Hot spring at Hveravik, Iceland

The system was not closed during cooling, so neither the pH nor CO2 fugacity (Figs. 23.2 and 23.3) returns to its original value at 250 °C. 

The predicted saturation states of the formation minerals (Fig. 23.5), furthermore, no longer identify a unique formation temperature. Whereas the temperatures suggested by albite, quartz, and potassium feldspar are quite close to the 250 °C formation temperature, those predicted by assuming that the fluid was in equilibrium with muscovite and calcite are too low, respectively, by margins of about 25 °C and 100 °C. To avoid error of this sort, we would need to determine the amount of gas lost from the sample and reintroduce it to the equilibrium system before calculating saturation indices. 

The spring yields a Na-Ca-Cl water at about 72 °C with a pH (measured at 

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. 

6 Calculated saturation indices (log QIK of aluminum-bearing minerals plotted versus temperature for a hot spring water from Gjogur, Hveravik, Iceland. Lines for most of the minerals are not labeled, due to space limitations. Sampling temperature is 72°C and predicted equilibrium temperature (arrow) is about 80°C. Clinoptilolite (zeolite) minerals are the most supersaturated minerals below this temperature and saponite (smectite clay) minerals are the most supersaturated above it. 

7 Calculated saturation indices (log QIK of silica minerals for Gjogur hot spring water. Chalcedony is approximately in equilibrium at 80°C, but quartz is supersaturated at this temperature. 

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