Thermodynamics of the Am Phi Boles - Anthophyllite-Ferroanthophyllite and the Ortho-clino Phase Loop

7/31/2019 Thermodynamics of the Am Phi Boles - Anthophyllite-Ferroanthophyllite and the Ortho-clino Phase Loop

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American Mineralogist , Volume 86, pages 640651, 2001

0003-004X/01/0506640$05.00 640

INTRODUCTION

Accurate representation of calculated phase boundaries for

mineral polymorphs in P-T-X phase diagrams requires that

molar Gibbs free-energy differences be known with high pre-

cision, preferably better than 100 or 200 J/mol. This degree

of precision is seldom attained in calorimetric measurements

of heats of solution or experimental bracketing of mineral re-

actions. Thus, constraints of the usual kinds on the free ener-

gies of polymorphs from mutually independent sources will

generally fail to represent their P-T-Xrelations correctly. Di-rect experimental reversal of polymorphic reaction boundaries

quickly becomes difficult if not totally impossible as tempera-

ture falls much below 1000 C, and extrapolation from high-

temperature conditions is not always possible. In these

circumstances, the occurrence and compositions of polymor-

phs in well-defined natural situations may provide the best in-

formation with regard to their relative stabilities. This seems

to be the case for the phase relations of the orthorhombic and

monoclinic ferromagnesian amphiboles.

* E-mail: [email protected]

Thermodynamics of the amphiboles: Anthophyllite-ferroanthophyllite and the ortho-clino

phase loop

BERNARD W. EVANS,*,1 MARK S. GHIORSO,1 HEXIONG YANG,2AND OLAF MEDENBACH3

1Department of Geological Sciences, Box 351310, University of Washington, Seattle, Washington 98195-1310, U.S. A.2Geophysical Laboratory, Carnegie Institution of Washington, 5251 Broad Branch Road, N.W., Washington, D.C. 20015-1305, U.S.A.

3Institut fr Mineralogie, Ruhr-Universitt Bochum, D-44780 Bochum, Germany

ABSTRACT

Ten new single-crystal X-ray structure refinements of unheated and heat-treated anthophyllite,

new measurements of the optical indicatrix of anthophyllite, and previously published data from

Mssbauer spectroscopy of heated anthophyllite, show that temperature-dependent long-range or-

der of Fe2+ and Mg on the M-sites of cummingtonite-grunerite and anthophyllite may be considered

identical for the purpose of thermodynamic modeling. The difference in solution properties be-

tween the monoclinic and orthorhombic series, as expressed in the composition (XFe) dependence of

lnKD in natural amphibole pairs, is accomodated through adjustment of an enthalpic term that is

independent of order-disorder.End-member thermodynamic properties of cummingtonite and ferroanthophyllite are derived

from those already known for anthophyllite and grunerite respectively, using intercrystalline KDdata and a fit of the T-XFephase loop to two critical field constraints: middle amphibolite-facies

amphibolites and upper amphibolite-facies metaperidotites. Amphibolites suggest a transition tem-

perature in the system FMSH at 555 C andXFe 0.3, whereas metaperidotites suggest a transition

temperature of 650 C at XFe 0.1. LnKD for Fe-Mg exchange between cummingtonite and

anthophyllite passes through zero atXFe 0.7, and as a result the T-XFe phase loop shows a minimum

at this composition.

Extrapolated end-member transition temperatures are estimated to be 800 C (Mg) and 450

C (Fe). At its breakdown to enstatite + quartz + H2O (790 C at 5 kbar), anthophyllite is marginally

stable with respect to end-member cummingtonite, and the addition of Ca renders the breakdown

reaction metastable. A stability field is possible for end-member ferroanthophyllite. Cummingtonite-

anthophyllite phase relations mirror those of the analogous clino- and orthopyroxene.

Compositionally the simplest of all amphibole groups, the

ferromagnesian amphiboles [(Mg,Fe)7Si8O22(OH)2] occur in a

broad range of metamorphic rock types in the form of mono-

clinic cummingtonite-grunerite (predominantly C2/m, but P21/

m in low-temperature cummingtonite) and orthorhombic

anthophyllite (Pnma). In this paper we attempt a quantitative

assessment of their mutual stabilities, using what we already

know about the thermodynamics of the cummingtonite-

grunerite series (Ghiorso et al. 1995; Evans and Ghiorso 1995).

Significant variables are pressure, temperature, and composi-tion [mole-fraction Fe/(Fe + Mg) or XFe]. Our goal is a set of

thermodynamic properties for anthophyllite-ferroanthophyllite

solutions, and an isobaric TXFe phase diagram, the phase loop,

for the dimorphs. Aluminous orthoamphiboles (Al-anthophyllite

and gedrite) will be considered in a later contribution.

Some preliminary reasoning allows us to conclude that both

ends (Fe and Mg) of the phase loop possibly occur at tempera-

tures found in the Earths crust. A compilation of Fe/Mg

intercrystalline partition data from natural parageneses, most

of which probably equilibrated between 500 and 650 C, shows

that lnKD for the exchange reaction:


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EVANS ET AL.: THERMODYNAMICS OF ANTHOPHYLLITE-FERROANTHOPHYLLITE 641

grunerite + anthophyllite = cummingtonite + ferro-anthophyllite

Fe7Si8O22(OH)2 Mg7Si8O22(OH)2 Mg7Si8O22(OH)2 Fe7Si8O22(OH)2

(1)

falls in the range 0.125 to + 0.025 for Al-free compositions

(Evans and Ghiorso 1995, Fig. 3). Assuming some cancella-

tion of non-ideal solution properties between ortho- and clino-amphibole, this range translates into a standard-state Gibbs

free-energy of exchange somewhere between +1.0 and 0.2

kJ/M-cation, perhaps near the average value of 0.4 kJ/M-cat-

ion. Equation 1 may be rearranged:

anthophyllite cummingtonite = ferro-anthophyllite gruneriteMg7Si8O22(OH)2 Mg7Si8O22(OH)2 Fe7Si8O22(OH)2 Fe7Si8O22(OH)2

(2)

to show that the same small standard reaction free-energy is

also a measure of the P or Tdifference between the end-mem-

ber ortho-to-clino reactions. For example, when the Fe end-

member transition (RHS of 2) is at equilibrium, the Mg

end-member transition will be out of equilibrium by an iso-

baric temperature increment Tgiven by:

T= G1/Str (2800)/(4 to 5)

adopting the average free-energy change and assuming that the

transition entropies Str (ortho to clino) are temperature inde-

pendent and both 4 to 5 J/K.mol (Ghiorso et al. 1995, Table 7;

Holland and Powell 1998, Table 5). This back-of-the-envelope

calculation suggests that the difference in temperature of the

ortho-to-clino transitions between the Fe and Mg end-mem-

bers is possibly on the order of 500700 C. Given the pub-

lished record of coexisting cummingtonite and anthophyllite

in nature, we have reason therefore to expect that one of the

end-member transitions, if not both, may be accessible under

the P and Tconditions found in the crust. The difference in the

Mg and Fe end-member transition temperatures for the analo-

gous, Ca-free pyroxenes (Pbca and C2/c) is just under 600 C

at 1 bar (Lindsley 1983; Sack and Ghiorso 1994).

