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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 Geochemistry 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, Fe 2+/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,
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
incompatible elements may be examined to deduce the nature of a magma'ssource(s).
Unfortunately, few trace elements are highly incompatible in silicic magmas, because of the great
variety of accessory minerals that contain them and because of the high partition coefficients for
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).
I
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118
E.H. Chrisp.ansen and J.D. Keith
100 r-;::~~=-=;:::;:::;:::::;:=:;-11
100
A I2fCh P~rtia/ijr:e/t1iJil
CI/Co =1/(O(1-F)+F)
10
I
B
Fractional Melting &
Fractional C stallization
10
i
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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
F (fraction of melt)
i
i'l
I
0.6
0.4
0.2
o
- 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
concentration of a trace element in a melt (Cl) relative to its original concentration (Co) is shown. Co is the trace
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
500
E
Co
Co
~=:~!i!!:~~~U
~(je
~\(j'(\
300
'O\~e'
<0'1
200
cr::
i
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Sov.
ell
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20
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,
.
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••
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~
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~(je
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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. For a source with 20% biotite and higher
concentrations of incompatible elements, points Yo and Ymelt represent the source and a 20% partial melt.. Because
of the low degree of melting, incompatible element concentrations are higher in melt Xmelt than in the melt Ymelt,
derivedfrom the enriched source. Points on curves represent 10% melt increments. DNb=0.1 and DRb = 0.05 for
both sources.
119
80
F
0.1
800
A
60
.3;40
.c
-500
E
o.
.3;400
.c
0:: 300
20
200
C-
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Original Composition. --.J
0.2
1
0.5
1.0
B
0.1
ONb = 0.3
600
E
0
0.1
F
700
r--
___ 0.2
ORb
0.3
0.6
0.9
100
Or\ginal Composition
2
5
10
0
0.1
20
0.5 . 1
0.2
Rb/Nb
2
5
10
20
Rb/Nb
Figure 3 Incompatible element ratios are relatively unchanged by variable degrees of (A) batch partial melting or
(B) fractional crystallization. Dl{b = 0.3 and was held constant. DRb was vaned as shown. F is the fraction of melt
in the system. Except when the partition coefficients are very different (e.g., DRb=0.9), the RbINb ratios change by a
factor of2 or less between F=l and F=O.l.
300~--------------------------~~---'
S-type Granites
100
A-type
Granites
30
E
a.
.3:
.c
0::
10
3
1
Batch Partial
0.3
0.1
AII""lti",,.,
1.0
2
3
5
10
20
30
50
100
Nb +Y (ppm)
Figure 4 The effect of mantle metasomatism on incompatible trace element concentrations during partial melting of
the mantle (upper curve) compared
to melting of normal mantle (lower curve). Rb represents the soluble elements
I
introduced into the mantle wedge by metasomatism above a subducting slab of oceanic lithosphere and Nb+Yrepresent the insoluble high field strength elements. DNb = DY = 0.1 and DRb= 0.05 are the bulk distribution
coefficients for the elements and are assumed to be the same for both sources. F is the fraction of melt in the system.
Field boundaries are from Pearce et al. (1984). Shaded region shows composition ofI-type rhyolites from Macdonald
et al. (1992). Batch partial melting of metasomatized mantle produces magmas with the characteristics of I-type
granites.
zi
E.H. Christiansen and J.D. Keith
I :
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, ,'I
The resulting magmas are enriched in the elements carried by the fluid, here represented by Rb, and
do not have trace element patterns like the source before fluid interaction.
Restite unmixing, where residual solids are variably separated from the partially molten mass, is
important at the site of magma generation and controls the initial composition of the separated
magma (crystals plus liquid). After magma segregation, restitic mineral phases may become
dissolved at lower pressure or after continued reaction with the melt. However, restite unmixing is
probably not important in the generation of the observed elemental trends in rhyolites and granites,
except some S-types thought to be derived from sedimentary parents, that show abundant
petrographic evidence for residual materials. In general, granitic magmas, especially the mineralized
ones that are highly evolved, have cleared themselves of restite. The eruption of large volumes of
nearly crystal-free rhyolite attests to this. Most of the chemical variation in granitic plutons probably
is caused by fractional crystallization rather than variable proportions of minimum melt and restite
from the source (e.g., McCarthy and Hasty, 1976; Bateman and Chappell, 1979; Lee and
Christiansen, 1983). During pluton solidification, cumulates and melts may be mixed in variable
proportions.
Application of inverse techniques, such as those described for batch partial melting by Minster
and Allegre (1978), to estimate extent of partial melting and source composition, has not generally
been successfUl with silicic igneous rocks. This failure is probably because most silicic magmas
experience extensive fractionation andlor contamination during their histories and because some
crustal melts may entrain substantial amounts of residual solids and cumulus phases.
All these simple models assume Henry's law behaviour of trace elements and complete
equilibration of melt and coexisting residual solids. Because of the relatively low temperatures and
highly polymerized rhyolitic melts produced by crustal melting, diffusion is too slow to remove
compositional gradients before melting occurs. Consequently, disequilibrium melting may be more
common in the crust than in the mantle. Bea (1995) questioned the applicability of simple
mathematical models to the partial melting of crustal rocks. He contends that the effective partition
co~fficients for trace elemet:lts in the major phases approach one at the low temperatures
characteristic of crustal melting. This is so because the melting rate may exceed the diffusion rate
for many trace elements. For minerals with diameters of 2 mm and trace element diffusion rates of
about 10-16 cm 2s- 1, it takes ).6 m.y. to achieve equilibrium between melt and residual crystals. If
this happens, then silicic melts would be more direct images of their sources than generally
imagined; the concentration of a trace element in the melt would be simple function of the
composition of the minerals that participated in the melting reaction. Another cause of non-Henrian
behaviour of trace elements during crustal melting arises because some trace elements are essential
structural components of accessory phases. The concentrations of such elements in melts are
controlled by the solubility of minerals, not by simple Henry's law behaviour. The best example of
this is Zr in zircon, but it probably applies to U and REE (rare-earth elements) in monazite, and Y
and REE in xenotime. Disequilibrium processes may also disturb the Sr and N d isotopic ratios of
crustally derived magma, so that the melt's isotopic composition no longer reflects that of the
source. This may happen if Sm and Nd are largely held in residual accessory phases with steep
REE patterns, such as monazite, that do not equilibrate with the magma before it separates
(Creaser, 1995). Alternatively, if magma produced is caused by the breakdown of hydrous
minerals, the decomposition of biotite could release highly radiogenic Sr to a small melt fraction.
4
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t
41
41
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=
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=
121
These problems probably only apply to magmas produced by very small degrees of melting.
Fractional Crystallization
The compositional evolution of most magmas is driven largely by cooling and fractional
crystallization. The extent of crystallization of a magma chamber is an important control on the trace
element concentrations of the residual melt and the solids. The effect of fractional crystallization,
with or without complete removal of crystals, can be modelled using the Rayleigh fractionation
equation (e.g., Allegre and Minster, 1978; Fig. 1). If high quality analyses of cogenetic igneous
rocks are available, inverse modelling techniques (Allegre et aI., 1977) can be used to deduce the
nature of the fractionating mineral assemblage and extent of frac~ionation; reasonable values for
indiVidual mineral/liquid partition coefficients must be used. The important role of accessory phases
in controlling the abundances of many trace elements makes the inverse approach more difficult for
silicic magmas than for mafic magmas, but has yielded reasonable results (e.g., Christiansen et aI.,
1984).
Obviously, rocks derived from highly fractionated magmas are compositionally quite different
from those magmas that have experienced only modest amounts of fractional crystallization.
Although concentrations of ore elements in silicic magmas or cumulates are seldom, if ever, of ore
grade, significant enrichments may occur that set the stage for subsequent ore-generating
processes. The extent of fractional crystallization is especially critical for those elements that have
either very high or very low bulk partition coefficients. Compatible elements with high partition
coefficients become sequestered in solid phases where they may become subsequently remobilized,
depending on the nature of the host mineral and its reactions with later fluids. If fractionation
proceeds extensively, magmatic concentrations of compatible elements may drop below the critical
level needed for the generation of ore (Candela, 1989). Copper behaves as a compatible element in
most felsic magmas, largely because its partitions strongly into the small proportion of immiscible
sulphides that eventually fractionate. Consequently, porphyry Cu deposits are typically associated
with less fractionated igneous rocks where the Cu concentrations remained relatively high. Copper
deposits are rarely found in association with highly fractionated leucogranites, which have
inherently low concentrations of Cu. On the other hand, incompatible elements are concentrated in
residual melts, locally setting the stage for subsequent processes to enrich the element further. The
behaviour of Be exemplifies that of an incompatible element. Be deposits_ are typically associated
with highly evolved granites, rhyolites, and pegmatites that had extensive histories of fractionation
to concentrate Be.
In highly evolved rhyolitic magmas, the effect of fractional crystallization is most obvious in
trace element rather than major element variations. Fractional crystallization drives a silicic magma
toward the composition of a minimum melt in the granite system. In high-Si02 rhyolite magma, the
/
coexisting solids (subequal proportions of plagioclase, sanidine and quartz) and the melt have
nearly the same major element composition and large amounts of fractionation can occur without
large changes in major element compositions. However, trace elements continue to change in
accord with Rayleigh fractionation. For example, the zonation in the Bishop Tuff in eastern
California spans a 2% change in Si02, but compatible and incompatible element abundances vary
by several hundred percent (Hildreth, 1979). These changes are consistent with fractional
crystallization (Christiansen, 1983; Michael, 1983), but have been attributed to other poorly
,I'
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122
RH. Christiansen and J.D. Keith
••
••
'"•
~
~
constrained processes such as volatile complexing, "thermo gravitational diffusion," or "complex
interactions of petrogenetic processes" (Hildreth, 1979; Macdonald et aI., 1992). The fact that trace
element enrichment and depletion patterns in rhyolites and trachytes correspond to those produced
by fractionation of their mineral assemblages, including distinctive accessory minerals, argues
persuasively for the comparatively simple process of crystal and liquid separation (Wolff and
Storey, 1984).
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~
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~
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Mixing and Assimilation
Once a batch of magma is produced, it immediately begins to interact with its surroundings in a
long process of trying to come to chemical, thennal, and mechanical equilibrium. Compositional
changes in the magma result from open-system behaviour, or contamination. This process becomes
especially important in the crustal entrapment zones above subduction zones where mantle-derived
magmas stall because of relatively high densities, partially crystallize to induce wall-rock melting,
and then mix with the crustal magma while being reintruded by fresh draughts of mantle-derived
magma.. Contamination continues to be important when a magma batch achieves buoyancy and
moves out of its source region and into totally different rock types in the overlying crust and/or
mantle.
