Journal of Petrology Advance Access published online on February 22, 2007
Journal of Petrology, doi:10.1093/petrology/egm001
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Liquidus Equilibria in the System K2O–Na2O–Al2O3–SiO2–F2O–1–H2O to 100 MPa: I. Silicate–Fluoride Liquid Immiscibility in Anhydrous Systems
*Department of Earth and Planetary Sciences, Mcgill University, Montreal, QC H3A 2A7, Canada
Received October 7, 2005; Revised typescript accepted January 9, 2007
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ABSTRACT |
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TOP
ABSTRACT
INTRODUCTIONLiquidus relations in the four-component system Na2O–Al2O3–SiO2–F2O–1 were studied at 0·1 and 100 MPa to define the location of fluoride–silicate liquid immiscibility and outline differentiation paths of fluorine-bearing silicic magmas. The fluoride–silicate liquid immiscibility spans the silica–albite–cryolite and silica–topaz–cryolite ternaries and the haplogranite-cryolite binary at greater than 960°C and 0·1–100 MPa. With increasing Al2O3 in the system and increasing aluminum/alkali cation ratio, the two-liquid gap contracts and migrates from the silica liquidus to the cryolite liquidus. The gap does not extend to subaluminous and peraluminous melt compositions. For all alkali feldspar–quartz-bearing systems, the miscibility gap remains located on the cryolite liquidus and is thus inaccessible to differentiating granitic and rhyolitic melts. In peralkaline systems, the magmatic differentiation is terminated at the albite–quartz–cryolite eutectic at
770°C, 100 MPa, 5 wt % F and cation Al/Na = 0·75. The addition of topaz, however, significantly lowers melting temperatures and allows strong fluorine enrichment in subaluminous compositions. At 100 MPa, the binary topaz–cryolite eutectic is located at 770°C, 39 wt % F, cation Al/Na 0·95, and the ternary quartz–topaz–cryolite eutectic is found at 740°C, 32 wt % F, 30 wt % SiO2 and cation Al/Na 0·95. Such location of both eutectics enables fractionation paths of subaluminous quartz-saturated systems to produce fluorine-rich, SiO2-depleted and nepheline-normative residual liquids. KEY WORDS: silicate melt; granite; rhyolite; fluorine; liquid immiscibility
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INTRODUCTION |
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Fluorine is the most abundant and compatible volatile element in highly evolved granitic and rhyolitic magmas (Webster, 1990
; London, 1997; Webster et al., 1997; Thomas & Webster, 2000; Thomas et al., 2005). The average fluorine concentrations in natural silicic suites increase from 0·11 wt % F in biotite and two-mica granites through 0·99 wt % F in topaz granites, rhyolites and ongonites (e.g. Kovalenko & Kovalenko, 1976; temprok, 1991; Dergachev, 1992) to 3·9 wt % F in quartz topazites (e.g. Eadington & Nashar, 1978; Kortemeier & Burt, 1988; Zhu & Liu, 1990; Johnston & Chappell, 1992; Antipin et al., 1999). This variability of fluorine abundances by nearly two orders of magnitude requires extreme levels of crystal–liquid fractionation (99% solidified), if that process alone is responsible for the generation of the F-rich melts.
An alternative mechanism for generating high fluorine concentrations in residual magmatic liquids is provided by fluoride–silicate liquid–liquid immiscibility (e.g. Gramenitskiy & Shchekina, 1994; Veksler, 2004; Veksler et al., 2005). Previous experimental studies, however, led to contradictory results concerning the presence and location of fluoride–silicate miscibility gaps (Table 1). These disagreements are found in simple systems, e.g. albite–NaF (Koster van Groos & Wyllie, 1968; Rutlin, 1998), and in multicomponent granitic systems (Kovalenko et al., 1975; Glyuk & Trufanova, 1977; Kovalenko, 1977; Wyllie, 1979; Danckwerth, 1981; Webster et al., 1987; Gramenitskiy & Shchekina, 1994; Xiong et al., 2002). These controversies are likely to result from: (1) misinterpretation of round fluoride crystals as immiscible globules (Gramenitskiy & Shchekina, 1994; Koreneva & Zaraiskiy, 2001); (2) misinterpretation of stable or quench fluid inclusions with high solute content as immiscible liquids (microliquation; e.g. Anfilogov et al., 1973; Gluyk & Anfilogov, 1973a, 1973b); (3) use of variable fluid or melt proportions during experiments, which causes significant departures from initial rock (melt) composition as a result of the fluid–melt partitioning; (4) mass loss and shift in melt composition as a result of fluoride vaporization in 1 atm experiments (Kogarko & Krigman, 1975; Siljan, 1990); (5) inappropriate choice of fluoride additives. The use of HF, alkali fluorides, NaF + AlF3 mixtures or fluoride minerals produces very distinct compositional effects and such silicate–fluoride sections diverge from liquid lines of descent or intersect the Alkemade compatibilities; that is, they may penetrate potential thermal barriers.
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Figure 1 illustrates mineral compatibilities between aluminosilicates and fluorides in the system (Na2O + K2O)–Al2O3–SiO2–F2O–1 at 700°C and 100 MPa (Dolej
& Baker, 2004). Increasing the chemical potential of F2O–1 is equivalent to adding fluorine in the form of HF, which does not affect the major-oxide composition of the system. The phase diagram intersects numerous low- and high-fluorination assemblages (Anovitz et al., 1987; Dolej & Baker, 2004). However, only the low-fluorination assemblages are stable with alkali feldspar and therefore accessible to natural granitic magmas. Fluorination by adding alkali fluorides (NaF, KF) intersects the quartz–cryolite (elpasolite) tie-line where alkali feldspar becomes unstable and the resulting strongly peralkaline compositions are not representative of natural magmas. Addition of NaF and AlF3 mixtures to balance the alumina–alkali ratio causes departure towards SiO2-poor (feldspathoidal) compositions and promotes the metastable absence of topaz. Thus, understanding chemical and mineral compatibilities is essential in experimental design and to closely approach the differentiation paths of silicic magmas.