The end-member Gibbs free energies of anthophyllite-Pnma

and grunerite-C2/m are known from bracketing experiments,

and have been incorporated into thermodynamic databases (e.g.,

Berman 1988; Robie and Hemingway 1995; Gottschalk 1997;

Holland and Powell 1998; Chernosky et al. 1998). The solu-

tion properties of the cummingtonite-grunerite series have been

evaluated from data on temperature-dependent, long-range Mg-

Fe order-disorder and experimental equilibria with

orthopyroxene, quartz, and H2O (Ghiorso et al. 1995). In this

paper, we use new single-crystal X-ray structure refinements

of anthophyllite, measurements of its optical indicatrix, and

previously published Mssbauer spectra, to compare long-range

Mg-Fe order in the orthorhombic series with that of the mono-

clinic cummingtonite-grunerite series. If the temperature-de-

pendent Mg-Fe order is the same in the two series, it greatly

simplifies the problem.

Anthophyllite exhibits long-range, non-convergent order of

Fe2+ and Mg over the four sites M1, M2, M3, and M4. The

strong preference of Fe2+ for the M4 site in anthophyllite has

long been known from X-ray structure refinements and

Mssbauer and infrared spectra, as reviewed in Deer et al.

(1997). The temperature dependence and kinetics of the order-

disorder process were studied by Seifert and Virgo (1975) and

Seifert (1977, 1978). The temperature and composition depen-

dence of order-disorder of Fe2+ and Mg on the M-sites of

cummingtonite-grunerite are also known (Hirschmann et al.1994). We shall show below that systematic differences in the

degree of Mg-Fe2+ order-disorder between the two series are

not detectable by the methods currently available. This means

that the real, small difference in solution properties between

the two series, evident from the compositional (XFe) dependence

of lnKD, the intercrystalline partition coefficient of Mg and Fe

(Evans and Ghiorso 1995), can be accommodated by minor

adjustment of one energy parameter in our solution model that

is not a function of ordering state.

SINGLE-CRYSTALREFINEMENTSOFNATURALANDHEAT-TREATEDANTHOPHYLLITE

X-ray single-crystal refinements were conducted on two

samples of anthophyllite (7.3.71.10, withXFe = 0.11, and W82-

009, withXFe = 0.22), both unheated and after separate hydro-

thermal heat-treatment at 600, 700, and 800 C followed by

rapid quenching. Anthophyllite sample 7.3.71.10 is the same

material as studied calorimetrically by Krupka et al. (1985).

An X-ray diffraction (XRD) analysis of a third sample (BM

93327, with XFe = 0.25) was carried out for us by F.C.

Hawthorne, before and after heat-treatment at 700 C. The heat-

treatments were done at 2 kilobars H2O pressure on the C-CH4buffer, for one month (600 C), two days (700 C), and 3 hours

(800 C). Microprobe analyses of the crystals studied by XRD

methods are set down in Table 1. The analyses were done on a

fully automated JEOL 733 Superprobe at the University of

Washington, Seattle, using our library of natural mineral stan-

dards, 15 kV accelerating potential, integration times of 10 to

40 seconds, and the correction factors of Armstrong (1988).

Formula contents (Table 1) are calculated on the basis of an

anion charge of 46 per formula unit (pfu) and all Fe as Fe2+.

This standard procedure tends to propagate error particularly

in the Si determination to the assignment of all other cations to

the C, B, and A groups in the amphibole formula. However,

our average calculated ratio VIAl/Al is 0.38, which we sug-

gest is very reasonable given that this ratio is 0.43 in the ideal

gedrite end-member (Robinson et al. 1971), a composition be-

lieved appropriate for Mg-rich samples (Spear 1980). Of all

the formula assignments, the least reliable is that of Na on the

B and A sites.

The assumption of exclusively Fe2+ in Mg-rich members of

the anthophyllite series is supported by Mssbauer absorption

spectra, in which Fe3+ is typically barely detectable (Bancroft

et al. 1966; Barabanov and Tomilov 1973; Seifert 1978; Stroink

et al. 1980; Law 1989; Ferrow and Ripa 1990). The spectrum

of only one of four anthophyllites studied by Law (1989), his

most Fe-rich sample, with 2.31 atoms pfu Fe, had resolvable

Fe3+ (Fe3+/Fe = 0.024).

Single-crystal X-ray refinements were done at the Univer-

sity of Washington, Seattle, the Geophysical Laboratory (GL),

Washington, D.C., and the University of Manitoba, Winnepeg


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EVANS ET AL.: THERMODYNAMICS OF ANTHOPHYLLITE-FERROANTHOPHYLLITE642

(UM). At the University of Washington, we used a Huber 512

four-circle diffractometer with graphite monochromatized

MoK radiation. XRD intensity data from one quadrant of re-

ciprocal space up to 2 = 65 were collected in the -scan

mode at a scan speed of 3 per minute. Three standard reflec-

tions were checked after every 97 reflections; no systematic or

significant variation in the intensities of the standard reflec-

tions was observed. All XRD intensity data were corrected for

Lorentz and polarization effects, and for absorption by the semi-

empirical method of North et al. (1968). Only reflections hav-

ing intensities 2I were considered as observed and included

in the refinements, where I is the standard deviation deter-

mined from the counting statistics. The laboratories at GL and

UM both used a Bruker SMART CCD (charge-coupled device)

X-ray diffractometer with graphite monochromatized MoK

radiation, but data collection techniques varied slightly: frame

widths 0.3 at GL and 0.2 at UM, frame time 30 s at GL and

60 s at UM, a hemisphere of three-dimensional X-ray data col-

lected up to 60 2 at GL and a sphere of data to 60 2 at UM.

Unit-cell parameters were refined on the basis of all integrated

reflections (>10), using a least-squares technique (GL and

UM). The three-dimensional data were reduced and corrected

for Lorentz, polarization, and background effects using the

Bruker program SAINT (GL and UM). An empirical correc-

tion for X-ray absorption was made (GL and UM) using the

program SADABS (G. Sheldrick, unpublished computer pro-

gram). Equivalent reflections were merged for the final refine-

ments (GL).

Preparatory to structure refinement, the following assump-

tions were made for the assignment of atoms among the crys-

tallographically distinct sites: (1) IVAl = 8.0 Si pfu with allIVAl ordered on T1; (2) VIAl (= Altotal

IVAl), Cr, and Ti were

confined to M2; (3) Ca and BNa were confined to M4; (4) ANa

and K when present were confined to the A-site; and (5) the

fractionation of Fe + Mn among the four M cation sites was

determined from the refinements. After structure refinement,

Mn was proportioned among the M-sites in the same manner

as Fe.

