Some aspects of experimental petrology

Earth-Science Reviews - Elsevier Publishing C o m p a n y , A m s t e r d a m - Printed in T h e N e t h e r l a n d s

SOME ASPECTS OF E X P E R I M E N T A L P E T R O L O G Y W. S. F Y F E AND W. S. M A C K E N Z I E Manchester University, Manchester (Great Britain)

SUMMARY

Experiment in petrology is by no means new but during the past twenty years or so it has become an integral part of most significant research in petrology. It is a part only, but a powerful one when added to field and tectonic description and the array of modern techniques of describing the chemical and structural features of the mineral phases which make up the rocks. In a review of this length we have made no attempt to summarize all recent work. Rather we have selected certain topics to illustrate petrological problems where experiment is essential to our appreciation of the problem; some may serve to illustrate the difficulties and limitations of experiment while others point the way to a much fuller understanding of petrogenetic factors. METHODS

The vast majority of experimental studies of chemical systems related to rocks have been concerned with finding the stable phases in the system under given values of pressure and temperature. In other words, we have been concerned with finding phase diagrams describing stable equilibrium. Little attention has been given to the rate processes involved in the attainment of equilibrium. All equilibrium phase diagrams express free energy or chemical potential relations and thus a phase diagram can provide or in some cases, be derived from thermodynamic data. Today we have apparatus in which we can carry out experiments under pressure-temperature conditions appropriate for conditions in all the crust of the earth and for part of the upper mantle and, as far as we know, there are no minerals which could not be synthesized in the laboratory for lack of pressure or temperature. Corrosion and control of oxidation states still presents difficulty in some important systems. We cannot duplicate geological time and of this fact we must be constantly aware. The general problem the experimenter faces is as follows. A given chemical system may exist in an array of states, A, B, C etc. We wish to find the conditions where each state represents the minimum chemical potential and hence where all other states will react to form that state. Without the evidence provided by Earth-Sol. Rev., 5 (1969) 185-215

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W. S. FYFE AND W. S. MACKENZIE

rocks, the task is formidable or even impossible for we are little concerned with "making things" although this aspect of experiment has useful applications. Thus in the petrological context, we are concerned most frequently with the task of putting numbers on the conditions of reactions known to occur in nature. If an experimentally determined boundary between two assemblages A and B represents equilibrium between the assemblages, then reversable transformation should be possible across this boundary and unless this is shown experimentally, there is little reason to suppose the boundary represents equilibrium. At present, this criterion of the validity of an experimental study, can be applied to surprisingly few phase diagrams. The fact that many "equilibrium" phase diagrams tend to change with time is thus not surprising. But with experience this situation is rapidly changing and reliable data are being produced at an impressive rate and problems of long standing are being solved. The geometry of a phase boundary must often conform to rather simple thermodynamic relations and in experimental studies, methods of thermodynamic extrapolation and interpolation may be of immense value in limiting the number of necessary observations. It perhaps should be stressed that a small number of well located points, interpolated by thermodynamic methods, may often be of greater significance than a shot-gun approach to a phase diagram. The treatment of the thermodynamics of systems involving liquids (melts) is more difficult than solidsolid or solid-gas systems but this situation should be improved rapidly, particularly with modern computor techniques. In what follows we shall consider some examples relevant to these remarks. In almost every case of equilibria related to metamorphic systems discussed here, the phase relations are not based on a synthetic approach, i.e., use is not made of high free energy starting materials. Ten years ago such materials were widely used. Again, the precision of silicate calorimetry is being improved by the introduction of superior solvents in the form of fused salts (HOLM and KLEPPA, 1966). Many solid state equilibria of the type A % B are being studied by using A-B mixtures and measuring the direction of transformation under conditions where no nucleation barriers exist, barriers which plagued much early work. In cases where some phases show very sluggish growth kinetics, techniques have been used for studying small amounts of transformations on single crystals (FYFE and HOLLANDER, 1964; EVANS, 1965). Solubility measurements in fused salts (WEILL, 1966) and aqueous systems (REESMAN and KELLER, 1965) show great promise as approaches to obtaining accurate thermodynamic quantities in silicate systems and such possibilities will increase our understanding of the liquid state. METAMORPHIC SYSTEMS The system AI2SiOs-SiO2-H20 This system which includes the phases sillimanite, andalusite, kyanite, Earth-ScL Rev., 5 (1969) 185-215

SOME ASPECTS OF EXPERIMENTAL PETROLOGY

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mullite, pyrophyllite and clay minerals, is of importance to almost all metamorphic facies. A steady flow of papers have been concerned with various aspects of the system but it now appears that many of the problems are becoming solved. We shall mention here only recent and more critical data. Metamorphic petrologists have attached great significance to the occurrence of the three polymorphs of AI2SiO 5, for if the polymorphs are indeed identical in composition, the conditions of their formation are independent of any chemical complexity of the entire system containing them. MIYASmRO (1953) constructed a phase diagram based on natural occurrence which we now know to be of the correct form. But placing numbers on this diagram has given experimentalists great difficulty because reactions between the phases are slow and free energy differences very small. The form of the phase diagram can be based on volume relations (SKINNER et al., 1961) and the high temperature entropy data of PANKRATZ and KELLEY (1964). These data show that over the temperature range of interest: VAnd > Vsm > V~:y and: Ssill > SAnd > SKy

Thus sillimanite must be the high temperature polymorph, kyanite the high pressure form; the boundary between kyanite and sillimanite will have positive slope while the boundary between sillimanite and andalusite must have negative slope. Entropy data indicate that equilibria between A and K or S and K may be easier to establish than the equilibrium between A and S, because the entropy difference of andalusite and sillimanite is only of the order of 0.3 cal./tool./°C. This means that at even quite large distances from the equilibrium temperature, the free energy of reaction will only be a few calories per mole and growth and nucleation, which both require a finite supersaturation, are likely to be very sluggish. The same small entropy difference means that if calorimetric methods are to be used, the precision required is enormous and needs to be two orders of magnitude better than has been commonly achieved with silicates in the past. Direct reversal of the kyanite-sillimanite reaction has been achieved by CLARK (1961) at temperatures near 1,300°C and by NEWTON (1966a) at 750°C. If Clark's high temperature data are extrapolated using entropy and volume relations the lower limit of this extrapolation gives a value similar to that found by Newton. Newton's data can be considered the best presently available and have been confirmed by MATSUSHIMAet al. (1967). The method of pressure calibration used in these studies leaves little room for error. WEILL (1966) measured the solubility of all polymorphs and mullite in molten cryolite and by use of the Temkin model for obtaining chemical potentials of species in a fused salt, obtained relative free energies. Weill's derived figures for the Earth-Sci. Rev., 5 (1969) 185-215

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W. S. FYFE AND W. S. MACKENZIE

kyanite-sillimanite equilibrium are very close to Newton's values and this permits some confidence in the validity of the thermodynamic treatment adopted. Data on the andalusite-kyanite equilibria can again be obtained from direct study of the transition by NEWTON (1966b) and from Weill's data. Again agreement is satisfactory. The andalusite-sillimanite reaction has not been achieved in direct studies and the only relevant data is that due to Weill. The anticipated sluggish kinetics of this reaction are consistent with the observations of the frequent occurrence of both polymorphs in rocks (for example, see COMPTON, 1960). Further, because

I 10

/

't

/

8

/

/

/

/

/

/

P

s

Kilobcr,,4~

Kyonite I //

:

/

/

l

i

0

100

200

300

400

500

600

700

800

900

T°C Fig. l. A possible phase diagram for the system AI~SiO~ consistent with the data of NEWTON (1966a) and WEtLL (1966) and the entropy data. The double boundary lines indicate effects of uncertainty in entropy values. The dashed line is the upper limit of pyrophyllite based on KERRICK (1968).

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of the small free energy and entropy difference, this transition could be impuritysensitive and data are needed on this aspect of the problem. In natural systems, coexisting polymorphs may not be isochemical. Recently, HOLM and KLEPPA(1966) have redetermined the heats of transition by calorimetric methods using a solvent of a fused lead oxide-cadmium oxide-boric acid mixture. The superior solvent properties of this mixture compared to HF solutions enabled these first realistic measurements of the heats of transition. Holm and Kleppa's data indicate rather higher pressures for the sillimanite-kyanite transition than those found by Newton, but when entropy errors are included, the accuracy of resulting free energies is not high enough for the results to have great significance for the andalusite-sillimanite equilibrium. In Fig.1 we show a phase diagram for the polymorphs based on the data of Newton and Weill (see FYEE, 1967) and based on the assumption that this is an example of true polymorphisrn. All data derived from piston-anvil (squeezer) devices have been ignored as uncertainties due to pressure gradients and strained materials may be large and difficult to evaluate. The reaction: AI2Si,Olo(OH)2 ~ AI2SiO s + 3SiO 2 + HzO pyrophyllite ~ andalusite + quartz + water is of importance in limiting the temperature where aluminosilicates may form in quartz rich systems. Almost all previous data on this reaction have been derived from synthesis experiments and it is not surprising that the present information is not satisfactory. Early data of RoY and OSBORN (1954) and KENNEDY (1959) gave an upper limit for pyrophyllite at pressures of 3 kbar at about 575°C. If these data are placed on Fig.1 it is clear that the andalusite field becomes very small indeed. Recent data of MATSUSHIMA et al. (1967) and ALTHAUS (1966) indicate temperatures nearer 500 °C for the reaction. Equilibrium was not demonstrated by the former workers and the boundary given by Althaus is too steep and gives an unreasonable entropy of reaction. Recently, HEMLEY (1967) has studied SiO2 activity of assemblages in the system AI203-SiO2-H20. Consider the case of the above reaction. At temperatures in the stability field of pyrophyllite, the silica activity in the solution must be less than that of quartz, and given good solubility data for quartz, one should predict at what point quartz should precipitate. Using this technique, Hemley reports equilibrium between the assemblage andalusite-quartz and pyrophyllite at 1,000 bar and 400 ___ 15°C. This value may be compared with data of KERRICK (1968) who used weight changes on quartz and andalusite crystals and found values of 430 ° ± 10 ° at 3.9 kbar and 410 + 10 ° at 1.8 kbar. As silica in solution can be determined with considerable accuracy, it is obvious that Hemley's method is of great value because it relies only on one rate process, solution, being rapid.

