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

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

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

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.

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

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.

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.

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, biology fractionates this oxygen isotope ratio (organisms accumulate more of the light isotope) 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.

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 planktonic 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). Who has more 18 O? Cold, benthic forams… and their  18 O ratio will be more positive 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.

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

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

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.

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 (and more +, with 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.

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.

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.

If we correct for the changes in seawater due to ice sheets, we can 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.

Sediment cores provide climate records that go back several million years, at lower resolution than the ice cores.

foramifera

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

Low-resolution marine stable- isotope records of the PETM and the carbon isotope excursion, together with the seafloor sediment CaCO3 record. The carbon isotope (a) and oxygen isotope (b) records are based on benthic foraminiferal records and the from drill holes in the South Atlantic. Panel b shows temperatures. The decrease in sedimentary CaCO3 reflects increased dissolution and indicates a severe decrease in seawater pH (that is, ocean acidification). From Zachos et al. Nature, 2008

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.

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 (  13 C = -60): Seawater becomes richer in 12 C and depleted in 13 C,  13 C ratio is LOW and NEGATIVE. Sea water Sediment

High bio- productivity Lots of 12 C stored as ‘biology’ leaving 13 C behind in the seawater Methane spike Low bio- productivity This is the PETM ‘spike’ And this is the broad Eocene thermal maximum

The Paleocene-Eocene thermal maximum (PETM) (1) sea surface temperature rose by 5°C in the tropics; (2) by more than 7°C in the Antarctic and Arctic. (3) ocean acidification was strong (CCD was shallow). (4) with the extinction of 30 to 50% of deep-sea benthic formaminiferal species. A good ‘Rule of Thumb’ is that temperature changes in the polar regions are about TWICE those of the Global Average Temperature Change. That is => 7°C temperature increase in the Antarctic means about a 3.5°C increase in global temperatures. Or about the temperature increase expected over the next 100 years due to anthropogenic greenhouse gas emissions.

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 >1200 gigatons of 13 C depleted carbon, most likely in the form of methane, into the hydrosphere and atmosphere. The broad Eocene thermal maximum lasted only 210,000 to 220,000 years, with most of the decrease in  13 C occurring over a 20,000- year period (the PETM) at the beginning of the event.

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

During this massive methane release, the oxidation and ocean absorption of this carbon would have lowered deep-sea pH (increased ocean acidity dramatically). 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?

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

Using Eocene data to simulate future climate (Zachos et al, Nature, 2008) (a), ocean surface pH (b), ocean surface calcite saturation (c) and deep-ocean temperature changes (d) in response to the input of 5,000 Gt C of anthropogenic CO2 into the atmosphere, starting from pre- industrial CO2 levels. Blue and green are w/wo a silicate- weathering feedback. Projected changes in deep-ocean temperature in d assume a homogeneous warming following temperature sensitivities to a doubling of CO 2 concentration: short-dashed line, 4.5 °C; solid line, 3.0 °C; long-dashed line, 1.5 °C.