1. CARBON SINKS (from the surface of the earth) WEATHERING OF ROCKS (pulls CO 2 out of the atmosphere) RIVERS – transports soluble carbon to ocean. OCEAN / PHYTOPLANKTON (converts soluble carbon to insoluble solids – cell walls and fecal pellets). Which then fall to the seafloor. SEDIMENTS (carbon is buried and temporarily out of the loop). CARBON RECYCLED SEDIMENTS ARE SUBDUCTED, ORGANIC COMPOUNDS BROKEN DOWN BY HEAT AS OCEAN SLAB IS CARRIED INTO MANTLE, CO 2 EMITTED BY SUBDUCTION VOLCANOS.

2. What is an isotope? Same element with the same number of protons, but with a different numbers of neutrons:

3. Stable isotope abundances Out of every 100 atoms of Oxygen, 0.2 atoms would be 18 O and the rest would be 16 O.

4. Fractionation The partitioning of stable isotopes of an element among different coexisting phases is called FRACTIONATION and is a MASS and TEMPERATURE dependent process Fractionation leads to variation in the natural abundances of stable isotopes expressed as differences in ISOTOPE RATIOS, R ALWAYS: R = HEAVY ISOTOPE/ LIGHT ISOTOPE THAT IS: R = RARE ISOTOPE / ABUNDANT ISOTOPE e.g. D/H, 13 C/ 12 C, 15 N/ 14 N, 18 O/ 16 O, 34 S/ 32 S

5. Definitions - ambiguity “ 18 O-rich” “ 18 O-poor” “heavy oxygen” “light oxygen” “enriched oxygen” “depleted oxygen” “ 16 O-poor” “ 16 O-rich” Because these ratios are so small, chemists measure 18 O/ 16 O (=R), rather than 18 O or 16 O abundance And then report them as ratios compared to a standard.

6. R = R 0 * f (  -1) Rayleigh fractionation R = isotope ratio in diminishing reservoir  = isotope fractionation factor The isotope ratio (R) in a diminishing reservoir of a reactant is a function of the initial isotope ratio (R 0 ), the remaining fraction of the reservoir (f) and the fractionation factor (  )

7. Reservoir containing both 18 O and 16 O atoms 18 O 16 O 18 O 16 O Remove 18 O and 16 O atoms at a different ratio than the initial reservoir 18 O 16 O That changes the ratio of 18 O and 16 O in the original reservoir. THIS process is temperature dependent. Change the temperature and you extract different isotope ratios from the original reservoir. Biology (forams)Original seawater Modified seawater

8. OXYGEN ISOTOPES AS A PROXY FOR PALEOTEMPERATURE There are two stable isotopes of oxygen used in paleotemperature estimates: 16 O (about 99.8% of total) and 18 O (most of the rest). There are other oxygen isotopes, but they are not used for paleotemperatures. The ‘normal’ ratio of 18 O/ 16 O is about 1/400, so when we express variations in this ratio, it is usually multiplied by a large number (1000), so the values are small whole numbers. Define the  18 O ratio as… [note: heavy isotope over light isotope, always] Where ( 18 O/ 16 O) SMOW is a sample of surface ocean where  18 O = 0.

9. The ratio Oxygen isotopes 16 O and 18 O are used a proxy to obtain paleotemperatures in two main environments; 1.From the oxygen obtained from calcium carbonate shells of foraminifera in oceanic sediments, and 2.From the oxygen obtained from ice in Arctic and Antarctic ice cores. In the foram shells in sediments, the ratio of 16 O and 18 O in the carbonate records that ratio that is present in seawater, modified by the temperature of the sea water (through fractionation of the 18 O and 16 O isotopes). The 16 O and 18 O ratio of seawater also depends on the volume of ice sheets that are present on the surface of the earth.

10. TEMPERATURE FRACTIONATION OF OXYGEN ISOTOPES 18 O AND 16 O Planktonic foraminifera live in the upper 100 meters of the ocean. In the PRESENT DAY ocean, surface seawater has a  18 O near 0 (zero). And, like carbon, biology fractionates this oxygen isotope ratio (critters ‘like’ the light isotope better) during metabolism. BUT the amount of this fractionation is TEMPERATE DEPENDENT. Both LAB and FIELD studies show that the temperature dependence of this BIOLOGICAL fractionation is 1 0 / 00  18 O decrease for each 4.2°C increase in water temperature. or… 18 O becomes less abundant in the foram carbonate shells - with respect to 16 O - when the temperature increases.