The results of our ten single-crystal X-ray structure refine-

ments are given in Tables 2 and 3. Site-preferences for Fe2+ are

M4 >> M1, M3 > M2; they are smaller at higher temperatures

of equilibration. Preferences can be compared with those of

similarly equilibrated cummingtonites and grunerites

(Hirschmann et al. 1994) with the aid of plots of RTlnKD, the

negative ideal ordering energies - Gid, against macroscopic

XFe (Fig. 1). Because the energy of exchange of Mg and Fe2+

between the M1 and M3 sites is small, on the order of 1 kJ/M-

cation (Fig. 1A), we can combine the occupancies of the M1

and M3 sites for comparison with M2 and M4 according to a

3-site model (Fig. 1B). Finally, by using a 2-site model that

TABLE 1. Chemical analyses of refined crystals of anthophyllite with formulae

Sample no. 7.3.71.10 W82-009 BM 93327Lab. no. unheated C44 C49 C37 unheated C57a C25 C38 unheated C93SiO2 58.9 59.4 58.8 59.1 56.7 56.8 57.1 56.8 56.18 56.30TiO2 0.01 0.00 0.01 0.01 0.03 0.02 0.02 0.03 0.07 0.05Al2O3 0.73 0.51 1.10 0.95 1.09 1.12 1.13 1.09 2.32 2.25Cr2O3 0.15 0.03 0.13 0.04 0.19 0.19 0.17 0.18 n.d. n.d.FeO* 6.26 6.71 6.57 6.74 12.50 12.62 12.20 12.65 14.23 14.03

MgO 30.63 30.44 30.37 30.56 25.55 25.71 26.02 25.94 23.96 23.94MnO 0.12 0.15 0.15 0.14 0.35 0.35 0.35 0.35 0.42 0.40NiO 0.14 0.11 0.11 0.12 0.10 0.10 0.10 0.10 n.d. n.d.CaO 0.48 0.54 0.54 0.57 0.65 0.61 0.71 0.62 0.62 0.66Na2O 0.08 0.06 0.10 0.10 0.11 0.12 0.11 0.11 0.16 0.17K2O 0.00 0.00 0.00 0.00 0.01 0.00 0.00 0.00 0.00 0.00F 0.00 0.00 0.00 0.00 0.03 0.00 0.00 0.02 0.00 0.00Cl 0.00 0.00 0.00 0.00 0.01 0.00 0.01 0.00 n.d. n.d.H2O 2.22 2.24 2.23 2.24 2.15 2.17 2.18 2.16 2.16 2.15less O=F,Cl 0.00 0.00 0.00 0.00 0.02 0.00 0.00 0.01 0.00 0.00Total 99.72 100.19 100.11 100.57 99.45 99.83 100.10 100.04 100.12 99.95

Atomic proportions based on anion charge of 46Si 7.923 7.962 7.892 7.900 7.888 7.879 7.880 7.861 7.814 7.834IVAl 0.077 0.038 0.108 0.100 0.112 0.121 0.120 0.139 0.186 0.166 T 8.000 8.000 8.000 8.000 8.000 8.000 8.000 8.000 8.000 8.000VIAl 0.039 0.043 0.065 0.050 0.067 0.062 0.064 0.039 0.194 0.203Ti 0.001 0.000 0.001 0.001 0.003 0.002 0.002 0.003 0.007 0.005

Cr 0.016 0.003 0.014 0.004 0.021 0.021 0.019 0.020 0.000 0.000Fe 0.704 0.752 0.737 0.753 1.454 1.463 1.408 1.464 1.655 1.633Mg 6.142 6.083 6.076 6.090 5.299 5.311 5.353 5.352 4.968 4.966Mn 0.014 0.017 0.017 0.016 0.041 0.041 0.041 0.041 0.049 0.047Ni 0.015 0.012 0.012 0.013 0.011 0.011 0.011 0.011 0.000 0.000Ca 0.069 0.078 0.078 0.082 0.097 0.091 0.105 0.092 0.092 0.098BNa 0.000 0.012 0.000 0.000 0.007 0.000 0.000 0.000 0.035 0.046 (B + C) 7.000 7.000 7.000 7.009 7.000 7.002 7.003 7.022 7.000 6.998ANa 0.021 0.004 0.026 0.026 0.023 0.031 0.029 0.030 0.008 0.000K 0.000 0.000 0.000 0.000 0.002 0.000 0.000 0.000 0.000 0.000 A 0.021 0.004 0.026 0.026 0.025 0.031 0.029 0.030 0.008 0.000F 0.000 0.000 0.000 0.000 0.013 0.000 0.000 0.009 0.000 0.000Cl 0.000 0.000 0.000 0.000 0.002 0.000 0.002 0.000 0.000 0.000OH 1.992 2.003 1.996 1.997 1.995 2.002 2.007 1.994 2.004 2.000Note: n.d. = not determined.* Total iron. Calculated H2O.


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averages the occupancies of the M1, M2, and M3 sites, we can

incorporate into our comparison (Fig. 1C) the Mssbauer data

for heat-treated anthophyllite obtained by Seifert (1978).

In cummingtonite (but not in grunerite), there is a small

calculated temperature dependence of RTlnKD (see Fig. 3 in

Ghiorso et al. 1995), reflective of excess entropy related to or-

der-disorder. The precision and accuracy of site occupancies

measured by XRD diffraction, in both cummingtonite

(Hirschmann et al. 1994) and anthophyllite (this work), are not

sufficient to show this dependence over the range 600 to 750

C, although a decrease in RTlnKD for the 2-site model showsclearly in the Mssbauer data of Seifert (1978). More impor-

tant in the present context is that, allowing for 2 and all other

possible uncertainties, Figures 1A, 1B, and 1C show no sys-

tematic differences in Fe-Mg order between anthophyllite and

cummingtonite-grunerite, at least in the range ofXFe of the

anthophyllites studied by XRD and Mssbauer spectroscopy

(0.11 to 0.25). Except for exchange between the M1 and M3

sites, cummingtonite and anthophyllite in fact seem to share

small changes in Gid as the Mg end-members are approached.

This similarity may reflect energetic consequences of a com-

positional change that in cummingtonite eventually induces a

change from C2/m to P21/m symmetry.