Obviously, the composition of the contaminants is important to the final composition of the
mixture, but the particular type of contaminating process plays an equally important role. The
magma may assimilate wall rocks by bulk digestion if a block falls into the magma. For example,
Johnson et aI. (1990) found Proterozoic zircons ina tuff erupted from the Mo-mineralized Questa
magmatic system of northern New Mexico. Apparently, these were derived from nearly complete
dissolution of Proterozoic wall rocks. Commonly, assimilation of wall rocks forces the magma to
crystallize. The effects of simultaneous assimilation and fractional crystallization (AFC; DePaolo,
1981a) are contrasted with simple fractional crystallization in Figure 5. Moreover, a magma may
encounter other bodies of magma, from the same or different sources and, in its attempt to mix with
these extraneous magmas, large compositional variations may be created. If mixing proceeds tb
completion, the final composition depends on the mixing proportions. Commonly, mixing is also
accompanied by crystallization as new phases fractionate, once they are stabilized by the changing
compositional and thermal regime created by mixing. On a smaller scale, the original magma, itself
a heat source that can induce the production of more magma by partial melting, may become
contaminated by partial melts drawn from the wall rock. The composition of these contaminating
magmas is governed by appropriate partition coefficients, wall rock compositions, and the degree
to which equilibrium is achieved. More subtle magma - wallrock interaction may occur by diffusive
exchange driven by compositional and thermal gradients. Some speculate that shallow magma
chambers also can be contaminated by hydrothennal fluids during caldera collapse or resurgence
(Hildreth et at., 1991). These! processes of contamination occur throughout the history of all
magmas, from the moment of initial production, through rise and emplacement as a pluton or
eruption as .volcanic rock, and are partially responsible for obscuring the connection between
igneous rocks and their original sources.
'"~
~
..,...,~
..•
..
~
---~
100
A
-
.. . . . . . . .- .
300.-------------~_r.----------.
I AFC r=O.21
B
_
250
10
Fractional
Co
.e
~
.0
0.1
~ 150
0.01
100
0.001 1
0.8
0.6
0.4
0.2
0
.:.;
.!
Contaminant
j, ~
~,~
E
C3
.::.
Crystallization
Assimlation
...•..
& FC
_200
1
0
123
DNb=O.10
DRb=O.25
~~
DNb=O.25
DRb=O.10
50L----L--~~~~~----~~-J
1
F (fraction of melt remaining)
Figure 5 Fractional crystallization.and assimilation combined with fractional crystallization (AFC; DePaolo,
1981a) are contrasted in.these diagrams. The rate of assimilation/rate of crystallization is the same in both diagrams
(r=O.2). (A) Cl is the trace element concentration in the evolving liquid and Co is its original composition. The
numbers on the curves show the ratio between the trace element concentration in the assimilated material and the
original magma. Partition coefficients (D) are shown in boxes. Thin lines are for simple fractional crystallization.
(B) The extent to which incompatible trace element ratios are changed by assimilation and fractional crystallization
depends on the composition of the contaminant (here assumed to be a pelitic sediment) and upon the partition
coefficients (DNb and DRb) for the trace elements. Dots on lines represent 10% increments of crystallization. In the
case shown here, little change in incompatible element trace element ratios is evident.
=
•
I.,.
E
.E
...
...
t.
Volatile Fugacities
Volatiles play an important role in the evolution and eventual mineralization of silicic magmas.
Magmatic volatile fugacities and fugacity ratios depend on source compositions, fractionation
histories, and mixing and assimil.ation. However, the effect of each volatile is so diverse that they
are worth considering separately. The principal volatile components of interest are H20, HF, HCl,
C02 and the various S species. The fugacities of these species largely reflect those in their sources
(see Carmichael, 1991, for a discussion centred onj02), but various evolutionary processes in
open magmatic systems may change the values (pichavant and Hammouda, 1995). Dramatic and
consistent variations in j02 between the various types of silicic magmas have been identified.
Proportions of mantle-derived magma (with widely varying j02) and the dramatically reduced
(C-containing sedimentary rocks) and oxidized (sedimentary and altered igneous rocks that reacted
with the atmosphere) magmas derived from the continental crust determine the prevailing f02 of
silicic. magma. Dehydration of oceanic crust in subduction zones also releases an oxidized fluid that
passes this characteristic on to magmas formed in this environment. Munoz (1984) has also shown
that halogen fugacities vary systematically in silicic igneous rocks and associated mineral deposits.
Silicic rocks related to Cli mineralization have higher j HClIjHF than those associated with Mo or
Sn-W -Be. Moreover, volatile fugacities control the stability of many phases that are important for
ore element concentration or dispersion. For example, the solubility of sulphur as an immiscible
sulphide melt strongly controls the Cu concentrations of evolving silicic magmas. If magma
becomes sulphide saturated, Cu partitions strongly into these blebs and is commonly removed from
the magma by fractionation. Hence, porphyry Cu deposits are rarely developed from highly
, i
E
124
E.H. Christiansen and J.D. Keith
evolved rhyolitic magmas, but may be related to their more mafic precursors. Another good
example of the role of volatile fugacity controlling mineral stability and trace element evolution lies
in the effect of f02 on the oxidation state and consequently, magmatic and hydrothermal behaviour
of Sn. In magmas with 102 below the QFM (quartz-fayalite-magnetite) oxygen buffer, the
compatibility of Sn in mafic silicates and oxides is apparently lower and Sn becomes strongly
enriched in the residual liquid by fractional crystallization processes (Ishihara, 1977). Sn+2 is
apparently less compatible in minerals than Sn+4, which predominates at high j02. Moreover, low
j02 increases the solubility of Sn in hydrothermal fluids derived from, or equilibrating with,
reduced granites (Eugster, 1986). In contrast, high f02 produces high Mo+6 /Mo+4 and lowers the
partition coefficients for Mo. This probably explains why Mo and Sn deposits are both related to
highly evolved magmas, but do not commonly occur as ore deposits related to the same magma
)
(Keith et al., 1993).
Fluid saturation can be treated as a phase separation as well. Most silicic magmas are probably
fluid-saturated during much of their evolution as the result of the extremely low solubility of CO 2.
The effects of the migration and partitioning of trace elements into this kind of fluid are probably
minor. More dramatic changes may occur in magmas that reach water saturation. Some elements
preferentially partition into an aqueous fluid (CI, the alkalis and alkaline earths, Pb, Zn, Cu)
compared with others that prefer the melt (F and high field strength elements--those with high
charge to radius ratios like Nb, Ta, Zr, and Ti). The mobility and redeposition of these elements
significantly alter the patterns produced by crystal-melt equilibria and must be carefully avoided to
ascertain magmatic patterns. For example, high Flel ratios (>5) in magmas may be produced by the
separation of a Cl-rich aqueous fluid rather than by distinctive sources or melt evolution paths.
Tectonic Setting
.
Although there are broad correlations between tectonic setting and the geochemistry of magmas, in
detail the composition of a silicic magma is not simply nor directly tied to the tectonic setting in
wQ.ich it was formed, differentiated, and emplaced (e.g., Twist and Harmer, 1987). Rather, the
composition is the product of its partial melting history, source(s), crystallization and contamination
processes, and prevailing volatile fugacities. Consequently, the trace element signatures of granites
do not uniquely identify the~r tectonic setting, but point directly to their melting and crystallization
histories. It is this latter inference that is most important for the association of ore with rhyolitic
magmas. We emphasize the compositional characteristics of granites, rather than their tectonic
settings, in this paper.~ We use trace element characteristics and so-called tectonic discrimination
diagrams to help identify magma sources and evolutionary processes.
However, to the extent that these controls on magma chemistry are linked to plate tectonic
setting, tectonic discrimination is; meaningful and the tectonic settings of the principal types of
silicic magmas are discussed belo~. For example, silicic igneous rocks from mid-ocean ridges are
distinct from those associated with oceanic plumes, compositions, mineral assemblages, and
volatile fugacities are all distinctive. This is apparently the result of homogeneity in the processes
and magma sources available in these two distinctive tectonic settings. Magmas formed in
subduction zones appear to have consistently highj02,jHCI, and distinctive trace element patterns.
In oceanic island arcs, uniformity in source materials, magma producing reactions, and ascent paths
leads to uniformity in island-arc igneous rocks. As a result, the associated ores are consistent and
F=:=================================-:==~--.--~.-~-.-
.125
predictable. In contrast, where subduction-generated magmas interact with compositionally
heterogeneous continental crust, they take on elemental and volatile characteristics that reshape their
compositions and ore-generating capacities. Subduction-zone magmas may become heavily
contaminated with sedimentary and other continental materials along continental margins. This
typically changes the magma composition to become more peraluminous and less oxidized; two
factors that radically change the assemblage of fractionating accessory phases and the partition
coefficients of trace (including ore) elements. Likewise, the heterogeneity of continental crust and
its variable contributions to all types of magma (plume-, rift-, subduction-, or collision-generated)
that pass through it or that are produced in it, imparts dramatic variations to all continental magma
associations and complicates the association of tectonics and magma composition and contributes to
the formation of certain metallogenic provinces.
SILICIC MAGMA TYPES
A useful framework for the discussion of the trace element compositions of silicic magmas is
provided by the compositional criteria proposed by Chappell, White, and co-workers in a series of
papers beginning in the 1970's (e.g., Chappell and White, 1974, 1992; White and Chappell, 1983;
Collins et al. 1982; Whalen et al., 1987). Although the origins of these compositional types are still
controversial, the application of the criteria for distinguishing various kinds of granites and
rhyolites is more clear cut. Ishihara (1977) identified distinctive ilmenite- and magnetite-series
granites and emphasized the importance of oxygen fugacity in the trace element character and ore
associations of granitic magmas. The consistent identification of distinctive granite types in specific
tectonic settings led to the idea that granite compositions and tectonic settings were genetically
related. In his paper, "Granite type and tectonic environment,:' Pitcher (1982) espoused this
connection and extended the subdivision of subduction-related granites into several groups. Pearce
et al. (1984) used trace elements to distinguish various types of granites, and divided granitic rocks
into groups that are similar to those of Chappell and White. However, Pearce et al. (1984)
distinguished the granite types as products of their tectonic setting; rather than as compositional
types that reflect their source components and magmatic processes. In this review, we discuss the
compositional characteristics. and ore associations of four principal types of rhyolitic/granitic
magma as follows, OR (ocean ridge)-, I-, S-, and A-type (Table 1). The letters representing the last
three categories were intended by Chappell and White to describe the crustal source of each type of
magma in a rather strict sense. I-types from igneous source materials, S-types from sedimentary
materials, and A-types from anhydrous, melt-depleted rocks. The letters are retained because of
their familiarity, but the connotation of strictly representing a crustal granite source is abandoned
here. The various groups ~e simply compositional types of granites that reflect their protracted
evolutionary histories, including, the extent and style of partial melting, their sources, degree of
fractionation and/or contamination, and inherited volatile fugacities. Because these compositional
controls are in turn affected by the plate tectonic setting, the general tectonic setting of each type and
some problems involved when comparing tectonic setting and granite composition are discussed.
