Download Stable Isotopes II High Temperature Applications

Geol. 656 Isotope Geochemistry
Chapter 9
Spring 2011
Stable Isotope Geochemistry II: High
Temperature Applications
INTRODUCTION
Stable isotopes have a number of uses in high temperature geochemistry (i.e., igneous and metamorphic geochemistry), which we will treat in the next several lectures. Perhaps the most important of
these is geothermometry, i.e., deducing the temperatures at which mineral assemblages equilibrated.
This application makes use of the temperature dependency of fractionation factors. Other important
applications include reconstructing ancient hydrothermal systems, detecting crustal assimilation in
mantle-derived magmas, and tracing recycled crust in the mantle. These applications primarily involve
O isotopes. Before discussing these subjects, let's briefly review the factors governing isotopic fractionation.
TEMPERATURE DEPENDENCE OF EQUILIBRIUM FRACTIONATIONS
In Chapter 8, we found that the translational and rotation contributions to the partition function do
not vary with temperature. In our example calculation at low temperature, we found the vibrational
contribution varies with the inverse of absolute temperature. At higher temperature, the e-h /kT term in
equation 8.35 becomes finite and this relationship breaks down. It higher temperature, the equilibrium
constant becomes generally proportional to the inverse square of temperature:
ν
ln K = A +
B
T2
9.01
Figure 9.01. a. Comparison of quartz-mineral fractionation factors estimated from the difference in
oxygen site potential (Vqtz – Vmineral) and experimentally observed fractionation factors at 1000 K. b.
Comparison of fractionation factors estimated through the increment method, which also considers
cation mass, and experimentally observed fractionation factors at 1000 K. From Chacko, et al. (2001).
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where A and B are constants. At infinite temperature, the fractionation is unity; i.e., ln K ≈ 0. Because
of the nature of this temperature dependency, fractionation of stable isotopes at mantle temperatures
will usually be small. This is one reason why stable isotopes are useful tracers of crustal assimilation
and recycling.
It must be emphasized that the simple calculations in Chapter 8 are applicable only to a gas whose vibrations can be approximated by a simple harmonic oscillator. Real gases show fractionations that are
complex functions of temperature, with minima, maxima, inflections, and crossovers.
Compositional and Structural Dependence of Equilibrium Fractionations
The nature of the chemical bond is of primary important in determining isotope fractionations. In
general, bonds to ions with a high ionic potential and low atomic mass are associated with high vibrational frequencies and have a tendency to incorporate the heavy isotope preferentially. This point is illustrated by the
site-potential method of estimating fractionation factors (Smyth, 1989). The site potential is simply the
energy required (e.g., in electron volts) to remove an atom from its crystallographic site and is a measure of bond strength. Figure 9.01a shows that the there is a strong correlation between the difference in
oxygen site potential in minerals and the fractionation factor between those two minerals. The solid
line shows that silicates plot along a line with the equation ∆1000K (qtz-mineral) = 0.751 (Vqtz–Vmineral). Oxides (and to a less extent, apatite and calcite) fall off this correlation. The deviation in the case of calcite
and apatite probably reflects the more strongly covalent nature of oxygen bonds in those minerals. In
the case of the oxides, it reflects the varying mass of the cation, as cation mass affects bond strength. In silicates, oxygens
are primarily bound to Si and secondarily
to other cations. However, in oxides such
as rutile, perovskite, and magnetite, oxygen is bound primarily to Fe or Ti, which
of course have very different masses than
Si. The “increment method” of estimating fractionation factors (e.g., Zhang,
1999) takes account of the cation ion
mass.
As Figure 9.01b shows, this
method produces and improved agreement of calculated and experimentally
observed fractionation factors for the oxides.
Substitution of cations in a primarily
ionic site in silicates has only a minor effect on fractionation factors. Thus, for
example, we would expect relatively little
O isotopic fractionation between Kfeldspar and Na-feldspar. This turns out
to be the case: the fractionation is of the
order of 0.1 per mil at room temperature.
The substitution of Al for Si in plagioclase is somewhat more important (substitution of Ca for Na is much less important), leading to a 1.1 permil fracFigure 9.02. Calculated oxygen isotope fractionation for
tionation between anorthite and albite at
several mineral pairs as a function of temperature (from
room temperature. Table 9.01 lists the
O’Neil, 1986).
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parameters for the temperature dependence of silicate and oxide fractionation factors at low temperatures.
Because oxygen occupies a generally similar lattice site in virtually all mantle minerals, it is covalently bonded to silicon and ionically bonded to other cations (e.g., Mg, Fe, Ca, etc.), fractionation of
oxygen isotopes between these phases and during melting are relatively small, although still significant.
Carbonates tend to be very 18O rich because O is bonded to a small, highly charged atom, C4+. The
fractionation relative to water, ∆18O (the ∆ is the difference rather than the fraction) is about 30 for calcite. The cation (i.e., Ca or Mg in carbonate) has a secondary role (because of the effects of its mass on
vibrational frequency). The ∆18O decreases to about 25 for Ba (about 3 times the mass of Ca).
Crystal structure usually plays a secondary role. The ∆18O between aragonite and calcite is of the order of 0.5 permil. But there apparently is a large fractionation (10 permil) of C between graphite and
diamond at room temperature.
Pressure effects turn out to be small, 0.1 permil at 20 kbars and less. The reason should be fairly obvious: there is no volume change is isotope
exchange reactions, and pressure effects TABLE 9.01. COEFFICIENTS OXYGEN ISOTOPE
depend on volume changes. The volume FRACTIONATION AT LOW TEMPERATURES:
of an atom is entirely determined by its
ΔQZ-φ = A + B × 106/T2
electronic structure, which does not deφ
Α
Β
pend on the mass of the nucleus. On the
Feldspar
0
0.97 + 1.04b*
other hand, there will be some minor fracPyroxene
0
2.75
tionation that results from changes in viGarnet
0
2.88
brational frequency as crystals are comOlivine
0
3.91
pressed.
Muscovite
–0.60
2.2
Geothermometry
Amphibole
–0.30
3.15
One of the principal uses of stable isoBiotite
–0.60
3.69
topes is as geothermometers. Like convenChlorite
–1.63
5.44
tional chemical geothermometers, stable
Ilmenite
0
5.29
isotope geothermometers are based on the
Magnetite
0
5.27
temperature dependence of the equi* b is the mole fraction of anorthite in the feldspar. This term
librium constant (equation 9.01). In actualtherefore accounts for the compositional dependence discussed
ity, the constants A and B in equation 9.01
above. From Javoy (1976).
are slowly varying functions of temperature, such that K tends to zero at abso- TABLE 9.02. COEFFICIENTS FOR OXYGEN ISOTOPE
lute 0, corresponding to complete separa- FRACTIONATIONS AT ELEVATED TEMPERATURES
tion, and to 1 at infinite temperature, corre- (600° – 1300°C)
Cc
Ab
An
Di
Fo
Mt
sponding to no isotope separation. We can
0.38
0.94
1.99
2.75
3.67
6.29
obtain a qualitative understanding of why Qz
0.56
1.61
2.37
3.29
5.91
this as so by recalling that the entropy of a Cc
1.05
1.81
2.73
5.35
system increases with temperature. At infi- Ab
0.76
1.68
4.30
nite temperature, there is complete dis- An
0.92
3.54
order, hence isotopes would be mixed ran- Di
2.62
domly between phases (ignoring for the Fo
moment the slight problem that at infinite Coefficients are for mineral pair fractionations expressed as:
temperature there would be neither phases 1000α = B×106/T2 where B is given in the Table. Qz: quartz,
nor isotopes). At absolute 0, there is per- Cc: calcite, Ab: albite, An: anorthite, Di: diopside, Fo: forsterite,
fect order, hence no mixing of isotopes be- Mt: magnetite. For example, the fractionation between albite
tween phases. A and B are, however, suffi- and diopside is 1000α = 1.81×106/T2 (T in kelvins). From
Αλ−Δι
Chiba, et al. (1989).
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ciently invariant over a range of geologically interesting temperatures that as a practical matter they
can be described as constants. We have also noted that at temperatures close to room temperature and
below, the form of equation 9.01 changes to K ∝ 1/T.
Because of the dependence of the equilibrium constant on the inverse square of temperature, stable
isotope geothermometry is employed primarily at low and moderate temperatures, that is, non-magmatic temperatures. At temperatures greater than 800°C or so (there is no exact cutoff), the fractionations are often too small for accurate temperatures to be calculated from them.
In principal, a temperature may be calculated from the isotopic fractionation between any phases
provided the phases achieved equilibrium and the temperature dependence of the fractionation factor
is known. Indeed, there are too many isotope geothermometers for all of them to be even mentioned
here. We can begin by considering silicate systems. Figure 9.02 shows fractionation factors between
various silicates and oxides as a function of temperature. Tables 9.01 and 9.02 list coefficients A and B
for temperature dependence of the fractionation factor between quartz and other common silicates and
oxides when this temperature dependence is expressed as:
Δ ≅ 1000 ln αQz−φ = A +
B
×10 6
2
T
9.02
with temperature expressed in Kelvins. Recall that fundamental rule of thermodynamics states that if
phases A and C and A and B are in equilibrium with each other, then C is also in equilibrium with B.
Thus Table 9.01 may be used to obtain the fractionation between any of the two phases shown.
The other isotope€that has been used extensively for geothermometry of igneous and
metamorphic rocks is sulfur. Its principal
application has been in determining the temperature of deposition of sulfide ores, most of
which precipitate from hydrous fluids. Sulfur may be present in fluids as any one of
several species. Since isotope fractionation
depends on bond strength, the predicted order of 34S enrichment is: SO 24 − > SO 23 − > SO2 >
SCO > Sx ~ H2S ~HS1– > S2– (Ohmoto and Rye,
1979). Figure 9.03 shows the temperature
dependence of fractionation factors between
H2S and other phases, and Table 9.03 lists coefficients for the equation:
Δ ≅ 1000 ln α φ −H 2 S = A +
€
B
×10 6
2
T
9.03
CO2 and other carbon–bearing species are
ubiquitous in meteoric and hydrothermal
waters. Carbonates often precipitate from
such solutions and the fractionation between
carbon-species provides yet another opportunity for geothermometry.
Figure 9.04
shows carbon isotope fractionation factors
between CO2 and other carbon bearing species as a function of temperature. The figure
includes fractionation factors both calculated
from theory and observed vibrational frequencies (calcite, carbonate ion, carbon mon-
Figure 9.03. Relationship of S isotope fractionation between H2S and other sulfur-bearing species and temperature. From Ohmoto and Rye (1979).
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oxide, methane) and empirical determined
values (dolomite, bicarbonate ion, and carbonic acid). Table 9.04 lists coefficients for
a third degree polynomial expression of
temperature dependence.
The Importance of Equilibrium
All geothermometers are based on the
apparently contradictory assumptions that
complete equilibrium was achieved between phases during, or perhaps after,
formation of the phases, but that the phases
did not re-equilibrate at any subsequent
time. The reason these assumptions can be
made and geothermometry works at all is
the exponential dependence of reaction
rates on temperature. Isotope geothermometers have these same implicit assumptions about the achievement of equilibrium as other systems.
The importance of the equilibrium basis
of the geothermometry must be emphasized. Because most are applied to relatively low temperature situations, violation
of the assumption that complete equilibrium was achieved is common. We have
seen that isotopic fractionations may arise
from kinetic as well as equilibrium effects.
If reactions do not run to completion, the
isotopic differences may reflect kinetic effects as much as equilibrium effects. Other
problems can result in incorrect temperature as well: the system may partially reequilibration at some lower temperatures
during cooling; free energies of isotope exchange reactions are low, meaning there is
little chemical energy available to drive the
reaction to equilibrium. Indeed, isotopic
equilibrium probably often depends on
other reactions occurring which mobilize
the element involved in the exchange.
Solid-state exchange reactions are particularly slow at temperatures well below the
melting point. Equilibrium between solid
phases will thus generally depend on reaction of these phases with a fluid. Of
course, this is true of ‘conventional’ chemical reactions as well, and metamorphism
generally occurs in the presence of a fluid.
Spring 2011
TABLE 9.03. COEFFICIENTS FOR SULFUR ISOTOPE
FRACTIONATION:
Δφ−H2S = A + B × 106/T2 (T IN KELVINS)
φ
CaSO4
SO2
FeS2
ZnS
CuS
Cu2S
SnS
MoS2
Ag2S
PbS
Β
A
6.0±0.5 5.26
–0.5±0.5 4.7
0.4±0.08
0.10±0.05
–0.4±0.1
–0.75±0.1
–0.45±0.1
0.45±0.1
–0.8±0.1
–0.63±0.05
T°C
Range
200-350
350-1050
200-700
50-705
50-700
From Ohmoto and Rye (1979)
Figure 9.04. Fractionation factors for distribution of carbon isotopes between CO2 and other carbon-bearing
species as a function of temperature. From Ohmoto and
Rye (1979).