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W . S. FYFE A N D W . S. M A C K E N Z I E

The glaucophane-lawsonite schist facies Over the past few years, the glaucophane-lawsonite schist facies has attracted considerable attention from petrologists. This interest has arisen from varied causes including the occurrence of numerous high-pressure phases, a close relationship to eclogites, the old problem of the Steinmann Trinity, and the geophysical implications in mantle convection and ocean floor spreading. The general position and significance of this facies was far from clear ten years ago but now we possibly know more of this than of many other more common facies. Previously, widely divergent opinions were held as to the conditions necessary to produce the spectacular schists containing garnet-lawsonite-glaucophane etc., which were often associated with eclogites. These opinions ranged from shallow metasomatic conditions near serpentinites, the influence of trapped saline solutions, to conditions of deep burial metamorphism. The finding of metasediments with similar mineralogy and including jadeite-quartz-lawsonite-aragonite-glaucophane metagraywackes, has clarified the general situation by showing that this type of metamorphism is regional in extent in the sense that cubic miles of rock have been subjected to the physical environment necessary to produce this type of mineralogy. It is now also known that glaucophane schists show some characteristics transitional to the zeolite facies, eclogite facies, greenschist facies and even the amphibolite facies. The general situation has been summarized by ESSENEet al. (1965). It appears that some simple and characteristic reactions of this metamorphic facies include the following changes in response to increasing burial: (1) calcic zeolites (laumontite) ~ lawsonite; (2) plagioclase ~ lawsonite + albite; (3) calcite ~ aragonite; (4) albite ~ (impure)jadeite + quartz. In rocks of almost all stages bearing lawsonite, soda pyroxenes including omphacite are common (EssENE and FYFE, 1967). We may briefly consider some of the experimental studies which bear on the problem of origin of these rocks. Phase relations are summarized in Fig.2. Studies of the reactions: NaA1Si20 6 • H 2 0 + SiO 2 --+ NaA1Si30 8 + H 2 0 analcite + quartz ~ albite + water by synthesis methods indicated that analcite + quartz might be stable to at least 300 °C and WINKLER (1965) used this figure as a measure of minimum temperatures for the facies. Studies by CAMPBELL and FYvE (1965) of this reaction, using a method where the apparent solubility of albite in analcite-quartz mixtures was measured, indicated that the upper limit of the analcite-quartz assemblages in the presence of liquid water was near or below 190 °C. With this data and the knowledge of entropies and volumes of all phases, the boundary shown in Fig.2 was constructed. Analcite-quartz assemblages thus have a very restricted field of stability

Earth-Sci. Rev., 5 (1969) 185-215

19l

SOME ASPECTS OF EXPERIMENTAL PETROLOGY

,

,i

/

/

I

/L

JJ/z /

Zoisi~e+ Kyanite + Quartz / /

#J j'

//

/

/

///

//[awsonile

/

~,northite

\

~,nalcite + QuailT.

Albile

20Q

4110 T[MPERATURE (°C)

600

Fig.2. Phase boundaries related to the formation of glaucophane-lawsoni~e schist facies.

and if solutions are saline the field becomes even smaller. These remarks do not apply to silica deficient rocks where the field of analcite is much enlarged. The calcic zeolite laumontite is not uncommon in less altered rocks bordering the glaucophane schist facies. No direct experimental data are available for the reaction: CaA12Si~O12 • 4HzO ~ CaAlzSizOv (OH)2 • H 2 0 + 2HzO + 2SiOz laumontite ~ lawsonite + water + quartz but thermodynamic data (CRAWVORD and FYVE, 1965) indicate a stability field as indicated in Fig.2. These data require direct measurement as well and again the single crystal experiment method may be quite suitable. The upper temperature limit of lawsonite has been measured by NEWTON and KENNEDY (1963) and CRAWFORD and FYFE (1965). The former workers determined the field of zoisite-kyanite and quartz while Crawford and Fyfe studied only the lawsonite-anorthite equilibrium (the single line of Fig.2). In both cases mixtures Earth-Sei. Rev., 5 (1969) 185-215

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W . S. FYFE A N D W . S. M A C K E N Z I E

of actual phases were used, not synthetic methods. Agreement between the two sets of results is satisfactory. These data provide an upper limit only but it will be noted that unless very high pressures are considered, they show that lawsonitebearing rocks must form at temperatures below 400 °C. In systems of more complex chemistry phases such as prehnite, pumpellyite and epidote must encroach on the lawsonite field and reduce it so that in general the laumontite and lawsonite fields occupy a rather small area of the P - T grid. The equilibrium between calcite and aragonite is well established by numerous studies (see JAMIESON, 1953; CLARK, 1957; CRAWFORO and FYEE, 1965). The boundaries determined by study of direct conversion and by solubility measurements are in close agreement. This is probably the only metamorphic facies where aragonite crystallizes as a stable phase and is preserved. Studies of the kinetics of the aragonite-calcite reaction (BROWN et al., 1962; Davis and ADAMS, 1965) indicate that metamorphic aragonite will survive unloading only if the thermal gradient during this process is of the order of 10 °/km or less and much larger values (e.g., 13 °/km) are quite impossible if aragonite is to survive. The above remarks refer to the reaction in a dry system but the rates are more rapid in aqueous systems and the survival provides clear evidence that after the close of prograde metamorphism hydrothermal activity is often virtually finished. Thus the data on the minerals aragonite and lawsonite indicate a P - T region bounded by pressures between 5-10 kbar and temperatures of 300°C or less. The conditions of the reaction: albite ~ jadeite + quartz were investigated by BIRCH and LECOMTE (1960) at high temperatures and more recently at lower temperatures by NEWTON and SMITH (1967). The curve shown in Fig.2 is based on data of Newton and Smith. These data are not far removed from those deduced by FYFE and VALPV (1959) on the basis of the reaction: jadeite --, nepheline + albite and on data on the stability of analcite. It will be noted that the succession of reactions which would be anticipated from all these data is exactly what is observed in rocks and this satisfactory correlation leaves little doubt as to the position of this metamorphic facies. It is doubtful if these conclusions could ever have been reached by classical methods of study alone. There are still problems involving variability of load and water pressures and complexity of solid solution equilibria, but it is unlikely that consideration of these complexities will cause large changes in our conclusions. Eclogites and glaucophane schists The intimate association of eclogites and glaucophane schists has long been noted. Many eclogites show evidence of being transformed by retrograde processes Earth-Sci. Rev., 5 (1969) 185-215

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SOME A S P E C T S O F E X P E R I M E N T A L P E T R O L O G Y

to glaucophane schists but in other cases the two appear to coexist. The position of the eclogites has been to some extent clarified by the findings: (a) that omphacite is a common mineral in many rocks of the glaucophane schist and (b) that amphibolites are also common associates of eclogites and glaucophane schists (EssENE and ]z'YFE, 1967; ESSENE, 1967). Sodic pyroxenes of all types in the system diopsideacmite-jadeite have been found and the P - T conditions for their formation seem to be closely in accord with what might be expected on an ideal mixing model. it seems that the composition of sodic pyroxenes in equilibrium with albite and quartz in such rocks may be a sensitive pressure indicator and they may be useful for plotting isobaric surfaces in rocks. As they are frequently present in the metasomatic aureole rocks of serpentinites, they may also serve as useful pressure indicators during the serpentinization process (see also LEONARDOSand FYVE, 1967).

/ 30

--

25

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/

20--

/

|

/ /

15--

/

/

/

10--

f

/ Y

J J

/

/

0

k 2OO

1

40Q

I~_ 60Q

~ 800

1 ~

I _ _ 1290

i 1400

1600

TEMPERATOnE ~°C

Fig.3. The basalt-eclogite transition region (shaded) after GREENand R~NGWOOO(1967b), It is clear from the work of GREEN and RINGWOOD (1967b) that eclogites require no higher pressures for formation than the jadeite metagraywackes of the Franciscan of California. Their data for the dry basalt-eclogite transition are shown in Fig.3. A point which requires clarification involves the stability of eclogites under conditions where water pressure equals total pressure. YODER and T1LLEY Earth-Sci. Rev., 5 (1969) 185-215

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W. S. FYFE AND W. S. MACKENZIE

(1962) have indicated that eclogites are unstable in the presence of water, a conclusion based on their findings that amphibolites occur in the high P - T region of the basalt phase diagram. But the fact that all the eclogite facies minerals occur in the wet glaucophane schist environment might indicate that eclogites themselves are stable. Such a possibility cannot be ruled out by existing data. Eclogites are much denser tham amphibolites or greenschists and the condition holds: VAmphib" > VEclogit e

Jr- V H z O

so that amphibolites can be dehydrated at high water pressures. ESSENE and FYFE (1967) have suggested phase relations similar to those of Fig.4 but this requires experimental confirmation. '

1

~--T

1

l

t

T

GL 10

AMPH

GN

t..a=a t=a..a

5

ZEO

200

400 600 TEM~RATIIRE (°0)

800

Fig.4. Regions of formation of the main metamorphic facies as proposed by MIYASHIRO (1961).

