1. Cavity Ringdown Spectroscopy, instrument cost ~$90K vs. ~$400K for mass spec from Gupta, P. 2009, Chapman Conf. poster

2. Cavity Ringdown Spectroscopy, principles -idea is to compare empty chamber and full-chamber ring-down across several absorption lines -must determine unknowns against a calibration of ring-downsof known standard values from picarro.com

3. Geothermometry & paleoclimate proxies10/10/12 The JOIDES Resolution drillship

4. Temperature-dependent fractionation - recap Equilibrium fractionation is temperature-dependent, always. - we’ve discussed the liquid-vapor fractionation for precipitation - today: carbonate-liquid fractionation red=warm blue=cold; arrows track movement of 18 O through phase changes any solid phase

5. Carbonate  18 O – introduction Minerals (e.g. carbonate, quartz, barite, etc) form from super-saturated solution.  18 O of these minerals is a fxn of  18 O of solution and temperature of solution Remember: the  18 O of a solid phase is usually reported in PDB (heavy standard) while  18 O of liquid phase is usually reported in SMOW (light standard) interconversion equation: Important: You need to know the  18 O of the solution to derive temperature from  18 O solid The ocean  18 O is defined as 0‰ (Standard Mean Ocean Water), and it’s a big volume, so how do you change  18 O of seawater? Friedman and O’Neil (1977)

6. The relationship between water-  18 O, temperature, and the equilibrium  18 O of calcite was determined empirically by Sam Epstein et al., (1953) and later modified by Craig (1965): O’Neil et al. (1969) determined an experimental relationship for the temperature-dependence of  for the calcite-water system: NOTE  c must be wrt PDB,  w must be wrt SMOW good for low T, paleoceanography T in Kelvin good for high T Carbonate  18 O – temperature relationships

7. Aragonite  18 O – temperature relationships Why is the  18 O arag -water  different than the  18 O cal -water  ? Is the  larger for aragonite or calcite? Zhou & Zheng, GCA, 2003 T-dependence of  (T in Kelvin):

8. Glacial-Interglacial foraminifera  18 O, revisited Data from deep-sea (benthic) foraminifera show +1.5‰  18 O shift during LGM LGM The million-dollar question in paleoceanography: How much of this shift was due to ice volume (sea level change) and how much was due to temperature change? Schrag modelled the glacial-interglacial shift in porewater  18 O (~1.0‰), so we have 0.5‰ left over for temperature change. How much did bottom water temperatures change during the LGM? (problem set) Or you could measure temperature (trace metal concentrations in carbonates), and obtain a “residual”  18 O that gives you the  18 O SW change.

9. Complications: Kinetic effects, vital effects and carbonate  18 O Fact: very few organisms precipitate carbonate in isotopic equilibrium with the surrounding water (the vital effect) One problem: skeletons are precipitated in super-saturated “micro-environments”, with sources from surrounding water & metabolic products Another problem: isotopic exchange may be rate-limited in biological reactions Kinetic isotope effects underlie vital effect Can track kinetic effects with isotope-isotope plot (  13 C vs.  18 O), check for slope = 2

10. Spero et al., Nature 1997 Carbonate ion effect on foram δ 18 O Results from culturing living forams: increase CO 3 2-,  18 O foram decreases * Not a very big effect, but casts further doubt on inferring G-I T changes from forams

11. In order to reconstruct surface temperatures from carbonate  18 O formed during the LGM, you need to 1)remove the ice-volume effect 2)constrain the  18 O of your local water mass 3)apply the paleo-temperature equation Glacial-Interglacial climate reconstruction However, people can use other proxies to get at temperature: 1)foraminifera assemblage data (CLIMAP) 2)tree lines and snow lines will be lower during cold times 3)trace metals in carbonates (Mg/Ca in forams or Sr/Ca in corals) 4)alkenones (saturation index of long-chained alkanes in coccolithophores)

12. δ 18 O as tracer of igneous processes spectrometer light intake A “black smoker” from the East Pacific Rise Applications of oxygen isotopes in igneous rocks: 1)determine temp of formation (water-mineral or mineral-mineral pairs) 2) quantify “water-rock” ratios of altered rocks

13. composition of lunar rocks, carbonaceous chondrites, and MORB Oxygen Isotopic compositions of geological materials What scale is this? Why is the ocean so light compared to MORB? If carbonates precipitate from a “light” ocean, why are they so heavy? Lunar rocks MORB basic lavas mantle nodules eclogites andesites ophiolites rhyolites & tuffs granitic rocks altered igneous rocks metamorphic rocks clastic sediments marine limestones Why does eclogite have a heavier  18 O than MORB? Why do metamorphic rocks exhibit such a range of  18 O?