..
127
I -Type Silicic Magmas
Rhyolites and granites generated at subduction zones are diverse, but most have a broad range of
similar features, which together distinguish I-type silicic magmas from other types (Table 1). Some
of the most important compositional features ofI-type granites and rhyolites are their high ratios of
large ion lithophile elements to high field strength elements (Pearce and Cann, 1973) and high
fugacities of 02, H 20, Hel, and oxidized sulphurous species. In mafic rocks, high BaJNb ratios
are employed to examine the relative depletion of high field strength elements as compared with
other equally incompatible elements. In evolved felsic magmas v:here sanidine and biotite are
stable, Ba behaves as a strongly compatible element and the BaJNb ratios drop. In such silicic
magmas, Rb/Nb and Th/Nb ratios give better indications of this depletion in the source. These
differences are especially clear when I-type rhyolitic magmas are compared with those from the
ocean basins, including oceanic plumes and mid-oceanic ridges (Fig. 6). Rb/Nb ratios should be
source related because they do not change dramatically as a result of closed system fractionation or
variations in degree of partial melting (Fig. 3). However, contamination with upper continental
crustal materials may increase this ratio in open magma systems (Fig. 5). Because of low Nb (+Y)
concentrations, I-type rhyolites and granites fall in the volcanic arc field of Pearce et al. (1984),
which we have relabeled here to emphasize compositional rather than tectonic characteristics (Fig.
7). As expected from these characteristics, I-type magmas have deep negative Nb anomalies on
chondrite-normalized trace element diagrams (Fig. 8). Most I-type magmas have highj02 (several
log units above QFM; Ewart, 1979) and consequently high Mg/Fe and Fe20~eO ratios in mafic
silicates and bulk rocks (Fig. 9). Many are calc-alkalic in the classification of Miyashiro (1974),
. which is based on this parameter, but some range to lower MglFe ratios and are.tholeiitic. Early
fractionation of magnetite, because of high j02, and other phase relations caused by high jH20
drive the characteristic calc-alkalic fractionation trend marked by Si02 enrichment and Fe-depletion.
This process has operated over time to produce Earth's unique continental crust. High water
fugacities are shown by the presence of hydrous minerals in relatively mafic rocks. High jHCVjHF
is shown by biotite and amphibole compositions (Munoz, 1984; Ague and Brimhall, 1988) and
high F/CI ratios in melt inclusions (Lowenstem, 1995; Fig. 10). Volcanic and many plutonic
associations are unimodal with broadly continuous trends on Si02 variation diagrams. Among
I-type magmatic series, intermediate compositions, andesite or granodiorite, are commonly the most
abundant rock types. These simple characteristics are complicated by the extent of fractionation and
by the nature of the lithospheric contaminants. Consequently, wide ranges of compositions appear
that correlate with the type of overriding plate- whether it includes continental or oceanic crust,
thick sedimentary basins or granulite facies metamorphic terranes, ancient metasomatized mantle or
young depleted lithospheric mantle.
The key features of I-type magmas are tied to the evolution of oceanic crust. Mter its formation
at an oceanic'ridge, the basaltic crust becomes hydrated, oxidized, and mineralized as ocean water is
pumped through it in large hydrothermal convection cells. Sediment accumulates above the igneous
crust. As this altered oceanic crust sub ducts and dehydrates, oxidized aqueous fluids carry soluble
elements into the overlying mantle. This creates magmas (and the continental crust derived from it)
that are oxidized and have high abundances of the water soluble elements (e.g., Cl, S, Cu, Ba, Rb
and·K) compared with oceanic basalts. The characteristic Nb-Ta-Ti depletion of subduction-related
I-type magmas is probably the result of residual titanates in dehydrating (or melting) oceanic crust
-----~-------~
RH. Christiansen and J.D. Keith
128
(e.g., Ryerson and Watson, 1987). The released fluid is depleted in elements stored in relatively
insoluble titanates like rutile, a common mineral in eclogites. Alternatively, partial melting of the
basaltic slab may produce a low temperature siliceous melt that is rutile-saturated. The rutile is left
behind when the magma rises into the overlying mantle wedge. There it reacts with peridotite to
produce the more common mafic I-type magmas, or rarely erupts as adakite (Yogodzinki et ai.,
1995).
1,200
•• •
o •
,
-O~
1,000
~i
E
o
600
S
.0
•
0
•
~
A-type
1- • •
,."
0::
400
,1
S-type
OR-type
§D
0 1
,Q.
e
•
'"
0
0
800
Figure 6 RbINb ratios of various types
of silicic igneous rocks. Data are from
Macdonald et al. (1992), the authors'
analyses of topaz rhyolites from the
western United States, and sources cited in
text for discussed localities. RhlNh ratios
are lowest in rhyolites from ocean ridges
(OR-type) and highest in S- and I-types.
Subalkalic A-type' rhyolites from the
continents have slightly higher RbINb
ratios than other A-types, including
peralkalic rhyolites and those from oceanic
plumes.
I-type
8
0
Peralkaline
<>
Plume
0,
200
0.1
0.2
0.5
1
2
5
10
20
50
Rb/Nb(ppm)
1000
1000
S-Type
Granite
300
E
"'-
• •
•
100,
.s
-<J
a:::
30
S-type
A-Type
Granite
I-Type
Granite
10
•
100
Bingham (Cu)
• Pine Grove (Mo)
•
Macusani
o
o
CanMacrN)
Phukel (Sn)
•
Spor MIn (Be)
E
"'-
.s
-<J
a:::
10
OR-type
~o~
'o~7>
~~Q,eJ? 'T:~7>rjIJ
~P~;'~9.
--i'?fo~
3
2
5
10
20
50
100 200
Nb+Y(ppm}
500 1000
2
5
10
20
50
100
200
500 1000
Nb+Y(ppm}
Figure 7 (A) Left: compositions of selected igneous rock suites that are related to mineralization plotted on Pearce
et al. (1984) tectonic discrimination diagram. The field names have been changed to emphaSize compositional, rather
than, tectonic characteristics. Datafrom sources cited in text. (B) Right: the effect o/various processes and source
composiiions on trace element trends anli positions on this diagram.
Sedimentary materials may contribute to the genesis of subduction zone magmas in two ways.
First, oceanic sediment is subducted along with the oceanic crust in many subduction zones.
Cosmogenic lOBe is produced in the atmosphere and becomes enriched in oceanic sediment but
decays away rapidly with a half-life of 1.5 m.y. Measurable activities of lOBe (Morris and Tera,
129
1989) in modern arc lavas demonstrate that substantial amounts of young subducted sediment
contribute to I -type magmas fonned in subduction zones. Sedimentary materials in the continental
crust may also contribute to I-type magmas during assimilation of wall rock by mantle-derived
magma. DePaolo (1981b) and Ague and Brimhall (1988) show the effect of this kind of crustal
interaction for many plutons in California, including the Sierra Nevada and Peninsular Ranges
batholiths. Assimilation of graphitic pelites created strongly reduced ilmenite-series granites in the
western portions of these batholiths. Assimilation of more oxidized crustal materials occurred to
,...
....•
,.
••
....•
Ie
Figure 8 Normalized trace element patterns for the principal types of silicic igneous rocks. The strongly compatible
elements, Sr, P, and Ti, have been omitted. Normalizing factors are mostly cho1ulritic abundances and taken from
Thompson (1984). Data from Macdonald et al. (1992) and sources cited in text. The pattern of an intraplate alkalic
basalt is shown for reference with a dashed line.
~
ss:::i0 -
I
I'
E.H. Christiansen and J.D. Keith
130
5
2
;?
0
1
!
'.
I-type
•
S-type
•
0.5
0
-
[JO
<U
u.
C'"J
0.2
0
N
@
Q)
u.
A-type
..• ~. 1f ..~
0.1
·
0;05
".r
':.
o
[email protected].. .....•:--•:..• ••••
...".
~..
~'"!
~
.~ . . .
Peralkaline
o
'~.
Plume
o
·
.1·
.
••
•I•·
®g • . .
Figure 9 Fe2031FeO ratios
for silicic igneous rocks of
various types. A- and J- type
data from obsidians of
Macdonald et al. (1992). S-type
granites from eastern Australia
from Blevin and Chappell
(1992).
•
0
1
2
3
4
6
5
FeO total (wt%)
the east and is evidenced by increases in incompatible element abundances and jHPljH20 and
jHPljHCl as indicated by biotite compositions; here, magmatic j02 was not strongly affected. In
many continental magmatic arcs, progressive assimilation of crustal sediments by mantle-derived
magma created S-type magmas that may be indistinguishable from those derived purely by crustal
anatexis.
0.3
S-type
•
......
0.1
~
0
!
A-type
0.·
C3
Peralkaline
()
0.03
Plume
o
0.01
0.005
0.01
0.03
0.1
0.3
F (wt%)
1
3
Figure 10 F and CI
abundances in various types of
glassy rhyolites and glass
inclusions in phenocrysts. Data
are from obsidian data of
Macdonald et al. (1992), Bailey
.J1980), 'Christiansen et al.
(1986) and glass inclusion data.
summarized in Lowenstern
(1995). B-Bingham (Cu) lava
flows, Utah; PG - Pine Grove
(Mo) tuff, Utah, glass
inclusions.
5
Tectonic Setting
I-type silicic magmas are found overwhelmingly at convergent plate margins, including oceanic
island arcs and continental magmatic arcs. As a group, I-type rhyolites and granites are largely the
same as those identified by Pearce et al. (1984) as volcanic-arc granites. We have not separately
distinguished sub-types of these granites (e.g., Pitcher, 1982; Macdonald et al., 1992) in this
review because they share so many common characteristics. The sub-types probably represent
variations in the amount of contamination by continental materials.
)~.
/
Example of I-type Magmatism at the Bingham Porphyry Copper Deposit.