The association of jadeite metagraywackes, eclogites and amphibolites follows from the facies relation of M1YASHmO (1961; Fig.4). Recent data on the phase relations of the aluminosilicates and jadeite would lead us to reduce the Earth-Sol. Rev., 5 (1969) 185-215

SOME ASPECTSOF EXPERIMENTALPETROLOGY

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pressure scale shown on this diagram and further we might change the pressure coordinate to water pressure but the relative position of the facies boundaries seems reasonable. If thermal perturbations are caused by deep intrusion ofgabbroic material into sediments in the glaucophane schist region, on the basis of this diagram we would expect the aureole rocks to be eclogites and amphibolites as observed, and heat flow theory (see JAEGER, 1959) indicates a scale of formation of such types in accord with field observations. It thus appears that we are reaching a moderately reasonable explanation of this unusual association; an explanation which would have been difficult to construct without experimental data. Stability of muscovite Among critical dehydration reaction involved in our understanding of high grade metamorphism and granite genesis must be included reactions such as: KAlz(AISi3Oxo)(OH)z --. KA1Si308 + A1203 + H 2 0 muscovite --, K-feldspar + corundum + water and muscovite + quartz ~ K-feldspar + AIzSiO 5 + water. Earlier work on these equilibria (YoDER and EUCSTER, 1955) showed the difficulties in obtaining equilibrium by methods of synthesis. The situation has now been clarified by the work of EVANS (1965) and VELDE (1966). Evans used a sensitive method for detecting the rate and direction of the reaction. Weight changes (growth or reaction) on single crystals of corundum in the presence of powdered muscovite and K-feldspar were measured. Dehydration curves were constructed for all the major reactions of muscovite and discrimination was possible between the aluminosilicates, kyanite or andalusite-sillimanite. The stabilities of the latter pair are so similar and their free energies so similar that it would be predicted that this method would not discriminate between this pair. Evans' data have been confirmed by studies of direct conversion by VELDE (1966) and their curves are in part coincident. By using available thermodynamic data and the Clausius Clapeyron equation, it is possible to show that Velde's curves are too steep and Evans' curves are preferable. This single crystal method of detecting equilibrium appears to be most fruitful in many cases of sluggish reactions. It has its particular use when one phase in a reaction has very low rates of growth. Water pressure and load pressure in metamorphism: biotite stability One of the ultimate problems in our study of metamorphic rocks is to be able to determine, for a given environment, all physical parameters operating in that environment. The values of these parameters which include Ptota~, P,2o, Po2, P.~,s etc., T, shear stress and time, must in some way be derived from information contained in the frozen mineral assemblages. The value of such detailed knowledge is obvious in future developments of our complete analysis of the dynamics of the earth. Earth-Sci. Rev., 5 (1969) 185-215

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W. S. FYFE AND W. S. MACKENZIE

A start in this direction has been made in the recent studies of biotite stability by WONES and EU6STER (1965) which illustrates the general features of the approach. i f in a rock we can find a good temperature indicator (say a pressure insensitive solid solution pair or from oxygen isotope fractionation), a hydration dehydration pair and an oxidation-reduction pair then we can simultaneously determine the variables PH2o, Po2, and T. Other reactions, involving polymorphism or solid solution could also determine the total pressure. Wones and Eugster studied reactions of the type: K Fe3AISi3Olo(OH)2 --+ K AlSi308 + F e 3 0 annite ~ sanidine + magnetite + hydrogen

4 +

H2

and extended these studies to the biotite series with magnesium and ferric iron. One end product of their studies was an equation for the equilibrium of an impure K-feldspar and impure magnetite of the form: 3428 - 4212 (1 1ogfH20 =

T

-

x1) 2

--

+ 8.23 -- log aKAlSqO8 -- log

+ log Xl + ~ 1og.fo 2

av%o,

where xl is the mole fraction of annite in the biotite, f the fugacity, and a the activity of the given component in a solid solution. To use this equation to determinefH2o we need to know T a n d J o 2. The equation is to some extent insensitive to pressure. Wones and Eugster applied this equation to some amphibolites. In these rocks use was made of studies in the F e - T i - O system to limitfo 2 - T conditions. The conclusions reached are that during the amphibolite metamorphism water pressures may have been as low as 0.1-10 bar in some regions. These conclusions are striking, for most metamorphic petrologists have considered the general hypothesis that water and load pressures must often be approximately equivalent, at [east until zones of partial melting are reached. Before great weight is placed on the results in this case, it seems necessary that a careful analysis is made of all possible errors in all parts of the analysis and at present this may be a matter of some difficulty. In part, the equation used above stems from an earlier study of annite stability (Eu6STER and WONES, 1962) and the data concerning the equilibrium relations between the minerals and oxygen buffers may leave something to be desired. But the approach is only a beginning and with better buffer systems, and the microprobe to determine compositions of synthetic phases, improvements are within reach and it seems quite certain that reliable data can be obtained. CARBONATE METAMORPHISM The progressive metamorphism of siliceous limestones and dolomites continues to attract attention. Data derived from these rocks may have considerEarth-Sci. Rev., 5 (1969) 185-215

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PETROLOGY

able bearing on the problem of C 0 2 / H 2 0 ratios in metamorphic fluids and because of the larger number of quite closely spaced reactions, these rocks may provide data on thermal gradients in regional and contact metamorphism. The subject has been recently reviewed by TURNER (1967) and readers are referred to this review. Turner has emphasized the value of thermodynamic analysis in such systems, where in general the properties of the solid phases are less complicated than in silicate minerals. Data summarizing some of Turner's conclusions are shown in Table I. It should be noted that calorimetric approaches may be quite TABLE 1 POSSIBLE REACTION TEMPERATURES OF PROGRESSIVE METAMORPHISM

Limiting temperature ('~C) at P below

Index minera& first a p p e a r

disappear

Pco2(water absent)

PH20 -

500 bar

1,000 bar

1,400 bar

500 bar

1,000 bar

1,400 bar

440 440

490 490

520 520 670 730

400 400 490 570 570 640

450 450 540 630 630 700

480 480 580 670 670 730

570 640

630 700

770 830

850 900

900 950

770 830

850 900

900 950

Tremolite dolomite-quartz Diopside Wollastonite Forsterite Monticellite Periclase

tremolite calcite-quartz dolomite-diopside calcite-diopsideforsterite dolomite

Pcoz

good in carbonate systems at low pressures, Large entropy changes mean that free energies do not need to be fixed with too great accuracy and low pressure CO2 (and H 2 0 ) is a poor reaction medium for direct studies. Thermodynamic data f r o m solubilities

As mentioned earlier, relative solubilities of competing mineral assemblages can provide data on relative stabilities. We have seen this approach used in the study of calcite-aragonite equilibria (JAM1ESON, 1953), AIzSiO 5 equilibria (WEILL, l 966) and m a n y other cases could be mentioned. Solubilities may also provide information on complex formation in solutions and these data are of importance in all transport problems in petrology. To obtain data on the free energies or chemical potentials of species in solution and in equilibrium with the solution requires some model of the solution and this problem is difficult except in the cases of simple strong electrolytes and monomeric molecular solutions. Gradually, the treatment of fused salts is improving (TEMKIN, 1945; B~ADLEY, 1962, 1964; Earth-Sei. Rev., 5 (1969) 185-215

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W. S. EYFE AND W. S. MACKENZIE

1968). Our knowledge of high temperature aqueous solutions is also gradually improving (ELLISand FYEE, 1957; COBBLE, 1964) and soon we should be able to obtain useful information from such studies. An example of this method of study is available from the low temperature studies of REESMAN and KELLER (1965). In this work equilibria of the type: BLANDER,

KA1E(A1Si3OIo)(OH)2 + 6H20 ~ K + + 3AlO 2 + 3H4SIO4 + 2H + were studied. If the solubility is known, the AG of solution is given by an expression such as: ,d G =

- R TInK

where K = [K +] [H +]2 [A10 ~]3 [H~SiO4] 3. As standard free energies are available for all the aqueous species, a free energy of formation of muscovite can be estimated. The possibilities and the difficulties with such a method are obvious. One must be sure of solution equilibrium, equilibrium between the species in solution, and the nature of the species. Similar methods could be applied at higher temperatures where equilibrium may be attained more rapidly. Observations on mineral reactions indicate that solution steps are always quite rapid while the same may be far from true of nucleation and growth steps. But before we can use such methods more data are needed on solution species. We already know that high temperature solutions tend to be molecular but little is known yet about states of polymerization. Even for quartz solutions the situation is confused (see WEILL and FYFE, 1964; ANDERSON and BURNHAM, 1965). IGNEOUS SYSTEMS

From the earliest period of experimental studies of silicate systems a great deal of effort was directed to systems which are of interest in rock forming processes. Considerable progress was achieved in interpreting the course of crystallization of igneous rock by studies of such systems as MgO-FeO-SiO/, CaMgSi206Mg2SiO4-SiO 2, CaMgSi206-NaA1Si3Os-CaAlz Si208, NaA1SiO4-KA1SiO4-SiO 2, etc., but there was always the reservation that these systems were too simple for comparison with natural rock systems, particularly in the case of basic rocks where at least ten oxides might have to be considered. The possibilities of carrying out systematic phase equilibrium studies in such complex systems must have seemed very remote even as recently as fifteen years ago. In addition it was becoming clear from knowledge of possible temperature and pressure gradients within the earth's crust and mantle that magma generation might be expected to occur at depths whose corresponding pressure was such that the phases stable at atmospheric pressure were no longer stable and thus melting and crystallization would have to be considered in terms of high pressure phases. Earth-Sci. Rev., 5 (1969) 185-215