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Figure 9.05 compares sphalerite–ga- TABLE 9.04 ISOTOPE FRACTIONATION FACTORS OF
lena sulfur isotope temperatures with CARBON COMPOUNDS WITH RESPECT TO CO
2
fluid-inclusion homogenization tem- 1000 Ln α = A × 108/T3 + B × 106/T2 + C × 103/T + D
peratures. Excluding the Pine Point
φ
Α
B
C
D
T°C
data, the best fit to the data is fairly
Range
close to that expected from the fractionation factors listed in Table 9.03:
-8.914
8.737 -18.11
8.44
≤600
Δsp-gn=0.73×106/T2.
Many of the CaMg(CO3)2
Ca(CO
)
-8.914
8.557
-18.11
8.24
≤600
3
points fall off the expected curve, in1–
HCO
0
-2.160
20.16
-35.7
≤290
dicating disequilibrium. Ohmoto and
2–3
CO
-8.361
-8.196
-17.66
6.14
≤100
3
Rye (1979) noted a number of factors
H
CO
0
0
0
0
≤350
2
3
that may contribute to the lack of fit,
CH
4.194
-5.210
-8.93
4.36
≤700
4
such as impure mineral separates used
CO
0
-2.84
-17.56
9.1
≤330
in the analysis; for example, 10% of the
C
-6.637
6.921
-22.89
9.32
≤700
galena in sphalerite and visa versa
From
Ohmoto
and
Rye
(1979).
would result in an estimated temperature of 215°C if the actual
equilibration temperature was
145°C.
Different minerals may
crystallize at different times and
different temperatures in a hydrothermal system and hence
would never be in equilibrium. In
general, those minerals in direct
contact with each other give the
most reliable temperatures. Real
disequilibrium may also occur if
crystallization is kinetically controlled. The generally good fit to
the higher temperature sulfides
and poor fit to the low temperature
ones suggests kinetics may indeed
be an important factor.
Isotope geothermometers do
have several advantages over conventional chemical ones. First, as Figure 9.05. Comparison of temperatures determined from
we have noted, there is no volume
sphalerite–galena sulfur isotope fractionation with fluid-inchange associated with isotopic exclusion homogenization temperatures. J: Creede, CO, S: Sunchange reactions and hence little
pressure dependence of the equi- nyside, O: Finlandia vein, C: Pasto Bueno, G: Kuroko. From
librium constant (however, Rum- Ohmoto and Rye (1979).
ble has suggested an indirect pressure dependence, wherein the fractionation factor depends on fluid composition which in turn depends
on pressure). Second, whereas conventional chemical geothermometers are generally based on solid
solution, isotope geothermometers can make use of pure phases such as SiO2, etc. Generally, any dependence on the composition of phases in isotope geothermometers involved is of relatively second order importance. For example, isotopic exchange between calcite and water is independent of the concentration of CO2 in the water. Compositional effects can be expected only where composition affects
bonds formed by the element involved in the exchange. For example, we noted substitution of Al for Si
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in plagioclase affects O isotope fractionation factors because the nature of the bond with oxygen. The
composition of a CO2 bearing solution, however, should not affect isotopic fractionation between calcite
and dissolved carbonate because the oxygen is bonded with C regardless of the presence of other ions
(if we define the fractionation as between water and calcite, some effect is possible if the O in the carbonate radical exchanges with other radicals present in the solution).
STABLE ISOTOPE COMPOSITION OF THE MANTLE
Before we can use stable isotope ratios as indicators of crustal assimilation and tracers of crustal recycling, we need to define the stable isotopic composition of “uncontaminated” mantle. It is, however,
important to recognize from the outset that, in a strict sense, there may be no such thing. We found in
our consideration of radiogenic isotope ratios that no samples of “primitive” mantle have been recovered: the mantle, or at least that portion
sampled by volcanism, has been pervasively processed. A very considerable
amount of oceanic crust has been subducted during this time, perhaps accompanied by sediment. As we shall
see, the stable isotopic composition of
the oceanic crust is extensively modified by hydrothermal processes and
low temperature weathering.
Subduction of this material has the potential for modifying the stable isotopic
composition of the mantle. Thus while
we will attempt to use stable isotope
ratios to identify “contamination” of
mantle by subduction, we must recognize all of it may have been “contaminated” to some degree.
Other problems arise in defining the
stable isotope composition of the mantle. We relied heavily on basalts as
mantle samples in defining the radiogenic isotope composition of the mantle. We could do so because radiogenic
isotope ratios are not changed in the
magma generation process. This will
not be strictly true of stable isotope ratios, which can be changed by chemical
processes. The effects of the melting
process on most stable isotope ratios of
interest are small, but not completely
negligible.
Degassing does significantly affect stable isotope ratios, particularly those of carbon and hydrogen,
which compromises the value of magmas as a mantle sample. Once oxides
Figure 9.06. Oxygen isotope ratios in olivines and clibegin to crystallize, fractional crystallinopyroxenes from mantle peridotite xenoliths. Data from
zation will affect oxygen isotope ratios,
Mattey et al. (1994).
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although the resulting changes are at most a few per mil. Finally, weathering and hydrothermal processes can affect stable
isotope ratios of basalts and other igneous rocks. Because hydrogen, carbon, nitrogen, and sulfur are all trace elements in
basalts but are quite abundant at the Earth’s surface, these
elements are particularly susceptible to weathering effects.
Even oxygen, which constitutes nearly 50% by weight of silicate rocks, is readily affected by weathering. Thus we will
have to proceed with some caution in using basalts as samples
of the mantle for stable isotope ratios.
Oxygen
Assessing the oxygen isotopic composition of the mantle,
and particularly the degree to which its oxygen isotope composition might vary, has proved to be more difficult than expected. One approach has been to use basalts as samples of
mantle, as is done for radiogenic isotopes. Relatively little isotope fractionation occurs during partial melting, so the oxygen
isotopic composition of basalt should the same as that in the
mantle source within a few tenths per mil. However, assimilation of crustal rocks by magmas and oxygen isotope exchange
during weathering complicate the situation. An alternative is
to use direct mantle samples such as xenoliths occasionally
found in basalts, although these are considerably rarer than
are basalts.
Figure 9.06 shows the oxygen isotope composition of olivines and clinopyroxenes in 76 peridotite xenoliths analyzed
by Mattey et al. (1994) using the laser fluorination technique.
The total range of values observed is only about twice that expected from analytical error alone, suggesting the mantle is
fairly homogeneous in its isotopic composition. The difference
between co-existing olivines and clinopyroxenes averages
about 0.5 per mil, which is consistent with the expected fractionation between these minerals at mantle temperatures.
Mattey et al. (1994) estimated the bulk composition of these
samples to be about +5.5 per mil.
Figure 21.7 of MORB (+5.7). After
Figure 9.07 shows the distribution of δ18O in selected basalts
Harmon and Hoefs (1994).
from 4 different groupings. To avoid the weathering problems
we discussed above, Harmon and Hoefs (1994) included only submarine basaltic glasses and basalts
that had less than 0.75% water or had erupted historically in their compilation. There are several points
worth noting in these data.
MORB are significantly more homogeneous than are other basalts. MORB have a mean δ18OSMOW of
+5.7‰ and a standard deviation of ±0.2‰. Thus the depleted upper mantle appears to be a comparatively homogeneous and well-mixed reservoir for oxygen, just as it is for other elements. Oceanic island basalts, which presumably sample mantle plumes, are slightly less enriched in 18O (mean δ18OSMOW
= +5.5‰), and are also more variable (1σ = 0.5‰). The histogram shown excludes Iceland, because Icelandic basalts are quite anomalous in their low δ18O (mean ~ 4.5‰). This has been shown to be due
primarily to assimilation of older basaltic crust that has equilibrated with meteoric water, which is quite
18
O depleted at the latitude of Iceland. There is, however, some evidence to suggest even primary Ice-
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landic basalts are depleted in 18O relative to MORB. Subduction-related basalts (i.e., island arc basalts
and their continental equivalents) are shifted to more positive δ18O values. This may well reflect contributions from the subducting slab, and we shall explore this further in a later lecture. Continental subduction related basalts are more 18O rich than their oceanic equivalents, most likely due to assimilation
of continental crust. Finally, continental intraplate volcanics are more enriched in 18O than are OIB,
again suggestive of crustal assimilation.
Carbon
The stable isotopes of H, C, N, and S are much more difficult to analyze in igneous rocks. These elements are generally trace elements and are volatile. With rare exceptions, they have a strong tendency
to exsolve from the melt and escape as gases as magmas approach the surface of the earth. Not only are
these elements lost during degassing, but they can be isotopically fractionated by degassing. Thus
there is far less data on the isotopes of these elements in basalts, and the meaning of this data is somewhat open to interpretation.
Most carbon in basalts is in the form of CO2, which has limited solubility in basaltic liquids at low
pressure. As a result, basalts begin to exsolve CO2 before they erupt. Thus virtually every basalt sample has lost some carbon, and subareal basalts have lost virtually all carbon (as well as most other volatiles). Therefore, only basalts erupted beneath several km of water provide useful samples of mantle
carbon. As a result, the data set is restricted to MORB, samples recovered from Loihi and the submarine part of Kilauea’s East Rift Zone, and a few
other seamounts.
The question of the isotopic composition of
mantle carbon is further complicated by fractionation and contamination.
There is a
roughly 4‰ fractionation between CO2 dissolved in basaltic melts and the gas phase, with
13
C enriched in the gas phase. If Rayleigh distillation occurs, that is if bubbles do not remain
in equilibrium with the liquid, then the basalt
that eventually erupts may have carbon that is
substantially lighter than the carbon originally
dissolved in the melt. Furthermore, MORB are
pervasively contaminated with a very 13C-depleted carbon. This carbon is probably organic
in origin, and recent observations of an eruption on the East Pacific Rise suggest a source.
Following the 1991 eruption at 9°30’ N, there
was an enormous ‘bloom’ of bacteria stimulated by the release of H2S. Bacterial mats covered everything. The remains of these bacteria
may be the source of this organic carbon. Fortunately, it appears possible to avoid most of
this contamination by the step-wise heating
procedure now used by most laboratories.
Most of the contaminant carbon is released at
Figure 21.8. Carbon isotope ratios in mantle (red)
temperatures below 600°C, whereas most of the
and mantle-derived materials (gray). After Mattey
basaltic carbon is released above 900°C.
(1987).
Figure 9.08 shows δ13C in various mantle and
mantle-derived materials. MORB have a mean
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δ13C of -6.5 and a standard deviation of 1.7. Hawaiian basalts appear to have slightly heavier carbon.
Xenoliths in oceanic island basalts are also slightly heavier than MORB. Whether this reflects a real difference in isotopic composition or merely the effect of fractionation is unclear. The most CO2-rich
MORB samples have δ13C of about –4. Since they are the least degassed, they presumably best represent the isotopic composition of the depleted mantle (Javoy and Pineau, 1991). If this is so, there may
be little difference in carbon isotopic composition between MORB and oceanic islands sampled thus far
(which include only Hawaii, Reunion, and Kerguelen). Gases released in subduction zone volcanos
have δ13C that ranges from 0 to –10‰, with most values being in the range of –2 to –4‰, comparable to
the most gas-rich MORB (Javoy, et al., 1986). Continental xenoliths are more heterogeneous in carbon
isotopic composition than other groups, and the meaning of this is unclear. Carbonatites have somewhat lighter carbon than most MORB.
Diamonds show a large range of carbon isotopic compositions (Figure 9.08). Most diamonds have
δ13C within the range of -2 to -8‰, hence similar to MORB. However, some diamonds have much
lighter carbon. Based on the inclusions they contain, diamonds can be divided between peridotitic and
eclogitic. Most peridotitic diamonds have δ13C close to –5‰, while eclogitic diamonds are much more
isotopically variable. Most, though not all, of the diamonds with very negative δ13C are eclogitic. Many
diamonds are isotopically zoned, indicating they grew in several stages.
Three hypotheses have been put forward to explain the isotopic heterogeneity in diamonds: primordial heterogeneity, fractionation effects, and recycling of organic carbon from the Earth’s surface
into the mantle. Primordial heterogeneity seems unlikely for a number of reasons. Among these is the
absence of very negative δ13C in other materials, such as MORB, and the absence of any evidence for
primordial heterogeneity from the isotopic compositions of other elements. Boyd and Pillinger (1994)
have argued that since diamonds are kinetically sluggish (witness their stability at the surface of the
Earth, where they are thermodynamically out of equilibrium), isotopic equilibrium might not achieved
during their growth. Large fractionations might therefore occur due to kinetic effects. However, these
kinetic fractionations have not been demonstrated, and fractionations of this magnitude (20‰ or so)
would be surprising at mantle temperatures.
