In geochemistry, paleoclimatology and paleoceanographyδ18O or delta-O-18 is a measure of the ratio of stable isotopesoxygen-18 (18O) and oxygen-16 (16O). It is commonly used as a measure of the temperature of precipitation, as a measure of groundwater/mineral interactions, and as an indicator of processes that show isotopic fractionation, like methanogenesis. In paleosciences, 18O:16O data from corals, foraminifera and ice cores are used as a proxy for temperature.
The definition is, in "per mil" (‰, parts per thousand):
where the standard has a known isotopic composition, such as Vienna Standard Mean Ocean Water (VSMOW). The fractionation can arise from kinetic, equilibrium, or mass-independentfractionation.
Foraminifera shells are composed of calcium carbonate (CaCO3) and are found in many common geological environments. The ratio of 18O to 16O in the shell is used to indirectly determine the temperature of the surrounding water at the time the shell was formed. The ratio varies slightly depending on the temperature of the surrounding water, as well as other factors such as the water's salinity, and the volume of water locked up in ice sheets.
δ18O also reflects local evaporation and freshwater input, as rainwater is 16O-enriched—a result of the preferential evaporation of the lighter 16O from seawater. Consequently, the surface ocean contains greater proportions of 18O around the subtropics and tropics where there is more evaporation, and lesser proportions of 18O in the mid-latitudes where it rains more.
Similarly, when water vapor condenses, heavier water molecules holding 18O atoms tend to condense and precipitate first. The water vapor gradient heading from the tropics to the poles gradually becomes more and more depleted of 18O. Snow falling in Canada has much less H218O than rain in Florida; similarly, snow falling in the center of ice sheets has a lighter δ18O signature than that at its margins, since heavier 18O precipitates first.
Changes in climate that alter global patterns of evaporation and precipitation therefore change the background δ18O ratio.
Based on the simplifying assumption that the signal can be attributed to temperature change alone, with the effects of salinity and ice volume change ignored, Epstein et al. (1953) estimated that a δ18O increase of 0.22‰ is equivalent to a cooling of 1 °C (or 1.8 °F). More precisely, Epstein et al. (1953) give a quadratic extrapolation for the temperature, as
where T is the temperature in °C (based on a least-squares fit for a range of temperature values between 9 °C and 29 °C, with a standard deviation of ±0.6 °C, and δ is δ18O for a calcium carbonate sample).
Lisiecki and Raymo (2005) used measurements of δ18O in benthic foraminifera from 57 globally distributed deep sea sediment cores, taken as a proxy for the total global mass of glacial ice sheets, to reconstruct the climate for the past five million years.
The stacked record of the 57 cores was orbitally tuned to an orbitally driven ice model, the Milankovitch cycles of 41 ky (obliquity), 26 ky (precession) and 100 ky (eccentricity), which are all assumed to cause orbital forcing of global ice volume. Over the past million years, there have been a number of very strong glacial maxima and minima, spaced by roughly 100 ky. As the observed isotope variations are similar in shape to the temperature variations recorded for the past 420 ky at Vostok Station, the figure shown on the right aligns the values of δ18O (right scale) with the reported temperature variations from the Vostok ice core (left scale), following Petit et al. (1999).[clarification needed]
On the previous Isotope page, you learned a bit about what isotopes are, how they are obtained, and how the isotopes of certain elements are measured. Here, we’ll elaborate on how to read carbon and oxygen isotope data and how the values are often interpreted by paleoclimatologists.
Reading Isotope Data
Measurements of carbon and oxygen isotope values of a sample obtained using a mass spectrometer are compared to a sample of known isotopic values, called a reference standard. The resultant isotopic signature of a sample is expressed using a delta (δ) followed by the isotope number and the symbol of the element being measured. Oxygen isotope measurements are read as δ18O, or delta oxygen eighteen, and carbon is read as δ13C, or delta carbon thirteen. These values are expressed as per mil (‰). The definitions of δ13C and δ18O are as follows:
Thus, delta values of carbon and oxygen can be either positive or negative.