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SOME ASPECTS OF EXPERIMENTAL PETROLOGY

Basalts

The study of the origin of basalts has occupied a large proportion of the attention of experimental petrologists during the last few years. Since the subject has been admirably reviewed in a previous issue of this journal by O'HARA (1968) only a brief commentary on this work need be given here. The pioneering work of Yoder and Tilley at the Geophysical Laboratory in this field began in 1956 when they started systematic work on the melting and crystallization behaviour of basalts using natural rock powders as starting materials. The first studies were made at atmospheric pressure and attempts were made to prevent oxidation of the starting materials by sealing the charges in platinum capsules. YODER and TILLEY (1962) found that, depending on the composition of the basalt selected, the primary phase might be olivine, pyroxene or feldspar but the temperature of appearance of the primary phase fell within a fairly small range of temperature and, in each of the rocks, all major phases had appeared in a fairly small temperature range (80°). This led Yoder and Tilley to the conclusion that basalts must be the products of fractional melting of some more basic rocks. Cpx

Ne~

Cox

Cpx

PI

PI

\ px

Opx

ol ol Fig.5. Exploded view of the "basalt" system showing the names of the rocks whose major normative constituents are within the tetrahedra. Alkali-basalt is within the tetrahedron Ne Cpx-OI-P1 but close to the Cpx-O1-P1 plane. (Simplified from YODERand TILLEY, 1962). By analogy with the synthetic system D i - N e - F o - Q , these authors constructed a tetrahedron, the corners of which represented clinopyroxene, nepheline, olivine and quartz and they discussed the main basaltic rock types with reference to this tetrahedron (Fig.5). Two important planes in the tetrahedron are: (a) the plane of Earth-Sci. Rev., 5 (1969) 185-215

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W. S. FYFE AND W. S. MACKENZIE

silica saturation (clinopyroxene-orthopyroxene-plagioclase) which divides the tholeiites from the olivine tholeiites; and (b) the critical plane of silica undersaturation (clinopyroxene-plagioclase-olivine) which divides the olivine tholeiites from the alkali basalts. The latter plane is of particular importance since under conditions of perfect equilibrium it is not possible, from a consideration of the major phases only, for a liquid to pass through this plane at a pressure of one atmosphere and produce both quartz bearing and nepheline bearing products since this plane contains maxima on all liquidus surfaces in the corresponding synthetic system. Yoder and Tilley also studied the melting relations of basalt in the presence of water at pressures up to 10 kbar, and by a short extrapolation of their data, they concluded that for all of the compositions investigated, an amphibole would be the primary phase at a pressure of about 12 kbar. They reasoned that the settling and resorption of an amphibolite could be a possible mechanism for converting an alkali basalt magma to a tholeiitic magma or vice versa depending on the composition of the amphibole. BOWEN (1928) suggested such a mechanism for the formation of alkalic liquid and Yoder and Tilley considered that their experiments added support to Bowen's hypothesis if the amphiboles were Ne-normative. They were, however, unable to be certain whether a tholeiitic liquid could precipitate the requisite Ne-normative amphibole since the compositions of the amphiboles formed in their experiments were unknown. Yoder and Tilley then considered the melting relations of a number of eclogites at pressures up to 10 kbar both dry and in the presence of water. In one case they found that at 10 kbar the charge consisted entirely of clinopyroxene 75 below the liquidus temperature. Otherwise these rocks behaved in a similar fashion to the basalts of similar composition. At pressures in the range 20-40 kbar they were able to synthesize garnet-clinopyroxene assemblages from each of the basalts they studied except in the case of an olivine nephelinite which produced a spinel-clinopyroxene-mica assemblage. They deduced that the two main magma types could be produced by partial removal of either the omphacite to yield a tholeiitic liquid or the garnet to yield an alkali-rich liquid. One of the most significant results of this work was to show that the thermal divides which exist at low pressure will be replaced by other barriers at high pressures. This topic has recently been dealt with in considerable detail by O'HARA and YODER (1967). YODER and TILLEY(1962) found that in the eclogites studied at high pressures the garnet and clinopyroxene begin to me)t at about the same temperature and that the melting interval is small and they concluded that the eclogites like the basalts must be a partial melting product of a more primitive rock. O'HARA (1965) has noted that the melting interval of eclogites is really relatively large and that it would be a remarkable coincidence if the partial melting products of a garnet peridotite at high pressure were chemically similar to those partial melting products which might be produced at low pressure since the phase assemblages Earth-Sol. Rev., 5 (| 969) 185 215

201

SOME ASPECTS OF EXPERIMENTAL PETROLOGY

are completely different. O'Hara has cast doubt on the possibility that basalts or eclogites are the products of liquids formed by partial melting of peridotite and in the case of eclogites, he interprets these as the crystal accumulates formed from the partial melting of garnet peridotite. The hypothesis of orthopyroxene fractionation as a means of producing nepheline-normative rocks from a tholeiitic magma has been recently revived by GREEN and RINGWOOD(1964, 1967a) since these workers found that orthopyroxene is the primary phase in certain tholeiitic compositions at pressures of 13.5-18 kbar. O'HARA and YOt)ER (1967) have been somewhat critical of this proposal and note that at the beginning of melting of a garnet lherzolite at high pressure the following reaction takes place: garnet + clinopyroxene + olivine ~ orthopyroxene + liquid If this liquid lies in the phase volume garnet-clinopyroxene-orthopyroxene-olivine, fractional crystallization of orthopyroxene would not result in a very significant change in the composition of the resulting liquid. However, the reaction relationship of orthopyroxene means that a liquid fractionating from a partial melt of garnet Iherzolite may produce a garnet-clinopyroxene rock if olivine does not crystallize. (These garnet-clinopyroxene rocks may be crystal accumulates from high pressure fractionation of garnet-lherzolites as mentioned above.) This liquid could yield compositions corresponding to tholeiites by fractionation of olivine at low pressure. At high pressure they cannot produce silica-rich liquids since they cannot pass through the thermal divide garnet-clinopyroxene-orthopyroxene. The restriction of the occurrence of peridotite nodules to alkali basalt types has been a topic of much interest to both field investigators and experimentalists. GREEN and R~NGWOOt~ (1967a) state that the common mineral assemblages of these nodules, viz., olivine, clinopyroxene, orthopyroxene, at the depth at which magma is formed (35-70 kin), is in equilibrium with an alkali olivine basalt magma whereas this assemblage would be unstable in contact with a tholeiitic magma under these conditions since although the olivine and orthopyroxene could remain, the clinopyroxene would be completely dissolved. This seems to be true since, if one looks at GREEN and RINGWOOD'S diagrams (1967a), at 11 kbar, olivine, orthopyroxene and clinopyroxene are all stable at the liquidus of their alkali-olivine basalt, whereas in the case of an olivine tholeiite melt under the same conditions, only olivine and orthopyroxene are stable at the liquidus. If one looks at their data for higher pressures in the region where magma formation might be expected it is difficult to reconcile these observations of their alkali-olivine basalt, e.g., at 13.8 and 18 kbar. Olivine is not a stable phase at the liquidus according to their data and hence the liquid cannot be said to be saturated with olivine--on the other hand, Green and Ringwood note that the olivine tholeiite composition is almost saturated with olivine between 12 and 20 kbar since their picrite composition has olivine as the primary phase under these conditions. Earth-Sci. Rev., 5 (1969) 1 8 5 - 2 | 5

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W. S. FYFE AND W. S. MACKENZIE

Green and Ringwood also note that garnet is not a liquidus phase in any of the basaltic compositions which they have studied below 27 kbar pressure and so garnet would not, in their view, play a significant role in the genesis of magmas by fractional melting at a depth smaller than 100 kin. This they put forward as a serious objection to the ideas of YODER and TILLEY (1962) and O'HARA (1965) that garnet peridotite is the parental material from which basaltic magmas are derived by partial melting. Again from Green and Ringwood's diagrams the situation is not quite as extreme as they appear to have made it since no data are given for the pressure range between 27 kbar and 18 kbar for the olivine basalt and the alkali olivine basalt. In the case of the olivine tholeiite at 22.5 kbar garnet appears 40°C below the liquidus temperature and in the picrite, garnet appears 30°C below the liquidus, and it is possible that garnet might be a near liquidus phase or the primary phase at a pressure of 22.5 kbar for the alkali basalt or olivine basalt corresponding to a depth of about 70 km. More recently GREEN and RINGWOOD (1968) have studied the melting relations in a suite of calc-alkaline igneous rocks under both dry and wet con-

Pressure 30kbar , clinopyroxene • garnet • quartz

1500

"

\

11 "

\

\

so

s;

"if,

60 wr %si%

,#

6k

I'

"

Fig.6. Extrapolated liquidus temperature for a series of rock compositions at 30 kbar pressure dry, with sequence of crystallization of the major phases. (From GREENand RIN~WOOD, 1968.) ditions and find that at 30 kbar, the liquidus temperature of an andesite containing about 62 % SiO2 is lower than that of a dacite (65 % SiO2) and a rhyodacite (69.6 % SiO2) under dry conditions (Fig.6)i These latter rocks have quartz as a liquidus phase, whereas the more basic rocks have a garnet primary phase. Under wet conditions garnet is also the primary phase in the dacite at 27 kbar. Although Earth-Sci. Rev., 5 (1969) 185-215

SOME ASPECTS OF EXPERIMENTAL PETROLOGY

203

lhey do not remark on it particularly, it seems from their data that even at 18 kbar andesite has a lower liquidus temperature than dacite. T~LLEY et al. (1967) have noted a remarkable similarity in the liquidus temperatures in a suite of Paricutin lavas of intermediate composition and in some Thingmuli lavas, despite the iron and alkali differences in the later members of these two series. It seems clear that a considerable amount of valuable information is to be gained by systematic experimental study of natural rock series to complement the recent work on synthetic systems related to the origin of basalt such as the studies by OSBORN (1959), PRESNALL (1966), ROEDER and OSBORN (1966), and others. It does seem to the present writers that the greatest problem the field petrologist has to face is how to relate the studies of one group of experimentalists to those of another group since the bulk chemical compositions used have so many variables. Since rocks cannot be classified in the same rigid fashion as with minerals, it may be difficult to simplify this problem but the writers feel that the time may not be far off when agreement may have to be reached on a few standard basic and ultrabasic rock compositions to which reference can be made. Ultrabasic rocks BOWEN and TUTTLE (1949) in a study of the system M g O - S i O 2 - H z O found no liquid at temperatures below 1,000 °C and Pn2o = 2 kbar and, of course, they were forced to the conclusion that magmatic intrusion could not be responsible for masses of ultrabasic rocks which showed little sign of metamorphosing the envelope of rocks in which they occur. In 1961 CLARK and FVFE described some experiments on the melting of a natural serpentinite and found that "near 1,300 °C the first suggestions of melting were indicated" at P,2o = 1,000 bar. These authors also pointed out that in the case of serpentine masses the demand for water for the serpentine reaction may cause abnormally low heating effects at the contacts. WYLLIE (1960) noted from the unpublished work of RICKER (1952) on the system C a O - M g O - F e O - S i O 2 , that there was a pronounced plateau on the liquidus surface of the Fe2SiO~-Mg2SiO4-Ca2SiO ~ and the effect of this would be for a considerable amount of olivine to dissolve in a basic melt with only a very slight increase in temperature of the melt. He used this argument in support of the belief that magmas containing a considerable amount of dissolved forsteritic olivine may be geologically quite important and it is not necessary to appeal to crystal accumulation of olivine when the olivine content is greater than 10-15 ~ of the rock. SEIFERT and SCHREYER (1966) have made a significant contribution to the experimental study of ultrabasic rocks by their findings in the system K z O - M g O SiO2-H20 at P,2o = 1 kbar. They found that compositions in this system containing fairly small amounts of K 2 0 could have liquid present at temperatures as low as 700~C and silica-rich compositions containing I 0 - 1 5 ~ K 2 0 w e r e completely liquid at 900°C. Except in the case of mica-peridotites the alkali content of which may have been subsequently removed in the gas phase, the K 2 0 content Earth-Sci. Rev., 5 (1969) 185-215

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w. s. FYFE AND W. S. MACKENZIE

of ultrabasic rocks is very low. Certainly this study serves to draw attention to the dramatic effects on temperatures of beginning of melting which may be produced by small amounts of low melting components, whose possible presence is sometimes ignored. The descriptions of the aureoles surrounding some peridotites give evidence that these may have been emplaced at relatively high temperatures of about 1,000 °C (GREEN, 1964) although for a long time it was believed that peridotite aureoles were of universally low temperature. Granites

The now classic studies of TUTTLE and BOWEN (1958) on the system NaAISi308-KA1Si308-SiOz-H20 have been extended in recent years, firstly by L u T n e t al. (1964), who have studied the phase relations in the pressure range from 4 kbar up to 10 kbar Pn2o. These authors found that the eutectic composition moves almost directly towards the Ab apex of the system from its position at lower pressure when it is a minimum rather than a eutectic. LuTn et al. (1964) noted that there is a slight difference in the average composition of aplites and pegmatites compared with granites when their normative constituents are plotted in the triangle Ab-Or-Q. Since the compositions of aplites and pegmatites concentrate slightly in the direction of the Ab corner, these authors suggest that such rocks were formed from liquids saturated with water at the beginning of crystallization and so crystallization may have been accomplished at rather high values of PH2o, thus explaining their slightly sodic character as compared with average granites. One may infer from this that these authors believe that the accompanying granites are not saturated with water at the beginning of crystallization and that only in the later stages, when pegmatite or aplite formation is likely to take place, the liquids are nearly saturated with water. VON PLATEN (1965), WINKLER (1965) and JAMES and HAMILTON(1968) have investigated the effects of the addition of CaAlzSi208 to the "granite" system but so far the PH20 limit is 2 kbar. The main conclusion of all of these studies is that the plagioclase-alkali feldspar-quartz univariant curve passes from near the A b - O r Q face of the tetrahedron A b - A n - O r - Q to the A n - O r - Q face and at 1 kbar meets this face at about 1 0 ~ An. As a result of this the composition of liquids formed by partial melting of sediments and other rocks is very strongly influenced by the CaO content of the initial composition; thus the richer in CaO the rock the greater the Or content of the first liquid formed. It is clear from these studies that the effect of the addition of CaAI2Si208 as a component has almost the opposite effect to that of increasing pressure on the "granite" system and there is an obvious need for studying the system O r - A b A n - Q - H 2 0 at pressures greater than 2 kbar especially since current petrologic thinking is tending more towards the view that the larger granitic bodies are the result of partial melting or remelting of pre-existing rocks. Earth-Sol. Rev., 5 (1969) 185-215

SOME ASPECTS OF EXPERIMENTAL PETROLOGY

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The most surprising results which have been obtained in studies of the melting of granite have been obtained at very high pressures (BOETTCHER and WYLLIE, 1968). These authors, using a natural granite containing approximately equal amounts of quartz, alkali feldspar and plagioclase, determined the beginning of meRing at HzO pressures up to 30 kbar. The temperature of beginning of melting falls to 605 °C at 15 kbar and thereafter begins to increase so that the curve representing the beginning of melting acquires a positive slope. The authors relate this to the albite --* jadeite + quartz transition and a further change in slope is again found at the quartz ~ coesite transition. That a minimum temperature of the beginning of melting is to be expected has been predicted by BARTH (1962) and SMITH (1963) but no experimental confirmation of this had been made and in the present case, the change in slope of the melting curves is due to the intervention of the transition albite -~ jadeite + quartz with a large negative A V, and not to the effect of pressure as such, overwhelming the effect of the addition component (H20) added to the system. Extrapolation of the basalt-water solidus to 15 kbar places it lower than the beginning of melting of granites and there is no available evidence that the basalt - H 2 0 temperature of beginning of melting is raised by the intervention of a phase transition. Peralkaline rocks A beginning was made to the experimental study of peralkaline rocks by CARMXCHAEL and MACKENZIE (1963) when it was shown that the addition of acmite and sodium metasilicate to the system NaAISi3Os-KA1Si3Os-SiOz-H20 resulted in a shift in the thermal valley in the "granite system" in such a manner that the pantellerites and comendites appeared to lie in this thermal valley. These authors following BOWEN (1945) concluded that, as the result of the operation of the plagioclase effect, the normal differentiation products of a trachyte would be either rhyolitic or phonolitic, but these would have a peralkaline character. BAILEY and SCHAmER (1964) criticized the work of Carmichael and MacKenzie chiefly on the grounds that the projections of normative amounts of Q, Or and Ab in peralkaline rocks or in synthetic compositions give a distorted view of the relations since it indicates that the pantelleritic liquids appear to become more potassic whereas they tend to become more sodium rich. Bailey and Sehairer introduced a new concept, which they described as the "orthoclase-effect" which implies that the crystallization of an alkali feldspar will tend to fractionate K 2 0 and increase the NazO content of the liquid preferentially. This was based on the erroneous assumption made by Bailey and Schairer that the KzO/Na20 ratio of feldspars in peralkaline rocks is always higher than that in the liquids from which they precipitate, and that the alkali ratio of the crystals was related in a simple fashion to the alkali ratio in the liquid. That this is not the case was shown by THOMPSON and MACKENZIE (1967). In a study of the system NazO-AlzO3-FezO3-SiO2, BAILEY and SCHAIRER Earth-Sci. Rev., 5 (1969) 185-215

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W. S. FYFE AND W. S. MACKENZIE

(1966) found that there were two quaternary eutectics which could be considered as corresponding to peralkaline rhyolites and peralkaline phonolites respectively; both of these eutectics are rich in sodium silicate and are the end stages of fractionation of liquids which are slightly deficient in A1203. Regarding the passage of slightly over-saturated liquids to undersaturated liquids, Bailey and Schairer suggest that this could be effected by the incorporation of iron as Fe 3+ in the feldspars instead of in acmite and in this way silica is abstracted from the liquid. This is a very attractive proposal since even though the amounts of iron which substitute for aluminium in natural feldspars are small, the dominant phase is feldspar and it will have considerable influence on the subsequent composition of the liquid. Although it has been known for a long time that Fe 3 + may replace AI in K-feldspar, the evidence that it could replace AI in Na-feldspar was lacking (FAUST, 1936). Bailey and Schairer obtained synthetic iron albites having refractive index values in the region of 1.550 in their dry synthesis whereas NOLAN (1966) obtained evidence that there was no substitution or extremely limited substitution of Fe 3+ in albite in hydrothermal experiments at Pn2o = 1,000 bar. It is clear that the replacement of A1 by Fe 3+ in alkali feldspars of intermediate composition requires further investigation before this hypothesis can be supported by experimental evidence. OXIDATION-REDUCTION REACTIONS

One of the most significant advances in experimental petrology since the development of hydrothermal experimentation has been the ability to control partial pressures of oxygen in experimental studies under high pressures of water. The first experiments of this type were reported by EUGSTER (1957) and the technique he employed consisted of using as a buffer an assemblage of solid phases which were related to one another by a reaction of the type: A + B = C