Only rarely is anthophyllite encountered in nature more Fe-

rich thanXFe 0.30 without it containing significant amounts

of Al. More Fe-rich natural anthophyllite trends composition-

ally through Al-anthophyllite toward gedrite, although a broad

solvus gap below about 600 C separates the two (Robinson et

al. 1971; Spear 1980). X-ray and Mssbauer data (Papike and

Ross 1970; Seifert 1978) show that the uptake of Al in

orthoamphibole is accompanied by increasing Fe2+-Mg disor-

der. Thus, a direct comparison of order-disorder in the simple

system MFSH cannot be taken to more Fe-rich compositions,

unless we resort to the use of synthetic anthophyllite. A rangeof intermediate Fe, Mg-compositions with orthorhombic sym-

metry have been synthesized in the system FMSH (Cameron

1975; Popp et al. 1976), but their space group symmetry was

not determined unambiguously; they are not anthophyllite-

Pnma. This material might be worthy of study in the future.

THEOPTICALINDICATRIXOFANTHOPHYLLITEANDCUMMINGTONITE-GRUNERITE

The optical properties of anthophyllite and cummingtonite

provide supporting evidence for their indistinguishable intrasite

exchange energies. Evans and Medenbach (1997) showed that

the shape of the optical indicatrix (2V) of cummingtonite-

TABLE 2. Cell dimensions and refinement statistics of anthophyllite crystals

Sample T(C) a() b() c() V(3) Refls. R Rw Note7.3.71.10. R.T 18.5219(9) 17.9740(8) 5.2725(4) 1755.3 1779 0.036 0.031 GLC44 600 18.5008(30) 17.9695(22) 5.2738(6) 1753.3 1971 0.034 0.044 UWC49 700 18.5004(25) 17.9646(29) 5.2710(7) 1751.8 1343 0.033 0.043 UWC37 800 18.4908(36) 17.9785(24) 5.2743(6) 1753.4 2000 0.037 0.049 UW

W82-009 R.T. 18.5705(9) 18.0361(8) 5.2876(2) 1771.0 2061 0.030 0.028 GL

C57 600 18.5654(33) 18.0271(21) 5.2845(6) 1768.6 1941 0.036 0.045 UWC25 700 18.5760(10) 18.0367(10) 5.2847(3) 1770.6 1749 0.036 0.031 GLC38 800 18.5668(9) 18.0306(8) 5.2867(3) 1769.8 1979 0.031 0.028 GL

BM 93327 R.T 18.5716(9) 18.0307(8) 5.2903(2) 1771.5 1401 0.029 0.023 UMC93 700 18.5711(6) 18.0256(6) 5.2885(2) 1770.4 1794 0.029 0.024 UMNote: GL, UW, UM, X-ray data collected at the Geophysical laboratory, the University of Washington, and the University of Manitoba.

TABLE 3. Site occupancies of anthophyllite

7.3.71.10 C44 C49 C37 W82-009 C57 C25 C38 BM 93327 C93T(oC) 600 700 800 600 700 800 700M1 Fe 0.017(2) 0.049(3) 0.063(3) 0.065(3) 0.049(2) 0.114(3) 0.119(3) 0.132(2) 0.087 0.151(3)

Mg 0.983 0.950 0.936 0.934 0.950 0.883 0.878 0.864 0.910 0.845Mn 0.000 0.001 0.001 0.001 0.001 0.003 0.003 0.004 0.003 0.004

M2 Fe 0.005(2) 0.024(3) 0.022(3) 0.038(3) 0.014(2) 0.049(4) 0.059(2) 0.071(2) 0.019 0.078(4)Mg 0.967 0.952 0.938 0.934 0.940 0.908 0.897 0.896 0.880 0.815

Al 0.020 0.022 0.033 0.025 0.034 0.031 0.032 0.020 0.095 0.100Mn 0.000 0.001 0.000 0.001 0.000 0.001 0.002 0.002 0.001 0.002Cr 0.008 0.001 0.007 0.002 0.011 0.010 0.009 0.010Ti 0.001 0.001 0.001 0.001 0.001 0.005 0.005

M3 Fe 0.022(4) 0.032(5) 0.048(5) 0.046(5) 0.034(3) 0.093(5) 0.097(4) 0.115(3) 0.058 0.126(4)Mg 0.978 0.967 0.951 0.953 0.965 0.904 0.900 0.882 0.940 0.870Mn 0.000 0.001 0.001 0.001 0.001 0.003 0.003 0.003 0.002 0.004

M4 Fe 0.328 0.291 0.266 0.258 0.672 0.528 0.483 0.474 0.704 0.505Mg 0.631 0.657 0.689 0.696 0.257 0.412 0.451 0.467 0.215 0.405Mn 0.006 0.007 0.006 0.005 0.019 0.015 0.014 0.013 0.021 0.015Ca 0.035 0.039 0.039 0.041 0.049 0.045 0.052 0.046 0.045 0.050Na 0.006 0.003 0.015 0.025

A Na 0.021 0.004 0.026 0.026 0.023 0.031 0.029 0.030 0.008 0.000K 0.002

Note: Assignment of Cr to the M3 as well as the M2 site (e.g., Fialips-Gudon et al. 2000) markedly improves the consistency of intrasite Gid valuesin samples C44, C49, and C37, where Cr is variable.


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grunerite is a sensitive indicator of Fe-Mg order-disorder in a

sample of known composition (XFe). Figure 2 incorporates new

measurements made with a double spindle-stage (Table 4) and

shows that, in the range XFe = 0.0 to 0.3, the 2Vz of natural

anthophyllite is indistinguishable from that of natural

cummingtonite. In this compositional range, the state of order

for natural samples, as measured by s123, the 2-site order pa-

rameter where s123 =XFeM4 average (XFeM1M2M3), changes from

0.0 (end-member cummingtonite) to slightly more than 0.8 at

XFe = 0.30 (Evans and Medenbach 1997). The coincidence in

the behavior of 2Vz vs.XFe is interpreted as indicative of identi-

cal order-disorder in slowly cooled anthophyllite and

cummingtonite. The 2Vz of more Fe-rich natural anthophyllite

is smaller than that of cummingtonite (Fig. 2), but the differ-

ences may reflect the larger amounts of the gedrite component,

as well perhaps as different cooling rates. Significantly,

anthophyllite shows a minimum 2Vz at XFe 0.3, just like

cummingonite, and for the same reason (Evans and Medenbach

1997).

