Journal of Petrology Advance Access originally published online on December 7, 2005
Journal of Petrology 2006 47(3):505-539; doi:10.1093/petrology/egi084
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The Genesis of Intermediate and Silicic Magmas in Deep Crustal Hot Zones
1 SECTION DES SCIENCES DE LA TERRE, UNIVERSITÉ DE GENÈVE, 13 RUE DES MARAÎCHERS, 1205 GENÈVE, SWITZERLAND
2 DEPARTMENT OF EARTH SCIENCES, UNIVERSITY OF BRISTOL, WILLS MEMORIAL BUILDING, BRISTOL BS8 1RJ, UK
RECEIVED APRIL 14, 2005; ACCEPTED OCTOBER 17, 2005
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ABSTRACT |
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TOP
ABSTRACT
INTRODUCTIONA model for the generation of intermediate and silicic igneous rocks is presented, based on experimental data and numerical modelling. The model is directed at subduction-related magmatism, but has general applicability to magmas generated in other plate tectonic settings, including continental rift zones. In the model mantle-derived hydrous basalts emplaced as a succession of sills into the lower crust generate a deep crustal hot zone. Numerical modelling of the hot zone shows that melts are generated from two distinct sources; partial crystallization of basalt sills to produce residual H2O-rich melts; and partial melting of pre-existing crustal rocks. Incubation times between the injection of the first sill and generation of residual melts from basalt crystallization are controlled by the initial geotherm, the magma input rate and the emplacement depth. After this incubation period, the melt fraction and composition of residual melts are controlled by the temperature of the crust into which the basalt is intruded. Heat and H2O transfer from the crystallizing basalt promote partial melting of the surrounding crust, which can include meta-sedimentary and meta-igneous basement rocks and earlier basalt intrusions. Mixing of residual and crustal partial melts leads to diversity in isotope and trace element chemistry. Hot zone melts are H2O-rich. Consequently, they have low viscosity and density, and can readily detach from their source and ascend rapidly. In the case of adiabatic ascent the magma attains a super-liquidus state, because of the relative slopes of the adiabat and the liquidus. This leads to resorption of any entrained crystals or country rock xenoliths. Crystallization begins only when the ascending magma intersects its H2O-saturated liquidus at shallow depths. Decompression and degassing are the driving forces behind crystallization, which takes place at shallow depth on timescales of decades or less. Degassing and crystallization at shallow depth lead to large increases in viscosity and stalling of the magma to form volcano-feeding magma chambers and shallow plutons. It is proposed that chemical diversity in arc magmas is largely acquired in the lower crust, whereas textural diversity is related to shallow-level crystallization.
KEY WORDS: magma genesis; deep hot zone; residual melt; partial melt; adiabatic ascent
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INTRODUCTION |
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A key question in igneous petrology concerns the origin of intermediate to silicic magmatic rocks, such as voluminous Cordilleran granite batholiths (diorites, tonalites, granodiorites and granites) and the evolved volcanic rocks (andesites, dacites and rhyolites) of destructive plate margins. The continental crust has an estimated silicic andesite to dacite composition, with a vertical stratification from mafic lower crust to more evolved granite-dominated upper crust (Rudnick & Fountain, 1995
). The origin of intermediate to silicic igneous rocks is, therefore, central to understanding the evolution of the continental crust.
In subduction settings melt is generated by partial melting in the mantle wedge where primary mafic magmas form by some combination of addition of H2O-rich fluids or melts released from the subducted slab (e.g. Davies & Stevenson, 1992; Tatsumi & Eggins, 1995; Schmidt & Poli, 1998; Ulmer, 2001; Grove et al., 2002; Forneris & Holloway, 2003) and mantle decompression resulting from subduction-induced corner flow (e.g. Sisson & Bronto, 1998; Elkins-Tanton et al., 2001; Hasegawa & Nakajima, 2004). Experimental studies of mantle melting (e.g. Ulmer, 2001; Parman & Grove, 2004; Wood, 2004), and observations of the petrology and geochemistry of mafic arc magmas, indicate that primary, mantle-derived magmas range in composition from basalts to magnesian andesites (Tatsumi, 1982; Tatsumi & Eggins, 1995; Bacon et al., 1997; Conrey et al., 1997; Carmichael, 2002, 2004; Grove et al., 2002). In terms of liquidus temperatures and dissolved H2O contents there is a range from dry and hot magmas to wet and cool varieties, even within a single volcanic arc (e.g. Sisson & Layne, 1993; Baker et al., 1994; Elkins-Tanton et al., 2001; Pichavant et al., 2002a). Volcanic rocks with MgO-rich compositions that could be in equilibrium with the mantle wedge are rare in continental arcs and only a minor component of island arcs, an observation attributable to density filtering and intracrustal ‘processing’ of ascending magmas. This processing accounts for the predominance of evolved (silica-rich) volcanic rocks and granitic plutonic rocks in continental and mature island arcs.