The felsic volcanic rocks associated with the Bingham porphyry copper deposit are examples of
131
strongly mineralized, subduction-related magmas. Bingham, Utah, is one of the largest known
porphyry Cu deposits with reserves plus past production of more than 2,000 Mt (Babcock et al.,
1995). Mo, Au, Ag, Pb and Zn have also been recovered. The deposit lies in the northeastern Basin
and Range province and formed far inland in a wide continental magmatic arc that cut across
Archean and Proterozoic basement buried by miogeoclinal sedimentary rocks. The mineralized
stock lies in the central part of an east-tilted, basin-and-range fault block; this extension and tilting
occurred during the late Cenozoic and are unrelated to the generation of the ore deposit. Middle
Tertiary .volcanic rocks, exposed east of the mineralized intrusion, are cogenetic with the Cu deposit
(Keith et at., 1995). The volcanic rocks consist of an older (40-37 Ma) and a younger (35-34 Ma)
series. The older series of rocks contains a diverse assemblage of mafic alkalic lavas
(biotite-bearing melanephelinites) to high-potassium, I-type lava flows, breccias and lahars that
include shoshonites, latites and dacites (K. Waite, unpublished analyses). The younger series of
breccias and lava domes has similar major element compositions, but includes more silicic lavas
including high-silica rhyolite. The Bingham copper deposit developed around a shallowly emplaced
quartz monzonite sto~k that, based on composition and age, appears to represent a subvolcanic
chamber for part of the older series of volcanic rocks. The younger series is apparently unrelated to
mineralization and must have been erupted from vents only slightly farther to the east.
The compositions of the fresh igneous rocks at Bingham show that they have most of the
characteristics of typical volcanic arc rocks, including high large ion lithophile abundances (K, Ba,
Rb), high BaINb and Rb/Nb ratios and high MglFe ratios because of high magmatic f02. The lavas
and fresh plutonic rocks are metaluminous (molar Al203/(CaO+Na20+K20) < 1.0, but molar
Al203/(Na20+K20) > 1.0). In most respects, they are like continental arc rocks worldwide. The
mineralized rocks are not highly fractionated, consistent with their association with Cu
mineralization; Cu is a highly compatible element Glassy rocks that could have preserved magmatic
F/CI ratios are rare in the older volcanic series at Bingham, but whole-rock data have F/CI ratios
that are typically between 1 and 3 (Fig. 10). These ratios are lower than those found in glass from
sub alkalic A-type rhyolites, such as topaz rhyolites. Volcanic glasses from other subduction-related
felsic magmas extend to even lower FICl ratios. Loss of easily mobilized CI after eruption could
account for the higher F/CI ratios in the Bingham lavas when compared with glassy I-type rocks.
For example, biotites from the hydrothermally altered rocks at Bingham show that the
orthomagmatic deposit formed at highjHClIjH20 and low jHFijHCI (Munoz, 1984).
The characteristics of the mineralized series of lavas at Bingham. can be largely, if somewhat
generally, explained as the result of oceanic slab dehydration in a subduction zone, rise of the
buoyant fluid into the overlying mantle inducing partial melting to form diverse mafic magmas that
rose into the crust, assimilating large proportions of the heterogeneous continental 'crust,
fractionating, and experieI}cing repeated magma recharge with diverse magma types limiting the
extent of fractionation. Magma mixing seems to have been especially important for the generation of
the ore deposits at Binghain. Alkalic mafic magmas appeared to have mixed with more silicic I-type
magmas to produce potassic intermediate magmas with high Cu, Cr and Ni contents. These mixed
magmas erupted as shoshonites and latites whose parental magma chambers crystallized to form
quartz monzonite that became Cu (Mo, Au)-mineralized. The elemental effect of mixing is most
clearly seen in plots of highly compatible elements like Cr and Ni against an incompatible element
like Rb (Fig. 11). The alkalic magmas, presumably derived from small degree partial melts of
: I
t
t
E.H. Chrispansen and J.D. Keith
132
metasomatized Precambrian mantle lithosphere, added Cr, Cu, Ni, S, H20, Nb, K, Ba, and
perhaps Cl and C02 to the more fractionated magmas. Alkalic lavas at Bingham have the highest Cu
contents of any of the nearly 300 fresh mafic to intermediate composition lavas from the northern
Basin and Range province (Barr, 1993). Here, it seems that Cu, Cl, and S may have come from
ancient subduction-generated materials stored in veins or dykes in the lithospheric mantle formed in
the Precambrian and returned to much younger subduction-generated magmas of Cenozoic age.
Dehydration of oceanic crust in the Tertiary may have also contributed.
Ores Associated with I-type Magmas
The most important metallic ores associated with I-type igneous rock suites are Cu, Pb, Zn, Au,
and Ag. Skarn deposits of Ware also an important association with I-type plutons on the
continents. Continental I-types are also enriched in Mo compared with those in·island arcs and
granodiorite-type Mo deposits (Mutschler et al., 1981) are commonly developed in I-type plutons.
Moreover, porphyry Cu deposits on the continents, such as Bingham, have substantial amounts of
Mo. Climax-type Mo deposits are associated with siliceous, incompatible element and F-rich,
oxidize!i magmas with some characteristics of A-type granites (White et al., 1981), including high .
F, U, Be, and Rb. In fact, compositions of rhyolites associated with Climax-type Mo deposits in
the western United States (Pine Grove, Utah, and Henderson, Colorado) plot in the A-type field on
1400
•
1200
• •• •
1000
Mixing
.E'
a.
a.
---0
"-
Bingham Utah
•
•
o Younger Series
f~\
til
!
• Minette
• Shoshonite
Trends~\
800
Older Series
!.
\, .
f
600
Fractional
Crystallization
Trend
400
200
f+.
\
f+.
\
;~~
oL-~--~~~~~~~~~~~~~~
.0
50
100
150
Rb (ppm)
200
250
Figure 11 Cr and Rb
concentrations in volcanic
rocks associated with the
Bingham porphyry copper
deposit, Utah (K. Waite,
unpublished data).
The
younger volcanic series is
unrelated to mineralization and
evolved principally by crystal
fractionation, fractionation
trends are lines with ticks.
The older volcanic series is
contemporaneous
with
mineralization and appears to
have resulted from mixing
between Cu-rich mafic alkalic
magmas .(minette) and a
fractionating sequence of more
normal I-type magmas.
Mixing produced shoshonites
and latites that are enriched in
compatible trace elements.
Pearce diagrams (Fig. 7). The distinction between evolved A-types and I-types is difficult because
extensive fractionation produces common features in both types of rhyolites (see Pearce et al.,
1984). However, the highf02 (and the attendant mineralogy), spatial and temporal association with
normal calc-alkalic rhyolites, and low FICl ratios (Fig. 10) in melt inclusions from the
Mo-mineralized Pine Grove magma system (Lowenstern, 1995; Keith and Shanks, 1988), suggest
•
41
41
41
41
41
41
••
t
t
t
t
•••
•
••
••
t
(I
••
••
•••
••
••
••
••
••
•••
••
••
••
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••
133
....
.•.
•,.
•..
•
'"
••
••
••
•
••
•
••
••
••
••
••
••
••
III
that these magmas may be highly fractionated I-type magmas, rather than A-type magmas.
Blevin and Chappell (1992) and Blevin et al. (1995) persuasively show the relationships
between I-type plutons and Cu, Pb, Zn, Mo, W, and Au are the result of source, oxygen fugacity,
and extent of fractionation. For example, source control is illustrated by the observation that I-types
are much more strongly associated with Cu-Au-Mo than with Sn, which is associated with S-types.
In this review, the sources (and contaminants) of most I-type magma series are taken to be mixtures
of metasomatic fluid derived from a subducting slab (high Cl, some metals, jHzO, and fOz),
mantle-derived magmas (high Cu, Ni, Au), and continental crust (probable source ofMo and Pb).
Varying proportions of these metal sources control the potential ore deposits that can form from a
magma. High magmatic concentrations of HzO and Cl help subduction zone magmas produce
ore-laden hydrothermal fluids. The role of magmatic fractionation can be seen in the association of
Cu deposits with the least fractionated, most oxidized rocks, the association of W deposits with
moderately fractionated rocks, and Mo and Sn with the most fractionated rocks (Fig. 12) .
Apparently, Cu is depleted by fractional crystallization processes and W, Mo, and Sn are enriched.
During the crustal residence of subduction zone magmas, there are great opportunities for
hybridization or miXing with repeated injections of mantle-derived magma, forestalling magma
evolution by fractional crystallization and keeping compatible element concentrations (e.g., Cu)
high enough for mineralization and reducing the opportunity for development of incompatible
element (e.g., Mo) ores. Finally, the strong enrichment of Mo in oxidized I-types and lack of
enrichment of Sn is consistent with the inferred effect of oxygen fugacity on the behaviour of these
elements in fractionating magmas .
S-Type Silicic Magmas
S-type granites occur as isolated plutons or in batholithic belts. Some plutons in both associations
have wide contact metamorphic aureoles and were characteristically emplaced at relatively great
depths. Strongly peraluminous rhyolites of Cenozoic age are rare. Our understanding of the "
compositional characteristics of this important group of magmas is limited to studies of granites and
rare, mostly altered volcanic rocks (e.g., Clemens and Wall, 1984; Zeck, 1970). The glassy
volcanic rocks from the Macusani region of Peru (Noble, 1984; Pichavant et al., 1988a, b) are
exceptions and are used here to represent some characteristics of fractionated S-type magmas.
The single most important elemental characteristic of S-type silicic magmas is that they are
strongly peraluminous (molar Al z0 3/(CaO+NazO+KzO) >1.1). This characteristic is revealed in
their distinctive mineral assemblages including cordierite, andalusite, muscovite, or garnet and as
accessories tourmaline, ilmenite, monazite, apatite and zircon (Clarke, 1981; Zen, 1988). Biotites
are typically more Fe-rich and aluminous than those of 1- or A-type magmas. S-type silicic magmas
are characteristically enriched in K (and Rb) and depleted in Na compared with I-types of similar
SiOz content (Chappell and White, 1992). The depletion ofNa explains why they are peraluminous
and is thought to be inherited from Na-depleted shales or greywackes in their sources. During
weathering at Earth's surface, Na (as well as Ca and Sr) is extracted from solids and transported to
seawater. At least some K and Rb is stored in clay minerals in shales and other sediments.
Consequently, sedimentary rocks commonly have higher KINa and ~blNa than their unweathered
protoliths (Taylor and McLennan, 1985). When these same sedimentary rocks participate in magma
genesis, as magma sources or contaminants, these characteristics are passed on to the derivative
.1
4
E.H. Christiansen and J.D. Keith
8r-----------------------~
6 I-
E
a.
a.
~4
4
4
4
,
••
,,•
•
t
••
••
••
•••
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••
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••
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4
411
2
66
70
68
72
74
76
78
Si02(wt%)
~r-----------------------~
o
16
o
(]])
o
<>
<>
<>
8
o
4
66
68
9 o0
o
•
70
72
74
76
78
Si02(wt%)
100
r------------------.
I-type
•
o
o
80
S-type
•
E
.9;
A-type
60
o
.c
a..