On the other hand, several lines of evidence support the idea that isotopically light carbon in some
diamonds had its origin as organic carbon at
the Earth’s surface. First, such diamonds are
primarily of eclogitic paragenesis, eclogite is
the high pressure equivalent of basalt. Subduction of oceanic crust continuously carries
large amounts of basalt into the mantle.
Oxygen isotope heterogeneity observed in
some eclogite xenoliths suggests these eclogites do indeed represent subducted oceanic
crust. Second, the nitrogen isotopic composition of isotopically light diamonds are
anomalous relative to nitrogen in other mantle materials yet similar to nitrogen in sedimentary rocks.
Hydrogen
Figure 9.09. Effect of degassing and post-eruptive
processes on the water content and δD of basalts.
From Kyser and O’Neil (1984).
239
Like carbon, hydrogen can be lost from basalts during degassing. On the one hand, the
problem is somewhat less severe than for
carbon because the solubility of water in basalt is much greater than that of CO2. Basalts
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erupted beneath a kilometer of more of water probably retain most of their dissolved
water.
However, basalts, particularly
submarine basalts, are far more readily
contaminated with hydrogen (i.e., with water) than with carbon. Furthermore, the effect on hydrogen isotopic composition depends on the mode of contamination, as
Figure 9.09 indicates. Direct addition of
water or hydrothermal reactions will raise
δD (because there is little fractionation during these processes), while low temperature weathering and hydration will lower
δD, because hydrogen, rather than deuterium, is preferentially incorporated into alteration phases. Loss of H2 and CH4, which
may partition into a CO2 gas phase when it
forms, could also affect the hydrogen isotopic composition of basalts. However, the
available evidence suggests that these species constitute only a small fraction of the
hydrogen in basalts, so this effect is likely Figure 9.10. δD in MORB and in mantle phlogopites
to be minor.
and amphiboles. The MORB and phlogopite data sugAs Figure 9.10 indicates, MORB have gest the mantle has δDSMOW of about –60 to –90.
mean δDSMOW of about –67.5‰ and a standard deviation of ±14‰. How much of this variability reflects the processes shown in Figure 9.09 and
how much reflects true heterogeneity in the mantle is unclear. Kyser (1986) has argued that mantle hydrogen is homogeneous with δDSMOW of –80‰. The generally heavier isotopic composition of MORB,
he argues, reflects H2O loss and other processes. However, Poreda, et al., (1986) found that δD in basalts from the Reykjanes Ridge south of Iceland correlated significantly with La/Sm and other trace
element ratios, suggesting at least some of the isotopic variation of hydrogen in basalts reflects real
variations in the mantle. Submarine basalts from Kilauea’s East Rift Zone have higher δD than MORB.
Kyser and O’Neil (1984) argued that these higher values result from the addition of water to the magma
in the rift zone. Hawaiian submarine basalts analyzed by Garcia et al. (1989) have δD very similar to
MORB.
Hydrous minerals in xenoliths also provide a sample of mantle hydrogen. As Figure 9.10 shows,
phlogopites have δD that is generally similar to that of MORB, though some lighter values also occur.
Amphiboles have much more variable δD and have heavier hydrogen on average. Part of this difference probably reflects equilibrium fractionation. The fractionation between water and phlogopite is
close to 0‰ in the temperature range 800°-1000°C, whereas the fractionation between water and amphibole is about –15‰. However, equilibrium fractionation alone cannot explain either the variability
of amphiboles or the deference between the mean δD of phlogopites and amphiboles. Complex processes involving in amphibole formation that might include Rayleigh distillation may be involved in
the formation of mantle amphiboles. This would be consistent with the more variable water content of
amphiboles compared to phlogopites.
Nitrogen
Figure 9.11 summarizes the existing data on the nitrogen ratios in the crust and mantle. There is far
less data than for other stable isotope ratios because of the low concentrations and pervasive con-
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tamination problems. The solubility of
N2 in basalts is very limited, though
much of the nitrogen may be present as
+
NH 4 , which is somewhat more soluble.
Hence of volcanic rocks, once again only
submarine basalts provide useful samples of mantle N. There are both contamination and analytical problems with
determining nitrogen in basalts, which,
combined with its low abundance (generally less than a ppm), mean that accurate measurements are difficult to make.
Measurements of δ15NATM in MORB range
from about –2 to +12‰. The few available analyses of Hawaiian basalts range
up to +20. At present, it is very difficult
to decide to what degree this variation
reflects contamination (particularly by
organic matter), fractionation during degassing, or real mantle heterogeneity.
Perhaps all that can be said is that nitrogen in basalts appears to have positive
δ15N on average.
Diamonds can contain up to 2000 ppm
of N and hence provide an excellent
sample of mantle N. As can be seen in Figure 9.11. Isotopic composition of nitrogen in rocks and
Figure 21.11, high δ13C diamonds (most minerals of the crust and mantle. Modified from Boyd and
common, and usually of peridotitic par- Pillinger (1994).
agenesis) have δ15N that range from -12 to +5 and average about -3‰, which contrasts with the generally positive values observed in basalts. Low δ13C diamonds have generally positive δ15N. Since organic matter and ammonia in crustal rocks generally have positive δ15N, this characteristic is consistent
with the hypothesis that this group of diamonds are derived from subducted crustal material. However, since basalts appear to have generally positive δ15N, other interpretations are also possible. Fibrous diamonds, whose growth may be directly related to the kimberlite eruptions that carry them to
the surface (Boyd et al., 1994), have more uniform δ15N, with a mean of about -5‰. Since there can be
significant isotopic fractionations involved in the incorporation of nitrogen into diamond, the meaning
of the diamond data is also uncertain, and the question of the nitrogen isotopic composition of the mantle remains an open one.
Sulfur
There are also relatively few sulfur isotope measurements on basalts, in part because sulfur is lost
during degassing, except for those basalts erupted deeper than 1 km below sealevel. In the mantle, sulfur is probably predominantly or exclusively in the form of sulfide, but in basalts, which tend to be
somewhat more oxidized, some of it may be present as SO2 or sulfate. Equilibrium fractionation
should lead to SO2 being a few per mil lighter than sulfate. If H2S is lost during degassing the re2−
maining sulfur would become heavier; if SO2 or SO 4 is lost, the remaining sulfur would become
34
lighter. Total sulfur in MORB has δ SCDT in the range of +1.3 to –1‰, with most values in the range 0 to
+1‰. Sakai et al. (1984) found that sulfate in MORB, which constitutes 10-20% of total sulfur, was 3.5
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to 9‰ heavier than sulfide. Basalts from Kilauea’s East Rift Zone have a very restricted range of δ34S of
+0.5 to +0.8 (Sakai, et al., 1984).
Chaussidon et al. (1989) analyzed sulfides present as inclusions in minerals, both in basalts and in
xenoliths, and found a wide range of δ34S (–5 to +8‰). Low Ni sulfides in oceanic island basalts, kimberlites, and pyroxenites had more variable δ34S that sulfides in peridotites and peridotite minerals.
They argued there is a fractionation of +3‰ between sulfide liquid and sulfide dissolved in silicate
melt. Carbonatites have δ34S between +1 and –3‰ (Hoefs, 1986; Kyser, 1986). Overall, it appears the
mantle has a mean δ34S in the range of 0 to +1‰, which is very similar to meteorites, which average
about +0.1‰.
Chaussidon, et al. (1987) found that sulfide inclusions in diamonds of peridotitic paragenesis (δ13C ~
-4‰) had δ34S of about +1‰ while eclogitic diamonds had higher, and much more variable δ34S (+2 to
+10‰). Eldridge et al. (1991) found that δ34S in diamond inclusions was related to the Ni content of the
sulfide. High Ni sulfide inclusions, which they argued were of peridotitic paragenesis, had δ34S between +4‰ and –4‰. Low Ni sulfides, which are presumably of eclogitic paragenesis, had much more
variable δ34S (+14‰ to –10). These results are consistent with the idea that eclogitic diamonds are derived from subducted crustal material.
OXYGEN ISOTOPES IN MAGMATIC PROCESSES
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. Most of the oxygen in silicates and in silicate liquids is not present as free
ions, but is bound to silicon atoms to form silica tetrahedra, which will be linked to varying degrees
depending on the composition of the mineral or the composition of the melt. The silica tetrahedra and
the Si–O bonds in a silicate 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 limits their effect).
Crystallization of feldspars can
lead to either enrichment or
depletion of 18O, depending on
the temperature and the comFigure 9.12. Plot of δ18O versus fraction of magma solidified dur18
position of the feldspar and the
ing Rayleigh fractionation, assuming the original δ O of the
melt. Because quartz generally
magma was +6. After Taylor and Sheppard (1986).
only crystallizes very late, the
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general effect of fractional crystallization on magma is generally to increase δ18O slightly, generally not
more than a few per mil. On the other hand, water-rock interaction at low to moderate temperatures at
the surface of the Earth, produces much larger O isotopic changes, generally enriching the rock in 18O.
Consequently, crustal rocks show significant variability in δ18O.
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 those 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 crystallization. In this situation, the crystallizing minerals remain in isotopic equilibrium with the melt until 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 crystallization, the isotope ratio in the remaining melt will be related to the original isotope ratio as:
R
1
R
=
α= s
R0 f + α (1 − f ) ;
R
9.04
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 from mass balance, the
definition of α, and the assumption that the O concentration in the magma is equal to that in the crystals; an assumption valid to about 10%. Since we generally do not work with absolute ratios of stable
isotopes, it is more convenient to express 9.04, in terms of δ:
1
⎡
⎤
Δ = δ melt − δ 0 ≅ ⎢
− 1⎥ × 1000
⎣ ƒ + α (1 − ƒ) ⎦
9.05
where δmelt is the value of the magma
after a fraction ƒ-1 has crystallized
and δo is the value of the original
magma. For silicates, α is not likely
to be much less than 0.998 (i.e., ∆ =
δ18Omelt – δ18Oxtals ≥ 2). For α = 0.999,
even after 99% crystallization, the
isotope ratio in the remaining melt
will change by only 1 per mil.
Fractional crystallization is a process
analogous to Rayleigh distillation.
Indeed, it is governed by the same
equation (8.69), which we can rewrite
as:
Δ = 1000(ƒα −1 −1)
€
9.06
The key to the operation of either
of these processes is that the product
of the reaction (vapor in the case of
distillation, crystals in the case of
Figure 9.13. δ18O as a function of SiO2 in a tholeiitic suite from
the Galapagos Spreading Center (GSC) (Muehlenbachs and
Byerly, 1982) and an alkaline suite from Ascension Island
(Sheppard and Harris, 1985). Dashed line shows model calculation for the Ascension suite.
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crystallization) is only instantaneously in equilibrium with the original phase. Once it is produced, it is
removed from further opportunity to attain equilibrium with original phase. This process is more efficient at producing isotopic variations in igneous rocks, but its effect remains limited because α is generally not greatly different from 1. Figure 9.12 shows calculated changes in the isotopic composition of
melt undergoing fractional crystallization for various values of ∆ (≈ 1000(α–1)). In reality, ∆ will change
during crystallization because of (1) changes in temperature (2) changes in the minerals crystallizing,
and (3) changes in the liquid composition. The changes will generally mean that the effective ∆ will increase as crystallization proceeds. We would expect the greatest isotopic fractionation in melts crystallizing non-silicates such as magnetite and melts crystallizing at low temperature, such as rhyolites,
and the least fractionation for melts crystallizing at highest temperature, such as basalts.
Figure 9.13 shows observed δ18O as a function of temperature in two suites: one from a propagating
rift on the Galapagos Spreading Center, the other from the island of Ascension. There is a net change in
δ18O between the most and least differentiated rocks in the Galapagos of about 1.3‰; the change in the
Ascension suite is only about 0.5‰. These, and other suites, indicate the effective ∆ is generally small,
on the order of 0.1 to 0.3‰.
We can generalize the temperature dependence of stable isotopes by saying that at low temperature
(ambient temperatures at the surface of the Earth up to the temperature of hydrothermal systems, 300400° C), stable isotope ratios are changed by chemical processes. The amount of change can be used as
an indication the nature of the process involved, and, under equilibrium conditions, of the temperature
at which the process occurred. At high temperatures (temperatures of the interior of the Earth or magmatic temperatures), stable isotope ratios are only minimally affected by chemical processes; they can
therefore be used as tracers much as radiogenic isotope ratios are.