We also measured the optical properties of two anthophyllite

crystals after equilibration in the laboratory at high tempera-

ture, followed by quenching (Table 4). The 2Vz of sample W82-

009 increased from 83.7 to 93.1 when equilibrated at 600 C,

and the 2Vz of sample BM 93327 increased from 80.7 to 92.0

after treatment at 700 C. If the order-disorder behavior of

anthophyllite and cummingtonite is assumed to be identical,

then we can extract values of the two-site order parameter s123based on (1) the measured 2Vz and XFe of anthophyllite using

the cummingtonite Equation 1 of Evans and Medenbach (1997);

and (2) the isothermal fits to the cummingtonite ordering data

shown in Figure 2 of Evans and Ghiorso (1995). For sample

W82-009 at 600 C these values ofs123 are (1) 0.482 and (2)

0.488, whereas the measured value of anthophyllite from Table

3 is 0.477. All three values are in excellent agreement, sug-

gesting that the assumptions made above are reasonable. Cor-

responding values for sample BM 93327 equilibrated at 700

C are (1) 0.481, (2) 0.486, and a measured value of 0.434. In

this case, givenXFe and temperature, the 2Vz of anthophyllite is

behaving exactly like cummingtonite. The smaller measured

s123 of the crystal of BM 93327 is not fully understood; it may

reflect the greater amount of gedrite substitution (Al = 0.4

pfu in this sample), or possibly some interlaboratory differ-

ences. This sample plots off the general trend for cummingtonite

in Figure 1C.

SOURCEOFSOLUTIONPROPERTYDIFFERENCES

We are not able to compare the ordering behavior of more

Fe-rich anthophyllite with that of compositionally similar mem-

bers of the cummingtonite-grunerite series because low-Al

orthoamphibole with XFe > 0.30 is extremely rare in nature.

Such anthophyllite or ferroanthophyllite is largely unstable with

respect to cummingtonite-grunerite gedrite (see below) in

amphibolite-facies parageneses where it might occur; and, at

lower temperatures that would favor anthophyllite rather than

cummingtonite, both ferromagnesian amphiboles are unstable

with respect to more hydrated and carbonated equivalents, such

as minnesotaite, siderite or ankerite + quartz iron oxides.

Accordingly, for modeling solution properties, we shall make

the crystal-chemically reasonable assumption that the long-

range Fe-Mg ordering behavior of more Fe-rich anthophyllite

continues to be indistinguishable from its compositionally

FIGURE 1. Comparison of RTlnKD vs. macroscopic XFe for site

partitioning in heat-treated cummingtonite-grunerite and anthophyllite.

A = four-site model, B = three-site model, C = two-site model. Filled

circles, cummingtonite-grunerite (Hirschmann et al. 1994); open

circles, anthophyllite (this work); triangles, anthophyllite by Mssbauer

spectrometry (Seifert 1978). Error bars are 2 for instrumental

uncertainty (in most cases smaller than the symbol).

70

80

90

100

110

120

130

0.0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1.0

(Fe+Mn)/(Fe+Mn+Mg)

2Vz

FIGURE 2. 2Vz of anthophyllite-ferroanthophyllite (Table 4)

compared to metamorphic cummingtonite-grunerite (Evans and

Medenbach 1997). Symbols as in Figure 1. Included are data for two

samples of proto-ferroanthophyllite from Sueno et al. (1998).


6/12


equivalent monoclinic analog.

Except for M1 vs. M3, intersite (ideal) exchange energies

in both the monoclinic and orthorhombic ferromagnesian am-

phiboles are large in comparison to those related to the accu-

racy and precision of measured site occupancies, and to those

related to possible differences in order-disorder between

anthophyllite and cummingtonite-grunerite. The energetic con-

sequence of temperature dependent order-disorder plays a ma-

jo r ro le in de te rm in in g th e so lu ti on pr ope rt ie s of

cummingtonite-grunerite (see in Fig. 2 Ghiorso et al. 1995),

and the same must also be true for the anthophyllite series.

The two series are nevertheless not identical in their solu-

tion properties. They differ macroscopically because it can be

shown that intercrystalline Fe-Mg exchange between them is

not independent of their Fe contents. A compilation of data for

natural clino-ortho ferromagnesian amphibole pairs was fit to

the expression: lnKD = 0.125 + 0.150XFe + 0.046XFe wt%

Al2O3 (see in Fig. 3 Evans and Ghiorso 1995), where XFe and

wt%Al2O3 refer to the orthoamphibole. This expression implies

for Al-free pairs a difference of 0.15 in lnKD in going fromXFe= 0 toXFe = 1, and zero lnKD atXFe = 0.75. Others have shown

that the zero, or extremum, composition is close toXFe = 0.4

(e.g., Elliot-Meadows et al. 1999), but this value applies to

cummingtonite coexisting with Al-anthophyllite.

Although we cannot rule out small differences in order-dis-

order behavior between the two series, the macroscopic differ-

ence in solution properties is likely to be due, in our opinion,

to one or more other crystal-structure factors. The dimorphs

differ fundamentally in the manner of stacking parallel to c

between adjacent tetrahedral layers (Gibbs 1969; Hawthorne 1983);

the sequence in Pnma anthophyllite is (..++ ++ ..) whereas

in C2/m and P21/m cummingtonite it is (..++++..). There are

also important differences in bond lengths at the M4 site, and

kinking of the tetrahedral chains (Boffa Ballaran et al. 2001).

These considerations allow us to derive the solution prop-

erties of anthophyllite-ferroanthophyllite from those of

cummingtonite-grunerite in a straightforward and simple man-

ner. We shall assume that all the solution parameters that are

determined by order-disorder in cummingtonite-grunerite, as

calibrated by single-crystal X-ray refinements of heat-treated

crystals (Table 9 in Ghiorso et al. 1995), also apply to

anthophyllite-ferroanthophyllite.Ghiorso et al. (1995, Eq. 5)

expressed the vibrational molar Gibbs free energy of

cummingtonite-grunerite solution in terms of a truncated sec-

ond-order Taylor expansion in composition and ordering vari-

ables. Only three coefficients in this expansion are independent

of order-disorder. Two of them are determined from standard-

state properties. The third coefficient, G

*r,r, was determined

for cummingtonite-grunerite from heterogeneous phase equi-

librium constraints. G

*r,r differs between the two series by an

amount that is determined by the difference (0.15) in the lnKDfor intercrystalline exchange at the Mg and Fe extremes (Fig.

3). The required adjustment in G

*r,ris from 11.2 kJ/mol in

cummingtonite-grunerite to 12.2 kJ/mol in anthophyllite-

ferroanthophyllite. G

*r,r is equivalent to W/4 where Wis an

effective macroscopic regular solution parameter. The abso-

lute values of lnKD (the y-axis of Fig. 3) depend on end-mem-

ber standard-state properties.

END-MEMBERTHERMODYNAMICPROPERTIES

Our recommended thermodynamic end-member properties

are listed in Table 5. For end-member anthophyllite, we use

thermodynamic properties given in Ghiorso et al. (1995, Table

7), which are virtually identical to those of Chernosky et al.

(1998, Table 6), and consistent with other minerals in the

Berman (1988) database. The heat capacity, expansivity, and

compressibility of end-member ferro-anthophyllite are assumed

to be identical to those of grunerite. The reaction entropy for

the ortho clino transition of Fe7Si8O22(OH)2 is taken to be 5

J/Kmol, as was done for Mg7Si8O22(OH)2 (Ghiorso et al. 1995),

making the monoclinic structure the high-temperature form.

Hirschmann et al. (1994) showed that, corrected Ca- and Mn-

free, the cell volume of end-member anthophyllite, if at all, is

smaller than that of Mg-cummingtonite by no more than 0.05%.

The mean refractive indices (Table 2 and Evans and Medenbach

1997) for the two series betweenXFe = 0.1 and 0.3 differ by a

little more than 0.001, which, from the Gladstone and Dale

relation (Jaffe 1988), translates into a density difference of

0.15%. If we split the difference between these two estimates,

we arrive at molar volumes that differ by 0.03 J/barmol for

anthophyllite and cummingtonite respectively (Table 5). It is

not known whether the same volume change holds up to the Fe

TABLE 4. Optical properties of anthophyllite

Cations per formula unitSample no. IVAl VIAl Fe Mn Mg Ca Na XFe nx ny nz 2VzUnheated31 0.000 0.038 0.025 0.062 6.730 0.052 0.056 0.013 1.6040 127.07.31.71.10 0.049 0.041 0.779 0.015 6.081 0.077 0.020 0.115 1.6120 1.6248 1.6336 103.896-34 0.033 0.024 0.899 0.036 5.959 0.077 0.017 0.136 1.6147 1.6266 1.6359 99.0121-73 0.000 0.007 1.156 0.029 5.731 0.054 0.004 0.171 1.6290 91.3

W82-009 0.154 0.077 1.460 0.043 5.325 0.090 0.030 0.220 1.6263 1.6359 1.6475 83.7BM 93327 0.150 0.153 1.614 0.049 5.081 0.088 0.024 0.246 1.6292 1.6379 1.6507 80.76A9 0.503 0.439 2.282 0.110 3.986 0.119 0.194 0.375 1.6450 1.6510 1.6610 74.02H-348 0.571 0.480 3.277 0.081 3.076 0.048 0.167 0.522 1.6580 1.6670 1.6810 76.6

HeatedW82-009(600) 0.093 0.091 1.457 0.038 5.301 0.097 0.033 0.220 1.6241 1.6359 1.6482 93.1BM 93327(700)0.228 0.182 1.657 0.048 4.948 0.135 0.059 0.256 1.6289 1.6401 1.6515 92.0Notes: XFe = (Fe + Mn)/(Fe + Mn + Mg).Sources of samples. 3-1: Klein (1968), 7.3.71.10: Krupka et al. (1985), 96-34 and 121-73: Fabris and Perseil (1971) (alsoSeifert 1978, samples AG1 and AG2), W82-009: Hirschmann et al. (1994), BM 93327: British Museum of Natural History,6A9: Robinson and Jaffe (1969), 2H-348: Guiraud et al. (1996).In parentheses: temperature of heat-treatment.


7/12


end-members, but, for simplicity, we shall assume that it does.

These small volume uncertainties affect the pressure depen-

dence of lnKD by a minor amount on the order of 0.001/kbar.

Their effect on the temperature of the transition is small but

not trivial, however, although a meaningful estimate of the dP/dT

of the transition is not possible at the present time.

Still to be evaluated are the enthalpies of the end-members

ferro-anthophyllite and cummingtonite. The Gibbs free ener-

gies of these end-members are extremely close to those of

grunerite and anthophyllite, respectively, and, in the absence

of any experimental constraints, the only way to evaluate the

small differences involved is to analyze data from the field.

Two kinds of field data can be used to constrain the free ener-

gies of the solutions: (1) field information that serves to bracket

the ortho = clino inversion temperature for ideally Al-free

(quadrilateral) ferromagnesian amphiboles as a function ofXFe;

and (2) compositional data that give robust estimates of lnKDfor equilibrium pairs of cummingtonite + low-Al anthophyllite

for specific values ofXFe. These data in combination, together

with the MgFe solution properties of the two phases derived

above, and the remaining end-member properties (Table 5), will

constrain the end-member enthalpies.

FIELDCONSTRAINTS

It has long been known that low-Al anthophyllite is a me-dium-grade metamorphic mineral typical of Mg-rich bulk com-

positions (metamorphosed ultrabasic rocks), whereas

cummingtonite and grunerite are characteristic of intermediate

FeMg and Fe-rich compositions (low-Ca amphibolites, low-K

pelites, uralitized gabbros, and metamorphosed iron-formation

or BIF) equilibrated at metamorphic grades reaching into the

granulite facies. Cummingtonite (XFe from 0.29 to 0.5) also

occurs as phenocrysts in H2O-rich silicic volcanic rocks and as

a possible magmatic mineral in a small percentage of plutonic

rocks. With slow cooling, cummingtonite inverts to

anthophyllite and liberates a small quantity of Ca-amphibole

(Ross et al. 1969; Evans et al. 1974; Carpenter 1982). These

observations, and the sign of the partition coefficient lnKD,

imply that the transition temperature, at least at first, declines

with increasing XFe. The most useful constraints on the T-XFelocation of the transition in the FMSH system are supplied by

amphibolites and metaperidotites.

In Figure 4 we have plotted total Al (apfu) against theXFefor microprobe-analyzed anthophyllite and cummingtonite from

five well-documented examples of middle amphibolite-facies

metamorphism of low-Ca amphibolites and low K-pelites (Stout

1972; Spear 1980, 1982; Clark 1978; Early and Stout 1991;

Schneidermann and Tracy 1991). Cummingtonite and

anthophyllite commonly coexist in these samples and in some

cases gedrite and/or Ca-amphibole occur as well. Metamor-

phic temperatures estimated by these authors fall in the range

535 to 575 C, and pressures from 3 to 6 kbar. Possible field

boundaries for these conditions have been inserted on Figure

4. Except for the base of the plot and the three-phase assem-

blage anthophyllite-gedrite-cummingtonite (which is virtually

coplanar in composition space), phase boundaries are some-

what dependent on the compositions of coexisting minerals, as

FIGURE 3. Computed correlation of lnKD andXFe of cummingtonite-

grunerite for Mg/Fe partition between cummingtonite-grunerite and

anthophyllite-ferroanthophyllite. KD = (Fe/Mg)Ath/(Fe/Mg)Cum

TABLE 5. Internally consistent thermodynamic properties of end-members

anthophyllite cummingtonite ferro-anthophyllite gruneriteMg7Si8O22(OH)2 Mg7Si8O22(OH)2 Fe7Si8O22(OH)2 Fe7Si8O22(OH)2

H0

f,Tr,Pr (kJ) 12073.132 12067.920 9627.015 9623.550S

0Tr,Pr(J/K) 535.259 540.259 720.000 725.000

V0

Tr,Pr(J/bar) 26.31 26.34 27.81 27.84k0 1233.8 1347.83

k1 102 71.3398 93.5691k2 105 221.638 202.285k3 107 233.394 303.919v1 106 1.1394 1.6703v2 1012 8.68919v3 106 28.105 28.400v4 1012 62.894

Notes:C kk k k

T,P

o

r= + + +

0

1 2

2

3

3T T T

(J/K)

V

Vv P P v P P v T T v T T

T,P

o

T ,P

o r r r r

r r

= + ( ) + ( ) + ( ) + ( )1 1 22

3 4

2

(J/bar)

H0

f,Tr,Pr and S0

Tr,Pr of anthophyllite and cummingtonite are given with greater precision than individually known in order to maintain internal consistency.


