1. High Temperature Stable Isotope Geochemistry Lecture 38

2. Where does the water come from?

3. 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 δ 18 O is shifted to higher values. The shift in δ 18 O 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,

4. 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.

5. 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 δ 18 O will be associated with maximum W/R.

6. Example: Lane Co., Oregon Low δ 18 O in rocks, reflecting water/rock ratios, forms a bullseye around main area of mineralization and economic gold deposit.

7. Sulfur Isotopes Many ores are sulfides and sulfur isotopes provide important clues to their genesis, including temperatures of deposition. Overview of δ 34 S: o Mantle, bulk Earth value ~0 (same as meteorites) o modern seawater is +20 (has varied over Earth’s history with δ 13 C). o Sedimentary sulfide, generally the result of bacterial sulfide reduction, can have δ 34 S as low as -40.

8. 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.

9. Archean MIF Sulfide Most studies report only 34 S/ 32 S asδ 34 S, but sulfur has two other isotopes 33 S and 36 S. We expect δ 33 S, δ 34 S, and δ 36 S to all correlate strongly, and they almost always do (hence few bother to measure 33 S or 36 S). When Farquhar measured δ 33 S and δ 34 S in Archean sulfides, he found mass independent fractionations. o ∆ 33 S is the permil deviation from the expected δ 33 S based on measured δ 34 S. Experiments show that SO 2 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 SO 2. 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.

10. Stable Isotopes in the Mantle and Magmas

11. Oxygen in the Mantle δ 18 O 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 Igneous rocks with δ 18 O very different from ~5.6‰ show evidence of low-T surface processing. o At high-T, δ 18 O isotopes can effectively be used as tracers like radiogenic isotopes.

12. 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.

13. Carbon in the Mantle MORB and submarine erupted OIB have δ 13 C of close to -6‰. Most diamonds have similar δ 13 C, with average around - 5‰. Carbonatites have the same δ 13 C, 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 δ 13 C ≈ -25‰. o This carbon was likely organic in origin and was anciently subducted into the mantle.

14. δ 18 O 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 δ 18 O.

15. 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 δ 18 O. Note error in equ. 9.69 in book

16. Boron Isotopes Stable isotope geochemistry has been expanding beyond the traditional isotopes. The large mass difference between 10 B and 11 B results in large fractionations. Fractionation is mainly between trigonal (e.g., BOH 3 ) and tetrahedral (e.g., BOH 4 – ) forms. o Both forms in seawater. o Mainly borate (BO 3 ) in boron minerals like tourmaline; BOH 4 - in clays, probably substitutes for tetrahedral Si in other silicates. Mantle, chondrites, most basalts: δ 11 B ~ -5‰. Variable in crustal rocks and sediments. Island arc volcanics are heavier - evidence of a fluid or seawater component. δ 11 B = +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.

17. 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 + H 2 O ⇋ 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: δ 11 B SW = δ 11 B 3 ƒ + δ 11 B 4 (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: δ 11 B SW = δ 11 B 3 ƒ + δ 11 B 4 - δ 11 B 4 ƒ = δ 11 B 4 - ∆ 3-4 ƒ Solving for δ 11 B 4, we have: δ 11 B 4 = δ 11 B SW + ∆ 3-4 ƒ δ 11 B 4 depends on ƒ, which depends on pH. Boron is incorporated into carbonate by surface adsorption of B(OH) 4 -. Thus the δ 11 B in carbonates tracks δ 11 B 4, which in turn depends on pH, assuming δ 11 B in seawater is constant. What will pH of seawater depend on? Note error in book.

18. Seawater pH and Atmospheric CO 2 from δ 11 B Pearson and Palmer (2000) measured δ 11 B in foraminifera from (ODP) cores and were able to reconstruct atmospheric CO 2 through much of the Cenozoic. Surprisingly, atmospheric CO 2 has been < 400 ppm through the Neogene, a time of significant global cooling. Much higher CO 2 levels were found in the Paleogene. This has largely been confirmed by another paleo-CO 2 proxy, δ 13 C in 37-C diunsaturated alkenones (Section 12.8.2; Figure 12.43). Atmospheric CO 2 conc (397 ppm) is now higher than it has been for 35 million years.