Interpreting Carbon and Oxygen Isotopes
Because both carbon and oxygen isotopes are measured simultaneously from one sample, the data are usually interpreted together. Oxygen and carbon isotope data from samples that are plotted, either against age or depth, are called ‘curves’ by scientists. There are several factors that influence carbon and oxygen curves, but below, we will focus on carbon and isotope curves obtained from planktic and benthic foraminifera, and a few of the ways these curves can be interpreted. Carbon isotopes are a bit more complex, so we have only included a few of the ways they can be interpreted. As a reminder, planktic foraminifera live near the surface of the ocean (in the mixed layer or the upper thermocline), and benthic foraminifera live at the ocean bottom, with some species living within the sediment on the seafloor.
Evaporation and precipitation are two factors that most influence the ratio of heavy (oxygen 18; O18) to light (O16) oxygen in the oceans. When seawater evaporates, O16 is preferentially uptaken because it is lighter, while the heavier O18 is left behind. When water vapor condenses, the heavier oxygen leaves first, as precipitation, before the lighter oxygen.
During different times in Earth’s history, the oceans had more O18 relative to O16. We’ll briefly discuss how these oxygen isotope values are interpreted by paleoclimatologists.
Increased Oxygen 18 Values
Water evaporated from the ocean surface is enriched in O16 (it has a lighter or more negative δ18O signal). When this water precipitates out as snow at the poles, it becomes trapped on land, compacting over time to create ice. When Earth is cool enough year-round that ice caps are permanent, the ice on Earth’s surface is enriched in O16 relative to O18. This happens as more and more of the lighter O16 is locked up in ice (the water has a more positive δ18O signal). During colder times in Earth’s history, the oxygen isotope value extracted from the shells (tests) of benthic and planktic foraminifera shells have heavier, or more positive, δ18O values. The foraminifera will preferentially create their tests out of the lighter O16 unless it is not available and during these cold intervals, O18 is more abundant. Warmer intervals are indicated by lighter, or more negative, δ18O values.
Decreased Oxygen 18 Values
When the Earth begins to warm, ice caps begin to melt back, and if the Earth warms further, there are hardly any permanent ice sheets at the poles. When this happens, the O16-enriched ice returns to the ocean as water. A huge influx of water enriched with O16 creates ocean waters that then become diluted with respect to O18. On the chart on the right, warmer times are indicated by more negative (or ‘lighter’) O18 values (more negative δ18O), indicating a reduction in the amount of ice on the Earth.
The above scenario applies to the global oxygen values of the oceans. Isotopic values of oxygen can also be measured from different regions and compared for the same time in Earth’s history. Regionally negative δ18O signals (more O16) can indicate that an area experienced increased rainfall.
Because the ocean bottom water is more homogeneous, or well-mixed, compared to the surface ocean water, benthic foraminifera record ‘global’ isotope values. The figure at the right show the δ18O values from benthic foraminifera for the last 10 million years. Because we are geoscientists, we read time from the oldest to the youngest, just like you read English from the left to right. Thus, to interpret what is happening through time, begin at the bottom of the curve and ‘read’ up, towards younger dates. Notice at the bottom of the curve, values are in the red color with more negative values. This indicates that at this time in Earth’s history, the Miocene, the bottom ocean temperature was much warmer, and there was much less ice at the poles. As you move up the curve, the values fluctuate quite a bit, but notice how the curve, in general, trends to the right in the blue area. This means the values are becoming more positive, which indicates that there are more O18 molecules in the water, which is interpreted to mean the Earth was cooling down during this time, and there was more ice at the poles.
Oxygen curves from planktic foraminifera record a more local signal than benthic foraminifera. This is because the surface ocean is very dynamic (it’s always changing) compared to the deep ocean. In addition, it is not as homogeneous (well-mixed) as the bottom of the water column. Curves from planktic foraminifera can be interpreted the same as benthic foraminifera in general, but the effects of evaporation and precipitation can also affect the isotope signal. Evaporation causes a decrease of O16 ions and thus an increase of O18 in the water column, which causes the curve to move towards more positive δ18O values. Precipitation increases the amount of O16 ions in the water column, which causes the curve to move towards more negative δ18O values.
Carbon isotopes obtained from the shells of marine organisms are strongly influenced by photosynthesis, respiration, and upwelling of ocean waters, the process by which older, more C12-rich waters are brought from the bottom ocean to the surface. Therefore, when a planktic foraminifera builds its shell in an upwelling area, the δ13C signal will be more negative. There are two main isotopes of carbon that we are concerned with in this website: carbon 13 (C13) and carbon 12 (C12). Just like oxygen isotopes, the carbon isotope with more neutrons (C13) is heavier, whereas C12 is lighter.