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In this study, we present liquidus equilibria in several subsystems of the four-component space Na2O–Al2O3–SiO2–F2O–1 to locate the fluoride–silicate miscibility gap and its relation to the liquid lines of descent of granitic and rhyolitic magmas. First, we calculate stabilities of fluoride phases in the haplogranite system during progressive fluorination and predict saturating solid phases. Second, we experimentally investigate several binary, ternary and quaternary systems that correspond to stable silicate–fluoride assemblages. The silica–cryolite binary is expected to intersect the field of fluoride–silicate liquid immiscibility and we trace its extension into the albite–quartz–cryolite and quartz–topaz–cryolite ternaries. The quartz–topaz–cryolite assemblage represents the products of albite fluorination, and this ternary system will be tested for peritectic relationships. We have performed experiments at 1 atm and at 100 MPa because this pressure range represents the typical emplacement levels of topaz granites, subvolcanic ongonite dykes and topaz rhyolites (Christiansen et al., 1986
; temprok, 1991; Cuney et al., 1992; Price et al., 1992; Thomas & Klemm, 1997). In the second part of this study (Dolej & Baker, 2007), we report results on the hydrous albite–quartz and haplogranite systems with topaz and cryolite, which illustrate differentiation paths of silicic magmas and provide information on the maximum solubilities of fluorine and H2O. These experimental results provide a complete framework for phase equilibria in fluorine-bearing, Li-, Ca- and Fe-poor, silicic magmas that range from peralkaline to peraluminous compositions. |
COMPOSITION SPACE AND PHASE COMPATIBILITIES |
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Felsic igneous rocks span the composition space Na2O–K2O–Al2O3–SiO2 (+ H2O). Individual phases are defined in Table 2. Fluorine is a fifth component and it is a monovalent anion in all solid, liquid and gaseous phases in this study. The corresponding chemical component must contain fluorine in the same valence state; suitable options are NaF, KF, AlF3, SiF4 or F2O–1. Although use of the fluorine molecule, F2, may appear to be obvious, we shall demonstrate below that because its valence state (0) differs from that of fluorine in all geological phases (–1) it is a poor choice.
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To illustrate various choices of components, we will consider three-phase equilibrium of quartz, andalusite and fluortopaz. This assemblage contains three chemical components and one possible set is Al2O3, SiO2 and SiF4. The equilibrium between the three phases is given by equation 1 in Table 3. Another valid set of chemical components is Al2O3, SiO2 and F2O–1, and in this case the equilibrium has a simpler form (equation 2, Table 3). This is because the Al:Si ratio in andalusite and fluortopaz is the same and F2O–1 does not contain any cation. The use of F2O–1 does not change the proportions of cations in the multicomponent systems and it leads to simpler equilibrium expressions. A third choice of chemical components for our system is to split the exchange operator F2O–1 into two components, F2 and O2. This may appear intuitive but it has disadvantages because it introduces an additional component and the quartz–andalusite–fluortopaz assemblage must be described by Al2O3, SiO2, F2 and O2 (see equation 3, Table 3). Furthermore, the choice of F2 as a component requires that the oxygen fugacity of the system be defined. Oxygen is only needed to balance the difference between F– (in topaz) and F0 (in the F2 component) but it is not required for any mineral–mineral equilibrium. When a component with a fluoride anion such as SiF4 or F2O–1 is chosen the oxygen-fugacity constraint is eliminated. We prefer to use the F2O–1 component (Burt, 1972
, 1975; Burt & London, 1982), an exchange operator that corresponds to a replacement of one oxygen by two fluorine anions, because this component maintains charge balance and keeps ratios of all cations constant (Dolej & Baker, 2004). The use of the exchange operator F2O–1 is thermodynamically valid (Burt, 1972, 1975), results in a correct number of system components and does not produce any loss of generality.
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The anion exchange operator, F2O–1, is used here to describe fluorination of oxides and silicates into fluorides or topaz (Table 3). In H2O-bearing systems, increases in F2O–1 can also be visualized as addition of HF (see equation 4, Table 3). Also, HF and F2O–1 projection points are identical (Fig. 1). In anhydrous systems, F2O–1 is a hypothetical composition and this apex is physically inaccessible (upper part of Fig. 1).
To portray equilibria in the five-component system Na2O–K2O–Al2O3–SiO2–F2O–1 we use the reduced tetrahedron (Na2O + K2O)–Al2O3–SiO2–F2O–1 (Fig. 1). This pseudoquaternary system contains important rock-forming minerals (quartz, micas, feldspars, feldspathoids) as well as fluorine-bearing phases (topaz, villaumite, cryolite, chiolite, malladrite, etc.) and the composition space includes both peralkaline and peraluminous compositions. Projection points of fluorides are obtained by combining oxides with F2O–1 in stoichiometric proportions (equations 4–9, Table 3). Various trends of fluorination (i.e. addition of HF, NaF, AlF3, etc.) are easily illustrated with respect to phase locations and compatibilities. A complete analysis of fluorination reactions, silicate–fluoride equilibria and mineral compatibilities has been presented by Dolej & Baker, (2004).