The generation of intermediate and silicic arc magmas is widely attributed to two main processes: differentiation of primary magmas by crystallization within the crust or uppermost mantle (e.g. Gill, 1981; Grove & Kinzler, 1986; Musselwhite et al., 1989; Rogers & Hawkesworth, 1989; Müntener et al., 2001; Grove et al., 2002, 2003) and partial melting of older crustal rocks (e.g. Smith & Leeman, 1987; Atherton & Petford, 1993; Tepper et al., 1993; Rapp & Watson, 1995; Petford & Atherton, 1996; Chappell & White, 2001; Izebekov et al., 2004). These processes can occur simultaneously with the heat and volatiles (principally H2O) liberated from the primary magmas triggering crustal melting (Petford & Gallagher, 2001; Annen & Sparks, 2002). Additionally crustal rocks can be assimilated into mantle-derived magmas (DePaolo, 1981). The assimilated components may be much older than, and petrogenetically unrelated to, the arc magmas and possess distinctive trace element and isotope geochemistry. Partial melting can also occur in igneous rocks, including cumulates, that have formed from earlier arc magmas; in this case the assimilated components and arc magmas may have strong geochemical affinities (e.g. Heath et al., 1998; Dungan & Davidson, 2004). Evidence for crustal assimilation and mixing of melts and crystals from different sources is common (Grove et al., 1988, 1997; Musselwhite et al., 1989; De Paolo et al., 1992). These processes are central to models of assimilation and fractional crystallization (AFC; DePaolo, 1981) and mixing, assimilation, storage and hybridization (MASH; Hildreth & Moorbath, 1988).
A key question is at what depth chemical differentiation occurs. Although the existence of shallow sub-volcanic magma chambers is indisputable, based on geophysical evidence as well as petrological and geological observations, it is less clear that such chambers are the place where most chemical differentiation takes place. To produce igneous rocks that contain more than 60 wt % SiO2 by fractional crystallization, 60% or more crystallization of a typical primitive arc basalt is required (e.g. Foden & Green, 1992; Müntener et al., 2001). The volume of parental mafic magma that crystallizes is, therefore, typically twice as much as the evolved magma produced. As large granitoid batholiths and voluminous eruptions involve hundreds to thousands of km3 of silicic magma (e.g. Smith, 1979; Crisp, 1984; Bachmann et al., 2002), huge volumes of associated mafic cumulates are required. However, geological and geophysical evidence for the requisite large volumes of complementary dense mafic cumulates in the shallow crust is generally lacking. One resolution to this problem is density-driven sinking of mafic cumulate bodies into the lower crust (Glazner, 1994). Alternatively, if differentiation of basalt occurs at deep levels in the crust then the complementary dense mafic cumulates will be located in the lower crust (e.g. Debari & Coleman, 1989; Müntener et al., 2001) where they may eventually delaminate into the mantle below (Kay & Kay, 1993; Jull & Keleman, 2001) thereby progressively driving the bulk crust towards andesite composition. The silica-rich residual melts generated by deep-seated basalt differentiation can be extracted and ascend, either to erupt immediately or to stall to form shallow magma chambers. If unerupted, such shallow chambers consolidate to form granite plutons, with mafic igneous rocks being a minor component or absent.
Recent numerical simulations of heat transfer (Annen & Sparks, 2002) and high-temperature experiments (Müntener et al., 2001; Prouteau & Scaillet, 2003) suggest a model whereby silica-rich magmas can be generated by incomplete crystallization of hydrous basalt at upper mantle and/or lower crustal depths. These observations motivate our development of a model in which basalt emplacement into the lower crust leads to generation of intermediate and silicic melts (Fig. 1). Our model builds upon the concept of underplating (Raia & Spera, 1997), expands on models of differentiation of basalt at high pressure (Gill, 1981; Grove et al., 2002) and incorporates aspects of AFC (DePaolo, 1981) and MASH (Hildreth & Moorbath, 1988). We develop a quantitative model in which evolved melts are generated from H2O-rich parental basalts both by partial crystallization of the basalts themselves and by partial melting of surrounding crustal rocks through heat and H2O transfer from the cooling basalts. A key feature of our model is that melt compositions are determined by the depth of emplacement of individual basalt intrusions and thermal equilibration with the local geotherm. We refer to the site of basalt injection and melt generation in the lower crust as a deep crustal hot zone. Previous models of underplating (e.g. Huppert & Sparks, 1988; Bergantz, 1989; Raia & Spera, 1997; Petford & Gallagher 2001; Jackson et al., 2003) have concentrated almost exclusively on melt generated by heating of the crust, with less attention paid to the residual melt generated by partial crystallization of the underplated basalt intrusions. Here we develop the concepts proposed by Annen & Sparks (2002) and consider the full range of possible mechanisms of melt generation in the hot zone, including residual melt from basalt crystallization and partial melting of surrounding crustal rocks (Fig. 1). We then consider the evolution of these melts as they are extracted from their source rocks and ascend to shallow crustal levels, degassing and crystallizing en route. The model is developed primarily for application to the genesis of subduction zone volcanic and plutonic rocks, and we will refer collectively to this whole suite of intermediate and silicic rock types as ‘andesite’, except where a compositional or textural distinction is relevant. However, our model has general applicability to other tectonic settings, including continental rift zones where plume-related basaltic magmas are intruded into the base of the continental crust.
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SOURCES AND MECHANISMS FOR INTERMEDIATE AND SILICIC MAGMA GENERATION |
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There are five currently popular models for the generation of andesites (sensu lato), as follows.
Model I. Partial melting of harzburgite in the mantle wedge, fluxed by H2O-rich fluids or melts liberated from the subducting slab (e.g. Tatsumi, 1982; Hirose, 1997; Blatter & Carmichael, 2001; Carmichael, 2002, 2004; Parman & Grove, 2004).
Model II. Crystallization of mantle-derived basalt or basaltic andesite in shallow crustal magma chambers (e.g. Sisson & Grove, 1993; Grove et al., 1997; Pichavant et al., 2002b).
Model III. Crystallization of mantle-derived basalt or basaltic andesite in the deep arc crust at or close to the Moho (e.g. Müntener et al., 2001; Annen & Sparks, 2002; Mortazavi & Sparks, 2003; Prouteau & Scaillet, 2003).
Model IV. Dehydration partial melting of meta-basalts (amphibolites) in the lower or middle crust by intrusions of hot, mantle-derived magma (e.g. Smith & Leeman, 1987; Petford & Atherton, 1996; Jackson et al., 2003).
Model V. Mixing between silicic magmas and mantle-derived mafic magmas (e.g. Heiken and Eichelberger, 1980). In some cases the silicic component is generated by partial melting of crustal rocks (e.g. Druitt et al., 1999).
In this paper we focus on Models III–V, which take place in the middle or lower crust. Models I and II are briefly considered first. Generation of andesite by mantle melting (Model I) has been demonstrated experimentally (Tatsumi, 1982; Hirose, 1997; Grove et al., 2002, 2003; Parman & Grove, 2004) and calculated thermodynamically (Carmichael, 2002, 2004). The andesites produced in this way have elevated MgO contents and high mg-numbers, a requirement for equilibrium with the Mg-rich olivines of mantle harzburgite. Boninite series magmas are widely thought to originate by H2O-fluxed melting of harzburgite (Falloon & Danyushevsky, 2000; Parman & Grove, 2004), whereas the generation of ‘high-Mg andesites’ may involve reactions between ascending slab-derived silicic melts and mantle peridotite (Yogodzinsky & Kelemen, 1998). However, high-Mg andesites and boninites are not the dominant rock types of volcanic arcs; typical arc andesites, with low mg-numbers, could not have been in direct equilibrium with mantle rocks.
Model II is widely favoured. Basalt and basaltic andesite lavas occur at many arc stratovolcanoes and occasionally contain xenoliths of cumulate origin (e.g. Arculus & Wills, 1980). Several experimental studies demonstrate that andesite can be generated by fractional crystallization of H2O-saturated basalts and basaltic andesites at pH2O = Ptot of 200–400 MPa and temperatures of 950–1050°C (Sisson & Grove, 1993; Grove et al., 1999, 2003; Pichavant et al., 2002b) by crystallizing an assemblage of plagioclase (An60–90) + clinopyroxene + amphibole + oxides ± orthopyroxene ± olivine. One constraint on the origin of andesites is that they typically contain <19% Al2O3 (Fig. 2), indicating that by the time residual melts have attained >57 wt % SiO2 they have become saturated in an aluminous phase. In Model II crystallization of plagioclase serves to limit Al2O3 enrichment in residual melts. The lack of abundant dense complementary mafic to ultramafic cumulate rocks in the shallow crust is problematic for Model II unless the associated mafic cumulates are removed by sinking (Glazner, 1994).
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Model III involves fractional crystallization of similar parental magmas to Model II, but at higher pressure, thereby obviating the problem of the missing mid- or upper-crustal mafic cumulates. Mantle-derived magmas intruded into the deep crust cool and crystallize producing evolved residual melts. The principal difference between high- and low-pressure crystallization of hydrous basalt lies in the nature of the crystallizing assemblage. At higher pH2O garnet (e.g. Wolf & Wyllie, 1994
; Rapp, 1995) and aluminous amphibole (Grove et al., 2003) are stabilized and can contribute to minimizing Al2O3 enrichment in residual melts. Conversely, plagioclase stability is reduced and liquidus plagioclase is anorthite-rich, a common finding in arc-related cumulate nodules (e.g. Arculus & Wills, 1980). In terms of melt chemistry, it is very hard to distinguish between residual melts produced by crystallization of An-rich plagioclase and pyroxenes from H2O-undersaturated basalt at 
1·0 GPa (Kawamoto, 1996) and those produced from H2O-saturated basalt at 0·2–0·4 GPa (e.g. Sisson & Grove, 1993; Pichavant et al., 2002b). The appearance of garnet as the liquidus aluminous phase in andesite and dacite melts at pressures over 1·1 GPa (Wolf & Wyllie, 1994; Rapp, 1995) imparts a distinctive trace element chemistry to residual melts (e.g. high Sr/Y), which provides a clear indication of high-pressure differentiation (e.g. Smith & Leeman, 1987; Feeley & Davidson, 1994; Feeley & Hacker, 1995).
