• CARBON SINKS (from the surface of the earth)

• WEATHERING OF ROCKS (pulls CO2 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,

• CO2 EMITTED BY SUBDUCTION VOLCANOS.

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 18O and the rest would be 16O.

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, 13C/12C, 15N/14N , 18O/16O, 34S/32S

Definitions - ambiguity

“18O-rich” “18O-poor”

“heavy oxygen” “light oxygen”

“enriched oxygen” “depleted oxygen”

“16O-poor” “16O-rich”

Because these ratios are so small, chemists measure 18O/16O (=R), rather than 18O or 16O abundance

And then report them as ratios compared to a standard.

R = R0* 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 (R0), the remaining fraction of the reservoir (f) and the fractionation factor ()

Reservoir containing both 18O and 16O atoms

18O

16O

18O

16O

Remove 18O and 16O atoms at a different ratio than the initial reservoir

18O

16O

That changes the ratio of 18O and 16O 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:

16O (about 99.8% of total) and 18O (most of the rest).

There are other oxygen isotopes, but they are not used for paleotemperatures.

The ‘normal’ ratio of 18O/16O 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 18O ratio as… [note: heavy isotope over light isotope, always]

Where (18O/16O)SMOW is a sample of surface ocean where 18O = 0.

The ratio Oxygen isotopes 16O and 18O 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 16O and 18O in the carbonate records that ratio that is present in seawater,

modified by the temperature of the sea water (through fractionation of the 18O and 16O isotopes).

The 16O and 18O 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 18O AND 16O

Planktonic foraminifera live in the upper 100 meters of the ocean. In the PRESENT DAY ocean, surface seawater has a 18O 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 18O decrease for each 4.2°C increase in water temperature.

or… 18O becomes less abundant in the foram carbonate shells - with respect to 16O - 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 18O ratio of seawater!

Examples.

Tropical plantonic foram shells that grow near 21°C have a 18O of about -1 0/00. (lower than the seawater value).

But benthic forams living in the deep ocean (near 2°C) have a 18O 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.

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

n =1000100 H2

18O900 H2

16O

RAIN n =100

20 H218O

80 H216O

n = 90080 H2

18O820 H2

16O

RAIN n =100

10 H218O

90 H216O

n = 80070 H2

18O730 H2

16O

R= 0.111 R= 0.0975 R= 0. 095

R= 0.25 R= 0.11

EVAP.n =1000

> >

>

> >n = number of H2O molecules cloud = diminishing reservoir

Global Meteoric Water Line

Product of D and 18O 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 (H2O, containing oxygen of both isotopes) in the atmosphere results in a latitudinal variation in the 18O of the water in different places.

Light water (water with 16O) evaporates more easily than water with a lot of 18O.

This ‘light water’ evaporates near the equator and is transported toward the poles through many evaporation/ppt cycles.

The 18O/16O 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, 18O values of seawater become larger (more 16O 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 18O ratio) tends to precipitate first, leaving the residual water vapor in the atmosphere enriched in light water (water with more 16O).

Each step of this evaporation/ ppt/ transport cycle decreases the 18O 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 (16O) is being transported preferentially to the poles from the equator – and there stored as ice.

This leaves water with ‘excess’ 18O (high values of 18O) left as seawater.

NOTE: if the temperature dependence of evaporation/precipitation were the only process working, seawater at HIGH latitudes would have very high 18O 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 18O near zero, similar to the tropics.

In the oceans, we use the oxygen isotope ratio determined from the calcium carbonate shells of forams.

Temperature dependence (from text) for the proxy 18O is

T = 16.9 – 4.2 (18Oc – 18Ow)

Where T is temperature in °C,

18Oc is the 18O measured in calcite shells, and

18Ow is the 18O value of seawater when shells formed.

An alternate form of the expression (see text, page 153) is

18Oc = 18Ow – 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.

An example you have seen before.

Remember: when 18O goes negative, that means that the seawater temperature is getting WARMER.

The ratio of these isotopes is expressed in relation to a standard (PeeDeeBelemnite) as

13C = [(Rsample/Rstandard) -1] x 1000 where R = (13C/12C).

As 13C values increase, the abundance of the heavier isotope (13C) increases.

Biological activity fractionates in favor of 12C

12C enriched

High 12C input

This enrichment of 12C within the biological reservoir, depletes the 12C in the exterior seawater, and the 13C ratio of the SEAWATER becomes HIGHER.

There are two common stable isotopes of carbon: 12C and 13C.

High 12C input to biology:

Seawater becomes depleted in 12C,

Sea water 13C ratio becomes HIGH and POSITIVE

12C enriched

XX

12C enriched

High 12C output to seawater as methane:

Seawater becomes richer in 12C and depleted in 13C,

13C ratio is LOW and NEGATIVE.

Sea water

Sediment

High bio-productivity

Methane spike

Low bio-productivity

Paleocene-Eocene Thermal Maximum (PETM)~100,000 years in length

What caused the PETM event?

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.

The initiation of the PETM is marked by

an abrupt decrease in the 13C proportion of marine and terrestrial sedimentary carbon, which is consistent with the rapid addition of >1500 gigatons of 13C 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 13C occurring over a 20,000-year period at the beginning of the event.

Consider what 1200 gigatons of methane is. Over 20,000 years.

We are (society is) presently emitting 34 Tg of CH4 per year – now.

Where 1 Tg = 109 kg = 106 metric tons.

So present rate is 34 Tg/year = 34 X 106 metric tons/year.

Over the past 50 years, that is a total of 1700 X 106 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 106 metric tons per year.

Which is only about 2X our present rate of CH4 emission…

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?

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

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

Carbon reservoirs

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

Methane hydrate (methane clathrate)

•δ13C of methane hydrates is ~ -60‰. (why?)

•Methane oxidizes within 10 years to CO2 in the ocean and atmosphere.

•Methane hydrate stores an enormous amount of reduced carbon (~7.5 – 15 x1018 g).

•A release of 1.1 to 2.1 x 1018g of carbon from methane is sufficient to explain the -2 to -3‰ excursion in the δ13C of the ocean/atmosphere inorganic carbon reservoir.

•The stability of methane hydrate depends on temperature and pressure.

Methane Hydrate – what is it?CH4 + H2O = (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 (CH4) is trapped inside the H2O ice structure. Sometimes, the organic molecule is heavier than methane (pentane, butane, etc), but not often.

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…

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.

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.

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.

•Initial CO2 release from biomassdestruction and volcanism.

•Warming temperatures graduallywarm the oceans causing rapid (104yrs)release of methane from coastal methane hydrate reservoirs.

•Methane oxidizes quickly (10 yrs) intoatmospheric CO2, adding to increased warming of the oceans, releasing more methane into the atmosphere.

•Mechanism stops when all availablemethane hydrate at that specific ocean temperature is released or ocean circulation change thereby stabilizingthe ocean temperatures.

•During the 104yrs of methane release, oceans acidifies?

Positive feedback of GHG release leading to the PETM

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

Other global disasters during the (troubled) Eocene.

The Chesapeake Bay Boloid impact….

CONTINENTAL EFFECT

ALTITUDE EFFECT