30
B. Fractionation Mechanisms of Selected Elements
It is reasonable to assume that due to the described consistency and
to other observations discussed below, oxygen isotope equilibrium is not
always attained in a given assemblage. However, in many cases there is a
close approach to isotopic equilibrium. On the basis of these systematic
tendencies in the ISOrO ratios of many minerals, it has become apparent that significant temperature information could be obtained up to
temperatures of 1000° C and even higher in natural assemblages, if calibration curves could be worked out for the various mineral pairs. In an
assemblage of n-phases, we can obtain n -1 independent "temperatures
of formation", one temperature for each mineral pair. Barring coincidence, if each mineral pair gives the same temperature, we can be certain
that oxygen isotope equilibrium was attained in the mineral assemblage,
and that equilibrium was "frozen in" at the same temperature in every
mineral of the assemblage. If one mineral introduces inconsistent temperatures offormation among a set of concordant equilibrium pairs, it is
certain either that it was not equilibrated with the rest of the assemblage,
or that it has undergone oxygen exchange subsequent to the formation
of the original assemblage. In order for a mineral pair to serve as a good
isotope geothermometer, SCHWARCZ et al. (1970) gave the following
criteria:
1) they should be abundant and commonly found together,
2) they should be stable together over a wide range of pressures and
temperatures,
3) they should exhibit limited compositional variation, since isotopic
fractionation between A and B depends on the chemical composition of
A and B.
Compared with other geothermometers, the advantage of isotope
thermometers is to be not pressure-sensitive. The basic reason for this
independence of pressure changes is that the volumes of isotopically
substituted molecules and crystals do not differ appreciably.
It is important to note that the oxygen isotopic composition of a
mineral is dependent upon the chemical composition of the mineral.
This was for instance experimentally demonstrated for the plagioclase
series by O'NEIL and TAYLOR (1967), who showed that the plagioclaseH20 fractionation depends strongly upon the anorthite content of the
plagioclase (see Fig. 11). This method has been called "internal thermometry" (ONUMA et al. 1972b). On the other hand, the paleotemperature
method of UREY et al. (1951) is based on the fractionation of oxygen
isotopes between calcium carbonate and seawater, so that a temperature
based on isotopic measurement of fossil shells requires an additional
assumption about the isotope ratio in the ocean water at the time of
precipitation. This method, in which one of the phases is not available
B. Fractionation Mechanisms of Selected Elements
It is reasonable to assume that due to the described consistency and
to other observations discussed below, oxygen isotope equilibrium is not
always attained in a given assemblage. However, in many cases there is a
close approach to isotopic equilibrium. On the basis of these systematic
tendencies in the ISOrO ratios of many minerals, it has become apparent that significant temperature information could be obtained up to
temperatures of 1000° C and even higher in natural assemblages, if calibration curves could be worked out for the various mineral pairs. In an
assemblage of n-phases, we can obtain n -1 independent "temperatures
of formation", one temperature for each mineral pair. Barring coincidence, if each mineral pair gives the same temperature, we can be certain
that oxygen isotope equilibrium was attained in the mineral assemblage,
and that equilibrium was "frozen in" at the same temperature in every
mineral of the assemblage. If one mineral introduces inconsistent temperatures offormation among a set of concordant equilibrium pairs, it is
certain either that it was not equilibrated with the rest of the assemblage,
or that it has undergone oxygen exchange subsequent to the formation
of the original assemblage. In order for a mineral pair to serve as a good
isotope geothermometer, SCHWARCZ et al. (1970) gave the following
criteria:
1) they should be abundant and commonly found together,
2) they should be stable together over a wide range of pressures and
temperatures,
3) they should exhibit limited compositional variation, since isotopic
fractionation between A and B depends on the chemical composition of
A and B.
Compared with other geothermometers, the advantage of isotope
thermometers is to be not pressure-sensitive. The basic reason for this
independence of pressure changes is that the volumes of isotopically
substituted molecules and crystals do not differ appreciably.
It is important to note that the oxygen isotopic composition of a
mineral is dependent upon the chemical composition of the mineral.
This was for instance experimentally demonstrated for the plagioclase
series by O'NEIL and TAYLOR (1967), who showed that the plagioclaseH20 fractionation depends strongly upon the anorthite content of the
plagioclase (see Fig. 11). This method has been called "internal thermometry" (ONUMA et al. 1972b). On the other hand, the paleotemperature
method of UREY et al. (1951) is based on the fractionation of oxygen
isotopes between calcium carbonate and seawater, so that a temperature
based on isotopic measurement of fossil shells requires an additional
assumption about the isotope ratio in the ocean water at the time of
precipitation. This method, in which one of the phases is not available