1. Stable Isotope Geochemistry III
Lecture 32

2. The Antarctic Ice Record
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.

3. 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.

4. 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 60˚N 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

5. The Next Ice Age?
From Marcott et al. (2013) Science, 339: 1198

6. 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.

7. 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.

8. Stable Isotopes in High Temperature Geochemistry

9. Where does the water come from?

10. 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 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 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,

11. 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.

12. 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 δ18O will be associated with maximum W/R.

13. Example: Lane Co., Oregon
Low δ18O in rocks, reflecting water/rock ratios, forms a bullseye around main area of mineralization and economic gold deposit.

14. δ18O in Hydrothermal Systems
Because of the temperature dependence of fractionation, the effect of water-rock interaction at low and high temperature can be quite different.
As seawater is heated, it exchanges O with the surrounding rock. At temperatures in the range of ° 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

15. Sulfur Isotopes
Many ores are sulfides and sulfur isotopes provide important clues to their genesis, including temperatures of deposition.
Overview of δ34S:
Mantle, bulk Earth value ~0 (same as meteorites)
modern seawater is +20 (has varied over Earth’s history with δ13C).
Sedimentary sulfide, generally the result of bacterial sulfide reduction, can have δ34S as low as -40.

16. 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.

17. Archean MIF Sulfide
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.
∆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.

18. Stable Isotopes in the Mantle and Magmas

19. 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.
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.

20. Hydrogen in the Mantle
Mantle sample restricted in hydrous minerals in xenoliths and submarine erupted basalts.
Mean δD in solid Earth is about -70‰.
Some variation in the mantle, but hard to pin down, partly because of fractionation during degassing.

21. 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 mantle-derived, not from sediments.
A subclass of diamonds, those with an eclogitic paragenesis, have much lighter carbon, with peak around δ13C ≈ -25‰.
This carbon was likely organic in origin and was anciently subducted into the mantle.

22. 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:
In both theory and observation, there will be not much more than 1 or 2‰ change in δ18O.

23. 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, 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 in book

24. 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: δ11B ~ -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.
Fractionation, mainly as a result of adsorption of light B on clays, drives seawater to extreme isotopic composition.

25. 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:
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?
Note error in book.

26. Seawater pH and Atmospheric CO2 from δ11B
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 ; Figure 12.43). Atmospheric CO2 conc (397 ppm) is now higher than it has been for 35 million years.