Download 4550-15Lecture32 - Earth and Atmospheric Sciences

Stable Isotope
Geochemistry III
Lecture 32
The Antarctic Ice Record
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Much subsequent
paleoclimate effort has
focused on δD in ice cores
from Antarctica and
Greenland.
The Vostok core from
Antarctica went back 400 ka.
Subsequent work shifted to
the EPICA core which went
back >800 ka.
Complications in
interpretation arise here too
because of changes in δD of
the oceans and changes in
atmospheric circulation result
in complex relationship
between T and δD, but
temperatures can be worked
out.
Overall, agreement between
the marine and Antarctic
records is excellent, but shows
some differences between
Antarctic and global climate
change.
Greenland Ice Record
• Ice records from
Greenland are not as
long, but provide finer
details of the last glacial
cycle.
o Greenland is ‘ground zero’ of
glaciation.
• They reveal extremely
variable climate in the last
ice age -DansgaardOeschager events - likely
related to iceberg events
documented in deep-sea
cores.
Feedback Factors
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Milankovitch variations
provide only a weak climate
signal that has been
apparently greatly amplified
in the Quaternary by
feedback factors.
June insolation at 60˚N
appears to be the key
sensitivity.
Feedbacks include:
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Albedo
Shift of CO2 from atmosphere to
oceans with consequent change in
greenhouse effect
Changes in ocean circulation,
particularly with delivery of heat to the
North Atlantic (ground zero for
continental ice sheets).
The role of CO2 is well
documented by CO2
concentrations in bubbles in
Antarctic ice.
Figure 12.45
The Next Ice Age?
From Marcott et al. (2013) Science, 339: 1198
Soil Paleoclimate Proxies
• Hydrogen and Oxygen
isotopes in soil clays
reflect (with
fractionation), the
isotopic composition of
meteoric water.
• This allows
reconstruction of
paleoprecipitation
patterns - Cretaceous
precipitation in N.
America in this figure.
Pedogenic Carbonate
• δ18O in pedogenic
carbonate also reflects
composition of
meteoric water (with
fractionation).
• In Pakistan, δ18O in
paleosol carbonates
record the evolution of
the monsoons.
Stable Isotopes in High
Temperature Geochemistry
Where does the water come
from?
Hydrothermal Systems
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One of the first of many important
contributions of stable isotope
geochemistry to understanding
hydrothermal systems was the
demonstration by Harmon Craig
(another student of Harold Urey) that
water in these systems was meteoric,
not magmatic.
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 hightemperature reaction (≲300°C) of the
local meteoric water with hot rock.
Acidic, sulfur-rich waters from
hydrothermal systems can have δD
that is different from local meteoric
water. This shift occurs when
hydrogen isotopes are fractionated
during boiling of geothermal waters.
The steam mixes with cooler meteoric
water, condenses,
Importance of Hydrothermal
Systems
• Hydrothermal systems
are the source of many
ore deposits, including
base metals (Pb, Zn,
Cu), gold, tin, and
many others.
• Hydrothermal activity is
also important in the
chemistry of the
oceans, the oceanic
crust, and the plate
tectonic cycle.
Water-rock ratios
• For a closed system:
D = d wf - d rf
• from which we can derive:
W
d rf - d ri
c
= i
´ r
f
R d w - d w - ∆ cw
• For an open system in which
water makes 1 pass through
the rock we start with
Rcr dd r = (d wi -[∆ + d r ])cw dW
• and derive:
æ d ƒ -d i
öc
W
= ln ç ƒ r i r +1÷ r
R
è -d r + d w - ∆ ø cw
• Point is that maximum
change in δ18O will be
associated with maximum
W/R.
Example: Lane Co., Oregon
• Low δ18O in rocks,
reflecting water/rock
ratios, forms a bullseye
around main area of
mineralization and
economic gold
deposit.
δ18O in Hydrothermal Systems
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Because of the temperature
dependence of fractionation, the effect
of water-rock interaction at low and high
temperature can be quite different.
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As seawater is heated, it exchanges O
with the surrounding rock. At
temperatures in the range of 300-400° C,
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.
•
At low-temperature fractionations are
quite large, ~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.
ODP Site 1256, Eastern Pacific
Sulfur Isotopes
• Many ores are sulfides
and sulfur isotopes
provide important clues
to their genesis, including
temperatures of
deposition.
• Overview of δ34S:
o Mantle, bulk Earth value ~0
(same as meteorites)
o modern seawater is +20 (has
varied over Earth’s history with
δ13C).
o Sedimentary sulfide, generally the
result of bacterial sulfide
reduction, can have δ34S as low
as -40.
Mississippi Valley Sulfide
Deposits
• Mississippi Valley type
Pb-Zn deposits are
sediment-hosted (often
carbonate) sulfides
deposited from low-T
hydrothermal solutions.
• Source of sulfide is
generally formation
brine or evaporite
sulfate (of ultimate
seawater origin) that is
subsequently reduced.
Archean MIF Sulfide
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Most studies report only 34S/32S as δ34S,
but sulfur has two other isotopes 33S and
36S.
We expect δ33S, δ34S, and δ36S to all
correlate strongly, and they almost
always do (hence few bother to
measure 33S or 36S).
When Farquhar measured δ33S and δ34S
in Archean sulfides, he found mass
independent fractionations.
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∆33S is the permil deviation from the expected δ33S
based on measured δ34S.
Experiments show that SO2
photodissociated by UV light can be
mass-independently fractionated.