11. But this oxygen isotope paleo-thermometer has a major complication – the amount of ice on the continents. The formation of large ice caps changes the  18 O ratio of seawater! Examples. Tropical plantonic foram shells that grow near 21°C have a  18 O of about -1 0 / 00. (lower than the seawater value). But benthic forams living in the deep ocean (near 2°C) have a  18 O value of about +5 0 / 00 (higher than the ambient seawater). Paleoclimate scientists can use this to determine the difference between the temperature of surface seawater and the temperature of bottom water at the same site, using a single sediment core – that includes both benthic and pelagic forams.

12. While water passes through the Hydrological Cycle, there is continuous oxygen isotope fractionation

13. n =1000 100 H 2 18 O 900 H 2 16 O RAIN n =100 20 H 2 18 O 80 H 2 16 O n = 900 80 H 2 18 O 820 H 2 16 O RAIN n =100 10 H 2 18 O 90 H 2 16 O n = 800 70 H 2 18 O 730 H 2 16 O R= 0.111 R= 0.0975R= 0. 095 R= 0.25 R= 0.11 EVAP. n =1000 >> > > > n = number of H 2 O molecules cloud = diminishing reservoir

14. Global Meteoric Water Line Product of  D and  18 O values for precipitation from all over the world. Slope of 8 approx. equal to value of Rayleigh condensation in rain. More heavy isotopes More light isotopes

15. How does this work? Oxygen isotopes are non-uniformly distributed over the surface of the earth. The process of evaporation, precipitation and transport of water vapor (H 2 O, containing oxygen of both isotopes) in the atmosphere results in a latitudinal variation in the  18 O of the water in different places.

16. Light water (water with 16 O) evaporates more easily than water with a lot of 18 O. This ‘light water’ evaporates near the equator and is transported toward the poles through many evaporation/ppt cycles. The 18 O/ 16 O ratio will be more negative in the snow that falls on a glacier than it is in the ocean from which the water evaporated. As the world's glaciers grow in volume, 18 O values of seawater become larger (more 16 O stored in ice). The oxygen isotope ratio of seawater (or ice core water) is now recording the size of the global ice sheets. FRACTIONATION OF OXYGEN ISOTOPES DUE TO EVAPORATION, PRECIPITATION AND TRANSPORTATION.

17. During precipitation as snow or rain, ‘heavy water’ (water with a higher 18 O ratio) tends to precipitate first, leaving the residual water vapor in the atmosphere enriched in light water (water with more 16 O). Each step of this evaporation/ ppt/ transport cycle decreases the  18 O value of the water vapor being transported from the equator to the poles by about 10 0 / 00. The result is that water with the light isotope of oxygen ( 16 O) is being transported preferentially to the poles from the equator – and there stored as ice. This leaves water with ‘excess’ 18 O (high values of  18 O) left as seawater.

18. NOTE: if the temperature dependence of evaporation/precipitation were the only process working, seawater at HIGH latitudes would have very high  18 O values (near +5 0 / 00 ). The fractionation between isotopes is higher at low temperatures! But the polar regions don’t. RAIN and runoff from ice/rivers produces seawater in the polar regions that has a  18 O near zero, similar to the tropics.

19. In the oceans, we use the oxygen isotope ratio determined from the calcium carbonate shells of forams. Temperature dependence (from text) for the proxy  18 O is T = 16.9 – 4.2 (  18 O c –  18 O w ) Where T is temperature in °C,  18 O c is the  18 O measured in calcite shells, and  18 O w is the  18 O value of seawater when shells formed. An alternate form of the expression (see text, page 153) is  18 O c =  18 O w – 0.23  T Where  means ‘change in’. This relationship allows paleoclimatologist to determine the temperature of the seawater at the time when the forams lived.

20. An example you have seen before. Remember: when  18 O goes negative, that means that the seawater temperature is getting WARMER.

21. The ratio of these isotopes is expressed in relation to a standard (PeeDeeBelemnite) as  13 C = [(R sample /R standard ) -1] x 1000 where R = ( 13 C/ 12 C). As  13 C values increase, the abundance of the heavier isotope ( 13 C) increases. Biological activity fractionates in favor of 12 C 12 C enriched High 12 C input This enrichment of 12 C within the biological reservoir, depletes the 12 C in the exterior seawater, and the  13 C ratio of the SEAWATER becomes HIGHER. There are two common stable isotopes of carbon: 12 C and 13 C.

23. High 12 C input to biology: Seawater becomes depleted in 12 C, Sea water  13 C ratio becomes HIGH and POSITIVE 12 C enriched X X High 12 C output to seawater as methane: Seawater becomes richer in 12 C and depleted in 13 C,  13 C ratio is LOW and NEGATIVE. Sea water Sediment

24. High bio-productivity Methane spike Low bio-productivity

25. Paleocene-Eocene Thermal Maximum (PETM) ~100,000 years in length What caused the PETM event?

26. The Paleocene-Eocene thermal maximum (PETM) (1) sea surface temperature rose by 5°C in the tropics; (2) by more than 6°C in the Arctic. (3) ocean acidification was strong (CCD was shallow). (4) with the extinction of 30 to 50% of deep-sea benthic formaminiferal species.