The formation of fluorine-bearing minerals in granitic systems corresponds to progressive fluorination of rock-forming silicates, defined by increasing chemical potential of F2O–1. With increasing µ(F2O–1), rock-forming minerals are converted to topaz, cryolite, chiolite and other fluoride minerals or gases (Fig. 2; equations 2 and 10–17, Table 3). In granitic rocks, fluorine concentrations are buffered by quartz, feldspar, topaz and cryolite (equation 10, Table 3) which corresponds to the first fluorination equilibrium in Fig. 2. The formation of the ‘high-fluorination’ phases (e.g. chiolite, malladrite) occurs after complete breakdown of feldspars. During magmatic crystallization, the accessibility of these ‘high-fluorination’ assemblages depends upon whether boundaries between specific phase assemblages (Fig. 2) become thermal barriers or peritectic transitions. If they form thermal barriers, magmatic differentiation will be restricted to relatively low fluorine contents, buffered by topaz and/or cryolite in the presence of quartz and feldspar (compare topaz granites, cryolite granites), whereas if they become peritectic transitions, residual magmas will achieve very high fluorine contents and become feldspar-absent (compare quartz topazites). To know which topology is relevant, we must determine liquidus relations in degenerate sections (e.g. the cryolite–topaz–quartz ternary, Fig. 1) and test for the presence of pseudoternary phases (e.g. albite).
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FLUORIDE VAPORIZATION |
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During experimental studies at high concentrations of fluorine, vapor pressures of aluminofluorides and silicofluorides increase, these species variably vaporize, and the loss of elements from high-temperature fluorosilicate liquids causes departures from the initial bulk composition (e.g. Snow & Welch, 1972
; Siljan, 1990; Pruttskov & Krivoruchko, 1997). To approach this problem, we have evaluated vapor pressures in equilibrium with relevant solid and liquid phases. We use several examples in the system albite–quartz–cryolite–topaz to illustrate the necessary thermodynamic approach to vapor–solid and vapor–liquid fluorosilicate systems.
In each compatibility subtetrahedron (Fig. 1), the four coexisting phases at the pressure and temperature of interest uniquely define chemical potentials of four independent components. For example, the coexistence of albite, quartz, cryolite and topaz determines the chemical potentials of Na2O, Al2O3, SiO2 and F2O–1, and their values are found by solving sets of linear equations containing the Gibbs free energies of the stable phases at the pressure and temperature of interest (equations 18–21, Table 3; Korzhinskii, 1959; Connolly, 1990; Dolej & Baker, 2004). The chemical potentials of the independent components define the chemical potential (or fugacity) of any gaseous species of interest (equations 5–9, Table 3). The calculated fugacities of the most abundant fluoride species are plotted in Fig. 3. Over the temperature range of 600–1200°C, the abundance of individual fluoride compounds varies as follows: SiF4 > NaAlF4 > AlF3 > NaF > (NaF)2. Individual fugacities (vapor pressures) differ by several orders of magnitude and SiF4 closely defines the total vapor pressure. The gas fugacities increase with increasing temperature. At low temperatures, the total vapor pressure is less than 1 atm and gaseous fluorides will constitute only a small fraction of the space in the capsule during the experiments. At high temperatures, the experimental design requires confining pressure. Our calculations demonstrate that an applied pressure of 100 MPa in our experimental procedure completely eliminates problems with fluoride vaporization (Fig. 4).
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In contrast to stoichiometric solid phases, the fluorosilicate melt composition is not constant and may continuously change during fluoride vaporization. We will illustrate this for melts along the silica–cryolite binary, which were studied previously (Weill & Fyfe, 1964
; Kogarko & Krigman, 1975). Because SiF4 is the predominant gaseous species (Kogarko et al., 1968; Snow & Welch, 1972; Siljan, 1990; Dolej & Baker, 2004; Fig. 3), the melt composition deviates from the quartz–cryolite binary system to the ternary space where the third component is SiF4. The removal of SiF4 from the SiO2–Na3AlF6 binary shifts the melt composition within the SiO2–Na3AlF6–SiF4 plane away from the SiF4 coordinate. That is, the melt composition leaves the SiO2–Na3AlF6 tie-line in the Na2O–Al2O3–SiO2–F2O–1 tetrahedron and enters one of the quaternary compatibility tetrahedra. By using the thermodynamic databases of Holland & Powell (1998) and Dolej & Baker, (2004), the devolatilization equilibria are expressed by equation 22 (at <896°C and 1 atm) or equation 23 (at >896°C and 1 atm), respectively, in Table 3.
The quaternary assemblage is temperature-dependent, and these solid phases represent the solidus assemblage after crystallization of the degassed melt. These phases were observed in experimental run products (e.g. Snow & Welch, 1972; Pruttskov et al., 1989) thus confirming the selective loss of fluorides to the vapor phase. To calculate fugacities of all fluoride species, the melt thermodynamics must be described by a non-ideal quaternary mixing model. We have calibrated a simple model for the NaAlSi3O8–SiO2–Na2Si2O5–Na3AlF6 melt, by expanding the formulation of Holland & Powell (2001) and including experimental data in Na2Si2O5- and Na3AlF6-bearing subsystems (Morey & Bowen, 1924; Kracek, 1930; Rutlin, 1998). The binary, asymmetric Margules parameters have the following values in the eight-anion formulation (Burnham, 1997; Holland & Powell, 2001): WSi4 O8 – NaAlSi3 O8 = WNaAlSi3 O8 – Si4 O8 = 12 kJ, WSi4 O8 – Na3·2 Si3·2 O8 = WNa3·2 Si 3·2 O8 – Si4 O8 = 3 kJ, WSi4O8 – Na4 Al1·33 F8 = 20 kJ, WNa4Al1·33 F8} – Si4 O8 = 85 kJ, WNaAlSi 3 O 8 – Na 4 Al1·33F 8= 30 kJ, WNa4 Al1·33 F 8 – NaAlSi 3 O8 = – 33 kJ; all remaining interaction terms are zero. At the temperature and pressure of interest, the degree of SiF4 depletion in the melt is expressed by the extent of the devolatilization reaction (equation 22 or 23, Table 3), and the resulting melt composition is recast into four components (Si4O8, NaAlSi3O8, Na2Si2O5 and Na3AlF6) whose activities are now defined by the non-ideal mixing model. Chemical potentials of the four melt components uniquely define the chemical potentials of the system components (Na2O, Al2O3, SiO2, F2O–1) by linear-algebraic manipulation, and the fugacity of any fluoride gaseous species, coexisting with the fluorosilicate melt, is determined as above (equations 5–9, Table 3). We have portrayed the vapor pressures of gaseous species and compositional departures for various extents of devolatilization of the cryolite–silica eutectic melt (95 wt % Na3AlF6 and 5 wt % SiO2) at 1050°C in Fig. 4.
