Oxygen-Isotope Dynamics to Paleoclimate
|Oxygen Isotopes (Specifically)
||The Practice of
Without a doubt, an understanding of ancient
climate is one of the most critical studies underway in Quaternary
geology. What were temperatures like during the last ice
age? How was vegetation distributed across North America during
the late Wisconsin glaciation? What caused camelids to migrate to
North America during the Pleistocene Epoch? These and countless
other questions are the types of inquiries that drive Quaternary
geology. Furthermore, these are the types of questions that
cannot be answered without an understanding of paleoclimate.
Unfortunately, no method exists to measure ancient climate directly
(Aber, 2006). As such, science has developed many proxy methods
for conducting paleoclimatic studies. This webpage is dedicated
to one of these methods, namely, oxygen-isotope ratios. In the
paragraphs that follow, the subject of oxygen-isotope ratios will be
systematically dissected and explained. Subtopics include:
isotopes (generally), oxygen-isotopes (specifically), oxygen-isotope
dynamics, application of oxygen-isotope dynamics to paleoclimate, and
the practice of oxygen-isotope studies.
Naturally, the inaugural question in any study
such as this is the one regarding definition. That is, "What is
an isotope?" One of the better extant definitions is that
proffered by the INEEL
Environment Surveillance, Education, and
Research Program of the Idaho National Laboratory. The INEEL
defines isotopes as, "Two or more forms of an element having the same
number of protons in the nucleus (or the same atomic number), but
having different numbers of neutrons in the nucleus (or different
atomic weights). Isotopes of single elements possess almost identical
chemical properties. An example of isotopes are plutonium-238,
plutonium-239, plutonium-240, and plutonium-241, each acts chemically
like plutonium but have 144, 145, 146, and 147 neutrons, respectively"
(Case et al., 2003). In other words, an isotope is an element
that is like other elements with the same atomic number except that it
has a different number of neutrons. Furthermore, since the number
of neutrons is directly proportional to the atomic mass of an element,
isotopes with a greater number of neutrons are weightier than those
with fewer neutrons. Using the INEEL illustration above, then,
one can conclude that plutonium-240 is more massive than
plutonium-238; that is, plutonium-240 weighs more than
plutonium-238. This principle will become important during
the discussion on oxygen-isotope dynamics.
Isotopes can be either stable or unstable.
Stable isotopes are those isotopes that do not experience radioactive
decay, whereas unstable isotopes are those that do experience
radioactive decay. Examples of unstable isotopes abound and
include such infamous elements as Uranium-238, Plutonium-239,
Carbon-14, and Phosphorus-32. Over time these elements emit
alpha, beta, or gamma particles to become wholly different elements
elements with a lower energy state; that is, they radioactively
decay. An alpha particle is nothing more than a helium (He)
nucleus; therefore, alpha decay is when a radioactive isotope emits 2
protons and 2 neutrons (i.e. one He atom) from its nucleus
2000). For example, uranium decays to
thorium via alpha decay:
|23892U ----> 23490Th + 42H|
Beta particles, on the other hand, are electrons that are emitted from the nucleus. These nucleus-derived electrons are produced when a neutron decays to a proton in a process called weak interaction (Beta Reactivity, 2006). One can think of this process as the splitting of a neutron into one proton and one electron. The proton remains in the nucleus, thus converting it to a new element. The electron, however, is emitted. The following is an example of beta decay:
|23490Th ----> 23491Pa + 0-1e|
Finally, gamma decay is the process whereby an
element's nucleus goes from a higher energy state to a lower one via
the emission of electromagnetic radiation (i.e. photons of
energy). Since the number of
protons and neutrons does not change during gamma decay, the parent and
daughter elements are the same (Gamma
Decay, 2000). The
difference is that after gamma decay, the daughter element is less
energetic. Figure 1 is an illustration of gamma decay.
Figure 1: Gamma Decay of the element Dysprosium (Dy). Taken from Gamma Decay (2000).
Radioactive isotopes can be extremely useful in
dating of natural materials. The only information needed is the
half life of the element and its original concentration in the
substance tested. Carbon dating is a common radioactive isotope
test performed on organic substances. In Quaternary studies,
however, scientists have found stable isotopes (i.e. those that do not
decay) much more useful. Two of the more common stable isotopes
used in Quaternary geology are those of hydrogen and oxygen. It
is the latter that will be discussed herein.