8/12


illustrated, for example, by the reaction: Cum + Plag = Al-

anthophyllite + Ca-amphibole (Robinson and Jaffe 1969; see

also Spear 1982). Extrapolating to zero Al, the data points sug-

gest a narrow gap between anthophyllite and cummingtonite,

centered onXFe 0.30. From KD data we know the gap here is

about 2 mol% wide. The field of (subsolvus) Al-anthophyllite

terminates at these temperatures atXFe 0.40 (Fig. 4), beyond

which it reacts to cummingtonite + gedrite (Stout 1972). The

many examples of anthophyllite that plot in Figure 4 in the

field of cummingtonite + gedrite (and the overlap of

anthophyllite and cummingtonite compositions) are interpreted

to represent late-stage growth of anthophyllite at lower than

maximum metamorphic temperatures, by one (or more) of the

following mechanisms; (1) grain-scale exsolution from gedrite;

(2) from the reaction gedrite + cummingtonite anthophyllite;

or (3) simple inversion from cummingtonite to anthophyllite.

Spear (1982) noted in some of his samples that anthophyllite

rims gedrite and was the last amphibole to grow. Also,

cummingtonite surrounded by a rim of anthophyllite has been

observed in many descriptions of amphibolites (Eskola 1914;

Tilley 1937; Tella and Eade 1978; Elliott-Meadows et al. 1999).

There are, in general, far too many reported examples of two

or three coexisting ferromagnesian amphiboles for them all to

be considered to represent equilibrium frozen at one set of

metamorphic conditions; collectively they represent a very

small portion of potential amphibole composition space

(Robinson et al. 1982). Because anthophyllite can form from

cummingtonite on cooling, Figure 4 should probably be used

to extract a maximum XFe for the phase boundary under the

average ofP-T conditions estimated for the five field areas.

Roughly half the samples plotted in Figure 4 were saturated in

Ca-amphibole. In these, the anthophyllites and cummingtonites

contain 23.5 and 36% Ca-amphibole (actinolite or horn-

blende) in solid solution, respectively.

In various kinds of metaperidotites, low-Ca amphibolites,

and metasomatic rocks, numerous analyses in the literature (e.g.,

Rabbitt 1948; Deer et al. 1997) show that low-Al anthophyllite

(


9/12


H2O and Tlc + Fo = 5 En + H2O. Although these results are not

definitive with respect to the relative stability of cummingtonite

and anthophyllite in the system MSH, they are not in conflict

with our field-based boundary. One must bear in mind, how-

ever, that the apparently orthorhombic Fe, Mg-amphiboles syn-

thesized by Cameron (1975) and Popp et al. (1976) have

compositions that place them (from the field evidence) clearly

in the T-XFe stability field of the monoclinic form. An occur-

rence of virtually end-member anthophyllite at the Wight mine,

Balmat, New York (sample 31, Klein 1968, with 0.15% F thiswork), believed to have formed at a maximum of 675 25 C

(E.J. Essene, personal communication), provides a low-tem-

perature field bracket on the end-member transition tempera-

ture.

The constraint on the phase loop supplied by amphibolites

(555 C, XFe 0.30) seems to agree better with the MFSH

loop than the CMFSH loop, even though hornblende is present

in many of the samples. An MSH transition set at 850 rather

than 800 C would fit the amphibolite data better, but the

metaperidotites less well. An overall flatter dT/dXFe slope for

the phase-loop might reasonably fit both the metaperidotites

and the amphibolites, but then the standard-state Gibbs free

energy of the exchange reaction, as we found by trial and error,would be too small to satisfy the natural KD data, because it

would shift the isotherms (Fig. 3) to more positive values of

lnKD.

Accepting our estimate of the transition temperature for the

Ca-free, Mg end-members (800 C, 5 kbar), the remaining un-

known, the Gibbs free energy of ferro-anthophyllite, can be

obtained from the standard-state Gibbs free energy of the ex-

change reaction (equivalent to the enthalpy of reaction), which

is tightly constrained by the natural KD data, in particular, those

from metaperidotites. This provides leverage on the dT/dXFe of

the phase loop and the transition temperature at the Fe-end. A

value of lnKD = 0.1 for coexisting anthophyllite and

cummingtonite in metaperidotites (XFe = 0.10.2) may be read

from the fit expression in Figure 3 of Evans and Ghiorso (1995),

Figure 3 here. This value of lnKD is independently supported

by a comparison of Fe-Mg partition in a large number of ferro-

magnesian amphibole + olivine pairs in metaperidotites (Fig.

6). Despite considerable scatter resulting mostly from a lack of

perfect exchange equilibrium, Figure 6 shows unmistakablythat, for a given olivine composition, cummingtonite is slightly

richer in Fe than anthophylliteby an amount corresponding

to an average lnKD of 0.1. It should be noted that our enthalpy

for end-member cummingtonite is 250 J/mol more negative

than adopted before (Table 7 in Ghiorso et al. 1995).

Our calculated binary two-phase loop for MFSH (Fig. 5)

has a temperature minimum close toXFe = 0.7 and an Fe end-

member transition temperature of 450 C, or about 35 C lower

in the presence of ferroactinolite. This result is consistent with

the common occurrence of cummingtonite-grunerite (ranging

inXFe from 0.4 to 0.98) in metamorphosed iron-formation (Klein

1982). The first occurrence of grunerite in the prograde meta-

morphism of iron-formation is close to the biotite isograd (Klein

1978; Haase 1982), that is, not much higher than 400 C. The

only anthophyllite reported in metamorphosed iron-formation

occurs in hematite-bearing rocks with high Mg/Fe2+ ratios

(Klein 1972). A flatter dT/dXFe slope would be in poorer agree-

ment with the widespread grunerite in BIF as well as the lnKDconstraints. In the Introduction, we guessed at a standard-state

exchange reaction free-energy of 0.4 kJ/M-cation. Our recom-

mended dataset (Table 5) has G1 = H1 = 1.75 kJ/mol or 0.25

kJ/M-cation, equivalent to a 350C difference in transition tem-

peratures from the Mg to the Fe end of the diagram.