In Models II and III, Al2O3 enrichment in derivative melts is further minimized if the primitive basalt itself has relatively low Al2O3. Circumstances for generation of such magmas are inferred in many arcs with a relatively depleted mantle wedge (Grove et al., 2003; Parman & Grove, 2004). For example, primitive arc basalts with only 14–15% Al2O3 have been described for Klyuchevskoy volcano, Kamchatka (Ozerov, 2000).
When mafic magmas are intruded into the arc crust they transfer heat and volatiles (principally H2O) into the surrounding crust, which can lead to partial melting of the wall-rocks. The deep crustal hot zone is, therefore, envisaged as a mixture of partially crystallized basalt, partially molten crustal rocks and H2O liberated from the solidifying basalts (Fig. 1). Geophysical evidence is consistent with these concepts. In the Cascades, for example, the release of significant volumes of H2O from deeply intruded basalts may account for the presence of a highly electrically conductive layer at 10–30 km depth (Stanley et al., 1990), and in the central Andes a broad conductive zone (Brasse et al., 2002) is associated with a low-velocity zone at depths of 20–40 km (Yuan et al., 2000), interpreted as a laterally extensive region of partial melt, capped by a silicic magma body 1 km thick (Chmielowski et al., 1999). Below volcanoes in the Japan arc broadband seismometers have recorded low-frequency tremors and micro-earthquakes at 30–50 km depth (Obara, 2002; Katsumata & Kamaya, 2003). These can be explained by deformation associated with magma intrusions (S. Sachs, personal communication, 2003) and their low frequency is consistent with the presence of a fluid phase. Finally, beneath central North Island, New Zealand, a seismically highly reflective layer at 35 km depth, interpreted as a body of partially molten rock (Stratford & Stern, 2004), suggests that beneath some arcs the hot zone may be located in the uppermost mantle, rather than within the crust, which is only 16 km thick in this region.
The partially molten crust surrounding the basalt may be older intrusions of related mantle-derived hydrous basalt (or amphibolite) or unrelated metamorphic arc crust. This is Model IV. The volume and composition of the partial melt produced depends on the intrusion rate (heat flux) of the mantle-derived basalts, the prevailing geotherm and the extent to which the melting region is fluxed by H2O liberated from the crystallizing basalt. Chemically hybrid melts can be formed if the residual melts from basalt crystallization are mixed with crustal partial melts during extraction, ascent and shallow intrusion; this is Model V.
Models III and IV both involve partially molten hydrous basaltic rocks in the lower crust produced, respectively, by crystallization and melting. Deep-seated crystallization of hydrous basaltic magmas differs from dehydration melting of the lower crust, as modelled by Raia & Spera (1997), Petford & Gallagher (2001) and Jackson et al. (2003), in one fundamental regard, the availability of H2O. In dehydration melting the H2O content of the source rock is strictly limited by the amount of H2O that can be structurally bound in hydrous minerals such as amphibole and mica. For a mafic amphibolite with 40% amphibole, this amounts to 0·8 wt % H2O. Greater quantities of H2O can be involved only if the heat source efficiently fluxes the source region with H2O. Although this is likely, no extant models of crustal melting consider this process, largely because it is uncertain whether H2O passing through a low-porosity source rock triggers melting or is simply carried away along fractures. By contrast, deep-seated crystallization of hydrous arc basalt magmas has no such upper limit on H2O content. Studies of melt inclusions in primitive arc magmas, together with high-pressure experiments, indicate dissolved H2O contents from almost zero to 10 wt % (e.g. Sisson & Layne, 1993; Carmichael, 2002; Pichavant et al., 2002a; Grove et al., 2003). The wide range of H2O contents and bulk compositions of parental arc basalts ensures that crystallization of hydrous basalt can generate a wide diversity of residual melt compositions, as demonstrated experimentally by Sisson et al. (2005).
Dehydration melting (Model IV) requires a heat source. In arcs the widespread association of evolved igneous rocks with mantle-derived basalt strongly suggests that mafic magmas provide the heat source (Hildreth, 1981). However, herein lies a problem: models of heat transfer show that arc basalts emplaced into the base of the crust at temperatures of 1100–1240°C (see Ulmer, 2001; Pichavant et al., 2002a) cannot provide enough heat to melt amphibolite lower crust extensively (Petford & Gallagher 2001; Annen & Sparks, 2002), because of the high dehydration melting temperature of amphiboles in mafic rocks (950°C). More fertile upper crustal pelitic protoliths can be melted more efficiently, but large amounts of basalt are still needed as a heat source (Annen & Sparks, 2002). In addition, silicic rocks in arcs are typically calc-alkaline and metaluminous, which places limits on the amount of pelite that can be melted. The isotopic and geochemical signatures of evolved plutonic and volcanic arc rocks clearly indicate contribution from pelitic crust in some cases (DePaolo et al., 1992), but significant amounts of basalt or meta-basalt (amphibolite) must be involved in their petrogenesis. The problem in arcs is how to generate large volumes of metaluminous, calc-alkaline evolved melts when the proposed amphibolite source is too refractory to undergo significant dehydration melting at plausible temperatures. This paradox can be solved if crystallization of H2O-bearing mantle-derived basalt is the principal source of the evolved melts.