These generalizations lead to a final axiom: igneous rocks whose oxygen isotopic compositions show significant variations from the primordial value (~6) must either have been affected by low temperature processes, or
must contain a component that was at one time at the surface of the earth (Taylor and Sheppard, 1986).
Rocks that have equilibrated with water at the surface of the Earth, e.g., sediments, tend to have δ18O
values significantly higher than +6 (generally greater than +10). Waters on the surface of the Earth
typically have δ18O significantly lower than +6 (0 for seawater, generally lower for meteoric, or fresh,
waters). Interestingly, the δD values of
sedimentary rocks and mantle-derived
igneous rocks are rather similar (- 50 to 85). This may be coincidental since the
δD of sediments are controlled by fractionation between minerals and water,
whereas δD of igneous rocks reflects the
isotopic composition of mantle hydrogen
(mantle water). One the other hand, it is
possible that it is not coincidental. Plate
tectonics results in the return of waterbeing rocks from the surface of the Earth
to the mantle (i.e., via subduction). It
may well be that after 4.5 billion years,
the subduction process essentially controls the isotopic composition of H in the
upper mantle.
Combined Fractional
Crystallization and Assimilation
Figure 9.14. Variation in δ18O of a magma undergoing AFC
vs. amount crystallized. Initial δ18O of the magma is +5.7.
Because oxygen isotope ratios of man- After Taylor (1980).
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tle-derived magmas are reasonably uniform and generally different from rocks that have equilibrated
with water at the surface of the Earth, oxygen isotopes are a useful tool in identifying and studying the
assimilation of country rock by intruding magma. We might think of this process as simple mixing between two components: magma and country rock. In reality, it is always at least a three-component
problem, involving country rock, magma, and minerals crystallizing from the magma. Magmas are essentially never superheated; hence the heat required to melt and assimilate surrounding rock can only
come from the latent heat of crystallization of the magma. Approximately 1000 J/g would be required
to heat rock from 150° C to 1150°C and another 300 J/g would be required to melt it. If the latent heat
of crystallization is 400 J/g, crystallization of 3.25 g of magma would be required to assimilate 1 g of
country rock. Since some heat will be lost by simple conduction to the surface, we can conclude that
the amount of crystallization will inevitably be greater than the amount of assimilation (the limiting
case where mass crystallized equals mass assimilated could occur only at very deep layers of the crust
where the rock is at its melting point to begin with). The change in isotopic composition of a melt undergoing the combined process of assimilation and fractional crystallization (AFC) is given by:
Δ ⎞
⎛
δ m − δ 0 = ⎜ [δ a − δ 0 ] − ⎟ {1 − ƒ R /(1− R) }
⎝
R ⎠
9.07
where R is the mass ratio of material assimilated to material crystallized, ∆ is the difference in isotope
ratio between the crystals and the magma (δcrystals−δmagma), ƒ is the fraction of liquid remaining, δm is the
δ18O of the magma, δo is the initial δ18O of the magma, and δa is the δ18O of the material being assimilated. The assumption is made that the concentration of oxygen is the same in the crystals, magma and
assimilant, which is a reasonable assumption. This equation breaks down at R = 1, but as discussed
above, this is unlikely: R will always be less than 1. Figure 9.14 shows the variation of δ18O of a magma
with an initial δ18O = 5.7 as crystallization and assimilation proceed.
Combining Radiogenic and Oxygen Isotopes
A more powerful tool for
the study of assimilation
processes can result if O
isotopes are combined
with a radiogenic isotope
ratio such as 87Sr/86Sr.
There are several reasons
for this. First, radiogenic
isotopes and stable isotopes each have their own
advantages. In the case of
basaltic magmas, radiogenic elements, particularly Nd and Pb, often
have lower concentrations
in the magma than in the
assimilant. This means a
small amount of assimilant
will have a large effect on
the radiogenic isotope ratios. On the other hand,
18
87
86
oxygen will be present in Figure 9.15. δ O vs. Sr/ Sr showing range of isotopic composition in
the magma and assimilant various terrestrial rocks. After Taylor and Sheppard (1986).
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at nearly the same concentration, making calculation of the mass assimilated fairly straightforward.
Also, it is easier to uniquely characterize the assimilant using both radiogenic and stable isotope ratios,
as suggested in Figure 9.15.
The equation governing a radiogenic isotope ratio in a magma during AFC is different from 9.07 because we cannot assume the concentration of the element is the same in all the components. On the
other hand, there is no fractionation between crystals and melt. The general equation describing the
variation of the concentration of an element in a magma during AFC is:
⎛ R ⎞ C a
Cm
−z
z
=
f
+
⎜
⎟ 0 (1− f )
0
⎝ R −1 ⎠ zCm
Cm
9.08
where Cm and the concentration of the element in the magma, Co is the original magma concentration, ƒ
are as defined for equation 9.07 above, and z is as:
€
z=
R + D −1
R −1
9.09
where D is the solid/liquid partition coefficient The isotopic composition of the magma is the given by
(DePaolo, 1981):
⎛ R ⎞ C a
−z
0 −z 0
⎜
⎟ (1− f )ε a + C m f ε
⎝ R −1 ⎠ z
εm =
⎛ R ⎞ C a
−z
0 −z
⎜
⎟ (1− f ) + C m f
⎝ R −1 ⎠ z
9.10
where ε is the isotope ratio with subscripts m, a
and 0 denoting the magma, the assimilant, and
the original magma respectively and the other
parameters are €
defined as in equation 9.08. Figure 9.16 shows calculated AFC curves on a plot
of δ18O vs. 87Sr/86Sr. Note that except in the case
where β = 1, where the problem simplifies to one
of simple mixing, the mixing lines end at ƒ = 0.01
(99% crystallized).
Figure 9.17 shows an actual case, the Adamello
Massif in Italy. Actual analyses are plotted as
dots with the concentrations of Sr in the sample
shown adjacent to the dot. AFC lines are computed assuming the original magma had δ18O =
5.6, 87Sr/86Sr = 0.704, and 750 ppm Sr, and the
country rocks have 87Sr/86Sr = 0.736, δ18O = +13.6,
and 150 ppm Sr. Dashed lines are contours of calculated Sr concentrations in the magma. There is
reasonably good agreement between the calculated model and the actual data, if we assume the
bulk partition coefficient varied a bit (which
would certainly be the case).
Figure 9.16. Variation of δ18O with 87Sr/86Sr during
AFC for a magma with an initial δ18O = 5.7 and
87
Sr/86Sr = 0.703, and an assimilant with 87Sr/86Sr =
0.735 and δ18O = +19. All curves are for R = 0.2
(5:1), except for one with D = 2 for which R = 0.11
(labeled 9:1). Dashed red lines are calculated Sr
Sediment Subduction vs. Assimilation
concentrations (ppm) assuming an initial concentraMagmas erupted and emplaced in subduction tion of Sr in the magma of 500 ppm. Where D = 1,
zones have unique geochemistry that is apparent the problem simplifies to one of simple mixing.
in both their trace element and radiogenic iso- After Taylor (1980).
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topic compositions. Most often, this unique
geochemistry is taken as evidence that the
sources of these magmas contain subducted
sediments. However, in many instances, this
interpretation is non-unique. The geochemical signature of subducted sediment is, in
many respects, simply the signature of continental crust. In general, it is just as plausible
that magmas have acquired this signature
through assimilation of continental crust as
by partial melting of a mantle containing
subducted sediment. Magaritz et al. (1978)
pointed out that by combining radiogenic
isotope and oxygen isotope analyses, it is
possible to distinguish between these two
possibilities.
Magaritz et al. reasoned as follows. First,
many continental materials have oxygen isotope ratios that are quite different from man18
87
86
tle values (e.g., Figure 9.18). Then they noted Figure 9.17. Variation of δ O with Sr/ Sr in the Adathat all rocks have similar oxygen concentra- mello Massif in Italy compared with model AFC process
tions. Because of this, addition of 10% sedi- using equation 9.07- 9.10. After Taylor and Sheppard
ment with δ18O of +20 to the mantle (+6) fol- (1986).
lowed by melting of that mantle produces a
magma with the same δ18O (+8.4), as adding
10% sediment directly to the magma (i.e., in
this case the magma also ends up with δ18O =
+8.4). However, addition of 10% sediment to
the mantle, followed by melting of that mantle
produces a magma with a very different Sr isotope ratio than does adding 10% sediment directly to the magma. The reason for this difference is that the Sr concentration of the magma
is much higher (an order of magnitude or more
higher) than that of the mantle. Thus addition
of sediment to magma affects the Sr isotope ratio less than addition of sediment to the mantle.
Indeed, most subduction-related magmas are
richer in Sr (~ 400 ppm) than most continental
Figure 9.18. Mixing curves on a plot of δ18O vs.
materials and sediments; but sediments in87
Sr/86Sr. The labels on the lines refer to different
variably have more Sr than does the mantle
ratios of Sr concentration in the two end-members.
(10-40 ppm).
After James (1981).
This point is illustrated in a more quantitative fashion in Figure 9.18. We discussed mixing trajectories earlier and noted that when two materials are mixed in varying proportions, a curve results on a plot of one ratio against another. The degree and sense of curvature depended on the ratio of
the ratio of the two denominators in the two end-members. Since the concentration of oxygen is approximately the same in most rocks, the ratio of 16O in the two end-members can be taken as 1. So, for
example, in plot of 87Sr/86Sr vs. oxygen isotope ratio, the degree of curvature depends only on the ratio
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of Sr in the two end-members. In Figure
9.18, when the ratio of Sr in end-member
M (magma or mantle) to Sr in endmember C (crust or sediment) is greater
than 1, a concave downward mixing
curve results. If that ratio is less than 1
(i.e., concentration of Sr in C is greater
than in M), a concave upward curve results. Assimilation of crust by magma essentially corresponds to the concave
downward case (SrM > SrC) and mixing of
subducted sediment and mantle corresponds to the concave upward case (SrM
< SrC). Thus these two processes are
readily distinguished on such a plot.
Figure 9.19 shows the data of Magaritz
et
al. (1978) for the Banda arc. These data
Figure 9.19. Oxygen and Sr isotopic composition of
fall
roughly along a mixing curve for
Banda arc magmas. Mixing curve constructed assuming
Sr
/Sr
C
M = 3. The samples are mainly baa concentration ratio of SrC/SrM = 3. End member M has
saltic
andesites
and andesites and have
87
86
18
Sr/ Sr = 0.7035 and δ O = +6; end member C has
reasonably
high
Sr concentrations (15087
86
18
Sr/ Sr = .715 and δ O = +14. Ticks indicate the per500
ppm).
Magaritz
et al. (1978) argued
centage of end member C in the mixture.
that the apparent 3:1 concentration ratio
between the C and M members was more
consistent with mixing of pelitic sediments (which have 100-150 ppm Sr) with mantle than assimilation
of crust by magma. The two triangles represent the isotopic composition of corderite-bearing lavas
from the island of Ambon. The petrology and composition of these lavas suggest they are produced directly by melting of pelitic sediments in the deep crust. Thus their isotopic compositions were taken as
potentially representative of the assimilant.
Considerable care must be exercised in the interpretation of oxygen isotope ratios in igneous rocks,
particularly when the rocks come from a tropical region such as the Banda Arc. In such regions, rocks
quickly begin to react with water and primary phases in the groundmass are replaced with various secondary minerals. The isotopic composition of such minerals will be much higher (δ18O ≈ + 20 to +30)
than the original ones. Thus the O isotopic composition of weathered igneous rocks will be higher than
that of the original magma. Thus in their study, Magaritz et al. (1978) also analyzed plagioclase phenocrysts. Phenocrysts generally weather less rapidly than does the groundmass and alteration effects
can be easily identified visually, which is not always the case with the groundmass. Thus phenocrysts
often provide a better measure of the oxygen isotopic composition of the original magma than does the
whole rock. This is true even though there might have been a small equilibrium fractionation between
the magma and the phenocryst.
The amount of subducted sediment implied by the mixing curve in Figure 9.19 is much greater than
has been inferred to be involved in other arcs (about 1-3%). The Banda arc is directly adjacent to the
Australian continent, so extensive involvement of sediment and crustal material is perhaps to be expected. On the other hand, the amount of sediment, or crustal material in the mixture is sensitive to assumptions about the oxygen isotope composition of that material. If the δ18O of the crustal material is
assumed to be +22 rather than +14, the calculated percentage of sediment involved would decrease by
50%.
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Figure 9.20. Sr, Nd, and O isotopes in lavas from Nevado del Ruiz Volcano in Columbia and an AFC
model calculation. Dashed line shows AFC trajectory for another Colombian volcano, Galeras. After
James and Murcia (1984).