Total vapor pressure at 1050°C ranges over several orders of magnitude and SiF4 remains the predominant gaseous species. Both the total vapor pressure and the SiF4 fugacity decrease with increasing extent of devolatilization (equation 23, Table 3) whereas fugacities of NaAlF4, NaF, (NaF)2 and AlF3 change only insignificantly (Fig. 4a). The chemical potentials of system components (Na2O, Al2O3, SiO2, F2O–1) vary only negligibly over the small extent of the devolatilization equilibrium, x = 10–6 to 10–2 (Fig. 4a). This means the fugacities of alkali fluoride and aluminofluoride species are relatively insensitive to x. In contrast, SiF4 occurs in the devolatilization equilibrium (equation 23, Table 3) and is subject to the law of mass action, which includes other liquid species. As SiO2 and Na3AlF6 react to produce SiF4, the mole fractions (and activities) of NaAlSi3O8 and Na2Si2O5 increase in the melt but are initially very small numbers. As the equilibrium constant for equation 23 (Table 3), K = f(SiF4)3.a(NaAlSi3O8)2.a(Na2Si2O5)2/[a(Na3AlF6)2.a(SiO2)13], is a function of pressure and temperature only, small activities of NaAlSi3O8 and Na2Si2O5 must be accommodated by high activity (fugacity) of SiF4. Therefore, the vapor pressure is very high at initial stages of devolatilization but decreases as the equilibrium (equation 23, Table 3) progresses to the right. It is also noteworthy that the relative abundance of gaseous species is rather insensitive to the bulk SiO2–Na3AlF6 proportions. According to calculations, SiF4 remains the predominant gaseous species with increasing SiO2 content up to tridymite saturation (see also Fig. 3).
Our thermodynamic calculations can also serve to interpret vapor-phase saturation in experiments. The intersection of the total vapor pressure and the experimental pressure (Fig. 4) defines a phase boundary, L (V), i.e. an equilibrium extent of devolatilization at the pressure of interest. When the total vapor pressure is lower than the confining pressure, the assemblage is vapor-undersaturated, i.e. in the L field. When the total vapor pressure is higher, a free vapor phase coexists with the liquid. The degrees of devolatilization at 100 MPa and 1 atm are shown in Fig. 4a and the corresponding mass changes in the fluorosilicate liquid in Fig. 4b. At 100 MPa, the extent of devolatilization is very small, and the changes in element concentrations in the melt are negligible (less than 0·05% relative). At 1 atm, the extent of devolatilization is greater, with weak SiO2 (4·6%) and F (0·15%) depletions and, consequently, Na2O and Al2O3 enrichments (0·35% each).
This approach provides estimates of the compositional shifts in the liquid. In our study, we have limited all experiments to a maximum of 1100°C, as the magnitude of vaporization rises with temperature (Fig. 3). In addition, experiments were frequently performed at both 0·1 MPa and 100 MPa. Comparison of run products at the two pressures revealed no discrepancies.
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EXPERIMENTAL AND ANALYTICAL METHODS |
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The silicate–fluoride melting equilibria were studied by quenching techniques and differential thermal analysis. Starting materials were synthetic glasses and pure natural minerals. Base glasses (HPG-2, AQ-1) represent compositions of the haplogranite minimum and the albite–quartz eutectic, respectively, at 100 MPa and H2O saturation (Tuttle & Bowen, 1958
). These were prepared by weighing of reagent-grade K2CO3, Na2CO3, Al2O3 and SiO2 in desired proportions into an agate mortar, followed by grinding for 1 h under alcohol or acetone. The slurry was dried overnight and transferred to the platinum crucible. The mixture was decarbonated at a heating rate of 150°C/h and held at 1020°C for 8 h. Melting was carried out in several cycles (1 or 2 h) at 1400–1600°C with intermittent crushing. After each cycle, an aliquot of glass was analyzed by electron microprobe to verify alkali loss and to monitor compositional homogeneity. After the last melting, crushed chips were ground in an agate mortar for 1 h (dry) and the resulting glass powders were stored at 120°C until use. Starting crystalline phases were a mixture of quartz and tridymite (99·99 wt % SiO2), natural albite (Amelia Court House, Virginia), natural topaz (Topaz Mountain, Utah), and natural cryolite (99·5%, Alfa Aesar). All substances were ground in an agate mortar (dry) and stored at 120°C. Compositions of all starting materials are given in Table 4. Starting mixes were prepared by weighing the base glasses and crystalline phases into an agate mortar and grinding for 1 h (dry). The as-weighed compositions are listed in Table 5 and are accurate to 0·1 wt %.
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Experiments were performed in gold or platinum capsules depending upon the pressure–temperature conditions. Seamless tubing (Au: 2·0–2·2 mm OD; Pt: 3·0–4·0 mm OD) was cut into segments 10 mm long, cleaned in concentrated hydrofluoric acid, repeatedly rinsed with distilled water, ultrasonically cleaned with alcohol for 4 min, and annealed over the gas burner to yellow–orange heat. Capsules were flat-welded and loaded with starting materials (Au: 8–11 mg; Pt: 20–35 mg). For 1 atm experiments, the crimped capsules were stored at 300°C for 1 h to remove traces of moisture and welded immediately; the weight loss after heating was less than 0·3%.