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CRYSTALLIZATION OF ANDESITE IN THE SHALLOW CRUST |
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Once generated in the deep crust andesite and dacite residual melts can detach and ascend into the shallow crust. Subduction-related andesites and dacites are commonly porphyritic, with phenocrysts of plagioclase plus various proportions of hornblende, clinopyroxene, orthopyroxene, biotite and oxides; the exact ferromagnesian assemblage depends on magma composition, partial pressure of volatiles (especially pH2O), oxygen fugacity (fO2) and temperature (e.g. Rutherford et al., 1985
; Rutherford & Devine, 1988; Blatter & Carmichael, 1998, 2001; Moore & Carmichael, 1998; Scaillet & Evans, 1999; Pichavant et al., 2002b; Izebekov et al., 2004). Invariably the groundmass or matrix glass in porphyritic andesites and dacites is rhyolitic in composition. The phenocryst assemblages commonly have complex textures and zoning patterns, which indicate that magmatic evolution can involve processes such as: repeated mixing of different batches of magma (e.g. Heiken & Eichelberger, 1980; Clynne, 1999); entrainment of old crystals from previously consolidated magma batches (Davidson et al., 1998, 2001, 2005; Heath et al., 1998; Cooper & Reid, 2003; Reagan et al., 2003; Dungan & Davidson, 2004) or from assimilation of crustal rocks (Ferrara et al., 1989); convective stirring (Couch et al., 2001); crystal growth induced by degassing (Blundy & Cashman, 2001). Whereas some of these phenocrysts grew from the magma in which they are found, others are entrained xenocrysts from earlier magma pulses or from chemically unrelated wall-rocks (e.g. Izebekov et al., 2004; Davidson et al., 2005). Detailed studies of volcano evolution (e.g. Bacon, 1983; Bacon & Druitt, 1988; Druitt & Bacon, 1989; Harford et al., 2002) and constraints on timescales for crystallization (e.g. Zellmer et al., 1999, 2003a, 2003b; Harford & Sparks, 2001) suggest that these various processes are the consequence of amalgamation of shallow magma bodies in the upper crust through many episodes of magma ascent from greater depths, sometimes accompanied by eruption. Field and geochronological evidence from calc-alkaline plutonic rocks (‘granites’, sensu lato) also supports their formation by amalgamation of many small intrusions, often of magmas with very similar bulk chemical composition but subtle textural differences (e.g. John & Blundy, 1993) or radiometric ages (e.g. Coleman et al., 2004; Glazner et al., 2004).
Our main concern here is to establish under what conditions the common phenocryst assemblages in andesites and granites are formed. Central to this issue are the H2O contents and temperatures of andesite magmas. The importance of these two variables in interpreting the phenocryst assemblages and compositions of andesites has been investigated for over 30 years in a large number of experimental studies at pH2O (
Ptot) of 0·1 to 
400 MPa (Eggler, 1972; Green, 1972; Eggler & Burnham, 1973; Maksimov et al., 1978; Sekine et al., 1979; Rutherford et al., 1985; Rutherford & Devine, 1988, 2003; Luhr, 1990; Foden & Green, 1992; Sekine & Aramaki, 1992; Sisson & Grove, 1993; Kawamoto, 1996; Grove et al., 1997, 2003; Barclay et al., 1998; Blatter & Carmichael, 1998, 2001; Moore & Carmichael, 1998; Cottrell et al., 1999; Martel et al., 1999; Sato et al., 1999; Scaillet & Evans, 1999; Pichavant et al., 2002b; Couch et al., 2003; Prouteau & Scaillet, 2003; Barclay & Carmichael, 2004; Costa et al., 2004; Izebekov et al., 2004). Although many of these studies are focused on rocks from a specific volcano, some general conclusions can be drawn regarding subduction-related andesites and dacites, as follows.
(1) Eruption temperatures, as determined by geothermometry, are consistently less than low-pressure (<300 MPa) andesite liquidus temperatures even under H2O-saturated conditions. In many cases the difference is several tens of degrees and can be as much as 200°C (e.g. Blatter & Carmichael, 1998; Barclay & Carmichael, 2004).
(2) The liquidus phases at low pH2O often include minerals (e.g. olivine, clinopyroxene) that are absent from the phenocryst assemblage in the natural rocks (e.g. Blatter & Carmichael, 1998, 2001; Scaillet & Evans, 1999; Costa et al., 2004).