The mean δ18O of Marianas lavas is, in contrast to Banda, about +6.2 (Ito and Stern, 1986; Woodhead
et al., 1987). This implies the amount of sediment involved in the source is less than 1%, which is consistent with the amount inferred from radiogenic isotope studies.
In contrast to this interpretation, James and Murcia (1984) have used radiogenic and stable isotope
systematics to argue for extensive assimilation of crust by magmas in the northern Andes. As shown in
Figure 9.20, andesites from Nevado del Ruiz and Galeras volcanoes in Columbia define a steep array on
plots of δ Ο vs. 87Sr/86Sr and 143Nd/144Nd. Comparing this with Figure 9.18, we see that such steep arrays imply mixing between magma and crust rather than mantle and sediment. As we observed in the
last lecture, assimilation will inevitably be accompanied by fractionation crystallization. James and
Murcia (1984) modeled this assimilation using the equations similar to 9.07- 9.10 and the assumed concentrations listed in Table 9.05. The model fits the data reasonably well and implies assimilation of up
to about 12% country rock. An R of 0.33 provided the best fit, though James and Murcia (1984) noted
this parameter was not well constrained. A high value of R suggests the assimilation is occurring at
moderate depth (< 10 km), because it suggests only moderate input of heat in melting the country rock.
On a broader basis, Harmon et al. (1984) have used O isotope ratios to show that extent of crustal assimilation in Andean magmas varies regionally and correlates with crustal thickness. The Andes are
divided into 3 distinct volcanic provinces. In the north, the subduction zone is steeply dipping, and the
volcanoes are located approximately 140 km over the Benioff
Table 9.05: Parameter for Ruiz Assimilation Model
zone. The crust in the Northern
Original Magma
Assimilant
Volcanic Zone (NVZ), located in Coδ18O
6.5‰
11‰
lumbia and Ecuador, is approxi∆
0.5‰
mately 40 km thick and is mainly
87
Sr/86Sr
0.7041
0.710
Cretaceous and Cenozoic in age.
βSr
0.2
Volcanics are primarily basaltic
Sr, ppm
650
100
andesites and andesites. In the
143
Nd/144Nd
0.51288
0.5122
Central Volcanic Zone (CVZ), which
Nd, ppm
20
20
extends from southern Peru to
βNd
0.2
northern Chile and Argentina, the
R
0.33
crust is 50 to 70 km thick and Precambrian to Paleozoic in age. The
18
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Figure 9.21. Oxygen and Sr isotope ratios in Andean volcanics. Dashed line shows the mean
δ18O of MORB. After Harmon et al. (1984).
Benioff Zone lies approximately 140 km beneath these volcanoes. The volcanics range from andesite
through rhyolite, and include extensive ignimbrites. The Southern Volcanic Zone (SVZ) is located in
south-central Chile and overlies a shallower, more gently dipping Benioff Zone. These volcanoes lie
roughly 90 km above the Benioff Zone and atop 30-35 km of Mesozoic and Cenozoic crust. Here the
volcanics are mainly basaltic, with minor andesite and dacite.
Figure 9.21 shows that δ18O is most mantle-like in the SVZ, with a mean δ18O of about +6.0, just
slightly higher than the mean for MORB. 87Sr/86Sr ratios average about 0.704, just slightly higher than
the mean for intra-oceanic arcs. The NVZ has higher δ18O, with a mean about +6.9, and somewhat
higher 87Sr/86Sr. δ18O in the CVZ is highest, ranging from +7 to +14. This corresponds to the region of
the thickest and oldest crust in the Andes. Volcanics in the southern part of the CVZ, from 21-26° S,
have apparently experienced the greatest assimilation. Minimum δ18O values of around +7 in this region suggest that all magmas have assimilated
at least some crust. Figure 9.22 shows an example of data from one volcano in the CVZ, Cerro
Galan in northwest Argentina. Sr and O isotope
data define broad, slightly convex-upward array
that is reminiscent of the AFC curves in Figure
9.16 for a value of D ≈ 2.
Stable Isotopes as Indicators of Crust-toMantle Recycling
Figure 9.22. 87Sr/86Sr and δ18O in Cerro Galan,
and ignimbrite center in Northwest Argentina.
After Harmon et al. (1984).
250
The potential importance of the crust–to–
mantle flux resulting from subduction was first
pointed out by Armstrong in 1968. More specifically, Hofmann and White (1982) proposed
that mantle plumes consist in part of subducted
oceanic crust. As we shall see, the oceanic crust
undergoes substantial modification of its oxygen
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isotopic composition as a result of the hydrothermal activity and metamorphism that occurs virtually simultaneously with formation of the oceanic crust and as a result of low-temperature reactions with seawater. The
hydrothermal activity tends to lower the δ18O of the
middle and lower oceanic crust, while the lowtemperature weathering raises the δ18O of the upper
oceanic crust.
Interestingly, there is little or no net change in oxygen isotope composition of the oceanic crust as a result
of these processes. Thus subduction of the crystalline
portion of the oceanic crust should have no net effect
on the oxygen isotopic composition of the mantle as a
whole, but it might increase isotopic heterogeneity.
The lower average δ18O of OIB (Figure 9.07) is nevertheless interesting in the context of the Hofmann and
White hypothesis. Accepting the axiom of Taylor and
Sheppard (1986) that material with δ18O substantially
different from the mantle value (as a working definition of the mantle value, we will adopt the mean value Figure 9.23. Oxygen isotope ratios in eclogite
of MORB of +5.7), the lower average δ18O and the xenoliths from the Roberts Victor Mine kimgreater heterogeneity of δ18O in OIB strongly supports berlite.
Figure 9.24. δ18O in olivines from oceanic island basalts and
MORB. See figure 14.4 for the definition of EM1, EM2, and
HIMU. Shaded areas shows the range of δ18O in upper mantle peridotites (e.g., Figure 21.6). From Eiler et al. (1997).
the interpretation that the plumes
that produce these magmas contain
material that was once at the
surface of the Earth. One might
speculate that the lower mean
value for OIB arises because upper
oceanic crust is somehow less
likely to survive subduction into
the deeper mantle, perhaps because it partially melts in subduction zones (or has in the past when
the Earth was hotter).
The subduction process may also
influence the composition of the
subcontinental lithosphere.
As
may be seen in Figure 9.23,
eclogite* xenoliths from the Roberts
Victor
Mine
(South
Africa)
kimberlite have remarkably variable oxygen isotope compositions
(MacGregor and Manton (1986).
The range in δ18O is similar to that
of “mature” oceanic crust (i.e.,
* Eclogite is a rock consisting primarily of Na-rich pyroxene and garnet that forms only at high pressures. Compositionally these eclogites are similar to basalts, as are most.
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subsequent to hydrothermal low temperature
metamorphism), and contrasts with the
homogeneity of peridotite xenoliths (Figure
9.06), many of which also come from kimberlites. Since eclogites are the high pressure
equivalent of basalts, this suggests the
ecolgites come from subducted oceanic crust.
Further evidence to this effect is the
correlation observed between O isotope ratios,
Sr isotope ratios, and K and Rb concentrations,
as these will also be affected by metamorphism of the oceanic crust. Geobarometry
suggests these eclogites last equilibrated at
depths of 165-190 km. Radiogenic isotope
evidence suggests they are 2.47 Ga old.
We examined the radiogenic isotope
evidence for mantle heterogeneity in Chapter
6. One interpretation is that this heterogeneity
results from recycling of oceanic crust and
sediment into the mantle. Since sediment,
and, to a lesser degree, oceanic crust (as we
shall see in the next section) differs in its
oxygen isotopic compositions from normal
mantle material, we might expect this
recycling to produce variations in O isotope
composition of basalts. As Figure 9.24 shows,
this is indeed observed. EM2 basalts (defined
by their radiogenic isotopic compositions as
shown in Figures 6.11 and 6.12) show sys18
Figure 9.25. Relationship between δ O and K2O
tematically higher δ18O than MORB, while
and (La/Sm)N (the chondrite-normalized La/Sm
HIMU and low 3He/4He basalts (many HIMU
ratio) in MORB glass. Solid line shows the exbasalts have low 3He/4He) have systematically
pected result if recycled altered oceanic crust is
low δ18O. It is in EM2 basalts, such as those
mixed with depleted peridotite and then melted
from the Society Islands, where the best
10%. Dashed line shows the expected result if a
evidence for a recycled sedimentary
10% melt of eclogitic recycled oceanic crust is
component is seen (e.g. Figure 6.18).
mixed with a 10% melt of depleted peridotite.
Even among MORB, there is a correlation
From Eiler and Schiano (2000).
between incompatible element enrichment
and δ18O (Figure 9.24). Since δ18O should be effectively insensitive to the extent of partial melting and
fractional crystallization, this correlation is likely to reflect a correlation between these properties in the
MORB source. Since O isotope ratios can be changed only at low temperatures near the Earth’s surface,
Eiler and Schiano (2000) interpret this as evidence of a recycled component in MORB.
STABLE ISOTOPES IN HYDROTHERMAL SYSTEMS
Ridge Crest Hydrothermal Activity and Metamorphism of the Oceanic Crust
Early studies of “greenstones” dredged from mid-ocean ridges and fracture zones revealed they were
depleted in 18O relative to fresh basalts. Partitioning between of oxygen isotopes between various minerals, such as carbonates, epidote, quartz and chlorite, in these greenstones suggested they had equilibrated at about 300° C (Muehlenbachs and Clayton, 1972). This was the first, but certainly not the only,
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evidence that the oceanic crust underwent hydrothermal metamorphism at depth. Other clues included highly variable heat flow at ridges and an imbalance in the Mg fluxes in the ocean. Nevertheless, the importance of hydrothermal processes was not generally recognized until the discovery of low
temperature (~20° C) vents on the Galapagos Spreading Center in 1976 and high temperature (350° C)
“black smokers” on the East Pacific Rise in 1979. Various pieces of the puzzle then began to fall rapidly
into place and it was soon clear that hydrothermal activity was a very widespread and important phenomenon. Most of the oceanic crust is affected to some degree by this process, which also plays an important role in controlling the composition of seawater.
Hydrothermal metamorphism occurs because seawater readily penetrates the highly fractured and
therefore permeable oceanic crust. A series of chemical reactions occurs as the seawater is heated,
transforming it into the reduced, acidic, and metal-rich fluid. Eventually the fluid rises and escapes,
forming the dramatic black smokers. Fluid in many “black smokers” vents at temperatures of about
350° C*. This results from the density and viscosity minimum that occurs close to this temperature at
pressures of 200-400 bars combined with a rapidly decreasing rock permeability above these temperatures.
Of the reactions that occur in the process, only one, namely oxygen isotope exchange, concerns us
here. Seawater entering the oceanic crust has a δ18O of 0, fresh igneous rock has a δ18O of +5.7. As seawater is heated, it will exchange O with the surrounding rock until equilibrium is reached. At temperatures in the range of 300-400° C and for the mineral assemblage typical of greenschist facies basalt, the
net water-rock fractionation is small†, 1 or 2‰. Thus isotopic exchange results in a decrease in the δ18O
of the rock and an increase in the δ18O of the water. Surprisingly, there have only been a few oxygen
isotope measurements of vent fluids; these indicate δ18O of about +2.
At the same time hydrothermal metamorphism occurs deep in the crust, low-temperature weathering
proceeds at the surface. This also involves isotopic exchange. However, for the temperatures (~2° C)
and minerals produced by these reactions (smectites, zeolites, etc.), fractionations are quite large (something like 20‰). The result of these reactions is to increase the δ18O of the shallow oceanic crust and
decrease the δ18O of seawater. Thus the effects of low temperature and high temperature reactions are
opposing.
Muehlenbachs and Clayton (1976) suggested that these opposing reactions actually buffered the isotopic composition of seawater at a δ18O of ~0. According to them, the net of low and high temperature
fractionations was about +6, just the observed difference between the oceanic crust and the oceans.
Thus, the oceanic crust ends up with an average δ18O value about the same as it started with, and the
net effect on seawater must be close to zero also. Could this be coincidental? One should always be
suspicious of apparent coincidences in science, and they were.
