridotite samples from off-craton occurrences (Herzberg, 1993). This depleted source contains about 42% MgO and 8% FeO, similar to FeO in fertile peridotite.
Shown in Fig. 5a are the compositions of liquids formed by equilibrium melting of this depleted source with residual harzburgite and olivine. Again, we do not provide pressure information owing to a lack of experimental data. The compositions shown for olivine- and harzburgite-saturated assemblages are compared with those for equilibrium melting of fertile peridotite in Fig. 5b. In the 1 atm to 2 GPa pressure range, there is a similarity in FeO and MgO contents of harzburgite-equilibrated liquids for fertile and depleted sources having 8·0% FeO. However, at 3–5 GPa, harzburgite-equilibrated liquids of a depleted source with 8·0% FeO are similar in FeO and MgO to those for Fe-rich fertile peridotite with 9·0% FeO. These observations illustrate the harzards that can result from using the composition of an erupted magma to infer the FeO content of a peridotite source.
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FeO–MgO–F systematics for all forward models of equilibrium melting
It is sometimes convenient to consider a residue–liquid equilibrium rather than an olivine–liquid equilibrium, although both are obviously identical when olivine is the only residuum phase. If the mg-number of a residue is nearly identical to that of olivine in that residue, then equation (18) can be modified so that (FeO/MgO)Ol is replaced by (FeO/MgO)o, the initial source rock composition. Indeed, the arrays of FeO and MgO we compute for liquids on the solidus using source rock composition are nearly identical to those computed using olivine in equation (18), the difference being <0·1% FeO. The equation for the exchange coefficient can therefore be generalized as
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0, FeO–MgO solutions to these equations are essentially identical to those derived from equation (18) for the olivine–liquid equilibrium for peridotite compositions KR-4003, MM3 and PHN1611. For olivine–liquid and harzburgite–liquid equilibria, solutions to equations (21), (22) and (23) at elevated melt fractions yield FeO–MgO–F systematics that are identical to those that have been independently calculated from mass balance solutions to the batch melting equation for all oxide components. However, solutions to equations (21)–(23) are applicable to all phase assemblages and at all melt fractions. These equations have therefore been applied to FeO–MgO–F systematics for phase assemblages ranging from the solidus to [L + Ol + Opx] in Figs 4b and 5a for Fe-rich fertile peridotite and depleted peridotite.
A forward model for FeO–MgO–F systematics specific to isobaric accumulated perfect fractional melting (APFM) of a fertile peridotite source (FeO 8·0%)
Liquids formed by isobaric accumulated fractional melting were determined by first computing equilibrium liquid and residue compositions using equation (13) for equilibrium melting as discussed above. This provides a set of D and F values that were then used in Shaw’s (1970) equation for accumulated fractional melting:
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The computational procedure is discussed in more detail in Electronic Appendix 4 on the Journal of Petrology Web site at http://www.petrology.oupjournals.org. Results are shown in Fig. 6a.
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An apparent paradox of Shaw’s (1970) equations for fractional and accumulated fractional melting is that the MgO content of the residue becomes greater than the MgO content of olivine when olivine is the sole residue phase, a problem that is also discussed in Electronic Appendix 4. We therefore examine accumulated perfect fractional melting using the following incremental batch melting approach. We compute liquid compositions with 0·01 mass fractions of melting of the residue, the liquid is completely removed from its residue, the residue is melted again at successive 0·01 mass fractions, and each new melt is perfectly extracted and mixed with the previously formed melt. The process is repeated many times until the residue is completely melted, and each instantaneous melt is formed by equilibrium melting of an instantaneous residue whose composition is continuously changing owing to melt extraction. Each 0·01 mass fraction of melt extracted from the residue is called an ‘instantaneous’ drop of melt. This analysis is restricted to fertile peridotite KR-4003 for which we have pressure information. We consider first the case of isobaric melting at 1–6 GPa, and this is followed by consideration of polybaric melting situations.
The MgO content of the instantaneous liquid is computed from
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The amount of liquid produced at the first increment of melting is 0·01, and this becomes progressively lower than 0·01 with each subsequent increment of melting because the mass of residue is progressively being reduced from 1·0 to zero. We keep track of the melt fraction of each increment with respect to the initial source mass, and call this Fin. Therefore, at the first increment of melting Fin = FS, but during subsequent increments Fin < FS. For the total melt fraction
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For constant D, this incremental solution to the batch melting equation for perfect melt extraction yields results that are identical to those calculated using the equations of Shaw (1970). That is, the instantaneous batch liquid composition is identical to Shaw’s (1970) perfect fractional melt composition, and the aggregated melt compositions in both cases are also the same.
Results for fractional melting of fertile peridotite with variable D are shown in Fig. 6a and b. Liquids formed by accumulated fractional melting have higher FeO contents than those for liquids that form by equilibrium melting (Fig. 4a) at any specific melt fraction, a conclusion reached previously by Langmuir et al. (1992). The most important observation is that identical FeO–MgO–F systematics are obtained for accumulated fractional melts calculated with Shaw’s (1970) equation and incremental solutions to the batch melting equation. The utility of the incremental batch approach is that it can be easily applied to polybaric melting, as discussed below, and it can be modified to include imperfect melt extraction, a problem we will not consider further. However, a close examination of Fig. 6a and b reveals several minor but important differences. The incremental batch melting approach (Fig. 6b) provides pressure information whereas Shaw’s (1970) equation does not. Additionally, there are minor differences in the compositions of aggregate liquids that form when olivine is the sole residue phase, an issue that is discussed further in Electronic Appendix 4.