(3) Although amphibole is a common phenocryst it rarely occurs on the andesite liquidus at pH2O <400 MPa even under oxidizing conditions; where it is stable, amphibole typically appears 100°C below the liquidus (e.g. Rutherford & Devine, 1988, 2003; Blatter & Carmichael, 1998, 2001; Moore & Carmichael, 1998; Martel et al., 1999; Costa et al., 2004; Izebekov et al., 2004).
(4) Plagioclase is stabilized only at low pH2O and is rarely a true liquidus phase at pH2O >100–200 MPa, even though plagioclase is a ubiquitous phenocryst phase in most andesites (e.g. Eggler, 1972; Maksimov et al., 1978; Sekine et al., 1979; Sekine & Aramaki, 1992; Blatter & Carmichael, 1998, 2001; Moore & Carmichael, 1998; Martel et al., 1999; Grove et al., 2003).
(5) The anorthite (An) content of plagioclase increases with increasing pH2O (at constant temperature) and increasing temperature (at constant pH2O). For a given andesite, plagioclase phenocryst rims typically have considerably lower An contents (by up to 30 mol %) than the experimentally determined liquidus or near-liquidus plagioclase (e.g. Rutherford et al., 1985; Scaillet & Evans, 1999; Rutherford & Devine, 2003; Costa et al., 2004).
(6) The observed phenocryst assemblage, phase compositions and crystallinity typically are consistent with H2O-saturated conditions at pressures of 100–300 MPa and at sub-liquidus temperatures consistent with those obtained from mineral thermometry on the natural rocks (e.g. Blatter & Carmichael, 1998; Moore & Carmichael, 1998; Martel et al., 1999; Costa et al., 2004).
The similarity of phase proportions and compositions in both experiments and natural andesites (Fig. 3) indicates that, to a first approximation, these magmas have undergone near-closed system crystallization from an initial fully molten state to a porphyritic magma under conditions of low-pressure H2O-saturation. However, very few of the studied andesites contain their full complement of experimentally determined liquidus phases under these conditions, suggesting that either magma temperatures were never high enough to form a fully molten andesite liquid at low pressure or that the original liquidus phases were completely eliminated (or re-equilibrated) by reaction with the melt. The interpretation we favour is that andesite liquids, once formed and extracted from the deep crust, typically crystallize under polybaric conditions, at temperatures that do not significantly exceed their eruption temperature. Thus the initial fully molten state of an andesite is not a consequence of high temperature, but a consequence of high pH2O. All of the above observations [(1)–(6)] are consistent with this interpretation, as are the observed zoning patterns and rim compositions of plagioclase phenocrysts (Fig. 3). For example, the phenocryst assemblage and proportions of the Colima andesite (Fig. 3a) can be reproduced closely at 950–960°C (consistent with mineral thermometry on the natural lava) and pH2O from 70 to 150 MPa (Moore & Carmichael, 1998). The very calcic cores of some plagioclase phenocrysts (An85) were ascribed by Moore & Carmichael (1998) to the onset of crystallization at even higher pH2O but at essentially the same temperature. Using analyses of phenocryst-hosted melt inclusions, Blundy & Cashman (2005) advanced a similar argument for the silicic andesites of Mount St. Helens. They proposed that the observed phenocryst assemblage of the white pumice of 18 May 1980 crystallized in response to decompression from 233 to 140 MPa at a near-constant temperature of 900°C, whereas the subsequent microlite-bearing dome lavas continued to crystallize down to pressures as low as 9 MPa with negligible cooling. Another example is the Soufrière Hills andesite, Montserrat, where An50–60 plagioclase inclusions in the cores of amphibole phenocrysts (Higgins & Roberge, 2003), combined with experimental data (Couch et al., 2003; Rutherford & Devine, 2003), indicate protracted polybaric crystallization at temperatures sufficiently low to stabilize amphibole (840–880°C; Murphy et al., 2000; Devine et al., 2003; Rutherford & Devine, 2003). Major element chemistry of whole-rocks, phenocrysts and groundmass glass (Murphy et al., 2000; Harford et al., 2002) is consistent with crystallization of predominantly amphibole and plagioclase from a liquid whose initial andesite composition evolved to rhyolite as crystallization proceeded.
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All of the above examples suggest that decompression crystallization can play a major role in determining the crystallization sequence, assemblage and proportions. That is not to say that cooling is not important in some circumstances, nor that reheating caused by magma mixing does not occur: there is compelling evidence for both processes in many andesite magmas.
An attractive attribute of polybaric, decompression-driven crystallization is that it can be very rapid in comparison with the slow rates of crystallization expected for cooling-driven crystallization caused by heat loss from shallow magma chambers. For example, consider the case of H2O-saturated Colima andesite. To generate the observed phenocryst proportions by isobaric cooling alone would require a temperature drop of some 125°C at pH2O = 70 MPa (Moore & Carmichael, 1998). To attain the same crystallinity by isothermal decompression (at 960°C) would require a pressure drop of 60 MPa, equivalent to an ascent of 2 km. A pressure drop can be achieved much more rapidly than a temperature drop, as follows.