Let’s think about this a little. Let’s assume the net fractionation is 6, but suppose the δ18O of the ocean
was –10 rather than 0. What would happen? Assuming a sufficient amount of oceanic crust available
and a simple batch reaction with a finite amount of water, the net of high and low temperature basaltseawater reactions would leave the water with δ18O of –10 + 6 = –4. Each time a piece of oceanic crust is
allowed to equilibrate with seawater, the δ18O of the ocean will increase a bit. If the process is repeated
enough, the δ18O of the ocean will eventually reach a value of 6 – 6 = 0. Actually, what is required of
seawater–oceanic crust interaction to maintain the δ18O of the ocean at 0‰ is a net increase in isotopic
* While this is typical, temperatures of 400° C or so have also been found. Most low-temperature vents waters, such
as those on the GSC appear to be mixtures of 350° C hydrothermal fluid and ambient seawater, with mixing occurring at shallow depth beneath the seafloor. Although hydrothermal fluids with temperatures substantially above 400°
C have not been found, there is abundance evidence from metamorphosed rocks that water-rock reactions occur at
temperatures up to 700° C.
† While the mineral-water fractionation factors for quartz and carbonate are in the range of +4 to +6 at these temperatures, the fractionation factor for anorthite and chlorite are close to zero, and that for magnetite is negative.
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composition of seawater by perhaps 1–
2‰. This is because low-temperature
continental weathering has the net effect
of decreasing the δ18O of the hydrosphere. This is what Muehlenbachs and
Clayton proposed.
The “half-time” for this process, has been
estimated to be about 46 Ma. The halftime is defined as the time required for
the disequilibrium to decrease by half.
For example, if the equilibrium value of
the ocean is 0 ‰ and the actual value is 2 ‰, the δ18O of the ocean should increase to –1 ‰ in 46 Ma. It would then
require another 46 Ma to bring the
oceans to a δ18O of –0.5‰, etc. Over long
time scales, this should keep the isotopic
Figure 9.26. δ18O of the Samail Ophiolite in Oman as a funccomposition of oceanic crust constant.
We’ll see in a subsequent lecture that the tion of stratigraphic height. After Gregrory and Taylor
oxygen isotopic composition of the ocean (1981).
varies on times scales of thousands to tens of thousands of years due to changes in ice volume. The iso-
Figure 9.27. δ18O in marine carbonates (brachiopods, belemites, oysters, and foraminfera shells) over
Phanerozoic time. From Veizer et al. (1999).
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topic exchange with oceanic crust is much too slow to dampen these short-term variations.
That water-rock interaction produces essentially no net change in the isotopic composition of the oceanic crust, and therefore of seawater was apparently confirmed by the first thorough oxygen isotope
study of an ophiolite by Gregory and Taylor (1981). Their results for the Samail Ophiolite in Oman are
shown in Figure 9.26. As expected, they found the upper part of the crust had higher δ18O than fresh
MORB and while the lower part of the section had δ18O lower than MORB. Their estimate for the δ18O
of the entire section was +5.8, which is essentially identical to fresh MORB. Ocean Drilling Project
(ODP) results show much the same pattern as the Samail ophiolite, though no complete section of the
oceanic crust has yet been drilled. ODP results do show that hydrothermal alteration is not uniform. In
Hole 504B, the deepest hole yet drilled, the transition to the hydrothermally altered zone was found to
be quite sharp.
If the Muehlenbachs and Clayton hypothesis is correct and assuming a steady-state tectonic environment, the δ18O of the oceans should remain constant over geologic time. Whether it has or not, is
controversial. Based on analyses of marine carbonates, Jan Veizer and his colleagues have argued that
it is not. Figure 9.27 shows the variavariation in δ18O in marine carbonates
over Phanerozoic time. The isotopic
composition of carbonates reflect (1) the
composition of water from which they
precipitated and (2) the fractionation
between water and carbonate. The
latter is large (~30‰) and temperature
dependent. However, over reasonable
ranges of temperature, the fractionation
factor will vary by only a few per mil.
(As is convention, the data in Figure
9.27 are reported relative to PDB, which
is about +30‰ relative to SMOW. This
roughly matches the fractionation between water and carbonate: water with
δ18OSMOW= 0 should precipitate carbonate with δ18OPDB≈ 0.) The variations in
Figure 9.27 are much larger than this,
implying that seawater oxygen isotopic
composition has indeed varied significantly. Subsequent reaction and equilibration of carbonates with pore water
could shift the δ18O of the carbonates,
but these data are from carefully screen
samples that are least likely to be altered in this way. Some of the shortterm variations might be due to
changes in ice volume. Increasing ice
volume would shift δ18O positively, and
indeed, there is some evidence that δ18O
Figure 9.28. Change in sealevel and the isotopic composiwas higher when the climate was
tion of water in the model of Wallmann (2001) compared
colder, such as during the late Ordoviwith actually sea level and the Veizer et al. (1999) carboncian and Carboniferous glaciations. In
ate data. Model is the run at two different recycling effiaddition to the short-term variations
ciencies (r), 0.3 and 0.5. From Wallmann (2001).
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there is an overall increase from about -8‰ in the Cambrian to 0‰ at present. This clearly contradicts
the idea that ridge crest hydrothermal activity buffers the δ18O to a constant value.
Wallmann (2001) has produced a box model of the isotopic composition of ocean water that matches,
in a very qualitative way, the secular increase in δ18O observed by Veizer. He argues that the long-term
isotopic composition of seawater depends not only on ridge crest hydrothermal activity, but also on a
number of other inputs to and removals from the ocean, particularly those related to the deep subduction cycle. He also argues that the isotopic composition of subducted water in oceanic crust and sediment that is less that the isotopic composition of water degassed from the mantle, and further, that water has been subducted at a higher rate than it has been degassed from the mantle. Over the Phanerozoic, this, he argues, has produced a decrease in the mass of the oceans and in sealevel. His model this
shows an increase in δ18O and a decrease in sealevel (Figure 9.28). However, the controversy has not
been resolved. Based on δ18O of ancient ophiolites (the Oman ophiolite is Cretaceous), Muehlenbachs
continues to argue that the oxygen isotopic composition of seawater has not changed significantly
throughout Phanerozoic and Proterozoic time (e.g., Muehlenbachs and Furnas, 2003).
Meteoric Geothermal Systems
Hydrothermal systems occur not only in the ocean, but just about everywhere that magma is intruded into the crust. In the 1950’s a debate raged about the rate at which the ocean and atmosphere
were created by degassing of the Earth’s interior. W. W. Rubey assumed that water in hydrothermal
systems such as Yellowstone was
magmatic and argued that the ocean
and atmosphere were created quite
gradually through magmatic degassing.
Rubey turned out to be
wrong. One of the first of many important contributions of stable isotope
geochemistry to understanding hydrothermal systems was the demonstration by Craig (1963) that water in
these systems was meteoric, not magmatic. The argument is based upon the
data shown in Figure 9.29. For each
geothermal system, the δD of the
“chloride” type geothermal waters is
the same as the local precipitation and
groundwater, but the δ18O is shifted to
higher values. The shift in δ18O results
from “high” temperature (~300°C) reFigure 9.29. δD and δ18O in meteoric hydrothermal systems.
action of the local meteoric water with
Closed circles show the composition of meteoric water in the
hot rock. However, because concentravicinity of Yellowstone, Steamboat Springs (Colorado), Mt.
tion of hydrogen in rocks is nearly 0
Lassen (California), Iceland, Larderello (Italy), and The Gey(more precisely because ratio of the
sers (California), and open circles show the isotopic composimass of hydrogen in the water to mass
tion of chloride-type geothermal waters at those locations.
of hydrogen in the reacting rocks is exOpen triangles show the location of acidic, sulfide-rich geotremely high), there is essentially no
thermal waters at those locations. Solid lines connect the mechange in the hydrogen isotopic comteoric and chloride waters, dashed lines connect the meteoric
position of the water. If the water inand acidic waters. The “Meteoric Water Line” shows the corvolved in these systems was magmatic,
relation between δD and δ18O observed in precipitation
it would not have the same δD as local
shown in Figure 8.10.
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meteoric water (though it is possible that these systems contain a few percent magmatic water).
Acidic, sulfide-rich water from these systems does have δD that is different from local meteoric water. This shift occurs when hydrogen isotopes are fractionated during boiling of geothermal waters.
The steam produced is enriched in sulfide, accounting for their acidic character. The water than condenses from this steam the mixes with meteoric water to produce the mixing lines observed.
Water-Rock Reaction: Theory
Very often in geology it is difficult to observe the details of processes occurring today and our understanding of many of Earth processes comes from observing the effects these processes have had in the
past, i.e., the record they have left in the rocks. So it is with hydrothermal systems. In present systems,
we very often can observe only the water venting, we cannot observe the reactions with rocks or the
pattern of circulation. However, these, as well as temperatures involved and water-rock ratios, can be
inferred from the imprint left by ancient hydrothermal systems.
Estimating temperatures at which ancient hydrothermal systems operated is a fairly straightforward
application of isotope geothermometry, which we have already discussed. If we can measure the oxygen (or carbon or sulfur) isotopic composition of any two phases that were in equilibrium, and if we
know the fractionation factor as a function of temperature for those phases, we can estimate the temperature of equilibration. We will focus now on water-rock ratios, which may also be estimated using
stable isotope ratios.
For a closed hydrothermal system, we can write two fundamental equations. The first simply describes equilibrium between water and rock:
Δ = δ wf − δ rf
9.11
where we use the subscript w to indicate water, and r to indicate rock. The superscript f indicates the
final value. So equation 9.10 just says that the difference between the final isotopic composition of water and rock is equal to the fractionation factor (we implicitly assume equilibrium). The second equation is just a statement of mass balance for a closed system: the amount of the isotope present before reaction must be the same as after reaction:
cwW δ wi + cr Rδ ri = cwW δ wf + cr Rδ rf
9.12
where c indicates concentration (we assume concentrations do not change, which is valid for oxygen,
but perhaps not valid for other elements), W indicates the mass of water involved, R the mass of rock
involved and the superscript f denotes the final isotope ratio.
Substituting equation 9.10 and rearranging, we derive the following equation:
W
δ rf − δ ri cr
=
R δ wi − δ rf − Δ cw
9.13
The term on the left is the ratio of water to rock in the reaction. Notice that the r.h.s. does not include
the final isotopic composition of the water, information that we would generally not have. The initial
oxygen isotope composition of the water can be estimated in various ways. For example, we can determine the hydrogen isotopic composition (of rocks) and from that determine the oxygen isotope
composition using the δD– δ18O meteoric water line. The initial δ18O of rock can generally be inferred
from unaltered samples, and the final isotopic composition of the rock can be measured. The fractionation factor can be estimated if we know the temperature and the phases in the rock. For oxygen, the ratio of concentration in the rock to water will be close to 0.5 in all cases.
Equation 9.13 is for a closed system, i.e., a batch reaction where we equilibrate W grams of water with
R grams of rock. That is not very geologically realistic. In fact, a completely open system, where water
makes one pass through hot rock, would be more realistic. In this case, we might suppose that a small
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parcel of water, dW, passes through the system and induces and incremental change in the isotopic
composition of the rock, dδr. In this case, we can write:
(
)
Rcr dδ r = δ wi − [ Δ + δ r ] cw dW
9.14
This equation states that the mass of isotope exchanged by the rock is equal to the mass of isotope exf
changed by the water (we have substituted ∆ + δr for δw ). Rearranging and integrating, we have:
⎛ δ f − δ i
⎞ c
W
= ln⎜ f r i r +1⎟ r
R
⎝ −δr + δw − Δ ⎠ cw
9.15
Thus it is possible to deduce the water rock ratio for an open system as well as a closed one.
Using this kind of approach, Gregory and Taylor (1981) estimated water/rock ratios of ≤ 0.3 for the
gabbros of the Oman ophiolite. It should be emphasized, however, that this can be done with other isotope systems as well.
€ For example, McCulloch et al. (1981) used Sr isotope ratios to estimate water/rock ratios varying from 0.5 to 40 for
different parts of the Oman ophiolite.
The Skaergaard Intrusion
A classic example of a meteoric hydrothermal system is the Early Tertiary Skaergaard intrusion in East Greenland. The
Skaergaard has been studied for nearly 75
years as a classic mafic layered intrusion.
Perhaps ironically, the initial motivation
for isotopic study of the Skaergaard was
determination of primary oxygen and hydrogen isotopic compositions of igneous
rocks. The results, however, showed that
the oxygen isotope composition of the
Skaergaard has been pervasively altered by
hydrothermal fluid flow. This was the first
step in another important contribution of
stable isotope geochemistry, namely the
demonstration that most igneous intrusions have reacted extensively with water
subsequent to crystallization.
Figure 9.30 shows a map of the Skaergaard with contours of δ18O superimposed
on it. Figure 9.31 shows a restored cross
section of the intrusion with contours of
δ18O. There are several interesting features.
First, it is clear that circulation of water
was strongly controlled by permeability.
The impermeable basement gneiss experienced little exchange, as did the part of the
intrusion beneath the contact of the gneiss
with the overlying basalt. The basalt is
quite permeable and allowed water to flow
freely through it and into the intrusion.
