The application of trace elements to the petrogenesis of igneous rocks of granitic composition

26

Earth and Planetary Science Letters, 38 (1978) 26 43 © Elsevier Scientific Publishing Company, Amsterdam - Printed in The Netherlands

THE APPLICATION OF TRACE ELEMENTS TO THE PETROGENESIS OF IGNEOUS ROCKS OF GRANITIC COMPOSITION GILBERT N. HANSON Department of Earth and Space Sciences, State University of New York at Stony Brook, Stony Brook, N. Y. 11794 (USA)

Although trace element modeling has been used to great advantage for petrogenetic interpretations of basaltic systems, similar studies on igneous rocks of granitic composition have been fewer. In general the mineral/melt distribution coefficients for rare earth elements (REE) in granitic melts are equal to or greater than those for similar minerals in the basaltic system. Thus the effects of these minerals on the REE patterns of granitic melts during partial melting or differentiation are exaggerated as compared to basaltic systems, making detection of residual phases easier. For the K/Rb ratio, if neither a K-feldspar component nor biotitephlogopite is present in the residue, it is difficult to reduce the K/Rb ratio of the melt relative to the parent by a factor of two by either differentiation or partial melting. The petrogenesis of four distinctly different rocks are received: (1) an Archean tonalite presumably derived by partial melting of an Archean tholeiite at mantle depths, leaving a garnet plus clinopyroxene residue; (2) an Archean quartz monzonite presumably derived by partial melting of a short-lived graywackeargillite sequence at crustal depths; (3) a dacite from Saipan presumably derived by differentiation from a basaltic parent; and (4) a trachyte from Ross Island, Antarctica, presumably derived by differentiation from a basanitoid parent and contaminated by continental crustal components.

1. Introduction Although trace element modeling has been used with success in petrogenetic studies of basaltic rocks, similar studies on granitic rocks * of igneous origin have been relatively few. A number of studies have shown, however, that trace elements can be used to great advantage for determining the origin of granitic rocks (e.g. [ 1 - 6 ] ) . Whereas rocks of basaltic composition may be derived by partial melting of peridotitic or basaltic parents and may be modified by differentiation and possibly reaction with the crust, rocks of granitic composition may be derived by partial melting of the mantle, basaltic rocks, other granitic rocks, rocks of intermediate composition, and sedimentary rocks, over wide ranges of total pressure, water pressure, and temperature. The resulting granitic melts may have * By "granitic rocks" is meant felsic and sialic melts occurring as intrusions or extrusions whose compositions may be considered in the Q-Plag-Orsystem in which plagioclase has an An content of less than 50% (excluding anorthosites).

been modified by differentiation, mixing with other melts, or reaction with rocks of different composition. Although there are variations in the trace element content for the parents for basaltic rocks, the variation is not as great as it potentially is for the parents of granites. Due to the large possible variations in P, T, PH20 and possible parent compositions, and the sluggishness of reactions in silicic melts, existing experimental studies on granitic systems are not always adequate for petrogenetic interpretations. Also, the available mineral/melt distribution coefficient (Kd) data based on natural systems and experimental studies is limited. Since any petrogenetic study has as its main purpose the elimination of possible models, it means that for the granitic system it may not be as easy to restrict the number of viable petrogenetic models as it is for the basaltic system. This paper presents typical mineral/melt distribution coefficient data for granitic rocks, a review of and model calculations for the use of trace element data in petrogenesis, the effects of residual minerals on K, Rb, Sr, Ba and rare earth elements (REE) in granitic melts and a review of the data for four differ-

27 ent granitic melts and their petrogenetic interpretations. 2. Use of mineral/melt distribution coefficients in petrogenesis The advances in the use of trace elements in the last ten years has been mainly in shifting from the use of trace elements as "fingerprints" allowing comparison of rock types to the use of the trace elements to decipher the origin and evolution of igneous rocks. These advances have been made through the use of Kd's for trace elements in quantitative models for differentiation and partial melting. (Table l gives the symbols used in this paper.) A mineral/melt Kd is the measured weight fraction of a given trace element in a mineral divided by the measured weight fraction of that element in a coexisting melt. The melt and mineral are assumed to be in equilibrium or quasi-equilibrium. An example would be 600 ppm Ba in K-feldspar phenocrysts and 100 ppm in the glassy matrix giving a K-feldspar/ silicate melt Kd of 6 for Ba. Under equilibrium conditions the mineral/melt Kd is independent of relative proportions of mineral and melt and concentration of an element. Whether an element is treated as a trace element or not depends on whether the element follows Henry's law in the phases considered. An element may be present in trace quantities but actually be an essential structural constituent of a mineral phase. An example would be Zr in a granitic melt where it may have a concentration of only hundreds of parts per million yet zircon is present in which Zr is not a trace element but an essential structural constituent. The Kd's for a given element are dependent on the temperature and composition of the mineral and melt. For example, the mineral/melt Kd's for REE are generally higher in granitic melts than in basaltic melts. It is not always clear whether the variations are more dependent on temperature or composition as the composition of both liquid and minerals are significantly different from the granitic melts as compared to basaltic. The temperature dependence of a Kd for a given element is [7] : in Kd

= C 1 T -I + C 2

where C1 and C2 are constants.

(1)

Theoretical derivations of quantitative modeling of trace elements for differentiation and partial melting of igneous rocks have been presented by Neuman et al. [8], Schilling and Winchester [9], Gast I10], Shaw [ 11 ], Greenland [12], Albar~ de and Bottinga [ 13] and Hertogen and Gijbels [14]. Two models are commonly considered for partial melting-batch melting and fractional melting. In either model the melting is assumed to occur under equilibrium conditions; that is, melting occurs over a long enough period of time so that diffusion of'a given trace element in the melt and liquid allow proper distribution of the trace elements in the melt and residual mineral phases. Since diffusion coefficients are much higher in the melts than in residual minerals [15], the rate-controlling factor determining whether equilibrium is attained is the diffusion rate in the minerals. For example, equilibration of K and Rb in orthoclase [16] or Sr in phlogopite [15] for grain sizes of 1 cm take place on the order of a thousand years at 1000°C but on the order of a million years at 700°C. Since the tinre involved for equilibration is proportional to the square of the radius of a grain, parents with a smaller grain size equilibrate much more rapidly. For example, if the grain size is reduced by one-fourth, the time needed for equilibration is reduced by one-sixteenth. There is the possibility that for coarse-grained rocks at the lower temperatures of melting equilibrium may not be attained between the centers of grains and the melts. For those elements with low Kd's the effect on the melt will be to make it appear that the parent has a lower concentration of those elements than it actually has or that the extent of partial melting is greater than it actually is. Also, there may be some effects on the trace element ratios in the melts due to the different diffusivities of the varous trace elements in the minerals. For those trace elements with large Kd's the effect on the melt will be to make it appear that the residue has a lower proportion of the minerals with the high Kd's than it actually has. In batch melting the trace elements in the melt and solid residue are in equilibrium as melting proceeds until enough of the melt accumulates so that the lower density melt rises removing itself from the residual phases and forming a pluton, a hypabyssal body or volcanics. In fractional melting, as melting takes place, small fractions are continuously removed