Residue compositions of equilibrium and fractional melting are important for a complete mass balance analysis of peridotite melting, and bear on many problems of geological interest. Although a comprehensive discussion of residues is beyond the scope of this paper, there are several important aspects of this problem that require discussion. We consider only the mg-numbers of olivines in the residues. For the case of equilibrium melting of fertile peridotite, the mg-number of olivine increases from 89·5 to 95·9 from the solidus to the liquidus (Fig. 7a). For equilibrium melting of depleted peridotite, the mg-number of olivine increases from 90·2 to 96·2 from the solidus to the liquidus (Fig. 7a). The effect of perfect fractional melting is to produce residual olivines with mg-numbers that increase to 100 at the liquidus (Fig. 7b) for both depleted and fertile peridotite compositions. We consider now olivine phenocrysts with mg-numbers of 92·0 precipitated from an aggregate magma formed by accumulated perfect fractional melting of a fertile or depleted source. Olivines in the residue of the pooled liquid are in equilibrium with the last drop of perfect fractional melt, and these have mg-numbers of about 94·0 (Fig. 7b). A primary accumulated magma has a unique residue, but the accumulated liquid is not in equilibrium with that residue; only the final drop of liquid extracted is in equilibrium with the residue. This is the process treated by Gast (1968) and Shaw (1970) and specifies that the liquids that are integrated are produced from a progressively depleting residue and are then added in aliquots to the final accumulation.
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A forward model for FeO–MgO–F systematics specific to polybaric accumulated perfect fractional melting (APFM) of a fertile peridotite source (FeO 8·0%)
An examination is now made of perfect polybaric accumulated fractional melting of fertile peridotite KR-4003. Before proceeding with the FeO and MgO characteristics of such melts, we consider the range of temperatures and pressures under which decompression melting can occur, and we do this with aid of the phase diagram for Kettle River peridotite KR-4003 shown in Fig. 8a. The anhydrous solidus for mantle peridotite is from Herzberg et al. (2000), which is almost identical to that offered by Hirschmann (2000) except that it is higher by about 30°C at 3–4 GPa. Temperatures for olivine-in and orthopyroxene-in were computed using the method of Beattie (1993) for liquid in equilibrium with olivine ± orthopyroxene at 1 atm, and extended to high pressures by a recalibration with experimental data summarized above for KD. Using 322 data points from 0·2 to 7·0 GPa and 1100 to 1950°C, we have obtained minimum residuals between observed and calculated temperatures with
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level for pressures up to 7 GPa; this is only marginally higher than that obtained by Beattie (1993) for a much smaller dataset to 4 GPa. Temperatures for Cpx- and Gt-in were calculated for harzburgite-saturated liquids at the point of saturation in Cpx and Gt as defined by Fig. 1. The difference between calculated temperatures and those in all experiments conducted by Walter (1998) at 3–7 GPa on this same peridotite composition averages 16°C.
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We examine the FeO and MgO contents of liquids that are formed along the five T–P paths shown in Fig. 8a, where melting initiates at 3, 5, 7 and 10 GPa. In one case the melting path is isothermal, and the other four approximate adiabatic gradients from Iwamori et al. (1995) but modified to remove ‘bumps’ that arose from strong T–P curvature to their solidus for spinel-lherzolite (Herzberg et al., 2000). The other adiabatic T–P paths shown in Fig. 8a are from Asimow et al. (2001). Although we consider only linear paths in T–P space, the conclusions that follow are applicable to more realistic decompression melting paths. Fractional melt compositions were computed using equations for isobaric incremental batch melting, except that DMgO,in was parameterized as a polybaric function of CMgO,Sin; values of DMgO,in and CMgO,Sin at each pressure along the polybaric path were parameterized from solutions to isobaric equilibrium melting (Fig. 4a). Instantaneous residue compositions vary continuously along each melting path owing to perfect melt extraction. Compositions of the accumulated melts were calculated from perfect fractional melts along each path between the solidus and the final melting pressure. Final melting pressures vary at 1 GPa increments below the initial melting pressure. Results are shown in Fig. 8b.
In most cases there is a substantial drop in the FeO content of liquids with decompression, but MgO varies by no more than 2% despite considerable differences in the melting paths (Fig. 8b). The reason why there is little change in MgO is because all T–P melting paths are approximately coincident with T–P isopleths of MgO in liquids defined by olivine saturation surfaces (Fig. 8a). Dissolution of ‘wall-rock’ olivine into ascending magmas should not be significant, and olivine–liquid equilibrium temperatures must closely reflect actual eruption temperatures. An important corollary is that the MgO content of a primary magma is similar to the MgO content of the initial melt on the solidus, and it reflects the pressure of initial melting.
There is also an important relationship between aggregate polybaric and isobaric melt compositions. Let us consider the case of decompression melting starting at 5 GPa and terminating at about 1·5 GPa along the path designated at 5’ in Fig. 7a and b. The aggregate melt composition has 20·5% MgO and 10·1% FeO and is formed at a melt fraction of 0·40. This same melt composition is formed isobarically at 4 GPa, also at an average melt fraction of 0·40 (Fig. 8b). The aggregate magma formed polybarically therefore records in its geochemistry an average pressure of 4 GPa, as defined by isobaric 4 GPa FeO–MgO–F relations.
For melting regimes in oceanic ridges and mantle plumes, there will be an infinite number of streamlines, with horizontal and vertical components, along which fractional melts are formed and in local equilibrium with their residues. The melting regime may then be visualized as a spider’s web of streamlines, and the primary aggregate magma is an integrated product of all fractional melts that collect in the melting regime. Any true forward melting model must therefore compute the compositions of all fractional melts along all streamlines, and then perform the integration within the melting regime. The complexity of this procedure is immediately apparent. The importance of isobaric FeO–MgO–F systematics is that they provide a simple means of constraining the composition of a primary aggregate melt for any melting regime, regardless of the polybaric streamline complexity. This is explored in the subsequent discussion.
Forward models for FeO–MgO–F systematics specific to accumulated perfect fractional melting (APFM) of depleted and iron-rich peridotite source
Liquids in equilibrium with olivine [L + Ol] and harzburgite residues [L + Ol + Opx] were discussed above (Fig. 5b). The bulk distribution coefficients D for FeO and MgO were used together with F and Shaw’s (1970) equation, (24), to calculate liquids formed by accumulated perfect fractional melting. Results are illustrated below in the discussions on Gorgona and Hawaiian magmatism. We cannot provide pressure information for the aggregate liquids because it is not known for equilibrium melting.
