What is an isotope?
Same element with the same number of protons, but with a
different numbers of neutrons:
ELEMENT
ISOTOPE
HYDROGEN
(Z=l)
l
2
Stable isotope
abundances
Out of every 100 atoms of
Oxygen, 0.2 atoms would
be 18O and the rest would
be 16O.
CARBON
(Z=6)
NITROGEN
(Z=7)
OXYGEN
(Z=8)
H (Protium)
H (D, for Deuterium)
ABUNDANCE
99.985
0.015
12
98.9
1.1
14
N
15
N
99.63
0.37
16
99.76
0.04
0.2
C
13
C
O
17
O
18
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,
13C/12C, 15N/14N
, 18O/16O, 34S/32S
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.
Definitions - ambiguity
“18O-rich”
“18O-poor”
“heavy oxygen”
“light oxygen”
“enriched oxygen”
“depleted oxygen”
“16O-poor”
“16O-rich”
Reservoir containing
both 18O and 16O
atoms
Original seawater
16O
Remove 18O and 16O
atoms at a different
ratio than the initial
reservoir
18O
Modified seawater
16O
18O
Biology (forams)
16O
18O
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.
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 d18O ratio as… [note: heavy isotope over light isotope, always]
Where (18O/16O)SMOW is a sample of surface ocean where d18O = 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 d18O 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 d18O 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.
Examples.
Tropical planktonic foram shells that grow near 21°C have a d18O of
about -1 0/00. (lower than the seawater value).
But benthic forams living in the deep ocean (near 2°C) have a d18O
value of about +5 0/00 (higher than the ambient seawater).
Who has more 18O?
Cold, benthic forams…
and their d18O 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.
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 d18O ratio of seawater!
While water passes through the Hydrological
Cycle, there is continuous oxygen isotope
fractionation
Global Meteoric Water Line
More heavy isotopes
Product of dD and d18O
values for precipitation
from all over the world.
Slope of 8 approx. equal to
value of Rayleigh
condensation in rain.
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 d18O 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, d18O values of seawater
become larger (and more +, with
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 d18O
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 d18O) left as seawater.
NOTE: if the temperature dependence of evaporation/precipitation were the
only process working, seawater at HIGH latitudes would have very high d18O
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 d18O 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 d18O is
T = 16.9 – 4.2 (d18Oc – d18Ow)
Where T is temperature in °C,
d18Oc is the d18O measured in calcite shells, and
d18Ow is the d18O value of seawater when shells formed.
An alternate form of the expression (see text, page 153) is
Dd18Oc = Dd18Ow – 0.23 DT Where D 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 d18O goes negative, that means that the seawater
temperature is getting WARMER.
Low-resolution marine stableisotope 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
There are two common stable isotopes of carbon: 12C and 13C.
The ratio of these isotopes is expressed in relation to a standard
(PeeDeeBelemnite) as
d13C = [(Rsample/Rstandard) -1] x 1000
where R = (13C/12C).
As d13C values increase, the abundance of the heavier isotope (13C) increases.
Biological activity fractionates in favor of 12C
High 12C input
12C
enriched
This enrichment of 12C within the biological
reservoir, depletes the 12C in the exterior
seawater, and the d13C ratio of the SEAWATER
becomes HIGHER.
Sea water
High 12C input to biology:
12C
enriched
Seawater becomes
depleted in 12C,
Sea water d13C ratio
becomes HIGH and
POSITIVE
High 12C output to
seawater as methane
(d13C = -60):
Sediment
12C
enriched
X X
Seawater becomes
richer in 12C and
depleted in 13C,
d13C ratio is LOW and
NEGATIVE.
And this is the broad
Eocene thermal maximum
Low bioproductivity
This is the PETM ‘spike’
Methane
spike
High bioproductivity
Lots of 12C stored
as ‘biology’ leaving
13C behind in the
seawater
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 d13C proportion of marine and
terrestrial sedimentary carbon,
which is consistent with the rapid addition of >1200
gigatons of 13C 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 d13C occurring over a 20,000year 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.
Acid Oceans?
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 deepsea sites, all within the geologic time frame of the PETM.
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 preindustrial CO2 levels.
Blue and green are w/wo a silicateweathering feedback. Projected
changes in deep-ocean temperature
in d assume a homogeneous
warming following temperature
sensitivities to a doubling of CO2
concentration: short-dashed line, 4.5
°C; solid line, 3.0 °C; long-dashed
line, 1.5 °C.