28 and stored in a magma chamber. Batch melting is probably the more realistic model because the first melt that forms will contain essentially all of the volatiles. When the melt leaves the residue the volatiles accompany the melt leaving a dry residue. The solidus of this residue will be significantly greater than that o f the original parent and melting will probably not proceed without the addition o f more heat or volatiles. A common misconception is that the trace element composition of a rock is controlled by the minerals present, for example, a plagioclase-rich rock has a high K/Rb ratio, positive Eu anomaly, etc. Except for cumulates, this is not necessarily the case. The nfineral composition of a rock is determined by the major element composition and the P, T, tgH20, etc., conditions. Assuming a closed system prior to melting, the trace elements of the original rock are distributed among the minerals according to the appropriate Kd's. As melting proceeds minor as well as major mineral phases react or melt and disappear. If there is equilibrium, as a phase disappears the concentrations of the trace elements in the residual minerals and melt continue to readjust. Shaw [1 1] developed a general equation for partial melting during batch melting, the only case considered here, in which the weight concentration of a given trace element in a melt, CL relative to the con-

TABLE l

Symbols used in this paper (the symbols are similar to those used in Shaw [ l 1 ] and l lcrtogen and Gijbels [ 14 ]) X i

=

Kd i =

D

=

weight fraction of phase i in the solid mineral/lnelt weight distribution coefficient of a trace element ~k~r phase i bulk distribution coefficient of a given trace element

for the residual mineral phases at time of separation of melt and residue; D = ~znXiKd i D O - bulk distribution coefficient of a given trace element at the onset of melting pi _ fractional contribution of phase i to tile liquid p

=

~lpiKd i

weight concentration of a trace element in parent C L = weight concentration of a trace clement in a derived melt Cs = weight concentration of a trace element in the residual mineral phases CO

=

D

= Cs/C L = weight

F

fraction of melt relative to original parent

centration of the element in the parent that is melting, Co, is given by:

CL/Co = 1/[Do + F ( I - P)]

(2)

where

Do = ~TX;Kct;

(3)

and

p = ~TpiKcl i

(4)

Do is the bulk distribution coefficient of the parent for a given trace element based on the mineral composition at the beginning of melting. X i is the weight fraction of a given mineral i in the original parent. Kd i is the mineral/melt distribution coefficient for a given trace element for mineral i. P is the bulk distribution coefficient o f the minerals that are making up the melt in which pi is the weight normative fraction o f mineral i in the melt. The symbols used in this paper are summarized in Table 1. Rearrangement of equation (2) gives:

CL/Co : 1/[(Do - PF) + F]

(5)

in which it can be seen that as melting proceeds the factor (Do PF) takes into account the extent of melting on the change of the bulk distribution coefficient of the residual phases. It does not, however, take into account the fact that some phases may have reacted or melted completely. So it only applies while the main mineral phases involved in the residue and melting are still present in the residue. The relationship in equation (5) may leave the impression that the composition o f a melt, CL, for a given element is dependent on the original mineralogy and the normative composition o f the melt. As will be shown this is not true. What the equation does show is that if the original mineralogy and the normative composition of the melt is known, it is possible to calculate CIJCo as long as the original phases remain. For the simple case of modal melting where the melt and the parent have the same normative composition, Shaw [11 ] gives the relation:

CL/C o : 1/[Do + F(1

Do)l

(6)

where Do substitutes for P. Rewriting equation (6) in the same form as equa-

29

IO0

tion (5) gives:

eL/Co = 1/[Do(1 - F ) + F ] .

(7)

I

t

Hertogen and Gijbels [14] have presented derivations to take into account the restriction for equation (2) that no phase may be used up, and to allow the Kd's and melting proportions to vary during melting. However, to use these equations it is necessary to know the correct Kd's as a function of temperature and mineral and melt composition and to have a very good understanding of the phase equilibria during the melting. Since this information for most rock systems is poorly known, it would be more appropriate to be able to consider only the possible parent trace element concentration, the mineral content of the residue at the time of removal o f the liquid from the residue, and the extent of partial melting. D, the bulk distribution coefficient o f the residue at the time of removal of liquid [11, eq. 12], is given by:

D

=

(D O

-

PF)I(1

-

F)

c5

0.1

(8) 001

(9)

Substituting in equation (4) gives the same relation as that of Schilling and Winchester [9, eq. 4] :

eL/C 0

=

1/[D(1 -

F) + F] .

C,/Co = D(I-F)+F D=O

I0

or:

D(1 - F) = D O - PF.

PARTIAL MELTING

(10)

This formulation shows that at the time of removal of a liquid from a residue the concentration of a given element in the liquid relative to the parent CL/C o is dependent on]y on the bulk distribution coefficient of the residue, D, and the extent of partial melting, F . Thus, if a mineral was originally present during the partial melting sequence but has since melted or reacted out, it has no influence on the trace element con]position of the melt. This, of course, assumes an equilibrium or a quasi-equilibrium situation for the given trace element between the minerals and melt. The concentration of a trace element in the residue, Cs, relative to the parent, Co, can be calculated from the relationship Cs/Co = eL/Co" D which is from the definition of a distribution coefficient, i.e. D = Cs/C L. eL/Co vs. F for partial melting is plotted in Fig. 1. I f D is larger than F, eL~Co approaches lID for F less than about 40% and approaches 1./D more closely the smaller the value o f F . As D approaches 0, eL/Co

-

0

-

O5 F

10

Fig. 1. The concentration of a trace element in a melt relative to its concentration in the parent rock, eL/Co, versus the weight fraction of partial melt F [ 10,1 1]. D is the bulk mineral/melt distribution coefficient tbr the minerals in the residue at the time of removal of the melt. It is assumed that the melt is in equilibrium with the residual mineral phases until separation of the melt and residue.

approaches 1/F and approaches it more closely as F increases. During differentiation * the Rayleigh fractionation law [17] as applied by Neuman et al. [8] can be used to describe the trace element concentration of the differentiated melt, e L , relative to the parent melt, Co:

eL/C 0 :

F (o-l)

(1 1)

F is the fraction of melt left and D is the bulk * Differentiation is the evolution of a melt derived by fractional crystallization. Fractional crystallization as used in the Rayleigh fractionation law is the essentially instantaneous equilibrium precipitation of an infinitesimally small amount of crystals which inamediately settle out of the melt and are immediately covered by another layer of crystals which prevent the earlier formed crystals from equilibrating with the evolving melt.

