Geol. 656 Isotope Geochemistry
Lecture 31
250
HIGH TEMPERATURE APPLICATIONS II: OXYGEN
ISOTOPES AS AN INDICATOR OF ASSIMILATION
INTRODUCTION
We noted earlier that the equilibrium constant of isotope exchange reactions, K, was proportional to
the inverse square temperature and that isotopic fractionation at high temperature will be limited. In
magmatic systems, another factor limiting the fractionation of stable isotopes is the limited variety of
bonds that O is likely to form. Silicate liquids have short-range structure. Most of the oxygen is not
present as free ions, but is bound to silicon atoms to form silica tetrahedra in the melt, which will be
linked to varying degrees depending on the composition of the melt. The silica tetrahedra, and the Si–
O bonds in the melt, are essentially identical to those in silicate minerals. Thus we would expect the
fractionation of oxygen isotopes between silicate liquids ( magmas) and silicate minerals crystallizing
from those liquids to be rather limited. We would expect somewhat greater fractionation when non-
silicates such as magnetite (Fe3O4) crystallize. In general, crystallization of quartz will lead to a
depletion of 18O in the melt, crystallization of silicates such as olivine, pyroxene, hornblende, and biotite
will lead to slight enrichment of the melt in 18O, and crystallization of oxides such as magnetite and
ilmenite will lead to a more pronounced enrichment of the melt in 18O (however, oxides such as
magnetite are generally only present at the level of a few percent in igneous rocks, which obviously
limits their effect). Crystallization of feldspars can lead to either enrichment or depletion of 18O, de-
pending on the temperature and the composition of the feldspar and the melt. Because quartz gen-
erally only crystallizes very late, the effect of fractional crystallization on magma is generally to
decrease δ18O slightly, generally not more than a few per mil. As we have seen, the range of δ18OSMOW in
the fresh, young basalts and other mantle materials is about +4.5 to +7. Because this range is narrow,
and the range of crustal materials is much greater, O isotope ratios are a sensitive indicator of crustal
assimilation. Isotope ratios outside this range suggest, but do not necessarily prove, the magmas have
assimilated crust (or that post-eruptional isotopic exchange has occurred).
OXYGEN ISOTOPE CHANGES DURING CRYSTALLIZATION
The variation in O isotope composition produced by crystallization of magma will depend on the
manner in which crystallization proceeds. The simplest, and most unlikely, case is equilibrium crys-
tallization. In this situation, the crystallizing minerals remain in isotopic equilibrium with the melt un-
til crystallization is complete. At any stage during crystallization, the isotopic composition of a mineral
and the melt will be related by the fractionation factor, α. Upon complete crystallization, the rock will
have precisely the same isotopic composition as the melt initially had. At any time during the crystalli-
zation, the isotope ratio in the remaining melt will be related to the original isotope ratio as:
R!
R0
=1
f+
!
(1"f)
;
!
=Rs
R!
31.1
where Rl is the ratio in the liquid, Rs is the isotope ratio of the solid, R0 is the isotope ratio of the original
magma, ƒ is the fraction of melt remaining. This equation is readily derived f rom mass balance, the
definition of α, and the assumption that the O concentration in the magma is equal to that in the crys-
tals; an assumption valid to about 10%. Since we generally do not work with absolute ratios of stable
isotopes, it is more convenient to express 31.1, in terms of δ:
!=
"
melt #
"
0$1
ƒ +
%
(1#ƒ)
#1
&
'
()
*
+,1000
31.2
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