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INVERSE MODELS |
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TOPABSTRACTINTRODUCTIONA HYBRID FORWARD AND...MAGMAS IN PROJECTIONCALCULATED MAGMA COMPOSITIONS:...INVERSE MODELSGORGONA KOMATIITES: AN EXAMPLE...HAWAII (KILAUEA)TERTIARY PICRITES AND KOMATIITES...MID-OCEAN RIDGE BASALTSPLUME THERMAL AND PETROLOGICAL...CONCLUSIONSREFERENCES |
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We restrict our inverse models to well-documented high-MgO aphyric and olivine-phyric lavas that are related to each other in space and time by fractional crystallization of olivine. We assume that olivine was the sole fractionating phase that relates all lavas to each other and all lavas to the primary magma from the melting regime. The inversion method consists of selecting a representative lava composition into which olivine is incrementally added or from which olivine is incrementally subtracted. The composition of olivine in local equilibrium with the liquid is computed by the method discussed above, where KDOl/LFeO/MgO is adjusted for changes in liquid composition, and each increment represents 1 wt % addition or subtraction of olivine. These inverse calculations produce a loci of compositions that can be compared with the compositions of liquids in the various forward models discussed above. A computational algorithm is provided in Electronic Appendix 5 on the Journal of Petrology Web site at http://www.petrology.oupjournals.org so that the reader can replicate our calculations.
Several workers, in attempts to compute mid-ocean ridge basalt (MORB) and hotspot primary magmas, have assumed that all Fe in published lava analyses was FeO (i.e. FeOT; Langmuir et al., 1992), whereas others have attempted to consider only the true amount of FeO in the lava that is exchangeable with olivine (Albarède, 1992; Nisbet et al., 1993; Larsen & Pedersen, 2000). We adopt the latter approach because the amount of Fe2O3 can be large, with Fe3+/Fe ranging from about 0·07 to 0·15 in most MORB and komatiite occurrences reviewed below. Only when both FeO and Fe2O3 are evaluated in a lava suite of interest can we make a comparison with our forward models where iron has been calculated strictly as FeO from the exchange coefficient KDOl/LFeO/MgO. Failure to do so can result in substantial errors. For example, in the case where MgO = 20·0%, FeOT = 10·0% and Fe3+/Fe = 0·10, it can be determined that Fe2O3 = 1·11% and FeO = 9·0%. Inspection of Fig. 4 reveals that use of 10% FeOT instead of 9·0% FeO would yield a pressure for an equilibrium primary magma that is about 2 GPa too high.
For Kilauea and Baffin Island lavas we use FeO and Fe2O3 that have been directly measured by wet chemistry (Fe3+/Fe = 0·1; Wright et al., 1975; Robillard et al., 1992). In all cases we treat lavas that move from the melting regime to the surface as a system that is closed to oxygen, and assume no change in Fe3+/Fe for both primary and derivative lavas. A small but additional source of uncertainty is that we compare lavas for which both FeO and Fe2O3 have been constrained with forward models that contain no Fe2O3, as for fertile peridotite KR-4003 and its Fe-rich analogue. Although it is possible to include Fe2O3 in the forward model computations for olivine- and harzburgite-saturated liquids, as has been done for the depleted peridotite composition, the partition coefficients are not well constrained for assemblages containing garnet and clinopyroxene. Primary komatiite magmas are expected to contain 1–2% Fe2O3, and this will result in a modest dilution of all oxide elements in our computed forward models. For example, a primary magma with 10% FeO we calculate in a forward model would actually contain 9·8% FeO if it also had 2% Fe2O3 (10% FeO x 98/100). Our forward and inverse models should therefore be useful in situations where the amount of Fe3+/Fe > 0·1.
In cases where FeO and Fe2O3 have not been directly measured in the lavas (i.e. West Greenland and Gorgona), they have been computed from FeOtotal using the method of Kress & Carmichael (1991) and estimates of oxygen fugacity (Canil, 1999; Larsen & Pedersen, 2000). Oxygen fugacities are compared with those for the following buffers: nickel–nickel oxide (NNO) from O’Neill & Pownceby (1993), fayalite–quartz–magnetite (FQM) from O’Neill (1987), and iron–wüstite (IW) from O’Neill & Pownceby (1993). For example, Christie et al. (1986) determined Fe3+/Fe = 0·07 ± 0·03 for 78 MORB glasses. An fO2 = 2·5 log units more reducing than the NNO buffer (i.e. NNO - 2·5) would yield Fe3+/Fe = 0·07 for MORB using the method of Kress & Carmichael (1991), although different parameters in Kilinc et al. (1983) provide the same Fe3+/Fe with NNO = -2·3 (Christie et al., 1986). Data of Wright et al. (1975) on eruptions between 1968 and 1971 exhibit Fe3+/Fe = 0·10, only slightly more oxidizing than MORB (Fe3+/Fe = 0·07; Christie et al., 1986). A ratio of Fe3+/Fe = 0·10 is characteristic of an fO2 = 1·3 log units more reducing than the NNO buffer (i.e. NNO - 1·3; O’Neill & Pownceby, 1993).
Another minor difficulty with the inverse method is that addition and subtraction of olivine will yield a primary magma with Fe3+/Fe that differs somewhat from the selected lava composition because olivine accepts no Fe2O3. For example, the primary magma we calculate for Kilauea by adding olivine to sample DAS69-5-1 (Wright et al., 1975) contains 1·01–1·03% Fe2O3, whereas it should contain 1·18–1·19% Fe2O3 at NNO - 1·3 (Table 4). We have made no corrections for this effect because the error is small and overwhelmed by uncertainties in KD. Another difficulty is that we have completely ignored the effects of chromite precipitation. Our estimates for Cr2O3 in primary magmas are therefore unreliable.
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GORGONA KOMATIITES: AN EXAMPLE OF A HYBRID FORWARD–INVERSE MODEL |
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TOPABSTRACTINTRODUCTIONA HYBRID FORWARD AND...MAGMAS IN PROJECTIONCALCULATED MAGMA COMPOSITIONS:...INVERSE MODELSGORGONA KOMATIITES: AN EXAMPLE...HAWAII (KILAUEA)TERTIARY PICRITES AND KOMATIITES...MID-OCEAN RIDGE BASALTSPLUME THERMAL AND PETROLOGICAL...CONCLUSIONSREFERENCES |
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Detailed geochemical descriptions of Cretaceous picrites and komatiites from Gorgona Island have been provided by Echeverria (1980), Aitken & Echeverria (1984), Echeverria & Aitken (1986), Kerr et al. (1996), Arndt et al. (1997) and Révillon et al. (2000). The terms ‘picrite’ and ‘komatiite’ have been used to describe hyaloclastic volcanic rocks without spinifex textures and lava flows with spinifex textures, respectively. We will retain these terms.