Figures 9.30 and 9.31 define zones of low
Figure 9.30. Oxygen isotope variations in the Skaergaard
Intrusion. LZ, MZ, and UZ refer to the ‘lower zone, ‘middle zone’ and ‘upper zone’ of the intrusion, which dips
20-25° to the southeast. UBZ refer to the ‘upper border
group’. The δ18O = +6 contour corresponds more or less
to the trace of the gneiss-basalt contact through the
intrusion (SW to NE). The gneiss is essentially impermeable, while the basalt is highly fractured. Thus
most water flow was above this contact, and the gabbro
below it retained its original ‘mantle’ isotopic signature
(+6). After Taylor (1974).
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1‰
Basalt
.
1‰ Border Group
4‰
>5.7‰
4‰
Layered Series
present
Topography
Gneiss
Figure 9.31. Restored cross-section of the Skaergaard intrusion with contours of δ18O.
Figure 9.32. Cartoon illustrating the hydrothermal system
in the Skaergaard intrusion. After Taylor (1968).
δ18O, which are the regions of hydrothermal upwelling. Water was apparently
drawn into the sides of the intrusion and
then rose above. This is just the sort of
pattern observed with finite element
models of fluid flow through the intrusion.
Calculated water-rock ratios for the
Skaergaard were 0.88 in the basalt, 0.52 in
the upper part of the intrusion and 0.003
for the gneiss, demonstrating the importance of the basalt in conduction the water
into the intrusion and the inhibiting effect
of the gneiss. Models of the cooling history of the intrusion suggest that each cm3
of rock was exposed to between 105 and 5
× 106 cm3 of water over the 500,000 year
cooling history of the intrusion. This
would seem to conflict with the water/rock ratios estimated from oxygen isotopes. The difference is a consequence of
each cc of water flowing through many
cc’s of rock, but not necessarily reacting
with it. Once water had flowed through
enough grams of rock to come to isotopic
equilibrium, it would not react further
with the rock through which it subsequently flowed (assuming constant temperature and mineralogy). Thus it is important to distinguish between W/R ratios
calculated from isotopes, which reveal
only the mass (or molar) ratio of water
and rock in the net reaction, to flow models. Nevertheless, the flow models demonstrate that each gram of rock in such a
system is exposed to an enormous amount
of water. Figure 9.32 is a cartoon illustrating the hydrothermal system deduced
from the oxygen isotope study.
Oxygen Isotopes and Mineral
Exploration
Oxygen isotope studies can be a valuable tool in mineral exploration. Mineralization is very often
(though not exclusively) associated with the region of greatest water flux, such as areas of upward
moving hot water above intrusions. Such areas are likely to have the lowest values of δ18O. To understand this, let’s solve equ. 9.15, the final value of δ18O:
δ = (δ − δ + Δ)e
f
r
€
i
r
i
w
−Wcw
Rcr
259
+ δwi − Δ
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If we assume a uniform initial isotopic
δr
composition of the rocks and the water,
then all the terms on the r.h.s. are con5
stants except W/R and ∆, which is a
function of temperature. Thus the final
values of δ18O, i.e., the values we
measure in an area such as the Skaer0
gaard, are functions of the temperature
of equilibration, and an exponential
function of the W/R ratio. Figure 9.33
-5
f
shows δ18Or plotted as a function of
i
W/R and ∆, where δ18Or is assumed to
i
0.5 -10 -4
be +6 and δ18Ow is assumed to be -13.
1
-2
1.5
0
Figure 9.34 shows another example W/R
2 Δ
of the δ18O imprint of an ancient hydro4
thermal system: the Bohemia mining
district in Lane County, Oregon, where
Tertiary volcanic rocks of the Western
Cascades have been intruded by a series of dioritic plutons. Approximately
$1,000,000 worth of gold was removed
from the region between 1870 and
1940. δ18O contours form a bull’s eye
18 f
pattern, and the region of low δ18O cor- Figure 9.33. δ Or as a function of W/R and ∆ computed
responds roughly with the area of pro- from equation 9.16.
phylitic (i.e., greenstone) alteration.
Notice that this region is broader than
the contact metamorphic aereole. The
primary area of mineralization occurs
within the δ18O < 0 contour. In this
area relatively large volumes of goldbearing hydrothermal solution, cooled,
perhaps to mixing with ground water,
and precipitated gold. This is an excellent example of the value of oxygen
isotope studies to mineral exploration.
Similar bull’s eye patterns are found
around many other hydrothermal ore
deposits.
SULFUR ISOTOPES AND ORES
Introduction
A substantial fraction of all economically valuable metal ores are sulfides.
18
district,
These have formed in a great variety of Figure 9.34. δ O variations in the Bohemia miing
18
Oregon.
Note
the
bull’s
eye
pattern
of
the
δ
O
contours.
environments and under a great variety of conditions. Sulfur isotope stud- After Taylor, 1968.
ies have been very valuable in sorting
out the genesis of these deposits. Of the various stable isotope systems we will consider in this course,
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sulfur isotopes are undoubtedly the most complex. This complexity arises in part because of
there are five common valence states in which
sulfur can occur in the Earth, +6 (e.g., BaSO4), +4
(e.g., SO2), 0 (e.g., S), –1 (e.g., FeS2) and –2 (H2S).
Significant equilibrium isotopic fractionations occur between each of these valence states. Each of
these valence states forms a variety of compounds, and fractionations can occur between
these as well. Finally, sulfur is important in biological processes and fractionations in biologically mediated oxidations and reductions are
often different from fractionations in the abio- Figure 9.35. δ34SCDT in various geologic materials
logical equivalents.
(after Hoefs, 1987).
There are two major reservoirs of sulfur on the
Earth that have uniform sulfur isotopic compositions: the mantle, which has δ34S of ~0 and in which sulfur is primarily present in reduced form, and
seawater, which has δ34S of +20 and in which sulfur
2is present as SO4 . Sulfur in sedimentary, metamorphic, and igneous rocks of the continental crust
may have δ34S that is both greater and smaller than
these values (Figure 9.35). All of these can be
sources of sulfide in ores, and further fractionation
may occur during transport and deposition of sulfides. Thus the sulfur isotope geochemistry of sulfide ores is remarkably complex.
Sulfur Isotope Fractionations in Magmatic
Processes
Figure 9.36. Fractionation of sulfur isotopes between fluid and melt (shown by dashed curves)
as a function of oxygen fugacity and temperature
for PH2O = 1 kB. Solid lines show equal concentration boundaries for quartz + magnetite ® fayalite (QM-F), H2S ® SO2, and Magnetite ® Hematite (M-H). After Ohmoto and Rye (1979).
261
Sulfur is present in peridotites as trace sulfides,
and that is presumably its primary form in the
mantle. At temperatures above about 400 °C, H2S
and SO2 are the stable forms of sulfur in fluids and
melts. In basaltic melts, sulfur occurs predominantly as dissociated H2S: HS–. It is unlikely that
significant fractionation occurs between these
forms during melting. Indeed, as we have seen, the
mean δ34S in basalts (~+0.1) is close to the value in
meteorites, which is presumably also the mantle
value. The solubility of H2S in basalt appears to be
only slightly less than that of water, so that under
moderate pressure, essentially all sulfur will remain
dissolved in basaltic liquids. The solubility of H2S
in silicate melts is related to the Fe content, decreasing with decreasing Fe.
As basalts rise into the crust, cool, and crystallize,
several processes can affect the oxidation state and
solubility of sulfur and the produce isotopic fractionations. First, the decreasing pressure results in
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Geol. 656 Isotope Geochemistry
Chapter 9
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some of the sulfide partitioning into the gas (or fluid) phase. In addition, H2 can be lost from the melt
through diffusion. This increases the ƒO2 of the melt, and as result, some of the sulfide will be oxidized
to SO2, which is very much less soluble in silicate melts than H2S. Decreasing Fe content as a consequence of fractional crystallization will also decrease the solubility of S in the melt, increasing its concentration in a coexisting fluid or gas phase. Isotope fractionation will occur between the three species
(dissolved HS–, H2S, SO2). The isotopic composition of the fluid (gas) will differ from that of the melt,
and can be computed as:
⎛ R ⎞
δ 34S fluid =δ 34Smelt − Δ HS − + Δ SO2 ⎜
⎟
⎝ R +1 ⎠
9.17
where ∆HS is the fractionation factor between HS– and H2S, ∆SO2 is the fractionation factor between H2S
and SO2. R is the molar ratio SO2/H2S and is given by:
€
R=
X SO2
X H 2S
=
Kν H 2 S ƒO3/22
Pf ν H 2O X H 2Oν SO2
9.18
where ν is the activity coefficient, Pf is the fluid pressure (generally equal to total pressure), ƒO2 is oxygen fugacity, and K is the equilibrium constant for the reaction:
3
H2S(g) + 2 O2 ® H2O(g) + SO2(g)
9.19
€
Figure 9.36 shows the sulfur isotope fractionation between fluid and melt calculated from equations
9.17 and 9.18 as a function of function of temperature and ƒO2 for PH2O = 1 kB. At the temperatures
and ƒO2 of most basalts, sulfur will be present primarily as H2S in the fluid (gas phase) and HS– in the
melt. The fractionation between these species is small (~ 0.6 ‰), so the isotopic composition of fluid
phase will not be very different that of the melt. For rhyolites and dacites, a significant fraction of the
sulfur can be present as SO2, so that greater fractionation between melt and fluid is possible.
An interesting feature of the above equations is that the fractionation between fluid and melt depends
on the water pressure. Figure 23.11 is valid only for PH2O = 1 kB. A decrease in Pf or XH2O (the mole
fraction of water in the fluid) will shift the SO2/H2S equal concentration boundary and the δ34S contours
to lower ƒO2. Conversely, an increase the in the water content will shift to boundary toward higher
ƒ O2 .
Both the eruptions of El Chichón in 1983 and Pinatubo in 1991 release substantial amounts of SO2.
The SO2–rich nature of these eruptions is thought to result from mixing of a mafic, S-bearing magma
with a more oxidized dacitic magma, which resulted in oxidation of the sulfur, and consequent increase
of SO2 in the fluid phase. There are a number of other processes that affect the solubility and oxidation
state of sulfur in the melt, and hence isotopic fractionation. Wall rock reactions could lead to either
oxidation or reduction of sulfur, crystallization of sulfides or sulfates could cause relatively small fractionations and additionally affect the SO2/H2S ratio of the fluid. Depending on the exact evolutionary
path taken by the magma and fluid, δ34S of H2S may be up to –13‰ lower than that of the original
magma. Thus variations in the isotopic composition of sulfur are possible even in a mantle-derived
magma whose initial δ34S was that of the mantle (~ 0‰). Variability of sulfur isotopic compositions do,
however, give some indication of the ƒO2 history of a magma. Constant δ34S of magmatic sulfides suggests ƒO2 remained below the SO2/H2S boundary; variability in isotopic composition suggests a higher
ƒ O2 .
Sulfur Isotope Fractionation in Low-Temperature Systems
Many important ores are sulfides. A few of these are magmatic, but most sulfide ores were deposited
by precipitation from aqueous solution at low to moderate temperature. At temperatures below about
1−
1−
1−
2−
400° C, sulfide species (H2S and HS–) are joined by sulfate (SO 4 , HSO 4 , KSO 4 , NaSO 4 , CaSO4 and
MgSO4) as the dominant forms of aqueous sulfur. The ratio of sulfide to sulfate will depend on the
262
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Geol. 656 Isotope Geochemistry
Chapter 9
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oxidation state of the fluid. Small fractionations among the various sulfide and sulfate species, but
there is a major fractionation between sulfide and sulfate. Neglecting the small fractionation between
H2S and HS– and among sulfate species, δ34SH2S of the fluid can be expressed as:
⎛ Rʹ′ ⎞
δ 34 S H 2 S = δ 34 S fluid − Δ SO 2− × ⎜
⎟
4
⎝ Rʹ′ +1 ⎠
where
9.20
Δ SO 2−4 is the fractionation between H2S and SO 24 − and R´ is the molar ratio of sulfide to sulfate:
ΣSO42−
9.21
Rʹ′ =
ΣH 2 S
€
In general, R’ will be a function of ƒO2, pH, fluid composition and temperature. Figure 9.37 shows the
difference between δ34S in sulfide and δ34S in the total fluid as a function of pH and ƒO2. Only under
conditions of low pH and low ƒO2, will the δ34S of pyrite (FeS2) be the same as the δ34S of the fluid from
which it precipitated. For conditions
of relatively high ƒO2 or high pH, substantial differences between
€
the δ34S of pyrite and the δ34S of the fluid from which it precipitated are possible. Figure 9.38 shows the
difference between δ34S in sulfide and δ34S in the total fluid as a function of the sulfate/sulfide ratio (R´)
and temperature. When the fluid is sulfide dominated, the δ34S of the sulfide and that of the bulk fluid
will necessarily be nearly identical. For conditions where the concentrations of sulfate and sulfide are
similar, large fractionations between sulfides and fluids from which they precipitate are possible.
