Geothermometry

Several chemical geothermometers are in widespread use. The silica geothermometer (Fournier and Rowe, 1966) works because the solubilities of the various silica minerals (e.g., quartz and chalcedony, Si02) increase monotonically with temperature. The concentration of dissolved silica, therefore, defines a unique equilibrium temperature for each silica mineral. The Na-K (White, 1970) and Na-K-Ca (Fournier and Truesdell, 1973) geothermometers take advantage of the fact that the equilibrium points of cation exchange reactions among various minerals (principally, the feldspars) vary with temperature. 

Using geochemical modeling, we can apply chemical geo thermometry in a more 

In this chapter, we explore how we can use chemical analyses and pH determinations made at room temperature to deduce details about the origins of natural fluids. These same techniques are useful in interpreting laboratory experiments performed at high temperature, since analyses made at room temperature need to be projected to give pH, oxidation state, gas fugacity, saturation indices, and so on under experimental conditions. 

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Arnorsson, S. and Gunnlaugsson, E. (1983) The chemistry of geothermal waters in Iceland, III. Chemical geothermometry in geothermal investigations. Geochim. Cosmochim. Acta, 47, 567-577. 

Fouillac, C. and Michard, G. (1981) Sodium/lithium ratio in water applied to geothermometry of geothermal reservoirs. Geothermics, 10, 55-70. 

Polythermal reaction models (Section 14.1), however, are commonly applied to closed systems, as in studies of groundwater geothermometry (Chapter 23), and interpretations of laboratory experiments. In hydrothermal experiments, for example, researchers sample and analyze fluids from runs conducted at high temperature, but can determine pH only at room temperature (Fig. 2.2). To reconstruct the original pH (e.g., Reed and Spycher, 1984), assuming that gas did not escape from the fluid before it was analyzed, an experimentalist can calculate the equilibrium state at room temperature and follow a polythermal path to estimate the fluid chemistry at high temperature. 

From a plot of the saturation states of the silica polymorphs (Fig. 23.7), the fluid s equilibrium temperature with quartz is about 100 °C. Quartz, however, is commonly supersaturated in geothermal waters below about 150 °C and so can give erroneously high equilibrium temperatures when applied in geothermometry (Fournier, 1977). Chalcedony is in equilibrium with the fluid at about 76 °C, a temperature consistent with that suggested by the aluminosilicate minerals. 

Land, L. S. and G. L. Macpherson, 1992, Geothermometry from brine analyses, lessons from the Gulf Coast, U.S.A. Applied Geochemistry 7, 333-340. 

Reed, M. and N. Spycher, 1984, Calculation of pH and mineral equilibria in hydrothermal waters with application to geothermometry and studies of boiling and dilution. Geochimica et Cosmochimica Acta 48, 1479-1492. 

The implications of diffusion kinetics for ionic exchange geothermometry have been discussed in detail by Lasaga (1983). 

Figure 5,69 Feldspar geothermometry. From Stormer (1975). Reprinted with permission of The Mineralogical Society of America.

Figure 6.15 gives examples of solid-liquid geothermometry relative to the olivine-liquid (A), orthopyroxene-liquid (A), and plagioclase-liquid (B) pairs. There is satisfactory agreement between experimental and calculated values, and the thermometric procedure may thus be applied with confidence. 

Some aspects of thermally activated diffusion were discussed in section 5.9.1 in regard to intracrystalline exchange geothermometry. Concerning more precisely isotope diffusion, there are two main aspects that may be relevant in geochemistry ... 

Although the scope of this book does not allow an appropriate treatment of stable isotope compositions of earth s materials (excellent monographs on this subject can be found in the literature—e.g., Hoefs, 1980 Faure, 1986), we must nevertheless introduce the significance of the various compositional parameters adopted in the literature before presenting the principles behind stable isotope geothermometry. 

Berger G. W. and York D. (1981). Geothermometry from " °Ar/ Ar dating experiments. Geo-chim. Cosmochim. Acta, 45 795-811. 

Boyd F. R. (1973). The pyroxene geothermometry. Geochim. Cosmochim. Acta, 37 2533-2546. 

Chiba H., Chacko T., Clayton R. N., and Goldsmith J. R. (1989). Oxygen isotope fractionations involving diopside, forsterite, magnetite, and calcite Application to geothermometry. Geochim. Cosmochim. Acta, 53 2985-2995. 

Dahl P. S. (1980). The thermal compositional dependence of Fe -Mg distributions between co-existing garnet and pyroxene Applications to geothermometry. Amer. Mineral, 65 852-866. 

Ganguly J. (1979). Garnet and clinopyroxene solid solutions and geothermometry based on Fe-Mg distribution coefficient. Geochim. Cosmochim. Acta, 43 1021-1029. 

Ghiorso M. S. and Carmichael L S. E. (1980). A regular solution model for meta-aluminous silicate liquids Applications to geothermometry, immiscibility, and the source regions of basic magma. Contrib. Mineral Petrol, 71 323-342. 

Herzberg C. T. (1978b). Pyroxene geothermometry and geobarometry Experimental and thermodynamic evaluation of some subsolidus phase relations involving pyroxenes in the system CaO-MgO-AljOj-SiOj. Geochim. Cosmochim. Acta, 42 945-957. 

Lasaga A. C. (1983). Geospeedometry An extension of geothermometry. In Advances in Physical Geochemistry, vol. 3, S. K. Saxena (series ed.). Berlin-Heidelberg-New York Springer-Verlag. 

Mori T. (1977). Geothermometry of spinel Iherzolites. Contrib. Mineral Petrol, 59 261-279. 

Powell R. and Powell M. (1977b). Geothermometry and oxygen barometry using iron-titanium oxides A reappraisal. Min. Mag, 41 257-263. 

Ross M. and Huebner J. S. (1975). A pyroxene geothermometer based on composition temperature relationship of naturally occurring orthopyroxene, pigeonite and augite. In International Conference on Geothermometry and Geobarometry, The Pennsylvania State University. 

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