30 distribution coefficient for the crystals settling out of the melt. Greenland [12] presents equations which take into account variations in the proportions of minerals or variations in the Kd's as a function of composition or temperature. It is also possible to use equation (11) incrementally to take into account the variations in mineral proportions and Kd's. For differentiation as D approaches zero, CtjCo approaches 1IF, similar to the case for partial melting (Fig 2). Unlike partial melting, however, i f D for a given element is larger than 1, fractional crystallization rapidly depletes that trace element from the melt. For example, i f D = l 0, 7.5% fractional crystallization reduces that element's concentration in the melt by a factor of two. During partial melting under equilibrium conditions, if a mineral has melted or reacted out prior to removal of the melt, the former presence of that mineral has no influence on the trace element com-

IO0

FRACTIONAL CRYSTALLIZATION-

position of the melt; that is, the trace elements are distributed between the melt and the remaining residual minerals according to the appropriate Kd's. During fractional crystallization, however, if a mineral had precipitated out earlier and were no longer crystallizing because of changing T, P, PH20 or com" position of the melt, it has left an effect on those trace elements in the melt for which the mineral/melt Kd is significantly greater than 0. Thus, for the differentiation model it is necessary to account for all of the minerals which may have precipitated and to consider variations in the Kd's as a function of temperature and composition. In natural magmas, the distribution of a trace element between a phenocryst and melt is complicated by zoning which is a result of the rapid growth of crystals and the relatively slow rate of diffusion of the trace elements in the crystal and melt [13]. Hart and Brooks [ 18] have shown that in a single phenocryst crystal the clear and cloudy portion have quite distinct trace element concentrations. They suggest that the best Kd values to use for differentiation studies are those for zoned phenocrysts as they seem

C,/Co=F(Dq)

I00

3=0

PARTIAL I " 10% MELTING / 80 E

ge~

,\

/ -'°"

ii a D~I

(5

40 U

30%

2

~

- 40 %

20

0%

0.1

q r

001

0

O5

Io

F

I:ig. 2. The concentration of a trace clement in a melt relative to its concentration in the parent melt, CL/Co, versus the weight fraction of melt remaining during fractional crystallization, F, for a given bulk distribution coefficient D. This model is based on the Rayleigh fractionation law.

£

go

~o

do

,3o

R ( D = 4 ) ppm Fig. 3. A plot of trace element concentrations in two suites of igneous rocks. One suite is derived by 1 0 - 4 0 % partial melting of a h o m o g e n e o u s parent, while the other suite is derived by 40% partial melting o f the same h o m o g e n e o u s parent followed by 0 - 5 0 % fractional crystallization ( F = 1 to 0.5 respectively). Since D for trace element Ce is 0 in both suites, Ce follows a 1/F relation for both partial melting and differentiation. D for trace element R is 4 in both suites. Ce = 25 p p m and R = 100 p p m for a melt derived by 40% melting of the h o m o g e n e o u s parent.

31 to represent the rule. In a study of differentiation in a sequence of volcanics the ideal solution would be to determine and use the Kd's based on the phenocrysts in those volcanics. For either differentiation or partial melting it is possible to determine the extent of fractional crystallization or partial melting by selecting a trace element which is suspected of having a very low D. For D approaching zero, CL/Co = l/F, which gives the minimum allowable extent of partial melting or minimum fraction of melt remaining due to fractional crystallization. It is also possible to distinguish whether a suite of comagmatic rocks is a result of partial melting or of fractional crystallization by considering those trace elements with D's significantly greater than 1 (see Figs. 1,2 and 3). If the concentration of the element remains relatively constant throughout the suite, the main process is partial melting. If the element with a large Kd has a large variation, the main process is differentiation.

3. Residual minerals in granitic melts The possible minerals involved as residual phases during partial melting or differentiation to form a granitic melt are plagioclase, K-feldspar, quartz, hornblende and other amphiboles, clinopyroxene, hypersthene, olivine, biotite or phlogopite, muscovite, cordierite, kyanite or sillimanite, scapolite oxides, sulfides, apatite, zircon and other accessory minerals. Kd's are not available for K, Rb, Sr, Ba and REE for quartz, muscovite, scapolite kyanite, sillimanite, cordierite, oxides or sulfides. Except for cordierite, scapolite and muscovite it is probably safe to assume, on the basis of crystal chemistry, that the Kd's for K, Rb, Ba, Sr and REE are close to zero for these minerals. Olivine also has very low Kd's for these elements and is probably important only for the derivation of a granitic melt by differentiation from a basalt. There is a controversy as to whether a granitic melt in the sense used here, can be derived directly by partial melting of peridotite or the mantle under high PH2 o conditions [ 1 9 - 2 1 ] . The feldspars potentially may be quite complex due to solid solution between plagioclase and Kfeldspar (see, for example, chapter 5 of Carmichael et al. [22]). At high water pressures where PH20 =/°load

or for relatively high An contents for the plagioclase, the minimum in the feldspar system intersects the solvus and there may be only limited solid solution of K in plagioclase. For low water pressures and plagioclase with low An contents the solidus may not intersect the solvus and there is complete solid solution between K-feldspar and plagioclase. For low PH20 conditions and parents with high plagioclase contents, because of the extensive solid solution, the first melts formed during partial melting may be significantly more plagioclase-rich than the minimum belt composition. Although K is a major element, in many cases it can be treated as a trace element in the melt, if there is no residual phase, such as biotite or K-feldspar in which K is an essential structural constituent. If biotite or K-feldspar is present with minimum solid solution of other cations for K, then the concentration of K in the melt will be constant, independent of extent of melting, as long as any of that phase and the other phases remain in the residue and T, P, and PH20 are not changing significantly. This can be understood from Fig. 4, a simple albite-orthoclasequartz ternary eutectic system with no solid solution. If the parent is in the plagioclase field, as melting begins, the composition of the first melt is at the eutectic composition. As melting proceeds the composition of the melt remains at the eutectic until all

Q

Ab

/

, Or

Fig. 4. A Q-Ab-Or ternary system with no solid solution. The lines drawn separating the mineral fields are coteetics and the intersections are eutectics. The proposed parent for partial melting is shown in the Ab field.

32 of the orthoclase in the residue is used up. Thus, as long as albite, orthoclase and quartz are present, the K content remains constant. As soon as the orthoclase is used up the melt moves along the quartz-albite cotectic and the K content of the melt is reduced by dilution with the continued addition of K-poor melt. In this section the individual minerals are considered with respect to how they as residual phases will qualitatively affect the concentrations of K, Rb, St, Ba, and REE in granitic melts derived by either partial melting or differentiation. The Kd data on which the effect of the minerals on melts is discussed are presented in Table 2 and Fig. 5. It should be realized in the following discussion of the individual minerals that the magnitude of a mineral's effect is directly related to its relative abundance as well as the magnitude of the Kd for a given element.