Throughout Earth’s history, the amount of C12 and C13 in the atmosphere and in the oceans has changed, which makes working with carbon isotopes a bit trickier than oxygen. Thus, we’ll just briefly discuss some of the ways carbon curves can be interpreted without bogging you down with too many technical terms and background information (although if you are interested in learning more about this, please contact us!).
Increased Carbon 13 Values
Increased, or heavier, δ13C values, generally indicate increased productivity in the ocean. Photosynthesizing organisms, such as algae and plankton, preferentially uptake C12 during photosynthesis, which leaves more C13 in the water column with which marine organisms build their shells. When there is a lot of photosynthesis happening in the water column geoscientists refer to this as a time of increased productivity. When there are times of increased productivity (lots of things growing and photosynthesizing in the water column), this usually means there is increased burial of carbon, or sediment accumulation on the seafloor (because there’s lots of things pooping and dying and falling down to the seafloor).
An increase in δ13C values can also indicate that erosion from land (the terrestrial realm) is decreased. Soil tends to have a more negative δ13C value because it contains remains of dead plants (that are generally made of C12 ions). When this soil washes into the ocean, it puts more C12 in the waters.
Decreased Carbon 13 Values
Alternatively, when there is very little photosynthesis occurring in the water column, there are more C12 molecules in solution, and thus more are incorporated into the shells of marine organisms. This can also correlate with times of decreased burial, because there are not as many organisms living in the water column (less poop and fewer dead organisms falling to the seafloor).
δ13C values can also become more negative due to increased erosion from land into the ocean (see above paragraph). CO2 from volcanoes tends to also be very enriched in C12, so times in Earth’s history that are characterized by intense volcanism can create more negative δ13C values.
Carbon curves created from benthic foraminifer shells tell geoscientists something about the bottom water conditions in the ocean. Respiration, or the decay of organic matter (remember the poop and dead organisms we talked about earlier?) tends to release more C12 into the bottom waters. As water masses move along the seafloor, they pick up this C12 signature, which gets incorporated into the benthic foraminifer’s shell. Thus, a benthic foraminifer’s δ13C value can tell us about the age of the bottom waters (older bottom waters that have picked up more C12= more negative δ13C values; younger bottom waters= more positive δ13C values). The δ13C curve at left is a global stack (lots of carbon curves from all over the world were plotted together) from benthic foraminifera. Again, we read the figure from the bottom (oldest) to the top (youngest). Notice that δ13C values, in general, from 10-0 million years become more negative. This signal is in part from the bottom ocean waters becoming older as the modern-day deep-water ocean circulation began (youngest waters sink in the Arctic and travel along the bottom of the water column into the Pacific Ocean).
Benthic foraminifera carbon values also tell us something about sea level through time, as related to weathering. When sea level is high, this generally leads to less weathering, or less transport of soils (which, remember, are enriched in C12), into the oceans, and thus is recorded as a more positive δ13C signal. Low sea level generally correlates to higher global erosion rates, and thus a more negative δ13C signal. On the chart at left, notice that during 10-6.5 million years ago, δ13C values are generally more positive. As you move into the modern towards 0 million years, the values are more negative. If you look back at the oxygen isotope chart, you’ll notice that these relatively negative carbon values correlate with times of cooler climate and more ice. When there is more ice, this leads to a drop in sea level. Thus, the more negative values on the carbon isotope curve are partly due to a sea level fall from increased ice sheets at the poles, which led to increased weathering.
Similar to oxygen isotopes, carbon isotopes obtained from planktic foraminifera are generally interpreted as a more local signal, as the surface layer in the ocean is not as homogeneous (well-mixed) as the bottom waters. Carbon isotope values in planktic foraminifera shells are mainly affected by photosynthesis and upwelling. When more organisms are photosynthesizing in the upper water column (increased productivity), this leads to more C13 ions available in the water for foraminifera to build their shells (because the C12 ions are being used for photosynthesis). When there is little to no photosynthesis happening in the water column, more C12 ions are available for the foraminifera to build into their shells, causing the δ13C signal to become more negative.
To learn more about oxygen and carbon isotopes, visit the following sites: