Re-Os & U-Th-Pb
Isotope
Geochemistry
Lecture 32
The Re-Os System
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187Re
decays to 187Os by β– decay with a half-life of 42 billion
years.
Unlike the other decay systems of geological interest, Re and
Os are both siderophile elements: they are depleted in the
silicate Earth and presumably concentrated in the core. The
resulting very low concentration levels (sub-ppb) make analysis
extremely difficult. Interest blossomed when a technique was
developed to analyze OsO4– with great sensitivity. It remains
very difficult to measure in many rocks, however. Peridotites
have higher concentrations.
The siderophile/chalcophile nature of these elements, making
this a useful system to address questions of core formation and
ore genesis.
Os is a highly compatible element (bulk D ~ 10) while Re is
moderately incompatible and is enriched in melts. For
example, mantle peridotites have average Re/Os close to the
chondritic value of 0.08 whereas the average Re/Os in basalts
is ~10. Thus partial melting appears to produce an increase in
the Re/Os ratio by a factor of >102. As a consequence, the
range of Os isotope ratios in the Earth is very large. The mantle
has a 187Os/188Os ratio close to the chondritic value of, whereas
the crust appears to have a a 187Os/188Os > 1. By contrast, the
difference in 143Nd/144Nd ratios between crust and mantle is
only about 0.5%.
The near chondritic a 187Os/188Os of the mantle is surprising,
given that Os and Re should have partitioned into the core
very differently. This suggests most of the noble metals in the
silicate Earth are derived from a late accretionary veneer
added after the core formed.
In addition, 190Pt decays to 186Os with a half-life of 650 billion
years. The resulting variations in 186Os/188Os are small.
Os Isotopes in the SCLM
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Since the silicate Earth appears to have a nearchondritic 187Os/188Os ratio, it is useful to define a
parameter analogous to εNd and εHf that measures
the deviation from chondritic. γOs is defined as:
- ( Os
Os )
Os )
g
´100
Os
( Os)
Studies of pieces of subcontinental lithospheric
Os
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(
=
187
Os 188
187
sample
187
188
188
Chond
Chond
mantle xenoliths show that much of this mantle is
poor in clinopyroxene and garnet and hence
depleted in its basaltic component. Surprisingly,
these xenoliths often show evidence of
incompatible element enrichment, including high
87Sr/86Sr and low ε . This latter feature is often
Nd
attributed to reaction of the mantle lithosphere with
very small degree melts percolating upward
through it (a process termed “mantle
metasomatism”).
This process, however, apparently leaves the Re-Os
system unaffected, so that 187Re/188Os and
187Os/188Os remain low.
Low γOs is a signature of lithospheric mantle.
Os Isotopes in Seawater
• Os isotopes in seawater
(tracked by measuring Os in
Mn nodules and black
shales) reveals a variation
much like that of 87Sr/86Sr.
• The reflects a balance of
mantle and crustal inputs.
• And, perhaps, meteoritic
ones. Very low ratios occur
at the K-T boundary. Ratio
was already decreasing
before then: Deccan traps
volcanism? (supports the hit
‘em while their down theory
of the K-T extinction).
U-Th-Pb
• In the U-Th-Pb system there are three decay schemes
producing 3 isotopes of Pb. Two U isotopes decay to 2
Pb isotopes, and since the parent and daughter isotopes
are chemically identical, we get a particularly powerful
tool.
• Following convention, we will designate the 238U/204Pb
ratio as μ, and the 232Th/238U ratio as κ. We can write two
versions of our isochron equation:
Pb æ
=
204
Pb çè
206
206
204
Pb ö
+ µ(el238t -1)
÷
Pb ø 0
235
Pb æ 207 Pb ö
U
= ç 204 ÷ + µ 238 (el238t -1)
204
Pb è Pb ø 0
U
207
o Conventionally, the 235U/238U was assumed to have a constant, uniform value
of 1/137.88. Recent studies, however, have demonstrated that this ratio varies
slightly due to kinetic chemical fractionation. Consequently, for highest
precision, it should be measured. In most cases, however, we can use the
revised apparent average value of 1/137.82.
Pb-Pb isochrons
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These equations can be rearranged
by subtracting the initial ratio from
both sides. For example:
Pb
= µ(el238t -1)
204
Pb
206
∆
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Dividing the two:
∆ 207 Pb / 204 Pb
=
∆ 206 Pb / 204 Pb
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U (el238t -1)
238
U (el235t -1)
235
the 235U/238U is the present day ratio and assumed
constant.
The left is a slope on a plot of
207Pb/204Pb vs 206Pb/204Pb. Slope is
proportional to time, and so is an
isochron.
The value is that we need not know or
measure the U/Pb ratio (which is
subject to change during
weathering).
In essence, we have 3 dating
methods: 235U-207Pb, 238U-206Pb and PbPb. An age is said to be concordant
when all 3 agree.
Pb Isotopic Evolution
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Because the half-life of 235U is much shorter
than that of 238U, 235U decays more rapidly
and Pb isotopic evolution follows curved
paths on this plot.
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The exact path depends upon µ.
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True only for the planet as a whole, not individual rock
formations.
Parts of the planet can lie to the right if U/Pb increased or
the the left if U-Pb decreased.
All systems that begin with a common initial
isotopic composition at time t0 lie along a
straight line today. This line is the Pb-Pb
isochron.
When the solar system formed 4.57 billion
years ago, it had a single, uniform Pb
isotope composition. All closed systems,
such as planets, that started with this value
at that time must lie on a 4.57 Ga isochron,
called the Geochron.
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The Earth as a whole must fall on this line if it
formed at the same time as the solar system
with the solar system initial Pb isotopic
composition.
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The problem is that Earth may be 100 Ma younger than the
‘solar system’ - because it took a long time to form large
terrestrial planets.
There is some flexibility in the exact position of the Geochron
because the age is not exactly known.
232Th-208Pb
• We can combine the growth equations for
208Pb/204Pb and 206Pb/204Pb in a way similar to our
207Pb-206Pb isochron equation We end up with:
∆ 208 Pb / 204 Pb
(el238t -1)
= k l235t
∆ 206 Pb / 204 Pb
(e -1)
o where κ is the 232Th/238U ratio.
• The left is a slope on a plot of 208Pb/204Pb vs
206Pb/204Pb and is proportional to t and κ.
o assuming κ has been constant (except for radioactive decay).
Pb Isotope Ratios in the Earth
Pb Isotope Ratios in the Earth
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Major terrestrial reservoirs, such as the upper mantle
(represented by MORB), upper and lower
continental crust, plot near the Geochron between
growth curves for µ = 8 and µ = 8.8, suggesting µ of
the Earth ≈ 8.5.
If a system has experienced a decrease in U/Pb at
some point in the past, its Pb isotopic composition
will lie to the left of the Geochron; if its U/Pb ratio
increased, its present Pb isotopic composition will lie
to the right of the Geochron.
U is more incompatible than Pb, so incompatible
element depleted reservoirs should plot to the left
of the Geochron, enriched ones to the right.
From the other isotopic ratios, we would have
predicted that continental crust should lie to the
right of the Geochron and the mantle to the left.
Surprisingly, Pb isotope ratios of mantle-derived
rocks also plot mostly to the right of the Geochron.
This indicates the U/Pb ratio in the mantle has
increased, not decreased as expected.
This phenomenon is known as the Pb paradox and
it implies that a simple model of crust–mantle
evolution that involves only transfer of incompatible
elements from crust to mantle through magmatism
is inadequate.
There is also perhaps something of a mass balance
problem - since everything should average out to
plot on the Geochron.