Komatiite lavas are shown in Fig. 9a and b for an fO2 at NNO - 1·3. This is similar to that for Kilauea, and to fO2 estimated by Canil (1999; NNO - 1·5 ± 1·5) for Gorgona. Olivine fractionation dominates major element geochemistry, and we select a representative composition from Aitken & Echeverria (1984; Table 4) for which olivine is added and subtracted. This is the inverse model component that defines the locus of possible primary magma compositions shown in Fig. 9a and b. The forward model component is shown also in Fig. 9a and b, that being for accumulated perfect fractional melting of an assumed fertile peridotite source, composition KR-4003 in Table 1. We can uniquely determine the primary magma composition by interfacing the projection in Fig. 9a with the FeO–MgO diagram in Fig. 9b by searching for a composition that provides a unique melt fraction F. That melt fraction F is 0·28, yielding 18·8% MgO and 10·0% FeO and an average pressure of 3·8 GPa (Fig. 9b; Table 4). For Gorgona and all other cases considered in this paper, the melt fraction provides the link between inverse and forward models.
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An average pressure of 3·8 GPa is obtained for the primary magma formed by accumulated perfect fractional melting of a fertile peridotite source. This pressure probably has little significance because it is specific to a fertile peridotite source, and there exist substantial uncertainties stemming from experimental calibrations. Also, as discussed above, no pressure information is obtainable for forward models involving depleted and Fe-rich peridotite.
The primary magma composition would precipitate olivine phenocrysts having an mg-number of 90·6 (±0·5; Table 4), and fractional crystallization would produce derivative liquids that would crystallize olivines with lower mg-numbers. Our calculated olivines are in good agreement with mg-numbers of about 90–91 that are commonly observed for skeletal olivines in spinifex-textured komatiites (Révillon et al., 2000). However, rare olivines with mg-numbers of about 91·5 have also been reported (Aitken & Echeverria, 1984; Révillon et al., 2000), and these can be in equilibrium with the locus of possible FeO and MgO contents in liquids shown in Fig. 9b. One way of explaining Ol91·5 is to have more MgO in the primary magma by adding more olivine to the derivative magma in the inverse model. This approach yields a primary magma with 21% MgO rather than 18·8% in our model, but with contradictory phase equilibrium properties (i.e. L + Ol in FeO–MgO space but L + Ol + Opx + Gt in projection). Another way of explaining Ol91·5 is by reducing the FeO content of the magma. Indeed, an aggregate magma that precipitates Ol90·6 can be a mixture of perfect fractional melts in equilibrium with Ol89·4 at the solidus to Ol91·5 at the maximum extent of melting (Fig. 7b). This range of olivine compositions could form during decompression from an initial melting pressure of about 4·2 GPa (Fig. 9b), and Ol91·5 could have precipitated from advanced fractional melts with lower FeO contents than the accumulated fractional melts (Fig. 9b). The implication is that the mixing of fractional melts within the melting regime was not a perfect process. This explanation can account for the outliers in FeO contents seen for komatiites (Fig. 9b), and it will be raised again in the discussion on Gorgona picrites.
It is reasonable to question the appropriateness of using an assumed fertile peridotite composition in the forward model because trace elements and isotopes for Gorgona komatiites indicate a depleted MORB-like source (Echeverria, 1980; Kerr et al., 1996; Arndt et al., 1997). We therefore use the depleted peridotite source in Table 1 in another forward model of accumulated fractional melting, and results are shown in Fig. 10a and b. Again, we determine the primary magma composition by interfacing the projection in Fig. 10a with the FeO–MgO diagram in Fig. 10b by searching for a composition that provides a unique melt fraction F in both cases. The melt fraction is 0·11, yielding a primary magma with 18·8% MgO and 10·0% FeO, which is identical to the primary magma composition for the fertile peridotite source. We consider six possible primary magma compositions from solutions for equilibrium and accumulated fractional melting for three different peridotite source compositions. Results are listed in Table 4. For fertile and depleted source compositions, the MgO contents are restricted to 18·8–20·9%, in agreement with other estimates for MgO (Nisbet et al., 1993; Kerr et al., 1996; Révillon et al., 2000). However, melt fractions vary considerably, from 0·11 for the depleted source to 0·28 for the fertile source.
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All primary magmas plot within the harzburgite stability field [L + Ol + Opx], near the melt fraction defined by Cpx-out in both projection and in FeO–MgO plots (Figs 9 and 10). Harzburgite is therefore the residue for these primary magmas. All primary magmas formed by accumulated fractional melting must display internal consistency in residuum mineralogy when viewed in projection and FeO–MgO plots.
Picrites at Gorgona are generally lower in FeO and higher in MgO than komatiites, and they have maximum mg-numbers in the 93·3–93·6 range (Echeverria & Aitken, 1986; Révillon et al., 2000). It can be seen in Figs 9b and 10b that the komatiites could not have formed from the picrites by fractional crystallization of olivine, a conclusion reached also by Arndt et al. (1997). It can also be seen that the picrites could not have formed from the primary komatiite by advanced accumulated fractional melting; higher melt fractions would yield a primary picrite magma with slightly higher FeO, the opposite of what is observed (Fig. 9b). Indeed, any picrite that precipitates Ol93·6 is too low in FeO to be an accumulated fractional melt of either fertile or depleted mantle peridotite.
Arndt et al. (1997) interpreted the rare earth element geochemistry of Gorgona lavas by a dynamic fractional melting process wherein a hot plume interior loses its melt and leaves behind a refractory source that is subsequently melted to form the picrites. This open-system melting process can be modelled with fractional melts that do not mix, using equations (25) and (26). Shown in Fig. 11 are perfect fractional melts formed polybarically along three adiabatic melting paths in Fig. 8a. The effect of advanced partial melting during decompression is to form perfect fractional melts with lower FeO contents without much change in MgO (Fig. 11). The mixing of perfect fractional melts with a wide range of FeO yields accumulated fractional melts with the average FeO contents shown in Fig. 8b. There exists a locus of possible FeO and MgO contents of perfect fractional melts that will be in equilibrium with olivine having an mg-number of 93·6 (Fig. 11). Of these, a primary picrite magma with about 26% MgO is possible and consistent with that estimated by Révillon et al. (2000). For an assumed fertile peridotite source, this primary magma would have required an initial melting pressure of about 10 GPa, a final melting pressure of about 6 GPa, and would have formed after 0·4 mass fractions of total melting of the initial source mass (Fig. 11). Another solution has initial melting at 7 GPa and final melting at 2·5 GPa where fractional melts will be in equilibrium with Ol93·6; in this scenario, the Gorgona picrites are mixtures biased with advanced fractional melts having a wide range in FeO but with about 23% MgO (Fig. 11). The range of permissible MgO contents is about 23–26%, demonstrating that olivine with an mg-number of 93·6 provides a limited constraint on primary magma composition. It is also not possible to rigorously constrain initial melting pressure for the picrites. However, it must have been higher than the 4·2 GPa that is required to explain the komatiites (Fig. 11). An initial melting pressure of 5 GPa is even too low as it could not yield perfect fractional melts with FeO contents low enough to yield Ol93·6. The picrites appear to have formed in the hottest and deepest part of the Gorgona plume axis, in agreement with conclusions drawn by Arndt et al. (1997).
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Our calculations provide strong support for the interpretation that the Gorgona picrites formed by the dynamic fractional melting model proposed by Arndt et al. (1997). In this model, the picrites formed from the hot plume axis after the removal of earlier and lower-degree melts. Arndt et al. (1997) noted that these earliest formed melts may not have reached the surface because they have not been recognized on Gorgona or elsewhere, and that they may have been removed to another part of the plume. We concur that the picrites formed as fractional melts that did not mix with other fractional melts in the melting regime, and the komatiites formed as reasonably well-mixed aggregate fractional melts. Therefore, there must have been important restrictions to magma mixing in the Gorgona plume.
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HAWAII (KILAUEA) |
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TOPABSTRACTINTRODUCTIONA HYBRID FORWARD AND...MAGMAS IN PROJECTIONCALCULATED MAGMA COMPOSITIONS:...INVERSE MODELSGORGONA KOMATIITES: AN EXAMPLE...HAWAII (KILAUEA)TERTIARY PICRITES AND KOMATIITES...MID-OCEAN RIDGE BASALTSPLUME THERMAL AND PETROLOGICAL...CONCLUSIONSREFERENCES |
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Recent estimates place the MgO content at 13–19% for Hawaiian primary magmas (Albarède, 1992; Eggins, 1992a; Hirschmann & Ghiorso, 1994; Baker et al., 1996; Hauri, 1996; Norman & Garcia, 1999), and picrite glasses with 15% MgO have been directly observed (Clague et al., 1991). We restrict our new analysis to an evaluation of the primary magma for historical eruptions on Kilauea, and use data reported by Murata & Richter (1966), Wright & Fiske (1971), Wright (1972, 1973), Wright et al. (1975), Clague et al. (1991) and Norman & Garcia (1999). Inverse modelling was performed on sample DAS69-5-1 from Wright et al. (1975), which contains 12·2% MgO. Solutions for the Kilauea primary magma composition are given in Table 5 for three assumed source compositions, and we offer in Electronic Appendix 6 on the Journal of Petrology Web site at http://www.petrology.oupjournals.org a detailed model for a fertile peridotite source. Primary Kilauea magmas contain 16·8–24·4% MgO (Table 5), but the upper end is for equilibrium melting, which is unlikely. Olivines that would crystallize from the remaining more likely primary magmas have mg-numbers in the 89·4–91·4 range (Table 5). These are in good agreement with maximum mg-numbers of 90·7 and 91·3 observed for olivines in Kilauea and Mauna Loa lavas, respectively (Norman & Garcia, 1999). Albarède (1992) estimated that olivines with similar mg-numbers precipitated from a primary magma containing 19·2% MgO, in good agreement with our estimates. We note that CaO contents of olivine phenocrysts are high (i.e. 0·20%; Norman & Garcia, 1999), which is consistent with an origin by precipitation from magmas (Table A1, Electronic Appendix 3) and inconsistent with an origin as accidental fragments of the mantle (e.g. Thompson & Gibson, 2000).
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The major difference between models is in the melt fractions. Primary magmas with 16·8–21·3% MgO can be produced by 0·14–0·31 mass fractions of total melting, demonstrating the sensitivity of F to the composition of the source. A melt fraction of 0·15 has been independently estimated from uranium series disequilibria (Sims et al., 1999), and depleted mantle has been identified as one of several isotopic components in the source (Feigenson, 1984).
It has been suggested that the Hawaiian source is lithologically heterogeneous, with eclogite or garnet pyroxenite veins within a peridotite host (Hauri, 1996; Lassiter & Hauri, 1998). This is at variance with conclusions drawn from combined hafnium–neodymium–thorium isotope systematics (Stracke et al., 1999) and Na/Ti systematics (Putirka, 1999). It is also in contrast to our models that yield primary magma from a peridotite source. We do not dismiss the potential importance of subducted crust. Indeed, we discuss in Electronic Appendix 6 that the Kilauea source may have 1–4% by mass of subducted pelagic sediment. The essential question is whether subducted crust identified geochemically occurs also as a distinct lithological entity in the form of eclogite veins, a discussion that is raised in Electronic Appendix 6. If indeed the Hawaiian source region is a lithological mixture of peridotite and eclogite, then our primary magma compositions will be in error. However, we invite the proponents of lithological heterogeneity to offer a model that predicts the precipitation of olivines with mg-numbers of about 91.
