Environmental Geochemistry
Name:______________________________
Lab 5 Isotopes
This lab includes a excerpt taken from Wally Broecker’s “Glacial World According to Wally”.
Wallace S. Broecker is currently the Newberry Professor in the Department of Earth and
Environmental Sciences at Columbia University. He developed the idea of a global "conveyor
belt" linking the circulation of the global ocean and made major contributions to the science of
the carbon cycle and the use of chemical tracers and isotope dating in oceanography. If you were
to just google Wallace S B his name appears first. The series of articles and books that he has
written would teach you everything you need to know about paleoclimate. For those of you that
have an acute interest in this topic, I recommend you go to the internet and download a PDF of
this book.
The Glacial World According to Wally
Indicators
Ice Volume and Oxygen Isotopes
“For every 500 water molecules in the sea, one bears the heavy isotope 18O. The 16O to 18O ratio
in sea water typifies our solar system and bears witness to the fact that nuclei (such as 16O)
consisting of an integral number of helium nuclei are manufactured with great preference in the
centers of stars. We know from measuring the ratio of light oxygen atoms (16O) to heavy oxygen
atoms (18O) in a wide variety of Earth substances that while the ratio is always close to 500, it
varies over a range of several percent. For instance, the calcium carbonate manufactured by
marine organisms has on the average four percent more heavy oxygen than sea water and the
snow falling at high elevations on the Antarctic ice cap has on the average four percent less
heavy oxygen than sea water. These differences reflect the fact that while chemically similar, the
isotopes of any given element are not chemically identical. For example, water made with the
heavy oxygen isotope has a one percent lower vapor pressure than water made with the light
oxygen isotope.
The separation of oxygen isotopes plays an important role in paleoclimatology providing
information both about past ice volumes and about past temperatures. As this section concerns
ice volume, we will consider this application first. On the time scale of one million years or less,
only one means exists to change the average isotopic composition of the water in the sea, i.e., the
growth and retreat of ice caps. The tie between the isotopic composition of ocean water and ice
volume stems from the fact that the snow which accumulates to make ice caps is depleted by
several percent in heavy oxygen. We know this from measurements on recent snow from
Antarctica and Greenland. We also understand why this is the case. This depletion is created
because the water which condenses to form precipitation in a cooling air mass is enriched in the
heavy isotope. The light isotope is correspondingly enriched in the residual vapor. Because of
this, the residual water vapor in the airmass becomes progressively depleted in the heavy isotope.
By the time air masses reach the very cold regions atop the Greenland and Antarctic ice caps,
this depletion in the heavy isotope has reached several percent. Because of this, the ice which
forms these caps (and also the ice which formed the great caps of glacial time) is depleted in
heavy water. These missing heavy molecules are stored in the sea, giving it a slightly higher 18O
to 16O ratio. Thus the greater the volume of ice on the continents, the higher the 18O to 16O ratio
in sea water. For example, if the snow which accumulated to form the ice sheets of last glaciation
were on the average 3.5 percent depleted in 18O (i.e., δ18O = -35‰), then the magnitude of the
enrichment in the sea is easily estimated. The ocean has a mean depth of 3800 meters, so the
removal of a 120 meter thick layer of water with 35‰ lower 18O content would leave the water
remaining behind enriched by about
(120/3800) x 35‰ or 1.1‰
We use per mil (i.e. parts per 1000) instead of percent (i.e. parts per 100) in this calculation to
conform to the standard notation adopted by those who make oxygen isotope measurements on
Earth materials (see figure 7 caption). Average ocean water constitutes the zero on this scale.
As the oxygen atoms in the CaCO3 deposited by marine organisms come from the oxygen in sea
water, these shells must record changes in the 18O to 16O ratio in the sea. Thus, down core series
of oxygen isotope ratios on the microscopic shells of foraminifera found in deep sea sediments
should provide a record of changes in ice volume on the continents! Indeed, changes in ice
volume constitute the largest contributor to the record kept by foraminifera shells.
Unfortunately, it is not the only contributor. Two other influences must be considered. The first
of these has to do with the fact that the ocean's 18O is not uniformly distributed. The same
meteorologic forces which cause the snow which falls on polar caps to be depleted in 18O
redistribute 18O across the sea surface. Those regions of the surface which lose more water to
evaporation than they gain from precipitation (and from the runoff from adjacent rivers) are
enriched in 18O as well as in salt. The opposite is true for areas of low salinity where
precipitation and runoff exceed evaporation. The range of these changes in today's open ocean
surface waters is as large as the signal imprinted by ice volume changes. Thus possible changes
in the distribution of 18O within the sea with changing climate must be considered.