CONCLUDINGREMARKS

We believe that our phase diagram (Fig. 5) successfully

accounts for the occurrence in various rock types of low-Al

anthophyllite, cummingtonite, and grunerite. In addition, the

diagram opens up the possibility of a stability field at low tem-

perature for pure ferro-anthophyllite. End-member

Fe7Si8O22(OH)2, usually somewhat manganoan, forms in na-

ture by cooling and hydration of fayalite, and has been described

FIGURE 5. Calculated TXFe phase loop at 5 kbar for coexisting

Al-free orthorhombic and monoclinic Fe, Mg-amphiboles (2 continuous

lines), and a phase loop (3 dashed lines) for three coexisting Al-freeamphiboles in CMFSH. Arrows denote constraints discussed in the

text.

FIGURE 6. Mg/Fe partition between natural cummingtonite,

anthophyllite, and olivine as a function of Fe/(Fe + Mg) in olivine.

Symbols as in Figure 1. KD = (Fe/Mg)0l/(Fe/Mg)Amph.


10/12


as grunerite (e.g., Bowen and Schairer 1935; Bonnichsen 1969;

Floran and Papike 1978; Vaniman et al. 1980; Janeczek 1989).

Recently, it has been shown that proto-ferroanthophyllite-Pnmn

(Sueno et al. 1998) and ferro-anthophyllite-Pnma (Bozhilov

and Evans 1999) are alternative possibilities. These observa-

tions are not fully understood at present (and MnO may play a

role in the occurrence of protoamphibole), but they certainlyhint at the possibility that an orthorhombic end-member is more

stable than grunerite at low temperature, but not so low a tem-

perature as to preclude growth of any Fe-amphibole. This lower

limit is about 340 C, because below this temperature we can

expect reaction of end-member Fe 7Si8O22(OH)2, whether

grunerite or ferroanthophyllite, to minnesotaite in the presence

of quartz and H2O (Miyano 1978; Rasmussen et al. 1998). Thus,

Fe end-member, Al-free orthorhombic amphibole may have a

stability field above 340 C and below 450 C. Probably the

best place to look for ferro-anthophyllite is in high Fe/Mg hy-

drothermal ore deposits; in low-grade metamorphosed BIF, the

required temperature may be too low and, furthermore, grunerite

will be favored in the presence of Ca.

Our reference-state Gibbs free-energies of formation for

monoclinic vs. orthorhombic FeMg-amphibole, not surpris-

ingly, are quite similar at 298 K and 1 bar, with GAthCum =

3.875 kJ/mol and GFathGru = 2.125 kJ/mol. When free ener-

gies are extracted from different sources along independent

paths, the differences, and the relative uncertainties, tend to be

considerably larger in the case of polymorphs. For example, the

corresponding reference-state transition free energies from Hol-

land and Powell (1998, Table 5) are 10.61 and 12.63 kJ/mol re-

spectively, and there is no stability field for either of the

monoclinic end-members.

None of the published papers on the subsolidus phase dia-

gram for MSH have considered the possibility of a stability

field for cummingtonite. Our data suggest that, close to its

breakdown to enstatite, quartz, and H2O (790 C at 5 kbar), the

stable amphibole may indeed be anthophyllite, but the uncer-

tainties are such that it could be cummingtonite. Additional

study of experimental run-products in the 800 C region in this

system, by HRTEM for example, would certainly be helpful.

In the presence of tremolite, the field of anthophyllite is suffi-

ciently reduced (Fig. 5) that it is unlikely to reach that of

enstatite, quartz, and H2O. In an attempt to check of our pro-

posed phase loop, D. M. Jenkins (personal communication,

October 1998) hydrothermally treated anthophyllite sample BM

93327 (XFe = 0.25, 12.5 wt% Al2O3) at 5 kbar. He found no

indication of conversion to cummingtonite at 700 C for 20

days, at 800 C for 12 days, and at 850 C for 5 days (at which

temperature partial breakdown to orthopyroxene took place).

The free-energy drive for the inversion reaction is so small that

direct experimental bracketing of the inversion curve will

clearly be extremely difficult. On the other hand, crystallization

experiments at high oxygen fugacity on hydrous, high Mg/Fe rhyo-

lite and dacite melts might be able to overcome the kinetic prob-

lem, in that ferromagnesian amphiboles grow readily from such

melts (e.g., Nicholls et al. 1992; Rutherford and Devine 1996;

Scaillet and Evans 1999).

Figure 4 shows the stabilizing effect of Al on anthophyllite,

but otherwise the diagram should only be regarded as a tenta-

tive illustration of the mutual relations of cummingtonite,

anthophyllite, and gedrite in the system NFMASH. For ex-

ample, it would be interesting to know if field evidence can be

brought to bear on the T-XFe behavior of the reaction terminal

to anthophyllite: Al-Ath = Ged + Cum.

The analogous inversion in Fe, Mg-pyroxene from

orthopyroxene to clinopyroxene shows the same behavior asthe ferromagnesian amphiboles in the T-XFe section, with a

minimum close to the Fe-end (Fig. 3 in Davidson et al. 1988,

Fig. 7 in Sack and Ghiorso 1994). The 350 C difference in the

transition temperature from the Mg to Fe end for Fe, Mg-am-

phibole may be normalized to the average temperature (in K):

350/898 = 0.390. This normalization yields a number that is

essentially identical to the corresponding value for pyroxene:

590/1478 = 0.399 (Fig. 7 in Sack and Ghiorso 1994). The cor-

respondence may be understood by multiplying both numera-

tor and denominator of this ratio by the appropriate averaged

heat capacity of the phase, i.e. (T) (CP,ave) / (Tave ) (CP,ave), which

may be written (H)/(Have). This demonstrates that the Tnor-

malized to the average temperature of the transition is a proxy

for the enthalpy difference between the Mg and Fe end-mem-

bers of the series, rendered as an intensive quantity through

division by the average enthalpy. The fact that structures and

transition mechanisms are similar between amphibole and py-

roxene suggests that normalized enthalpy differences between

Mg and Fe end-members of both series should compare favor-

ably. In fact, this calculation could have been used to predict

the transition temperatures in the amphibole series from knowl-

edge of those in the pyroxenes, and it could be argued that the

favorable correspondence is an independent test of the obser-

vations that led us to select the transition temperatures adopted

for the Fe, Mg-amphiboles.

ACKNOWLEDGMENTSWe thank J. Fabris, M. Guiraud, C. Klein, the Natural History Museum

London, J.S. Schneidermann, and R.J. Tracy for the loan of samples, M. Schindlerand F.C. Hawthorne for the refinements of sample BM 93327, and D.M. Jenkinsfor high PTwork on this sample. Critical reviews by G. Droop, J.C . Schumacher,and V. Trommsdorff are greatly appreciated. This study was supported by theNational Science Foundation, grant EAR 97-06326.

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Thermodynamics of the Am Phi Boles - Anthophyllite-Ferroanthophyllite and the Ortho-clino Phase Loop

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