Roughly fifteen isotopes of oxygen are known to
exist. These are Oxygen-12 through Oxygen-26. Of these
fifteen, however, only Oxygen-16, Oxygen-17, and Oxygen-18 are stable
isotopes. Furthermore, Oxygen-16 is the most abundant,
constituting 99.762% of the Earth's oxygen (Firestone, 2000; Gibson,
2005). Known as "light" oxygen, Oxygen-16 is composed of 8
protons and 8 neutrons (Herring, 2006). Applying the definition
of "isotopes" above, it is reasonable to conclude that Oxygen-17 and
Oxygen-18 have 8 protons/9 neutrons and 8 protons/10 neutrons,
respectively (see Figure 2 below). Of these latter two
oxygen-isotopes, though, only Oxygen-18 is found in sufficient
quanities to be of use to science. Known as "heavy"
oxygen--because its mass is approximately 12.5% more than the mass of
Oxygen-16--and having an abundance of around 0.200%, Oxygen-18 can be
used along with Oxygen-16 to determine ancient climate. This is
because the ratio of Oxygen-18 to Oxygen-16 in water changes with the
climate (Herring, 2006). Incidentally, when speaking of oxygen
isotopes throughout the remainder of this webpage, reference will be
to the Oxygen-16 and Oxygen-18 varieties.
Figure 2: Oxygen-16 and Oxygen-18 isotopes. White spheres are neutrons and blue spheres are protons. Taken from Herring (2006).
Water molecules are composed of two hydrogen
atoms and one oxygen atom. Because oxygen can be found
principally in two stable varieties, it stands to reason that some
water molecules will contain Oxygen-16 atoms and others will contain
Oxygen-18 atoms. As noted earlier, the isotope with a greater
number of neutrons will be more massive than the one with a lesser
number. Therefore, it is reasonable to conclude that water
molecules bearing Oxygen-18 will be slightly heavier than those
containing Oxygen-16. This is an important point in
oxygen-isotope dynamics, and one that will be discussed in the
The ratio of Oxygen-18 to Oxygen-16 in the Earth's
atmosphere is dependent upon
the dynamics of evaporation and
condensation that occur at the atmospheric boundary layer and within
the atmosphere itself. This is because water molecules containing
Oxygen-16 tend to evaporate more readily than those containing
Oxygen-18, while molecules containing Oxygen-18 tend to condense more
easily (Herring, 2006). As such, evaporation of water from the
ocean at lower latitudes tends to increase the atmospheric
concentration of Oxygen-16 in those regions due to higher rates of H216O
vaporization. However, as Hadley cells and other global/regional
atmospheric convections begin transporting the newly formed water vapor
poleward, cooler temperatures and upper level divergences eventually
cause condensation and precipitation to occur. During the process
of condensation, the first water molecules to precipitate are those
containing Oxygen-18. Therefore, initial precipitation during a
rain or snow event is typically enriched in H218O
and gradually becomes more and more "light." This process is
known as the "rainout effect," and it results in atmospheric Oxygen-18
decline (Gibson, 2005). In fact, it is the cumulative effects of
rainout that result in a lateral differentiation of 18O/16O
concentration in the atmosphere. That is, as one travels from the
equator to the poles, the atmosphere generally becomes more and more
depleted of Oxygen-18, until precipitation at the poles contains up to
5% less Oxygen-18 than does ocean water (Herring, 2006).
This decrease in Oxygen-18 levels with increasing latitude is known as
the "latitude effect" (Kendall et al., 2004). As a result, polar
ice sheets tend to evidence a significant reduction in Oxygen-18.
Figure 3 illustrates this point graphically.
Figure 3: Illustration of Oxygen-18 depletion with latitude. Taken from Herring (2006).
a contemporary perspective, the world
ocean is generally in a
state of isotopoic equilibrium. That is to say, though there is
certainly lateral variation in oxygen-isotope ratios (implied in Figure
3 above), the average18O/16O
concentration does not change overall. This is because
evaporation is balanced by condensation/precipitation worldwide.
In other words, enrichment of Oxygen-18 in equatorial regions due to
evaporation is balanced by enrichment of Oxygen-16 due to both affluent
glacial meltwater (see Figure 3) and cumulative global precipitation.
As will be
shown later, however, this is not necessarily the case over long
periods of geologic time.
oxygen-isotope concentration (d18O) of a water
sample is determined with respect to a datum known as standard mean
ocean water (SMOW). SMOW is a sample of well-mixed ocean water
that contains a precisely known 18O/16O
and is currently stored at the International Atomic Energy Agency in
Vienna, Austria (Groning
& Frohlich, 2006).
Oxygen-isotope concentrations are typically expressed in parts per
using the following mathematical equation:
|d18O = [(18O/16O) sample - SMOW / SMOW] X 1000|
d18O = 00/00 for SMOW
(Aber, 2006; Kennett, 1982). Presently, the world ocean has a d18O
= 0.00/00, and glacial
ice sheets exhibit d18O
values ranging from -200/00 to -300/00
2006--see Figure 4). The next section, however, will detail
conditions in which these numbers experience marked
Figure 4: Modern-day oxygen-isotope concentrations for the hydrosphere, atmosphere, and cryosphere. Values for "d" are given in parts per thousand (0/00). Information derived from Aber (2006) and Herring (2006).
|1. As the 18O/16O
ratio for landed ice decreases, the18O/16O
ratio for ocean water increases.