Peralkaline
40
<>
Plume
20
66
o
.0
68
70
72
74·
76
78
Si02 (wt"Io)
Figure 12 Concentrations of various ore elements as a junction of Si02 content for various types of rhyolitic
obsidians (Macdonald et al., 1992). The highly evolved Macusani glass is the only S-type obsidian shown and its
concentrations of Sn (250 ppm) and W (76 ppm) are too high and Mo content (0.3 ppm) too low to be shown.
A-type rhyolites are generally the most,enriched in the incompatible ore elements, except for Mo, where evolved
I-type rhyolites are just as enriched, anq the reduced Macusani glass most depleted.
.•.
.-
~
~
=
~
=
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-'
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•,..••
r'
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.
,
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magmas. Less soluble elements like Nb and Y, are not strongly mobilized by weathering processes,
but they may become enriched as residual elements in mature shales. Since most continental rocks
have a heritage that traces their origins to a subduction zone, most immature sedimentary sequences
also have the hallmarks of arc magmas, namely high Ba!Nb, Rb/Nb, and LalNb ratios.
Consequently, in the discriminant diagrams of Pearce et aI. (1984), S-type granites are
distinguished by high Rb compared with I-type granites. Nb and Y may be enriched compared with
I-types, by the residual processes noted above and by partial melting and fractionation processes
(Fig. 7).
The aluminous compositions of these magmas also explain other compositional features. For
example, many S-type magmas are strongly depleted in REE, Zr, c:md Th compared to comparably
siliceous I-types. These features are probably not source related but are the result of the low
solubility of accessory phases, such as zircon, monazite and thorite in peraluminous melts.
Depletion in these trace elements is probably a two-stage process, occurring during partial melting
events that leave accessories in the residue and continuing during fractional crystallization of the
magma. In highly evolved, peraluminous magmas, xenotime and columbite lave low solubilities
and may fractionate to further increase the already high Rb/(Nb+Y) ratio and, in extreme cases,
driving the residual melt composition to the left on Pearce diagrams (Fig. 7). Such a process may
explain the relatively low Nb and Y concentrations of Macusani rhyolites. On the other hand, apatite
is highly soluble in peraluminous siliceous magmas; this controls the concentration of P20S in the
melt (Pichavant et aI., 1992). For example, Chappell and White (1992) have shown that P20S
decreases during the fractionation of I-types much more rapidly than in S-types, which may even
show P enrichment (Fig. 13).
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Figure 13 Phosphorous contents of S-type granites and rhyolites are significantly higher than those of other
granite types because of greater solubility of phosphate minerals in peraluminous magma. Data from Chappell et
al.(1991), Macdonald et al. (1992), Pollard et al. (1995), Pichavant et al. (1988a), and Clemens and Wall (1984).
~;
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136
E.H. Christiansen and J.D. Keith
:
High concentrations of B and the presence of tounnaline are particularly good indicators of
S-type silicic magmas. B is strongly concentrated in sediments (approximately 100 ppm) compared
to the mantle (0.1 ppm; Ryan and Langmuir, 1993). Apparently this characteristic is passed on to
metasedimentary rocks and to magmas derived from them. Highly evolved A-type rhyolites have
much lower B and tourmaline is rare in their intrusive equivalents. The A-type Spor Mountain
vitrophyres of western Utah, for example, have less than 100 ppm B and whereas peraluminous
glasses from Macusani have over 2000 ppm (Pichavant et aI., 1988a), but both have comparable
enrichments of Rb and Be, elements with similar partition coefficients.
Another important characteristic of S-type granites and rhyolites is low f02. Many of these
granites lack magnetite and ilmenite is the only Fe-Ti oxide. This fact was recognized by Ishihara
(1977)~ but, as he points out, not all ilmenite-series granites are S-types, many are contaminated
and reduced I-types. Low oxygen fugacities can also be inferred from the low Fe203/FeO ratios
(Fig. 9; Blevin and Chappell, 1992) and high Fe/Mg in mafic silicates. The compositions of
coexisting minerals in the S-type Macusani volcanic rocks of Peru also show that they crystallized
at low f02, below QFM (Pichavant et aI., 1988b ).
Isotopically S-types have the characteristics of the crust from which they are largely derived,
including high 81 80 (reflecting enrichments during a weathering cycle; Table 1), and generally high
Sr and low Nd isotope ratios (Chappell and White, 1992). However, Sr and Nd isotopic
compositions are sensitive to the age of the source, not just the parent/daughter ratios.
Consequently, S-types derived from young sedimentary rocks do not have high Sr isotope ratios.
The S-types granites of the New England fold belt of eastern Australia illustrate this relationship
(Shaw and Flood, 1981).
Tectonic Settings
S-type granites are found in several settings that are generally related to plate convergence. Perhaps
the best constrained are the S-type granites that clearly formed because of the collision of India and
Asia. Moreover, many S-types are synkinematic and foliated. These facts led to the notion that
S-type granites are syn-collisiorial granites, as in the geochemical classification of Pearce et al.
(1984). However, abundant S-types are also found in other plate convergent settings, where they
formed as parts ofI-type suites. The S-type granites may slightly pre- or post-date the other granitic
rocks in a given area. Most of these associations formed far from the convergent plate margin
(pitcher, 1982) and in regions underlain by thick sequences of metasedimentary rocks including
metamorphosed shales, greywackes, and sandstones. Some members of the strongly peraluminous
granites of the western Cordillera of North America are examples of this setting (Miller and
Bradfish, 1980; Lee et al., 1981) and eastern Asia (Takahashi et aI., 1980), although some
muscovite:..bearing plutons are more appropriately seen as strongly contaminated I-types because
they lack all of the characteristics of S-type magmas. Pitcher (1982) separates these tectonic settings
into a Hercynotype (collisional) and Andinotype (subduction without large collisions). In the latter,
S-types form a much smaller proportion of the granites.
The cause of the melting that produces S-type silicic magmas varies by tectonic setting. Those
that form above still active subduction zones probably form as a result of melting induced by mafic
mantle-derived magmas, with which many undoubtedly mixed and hybridized both before and after
silicic magma was produced. Other S-types lack a clear association with such mafic magmas and
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probably formed as a result of volatiles flowing from a zone of metamorphic dehydration to higher
levels of the crust (Fyfe and Kerrich, 1985), as a result of heating following tectonic thickening of
continental crust during collision (pitcher, 1982), or as a result of decompression following tectonic
unroofing of collisional orogenic belts (Hodges, 1990).
Examples of S-type Rhyolites and Granites
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The granites associated with the large W skarns of western Canada, MacTung and CanTung, are
S-type granites, and contain biotite, muscovite, and gamet (van Middelaar and Keith, 1990).
MacTung reportedly contains 63 m.t. of ore with 0.95% W03 and CanTung contains over 6 m.t. of
ore at 1.6% W03 (Einaudi et ai., 1981). Granites that caused both deposits lack magnetite and have
Fe-rich biotites. Rb concentrations are moderately high, indicating extended fractionation and
potential enrichment in W set the stage for ore generation. CanTung granites are more fractionated
than those at MacTung and are related to the higher grade W deposit. Nb and other high field
strength elements, especially Zr, have low concentrations and the granites plotin the S-type field of
Figure 7. Biotite compositions reveal that these plutons crystallized and became mineralized at
relatively highjHFifHCI compared with I-type granites (van Middelaar and Keith, 1990).
The Sn-mineralized Phuket Supersuite of southwest Thailand displays many characteristics of
S-type granites and illustrates some processes that are important in the' generation and mineralization
of such granites. The evolved Sn-mineralized plutons of the suite have high 87Sr/86 Sr ratios
(>0.710), peraluminous compositions (molar Ah03/(CaO+Na20+K20) = 1.05-1.17), 10wNa/K
ratios, low 102 with ilmenite and rutile as the characteristic Fe-Ti oxides, and gamet, monazite and
tourmaline. Some granites are also topaz-bearing. The trace element characteristics of these granites
are like those of other S-type granites with high Rb/(Nb+Y) (Fig. 7), high B (20 to 1300 ppm),
low Zr (15 to 90 ppm), low REE concentrations, and high P205 (about 0.1 wt% in evolved rocks
with about 1000 ppm Rb). These arethe characteristics of a fractionated magma with substantial
proportions of a weathere~ sedimentary protolith. Such are cool peraluminous magmas with :low
solubilities of monazite and zircon and high apatite solubility (pichavant et ai., 1992). Nonetheless,
Pollard et al. (1995) conclude that the granites in this association resulted from extensive crystal
fractionation of I-type magmas and were not derived from a metasedimentary protolith. As Pollard
et al. (1995) argue, high 87Sr/86 Sr ratios may be derived from meta-igneous rocks, but it is difficult
to understand how such a protolith could produce strongly peraluminous magma with the observed
phase stabilities, fractionation trends, high B and P, very high RblNb ratios, and a f02 low enough
to preclude magnetite crystallization and induce Sn incompatibility. Without the inclusion of
sedimentary materials, closed system fractionation of an I-type silicic magma, produces residual
magmas that are metaluminous, low in B, P205, and Sn, and relatively oxidized. Unevolved and
unmineralized portions of the supersuite are oxidized I-type granites. Consequently, the
compositional characteristics of the mineralized plutons of the Phuket supersuite probably
devdoped as sedimentary materials were incorporated in mantle-derived I-type magmas. The
Phuket granites are probably typical of many suites that evolve from subduction-related I-type
magmas to S-type magmas as high proportions of weathered, carbon-bearing sedimentary rocks are
assimilated.
,
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138
E.H. Chri*ansen and J.D. Keith
Ores Associated with S-type Magmas
Many important ore deposits are associated with S-type granites. The most characteristic are Sn and
W. The Sn deposits of southeastern Asia are associated with S-type granites (Beckinsale et al.,
1979). And as noted above, W deposits in the Canadian Cordillera are associated with
peraluminous muscovite-garnet granites with the trace element characteristics of other S-type
granites. Be, Li, and Ta are important in phosphate-rich pegmatites derived from strongly
peraluminous magmas, such as the Greenbushes pegmatite, western Australia, which contains half
of the world's Ta resources (Partington et al., 1995).
. The principal controls on the ore associations of S-type granites include the modest enrichments
of these ore elements in the weathered continental crustal materials from which the granites are
domimintly composed. Most other granite types contain significant fractions of mantle-derived
dilutants to these generally incompatible elements. Perhaps the most important control for the
abundance of Sn deposits, however, is the low 102 and not any substantial Sn enrichment in the
source materials (Lehmann, 1982). This condition probably reflects the incorporation of organic
carbon derived from carbonaceous sedimentary materials. Sn behaves as an incompatible element
and becomes strongly enriched during fractionation of reduced magmas in contrast to strongly
fractionated but uncontaminated I-type granites. Magmatic fractionation must be an important
control on the enrichment of W in S-type granites, though the incompatibility of W is not as
sensitive to extreme oxidation state as either Sn or Mo (Keith et al., 1993). Another important
control is the peraluminous character of the magma which increases the solubility of phosphorous .