At magmatic temperatures, reactions generally occur rapidly and most systems appear to be close to
equilibrium. This will not necessarily be the case at lower temperatures because of the strong dependence of reaction rates on temperature. While isotopic equilibration between various sulfide species
and between various sulfate species seems to be readily achieved at moderate and low temperatures,
isotopic equilibration between sulfate and sulfide appears to be more difficult to achieve. Sulfatesulfide reaction rates
-32
have been shown to
H
depend on pH (reacM
tion is more rapid at
low pH) and, in par-34
ticular, on the presence
of sulfur species of in–20
–15
termediate
valences.
M –25.2
-36
–10
Equilibration is much
Py
–25
–5
more rapid when such
intermediate valence
0
34
34
-38 δ SH2S - δ Sfluid
species are present.
Presumably, this is because reaction rates between species of adja-40
+5
cent valance states
Po M
Py
(e.g., sulfate and sulPo
fite, SO32-) are rapid,
-42
but reaction rates be3
4
5
6
7
8
9
tween widely differing
pH
valence states (e.g.,
34
sulfate and sulfide) are Figure 9.37. Difference in δ S between H2S and bulk fluid as a function of pH
and ƒO2 at 250°C. Equal concentration boundaries are shown for magnetitemuch slower.
Low
temperatures hematite (M-H), magnetite-pyrite (M-Py), magnetite-pyrrhotite (M-Po), and
also lead to kinetic, pyrite-pyrrhotite (Py-Po). After Ohmoto and Rye (1979).
log ƒO2
.
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Geol. 656 Isotope Geochemistry
Chapter 9
Spring 2011
rather than equilibrium, fractionations. As we saw, kinetic fractionation factors result from different
isotopic reaction rates. Interestingly enough, the rates for oxidation of H232S and H234S appear to be
nearly identical. This leads to the kinetic fractionation factor, αk, of 1.000±0.003, whereas the equilibrium fractionation between H2S and SO4 will be 1.025 at 250° C and 1.075 at 25° C. Thus sulfate produced by oxidation of sulfide can have δ34S identical to that of the original sulfide. Kinetic fractionations for the reverse reaction, namely reduction of sulfate, are generally larger. The fractionation observed is generally less than the equilibrium fractionation and depends on the overall rate of reduction:
the fractionation approaches the equilibrium value when reaction rate is slow. Disequilibrium effects
have also been observed in decomposition of sulfide minerals. Figure 9.39 illustrates some interesting
possible effects that can arise as a result of disequilibrium.
If there is disequilibrium between sulfate and sulfide in solution, it is likely equilibrium will not be
achieved between mineral pairs involving pyrite and chalcopyrite (CuFeS2) even when isotopic equilibrium is attained between other sulfides such as galena (PbS), sphalerite (ZnS), and pyrrhotite (FeS).
This is because precipitation of the former involves reactions such as:
4Fe2+ + 7H2S + SO 24 ® 4FeS2 + 4H2O + 6H+
9.22
whereas the latter involve only simple combinations, e.g.:
+
Zn2+ + H2S + ® ZnS + 2H
9.23
Isotopic Composition of Sulfide Ores
Figure 9.38. Difference in δ34S between H2S
and bulk fluid as a function of temperature
and sulfate/sulfide ratio in a pH neutral
fluid. Equal concentration boundaries M:
magnetite, H: hematite, Py: pyrite, Po: pyrrhotite, Bn: bornite (Cu5FeS4), Cp: chalcopyrite. After Ohmoto and Rye (1979).
A number of important sulfide deposits apparently
were produced by reduction of sulfate that was ultimately derived from seawater. The expected isotopic
compositions of such deposits are shown in Figure
9.40. The isotopic composition of these sulfides depends on the reduction mechanism, temperature, and
whether the system was open or closed to sulfate and
sulfide. At temperatures less than 50° C, the only
mechanism for reduction of sulfate is bacterial. Optimal temperatures for such reduction are around 30
to 50° C. Deep euxinic basins such as the Black Sea
are good examples of systems that are open to SO4 but
closed to H2S and where reduction is bacterial. In
these cases, reduction occurs slowly at the bottom, but
SO4 is continuously supplied from the water mass. In
such environments, sulfides appear to have a δ34S of
40 to 60‰ less than that of contemporaneous seawater (the sulfur isotopic composition of seawater has
varied through time). A good example of such a deposit is the Kupfershiefer in Germany, where the
most common δ34S is about –40‰, which is about
50‰ less than Permian seawater (+10‰).
In systems closed to SO4 but open to H2S, the process is similar to Rayleigh distillation. This would be
the case, for example, where sulfate reduction occurs
much more rapidly than sulfate is supplied to the system, but sulfide is effectively lost from the system by
crystallization of sulfide minerals. The composition
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Geol. 656 Isotope Geochemistry
Chapter 9
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of the sulfate, as a function of the fraction of sulfate remaining, ƒ, is given by:
34
δ SSO 2− = δ S
4
9.24
4
9.25
−1000(α −1)
€
1
3
c
B
2
H2S
1
SO42–
0
-1
b
SO42–
ΣH2S
t
δ 34 S Ht 2 S = δ 34 SSO
2−
a
2
−1)
The composition of the sulfide at any
time, t, is related to that of sulfate by:
€
3
..
ΣSO4
+1000( f
1−α
0
SO42−
log
34
HS
0
SO42–
-1
2
where α is the effective sulfate-sulfide
H2S
fractionation factor. The effective frac-2
-2
tionation factor for reduction may be
A
greater than the equilibrium one because of the more rapid reaction of 32S
-3-30 -20-10 0 10 20 30 -10 0 10 20 30 -10 0 10 20 30-3
species than 34S ones. A typical value of
δ34S ‰
α under these circumstances might be
1.025.
Figure 9.39. Isotopic relationships between coexisting H2S
Systems closed to both species are and SO 2 as a function of sulfate/sulfide ratio. (a) Equilib4
analogous to equilibrium crystalliza- rium conditions at 250° C where δ34S
fluid = 0 ‰. (b) Non34
tion. In either case, the δ S of both sul- equilibrium oxidation of H S. Isotopic composition of H S
2
2
fide and sulfate increases during the remains constant, but that of sulfate changes due to addition
reduction process. Systems closed to of sulfate derived from non-equilibrium oxidation of H S. (c)
2
SO4 characteristically show a spread in Non-equilibrium mixing of H S-rich fluid A and SO -rich
2
4
34
δ S that is skewed toward positive val- fluid B. δ34S remains constant (equilibrium fractionation is
H2 S
34
ues, have δ S that increases in the later achieved during reduction), but δ34S
SO4 varies due to addistages, and have both minimum and tion of sulfate derived from oxidation of
sulfide. After Ohmodal values that are approximately moto and Rye (1979).
25‰ lower than the original sulfate
(e.g., contemporaneous seawater). Examples are the White Pine and Zambian copper deposits, which
apparently formed in shallow marine or brackish environments.
At temperatures above 50° C, thermal decomposition of sulfur-bearing organic compounds produces
H2S. The δ34S values of such H2S are typically 15‰ less than that of seawater. This reduction is accelerated by the presence of S species of intermediate valence. The extent of isotopic fractionation in
this process will depend on temperature.
In ridge crest hydrothermal systems, seawater sulfate is reduced by reactions with Fe2+ such as:
2
8Fe2+ + 10H+ + SO 4 ® H2S + 8Fe3+ + 4H2O
9.26
Reduction most likely occurs at temperatures above 250° C (sulfide was not produced in basalt-seawater experiments below this temperature), and it is likely that equilibrium is achieved in this process.
Modern seawater has a δ34S of +20, and values of H2S between –5‰ and +20‰ (the latter value is a result of complete reduction of sulfate) are expected. Consistent with this prediction, sulfide in active
seafloor hydrothermal vents has a δ34S of +3.5. This process produces the class of ores referred to either
as stratabound sulfides or volcanogenic massive sulfides. Isotopic compositions of some example deposits are shown in Figure 9.41, along with the composition of contemporaneous seawater. They typically have δ34S approximately 17‰ lower than contemporaneous seawater.
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Chapter 9
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The isotopic composib34S ‰
tions of example porphyry
-60
-40
-20
0
20
copper deposits are shown
in Figure 9.42. The sulEunixic
fides in these generally
S.W.
34
Environments
have δ S between –3 and
Open to H 2S
+1‰, which is close to the
Shallow
Marine
mantle value, and the fractionation between sulfides
and Brachish
Closed to H2 S
and coexisting sulfates
Environments
suggest equilibration temBacterial Reduction (T < 50° C)
peratures between 450° C
and 650° C, which is genDecomposition of
erally in good agreement
Organic Compounds (T > 50° C)
with other temperature estimates. The isotopic comOrganic Reduction (T > 80° C)
positions indicate the sulfur was derived from igneous sources, either as Figure 9.40. Expected distribution of δ34S values for sulfide ores produced
magmatic fluids or by dis- by sulfate reduction of seawater having an isotopic composition of δ34S =
solution of igneous sul- +20‰. After Ohmoto and Rye (1979).
fides. The low δ34S of Galore Creek suggests the
oxidation state of the magma was high or that some sedimentary sulfide was incorporated. The high
δ34S of Morococha suggests some sulfur was derived from the evaporites found in the surrounding
country rock. H and O isotopic compositions in these deposits are generally inconsistent with water
being of magmatic derivation. It is possible these isotope ratios reflect overprint of meteoric water circulation after mineralization.
Mississippi Valley type deposits are
carbonate-hosted lead and zinc sulfides
formed under relatively low temperature conditions. Figure 9.43 shows the
sulfur isotope ratios of some examples.
They can be subdivided into Zn-rich
and Pb-rich classes. The Pb-rich and
most of the Zn-rich deposits were
formed between 70 and 120° C, while
some of the Zn-rich deposits, such as
those of the Upper Mississippi Valley,
were formed at temperatures up to 200°
C. Co-existing sulfides of the Pb-poor
Upper Mississippi Valley deposits are
in isotopic equilibrium, whereas sulfur
isotope equilibrium was most often not
achieved in Pb-rich deposits. In the
former, δ34S values are quite uniform
Figure 9.41. δ34S in volcanogenic massive sulfide deposits.
over a large area, suggesting the oreArrows show the isotopic composition of contemporaneforming fluid supplied both metals and
ous seawater sulfate. After Ohmoto and Rye (1979).
sulfides and transported them over
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Geol. 656 Isotope Geochemistry
Chapter 9
Spring 2011
long distances. The high positive δ34S suggests the sulfur was ultimately derived from
ancient seawater, perhaps from formation
water or evaporites in deep sedimentary basins, and reduced by reaction with organic
compounds.
δ34S is correlated with Pb isotope ratios in
galenas of the S. E. Missouri district. This,
and the variability of δ34S, suggests there was
more than one source of the sulfur and lead.
Isotopically heavy sulfur was apparently carried with the metal by the ore forming fluid
while the isotopically light sulfur was derived from pyrite in the host rock. In Pine
Point, sulfide may have been locally produced by reaction between sulfate-bearing
fluids and organic matter in sediment. Local
production of sulfide shortly before ore
deposition may help to account for the lack
of isotopic equilibrium in this deposit, since
time is an element in the attainment of isoFigure 9.42. δ34S in porphyry copper sulfides and sultopic equilibrium.
fates. After Ohmoto and Rye (1979).
Figure 9.44 illustrates a generalized model
b34S, ‰
for the genesis of Mississippi Valley type de-20
0
+20
+40
posits. In most instances, metals and sulfur
appear to have been derived distant from
BaSO4
sedimentary units, perhaps particularly from
Pine Point
FeS2 Trend toward
evaporites, by hot, deep-circulating meteoric
ZnS later stage
water. In North America, most of these seem
PbS
to have formed during or shortly after the
late
Paleozoic
Appalachian-OuchitaSardina
Marathon Orogeny. Mixing of the hot saline
fluids with low salinity ground water was
probably the immediate cause of metal preUp. Mississippi Valley
cipitation in most instances. In others, such
as Pine-Point, local reduction of sulfate in the
fluids to sulfide may have cause precipitation.
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S. E. Missouri
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267
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Chapter 9
Spring 2011
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(±H 2S)
Zn, Pb
(±H 2S)
CO2CH4
Deep Meteoric
Water
Evaporite
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