Cooling of shallow magma chambers is controlled by conduction through the wall-rocks and convection within the magma body and the superjacent hydrothermal system (Carrigan, 1988). The cooling timescale is controlled by the magma chamber size and the vigour of hydrothermal convection. The world's most active geothermal systems associated with large silicic magma chambers have convective thermal fluxes of several W/m2 (Carrigan, 1988). Assuming that the magma chamber convects internally, then the heat loss from the chamber can be converted into the time required to cool the chamber to a given temperature by a heat balance calculation. For example, for a cylindrical chamber with 1 km radius and 1 km depth (a volume of 3 km3), the time to cool the magma internally by 100°C is calculated at 8400 years for a heat flow of 2·2 W/m2 and a heat loss of 220 kJ/kg assuming 20% crystallization, a latent heat of 419 kJ/kg and heat capacity of 1361 kJ/kg per K. Larger chambers or lower hydrothermal heat fluxes would increase crystallization times significantly.
In contrast, magma ascent into the shallow crust is envisaged to occur in dykes (Petford et al., 1993) at speeds of cm/s to dm/s. The time taken for an H2O-saturated andesite melt to ascend 2 km would be a matter of hours (Lister & Kerr, 1991; Petford et al., 1993), thereby generating a significant undercooling caused by gas exsolution, leading to rapid nucleation and growth of crystals. Rapid crystallization of phenocrysts in arc magmas is consistent with U-series data (Reagan et al., 2005) and diffusion dating studies of phenocrysts (Zellmer et al., 1999, 2003b; Costa et al., 2004), which suggest crystallization on timescales that are far more rapid than would be expected for crystallization driven by cooling alone. Rapid crystallization also provides an effective means of generating the near-closed system crystallization inferred from experimental studies, because the timescales are too short to permit significant crystal–melt segregation, for example by crystal settling. The physical consequences of decompression crystallization are discussed further in a later section.
Whatever the cause of crystallization, the experimental data present a compelling argument that the chemical composition of andesites is determined at depth, prior to magma emplacement in the shallow crust. Of course, this concept does not exclude subsequent processing of andesite magmas in shallow chambers, including magma mixing and more advanced fractional crystallization. For example, at Santorini, Greece, dacites and rhyolites can be demonstrably related to andesite by low-pressure fractional crystallization of orthopyroxene–clinopyroxene–plagioclase–oxide assemblages (Nicholls, 1971; Druitt et al., 1999), whereas at Crater Lake, USA, rhyolite magma accumulated prior to the climactic eruption of Mount Mazama by repeated injection of andesite magmas into a shallow chamber, and extraction of residual rhyolitic melts by filter pressing (Sisson & Bacon, 1999). Partially solidified andesitic bodies, or ‘proto-plutons’, with >50% crystals can also be remobilized by subsequent pulses of hot magma from below, as envisaged at Soufrière Hills (Couch et al., 2003) and Fish Canyon Tuff, USA (Bachmann & Dungan, 2002). There are also examples of zoned plutons, such as Boggy Plain, Australia (Wyborn et al., 2001) where in situ fractionation from andesite to more evolved magmas has occurred. Additionally, such magma bodies are likely to develop incrementally over long periods of time so that mixing occurs between rising batches of andesite from depth (Fig. 1). The key concept is that the starting point for shallow chamber processes (e.g. further fractionation, wall-rock assimilation, magma mixing, magma recharge, repeated remobilization, etc.) is andesite, itself generated at greater depths.
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EVIDENCE FOR HIGH H2O CONTENTS IN ARC MAGMAS |
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Observations (Anderson, 1979
; Murphy et al., 2000; Cervantes & Wallace, 2003) and experimental studies (e.g. Sisson & Grove, 1993; Pichavant et al., 2002a; Barclay & Carmichael, 2004) indicate that many arc basalts have H2O contents in the range 2–6 wt %. Evolved residual melt obtained by crystallization of such basalts will be even more H2O-rich provided that the pressure is high enough for H2O to remain in solution. For example, 60% crystallization of basalts with 2–6 wt % H2O can generate intermediate to silicic melts with H2O contents of 5–15 wt %. (The figure is only slightly less if amphibole or mica are crystallizing phases.) Estimates of H2O contents in calc-alkaline intermediate and silicic magmas commonly yield values of 4–6 wt % (Anderson, 1979; Green, 1982; Barclay et al., 1998; Devine et al., 1998; Carmichael, 2002, 2004; Blundy & Cashman, 2005), although andesite melt inclusions with up to 10% H2O have been reported (Anderson, 1979; Grove et al., 2003). These estimates are principally based on comparison of natural phenocryst assemblages with experimental products and/or melt inclusion studies. Both approaches provide good estimates of pre-eruption H2O contents during the later stages of magma crystallization, but do not necessarily constrain H2O contents at earlier stages of magma genesis. For example, Carmichael (2002, 2004) inferred from experimental phase equilibria and thermodynamic calculations that andesites erupted in west–central Mexico crystallized by decompression from a melt with an original H2O content of at least 6 wt %, and possibly as much as 16 wt %, almost all of which was lost during magma ascent and eruption.