(pichavant et al., 1992), allowing phosphate-rich rare element pegmatites to evolve by fractional
crystallization or very small degrees of partial melting. Low Nb/Ta ratios may be achieved by
extended fractionation, since niobium-rich minerals begin to crystallize in low-temperature
peraluminous magmas. Unfortunately, there are few peraluminous obsidians to compare with the
A- and I-type obsidians in Figure 12. The extremely high conc;entrations of Sn (250 ppm) and W
(76 ppm) in the Macusani glass (Macdonald et al., 1992) are probably not representative of most
S-type granitic magmas. On the other hand Mo contents (0.3 ppm) were much lower in the
fractionated but reduced Macusani magma as compared to fractionated and oxidized I-types, which
can contain as much as 8 ppm Mo (Fig. 12).
A· Type Silicic Magmas
Of all the silicic magma types considered here, A-type granites and rhyolites may be the most
diverse and, perhaps, the least understood. The A has variously connoted alkalic (as in peralkalic),
anorogenic (as in rifting- or plume- as opposed to subduction-related), or even anhydrous, although
each of these descriptive terms is questionable if applied universally. First, a major category of
these rockS is metaluminous or marginally peraluminous, not peralkalic (molar Al20yf(Na20+K20)
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< 1), although all have high alkali contents compared with 1- or S-types. Second, the use of the
term "extensional orogeny" makes it difficult to call silicic magmas "anorogenic," if they form
during episodes of crustal extension and rifting; the meaning is clear when used in the petrologic
context to contrast these rocks from "orogenic" andesites and rhyolites like those found above
subduction zones. Third, a large group of these rhyolites and granites are found in orogenic belts
and formed shortly after the cessation of contractional orogenies. These post-orogenic rocks are
difficult to categorize. Finally, highly fractionated A-type magmas are water-rich, although it is
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probably important that their sources were relatively water-poor.
Although we categorize rhyolites and granites of this class as peralkalic or sub alkalic
(metaluminous to slightly peraluminous), both classes share some important compositional
features, including high Si02, Na20+K20, Fe/Mg, and low Al203, CaO, and P205. Most A-type
magmatic rocks have Si02 contents greater than 68%. Compared with granites of similar Si02
content, common trace element characteristics are low Rb/Nb ratios (Fig. 6), the absence of deep
Nb-Ta anomalies on chondrite-normalized trace element diagrams (Fig. 8), enrichments of many
high-field strength elements (Nb, Ta, W, Y and especially Zr and Hf in the peralkalic varieties),
enrichments in Be, Zn, 9a, Sn, Pb, Th, U and the halogens (high F/CI in aluminous types and
lower F/CI in peralkalic types, Fig. 10), and characteristically lo~ concentrations of compatible
elements (Mg, Ca, Ti, Sr, Eu, Ba) (e.g., Collins et al., 1982; Anderson, 1983; Whalen et al.,
1987; Ramo and Haapala, 1995; Macdonald et al., 1992). Typical A-type silicic rocks also have
low Y/Nb and high GalAI ratios.
A wide range off02 values has been recorded for A-types, from near QFM to several log units
above that (e.g., Anderson and Bender, 1989). In the Macdonald et al. (1992) compilation of
obsidian analyses, Fe203/FeO ratios of nonperalkalic A-type rhyolites are comparable to those of
I-types (Fig. 9), although Fe/Mg ratios are consistently high. Peralkalic rhyolites probably have
high Fe3+/Fe2+ ratios because of the effect of alkalis in stabilizing Fe 3+ in silicate melts and not
because of an inherently highf02 (Kress and Carmichael, 1991). Volatile fugacities in peralkalic
rhyolites are poorly constrained because the mineral assemblages are largely anhydrous and Fe-Ti
oxides are not stable in many peralkalic melts. However, the f02 in most peralkalic rhyolites is
probably near QFM, as judged from Fe20~eO ratios and by the common appearance of fayalite
and quartz. Absolute, water concentrations in peralkalic rhyolites are controversial. Kovalenko et
al. (1994), from Pantelleria, and Webster et al. (1993), from Fantale, Ethiopia, have reported as
much as 4 wt% H20 in melt inclusions included in quartz in peralkalic volcanic rocks. However,
Lowenstern and Mahood (1991) found less than 2% H20 in their glass inclusions from Pantelleria.
Few hydrous minerals are stable in peralkalic melts; amphibole is one of the few and in sub alkalic
magmas amphibole is stabilized at higher dissolved water contents than biotite. Most A-type
rhyolites and granites have high concentrations of fluorine. For example, Precambrian A-type
granites from the southwestern United States have significantly higher concentrations of F than any
of the Phanerozoic granite types from that region (Christiansen and Lee, 1986). An important
characteristic of peralkalic magma is a low F/CI ratio. In an absolute sense, many peralkalic silicic
magmas have higher concentrations of F than many sub alkalic magmas, but they are consistently
more enriched in Cl and consequently have low FICI ratios (Fig. 10), sometimes as low as those in
glassy I-type magmas but IIl;ore generally like those of mafic magmas derived from mantle. plumes.
High FICI ratios are not m~aningful in granites or in most volcanic rocks, because halogens, like
water, have few mineralogic sites in fully crystalline rocks. Granites and crystalline rhyolites
generally have FICI ratios that are higher than the magma from which they crystallized because F is
more compatible in mineral structures than Cl. Fluorine is hosted by fluorite, micas, apatite, topaz,
and in other silicates, where it may replace 0, in greater quantities than Cl. Consequently, useful
comparisons of the halogen contents of silicic magmas can only be made with glassy volcanic
rocks, and even these may have lost halogens during eruptive degassing (Fig. 10). Glass inclusion
I
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140
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E.H. Christiansen and J.D. Keith
studies may provide better estimates of the magmatic concentrations of these mobile elements; those
done so far agree with bulk rock studies and show that sub alkalic magmas have FICI ratios that
exceed 1 (Webster and Duffield, 1991; X.Y. Zhang and E.H. Christiansen, unpublished).
Many of the sub alkalic group of A-type granites become enriched in heavy REE (HREE) relative
to light REE (LREE) during fractional crystallization. In most I-type magmas LREEIHREE ratios
increase during fractionation because of the smaller partition coefficients for the slightly larger
LREE. Consequently, when this. trend was first observed (Hildreth, 1979), liquid-state
differentiation processes were invoked, such as thermogravitational diffusion or the Soret effect.
These processes have been shown to be unimportant in natural silicic magmas. In A-type rhyolite
magma, the principal control on changing LREE IHREE ratios is the accessory mineral assemblage,
which commonly includes several minerals with strong affinities for LREE, such as allanite,
chevkinite, xenotirne, huttonite, or monazite.
Many features of A-type granites and rhyolites could result from extreme fractionation. Note the
overlap of many incompatible element concentrations on Figure 12. Thus,-the question naturally
arises, are these rhyolites simply fractionated I-types or S-types? Sometimes these disti:nctions
between A-types and these other granite types are clear, but ambiguities are more common. Whalen
et al. (1987) conclude that high GalAl ratios can be used to distinguish fractionated A-type granites
from the other types. Most A-type granites have 1000xGalAl > 2.6, although some fractionated Sand I-types may exceed this discriminating value. They also use Zr+Ce+Y > 250 ppm as a
dis9rirninating parameter. Although this feature may be characteristic of fractionated alkalic A-types,
the sub alkalic types have low solubilities for the accessory phases (zircon, allanite, monazite and
xenotime) that control the concentrations of these phases and Zr+Ce+Y commonly drops during
fractionation. We have found that low RbINb ratios (less than 7) are a simpler way to distinguish
most A-type granites from I- and S-types (Keith et al., 1993) and this ratio is less sensitive to the
effects of fractionation or contamination (Fig.s 3 and 5). However, niobates form in· some highly
fractionated A-type granites and may raise RbINb ratios to higher values. For example, highly
evolved rhyolites with topaz and complex Nb oxides such as the Honeycomb Hills rhyolites of
western Utah (Congdon and Nash, 1991) and some ongonites from the Mongolian type locality
(Kovalenko and Kovalenko, 1976) have Rb/Nb ratios that exceed 15. High FICI ratios and low j02
may also be helpful in distinguishing some A-type granites and rhyolites from I-types. Highly
fractionated S-type magmas may become F-rich and crystallize topaz, but they can usually be
distinguished from topaz-bearing A-types by much higher P20S and B contents that are manifest in
abundant phosphate minerals and tourmaline in S-type leucogranites and pegmatites.
Tectonic Settings
A-type rhyolites and granite;s are found in various tectonic settings. Such silicic magmas are
characteristic of those erupted above mantle plumes, both on the continents (the Snake River Plain
and Yellowstone Plateau) and in the ocean basins (Iceland). However, the most common
association of A-type silicic magmas on the continents is with rifting, where they are the silicic
end-members of bimodal volcanic associations. This association is especially common in the
western United States where lithospheric extension since about 20 Ma has been accompanied by the
eruption of bimodal suites of basalt and rhyolite, with rare intermediate magma (Christiansen and
Lipman, 1972). Peralkalic rhyolites and sub alkalic ,rhyolites and granites with topaz are the
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141
characteristic products. The tectonic relations of the Proterozoic granites of A-type affinities that
extend in a wide belt from southwestern North America to the Baltic Shield remain controversial,
although good evidence for crustal thinning, extension, and mafic dyke intrusion has been
described for the rapakivi granites of southern Finland (Ramo and Haapala, 1995). The tectonic
association of "post-orogenic" A-types is also problematic. For example, the late Precambrian
granites of the Arabian shield are A-types and formed during the [mal stages or immediately after
the collisional accretion of the Pan-African Orogeny (Stoesser and Camp, 1985). The A-types of
southeast Australia appear to have formed in a similar late orogenic association, but here evidence
of bimodal magmatism and at least local extension is found (Collins et at., 1982). The main item
these settings have in common is a lack of input of mafic magma from a subduction zone. Instead,
the mafic magmas are the result of processes related to lithospheric extension andlor mantle plumes.