Interpretation: prior to 2.3 Ga, UV light
was able to penetrate into the lower
atmosphere and dissociate SO2. In the
modern Earth, stratospheric ozone
restricts UV penetration into the
troposphere(sulfur rarely reaches the
stratosphere, so little MIF fractionation).
This provides strong supporting evidence
for the Great Oxidation Event (GOE) at
2.3 Ga.
Stable Isotopes in the
Mantle and Magmas
Oxygen in the Mantle
• δ18O in olivine in peridotites
is fairly uniform at +5.2‰.
• Clinopyroxenes slightly
heaver, ~+5.6‰.
• Fresh MORB are typically
+5.7‰
• Some OIB and IAV show
deviations from this.
• Bottom line: no more than
tenths of per mil
fractionations at high T.
o
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Igneous rocks with δ18O very
different from ~5.6‰ show
evidence of low-T surface
processing.
At high-T, δ18O isotopes can
effectively be used as tracers like
radiogenic isotopes.
Hydrogen in the Mantle
• Mantle sample
restricted in hydrous
minerals in xenoliths
and submarine erupted
basalts.
• Mean δD in solid Earth
is about -70‰.
o Some variation in the mantle,
but hard to pin down, partly
because of fractionation
during degassing.
Carbon in the Mantle
• MORB and submarine
erupted OIB have δ13C of
close to -6‰.
• Most diamonds have similar
δ13C, with average around 5‰.
• Carbonatites have the
same δ13C, indicating the
carbonate is mantlederived, not from sediments.
• A subclass of diamonds,
those with an eclogitic
paragenesis, have much
lighter carbon, with peak
around δ13C ≈ -25‰.
o
This carbon was likely organic in
origin and was anciently subducted
into the mantle.
δ18O in Crystallizing Magmas
• Fractionations between
silicates and silicate
magmas are small, but they
can be a bit larger when
oxides like magnetite and
rutile crystalize.
• We imagine two paths:
equilibrium and fractional,
the latter more likely.
• For fractional crystallization:
∆ = 1000( f a -1)
• In both theory and
observation, there will be
not much more than 1 or
2‰ change in δ18O.
Fractional CrystallizationAssimilation
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Magmas intruding the crust can
melt and assimilate crust
(because the magmas are hotter
than the melting temperature)
Energy to melt comes largely
from the ∆H of crystallization,
hence crystallization and
assimilation will be linked.
If R is the ratio of mass assimilated
to mass crystallized, the isotope
ratio will change as:
æ
∆ö
R
d m - d 0 = ç [d a - d 0 ] + ÷ {1- f - R/( R-1) }
è
ø
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where subscripts m, 0, and a refer
to the isotopic composition of
the magma, the original magma,
and the assimilant, ƒ is fraction of
liquid remaining and ∆ is
crystal/liquid fractionation factor.
This can lead to much larger
change in δ18O.
Note error in equ.
9.69 in book
Boron Isotopes
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Stable isotope geochemistry has
been expanding beyond the
traditional isotopes.
The large mass difference between
10B and 11B results in large
fractionations.
Fractionation is mainly between
trigonal (e.g., BOH3) and tetrahedral
(e.g., BOH4–) forms.
Both forms in seawater.
Mainly borate (BO3) in boron minerals like tourmaline;
BOH4- in clays, probably substitutes for tetrahedral Si
in other silicates.
Mantle, chondrites, most basalts: δ11B
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~ -5‰. Variable in crustal rocks and
sediments. Island arc volcanics are
heavier - evidence of a fluid or
seawater component.
δ11B = +39‰ in seawater. Seawater is
heavier than anything else.
o
Fractionation, mainly as a result of adsorption of light
B on clays, drives seawater to extreme isotopic
composition.
Boron in the Ocean &
Carbonates
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Boron is present in seawater both as B(OH)3, and B(OH)4-. The reaction between
them is:
B(OH)3 + H2O ⇋ B(OH)4- + H+
Note error in
The relative abundance of these two species depends on pH
B(OH )-4
log
= pK app + pH
B(OH )3
book.
The isotopic composition of these two species must vary with pH if the isotopic
composition of seawater is constant.
From mass balance we have:
δ11BSW = δ11B3ƒ + δ11B4(1-ƒ)
where ƒ is the fraction of B(OH)3 If the isotopic compositions of the two species are
related by a constant fractionation factor, ∆3-4, then:
δ11BSW = δ11B3ƒ + δ11B4 - δ11B4ƒ = δ11B4 - ∆3-4ƒ
Solving for δ11B4, we have:
δ11B4 = δ11BSW + ∆3-4ƒ
δ11B4 depends on ƒ, which depends on pH.
Boron is incorporated into carbonate by surface adsorption of B(OH)4-. Thus the δ11B
in carbonates tracks δ11B4, which in turn depends on pH, assuming δ11B in seawater
is constant.
What will pH of seawater depend on?
Seawater pH and
11
Atmospheric CO2 from δ B
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Pearson and Palmer (2000)
measured δ11B in foraminifera
from (ODP) cores and were able
to reconstruct atmospheric CO2
through much of the Cenozoic.
Surprisingly, atmospheric CO2 has
been < 400 ppm through the
Neogene, a time of significant
global cooling. Much higher CO2
levels were found in the
Paleogene.
This has largely been confirmed
by another paleo-CO2 proxy, δ13C
in 37-C diunsaturated alkenones
(Section 12.8.2; Figure 12.43).
Atmospheric CO2 conc (397 ppm)
is now higher than it has been for
35 million years.