Additional experimental evidence for elevated H2O contents in arc magmas comes from the presence of aluminous amphibole phenocrysts in andesites. At Mount Shasta, USA, Grove et al. (2003) showed that pargasitic amphibole (9–12 wt % Al2O3) overgrowth rims on magnesian olivine and pyroxenes are consistent with amphiboles produced experimentally from H2O-saturated magnesian basalt at 800 MPa. At this pressure the dissolved H2O content of the melts is estimated at 14 wt %. At Mount Pinatubo, Philippines, Prouteau & Scaillet (2003) observed aluminous cores (>11 wt % Al2O3) to some amphibole phenocrysts in the 1991 dacite. Amphiboles of similar composition were produced in H2O-undersaturated experiments on the same dacite at pressures of 960 MPa, under which conditions melt H2O contents exceed 10 wt %. Prouteau & Scaillet (2003) attributed the aluminous amphibole cores to generation of the 1991 dacite by crystallization of a basaltic parent melt near the base of the arc crust. The lower Al2O3 amphibole rims correspond to later crystallization at 200 MPa in the sub-volcanic magma chamber.
In summary, the available petrological and experimental data are consistent with the derivation of H2O-rich andesites by crystallization of hydrous, mantle wedge-derived basalt in a lower crustal hot zone.
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MODELLING DEEP CRUSTAL HOT ZONES I—METHODS |
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We address the thermal development of a crustal hot zone with specific attention paid to all potential sites of melt generation, including partially crystallized basalt and partially melted crust. In our model a hot zone develops by injection of numerous discrete basaltic sills at the Moho or within the crust (Fig. 4). The thermal evolution of the hot zone, as a result of heat transfer between successive basaltic intrusions and country rock, is computed using the heat balance equation
 | (1) |
is density, Cp is specific heat capacity, T is temperature, t is time, X is melt fraction, L is latent heat of fusion, k is thermal conductivity and x is (vertical) distance. The system is discretized into a one-dimensional array of cells, and equation (1) is solved by forward finite difference and iterative methods. The code was written with Delphi 4© in Object Pascal language. The resolution of the finite difference cells is 25 m.
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Between the liquidus and solidus the finite difference equivalent to equation (1) is solved by iterative approximation:
 | (2) |
t is the time step between the time p and the time p + 1. Cells i – 1 and i + 1 are below and above cell i, respectively. t is limited by the cell dimension, x, and rock diffusivity to <8·5 years. Above the liquidus or below the solidus, the latent heat is zero and the temperature of cell i at time p + 1 is
 | (3) |
). The basalt emplacement rate, the fertility of the crust, the temperature of the injected basalt, and the sill injection level all control the thermal evolution of the system, and the amount and composition of the melt generated (Annen & Sparks, 2002). The model is entirely conductive and static. The issue of melt segregation process is discussed below, but it is instructive to consider the static case first.
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Temperature and melt fraction
Application of equation (2) requires knowledge of the variation of melt fraction X with temperature T. The liquidus temperature (TL; X = 1) of a basalt varies with dissolved H2O and MgO contents (Ulmer, 2001
; Wood, 2004). The relationship between X and T was parameterized using experimental data. There is no single set of experiments on a primitive basalt with fixed H2O content against which to calibrate X–T relationships from solidus to liquidus. Instead we have spliced together two datasets, one at low temperature and one at high temperature, for two different basalt compositions obtained at slightly different pressures (Fig. 5). For high temperatures (X >0·5) we have used the 1·2 GPa experimental data of Müntener et al. (2001) for a Cascades basaltic andesite (sample 85-44; 10·8 wt % MgO, mg-number 0·71) with initial H2O contents of 5, 3·8 and 2·5 wt % (Fig. 5a), run at fO2 close to QFM (the quartz–fayalite–magnetite buffer). For lower temperatures (X <0·4) we used the 0·7 GPa experiments of Sisson et al. (2005) on Cascades basalt (87S35A; 6·5 wt % MgO, mg-number 0·54) with 2·3 wt % H2O (Fig. 5b). To match the fO2 of the two datasets we have only used those experiments of Sisson et al. (2005) that are within 0·5 log units of QFM.
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For the modelled basalt we assumed linear X–T relationships between TL and the onset of amphibole crystallization, Ta, and between Ta and the solidus, Ts:
 | (4a) |
 | (4b) |
 | (4c) |
 | (4d) |
to 1261°C (Fig. 5a). Following Wood (2004) TL varies with H2O according to the relationship
 | (5) |
T
is the liquidus depression relative to an anhydrous basalt of the same composition. TL was calculated with equation (5) to be 1225, 1285 and 1302°C for total H2O contents of 5, 2·5 and 1·5 wt %, respectively (Fig. 5a) in good agreement with the variation in TL of the experiments of Müntener et al. (2001). The solidus temperature (Ts; X = 0) of basalt also depends on H2O content. However, for any basalt that contains more H2O than can be accommodated in su