Examples of A-type Rhyolites and Granites
The rhyolites of the Spor Mountain Formation exemplify the sub alkalic group of A-type silicic
magmas. The Spor Mountain Formation consists of topaz-bearing rhyolite erupted about 21 Ma
(Lindsey, 1982) and marks the initiation of extension in the Basin and Range province in western
Utah. Hosted within a tuff and lava flow complex, is North .America' s largest deposit of Be. Davis
(1984) estimates that the total resource is at least several million tons of bertrandite ore that averages
about 0.7% BeO. Nearby, a topaz-bearing granite of the same age also contains topaz and magmatic
(?) beryL The rhyolite vitrophyres are enriched in Be (about 60 ppm) and F (as much as 2% in
glass inclusions preserved in quartz phenocrysts and 1.2% in vitrophyres) and consequently
devitrified rocks commonly have topaz as a vapour phase mineral. FICI ratios in the glasses are
characteristically high (Fig. 10). Magmatic phases include smoky quartz, ~anidine, oligoclase,
Fe-rich biotite and accessory zircon, monazite, thorite, uraninite and complex oxides of
Fe-Ti-Mn-W-Nb-Ta. The magmatic.fO.z was apparently low in the Spor Mountain rhyolite (near
QFM), but other aluminous A-type rhyolites were more oxidized and crystallized magnetite and
titanite. On Pearce discrimination diagrams (Fig. 7), these rhyolites plot within the A-type granite
field, but as a group Cenozoic topaz rhyolites from the region straddle the A-type and S-type
boundary. Isotopically the Be-mineralized rhyolite and granite have low 8 18 0 (6 to 10 per mil),
relatively low initial Sr isotope ratios (0.706), and high ENd (-6). This latter observation is
important because S-type, and even some I-type, granites in the same region have ENd as low as
-18 (Farmer and DePaolo, 1983) and interacted extensively with the early Proterozoic or Archean
crust that underlies the region. A large component of mantle-derived basalt must be present in the
magma, probably because of basalt intrusion in the crust followed by partial melting, although this
interpretation is non-unique. Extensional tectonics and attendant mafic magmatism, intimate
hybridization of the lower (?) continental crust, coupled with low degrees of partial melting «10%)
and an extended fractionation history, explains most of the isotopic and trace element characteristics
of sub alkalic A-type granites in general.
The peralkalic A-type rhyolites have even higher concentrations.ofhigh field strength elements
like Zr, Hf, Nb and REE than the subalkalic varieties (Fig. 8). These enrichments are partially the
result of source enrichments because peralkalic silicic magmas generally arise from mafic magmas
derived from mantle unaffected by subduction zone processes. As noted above, the oxidized,
hydrous, silicic (?) magmas generated above subduction zones are depleted in Nb, as compared
II
;11
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142
E.H. Christiansen and J.D. Keith
with other incompatible elements and consequently have negative anomalies on
chondrite-normalized trace element diagrams. The enrichment of high field strength elements,
especially Zr and Hf, also results from the high solubility of many accessory minerals in peralkalic
magmas (Watson and Harrison, 1983; Montel, 1986). For example, zircon solubility is increased
by 4 to 10 times in peralkalic rhyolite as compared with typical metalurninous rhyolite at the same
temperature and water content. This effect is obvious when Zr, with its variable behaviour, is
compared with an element like Rb that is incompatible in both peralkalic and sub alkalic silicic
magmas (Fig. 14). Likewise, monazite, apatite, allanite, tho rite and columbite-tantalite either are
unstable or have high solubilities in typical peralkalic magmas (e.g., Montel, 1986). Consequently,
LREE, Th, Nb, and Ta (trace elements accommodated in these accessory minerals) all behave as
strongly incompatible elements, in contrast to their moderately incompatible or even compatible
behaviour in many sub alkalic A-type magmas. Most peralkalic rhyolites show evidence for
extensive fractional crystallization. Ca, Mg and P are strongly depleted, almost to the point of being
trace elements in some strongly peralkalic magmas. Feldspar-compatible trace elements always have
extremely low concentrations.
The compositions of peralkalic rhyolites from the McDermitt Caldera, northern Nevada (Rytuba
and McKee, 1984), are shown in Figures 7 and 8. The Miocene McDermitt caldera complex formed
contemporaneously with the onset ·of Basin and Range extension in the region. RiftIng may have
been related to the initial rise of the plume speculated to presently lie beneath the Yellowstone
Plateau (Zoback et at., 1994). Most of the silicic magmas are mildly peralkalic ash-flow tuffs and,
lava domes and their compositions fall in the A-type granite field on Pearce discrimination diagrams
because of their high concentrations of both Y and Nb (Fig. 7): However, the Rb concentrations
are generally lower than in the aluminous A-types with comparable Y + Nb concentrations. FICI
ratios in glassy peralkalic samples are lower than in the aluminous varieties, and more like those
found in plume-related rhyolites and basalts (Fig. 10) Chondrite-normalized trace element patterns
illustrate the lack of Nb and Zr depletions (Fig. 8), that reflect the high solubilities of the controlling
. accessory minerals and a probable mantle derivation from non-orogenic magmas coupled with little·
continental contamination. Mantle-like Sr and Nd isotope ratios are consistent with this suggestion
. (Tegtmeyer and Farmer, 1990). Extensive fractionation of this mantle-derived magma apparently
caused the enrichments of incompatible elements. Barium, Sr and Eu are strongly depleted by
extensive feldspar fractionation from less siliceous parental magmas. The McDermitt caldera
complex is highly mineralized with deposits of Hg, Sb, Cs, Li and U. As a result of extensive
fractionation, the peralkalic tuffs had high contents of these elements and were the source rocks
from which these metals were leached by hydrothermal systems late in the development of the
caldera (Rytuba and McKee, 1984). Peralkalic A-types generally have lower U and Th than the
subcilkalic A-type rhyolites and granites of the continents and, in this regard, are similar to those
found on ocean islands.
Silicic igneous rocks found on ocean islands related to mantle plumes, though rare, have many
characteristics of A-type granites found on the continents including high concentrations of high field
strength elements, including especially high Nb and Ta, and volcanic series that include peralkalic
rhyolites. The compositions of these oceanic silicic rocks plot in the within-plate granite field of
Pearce et at. (1984) and the A-type field Whalen et at. (1987). They are nonetheless distinctive from
A-types found on the continents. A typical feature of these rhyolites is lower concentrations of large
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600
Rhyolite Obsidians
•
•
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i
•
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Peraluminous
Others
I,
400
200
200
400'
600
,..
800
I
Figure
14
Rb and Zr
concentrations in peraluminous
compared with peralkaline
rhyolites. Apparently, Rb is an
incompatible element in
peralkaline and peraluminous
silicic magmas, whereas Zr is an
incompatible element in
peralkaline magmas, but is a
compatible element in
peraluminous magmas in which
Zr concentrations (as low as 50
ppm) are limited by the
insolubility andfractionation of
zircon.
Modified from
Christiansen et al. (1984) by
addition ofdata from Macdonald
etal.(1992).
1000
Zr (ppm)
,.
lie
••
•
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•
..
....
..
..
•.•
lilt
ie
lilt
ion lithophile elements, like Cs and Rb, than other siliCIC magma type, excepting the rare granites
found at normal mid-ocean ridges. Consequently, these rhyolites occupy a distinctive part of the
RblNb-Rb diagram shown in Figure 6, with nearly chondritic Rb/Nb ratios. They are more
enriched in Rb than OR-type silicic magmas and have higher RblNb ratios. Other distinctive
features of oceanic plume-type rhyolites are their high REE (including Eu), Zr, Hf, Sn and Zn
contents in both peralkalic and sub alkalic rocks and low Pb, Li, Rb, an~ Cs.
Rhyolites from Iceland and the Galapagos islands typify the oceanic plume association. These
rhyolites are all metaluminous or peralkalic. Their characteristics are illustrated on the Pearce
discrimination diagrams (Fig. 7) and on chondrite-normalized trace element diagrams (Fig. 8). F/Cl
ratios in melt inclusions and obsidians from metaluminous silicic rocks of oceanic islands are about
1 or higher. Peralkalic rocks from these oceanic islands have lower FICI ratios of about 1 to 0.5,
typical of the more mafic rocks on oceanic islands (Schilling et al., 1980) consistent with a direct
link between the basalts and rhyolites by fractional crystallization. In fact, most studies of these
rhyolites emphasize the role of fractional crystallization from basaltic parents through trachytic
intermediate compositions. The low concentrations of large ion lithophile element, compared with
continental A-types (and even many 1- and S-types), are likely the result of the absence of
continental crustal contamination of subduction zone fluids.
Sources of A-type Magmas
These two varieties of A-type
silicic magma probably have fundamentally different proportions of
I
source materials (e.g., Christiansen et al., 1986; Eby, 1992). The peralkalic varieties almost
certainly are dominated by mantle materials. This is supported by their commonly low initial Sr
isotope ratios and high Nd isotope ratios (e.g., Kovalenko et al., 1995; Tegtmeyer and Farmer,
1990). From the perspective of trace element concentrations, it would be difficult to produce a
silicic magma with no Zr depletion by partial melting any type of crustal material. The characteristic
lack of Zr anomalies in peralkalic rhyolites is strong evidence that they evolved from mafic magmas
with little crustal input. Likewise, the absence of negative Nb anomalies in these rocks argues not
\.
. I
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, '
p
144
E.H. Christiansen and J.D. Keith
only against the involvement of contemporaneous subduction-derived magma, but also against the
incorporation of continental materials in generaL Moreover, the low FICI ratios characteristic of
peralkalic rhyolites are also consistent with a mantle origin. Continental crustal materials generally
have high FICI ratios because of the absence of mineralogical sites for Cl in the common hydrous
minerals. Biotite, amphibole, and apatite in crustal rocks are almost always F-rich. This is
particularly true of high-grade metamorphic rocks such as those proposed to be the source of
A-type granites (Filippov et al., 1974). Many sub alkalic A-types probably evolved from magmas
that included large proportions of crustal material (e.g., Collins et al., 1982; Christiansen et al.,
1988; Clemens et al., 1986; Anderson and Bender, 1989; Whalen et al., 1987; Ramo and Haapala,
1995). This is suggested by their Nd and Sr isotopic compositions, substantial depletions of Zr,
and high FICI ratios. However, the trace element patterns, similar incompatible element ratios in
silicic and adjacent mafic magmas, and isotopic compositions argue that partial melting occurred in
a zone where abundant mafic magmas lodged and hybridized with the crust before these small
degree melts were produced and extracted (e.g., Hildreth et al., 1991). Unless melt extraction
from, or metamorphic dehydration of crustal protoliths can produce both low RblNb and RblSr
ratios and high SmlNd ratios, a significant mantle input is suggested for most continental A-types.
The fact that these same features also characterize A-type silicic magmas in the ocean basins is
consistent with a mantle heritage. The extensional tectonic setting of many of these magma systems
allowed small melt fractions to segregate easily into dykes and did not promote extensive mixing
between crustally derived melts and the contemporaneous basalts, as is characteristic of I-type
magma series generated in subduction zones in which intermediate composition magmas are
common.