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TERTIARY PICRITES AND KOMATIITES FROM THE ICELANDIC PLUME |
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TOPABSTRACTINTRODUCTIONA HYBRID FORWARD AND...MAGMAS IN PROJECTIONCALCULATED MAGMA COMPOSITIONS:...INVERSE MODELSGORGONA KOMATIITES: AN EXAMPLE...HAWAII (KILAUEA)TERTIARY PICRITES AND KOMATIITES...MID-OCEAN RIDGE BASALTSPLUME THERMAL AND PETROLOGICAL...CONCLUSIONSREFERENCES |
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There are now many descriptions of Tertiary picrites and komatiites from Baffin Island and West Greenland (Drever, 1956; Drever & Johnston, 1957; Clarke, 1970; Clarke & Pedersen, 1976; Francis, 1985; Pedersen, 1985; Gill et al., 1992; Larsen et al., 1992; Robillard et al., 1992; Holm et al., 1993; Saunders et al., 1997; Larsen & Pedersen, 2000; Scarrow et al., 2000). Of these, inverse modelling was performed on sample 10014 from Baffin Island, an olivine-phyric lava with 11·85% MgO (Robillard et al., 1992), and sample 113325 from Disko Island in West Greenland, a lava with 10·0% MgO (Holm et al.,1993). Baffin Island lavas contain Fe3+/
Fe = 0·1, based on wet chemistry (Robillard et al., 1992), indicating an fO2 similar to that for Kilauea (i.e. NNO - 1·3). However, Larsen & Pedersen (2000) inferred an fO2 similar to NNO for West Greenland lavas, based on coexisting glass–chromite compositions. Solutions for primary magma compositions are listed in Tables 6 and 7 for various assumed peridotite source compositions, and we offer in Electronic Appendix 7 on the Journal of Petrology Web site at http://www.petrology.oupjournals.org a detailed description for a depleted source. A depleted source is consistent with Sr and light rare earth element depletions for Baffin Island and Disko Island lavas (Robillard et al., 1992; Holm et al., 1993). Primary magmas from depleted and fertile peridotite sources (i.e. with 8·0% FeO) contain 17–20% MgO, in excellent agreement with previous estimates (Francis, 1985; Pedersen, 1985; Holm et al., 1993; Larsen & Pedersen, 2000). The Fe-rich source composition yields primary magmas with MgO as low as 14% (Table 7).
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Baffin Island lavas contained zoned olivine xenocrysts with core–rim mg-numbers of 93–89, although the maximum mg-number of 91·9 seems more common, and Francis (1985) concluded that they must have precipitated from liquids with more than 18% MgO. The interpretations of Francis (1985) are in good agreement with our forward model of a depleted source (Baffin 2F; Table 6), wherein we obtain a primary magma with 17·9% MgO in equilibrium with olivine having an mg-number of 90·6. Olivines that would crystallize from our estimated West Greenland primary magmas also have mg-numbers of around 90·5 (Table 7). Olivines with these compositions are observed in lavas from the Vaigat Formation, but rare olivines with mg-numbers as high as 93·0 also occur (Larsen & Pedersen, 2000), similar to those from Baffin Island (Francis, 1985). We discuss in Electronic Appendix 7 that Ol93·0 could not have precipitated from any reasonable accumulated fractional melt. However, these high mg-number olivines are very similar to those for picrites from Gorgona. In our discussion above, we suggested that the picrites formed as fractional melts from the refractory Gorgona plume core, and a parallel explanation may hold true for the Icelandic plume. The difference, however, is that picritic melts with low FeO contents are not observed in the Icelandic plume lavas. We suggest that olivines with mg-numbers of 93·0 in Icelandic plume lavas may be wall-rock xenocrysts from the refractory plume core, and that the fractional melts that they were in equilibrium with may have been thoroughly mixed in the formation of the aggregate primary magma. Alternatively, they may be phenocrysts of perfect fractional melts that exited the melting regime and mixed with other melts during eruption. Without direct samples of these fractional melts, the MgO content cannot be constrained.
Ni partitioning experiments have been used to infer that primary magmas could not have MgO contents much greater than 13% because they would yield olivine-fractionated derivative liquids with Ni contents that are lower than those observed in lavas (Hart & Davis, 1978; Yaxley, 2000). We would not ordinarily add to the discussion that has followed (Clarke & O’Hara, 1979; Elthon & Ridley, 1979; Francis, 1985; Larsen & Pedersen, 2000), except that the Ni problem has been raised again for West Greenland volcanics (Yaxley, 2000). Contents of NiO and MgO were computed by incremental addition of olivine to lava 113325 (Holm et al., 1993) using MgO-dependent partition coefficients for NiO between olivine and liquid (Beattie et al., 1991; Table 2). Results shown in Fig. 12 demonstrate that a model primary magma with 17·2% MgO and 0·083% NiO yields calculated olivine phenocrysts, derivative liquids and olivine cumulates that are in good agreement with those that are observed in lava flows.
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MID-OCEAN RIDGE BASALTS |
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TOPABSTRACTINTRODUCTIONA HYBRID FORWARD AND...MAGMAS IN PROJECTIONCALCULATED MAGMA COMPOSITIONS:...INVERSE MODELSGORGONA KOMATIITES: AN EXAMPLE...HAWAII (KILAUEA)TERTIARY PICRITES AND KOMATIITES...MID-OCEAN RIDGE BASALTSPLUME THERMAL AND PETROLOGICAL...CONCLUSIONSREFERENCES |
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It is useful to compare primary magmas from plume associations with those from mid-ocean ridges. The only occurrences of olivine-phyric MORB that we are aware of are those from the Reykjanes Peninsula on Iceland (Jakobsson et al., 1978) and those from the Siqueiros transform fault that offsets the East Pacific Rise (Perfit et al., 1996). We restrict this analysis to the Siqueiros MORB because the substantial variability of FeO reported by Jakobsson et al. (1978) on Reykjanes MORB compromises primary magma solutions. The record of olivine as a sole liquidus phase has been erased in all other MORB by gabbro and troctolite crystallization in crustal magma chambers (O’Hara, 1968b). Procedures are available for backtracking these effects (Langmuir et al., 1992), but they are subject to some uncertainty (O’Hara & Herzberg, 2002).
Siqueiros MORB reported by Perfit et al. (1996) are fresh picritic basaltic lavas characterized by high-MgO glasses and 5–20 modal % olivine phenocrysts. Magmas that were primary to the MORB glasses were evaluated by incrementally adding olivine to MORB sample ALV2384-007B. This particular sample is a glass with the highest mg-number, and lowest Na2O and TiO2. Christie et al. (1986) determined Fe3+/Fe = 0·07 ± 0·03 for 78 MORB glasses, and this ratio has been used to calculate FeO and Fe2O3 for the Siqueiros database (Perfit et al., 1996).
