Christiansen, E.H., and Keith, J.D., 1996, Trace element systematics in silicic magmas: A metallogenic perspective: Geological Association of Canada Short Course Notes Volume 12, Trace Element Geochem- istry of Volcanic Rocks: 24-26 May 1996, Winnipeg, Manitoba Canada, p. 115-151. 116 RH. Christiansen and J.D. Keith Consequently, the composition of silicic magma reflects its tectonic setting, but in a secondary and sometimes inconsistent fashion. Magma Sonrces The "source" of a magma is more a hypothetical construct to aid the understanding and modelling of magma evolution than a physical reality. Nonetheless, many magmas appear to derive much of their mass from a distinctive reservoir and, herein, we envision the history of a magma as beginning with partial melting in a single reservoir. Almost immediately, the magma begins to communicate with its environment, exchanging heat and mass as a non-isolated open system; we consider these contributions to the magma in subsequent sections. To describe a magma source completely, we need to know its bulk composition, its mineralogy, the fugacities of the volatile species, and the values of the relevant partition coefficients. Each of these is complexly dependent on the others and on the prevailing temperature and pressure. The source of a magma places important constraints on the composition of any partial melt that is derived from it. At equilibrium, the activities of all of the components must be equal in the melt and coexisting solids. Thus, the degree of Al and Si saturation, Fe2+/Fe3+ ratios, various volatile fugacities, etc. are initially controlled by the composition of the source. In a primary melt, the mineral compositions are the same in the source residues and the magma. Moreover, simple partition coefficients can be used to describe the concentration ratios of each trace element in the meLt to that in the solid at equilibrium. Ratios of incompatible elements (e.g., RblNb) are similar in the source and in the magma. The types and abundance of mineral phases that remain in the source are reflected in compatible trace element depletions in the derived magma. On the other hand, enrichments of a specific trace element in the melt may indicate that a phase was eliminated in the partial melting reaction. Source composition can even control the extent of melting, as described below. In each of these ways, silicic magmas are images of their sources. However, subsequent processes greatly degrade this image. Partial Melting Processes The generation of magma involves partial melting of solids in the source. The extent and type of melting are important for the elemental composition of the resulting magma. To describe trace element systematics in magmas, it is useful to define the extent of melting as F (the weight fraction of melt present in a partially molten system) and the bulk partition coefficient as D (the ratio of the concentration of a trace element in the assemblage of crystals to its concentration in the coexisting melt). Because trace elements have different affinities for different minerals, the bulk partition. coefficient for each element depends on the types and proportions of the minerals that are in equilibrium with the melt. Compatible elements have D > 1. Sr, Ba and Eu all partition strongly into feldspars in .silicic melts and dre usually compatible elements. Rb, Li and Nb partition only weakly into the principal minerals found in silicic magmas and are usually incompatible elements, with D < 1. The main variable in the partial melting process is the extent, or degree, of melting. Whatever the type of melting process, low degrees (small F) of melting produce melts enriched in incompatible elements and depleted in compatible elements, compared to the source rock.. Larger degrees of melting cause incompatible elements to drop in concentration and compatible elements to increase 117 until at F = 1 (100% melting) the melt has the same composition as the source (Fig. 1). The extent of partial melting is controlled by the heat flux (either conductive or advective), the magnitude of decompression, the amount of fluxing volatiles introduced, or the proportions of hydrous minerals for dehydration melting. Regarding this latter point, a crustal protolith with abundant hydrous minerals may produce a magma with low incompatible element concentrations because a larger proportion of the protolith may melt. In contrast, a nearly dry protolith, even one that has experienced a previous episode of melt extraction, may produce magmas that are more enriched in incompatible elements by virtue of a small melt fraction. For example, decomposition of 5% biotite in a nearly anhydrous high-grade metamorphic rock produces about 5% partial melt (Burnham, 1979; Clemens, 1984) that is comparatively enriched in incompatible trace elements (Fig. 2). To understaild how different melting processes also affect trace element