One of the most important processes that creates the trace element characteristics of sub alkalic
A-type magmas is a small degree of melting of a hybridized crustal protolith. Melt fractions of less
than 15% have been invoked in several studies as leading to large initial enrichments in
incompatible trace elements (e.g., Christiansen et al., 1988). Assuming batch partial melting, only
m.oderate depletions of compatible elements are produced and consequently primitive A-type melts
may have high Ba for example. Small melt fractions may be the natural result of melting being
induced by the decomposition of small amounts of biotite or amphibole in a nearly hybridized with
volatile-poor mafic magII}a or in a high-grade metamorphic protolith. Subsequent strong
fractionation in closed or open systems appears to be required to create the observed depletions in
many A-type rhyolites and granites.
An extensional tectonic environment is probably critical for the segregation of small volumes of
melt from their protoliths. In such an environment, dykes could rapidly fill with newly formed
magma and be drained out of a crustal source region heated by basalt intrusion. These melts appear
to rise without substantial mixing with the mantle-derived magmatic heat sources because of the
presence of extension-related crustal fracture systems and, in some cases, relatively low flux of
mantle-derived magma. In compression environments, the accumulation of a relatively large volume
of buoyant magma may be necessary to overcome the strength of the rock and induce rise and
separation. Before this can occur, new mafic magmas may be added to~ the melting region.
Apparently in supra-subduction zone settings, magma mixing is common and suppresses the
chemical effects of fractionation.
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145
Ores Associated with A-type Granites and Rhyolites
Ore deposits directly related to peralkalic silicic magmas are sparse, but include a variety of
elements that become concentrated by fractional crystallization in peralkalic magmas but lack
insoluble mineralogical hosts. Notable ore deposits include, Hg-Sb-Cs-Li-U deposits in the
peralkalic McDermitt caldera of Nevada (Rytuba and McKee, 1984), epithermal Au deposits are
associated with sub alkalic rhyolites in Oregon and peralkalic rhyolite and trachtye intrusions in the
Black Hills of South Dakota, Be-Zr-Nb-REE mineralization is associated with peralkalic granite and
syenite at Thor Lake, Northwest Territories, and in peralkalic pegmatites of Khaldan-Buregtey,
Mongolia (Kovalenko et al., 1995).
Ores related to sub alkalic A-types include Be, U, Li, Sn, and Mo. Small Sn deposits are found
in topaz rhyolites in western North America (Huspeni et al., 1984; Eggleston, 1987) and larger
deposits are found associated with A-type granites in Jiangxi province, China (lowP topaz granites
of Zhu, 1990) and Nigeria. Both kinds of A-type granite are found in Nigeria, but the biotite
granites are more mineralized, here with Sn and lesser Nb-U-Zn-W (Bowden and FO.nnaird, 1978;
Imeokparia, 1982). It is important to note that A-type rhyolites are generally no more enriched in
Mo than the oxidized I-types; this contrasts with distinctly higher Be, Sn, Zn, U, etc. in the A-types
(Fig. 12) and is consistent with the control of Mo incompatibility by high f02. Be ores and gem
red-beryl (Keith et al., 1994) are extracted from separate topaz rhyolite dome complexes in western
Utah that are also related to small deposits ofU and Li. All of these ore elements are incompatible
elements that become strongly enriched by fractional crystallization and low degrees of partial
melting. Moreover, this suite of elements has moderately higher concentrations in the continental
crust, which appears to play an important role as source or contaminant in these non-peralkaline
silicic magmas.
The suite of enriched elements in A-type magmas then is a function of source (nonorogenic
mantle for peralkaline varieties versus a similar mantle plus crust for sub alkaline varieties), small
degrees of melting, coupled with extensive fractionation in whichj02 and peralkalinity control that
nature of the fractionating phases. Peralkaline series can become dramatically enriched in Zr, Nb,
U, and Th because minerals rich in these high field strength elements are much more soluble in
peralkaline than sub alkaline magmas. During fractionation of magmas with a continental crustal
component, reduced magma series can become enriched in Sn and oxidized variants in Mo. All will
become enriched in Be, U, Li and other incompatible elements. The apparent lack of continuous
mafic magma recharge into A-type magma chambers may allow fractionation to dominate the
diluting effect of magma mixing that is so common in I-type magma associations. High F
concentrations in peralkaline and sub alkaline varieties also assist in lowering the magma's so1.idus
temperature (Manning, 1981) and viscosity (Baker and Vaillancourt, 1995) to allow for more
effective crystal-melt separation and enhance the enrichment of incompatible elements. Peralkalic
melts have even lower viscosities because of the depo1ymerizing affect of the alkalis (Baker and
Vaillancourt, 1995).
OR-type Silicic Magmas
Rhyolites and granites found at ocean ridges are rare and small in volume, but have distinctive
features that contrast with other types. Consequently they are informative about the sourceS and
processes that are important for silicic magmas in generaL Most of these silicic rocks are
i
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146
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E.H. Christiansen and J.D. Keith
metaluminous to slightly peraluminous tonalites or quartz diorites and include oceanic
plagiogranites, such as those described by Coleman and Peterman (1975). These are sodic granites
that typically lack K-feldspar and are found in ophiolite complexes, some of which formed at
oceanic ridges. Pearce et al. (1984) call these ocean ridge granites, but compositionally similar
rocks are also found in ophiolites that may have originally formed in back arc basins above
subduction zones, such as the Troodos Massif of Cyprus (Aldiss, 1981).
The most characteristic features of OR-type granites are strong depletions oflarge ion lithophile
elements like Ba, Rb and K (Fig. 8). This feature must be inherited from the low concentrations of
these elements in the source of these ocean-ridge granites, presumably by fractionation from MORE
(mid-ocean ridge basalt) derived from the depleted upper mantle or by partial melting of these
basalts above ridge-centred magma chambers. Another important feature is the lack of a negative'
Nb-Ta-Ti anomalies in chondrite-normalized trace element patterns. Consequently, these silicic
magmatic rocks have the lowest RblNb ratios (Fig .. 6) of any considered here and occupy a
distinctive position on the discriminant diagrams (Fig. 7) of Pearce et al. (1984). No data on the
FICI ratios of glassy rocks or melt inclusions from silicic ocean-ridge rocks is available. However,
if they are fractionates of mid-ocean ridge basalt, they probably have FICI ratios near 1 to as high as
10 (Schilling et al., 1980; Byers et al., 1986). Basalts and peralkalic rhyolites related to mantle
plumes have FICI ratios from 1 to 0.1 (Fig. 10). The same kind of comparison suggests that oceanridge granites also crystallize at low j02.
:gxtensive fractionation may lead to enrichments of incompatible elements such as Pb, and
perhaps Zn, in the silicic magmas, but some other ore metals (notably Cu) are more enriched in the
associated mafic lavas. Moreover, Pb and Zn are more likely to be enriched in continental magma
series. The production of magma rich in ore elements that are incompatible at low j02 is difficult
because of the depleted nature of the MORE source materials.·Moreover, any enriched magmatic
fractionate would be very small in volume. The production of ore by secondary processes acting on
such small fractionates would also be difficult.
SUMMARY OF METAL ASSOCIATIONS
Many ore deposits considered here are related to compositional controls, that is to high
concentrations of ore elements in magma or to an intensive property of magma that enhances
transport and eventual concentration (e.g.,j02,jHCI, jH20). We have summarized the principal
associations of ore deposits (Fig. 15) with the different granite types on the trace element
classification diagram of Pearce et al. (1984). The systematic position of ore-related silicic magmas
on this diagram re-iterates the role of source composition, magmatic processes (such as partial
melting, fractional crystaIlizav.0n and contamination), volatile/fugacities and tectonic setting on the
concentration of ore metals irt silicic magmas. Ray et al. (1995) produced a similar diagram for the
association of granite types and skarn-related ores from British Columbia. Fe, Au and Cu deposits
are typically found associated with unfractionated I-type plutonic rocks. All these elements become
depleted during fractionation of silicic magmas. Moreover, increasing continental crustal
contamination dilutes the magmatic concentrations of these elements as well. Consequently, the
highest probability of ore development is with the least fractionated and least contaminated magmas.
More fractionated I-type magmas may produce Cu-Mo and then Mo-Cu deposits. The most
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w. Ta REE
MSn
Zr
Sn
100
o
Ag Zn
U
Nb
Hg
Figure 15 Trace element
compositions of granites and
rhyolites associated with various
types of mineralization.
Modified after Pearce et al.
(1984) and Ray et al. (1995).
Cu
30
Au
10
Fe
I-Type
Granite
OR-Type·
Granite
3
2
5. 10
20
50 100 200
5001000
Nb + Y (ppm)
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chemically evolved rocks of this sort are associated with Climax-type Mo deposits that lack Cu
mineralization. High f02 seems to be the most critical factor that makes Mo behave as a highly
incompatible element (Tacker and Candela, 1987; Keith etal., 1993). Consequently, highly
fractionated I-types with A-type affinities may host these ore deposits. In contrast, fractionation of.
reduced magma series such as those related to S-type granites with incorporated sedimentary
carbon, lead to Sn deposits. Sn is a compatible element in oxidized magmas and probably does not
reach the levels necessary to generate ore in oxidized magmas (Ishihara, 1977; Lehmann, 1~82).
Tungsten's behaviour is interinediate to that ofMo and Sn, and is enriched in fractionated I-types
that have suffered crustal contamination and in some granites with stronger S-type affinities ( Keith
et al., 1989). A-type magmas are most closely associated with incompatible elements like Be, F, U,
Th, Nb, and Zr, and perhaps Sn as well. The low degrees of crustal melting, contributions from
distinctive mantle reservoirs represented by mafic magmas from plumes and continental rifts, and
high degrees of fractionation of either crustal or mantle-derived magma are the primary causes of
this association with incompatible trace elements. Ocean-ridge granites are rarely associated with
mineralization.
The role of magma composition in setting the stage for ore generation must be met with another
set of conditions thai leads to the generation of ore. For example, a Mo-specialized granite may be
emplaced at sufficiently great depths that it reaches volatile saturation very late in its history and is
thus umi.ble to generate a hydrothermal ore deposit. Trace element characteristies may help us
identify the specialization of a granite, but will do little or nothing to help decipher the depth of
emplacement. Other important geologic conditions that would leave no imprint on the magmatic
concentrations are innumerable. Trace element geochemistry is only one of many facets that must be
considered in exploration for rhyolite- and granite-related deposits. No one parameter can be used
to ascertain the compositional character, tectonic setting, or metallogenic affinity of silicic ritagmas.
148
B.H. Christiansen and J.D. Keith
ACKNOWLEDGMENTS
As with any review paper, we have relied heavily on the data and ideas of others and apologize to
those whose contributions were ignored or misinterpreted. We are grateful for the review of M.G.
Best and discussions with many graduate students. We also thank K.Waite for permission to use
her unpublished data.
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