Solutions for the primary magma composition are given in Table 8 and discussed in detail in Electronic Appendix 8 on the Journal of Petrology Web site at http://www.petrology.oupjournals.org. Primary magmas for both fertile and depleted peridotite sources are restricted to 11·8–12·9% MgO, melt fractions range from 0·05 to 0·10, and computed olivines have mg-numbers of 89·9–90·5 (Table 8). The Siqueiros primary magma has somewhat more MgO and less FeO than those of N-MORB estimated by McKenzie and coworkers (McKenzie & Bickle, 1988; McKenzie & O’Nions, 1991; White et al., 1992), differences that are discussed in Electronic Appendix 8. Olivines that precipitate from perfect fractional melts or olivine wall-rock xenocrysts from the melting regime are expected to have mg-numbers no higher than about 90·8 (Fig. 7b). Observed olivine phenocrysts contain mg-numbers of 89·5–91·5. The effect of a ±1 uncertainty in KDOl/LFeO/MgO propagates to ±0·5 in olivine mg-number, indicating a complex origin for Ol91·5 observed in the lava flows. Indeed, Perfit et al. (1996) reported that the more magnesian grains show resorption and some of the larger more anhedral grains show deformation lamellae. Some melt inclusions in the olivine phenocrysts are unusual for MORB, and have been interpreted as examples of assimilated gabbroic cumulates + trapped liquid (L. V. Danyushevsky, personal communication, 2002). As many gabbroic cumulates are low in FeO, assimilation can potentially elevate the mg-number of the liquid and olivine precipitates. Localized crustal assimilation might explain why our computed olivine mg-numbers are somewhat lower than those in phenocrysts.
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PLUME THERMAL AND PETROLOGICAL STRUCTURE |
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TOPABSTRACTINTRODUCTIONA HYBRID FORWARD AND...MAGMAS IN PROJECTIONCALCULATED MAGMA COMPOSITIONS:...INVERSE MODELSGORGONA KOMATIITES: AN EXAMPLE...HAWAII (KILAUEA)TERTIARY PICRITES AND KOMATIITES...MID-OCEAN RIDGE BASALTSPLUME THERMAL AND PETROLOGICAL...CONCLUSIONSREFERENCES |
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Primary magma temperatures have been calculated with equation (30) for liquids in equilibrium with olivine. Results have been provided in Tables 4–8 at 1 atm and 4 GPa. The MgO content of the liquid has the most important effect on olivine–liquid equilibration temperature, and other oxide components have significantly lesser effects. The olivine–liquid equilibration temperature of a primary magma with a specific MgO content is therefore very similar to that for a liquid in equilibrium with fertile or depleted mantle peridotite at the same MgO content. And as the MgO contents of primary magmas do not change significantly during decompression melting (Fig. 8a), we can use the MgO contents of liquids in equilibrium with mantle peridotite as a proxy for the decompression melting T–P path. These are shown in Fig. 13.
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The intersection of the decompression melting T–P path with the anhydrous solidus can place constraints on the thermal characteristics of the melting regime as defined by the potential temperature (Tp; McKenzie & Bickle, 1988) and the pressure of initial melting on the anhydrous solidus (i.e. Po). For the case of a primary MORB magma having 12% MgO (Table 8), the eruption temperature is about 1275°C, Po = 2·5 GPa and Tp = 1380°C (Fig. 13). These conditions are similar to those inferred by Asimow et al. (2001) for the production of 7 km of oceanic crust, but they are hotter than the 1300°C potential temperatures preferred by McKenzie and coworkers (McKenzie & Bickle, 1988; McKenzie & O’Nions, 1991; White et al., 1992; Electronic Appendices 8 and 9 on the Journal of Petrology Web site at http://www.petrology.oupjournals.org). Primary magmas from Hawaii, Gorgona, Baffin Island and West Greenland with 18–20% MgO have eruption temperatures of about 1400–1440°C, intersect the solidus for fertile peridotite at 3·8–4·7 GPa (Fig. 13), and have potential temperatures of 1520–1570°C.
For Gorgona primary picrites with about 24% MgO we obtain Po = 8·0 GPa, Tp = 1700°C, similar to those inferred by Thompson & Gibson (2000). However, the MgO content is not well constrained for fractional melts in equilibrium with Ol93·6, and it can range from 23 to 26% (Fig. 11). When the MgO content of a primary fractional melt exceeds about 24% there can be very large uncertainties in both pressure of initial melting and the potential temperature because 
T/P for isopleths for MgO are nearly parallel to those for the solidus (Fig. 13). Similarly, changes in peridotite source composition can affect solidus absolute temperature and estimates of initial melting pressures (Electronic Appendix 9), although detailed experimental calibrations are lacking. These problems can have important consequences for interpretations of Archaean komatiites with 20–30% MgO, a discussion is that is well beyond the scope of this paper.
Pressures of melting can be significantly deeper than lithosphere thickness (Fig. 13), consistent with the plume model. The formation of hotspots by shallow passive melting in response to stresses and cracks in the lithosphere is an important alternative (Anderson, 2000), but it is not consistent with topography on the 410 and 660 km discontinuities below Hawaii and Iceland (Shen et al., 1998; Li et al., 2000).
Picrites from Baffin Island and West Greenland occupy the western peripheral edge of eruption sites with 58–62 Ma ages scattered over 2000 km of the North Atlantic Igneous Province (White & McKenzie, 1989; Saunders et al., 1997), a distribution that White & McKenzie (1989) related to the ancestral Icelandic plume. Computed plume simulations show considerable flattening of a plume head as it impacts the lithosphere, and that the plume head can stay hot over large horizontal distances (Larsen et al., 1999). Picrites and basalts from the British Tertiary Igneous Province are also likely to have been extracted from the eastern Icelandic plume periphery. Many are olivine-phyric lavas with FeOT contents that are higher than those for normal MORB (i.e. 10–13% FeOT; Scarrow et al., 2000), and they have elevated 3He/4He (Stuart et al., 2000). Similarly, the distribution of 88–90 Ma komatiite and picritic lavas throughout the Caribbean–Colombian plateau has been interpreted as eruption from a large plume head (Kerr et al., 1997). Our preferred model, shown in Fig. 14, has ultramafic magma production distributed throughout a large plume head within which liquids are extracted from a harzburgite residue (Herzberg & O’Hara, 1998), whereas it is restricted to the plume axis in the models of Campbell et al. (1989) and Campbell & Griffiths (1990).