concentrations in magmas, consider the difference in compatible trace element concentrations between batch partial melting and fractional melting, two end-member models used to describe partial melting processes. Partial melting that occurs in accord with the assumptions of batch partiru melting produces melts that are not extremely. depleted in the compatible elements, which are limited to concentrations proportional to lID (Hanson, 1978). For example, if the D for Sr is 5, the lowest concentration (at very small F) that can occur in a single batch of partial melt is 1/5 of the initial concentration in the solid source (Fig. 1). If a series of melts is related by varying degrees' of batch partial melting of a single source, then Sr would vary by no more than a factor of 5 from the most enriched to the most depleted melt. Most volcanic and intrusive suites thought to be cogenetic violate this limitation, I I • suggesting that variations in the degree of batch partial melting are not the dominant cause of elemental variation in silicic magmas. Of course, this conclusion has no bearing on the importance of batch partial melting in producing magmas that are subsequently modified by contamination or fractional crystallization. Although incompatible trace element concentrations are greatly affected by the extent of melting, their ratios are less changed (Fig.s 1 and 3). As a result, ratios of I' incompatible elements may be examined to deduce the nature of a magma'ssource(s). I : I' Unfortunately, few trace elements are highly incompatible in silicic magmas, because of the great , I I variety of accessory minerals that contain them and because of the high partition coefficients for I cool, highly polymerized silicic magmas (Nash and Crecraft, 1985; Mahood and Hildreth, 1983). Another important partial melting model is fractional melting, in which each bit of magma . escapes from the source as soon as it is produced. For fractional melting, there is no lower limit to the concentration of compatible trace elements in a partial melt (Fig. 1). The initial fractions of melt (low F) are extremely depleted in compatible elements and enriched in incompatible elements. Imagining the physical conditions that would allow the instantaneous separation of each small fraction of melt is difficult. In most cases, a certain amount of melt must be generated before the source becomes permeably. Natural melting processes probably include aspects of batch and fractional melting. More complicated mociels of partial melting include the contribution of an external fluid or melt to the production of magma (Allegre and Minster, 1978). This type of model may be especially applicable for the initial generation of subduction-zone magmas, where a fluid, derived by dehydration of descending oceanic crust, introduces soluble elements and, simultaneously induces partial melting in the overlying mantle wedge. The resulting magma is a composite of the fluid, melt from the mantle wedge, and solids modified by equilibration with the fluid and the melt (Fig. 4). .. • 118 E.H. Chrisp.ansen and J.D. Keith • 100 r-;::~~=-=;:::;:::;:::::;:=:;-11 100 A I2f Ch P~rtia/ijr:e/t1iJil B Fractional Melting & • Fractional C stallization • CI/Co =1/(O(1-F)+F) • I 10 10 • i I' i i • o o o o • --o --o • 0.1 • 0.1 • • 0.01 0.01 ..• 0.001 1 0.80.6 0.4 0.2 o 0.8 0.6 0.4 0.2 o • F (fraction of melt) - F (fraction of melt) .,• Figure 1 Mathematical models ofpartial melting andfractional crystallization processes can be used to understand the trace element concentrations of magmas. The relationship between the weight fraction of melt (F) and the i • concentration of a trace element in a melt (Cl) relative to its original concentration (Co) is shown. Co is the trace .. i'l I element concentration in the original solid source for .batch partial melting and is the trace element concentration in .. the original magma for fractional crystallization. D is the bulk crystalliiquid partition coefficient. (A) Batch partial' melting, in which the melt is in equilibrium with the solids until it separates. (B) Fractional melting, in which the • melt is continuously removed from a proto lith or source, and fractional crystallization, in which the solids are .. continuously removedfrom a magma, have identical effects on trace element concentrations. .. .., 1000 .. I ~=:~!i!!:~~~U Ymelt ~. j 500 ~(je Sov. .at! 300 ~\(j'(\ e' -E 'O\~ Co 200 <0'1 ~(je Co e.j Sov. = -.0cr:: oo~ ell i .'!..e'« .., 100 I <0\0\\ 50 Xo 30 5 10 20 30 50 100 Nb (ppm) Figure 2 The decomposition of a hydrous mineral such as biotite may control the degree of partial melting of a crustal source; in this case the amount of melt produced is directly proportional to the amount of biotite. For a biotite-poor (5%) source with low concentrations of incompatible elements, points Xo and Xmelt represent the compositions of the original solid source and a 5% partial melt.