14. Principle of Geothermometry in igneous applications The fractionation of oxygen or hydrogen in different minerals of a rock can be used as a geothermometer, provided that: 1. minerals deposited at same time, at equilibrium 2. no subsequent alteration 3. fractionation factors and T-dependence known experimentally NOTE: Using multiple mineral pairs will increase confidence in the calculated temperature, if the mineral pair temperatures agree – i.e. they are concordant. Remember from last lecture we talked about the high-T water-calcite equation? General form of geothermometry fractionation equations. Handy conversions: For phases m and n: T in Kelvin

15. T-dependent fractionation in various mineral pairs Where must these lines converge? NOTE: These slopes are different – so all you need to determine T is  m-n The highest fractionation is between quartz and magnetite. In general, 18 O is increasingly favored in higher-quartz minerals, and less favored in hydrous minerals (magnetite, amphibole, chlorite). How could we determine the slope of the Quartz-Muscovite fractionation?

16. Today’s Handouts: Tables of T-dependent Fractionation

17. Example Quartz, calcite, and chlorite were all precipitated in a hydrothermal vent setting. Measured  18 O’s: Quartz: 5.1‰ SMOW Calcite: 3.8‰ SMOW Chlorite: -1.5‰ SMOW Why does quartz have the heaviest  18 O, and chlorite the lightest? And why is the quartz only 5.1‰ heavier than SMOW? Did these minerals precipitate at the same temperature? How would you begin to solve this problem?

18. Metamorphism: Water-Rock interactions Fact: In several places it is possible to measure igneous rocks with  18 O values of -5‰! These are places were fluid has interacted with the rock (usually at high T) to change the isotopic composition of the rock. We can use a mass balance approach to calculate the amount of water that has reacted with a host rock (or “water/rock ratio”) over time (assuming equilibrium): is the equilibrium values for water and mineral, (need to know temperature independently) and mass balance equation: cw = conc. of O in water cr = conc. of O in rock W = mass water R= mass rock superscript I = initial superscript f = final and combining these: for a closed system What does a “closed system” mean?

19. Water-Rock interactions II for an open system Now we only have a small parcel of water (dW) interacting at any given time, but new water parcels are injected continuously in time, causing d  r in this scenario we need to integrate to calculate W/R ratios. Probably much more realistic, because water flows through the rock. In order to solve for W/R interactions, you need to know: 1.the temperature of the interaction (hopefully you can get that by a mineral-mineral pair) 2.the mineral phases that experienced fluid alteration 3.the isotopic composition of the water before it interacted with the rock (  D of rock… why?) 4.the isotopic composition of the rock before it interacted with the water (unaltered samples)

20. A characteristic signature of hydrothermal activity is a “bulls-eye” pattern of  18 O values, with low values in the middle. Alteration occurs along an established conduit of weakened structures. Most gems are the product of low-T, high-fluid metamorphism – $1M worth of gold mined in the Bohemia complex between 1870 and 1940 - happy hunting! A real-world example

21. A Cool Early Earth, (2002) Geology. 30: 351-354. A cool early Earth? Idea: measure U-Pb dates and  18 O of old zircons - if you find low  18 O relative to today’s ‘primitive’ mantle, then that implies interaction with meteoric waters at low temperatures This work is done using a laser flourination line plumbed to a dual inlet mass spec

22. So what’s causing the relatively high zircon  18 O values at 4.2Ga? time-line along bottom indicates: (1)accretion of the Earth, (2)formation of the Moon and the Earth’s core, (3)minimum age of liquid water based on high  18 O zircon, (4)Acasta gneiss, and (5)Isua metasedimentary rocks.

23. likely interaction with meteoric waters which implies a period of few impacts