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Stable Isotope Geochemistry III
Lecture 32 The Antarctic Ice Record
Much subsequent paleoclimate effort has focused on D 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 D of the oceans and changes in atmospheric
circulation result in complex relationship between T and D, 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. Greenland is ground zero
of glaciation. They reveal extremely variable climate in the last
ice age -Dansgaard-Oeschager events - likely related to iceberg
events documented in deep-sea cores. Feedback Factors Milankovitch
variations provide only a weak climate signal that has been
apparently greatly amplified in the Quaternary by feedback factors.
June insolation at 60N appears to be the key sensitivity. Feedbacks
include: 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 18O
in pedogenic carbonate also reflects composition of meteoric water
(with fractionation). In Pakistan, 18O in paleosol carbonates
record the evolution of the monsoons. Stable Isotopes in High
Temperature Geochemistry Where does the water come from?
Hydrothermal Systems 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 D of the chloride
type geothermal waters is the same as the local precipitation and
groundwater, but the 18O is shifted to higher values. The shift in
18O results from high-temperature reaction (300C) of the local
meteoric water with hot rock. Acidic, sulfur-rich waters from
hydrothermal systems can have D 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: from which we can derive:
For an open system in which water makes 1 pass through the rock we
start with and derive: Point is that maximum change in 18O will be
associated with maximum W/R. Example: Lane Co., Oregon
Low 18O in rocks, reflecting water/rock ratios, forms a bullseye
around main area of mineralization and economic gold deposit. 18O
in Hydrothermal Systems
Because of the temperaturedependence of fractionation, the effectof
water-rock interaction at low and hightemperature can be quite
different. As seawater is heated, it exchanges Owith the
surrounding rock. Attemperatures in the range of C,the net
water-rock fractionation is small,1 or 2. Thus isotopic exchange
results ina decrease in the 18O of the rock andan increase in the
18O of the water. At low-temperature fractionations arequite large,
~20. The result of thesereactions is to increase the 18O of
theshallow oceanic crust and decrease the18O of seawater. Thus the
effects of low temperature andhigh temperature reactions
areopposing. 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
34S: Mantle, bulk Earth value ~0 (same as meteorites) modern
seawater is +20 (has varied over Earths history with 13C).
Sedimentary sulfide, generally the result of bacterial sulfide
reduction, can have 34S 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 Most studies report only 34S/32Sas 34S, but sulfur has
two other isotopes 33S and 36S. We expect 33S, 34S, and 36S to all
correlate strongly, and they almost always do (hence few bother to
measure 33S or 36S). When Farquhar measured 33S and 34S in Archean
sulfides, he found mass independent fractionations. 33S is the
permil deviation from the expected 33S based on measured 34S.
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 18O
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. Igneous rocks with 18O very
different from ~5.6 show evidence of low-T surface processing. At
high-T, 18O 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 D in solid Earth is about -70. 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 13C of close to -6. Most diamonds have similar 13C, with
average around -5. Carbonatites have the same 13C, indicating the
carbonate is mantle-derived, not from sediments. A subclass of
diamonds, those with an eclogitic paragenesis, have much lighter
carbon, with peak around 13C -25. This carbon was likely organic in
origin and was anciently subducted into the mantle. 18O 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: In both theory
and observation, there will be not much more than 1 or 2 change in
18O. Fractional Crystallization-Assimilation
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:
where subscripts m, 0, and a refer to the isotopic composition of
the magma, the original magma, and the assimilant, is fraction of
liquid remaining and is crystal/liquid fractionation factor. This
can lead to much larger change in 18O. Note error in equ in book
Boron Isotopes 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: 11B ~ -5. Variable in crustal rocks and sediments. Island
arc volcanics are heavier - evidence of a fluid or seawater
component. 11B = +39 in seawater. Seawater is heavier than anything
else. Fractionation, mainly as a result of adsorption of light B on
clays, drives seawater to extreme isotopic composition. Boron in
the Ocean & Carbonates
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+ The relative
abundance of these two species depends on pH The isotopic
composition of these two species must vary with pH if the isotopic
composition of seawater is constant. From mass balance we have:
11BSW = 11B3 + 11B4(1-) where 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: 11BSW = 11B3 + 11B4 - 11B4 = 11B4
- 3-4 Solving for 11B4, we have: 11B4 = 11BSW + 3-4 11B4 depends on
, which depends on pH. Boron is incorporated into carbonate by
surface adsorption of B(OH)4-. Thus the 11B in carbonates tracks
11B4, which in turn depends on pH, assuming 11B in seawater is
constant. What will pH of seawater depend on? Note error in book.
Seawater pH and Atmospheric CO2 from 11B
Pearson and Palmer (2000) measured 11B 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, 13C in 37-C
diunsaturated alkenones (Section ; Figure 12.43). Atmospheric CO2
conc (397 ppm) is now higher than it has been for 35 million
years.