Fortunately, a means exists to avoid this problem. Rather than relying on the record kept by the
shells of planktonic foraminifera which inhabit surface water, attention is focused on the record
kept by the shells of benthic foraminifera which live on the sea floor. As most of the ocean's
water resides in the deep sea, redistribution of 18O will create much smaller changes in this
reservoir than at the ocean's surface. Also the 18O content of deep sea waters is far more
homogeneous than that of surface water. In particular, the deep Pacific (the largest of the ocean's
reservoirs) is uniform isotopically both with location and depth. Thus, measurements made on
benthic foraminifera shells picked from cores from the floor of the deep Pacific should not be
significantly affected by changes in the distribution of 18O within the sea.
The other source of change in the 18O content of CaCO3 precipitated in sea is temperature. We
have already mentioned that the 18O to 16O ratio in shells is on the average 40‰ higher than that
in the water in which the shell grows. The complication is that the magnitude of the isotope
separation between water oxygen and shell oxygen depends on temperature. For each degree
Celsius the temperature cools, heavy oxygen is enriched in the shell by an additional 0.25‰.
Hence a three degree centigrade cooling of the water in which the shells grow leads to nearly as
large an isotope shift (i.e., 3 x 0.25 or 0.75‰) as that thought to be produced by the growth of
the ice caps (i.e., 1.1‰). Since the planet was cooler during glacial time, ice volume and
temperature worked together to raise the 18O to 16O ratio in marine shells. Again the deep
Pacific proves to be the most favored place to find shells least subject to temperature
contribution to oxygen isotope shifts. Not only is temperature quite uniform with depth and
geographic location, but more important, it is the coldest of the deep ocean reservoirs. This
proves to be important because the already very cold polar regions from which the deep seas are
ventilated could not have been much colder during glacial time. Nevertheless, the results of
measurements on benthic foraminifera from the deep sea (see figure 10) show that shells formed
during full glacial time had 1.75‰ higher 18O to 16O ratio than shells formed during peak
interglacial time. This is 0.65‰ more than would be expected from the ice volume change alone
(1.10‰), suggesting that deep Pacific water was 2.5°C cooler during glacial time than now.
Unfortunately, the extent of cooling of the deep Pacific in all likelihood was not simply
proportional to the extent of ice growth. In other words, the proportion of the change resulting
from ice volume on the one hand and water temperature on the other was not always the same as
it was at the time of the glacial maximum. Hence, in order to untangle the temperature
component from the ice volume component in oxygen isotope records for benthic foraminifera
from the deep Pacific Ocean, additional information is needed.
Harvard’s Dan Schrag proposed a means to more firmly establish the 18O to 16O ratio change
attributable to ice volume. He showed that it is recorded in the pore waters of marine sediments.
At several sites on the sea floor, he measured oxygen isotope ratios on pore water squeezed from
sediments from a series of depths recovered by the Ocean Drilling Program. An example is
shown in figure 11. The maximum in δ18O at about 40 meters depth reflects the high 18O of
glacial time. The reason that it is only about one third the expected amplitude is that diffusion of
water within the pores has gradually replaced the glacial age water molecules with post glacial
ones. From a knowledge of the molecular diffusion rate and the time history of the ice volume, it
is possible to deconvolve the actual isotopic composition of the glacial age bottom water. Based
on these measurements, Schrag concluded that the ice volume effect is ice volume effect is
1.0 ± 0.1‰.
Hence he would conclude that the average isotopic composition of the glacial age ice was
–32 ± 3‰.”
Broecker, Wallace S. (1995), The glacial world according to Wally, Eldigio Press
Now, how can you use this information to solve a problem using 18O? Keep in mind that for
every 1oC change in temperature, the 18O/16O changes by 0.23o/oo Therefore you should be able
to account for ice volume and temperature changes and determine a number of things from the
isotopic signatures.
Problem:
You visit a distant planet that is in a glacial maximum. The sea level there drops 200 meters
during glacial periods. The depth of the ocean is 5,000 meters and the 18O of the ice is -40o/oo.
1) What is the change in 18O in the ocean due to the increase in ice volume on land?
2) You find the 18O difference between glacial and interglacial forams to be 4.6o/oo. Calculate
the temperature change associated with this difference. (Assume you have already corrected for
18O measurements for the change in ice volume.)