27. The initiation of the PETM is marked by an abrupt decrease in the  13 C proportion of marine and terrestrial sedimentary carbon, which is consistent with the rapid addition of >1500 gigatons of 13 C depleted carbon, most likely in the form of methane, into the hydrosphere and atmosphere. The PETM lasted only 210,000 to 220,000 years, with most of the decrease in  13 C occurring over a 20,000- year period at the beginning of the event.

28. Consider what 1200 gigatons of methane is. Over 20,000 years. We are (society is) presently emitting 34 Tg of CH 4 per year – now. Where 1 Tg = 10 9 kg = 10 6 metric tons. So present rate is 34 Tg/year = 34 X 10 6 metric tons/year. Over the past 50 years, that is a total of 1700 X 10 6 metric tons total Or about X 1000 less than the total PETM. BUT, the PETM was 20,000 years long, so the emission rate then was 1200 gigatons/20,000 years = 0.06 gigatons per year. Or about 60 X 10 6 metric tons per year. Which is only about 2X our present rate of CH 4 emission…

29. During this massive methane release, the oxidation and ocean absorption of this carbon would have lowered deep-sea pH. This low ocean pH would have led to rapid shoaling of the calcite compensation depth (CCD), followed by a gradual recovery. Evidence of a rapid acidification of the deep oceans would be evident in the abrupt transition from carbonate-rich sediment to clay, followed by a gradual recovery to carbonate. Samples of the ocean sediment from five South Atlantic deep- sea sites, all within the geologic time frame of the PETM. Acid Oceans?

30. During the Eocene, the CCD is inferred to have shoaled more than 2 km within a few thousand years.

31. Graphs of the core samples show an abrupt transition from carbonate-rich sediment to clay, followed by a gradual recovery (100K years) to carbonate.

32. Carbon reservoirs More than half the present day carbon is locked up in Gas hydrates.

34. Methane hydrate (methane clathrate) δ 13 C of methane hydrates is ~ -60‰. (why?) Methane oxidizes within 10 years to CO 2 in the ocean and atmosphere. Methane hydrate stores an enormous amount of reduced carbon (~7.5 – 15 x10 18 g). A release of 1.1 to 2.1 x 10 18 g of carbon from methane is sufficient to explain the -2 to -3‰ excursion in the δ 13 C of the ocean/atmosphere inorganic carbon reservoir. The stability of methane hydrate depends on temperature and pressure.

37. Methane Hydrate – what is it? CH 4 + H 2 O = (at low temperature and high pressure) methane hydrate Looks like ice, but is unstable at atmospheric pressure and room temperature. Also called ‘clathrate’. A molecule of methane (CH 4 ) is trapped inside the H 2 O ice structure. Sometimes, the organic molecule is heavier than methane (pentane, butane, etc), but not often.

38. NOTE: While this figure is required by law to be shown in all paleoclimate courses, do not try this at home. It is not dangerous, just extremely disappointing…

39. Phase diagram showing the water depths (and pressures) and temperatures for gas hydrate (grey area) stability. Many sediments lie within the range denoted by the box. The line shows the temperature in the Earth as a function of depth (geotherm). At greater depths in the sediments, the geotherm crosses from the hydrate zone (purple region) to the gas zone. This means that gas hydrate in sediments usually overlies free gas.

41. Drill core from Blake Ridge off SE U.S. Bubbles are methane gas. Natural gas (methane) pipe line, where water has leaked into the pressurized gas main.

43. Release of Methane hydrate Under stable ocean temperatures, methane hydrate can be released through slumping and permafrost warming due to rising arctic waters. Another way to release methane hydrate is through warming ocean temperatures.

44. Initial CO2 release from biomass destruction and volcanism. Warming temperatures gradually warm the oceans causing rapid (10 4 yrs) release of methane from coastal methane hydrate reservoirs. Methane oxidizes quickly (10 yrs) into atmospheric CO2, adding to increased warming of the oceans, releasing more methane into the atmosphere. Mechanism stops when all available methane hydrate at that specific ocean temperature is released or ocean circulation change thereby stabilizing the ocean temperatures. During the 10 4 yrs of methane release, oceans acidifies? Positive feedback of GHG release leading to the PETM

45. The trigger for the initiation of the PETM is a period of intense flood basalt magmatism (surface and sub-surface volcanism) associated with the opening of the North Atlantic, by generating metamorphic methane from sill intrusion into basin-filling carbon- rich sedimentary rocks

48. Other global disasters during the (troubled) Eocene. The Chesapeake Bay Boloid impact….

51. CONTINENTAL EFFECT

52. ALTITUDE EFFECT