+

02 or just A = B + 02

and as long as these phases are present at the end of the experiment, then the oxygen fugacity is constant for a fixed pressure and temperature. If water is present then the fugacity of hydrogen is also fixed. The separation of the solid materials of the buffer from those of the experiment was accomplished by enclosing the charge in a sealed platinum capsule which was surrounded by the buffer materials. Platinum is easily permeable by hydrogen. The buffer materials were enclosed in a gold tube which is not very permeable to hydrogen. In systems in which carbon is present the situation is complicated by the fact that the fugacity of COz must also be controlled and since platinum is not easily permeable by CO or CO 2 the technique of using sealed platinum capsules surrounded by a buffer cannot be used. FRENCH and EUGSTER (1965) evolved a Earth-Sci. Rev., 5 (1969) 185-215

SOME ASPECTS OF EXPERIMENTAL PETROLOGY

207

graphite buffer for the control of oxygen fugacities in such systems. The charge is contained in an open silver tube and this is surrounded by graphite held in a wider silver tube to which the COe has access. The CO2/CO ratio to which the sample is subjected is fixed by the values offo~, and the values offco~ andfco may be calculated so that the presence of graphite + gas can be used as an oxygen buffer. This is more important in the study of metamorphic reactions where graphite is a fairly common constituent but in certain classes of meteorites the presence of graphite is extremely significant in any attempt to discuss the conditions of formation of the meteorite. SHAW (1963) devised a method of controllingf.~ directly and so was able to dispense with the necessity of a solid-state buffer; in addition thejH 2 obtained does not depend on known solid-gas reactions as required by the oxygen buffer technique and so a continuous variation offn2 can be obtained. The sample is contained in a sealed capsule which is permeable to hydrogen as in the oxygen buffer system and argon is used as the pressure medium. The fH2 is controlled by introducing hydrogen in a small capillary tube, the end of which is covered by a metal permeable to hydrogen. SHAW (1967, pp.521-541) suggests the use of either platinum or a palladium-silver alloy AgToPd30, which gives a better rate of hydrogen diffusion than platinum. He suggests that runs of several days to a week or more are necessary to be able to show reversability of oxidation-reduction reactions. THE FUSED STATE

igneous petrology involves the interaction of silicate melts and solids and to appreciate the details of such processes, data on the structure and thermodynamic properties of both solid and liquid phases are required. While our knowledge of the solid phases advances steadily, the same cannot be said of the fused state but a start is being made and with the increasing availability of computors, detailed treatments are possible which would have been difficult previously. Here we mention only a few approaches of interest. Impressive results in reciprocal salt systems have been produced by BLANDER and his co-workers (1963, 1966, 1968). BLANDER and TOPOL (1966) have treated melting topology in systems of the type LiF-KF-LiCI-KC1. In this treatment, as with other similar treatments, the liquid is treated as a mixture of cations and anions and mixing functions involve such ionic mole fractions. In their treatment of the above system, all specific interaction terms are evaluated from the binary system alone and from data on the pure phases. From the binary data the entire system was calculated by computor. The two figures of Fig.7, presenting a calculated and a measured system, show impressive agreement. BLANDEk'S theory (1963) deals with what he terms "conformal ionic mixtures", these being mixtures in which the potential energy of all cation-anion pairs becomes a function of distances and radii of the ions involved. Earth-Sci. Rev., 5 (1969) 185-215

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W. S. FYFE AND W. S. MACKENZIE

LiF AA= [6,3 kcal./mole Z=6 1.0[ ~~ ~ I I ~i . 1I~ p ~ ~ I N ~1 N ,

\\

0 LiCl

O,I

0.2

0.5

0.4

O,5

KF

\

0.6 X~,

0,7

0.8

0,9

A

1.0 KCl

605

LiCI

554

KCI

772

B Fig."/. A comparison of calculated (A) and measured (B) phase diagram in the system

Li-LiC1 KF-KC1. (After BLANDER, 1968.)

Another interesting example of the use of fused salt theory can be found in WEILL'S (1966) use of the Temkin model in his studies of the solubility ofAlzSiO 5 polymorphs in molten Na3AIF6. These reactions are of the type: AI2SiO s + liquid ~ AI203 + S i O 2 (solution). The problem involves finding the silica activities as a function of concentration in the fused salt. The Temkin model assumes that the mixture consists of two interpenetrating lattices of cations and anions and that each can be treated as an ideal mixture. The activity of each ionic species then becomes a function of the ionic mole functions. In this case: nsi4+ "Si4+ -

"Si 4+

4- "A13+ + nNa+

no2 -

aO2- = ;'0 2- + " F and: "SiO2 = a s i 4 + ( a O 2 - ) 2 Despite all the approximations the treatment appears to yield reasonable results and has great flexibility in that, if necessary, various approximations are readily introduced to make it an n-lattice model as may be used in the treatment of a n-site solid. A different approach to some mineralogical examples has been used by Earth-Sci. Rev., 5 (1969) 185-215

SOME ASPECTS OF EXPERIMENTAL PETROLOGY

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BRADLEY (1962). This treatment considers the fused state as ionic and essentially writes a solubility product for the salt. Thus for a salt X+aX-b Bradley writes an expression K r = iX+]" iX-] b where the X + term is a mole fraction calculated on the basis of all ions ( + and - ) in the mixture. Bradley has used this method for fused mixtures such as NaC1-NaF, Ca(OH)2-H20, C a O - C O 2 - H 2 0 , Mg2SiO4-Fe2SiO4 and even highly polymerized systems such as CaSiO3-CaF2 (BRADLEY, 1964). SMITti (1964) has also used computor techniques for plotting liquidus surfaces in complex systems. WHITTAKER (1967) has considered applications of the BERNAL(1964) theory of liquids to problems of fractionation in magmatic crystallization. This approach seems most promising, particularly when today with diffraction techniques and uses of various types of absorption and Mossbauer spectra we are in a better position to look at site geometry and site populations in liquids just as we alreadyare able to do in solids. As such data accumulates we will be in a much stronger position for making revealing thermodynamic analyses of such problems. CRYSTAL-FIELD INFLUENCES

In 1959, WILLIAMS suggested that the fractionation pattern of some trace elements in the Skaergaard intrusion might reflect crystal-field influences. Crystalfield effects are concerned with the electronic energy terms which arise due to asymmetry of the electron arrangement in transition metals and it has been shown that these terms must be considered if we are to explain the differences between individual transition metals and similar non-transition metals. (See CURTIS, 1964; BURNS and FYFE, 1966, 1967.) These influences have important bearing on any problem (metamorphic or igneous) of distribution of such elements, where different site population and site geometry occurs in the coexisting phases. The terms may lead to non-ideal behaviour in solid solutions where ions except those with a d °, d 5 and d ~° configuration are concerned. The coordinate group of anions around a transition metal cation introduces degeneracy into the metal d orbitals. This degeneracy of splitting of the d levels leads to optical absorption and measurement of absorption spectra provides data on the chemistry and symmetry of the coordinate group. Thus absorption spectra measurements and the attendant pleochroic effects can provide data on site populations, ordering etc. We shall here review one much discussed distribution problem to illustrate some general features. It has been known for a long time that nickel is concentrated in earlyformed olivine when basic rocks crystallize. It is also known from melting studies in the system MgzSiO4-NizSiO 4 (RINGWOOD, 1956) that in the end member system, fractionation should not lead to early concentration of nickel. It appears Earth-Sci. Rev., 5 (1969) 185-215

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W. S. FYFE AND W. S. MACKENZIE

that the fractionation loop must invert on passing from the simple system to a rock of gabbroic composition. BURNS and FYFE (1964) studied site population in some silicate liquids by measurement of the absorption spectrum of nickel in glasses of ultrabasic to granitic compositions. It is found that in melts in the pyroxene-olivine end of the series, the transition metal ions are dominantly in octahedral holes in both liquid and solid phases. As feldspars are added, the liquid produces more tetrahedral sites and transition metal ions are distributed between both octahedral and tetrahedral sites. It is well known that most transition metal ions show a very distinct site preference (DuNn'Z and ORGEL, 1957) and it is possible to list site preference energies. We can now consider the problem along the following lines. Consider the two reactions: MgzSiO 4 + 6:4 liquid - NizSiO 4 + 6:4 liquid - -

solution solution

It is assumed that the liquid has a definite array of octahedral and tetrahedral sites. The "solubility" of the two olivines in the liquid will be determined by the entropies and heats of solution. Because magnesium ions show no electronic site preference, they will spread over the available sites with more facility than nickel ions. This effect will tend to make:

AHsol,. Mg z+ < AHsoln. Ni 2+ and

ASsol.. Mg 1+ > ASsol,. Ni 2+ relative to the same functions in the simple binary system. Both factors will lead to a lower solubility for the nickel compound. It is interesting to note that CARMICHAEL (1967) has suggested an inversion in fayalitic olivines in obsidians. The ferrous octahedral site preference energy is much smaller than that shown by nickel and if inversion occurred it would be expected only in the most silica- and alkali-rich rocks. CONCLUDING REMARKS Current research in metamorphic and igneous petrology has unique approaches and objectives but both can proceed only a limited way unless they are associated with experimental, geochemical and geophysical studies. In general, igneous petrology is concerned with the origin of silicate melts and hence the observable products of volcanism are of primary importance. Once their origin is understood, related plutonic rocks will fall into place. The major questions must concern such things as the depth of origin, the phases and bulk chemistry at the source depth, the mechanism of transport and the modifying influences of crystallization and assimilation on the path to the surface. Experimental petrology alone can provide some restrictions to the possible conditions, but coupled with geo-

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physical observations concerning the situation of sources and the density at the source, with suitable description of magma series in a given source area, eventually we should obtain reasonable models of the origin of melts. Ultimate problems must involve such things as thermal perturbations (noise) in the mantle (ELDER, 1968) and crustal re-cycling to great depths (ARMSTRONG, 1968). The latter mechanism, if significant on a large scale must confuse any simple models which involve clear distinction between crust and mantle. Modern metamorphic studies are becoming increasingly concerned with the lask of mapping the dynamic history of the crust and hence the underlying implications concerning the mantle. The time is approaching when we will examine a metamorphic rock and be able to list rather exactly the parameters (P, T, Pt~2o etc.) which operated during the formation of the preserved mineral assemblages. It seems now that the studies of oxygen isotope fractionation (EvSTEIN and TAYLOR, 1967) can provide a reasonable estimate of temperature, solid solution reactions of P and Pn2o etc. Study of the conditions under which all solid solutions are in detailed equilibrium (a null method) is today possible with the use of the microprobe. Dating techniques can fix the various metamorphic events and their relation to tectonic events. This approach opens up enormous opportunities in studies of crustal history and eventually, economic geology. Similar studies in migmatite zones should provide evidence as to the origin of some silicate melts. REFERENCES ALTHAUS, E., 1966. Die Bildung yon Pyrophyllit und Andalusit zwischen 2,000 und 7,000 bar H20 Druck. Naturwissenschaften, 53:105 106. ANDERSON, G. M. and BURNHAM, C. W., 1965. The solubility of quartz in supercritical water. Am. J. Sci., 263: 494-511. ARMSTRONG, R. L., 1968. A model for the evolution of strontium and lead isotopes in a dynamic earth. Rev. Geophys., 6: 175-199. BAILEY, D. K. and SCHAIRER, J. F., 1964. Feldspar-liquid equilibria in peralkaline liquids --- the orthoclase effect. Am. J. Sci., 262: 1198-1206. BAILEY, D. K. and SCHAIRER, J. F., 1966. The system Na20-AleO3 Fe~O:~-SiOz at 1 atm., and the petrogenesis of alkaline rocks. J. Petrol., 7:114-170. BARANV, R., 1962. Heats and free energies of formation of some hydrated and anhydrous sodiumand calcium-aluminium silicates. U.S., Bur. Mines, Rept. Invest., 5900: 1-17. BARTH, T. F. W., 1962. Theoretical Petrology. Wiley, New York, N.Y., 416 pp. BERNAL, J. D., 1964. The structure of liquids. Proc. Roy. Soc., A280: 299-322. BmCH, F. and LECOMTE, P., 1960. Temperature-pressure plane for albite composition. Am. J. Sci., 258: 209-217. BLANDER, M., 1963. Conformal ionic mixtures. J. Chem. Phys., 39: 2610-2616. BLANDER, M., 1968. The topology of phase diagrams of ternary molten salt systems. Chem. Geol., 3:33 58. BLANDER, M. and TOeOL, L. E., 1966. The topology of phase diagrams of reciprocal molten salt systems. Inorgan. Chem., 5: 1641-1645. BOETTCHER, A. L. and WYLLIE, P. J., 1968. Melting of granite with exccss water to 30 kbar pressure. J. Geol., 76: 235-244. BOWEN, N. L., 1928. The Evolution of the Igneous Rocks. Princeton Univ. Press, Princeton, 333 pp. BOWEN, N. L., 1945. Phase equilibria bearing on the origin and differentiation of alkaline rocks. Am. J. Sci., 243A: 75 89.

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BOWEN, N. L. and TUTTLE, O. F., 1949. The system MgO-SiO2-H20. Bull. Geol. Soc. Am., 60: 439~460. BRADLEY, R. S., 1962. Thermodynamic calculations on phase equilibria involving fused salts, Part 1. Am. J. Sci., 260: 374-382. BRADLEY, R. S., 1964. Thermodynamic calculations on phase equilibria involving fused salts, Part IlL Am. J. Sci., 262: 541-544. BROWN, W. H., FYFE, W. S. and TURNER, E. J., 1962. Aragonite in California glaucophane schists and the kinetics of the aragonite-calcite transformation. J. Petrol., 3: 566-582. BURNS, R. G. and FREE, W. S., 1964. Site preference energy and selective uptake of transitionmetal ions during rnagmatic crystallization, Science, 144:1001 1003. BURNS, R. G. and FYFE, W. S., 1966. Distribution of elements in geological processes. Chem. Geol., 1:49 56. BURNS, R. G. and FYFE, W. S., 1967. Crystal-field theory and the geochemistry of transition elements. In: P. H. ABELSON (Editor), Researches in Geochemistry. Wiley, New York, N.Y., 2: 259-285. CAMPBELL, A. S. and FYEE, W. S., 1965. Analcime albite equilibria. Am. d. Sci., 263: 807-816. CARMICPIAEL,l. S. E., 1967. The iron-titanium oxides of salic volcanic rocks and their associated ferromagnesian silicates. Contrib. Mineral. Petrology (Berlin), 14: 36-64. CARMICHAEL, ]. S. E. and MACKENZIE, W. S., 1963. Feldspar-liquid equilibria in pantellerites: an experimental study. Am. J. Sci., 261: 382-396. CLARK, R. H. and FYFE, W. S., 1961. Ultrabasic liquid. Nature, 191:158 159. CLARK, S. P., 1957. A note on the calcite-aragoniteequilibrium. Am. Mineralogist, 42: 564-566. CLARK, S. P., 1961. A redetermination of equilibrium relations between kyanite and sillimanite. Ant. d. Sci., 259: 641-650. COBBLE, J. W., 1964. The thermodynamic properties of high temperature aqueous solutions, VI. Applications of entropy correspondence to thermodynamics and kinetics, d. Am. Chem. Sot., 86: 5394-5401. COMPTON, R. R., 1960. Contact metamorphism in Santa Rosa Range, Nevada. Bull. Geol. Soc, Am., 71: 1383-1416. CRAWFORD, W. A. and FYFE, W. S., 1965. Lawsonite equilibria. Am. d. Sci., 263:262 270. CURTIS, C. D., 1964. Applications of the crystal-field theory to the inclusion of trace transition elements in minerals during magmatic differentiation. Geochim. Cosmochim. Acta, 28: 389-403. DAVIS, R. L. and ADAMS, L. H., 1965. Kinetics of the calcite-aragonJte transformation. J. Geophys. Res., 70: 433-441. DUNITZ, J. D. and ORGEL, L. E., 1957. Electronic properties of transition element oxides, II. Cation distribution amongst octahedral and tetrahedral sites. Phys. Chem. Solids, 3: 318 333. ELDER, J. W., 1968. Convection - - the key to dynamical geology. Sci. Progr. (London), 56: 1-33. ELLIS, A. J. and FYEE, W. S., 1957. Hydrothermal chemistry. Rev. Pure Appl. Chem., 7: 261-316. EPSTEIN, S. and TAYLOR, H. P., 1967. Variation of O18/O ~6 in minerals and rocks. In: P. H. ABELSON (Editor), Researches in Geochemistry. Wiley, New York, N.Y., 2; 29-62. ESSENE, E. J., 1967. Petrogenesis o f Franciscan Metamorphic Rocks. Thesis, Univ. Calif., Berkeley, Calif., 225 pp. ESSENE, E. J. and FYFE, W. S., 1967. Omphacite in Californian metamorphic rocks. Contrib. Mineral. Petrology (Berlin), 15: 1-23. ESSENE, E. J., FYFE, W. S. and TURNER, E. J., 1965. Petrogenesis of Franciscan glaucophane schists and associated metamorphic rocks, California. Beitr. Mineral. Petrog., 11 : 695-704. EUCSTER, H. P., 1957. Heterogeneous reactions involving oxidation and reduction at high pressures and temperatures. J. Chem. Phys., 26: 1760-1761. EUGSTER, H. P. and WONES, D. R., 1962. Stability relations of the ferruginous biotite, annite. d. Petrol., 3: 82-125. EVANS, B. W., 1965. Application of a reaction-rate method to the breakdown equilibria of muscovite and muscovite plus quartz. Am. d. Sci., 263: 647-667. FAUST, G. T., 1936. The fusion relations of iron-orthoclase, with a discussion of the evidence for the existence of an iron-orthoclase molecule in feldspars. Am. Mineralogist, 21 : 735 763. Earth-Sci. Rev., 5 (1969) 185-215

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FRENCH, B. M. and EUGSTER, H. P., 1965. Experimental control of oxygen fugacities by graphitegas equilibriums. J. Geophys. Res., 70: 1529-1539. FRY, N. and FYEE, W. S., 1969. Eclogites and water pressure. Contrib. Mineral. Petrology (Berlin), in press. FYFE, W. S., 1967. Stability of AI~SiO5 polymorphs. Chem. Geol., 2: 67-76. FYFE, W. S. and HOLLANDER, M. A., 1964. Equilibrium dehydration of diaspore at low temperatures. Am. J. Sci., 262: 709-712. FYFE, W. S., TURNER, F. J. and VERHOOGEN, J., 1958. Metamorphic reactions and metamorphic facies. Geol. Soc. Am., Mere., 73: 1-259. FYFE, W. S. and VALPY, G. W.~ 1959. The analcime-jadeite phase boundary: some indirect deductions. Am. J. Sci., 257: 316-820. GREEN, D. ]-I., 1964. The petrogenesis of the high-temperature peridotite intrusion in the Lizard area, Cornwall. J. Petrol., 5: 134-188. GREEN, D. H. and RINGWOOD, A. E., 1964. Fractionation of basalt magmas at high pressures. Nature, 201: 1276-1279. GREEN, D, H. and R/NGWOOD, A. E., 1967a, The genesis of basaltic magmas. Contrib. Mineral. Petrology (Berlin), 15: 103-190. GREEN, D. H. and RINGWOOD, A. E., 1967b. An experimental investigation of the gabbro to eclogite transformation and its petrological applications. Geochim. Cosmochim. Acta, 31 : 767-834. GREEN, D. H. and R[NGWOOD, A. E., 1968. Genesis of the calc-alkaline igneous rock suite. Contrib. Mineral. Petrology (Berlin), 18: 105-162. HEMLEY, J. J., 1967. Stability of pyrophyllite, andalusite and quartz at elevated pressures and temperatures (abs.). Trans. Am. Geophys. Union, 48: 224. HOLM, J. L. and KLEPPA, O. J., 1966. The thermodynamic properties of the aluminium silicates. Am, Mineralogist, 51 : 1608-1622. JAEGER, J. C., 1959. Temperatures outside a cooling intrusive sheet. Am. J. Sci., 257: 44-54. JAMES, R. and HAMILTON, D. L., 1968. Phase relations in the system NaAISi3Oa-KA1SiaOsCaAI2Si2Os-SiO2 at 1 kilobar water vapour pressure. Contrib. Mineral. Petrology (Berlin), 21:111-141. JAMIESON,J. C., 1953. Phase equilibria in the system calcite-aragonite. J. Phys. Chem., 21:13851390. KENNEDY, G. C., 1959. Phase relation in the system AI203-H20 at high temperatures and pressures. Am. J. Sci., 257: 563-573. KERRICK, D., 1968. Equilibrium decomposition of pyrophyllite at 1.8 kbar and 3.9 kbar water pressure. Am. J. Sei., 266: 204-214. LEONARDOS, G. H. and FYEE, W. S., 1967. Serpentinites and associated albitites, Mocassin Quadrangle, Calif. Am. J. Sci., 265: 609-618. LUTH, W. C., JAHNS, R. H. and TtrrTLE, O. F., 1964. The granite system at pressures of 4 to 10 kbar. J. Geophys. Res., 69: 759-773. MATSUSHIMA, S., KENNEDY, G. C., AKELLA, J. and HAYGARTH, J., 1967. A study of equilibrium relations in the systems AI203-SiO2-H20 and A1203-H20. Am. J. Sci., 265: 28M4. MIYASH1RO, A., 1953. Calcium poor garnet in relation to metamorphism. Geochim. Cosmochim. Acta, 4: 179-208. MIYASFIIRO, A., 1961. Evolution of metamorphic belts. J. Petrol., 2: 277-311. NEWTON, R. C., 1966a. Kyanite-sillimanite equilibrium at 750°C. Science, 151: 1222-1225. NEWTON, R. C., 1966b. Kyanite-andalusite equilibrium from 700 ° to 800"C. Science, 153: 170-172.

NEWTON, R. C. and KENNEDY, G. C., 1963. Some equilibrium relations in the join CaA1Si~OsH20. J. Geophys. Res., 68: 2967-2983. NEWTON, R. C. and SMJTH, J. V., 1967. Investigations concerning the breakdown of albite at depth in the earth. J. Geol., 75: 268-286. NOLAN, J., 1966. Melting relations in the system NaA1SiaOs-NaA1SiO4-NaFeSi~O~-CaMgSi206H20, and their bearing on the genesis of alkaline undersaturated rocks. Quart. J. Geol. Soc. London, 122:91-118. O'HARA, M. J., 1965. Primary magmas and the origin of basalts. Scot. J. Geol., 1: 19~-0.

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O'HARA, M. J., 1968. The bearing of phase equilibria studies in synthetic and natural systems on the origin and evolution of basic and ultrabasic rocks. Earth- Sci. Rev., 4: 69-133. O'HARA, M. J. and YODER Jr., H. S., 1967. Formation and fractionation of basic magmas at high pressures. Scot. J. Geol., 3:67-117. OS~ORN, E. F., 1959. Role of oxygen pressure in the crystallization and differentiation of basaltic magmas. Am. J. Sci., 257: 609-647. PANKRATZ, L. B. and KELLEY, K. K., 1964. High-temperature heat constants and entropies of andalusite, kyanite and sillimanite. U.S., Bur. Mines, Rept. Invest., 6370: 1-7. PRESNALL, D. C., 1966. The join forsterite-diopside-iron oxide and its bearing on the crystallization of basaltic and ultramafic magmas. Am. J. Sci., 264: 753-809. REESMAN, A. L. and KELLER, W. D., 1965. Calculation of apparent standard free energies of formation of six rock-forming minerals from solubility data. Am. Mineralogist, 50: 1729-1739. RJCKER, R. W., 1952. Phase Equilibria in the Quaternary System CaO-MgO FeO-SiO~. Thesis, The Pennsylvania State University. RINGWOOD, A. E., 1956. Melting relations in the system Mg2SiO4 NieSiO4. Geochim. Cosmochim. Acta, 10: 189-202. ROEDER, P, L. and OSBORN, E. F., 1966. Experimental data for the system MgO FeO Fe203 CaAI2-Si208-SiO2 and their petrologic implications, Am. J. Sci., 264: 428480. RoY, R. and OSBORN, E. F., 1954. The system A120:~-SiOz-H~,O. Am. Mineralo,eist, 39: 853-885. SEIEERT, F. und SCHREYER, W., 1966. Fluide Phasen im System K20-MgO-SiO2-H.,O und ihre m/3gliche Bedeutung ftir die Entstebung ultrabasic Gesteine. Ber. Bunseng,es, 70: 1045-1050. SHAW, H. R., 1963. Hydrogen-water mixtures: control of hydrothermal atmospheres by hydrogen osmosis. Science, 139: 1220-1222. SHAW, H. R., 1967. Hydrogen osmosis in hydrothermal experiments. In: P. H. ABELSON (Editor), Researches in Geochemistry. Wiley, New York, N.Y., 2: 521-541. SKINNER,B. J., CLARK, S. P. and APPLEMAN,D. E., 1961. Molar volumes and thermal expansions of andalusite, kyanite and sillimanite. Am. J. Sci., 259: 651-668. SMITH, F. G., 1963. Physical Geochemistry. Addison Wesley, Reading, Mass., 624 pp. SMITH, F. G., 1964. Computation of liquidus relationships in multicomponent silicate systems. Nature, 204: 370-371. TEMKIN, M., 1945. Mixtures of fused salts as ionic solutions. Acta Physiochem. U.S.S.R., 20: 411-420. THOMPSON, R. N. and MACKENZIE, W. S., 1967. Feldspar-liquid equilibria in peralkaline acid liquids: an experimental study. Am. J. Sci., 265: 714-734. TILLEY, C. E., YOOER JR., H. S. and SCHAmER, J, F., 1967. Melting relations of igneous rock series. Carnegie Inst. Wash. Yearbook, 66: 450-457. TURNER, F. J., 1967. Thermodynamic appraisal of steps in progressive metamorphism of siliceous dolomitic limestones. Neues Jahrb. Mineral., Monatsh., 1: 1-22. TUTTLE, O. F. and BOWEN, N. L., 1958. Origin of granite in the light of experimental studies in the system NaA1Si3Os-KalSi3Os-SiO2-H20. Geol. Soc. Am., Mem., 74: 1-153. VELDE, B., 1966. Upper stability of muscovite. Am. Mineralogist, 51: 924-929. VON PLATEN, H., 1965. Experimental Anatexis and Genesis of Migmatites. Controls ~f Metamorphism. Oliver and Boyd, Edinburgh, 368 pp. WEILL, D. F., 1966. Stability relations in the A1203-SiO2 system calculated from solubilities in the Al~O3 SiO2-Na~A1F6 system. Geochim. Cosmochim. Acta, 30: 223-237. WEILL, D. F. and FYFE, W. S., 1964. The solubility of quartz in H20 in the range 1,000-4,000 bar and 400-550 ~C. Geoehim. Cosmoehim. Acta, 28:1243-1255. WHtTTAKER, E. J. W., 1967. Factors affecting element ratios in the crystallization of minerals. Geochim. Cosmochhn. Acta, in press. WtNKLER, H. G. J., 1965. Petrogenesis of Metamorphic Rocks. Springer, Heidelberg, 220 pp. WrLUAMS, R. J. P., 1959. Deposition of trace elements in basic magma. Nature, 184: 44. WONES, D. R. and EUGSTER, H. P., 1965. Stability of biotite: experiment, theory, and application. Am. Mineralogist, 50:1228-1273.

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WYLLIE, P. J., 1960. The system CaO-MgO-FeO-SiO2 and its bearing on the origin of ultrabasic and basic rocks. Mineral. Mag,, 32: 459470. YODER, H. S. and EUGSTER, H. P., 1955. Synthetic and natural muscovites. Geochim. Cosmochim. Acta, 8: 225-280. YODER, H. S. and "['ILLEY,C. [~., 1962. Origin of basalt magmas: an experimental study of natural and synthetic rock systems. J. Petrol., 3: 342-532. NOTE ADDED IN PROOF Since this paper was written there have been many significant advances. Two we would mention include: The March 1969 issue of Am. J. Sci. (267: 257-456), devoted to the AI2SiO,~ polymorphs (see pp. 186-188) and a paper by FRY and FvvE (in press) on eclogites (see pp. 192-194). This latter paper shows that eclogites cannot be formed in the crust unless water pressure is much lower than load pressure. (Received October 15, 1968)

Earth-Sci. Rev., 5 (1969) 185--215