Plagioclase. Plagioclase generally depletes the melt in Sr. Typical Kd's for Sr for rhyolites are on the order of 4, but Z.E. Peterman (personal communication, 1974) has data suggesting values for Kd's for granitic melts as high as 30. The large positive Eu anomaly in the Kd pattern contributes to a negative Eu anomaly in the melt. The extent of the Eu anomaly in plagioclase increases with decreasingfoz and with decreasing T [23]. Although plagioclase has proportionally much higher Kd's for K than for Rb the absolute values of the Kd's for K and Rb are so low that plagioclase in the residue does not greatly affect the K/Rb ratio. Thus, except for Sr and Eu the other trace elements considered here are but little affected by plagioclase. However, if there is significant solid

solution of K in plagioclase, the Kd's for Ba and Rb are increased (see the anorthoclase in Table 2).

Potassium-feldspar. Potassium-feldspar is similar to plagioclase with regard to its high Kd's for Sr and Eu, and low Kd's for the REE other than Eu. It has, however, a high Kd for Ba. Potassium-feldspar can be distinguished from plagioclase, because it does not increase the Rb/Sr ratio in the melt as much as plagioclase does, but leads to a much greater Sr/Ba ratio in the melt. The K-feldspar, also contributes greatly to a reduction of the K/Rb ratio of the melt relative to the parent (see Fig. 6A). Anorthoclase. As might be expected anorthoclase, a solid solution of plagioclase and K-feldspar, has an effect comparable to plagioclase and K-feldspar. Garnet. Garnet has very small Kd's for K, Rb, Sr, Ba and the light REE and large Kd's for the heavy REE. When present in significant quantities it leads to melts with relatively depleted heavy REE and contributes to a positive Eu anomaly. Hypersthene. Hypersthene has relatively low Kd's for the trace elements considered here and does not greatly affect the ratios of these elements. It leads to melts with slightly more enrichment of the light REE than the heavy REE and contributes to a positive Eu anomaly. Clinopyroxene. Clinopyroxene in the residue leads to relative depletion of the middle and only somewhat

TABLE 2 Typical mineral/melt Kd's for common rock-forming minerals from dacitic and rhyolitic rocks * Element

K Rb Sr Ba

Garnet

Clinopyroxene (2)

Hornblende

Biotite

(1)

Hypersthene (1)

(2)

K-feldspar Anorthoclase (2) (3)

(1)

(1)

(1)

(1)

(2)

0.020 0.0085 0.015 0.017

0.0023 0.0027 0.0085 0.0029

0.037 0.032 0.516 0.131

0.081 0.014 0.22 0.044

0.065 0.0077 0.094 0.054

(5.63) ** 3.26 0:120 6.36

(1.49) ** 0.659 3.87 6.12

0.10 0.041 4.4 0.31

0.076 0.016 1.45 0.30

0.263 0.048 2.84 0.36

0.45 5.57 5.04

* From (1) Nagasawa and Schnetzler [391 ; (2) Philpotts and Schnetzler [25]; (3) Sun and Hanson [34]. ** These values are not actually Kd's as K is an essential structural constituent in biotite and K-feldspar.

Plagioclase

33 40C Zircon

....

o Garnet Apotite

///

/ Hornbmnde

less depletion of the heavy REE and contributes to a positive Eu anomaly.

Hornblende. Hornblende in the residue has little effect on K, Rb, Sr, or Ba. Although it could lead to some reduction in the K/Rb ratio the Kd's for K and Rb are relatively small. Like clinopyroxene the main effect of hornblende on the REE is in depleting the middle and, less so, the heavy REE, and contributing to a positive Eu anomaly. The Kd's for hornblende, however, are significantly greater than those for clinopyroxene.

I

///0\\\\ i/

Clinopyroxene

////

////0 ~

Oj ~

//t~0 Hypersthene Biotite

~~0~ ~

Biotite. Biotite, like hypersthene, has little effect on the REE distribution due to its relatively low Kd's for the REE. Its main effect will be on maintaining a constant K content in a melt under similar conditions, reducing the Rb/Sr ratio and increasing the Sr/ Ba ratio in the melt (see Fig. 6B).

~

Nid

SJmE'u Gd

D,

E'r

Zircon and apatite. Both zircon and apatite have large Kd's for the REE. If they occur in sufficient quantity, they may have a relatively large effect on REE patterns. The presence of zircon leads to a depletion of the heavy REE. The presence of apatite leads to relative depletion of the middle REE and contributes to a positive Eu anomaly.

Yb

3.1. Summary

~

Plogiock)se

0"0: o

oo." c',

.~

s" E'o ~

~

i E~

K - FeldSpar

;~

Fig. 5. Mineral/melt distribution coefficients for REE from dacites and rhyolites. These are the averagevalues from Arth and Hanson [4], which were selected from Nagasawa and Schnetzler [39], Higuchi and Nagasawa [40], Nagasawa [41], and Schnetzler and Philpotts [42]. The data for the anorthoclase are from Sun and Hanson [34].

(1) Potassium, Rb and Ba are retained by biotite and K and Ba by K-feldspar. (2) Strontium and Eu are retained by the feldspars. (3) The Rb/Sr ratio of the melt is increased in the melt by the presence of plagioclase, less so by Kfeldspar, and decreased by the presence of biotite in the residue. (4) The K/Ba and K/Rb ratios are not significantly affected by any of the minerals except K-feldspar or biotite. (5) The Sr/Ba ratio in the melt is reduced by plagioclase, increased slightly by K-feldspar, more so by biotite. (6) The REE have their greatest utility in that, although the values for the Kd's for a given mineral may vary widely as a function of temperature or composition, in general the shape of the pattern for a given mineral is consistent. Thus, a given mineral

34 will have a characteristic effect on a REE pattern of a melt which allows identification of that mineral as having been a residual phase either during partial melting or fractional crystallization, although it may not be possible to calculate the fraction of that mineral in the residue. Relative to the other REE: (1) The heavy REE are retained somewhat by clinopyroxene but to a greater extent by hornblende, garnet and zircon. (2) The middle REE are retained by clinopyroxene, hornblende and apatite. The light REE are not retained to any great extent by the common minerals except for apatite. (3) A positive Eu anomaly in the melt results from garnet, apatite, hornblende, clinopyroxene and hypersthene in the residue. (4) A negative Eu anomaly in the melt results from plagioclase or K-feldspar in the residue. (5) Equal proportions of plagioclase and clinopyroxene or twice as much plagioclase as hornblende in the residue produce melts with negligible Eu anomalies.

4. Trace element modeling In this section tire effects of various residual minerals o~1 the trace element contents of melts are deduced from calculated models. A parameter in which geochemists have been particularly interested is the K/Rb ratio (e.g. [24]). In general more basic rocks have higher K/Rb ratios and the more granitic rocks have lower K/Rb ratios. Let us consider how it might be possible by partial melting or differentiation of more basic rocks with low K to form granitic rocks with lower K/Rb ratios than the parent. Under partial melting conditions the K/Rb ratio in the melt relative to the K/Rb ratio in the parent is given by modifying equation (10) as: C L K/Rb DRb(I - F ) + F C O K/Rb DK(I - F ) + F

(12)

It can be seen in equation (12) that the K/Rb ratio of the melt relative to the parent is a function of D and F. IfD is smaller than F for both K and Rb, then the K/Rb ratio of the melt approaches that of the parent. The DK/DRb ratio is only important for affecting the

K/Rb ratio of the melt if the fraction of melting, F is low; that is, F has a much lower value than the numerical values for D. At F = 0, CL/Co = DK/DRb. For example, reducing the K/Rb ratio by a factor of two or more using the Kd's for granitic melts in Table 2, with only hornblende in the residue ( O R b -0.014,DK = 0.081) requires 5% or less melting. For differentiation, the K/Rb ratio in the melt relative to the parent melt is given by modifying equation (11) as"

C L K/Rb _ F(DK ORb) Co K/Rb

(13)

Again it can be seen that the K/Rb ratio in the melt relative to the parent is a function of both the fraction of melt remaining and the distribution coefficients. Reducing the K/Rb ratio in the melt by a factor of two or more with respect to the parent K/Rb ratio by fractional crystallization, using the same hornblende as in the partial melting example, requires that b must be 0.00003 or less. Thus, differentiation is much less effective at changing the K/Rb ratio in the melt than is a partial melting process for low D's. The concentrations of trace elements with D's significantly less than F in a suite of samples depend on the extents of partial melting or differentiation, but the ratios of these elements in the melt will reflect the ratio of these elements in the parent. If hornblende from a lanrprophyric rock with much higher Kd's for Rb and K were used where D K = 1.40 and DRb = 0.427 [25], a reduction of a factor of two or more in the K/Rb ratio in the melt relative to the parent would require F's of 35% and 49% or less for partial melting and differentiation, respectively. For the higher D values the differentiation process is more effective in changing the K/Rb ratio of melts. In order to affect the K/Rb ratio in a granitic melt it is necessary to consider other phases. Although plagioclase has high KdK/KdRb ratios, it's Kd's are even lower than those of hornblende (see Table 2). A relatively pure K-feldspar phase in the residue, however, would lead to a dramatic reduction of the K/Rb ratios as the Kd's for Rb are quite small. Depending on the extent of solid solution among K, Na and Ca in the K-feldspar, the feldspar could have a K content of about 13% or less. The weight ratio of

35 the K would about phase

POTASSIUM FELDSPAR

K/Rb

~ ..........

//

/

///

//

// / /

I0( -

~

Sr

Bo + I0

\\

Rb

\\\ \\

c o n t e n t o f K-feldspar relative t o t h e m e l t b e 3.3 or less, c o m p a r e d to the Kd for Rb o f 0.4. T h u s , K-feldspar o c c u r r i n g as a residual m a y have a significant effect o n r e d u c i n g t h e

K / R b ratios. Biotite as a residual p h a s e c o u l d lead to e i t h e r a h i g h e r or l o w e r K / R b ratio d e p e n d i n g o n t h e stability o f b i o t i t e . I f t h e f o 2 , f H 2 0 , fa, fF, etc., lead to stability o f b i o t i t e a n d b i o t i t e m e l t s c o n g r u e n t l y , i.e., b i o t i t e does n o t b r e a k d o w n t o K-feldspar plus ferrom a g n e s i u m phases d u r i n g m e l t i n g , t h e K c o n t e n t o f t h e m e l t c o u l d b e m u c h l o w e r t h a n t h e m i n i m u m in t h e Plag-Or-Q system. T h e K c o n t e n t o f b i o t i t e is t y p i c a l l y 6 . 5 - 8 . 0 % . Since t h e e u t e c t i c m e l t c o m p o s i t i o n h a s a K c o n t e n t o f a b o u t 4% at 4 kbars, this m e l t to b e d i s t i n g u i s h a b l e w o u l d have a K c o n t e n t o f less

\x x

K+IO00 ~"

Rb/Sr x I00

A

F

BIOTITE

IOOO

500

...........

K/Rb

J

Sr

Fig. 6. Plots of K, Rb, Sr and Ba concentrations (ppm) of melts and K/Rb and Rb/Sr ratios versus the fraction of partial melting, F, using the Kd's in Table 2. These two plots compare elements in melts derived from chemically identical parents. The mineral composition of the parent in A is 20% quartz, 10% K-feldspar, 0% biotite, 50% plagioclase, and 20% hornblende, and in B is 20% quartz, 0% orthoclase, 20% biotite, 50% plagioclase, and 10% hornblende. The concentrations are calculated assuming that biotite breaks down to an insignificant fraction of K-feldspar upon melting and that the K-feldspar and plagioclase are separate phases with minimum solid solution. During the first 30% of partial melting while a K phase is present the K content of the melt is buffered at the minimum melt composition, 4.2% K. The original composition of the parent for both graphs is: K = 15,000 ppm Rb = 30 ppm Sr = 500 ppm Ba = 600 ppm

K/Rb = 500 Rb/Sr = 0.06

while a K phase _is present the phases h~v.e the following K contents:'. '~ .... " ~ ' ~ " "" " ' • IO(3

Ba + ~0

Rb

50

../'~x x //////

~

K+ I000

~///

xxx~

Rb/Sr x I00

B

,6

~

~ F

,~

' 5O

quartz = 0% K biotite = 6.3% K K-feldspar = 12.2% K plagioclase = 0.42% K hornblende = 0.34% K The minimum melt consists normatively of 33% quartz, 33% plagioclase and 33% K-feldspar. Thus at greater than 30% partial melting for both cases the K phase is no longer present in the residue and the two cases are very similar at the higher percents of melting. For less than 30% partial melting the effects on the Ba, and Rb concentrations in the melt are quite different depending on whether K-feldspar or biotite is in the residue.

36 than about 3.5%. With a Kd for Rb of about 2 or 3 the K content of the biotite relative to that of the melt could become greater than the Kd for Rb. Biotite could then also contribute to a decrease in the K/Rb ratio in the melt. |f, however, biotite were not a stable phase during melting and melts incongruently, i.e., as melting proceeded biotite breaks down to Kfeldspar plus ferromagnesium phases, as has been suggested by Winkler [26, pp. 211-214] the melt would have a K content similar to that of the minimum melt composition in the Plag-Or-Q system. The biotite in the residue would then contribute to increasing the K/Rb ratio. In Fig. 6 are summarized the effects on K, Rb, Sr and Ba concentrations in melts of a eutectic composition derived by partial melting of similar parents one in which the K resides in orthoclase the other in biotite. The curves for K and Sr are similar in both cases. The curves for Rb and Ba are, however, significantly different in the two cases leading to significantly different trends for K/Rb and Rb/Sr at lower fractions of melting.

5. Late-stage residual fluids At some point during the cooling and crystallization of a granitic melt a volatile-rich phase may form. Such phases, whether silicate or aqueous liquids or supercritical fluids may be responsible for the formation of aplites or pegmatites found associated with granitic plutons. Also, the volatile-rich phases may seriously affect some trace element compositions of the body by removing or mobilizing the elements, and they may promote subsolidus growth of crystal phases, e.g. K-feldspar megacrysts. The importance of the late-stage fluids on granitic rocks has been emphasized, for example, by Taylor [27] and Buma et al. [1]. Hart et al. [43] have shown the importance of the effects of late-stage fluids enriched in K, Rb, Ba, and Cs on the concentration of these elements at different positions within a basaltic flow. In their study the central portion of a flow is relatively depleted in K, Rb, Ba, and Cs, whereas the upper portion is enriched in these elements. The depleted portion has much higher K/Rb and K/Cs ratios than the enriched portion. They conclude that a late-stage fluid enriched in K, Rb, Ba and Cs has been removed

from the central portion and migrated upward in the flow. The trace element ratios are affected probably because the late-stage fluids tend to prefer the larger ions. The late-stage fluids may have quite different properties depending on whether the body is a flow or an intrusion, and will probably be related to the depth of intrusion as well as the composition and relative amount of the fluid phase. Harris et al. [29] suggest that the depths of intrusion may be a function of the extent of the aqueous fluid dissolved in the magma. A water-saturated magma derived, for example, from wet sediments at 4 kbars with a minimum melt composition would intrude upward, cooling between the adiabatic decompression and the geothermal gradient curves, and would intersect the water-saturated solidus crystallizing at a depth equivalent to 2 kbars or more. A dry melt derived by partial melting of a relatively dry lower crust would, however intersect its solidus at a depth equivalent to much less than 2 kbars. If the melt is very dry, it could reach the surface before intersecting its solidus. Thus, a deepseated body might be expected to have a higher fraction of late-stage liquids than would a shallow intrusion or volcanic. At the temperatures of crystallization pure water is a supercritical fluid whose ability to dissolve is mainly a function of density [30, chapter 15]. At very low pressures the density is low and the supercritical fluid is a poor solvent. At higher pressures the supercritical fluid will have densities comparable to those of liquid water and its properties as a solvent will be similar to those of the liquid phase. The supercritical fluid apparently has no special properties as a solvent. Thus the depth of intrusion as it controls the density of a supercritical fluid, will also affect the characteristics of the fluid phase. Whether a supercritical fluid actually separates from a silicate melt depends on its solubility in the melt. The H20, CO2, C1, F concentrations as well as pH, and fo2 are probably also important in determining the solubility of cations in the fluids and their mineral/fluid Kd's. If the mineral/fluid Kd's are very large, very little of the trace element will go into solution and if the fluid phase forms only a very small fraction of the total body, its effects will be minor. If, however, the cations are very soluble, the mineral/fluid Kd's are relatively small, and the fluid phase forms a signifi-

37 cant fraction of the body its effects may be very important. This may be particularly true for K, Rb, Cs, and Ba in wet magmas. It is not clear how important complexing of REE is in a late-stage supercritical fluid. Mineyev [31 ] suggested that, although the REE as a group have very similar ionic species, at high pressures they form complexes of the form Na(REE)F4 with the stability of the heavy-REE complexes being far greater than those of the light-REE complexes. If this were true, fluoride-rich fluids would be expected to have higher concentrations of the heavy REE than of the light REE and depending on the magnitude of the Kd's this could lead to depletion of the heavy REE in the rock. Cullers et al. [32,33] in experimental studies determined that relative to supercritical aqueous solutions, the REE are strongly partitioned into silicate phases with Kd's on the order of 10 or greater for diopside and plagioclase and on the order of 100's for silicate glasses. They also found that Gd in an aqueous solution of 0.1 m NaF had the same mineralfluid Kd's as it had in a fluoride-poor aqueous solution. This would suggest that complexing may not be important for the REE, although the data are limited.

6. Examples of petrogenetic interpretations for granitic rocks The purpose of a petrogenetic study of an igneous rock or suite of igneous rocks is to determine the chemical and mineralogic composition of the parents at the time of melting, the history of the parents prior to melting, the extent of partial melting, the temperature and pressure conditions during partial melting, and modifications of the primary melt composition due to differentiation, and reaction of resulting melts or rocks due to mixing of melts; assimilation, metasomatism, zone refining or latestage residual fluids. Geologic relations of the rocks studied are necessary in order to determine what other rocks they are associated with, their geologic history, and whether they are extrusions or intrusions. If they are intrusions, it is important to estimate the depths and modes of emplacement. Field observations, careful selection of samples, and petrography are essential for estimating and possibly reducing the effects of later metasomatism, metamorphism, low-

grade alteration and weathering. Petrography is essential for determining the phases present during crystallization. Major element analysis along with petrography are essential for classification of the rocks and for allowing comparison with experimental studies. Geochronology studies are needed in collaboration with the field studies to determine whether assumed cogenetic samples are actually of the same age and to place time constraints on the history of the samples. Initial isotope ratios for Pb, Sr and Nd will allow an estimate to be made of the U/Pb, Rb/Sr and Sm/Nd ratios of the parent and will place constraints upon the length of time it had the given ratio. The stable isotopes of H, C, O, and S are particularly useful for determining the extent of hydrothermal alteration as well as for determining the character of the parent and the processes leading to the presently observed rock. Trace elements will allow an estimate of the trace element composition of the parent melted, the minerals in the residue at the time of removal of the melt, the sequence of minerals involved in differentiation, and along with the initial isotope ratios will allow an estimate of the extent of mixing or reactions with other melts or rocks. For the maximum information to be obtained from a petrogenetic study it is very important that trace element studies either be coordinated with ongoing projects on a sequence of rocks or are undertaken in areas which have been subject to intensive study. It is also very important in obtaining the maximum information to have data of the highest precision and accuracy attainable. There is always a chance that rock types in a sequence may superficially appear to have a common origin, whereas they may actually have quite unique histories. Therefore, it is always safest to assume that the individual rock units in a series may have unique histories until the evidence overwhelmingly supports a common history. Thus, it is very important in petrogenetic studies to do the analyses on single well-selected samples and not composite samples. An important consideration in determining the origin of an igneous rock is that the parent may be heterogeneous. For example, a typical homogeneous granitic intrusion within a batholith may have dimensions of about 5 km on a side or a volume of 125 km 3. If the pluton is derived by 20% partial melting of the crust this would imply that the parent

38 had a volume of 625 k m ~ which could be represented b y a block 8.5 k m on a side. Due to the large dimensions involved it is likely that the parent may consist o f a number o f rock types. The melt may, however, be derived from similar, least refractory portions of a heterogeneous parent. As a review o f trace element modeling applied to granitic igneous rocks the petrogenesis o f a tonalite, a dacite, a quartz monzonite and a trachyte are considered. Major element, Rb, St, and Ba data are shown in Table 3, the normative analyses on a Q-OrHag diagram in Fig. 7, and REE in Fig. 8. These rocks have been chosen because they each have quite distinctly different compositions and petrogenesis. This should not necessarily imply that rocks which are superficially similar to these rocks have a similar origin as well. The Archean tonalite (.4) was presumably derived

by partial melting o f a tholeiitic parent, most likely at mantle depths leaving an eclogite (clinopyroxene plus garnet) residue [2,3]. The two-mica quartz monzonite (B) was presumably derived by partial melting of a graywacke at crustal levels, leaving a residue o f quartz + garnet + plagioclase + biotite + amphibole -+ pyroxene -+ orthoclase [4]. The dacite from Saipan (C), from an island arc regime, is presumably a result o f differentiation o f plagioclase and clinopyroxene from a basic parent [5]. The trachyte from Ross Island, Antarctica (D), is presumably the result of differentiation o f olivine, spinel, clinopyroxene, amphibole, plagioclase, apatite, and anorthoclase from a light-REE-enriched alkali basalt parent and reaction o f the differentiated melt with continental crust components [34]. Although the tonalite (A) and the dacite (C) are K-poor rocks o f quartz dioritic composition, varia-

TABLE 3 Major and trace element analyses of an Archean tonalite derived by partial melting of an eclogite (A), an Archean quartz monzonite derived by partial melting of graywacke (B), a dacite derived by differentiation from a basic parent (C), and a trachyte derived by differentiation from an alkali basalt with interaction of the melt with continental crust (D) (oxides in wt.%; elements in ppm) A SiO2 TiO 2 A120 3 Fe20 3 FeO MgO CaO Na20 K20 H20 P205 MnO K Rb Sr Ba K/Rb Rb/Sr K/Ba Sr/Ba

B 63.3 0.28 19.4 1.03 1.40 1.45 5.30 6.12 0.90 0.51 0.13 0.03

7500 12.0 1127 405 625 0.011 18.5 2.78

C 73.8 0.11 14.9 0.30 0.7~ 0.26 0.93 3.60 5,12 0.08 0.02

42,500 210 156 651 202 1.35 65.3 0.24

D 80.23 0.09 10.61 0.83 0.18 0.12 1.07 3.83 1.47 0.73 0.02 0.03

12,200 20.1 74.7 135

58.1 1.00

15.9 2.72 5.06 0.90 2.40 5.28 4.98 2.19 0.22 0.23 41,300 104 48 520

607 0.269 90.4 0.55

A = sample 19 (DDH4) Archean tonalite from northwestern Ontario [4]. B = sample 31 (DLNWl0) Archean two-mica quartz monzonite from Minnesota [4]. C = sample $229 dacite from Saipan [5]. D = sample 13867 trachyte from Ross Island, Antarctica [36].

397 2.17 79.4 0.092

39 Q

PLAG

OF

Fig. 7. Normative Q-Plag(Ab + An)-Or plot of an Archean

quartz diorite from northwestern Ontario (A), an Archean quartz monzonite from northern Minnesota (B), a dacite from Saipan (C), and a trachyte from Ross Island, Antarctica (D). Projected cotectic lines and minimum melt composition for Ab/An = 2.9 and P H 2 0 = 4 kbars are shown for comparison [26].

IO0

z "- I0

I

I

Ce

Nd

I

~

I

I

Eu C~I

I

I

I

By

Er

Yb

Fig. 8. Chondrite-normalized REE plot using normalizing values o f Sun and Hanson [34] for an Archean quartz diorite

from northwestern Ontario (A), an Arehean quartz monzonite from northeastern Minnesota (B), a dacite from Saipan (C) and a trachyte from Ross Island, Antarctica (D).

tions in the major and trace elements clearly distinguish these rocks as having different origins. Barker et al. [5] note that the relations of the dacite from Saipan to the other rock types on Saipan are not clear and thus its petrogenesis is not easily discerned. This dacite has, however, distinctive features suggestive of its origin. For example, note its rather flat REE pattern, much lower CaO, Al203, and Sr content and the negative Eu anomaly. All of these features strongly suggest removal of a plagioclase component from a basic parent [5]. Hanson and Goldich [2], and Arth and Hanson [3,4] suggest that the 2700 m.y. old Saganaga Tonalite (A) with a 87Sr/86Sr initial ratio of 0.7009 + 0.0002 could not have been derived from a much older pre-existing granitic crustal rock. The low initial 87Sr/86Sr ratio suggests either a parent with a shortlived history in the crust or a basic or ultramafic parent. The low K content of 7500 ppm rules out a parent such as graywacke which has higher K contents, but is consistent with a basic or ultramafic source. The quartzose nature of the rock eliminates an ultramafic or quartz-undersaturated parent because partial melting of such a parent, except at high PH20, would lead to quartz-undersaturated melt [19]. Thus, the most reasonable parent would appear to be a quartz-saturated amphibolite or eclogite. The K/Rb, the Rb/Sr, and Sr/Ba ratios, similar to those for Archean basalts, and the depleted heavy-REE pattern (see Fig. 9) are consistent with derivation of the tonalite by partial melting of an Archean basalt at mantle depths leaving a residue of garnet and clinopyroxene. The 2700 m.y. old two-mica, quartz monzonite (B) from the Giants Range batholith in northeastern Minnesota discussed in Arth and Hanson [4] has an initial 87Sr/a6Sr ratio of 0.7002 -+ 0.0019 [35] suggesting that it also could not have been derived from pre-existing granitic crust. The rock has a composition near the minimum in the Q-Plag-Or system (Fig. 7). The low K/Rb, high Rb/Sr and low Sr/Ba ratios and relatively high K, Rb and Sr contents are not consistent with a basaltic parent with typical K, Rb and Sr contents of 2100, 5.9 and 175 ppm (average values for Archean basalts from Hart et al. [28]). Enriching Rb from 5.9 to 210 ppm would require a minimum of 97% fractional crystallization or a maximum of 3% partial melting, because the melt would

40

D A C I T E S AND SAGANAGA T O N A L I T E IO0

"
o3 uJ

35 % ~

_

o = io o 32 0

R_E_SI_D U E ..... /

PARENT THOLEIITE

0 ...........

35%

I i

Ce

i

Nd

i

i

i

Sm Eu Gd

i

Dy

i

Er

i

i

Yb Lu

Fig. 9. Chondrite-normalized plot of the field of Archean quartz diorites and dacites from northwestern Ontario and northeastern Minnesota of similar major and trace element composition to that of A, the proposed Archean tholeiite parent, calculated melts for 5 and 35% partial melting of the tholeiitic parent leaving a residue of garnet and clinopyroxene, and the pattern for the residue after 20% melting [4].

have to have been enriched over the parent some 36 times, i.e. CL/Co = 36. Assuming D = 0 for Rb, C f f Co = 1IF then F = 3%. However, in order to reduce the K / R b ratio from 360 to 202, residual phases with significant Kd's for K and Rb are necessary. Thus, an F o f 3% is a maximum and a much lower F is necessary. Similar arguments, as well as major element constraints, rule out a dacitic or tonalitic source. F o r the quartz monzonite (B) the only other abundant and short-lived rock type in the greenstone belt which could be the parent are the graywackeargillite sequences. The coarser fractions o f the graywacke-argillite sequences consist o f a mixture of dacitic and basaltic components, neither of which could be a parent for the quartz monzonite. In the sedimentary process, however, K and Rb were enriched in the argillitic fractions giving the graywacke-argillite sequence a higher K and Rb content and lower K/Rb ratio than the source rocks. It was concluded b y Arth and Hanson [4] that the best model for the origin o f the quartz monzonite was by partial melting o f graywacke under conditions of upper amphibolite or hornblende granulite facies

metamorphism. Twenty to fifty percent melting under these conditions produced melts whose calculated Rb, Sr, Ba, and REE contents were essentially identical to those of the quartz monzonite (see Fig. 10). The trachyte (D) is from a volcanic sequence on Ross Island, Antarctica, which varies from basanitoid to trachybasalt to phonolite and to minor trachyte. The major, minor and trace elements, and phenocryst mineralogy are consistent with the development of the sequence by crystal fractionation from a basanitoid parent [34,36]. The trace element analyses for the basanitoids suggest that they are a result of some 10% partial melting o f a garnet peridotite source enriched in light REE. Lead isotope data for the volcanics give a secondary isochron o f 1500 m.y. [37]. Fig. 11 is a plot of the REE and calculated models for a trachybasalt and the trachyte from Ross Island, normalized to a basanitoid assumed to be similar to the parent for this sequence. Table 4 gives the percentage of minerals fractionally crystallized in going from the basalt to the trachybasalt and from the trachybasalt to the trachyte. Table 5 presents a comparison of the data and calculated model values for Sr and Ba concentrations. The effect of precipitating minerals on the REE is more clearly evident when the samples are

10%

0

COMPOSITE

GRAYWACKE

GIANTS RANGE QUARTZ MONZONITES

I00

09 uJ FE:

i ~10 _ RESIDUE

,50°/°

I

I Co

I Nd

Slm

I ~ Eu G

I Dy

I Er

I Yb

i Lu

Fig. 10. Chondrite-normalized plot of the field of Archean quartz monzonites from northeastern Minnesota with major and trace element compositions similar to that orB, the proposed parent (a short-lived Archean graywacke-argillite composite sample), calculated melts, and the pattern for the residue after 30% melting under conditions of upper amphibolite or hornblende granulite facies metamorphism.

41 TRACHYTE

TABLE 5 Comparison of Sr and Ba data in ppm for Ross Island samples with calculated models [34]

7 ~'0

L~ Cle

I

NId

Sire Eu Gd

MODEL

C~~TRACHYBASALT

[

D,

[

Er

I

Nd

I I I

I

Dy

I

Er

I I

Yb Lu

Fig. 11. Rare earth element plots of trachyte and proposed intermediate tracliybasalt normalized to the proposed

basanitoid parent composition in A and calculated patterns for the differentiation sequence shown in Table 4 in B.

TABLE 4 Summary of percents of minerals used in quantitative trace element model calculations for differentiation going from basanitoid to trachybasalt and trachybasalt to trachyte. The trachyte represents 20% of the original melt [34] Basanitoid

23% clinopyroxene 15% olivine + spinel 2% plagioclase

Trachybasalt

40% total minerals Trachybasalt

25% clinopyroxene 12% olivine + spinel 18% plagioclase 7% anorthoclase 3% apatite 65% total minerals

Ba data

Ba model

1000 1200 48

1400 200

350 510 520

560 490

I

DIFFERENTIATION SEQUENCE

Sm Eo Gd

Sr model

Yb Lu

a I I ~ T R A C H Y T E ~C~

It LoI CeI

Basanitoid Trachybasalt Trachyte

Sr data

Trachyte

normalized to a proposed parent, or least fractionated sample in a series, because the patterns are easily compared to distribution coefficients. The increased enrichment of the lighter and less enrichment of middle and heavy REE in the trachybasalt mirrors the clinopyroxene/melt Kd's for the REE. The strong depletion o f the middle REE for the trachyte reflects the removal o f the small percent of apatite with large Kd's for middle REE whereas the negative Eu anomaly reflects a removal o f feldspar. The Ba and Sr data agree fairly well except for the higher calculated Sr abundances in the trachyte as compared to the actual data. Also, the calculated negative Eu anomaly is smaller than the actual Eu anomaly. This may be a result o f using too small a value for the plagioclase/ melt Kd's for Sr and Eu. The Pb and Sr isotopes for this sample are significantly different when compared to the rest o f the volcanics, which has been attributed to crustal contamination [37,38]. If the melt had reacted with plagioclase-rich rocks from the lower continental crust, the plagioclase could deplete the Eu and Sr without necessarily affecting the other trace elements. This might also help explain why this sample is quartz-normative.

7. Geodynamics and tectonic setting By determining the petrogenesis on a suite of rocks from a particular geologic regime it is possible to place limits on the geodynamic factors in the mantle which are responsible for the tectonic or crustal activity in the area at the time of formation of the suite o f igneous rocks. The igneous rocks formed in a given area at a given time probably reflect the source of the geodynamic disturbances. For example, in an island arc regime dacites may be derived by

42 partial melting of the ocean floor basalts in the downgoing slab leaving an eclogitic residue, or they may be derived by differentiation of basalts which are derived by partial melting of the mantle above the downgoing slab. If they are derived by partial melting of basalt converted to eclogite in the downgoing slab, they should have patterns similar to that of the Saganaga Tonalite (,4) in Fig. 8. If they are derived by differentiation of a basalt, they should be similar to the dacite from Saipan (C) in Fig. 8. Thus, a petrogenetic study has the potential of describing how the processes in the mantle are affecting the crustal developments, in this case an island arc. Another example is the origin of greenstone belts. They have been suggested to be either sites of essentially oceanic crust or surficial features on continental crust. The study of the granitic and basaltic rocks in the greenstone belt in northeastern Minnesota and northwestern Ontario [4], where the Saganaga Tonalite and the Giants Range Quartz Monzonite (A and B, Fig. 8) are from, suggests that it is unlikely that any of the igneous rocks in the greenstone belt have been derived by melting of continental crust, but rather their origin can be explained by partial melting of the mantle or of short-lived crustal material-derived from the mantle. This does not prove that the greenstone belt is on an oceanic crust, but strongly suggests it. For, if there were a continental crust underlying this greenstone belt, it would seem that some of the continental crust should have become partially melted and provided some of the volcanics or intrusions. It is important to note that this is an interpretation based on data from this one locality. An interpretation based on one suite of rocks should not necessarily become a model applied indiscriminately to other rocks, even though they may appear superficially similar in composition and geologic setting. At this stage of knowledge, it is best that the geodynamics and tectonic setting for each suite of rocks be interpreted based on its own geologic, petrographic and chemical characteristics.

Acknowledgements I am grateful for the helpful comments from C J . All~gre, S.R. Hart, C.H. Langmuir and R.D. Vocke. This work was supported by the National Science

Foundation grant GA 30747 (Earth Sciences Section) and grant OPP 72-00470 A02 (Office of Polar Programs).

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