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Model petrological plume structures are shown in Fig. 14. Both plume heads and axes can consist of a harzburgite core [L + Ol + Opx] and a garnet peridotite periphery [L + Ol ± Opx + Cpx + Gt] (Eggins, 1992b; Herzberg & O’Hara, 1998). The Gorgona and early Tertiary Icelandic plumes differ from Hawaii in consisting of more head than tail. They also had a deeper and hotter plume core that contained advanced fractional melts in equilibrium with olivines having mg-numbers of 93·0–93·6. Fractional melts from the plume core may not have mixed with fractional melts from the plume periphery if melt percolation streamlines were dominantly vertical. Indeed, isotopic heterogeneity observed in individual shield volcanoes in Hawaii (Lassiter & Hauri, 1998; Blichert-Toft et al., 1999) indicates that lateral melt movement is not signficant in a plume axis. The situation may be complex when material in the axis feeds a flattened plume head as it impacts the lithosphere (Fig. 14).
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CONCLUSIONS |
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TOPABSTRACTINTRODUCTIONA HYBRID FORWARD AND...MAGMAS IN PROJECTIONCALCULATED MAGMA COMPOSITIONS:...INVERSE MODELSGORGONA KOMATIITES: AN EXAMPLE...HAWAII (KILAUEA)TERTIARY PICRITES AND KOMATIITES...MID-OCEAN RIDGE BASALTSPLUME THERMAL AND PETROLOGICAL...CONCLUSIONSREFERENCES |
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A parameterization of a large experimental database provides forward models of potential primary magma compositions that exit the melting regime in the mantle, and these are represented in CMAS projections and plots of FeO–MgO. The effects of equilibrium melting, fractional melting and peridotite source composition on FeO and MgO contents of potential primary magmas have been evaluated. The parameterization can also be used in inverse models for computing the effects of olivine fractionation for any olivine-phyric lava suite. A hybrid forward and inverse model provides a unique primary magma composition, eruption temperature, melt fraction and residuum mineralogy for any assumed peridotite source composition and melting mechanism.
The major element geochemistry of primary magmas from Gorgona, Hawaii, Baffin Island and West Greenland cannot be used to infer major element composition of the source. The MgO contents of primary magmas are typically 18–20% MgO for wide variations in assumed fertile and depleted peridotite source compositions, and we estimate in Electronic Appendix 10 on the Journal of Petrology Web site (http://www.petrology.oupjournals.org) an uncertainty of ±1% MgO at the 1 level of confidence. However, MgO can drop to 14–17% for Fe-enriched sources, and increase to 26% for fractional melts from Gorgona. One serious geological casualty is the large uncertainty in crustal thickness because different source compositions predict melt fractions that range from about 0·1 to 0·4. Another casualty is the large uncertainty in estimates of the pressure of melt collection (Electronic Appendix 10).
There is a substantial drop in the FeO content of both fractional melts and accumulated fractional melts during adiabatic decompression melting. However, there is little change in MgO because adiabatic T–P paths are nearly coincident with olivine–liquid T–P saturation surfaces. Olivine–liquid equilibrium temperatures therefore closely reflect actual eruption temperatures and adiabatic decompression melting paths.
Primary magmas from Hawaii, Gorgona, Baffin Island and West Greenland with 18–20% MgO have eruption temperatures of 1400–1440°C, intersect the solidus for an assumed fertile peridotite composition at 3·8–4·7 GPa, and have potential temperatures of 1520–1570°C. For Gorgona picrites with 24% MgO, the potential temperature and initial melting pressure were about 1700°C and 8·0 GPa, respectively; melting was hot and deep, consistent with the plume model.
Our work provides strong support for the importance of fractional melting in plumes (Arndt et al., 1997). Primary magma compositions that we infer for accumulated fractional melting are in equilibrium with olivines with mg-numbers that are typically around 90·6, similar to observed maximum mg-numbers of olivines in lavas. Olivines with mg-numbers of 93·0–93·6 in lavas from Gorgona, Baffin Island and West Greenland could not have precipitated as phenocrysts from accumulated fractional melts or from equilibrium melts. Instead, they could have been in equilibrium with isolated advanced fractional melts having low FeO contents and MgO contents that are poorly constrained. In general, olivines with mg-numbers of 93·0–93·6 may be phenocrysts of perfect fractional melts that exited the melting regime or they may be wall-rock xenocrysts from the refractory plume core. The occurrence of olivines in lava flows with compositions that are expected from both well-mixed aggregate fractional melts and advanced fractional melts indicates that magma mixing in the melting regime of mantle plumes does not always occur, a conclusion reached independently by Slater et al. (2001) for Iceland.
The picrites from Gorgona are interpreted as advanced fractional melts that formed in the hot axis of a plume, a suggestion that was made previously by Arndt et al. (1997). Fractional melts with more than 24% MgO can have thermal characteristics that are difficult to constrain. A typical petrological plume structure consists of a harzburgite core [L + Ol + Opx] and a garnet peridotite periphery [L + Ol ± Opx + Cpx + Gt] located in both the head and axis.
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ACKNOWLEDGEMENTS |
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We are very grateful to N. Arndt, P. Asimow, L. Larsen, R. N. Thompson and M. Wilson for thorough and constructive reviews.
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FOOTNOTES |
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*Corresponding author. E-mail: Herzberg{at}rci.rutgers.edu
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REFERENCES |
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TOPABSTRACTINTRODUCTIONA HYBRID FORWARD AND...MAGMAS IN PROJECTIONCALCULATED MAGMA COMPOSITIONS:...INVERSE MODELSGORGONA KOMATIITES: AN EXAMPLE...HAWAII (KILAUEA)TERTIARY PICRITES AND KOMATIITES...MID-OCEAN RIDGE BASALTSPLUME THERMAL AND PETROLOGICAL...CONCLUSIONSREFERENCES |
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