Experiments were performed in tube furnaces (1 atm), cold-seal pressure vessels (up to 850°C, 100 MPa) or rapid-quench TZM pressure vessels (above 850°C, 100 MPa) with argon as a pressure medium. Temperatures were monitored by sheathed chromel–alumel thermocouples, calibrated against the melting point of NaCl (800·7°C, Dawson et al., 1963; Chase, 1998) and verified with a factory-calibrated thermocouple. Temperatures are accurate to ±2°C (tube furnaces and cold-seal vessels) and ±5°C (TZM vessels). Pressure was measured with Bourdon-tube gauges, calibrated against a factory-calibrated Heise gauge. Pressure measurements are accurate to ±5 MPa and precise to ±2 MPa. Experiments were quenched by either dropping the capsules from the 1 atm furnace into a cold-water bath (500°C/s), placing the cold-seal vessel into an air jet (150°C/min), or by free fall of the sample holder to the cooling collar in the TZM vessel (100°C/s). All capsules were checked for leakage, and charges were recovered immediately and stored at room conditions. Run products were studied optically in grain mounts and by electron microprobe; several chips from the same run were studied to avoid misinterpretation as a result of crystal settling. Large chips (0·5–1·5 mm) were mounted in epoxy, ground and polished in alcohol–oil mixtures, and carbon-coated for electron-microprobe analysis. All phase assemblages were verified by electron microprobe. Details of differential thermal analysis have been given by Dolej & Baker (2006).
The time spans necessary to achieve equilibrium vary over several orders of magnitude in systems ranging from molten ionic salts (Na3AlF6) to fully polymerized silicates (NaAlSi3O8, SiO2; Fig. 5). For cryolite, melting and crystallization occur within 6 and 2 min, respectively, as measured by differential thermal analysis. Experiments in the Na3AlF6–SiO2 system, including quartz dissolution and tridymite growth, are reversible within less than 60 min (Fig. 6a). In cryolite- and topaz-bearing systems, the formation of quench phases confirms that crystal nucleation has not been suppressed. Experiments across the silicate–fluoride join were usually carried out together to ensure identical run conditions, and consistency of their results from the silicate towards the fluoride end-member was verified. In the albite–quartz joins with fluorine-bearing minerals, different starting materials (aluminosilicate glasses vs crystalline albite and SiO2) were used simultaneously to verify the invariance of experimental run products. Run durations in excess of 7 days were necessary in anhydrous systems near the SiO2, albite–SiO2 or haplogranite end-member. The lack of equilibrium is demonstrated by the presence of incompletely reacted starting materials or inhomogeneous melt compositions (measured by electron microprobe). These runs were not considered further and the compositional regions with disequilibrium were indicated in run tables and on phase diagrams by cross symbols without outline.
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RUN PRODUCTS |
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We investigated a wide range of silicate to fluoride systems at temperatures between 520 and 1100°C. The overall appearance of the run products reflects nucleation and annealing rates related to individual phases, and growth rates that correlate with temperature. Run products contain stable equilibrium assemblages, but quench phases systematically occur in fluorine-rich compositions. Cryolite-rich or immiscible fluoride liquids did not form glasses and quenched to aggregates of solid phases (cryolite, sodium aluminosilicates; Fig. 6b). The proportion of quench aluminosilicates allows approximate estimation of the silicate components in the fluoride melt, which do not exceed
5 vol. %. On the other hand, fluorine-poor liquids quench to homogeneous glasses. At 10 wt % F in the melt, quench cryolite starts to appear in the form of branching agregates, elongate dendritic rods (up to 200 µm long, Fig. 6c) or micrographic intergrowths with aluminosilicate glass (10–80 µm). Above 25 wt % F, liquids quench to an inhomogeneous aggregate of cryolite grains, aluminosilicate phases and glass. Stable cryolite forms large round grains (10–70 µm) at all conditions (Fig. 6d and e) and can easily be distinguished from dendritic or micrographic cryolite that formed during quenching. These large cryolite grains are optically isotropic, single crystals with a stoichiometric composition and no quench-related inclusions. In several composition sections, experiments were performed slightly above the cryolite melting temperature to check for the different appearance of the liquid–liquid immiscibility (Fig. 6b).
Topaz develops subhedral crystal shapes and has a uniform grain size (10–30 µm); incomplete dissolution of topaz is revealed by angular morphologies and irregular fragment sizes. Alkali feldspar occurs as minute anhedral grains (10 µm) and its morphology tends to evolve to elongate laths (20–40 µm long) with increasing alkali or fluorine contents. The morphology of the silica polymorphs (quartz and tridymite) and their grain size are closely related to run temperature. Tridymite forms large subhedral grains with planar crystal faces (5–40 µm; Fig. 6e and f) or thin laths (30–90 µm long). In contrast, quench silica polymorphs in cryolite-rich melts form minute round grains (2–5 µm). Stable quartz occurs as subhedral prismatic grains (10 µm) and it becomes round and fine-grained (less than 5 µm) with decreasing temperature. Failure to attain equilibrium in fluorine-poor and silica-rich compositions can be identified by the presence of fragments of starting phases.
In all joins, the geologically relevant liquid lines of descent evolve to fluorine-rich residual melt compositions. The presence of quench phases, the decreasing amount of interstitial melt and the small grain size at low temperatures severely limit the systematic use of electron microprobe techniques for measuring the compositions of volatile-rich melts. Therefore, we have systematically studied binary and ternary sections and derived phase-diagram topologies by chemography.
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THE SILICA–CRYOLITE SYSTEM |
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Liquidus relations in the SiO2–Na3AlF6 system are the starting point for the general topology of silicate–fluoride systems (Table 6) and have additional applications to the study of cryolite attack on SiO2-based refractories during aluminum electrolysis (Snow & Welch, 1972
; Siljan et al., 2001). In previous studies, the location of the eutectic and the presence of the fluoride–silicate liquid–liquid immiscibility have been controversial (Weill & Fyfe, 1964; Grjotheim et al., 1971; Kogarko & Krigman, 1975, 1981). This binary system is characterized by a large stability field of silica polymorphs (quartz, tridymite; Fig. 7). This has been verified by a reversed run (number 341; Fig. 6a) with quartz single crystal surrounded by cryolite powder (50 wt % each). The run was placed at 1020°C and 1 atm for 1 h, then reversed to 935°C, kept at this temperature for 24 h and quenched. Its texture reveals the initial dissolution of quartz in the cryolite melt, followed by tridymite growth during cooling. These observations confirm not only the rapid approach to equilibrium but also provide unambiguous evidence for its location in the SiO2 + liquid field, which was hitherto unclear. The binary eutectic is located at 95 wt % cryolite, at 999°C (1 atm) and 1015°C (100 MPa), determined by differential thermal analysis.
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The asymmetric location of the binary eutectic can be explained by the disparate melting temperatures of silica and cryolite, augmented by strong positive deviations from ideality as a result of coordination differences (Dolej
& Baker, 2005). The location of the cryolite–silica eutectic on the cryolite side of the phase diagram is in contrast to other cryolite–silicate systems (Rutlin, 1998; Rutlin & Grande, 1999). The present experiments did not intersect the fluoride–silicate miscibility gap below 1100°C (Fig. 7), but the liquid–liquid immiscibility in the SiO2–Na3AlF6 binary has been documented at 1200°C and 1 atm (Kogarko & Krigman, 1975) and at 1200°C and 600 MPa (G. Robert & D. Dolej, unpublished experimental results). |
THE SILICA–ALBITE–CRYOLITE SYSTEM |
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The silica–albite–cryolite system serves as a model for peralkaline fluorine-bearing granites and rhyolites. The albite–cryolite binary has a very similar topology to the silica–cryolite binary, but the eutectic position is reversed to low fluorine contents (Rutlin, 1998
; Siljan et al., 2001).
We studied this ternary system along two parallel sections with 15 and 60 wt % cryolite, respectively, each from the SiO2 to the albite side of the ternary and spaced to provide additional joins from cryolite to Qz80Ab20 and Qz41Ab59 by weight, respectively (Fig. 8, Table 4); the Qz41Ab59 composition represents the albite–quartz eutectic at 100 MPa and H2O saturation (Tuttle & Bowen, 1958). The polybaric temperature–composition sections at 15 wt % and 60 wt % cryolite, respectively (Fig. 9a and b), are characterized by fluoride–silicate liquid–liquid immiscibility above 970°C (1 atm and 100 MPa). The miscibility gap is underlain by fields containing cryolite, silica polymorphs and/or albite, which coexist with fluorine-poor silicate liquid.
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The fluoride–silicate miscibility gap extends over the entire ternary system. It is located on silica liquidus in the SiO2-rich portion of the ternary, but it overlies the cryolite liquidus in the albite-rich part. Consequently, the location of binary eutectics shifts from the high-fluorine contents (95 wt % cry + 5 wt % SiO2) to low fluorine contents (<15 wt % cry + >85 wt % ab). Hence, finding the ternary albite–quartz–cryolite eutectic is fundamentally important for the maximum solubility of fluorine in peralkaline granitic melts and the accessibility of the fluoride–silicate liquid–liquid immiscibility to natural melts. In Fig. 9a, the down-temperature sequence of the stability fields in the vicinity of Ab59Qz41 is: Lsil, Lsil + cry and Lsil + cry + SiO2. That is, liquid with 15 wt % Na3AlF6 will first saturate with cryolite, and subsequent crystallization will drive liquid compositions away from the cryolite apex towards lower fluorine contents. All lower-temperature features, the SiO2–cry cotectic and the SiO2–ab–cry eutectic, must lie at less than 15 wt % cryolite.
The experimental results in the 15%-cryolite section allow for two distinct phase-diagram topologies, depending on the location of the qz/trd–cry–Lsil (Lfl) invariant point with respect to the investigated section (Fig. 9a). If the invariant point occurs at greater than 15 wt % cryolite, the sequence with increasing albite content of high-temperature fields will be SiO2 + Lfl, SiO2 + Lfl + Lsil, SiO2 + Lsil and Lsil; the field Lfl + Lsil will occur at very high temperatures only and the field cry + Lfl + Lsil will not exist. If the invariant point occurs at less than 15 wt % cryolite, the sequence of phase fields will be SiO2 + Lfl, SiO2 + Lfl + Lsil, Lfl + Lsil and Lsil (Fig. 9a). The second interpretation is preferred, as other invariant points, albite–quartz–cryolite and albite–cryolite eutectics, occur at less than 15 wt % cryolite as well. We emphasize that this choice does not affect the overall topology of the ternary system as it only concerns the relative position of an invariant point with respect to the composition section chosen for experimental work.
The liquidus relations in the SiO2–NaAlSi3O8–Na3AlF6 ternary are summarized in Fig. 10. The silica–cryolite cotectic represents a divide (thermal minimum) for the fluoride–silicate miscibility gap at 970°C (1 atm and 100 MPa). The ternary quartz–albite–cryolite eutectic occurs at 5 wt % F and 770°C (100 MPa). Peralkaline fluorine-bearing silicic magmas will fractionate along the quartz–albite cotectic with progressively increasing fluorine content in the melt until they reach the quartz–albite–cryolite eutectic and completely crystallize. As the fluoride–silicate miscibility gap is located on the cryolite liquidus in this region, the liquid–liquid immiscibility is inaccessible to peralkaline quartz–feldspar-precipitating melts.
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, locations of starting compositions. Concentrations of SiO2 and F are shown on the sides. |
THE SILICA–TOPAZ–CRYOLITE SYSTEM |
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Topaz is a common mineral in peraluminous fluorine-bearing granites, rhyolites and their differentiation products (ongonites and quartz topazites). From the thermodynamic viewpoint, topaz and cryolite are products of the first fluorination steps of feldspars (Fig. 2; equation 10, Table 3). Therefore, these minerals are potential fluorine buffers during crystallization of granitic and rhyolitic rocks. The SiO2–Al2SiO4F2–Na3AlF6 system covers a wide range of peralkaline to peraluminous compositions and is pierced by the albite–F2O–1 and nepheline–F2O–1 joins (Fig. 11). If a stability field of albite and/or nepheline was found in the silica–topaz–cryolite ternary, this would indicate the existence of the peritectic transition: alkali feldspar/feldspathoid + liquid
cryolite/chiolite + topaz + quartz. Such a peritectic point would allow for very high fluorine enrichments in residual melts and the formation of feldspar-free eutectic assemblages (compare quartz topazites).
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We studied several composition sections (isopleths) in this ternary (Fig. 11; Tables 7 and 8); the subaluminous SiO2–Tp47Cry53 section divides the peralkaline and peraluminous space, respectively, and intersects the other three joins at the following compositions: albite–F2O–1 (TCQ-1), nepheline–F2O–1 (TCQ-3) and cryolite–topaz (cation Al/Na = 1; TC-1).
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The phase diagram for the cryolite–topaz binary at 100 MPa is presented in Fig. 12. This system is characterized by simple binary behavior with a eutectic at 770°C (39·2 wt % F, cation Al/Na
0·95). In the topaz–cryolite–silica ternary, the temperature–composition sections cryolite–(SiO2)73Tp27 and silica–topaz with 60 wt % cryolite (Fig. 13) both intersect the field of the silicate–fluoride liquid–liquid immiscibility, which extends from the SiO2–Na3AlF6 binary. The two-liquid field is underlain by the ternary assemblage SiO2 + Lsil + Lfl and is located exclusively on the tridymite/quartz liquidus. The miscibility gap closes at 960°C, and it is not intersected by the silica–cryolite or silica–topaz cotectic curves. The large temperature interval (1000–740°C), occupied by silica- and/or cryolite-bearing fields (Figs 13 and 14) indicates prolonged existence of liquids migrating from high-temperature binary eutectics (silica–cryolite and silica–topaz, Fig. 7, Table 8) towards the ternary quartz–topaz–cryolite eutectic.
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This ternary quartz–topaz–cryolite system does not contain any additional (pseudoternary) phases (e.g. albite or nepheline). Thus the silica–topaz–cryolite system represents a thermal barrier in the Na2O–Al2O3–SiO2–F2O–1 composition space, and the stability fields of the ‘high-fluorination’ phases (chiolite, maladrite) are not accessible through magmatic fractionation.
Changes in composition of residual fluorine-rich liquids are illustrated by the liquidus projection of the SiO2–Al2SiO4F2–Na3AlF6 ternary (Fig. 14). In this diagram, the system aluminosity (cation Al/Na ratio) increases from the left to the right and fluorine concentrations increase towards the base. Cotectic curves converge to the ternary eutectic with composition Qz19Cry45Tp36 (31· 9 wt % F, Al/Na 0·95), located at 740°C (100 MPa). It is instructive to compare the changes of the eutectic location in several cryolite-bearing binaries. In the cryolite–silica binary, the eutectic occurs close to the cryolite end-member at 1015°C and 100 MPa (Fig. 7). In the cryolite–albite binary, the eutectic occurs close to the silicate end-member (900°C, 1 atm), and the liquid–liquid miscibility gap is inaccessible to the albite-saturated liquid line of descent (Fig. 10). In the cryolite–topaz binary, the eutectic is located in the center of the join (770°C, 100 MPa), without the liquid–liquid immiscibility (Fig. 12). The interactions of these four phases (quartz, albite, topaz and cryolite) control the course of differentiation in the ‘fluorohaplogranite’ system. Increasing melt aluminosity as a result of the addition of topaz causes the miscibility gap to close near the subaluminous composition (Fig. 14), and the paths of cotectic crystallization converge to high fluorine enrichment owing to the strong temperature depression near the cryolite–topaz join.
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THE SILICA–ALBITE–TOPAZ–CRYOLITE SYSTEM |
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The quaternary silica–albite–topaz–cryolite system represents an analogue for the fluorine-bearing granites and rhyolites. In this section, we synthesize the course of liquid lines of descent of quartz–feldspar-precipitating melts with variable alkali/aluminum ratios.
Melting equilibria in the silica–albite–cryolite system reveal that the fluoride–silicate miscibility gap is located on the cryolite liquidus and therefore inaccessible to fractionating peralkaline melts (Fig. 10). On the other hand, the liquidus relations in the silica–topaz–cryolite system indicate the down-temperature extension of the fluoride–silicate liquid immiscibility to subaluminous conditions (Fig. 14). To resolve the termination of the immiscibility for low-temperature subaluminous conditions in the presence of albite, the central portion of the quaternary system has been studied along two pseudobinary sections (Fig. 15): the subaluminous section starts from TCQ-1 (Table 5; Fig. 11) by adding albite (ATCQ-2, ATCQ-1) and the 60 wt %-cryolite isopleth, which emanates from peralkaline AQC-60 (Table 5; Fig. 8) by adding topaz until it reaches the subaluminous composition (ATCQ-3, Fig. 15). This approach allowed us to constrain liquidus and cotectic surfaces in the anhydrous quaternary system (Fig. 16).
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All experimental results from binary, ternary and quaternary sections under anhydrous conditions (Tables 5–9
) are combined in the liquidus projection in Fig. 16. This anhydrous four-component system consists of four liquidus volumes (quartz, albite, topaz and cryolite) and one liquid miscibility gap that extends continuously from the quartz stability field to the cryolite field. Liquid lines of descent follow cotectic surfaces and curves and their relationship to the miscibility gap determines the relevance of liquid–liquid immiscibility for geological compositions. The subaluminous plane (through the tetrahedron, Fig. 16) separates the peraluminous and peralkaline composition spaces and intersects the albite–quartz–topaz cotectic (EAQT–EAQTC curve, Fig. 16), in agreement with location of the quaternary albite–quartz–topaz–cryolite eutectic (EAQTC point) at a weakly peralkaline composition. With decreasing temperature (<700°C), melting in the anhydrous quartz–albite–topaz–cryolite system becomes extremely sluggish and the runs do not reach equilibrium in a reasonable time span. Therefore, a more accurate location and temperature of the anhydrous quaternary eutectic was not determined experimentally. This problem was eliminated by determining phase equilibria in the same system at hydrous conditions in our companion study (Dolej & Baker, 2007).
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Natural silicic magmas with low fluorine contents plot close to the albite–quartz binary eutectic (EAQ, Fig. 16). Magmatic fractionation of quartz and albite will place the liquid compositions on the quartz–albite cotectic surface, and depending on their initial alkali/aluminum ratio, the fractionation of quartz and albite will promote their peralkaline or peraluminous character. Peralkaline melts will reach the cryolite saturation surface and subsequently follow the albite–quartz–cryolite cotectic curve whereas peraluminous melts will saturate with topaz and further evolve along the albite–quartz–topaz cotectic curve. Whereas both ternary cryolite- or topaz-bearing ternary eutectics (EAQC, EAQT) are located at relatively low fluorine contents (less than 5 wt % F), the quaternary eutectic (EAQTC) is displaced to much higher fluorine levels. Therefore, the evolving melts, upon reaching cryolite or topaz saturation, will continue to fractionate along the univariant curves to high levels of fluorine enrichment (towards EAQTC) and their alkali/aluminum ratios will converge. In the absence of other rock-forming minerals (micas, aluminosilicates), which buffer the melt alumina saturation index (Shand, 1927
), the initial peralkaline or peraluminous trends will be reversed by topaz or cryolite saturation, respectively, and the residual melts will converge to a very weakly peralkaline quaternary eutectic. The fluoride–silicate miscibility gap is completely confined to quartz and cryolite stability volumes and it does not penetrate the albite–cryolite cotectic surface. The fluoride–silicate miscibility gap is thus not accessible to any feldspar-precipitating liquid line of descent. |
GEOLOGICAL IMPLICATIONS |
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Experimentally determined liquidus relations in the system Na2O–Al2O3–SiO2–F2O–1 illustrate the location and extent of the silicate–fluoride liquid–liquid immiscibility and describe liquid lines of descent of natural fluorine-bearing silicic magmas. Addition of fluorine to an albitic composition leads to the quartz–topaz–cryolite ternary, which does not contain any pseudoternary phases. Therefore, this system represents a thermal barrier in the quaternary Na2O–Al2O3–SiO2–F2O–1 space, and the stability fields of the ‘high-fluorination’ phases, chiolite and malladrite, are inaccessible to fractionating magmatic systems. We have determined the existence and location of silicate–fluoride liquid–liquid immiscibility in the quaternary silica–albite–cryolite–topaz system. The liquid immiscibility results from coordination differences between individual alkali–aluminofluoride polyhedra and polymerized aluminosilicate framework (Dolej
& Baker, 2005) and it extends from tectosilicate–cryolite binaries towards topaz-bearing systems. In the silica–cryolite binary system, the liquid–liquid miscibility gap is located on the tridymite liquidus above 1100°C at 0·1–600 MPa. In the albite–cryolite binary system, the two-liquid immiscibility occurs above 1000°C at 0·1 MPa (Rutlin, 1998) overlying the cryolite liquidus. Within the silica–albite–cryolite ternary, the miscibility gap closes at 970°C and its location is inaccessible to crystallization paths of alkali feldspar-saturated peralkaline magmas. Instead, magmatic crystallization will be terminated at the ternary eutectic by crystallization of cryolite at 770°C, 100 MPa and 5 wt % F.
Liquid lines of descent of subaluminous and peraluminous silicic magmas are represented by the silica–albite–cryolite–topaz quaternary. The two fluorine-bearing minerals, cryolite and topaz, form a binary eutectic at 770°C, 100 MPa with 39·2 wt % F and cation Al/Na ratio 0·95. The low temperature of this eutectic causes displacement of other ternary and quaternary eutectics towards this binary join. In the silica–cryolite–topaz ternary, the liquid lines of descent terminate at 740°C, 100 MPa with 31·9 wt % F and cation Al/Na ratio 0·95. This implies that residual melts in subaluminous systems can evolve to very high fluorine concentrations (in excess of 30 wt %) with a concomitant decrease in SiO2 (to 30 wt %), without intersecting the fluoride–silicate liquid–liquid immiscibility. In a companion paper (Dolej & Baker, 2007), we include K2O and H2O as additional components and apply experimental results directly to hydrous granitic and rhyolitic systems.
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ACKNOWLEDGEMENTS |
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This study was supported by the Natural Sciences and Engineering Research Council grants to D.R.B. and by the Geological Society of America and the Society of Economic Geologists student grants to D.D. Bob Loeffler provided topaz crystals from the Topaz Mountain, Utah. We would like to acknowledge critical reviews by Don Burt, Bob Linnen, John Longhi and Ron Frost that have led to significant improvements of the manuscript.
*Present address: Bayerisches Geoinstitut, University of Bayreuth, 95440 Bayreuth, Germany. Telephone: +49-(0)921-553718. Fax: +49-(0)921-553769. E-mail: david.dolejs{at}uni-bayreuth.de
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