2. As the 18O/16O ratio for landed ice increases, the 18O/16O ratio for ocean water decreases.
3. The18O/16O ratio for both landed ice and ocean water is atmospheric temperature dependent.
discussion of these principles is in order. This is best done by way of
other conditions remaining constant, suppose that the Earth's
atmosphere began to cool. In this scenario, landed ice
volume would obviously increase due to the equatorward extension of
freezing conditions. Considering latitude effect, this
equatorward movement of cooler temperatures would cause rain out to
begin at lower latitudes, thus decreasing the 18O/16O
ratio of the polar ice sheets (Herring, 2006; Kennett, 1982). In
short, cooler temperatures worldwide would result in more water being
trapped on land, and the water thus trapped would have a lower 18O/16O
ratio due to the combined effects of temperature distribution, rain
out, and latitude. Furthermore,
increased sequestration of water on land would cause the 18O/16O
ratio for ocean water to increase. That is, since more
Oxygen-16-rich water would be locked in polar ice sheets, there would
be less of it in the world ocean and, therefore, the ratio of 18O/16O
would necessarily increase (Hambrey & Alean, 2004; Aber,
2006). On the level of fundamental scientific assumption, what is
being described herein is nothing more than the conservation of
matter. Thus, for a cooling scenario, the interplay of the above
three principles is quite clear.
were the Earth's atmosphere to warm, landed ice volume would decrease
due to the poleward retreat of freezing conditions. This poleward
movement of cooler temperatures would result in rain out at higher
latitudes and a corresponding increase in 18O/16O
ratios for the polar ice sheets (Herring,
2006; Kennett, 1982). Finally, decreased volumes of landed ice
would return Oxygen-16-rich water to the world ocean, thus decreasing
the overall 18O/16O
ratio of seawater (Hambrey & Alean, 2004). Again, the
relationship of the above principles is easily discernable.
these illustrations, then, two sub-principles of principle #3 (above)
temperature decreases, the 18O/16O
ratio for landed ice decreases and the 18O/16O
ratio for ocean water increases.
3b. As atmospheric temperature increases, the 18O/16O ratio for landed ice increases and the 18O/16O ratio for ocean water decreases.
it is these latter two sub-principles that are directly applied to
paleoclimates. That is, when it is found in the terrestrial ice
record that 18O/16O
ratios are lower than presently observed or in the marine record that 18O/16O
ratios are higher than presently observed, then it is
assumed that paleotemperatures--and thus paleoclimates--were
cooler. On the other hand, if 18O/16O
ratios in the terrestrial record are found to be higher and 18O/16O
ratios in the marine record are determined to be lower than at present,
then warmer paleoclimates are posited.
There are essentially two geologic records from which scientists
have gathered and continue to gather oxygen-isotope data. These
are ice cores from glacial ice sheets (Figure 5) and sediments from the
marine environment. The former include cores gathered during the
Greenland Ice Core Project (GRIP)
and Greenland Ice Sheet Project 2
(GISP2), as well as the
famous Vostok core from Antarctica. The
Vostok core is, of course, the most famous ice core due to its
impressive length--some 3.7 km covering over 420,000 years of Earth's
history (Hambrey & Alean, 2004)! Analysis of glacial ice
easily yields oxygen-isotope ratios and has been of inestimable value
in the study of Quaternary geology.
Figure 5: Montage of photos taken during experimental ice core drilling in Greenland, 2005. Taken from Riebeek (2006b).
However, it is the study of marine sediments that has truly
revolutionized Quaternary geology. This is principally due to the
longer spans of geologic time available in the marine record.
Through the efforts of scientists such as C. Emiliani and N.
Shackleton, the analysis of oxygen-isotope ratios in marine sediments
has allowed geologists and climatologists to estimate paleotemperatures
for nearly the entire Quaternary period (Gibbard, 2006; Kennett,
1982). In fact, it was Shackleton who helped devise the equation
 that is still used today for estimating paleotemperatures from
fossil foraminifera in marine sedimentary deposits:
|T(oC) = 16.9 – 4.38(d18Oc – d18Ow) + 0.10(d18Oc – d18Ow)2|
oxygen-isotope ratio for calcite in foraminifera and d18Ow
oxygen-isotope ratio for ancient seawater (Kennett, 1982; Costa et al.,
2006). Foraminifera are used for oxygen-isotope analysis because
the shells of these organisms are made from calcium carbonate (CaCO3)--an
oxygen-isotope bearing compound. These fossils act like veritable
event recorders in nature. As planktonic and benthic foraminifera
build their shells, they incorporate oxygen in ratios proportional to the 18O/16O
ratio of the contemporary seawater. As such, when they die and
in marine sediment, they record indirectly
in their carcasses the 18O/16O
ratio of the
seawater at the time of deposition. The modifiers "proportional"
and "indirectly" are employed here because even though the18O/16O
ratio of ancient seawater can be determined from
foraminifera shells, it is not done so directly. Herring offers
the best explanation for this when he writes,
|"As the shells form, they tend to incorporate more heavy oxygen than light oxygen, regardless of the oxygen ratio in the water. The biological and chemical processes that cause the shells to incorporate greater proportions of heavy oxygen become even more pronounced as the temperature drops, so that shells formed in cold waters have an even larger proportion of heavy oxygen than shells formed in warmer waters, where the difference is less notable. This temperature-based skew effect means that the oxygen isotope make-up of shells would not precisely match the make-up of the ocean water in which they grew. Scientists must correct for this skew if they are to learn about the ratio of oxygen isotopes in the ocean waters where the shells formed" (Herring, 2006).|
As stated in the introduction, no method exists to measure
paleoclimates directly. However, through the use of
oxygen-isotope ratios, science has devised a means by which to estimate
ancient climate by proxy. Though by no means infallible, the
method of paleoclimate determination using oxygen-isotope ratios has
nonetheless revolutionized Quaternary geology. Today, scientists
are quite confident in their estimations of paleoclimate (see Figure
6). Countless research projects and scholarly works have been
completed that have as their foundational methodology the use of
oxygen-isotope ratios. What is more, the list continues to
grow. Where the use of oxygen-isotopes will take science, no one
can accurately predict; however, one thing is for certain:
oxygen-isotope studies will be part and parcel of Quaternary geology
for the forseeable future.
Figure 6: Example of a paleoclimate graph derived from oxygen-isotope ratios in marine sediments.
Adapted from Riebeek (2006a).
Aber, James S. (2006).
ES 331/767 Lecture 11: Paleoclimate Reconstruction.
Retrieved October 7, 2006, from the Emporia State University Web
(2006). Retrieved November 11, 2006, from the Georgia State
University Web site:
Case, M., Halford, D., & Martin, C. (2003). Appendix E: Glossary.
Retrieved November 11, 2006, from the INEEL Environmental Surveillance,
Education, and Research Program Web site:
Costa, K. B., Toledo, F. A. L., Pivel, M. A. G., Moura, C. A. V.,
Chemale, F. (2006). Evaluation of Two Genera of Benthic
Foraminifera for Down-Core Paleotemperature Studies in the Western
South Atlantic. Brazilian
Journal of Oceanography, 54(1), 75-84.
Firestone, Richard B. Isotopes
of Oxygen. Retrieved November 11, 2006, from the Lawrence
Berkeley National Laboratory Web site:
(2000). Retrieved November 11, 2006, from the Lawrence Berkeley
National Laboratory Web site:
Garrison, Tom. (2006). Essentials
of Oceanography (4th ed.). Belmont, CA: Thomson
Gibbard, Philip. (2006). Professor Sir Nicholas
Shackleton. Boreas, 35,
Gibson, John. (2005). Oxygen.
Retrieved November 11, 2006, from the Sustainability of Semi-Arid
and Riparian Areas Web site:
GISP2--Greenland Ice Sheet Project
2. (2006). Retrieved November 11, 2006, from the
University of New Hampshire Web site: http://www.gisp2.sr.unh.edu.
Groning, M., & Frohlich, K. (2006). Part II: Example of Reference
Materials Certified for Stable Isotope Composition.
Retrieved November 11, 2006, from the International Atomic Energy
Agency Web site:
Hambrey, M., & Alean, J. (2004). Glaciers (2nd ed.).
Cambridge, England: University Press.
Herring, David. (2006). Paleoclimatology: the Oxygen Balance.
Retrieved November 11, 2006, from the National Aeronautics and Space
Administration Web site:
Kendall, C., Caldwell, E., & Snyder, D. (2004). Resources on Isotopes, Periodic
Table: Oxygen. Retrieved November 11, 2006, from the
United States Geologic Survey Web site:
Kennett, James. (1982). Marine Geology. Englewood
Cliffs, NJ: Prentice-Hall.
(2000). Retrieved November 11, 2006, from the Boston University
Riebeek, Holli. (2006a). Paleoclimatology:
Explaining the Evidence. Retrieved November 27, 2006, from
the National Aeronautics and Space Administration Web site:
________. (2006b). Paleoclimatology:
Core Record. Retrieved November 11, 2006, from the
Aeronautics and Space Administration Web site:
Stauffer, Bernhard, et al. (2006). Greenland Ice Core Project: An ESF
Research Programme--Final Report. Retrieved November 11,
2006, from the National Oceanic and Atmospheric Administration Web