Paleoclimatology

Paleoclimatology (also Palaeoclimatology) is the study of climate change taken on the scale of the entire history of Earth. It uses records from ice sheets, tree rings, sediment, and rocks to determine the past state of the climate system on Earth.

Reconstructing ancient climates

Paleoclimatologists employ a wide variety of techniques to deduce ancient climates.

Ice

Mountain Glaciers and the polar ice caps/ice sheets are a widely employed source of data in paleoclimatology. Recent ice coring projects in the ice caps of Greenland and Antarctica have yielded data going back several hundred thousand years -- over 800,000 years in the case of the EPICA project.

• Inside of these layers scientists have found pollen, allowing them to estimate the total amount of plant growth of that year by the pollen count. The thickness of the layer can help to determine the amount of precipitation that year. Certain layers contain ash from volcanic eruptions.
• Air trapped within fallen snow becomes encased in tiny bubbles as the snow is compressed into ice in the glacier under the weight of later years' snow. This trapped air has proven a tremendously valuable source for direct measurement of the composition of air from the time the ice was formed.
• Because evaporation rates of water molecules with slightly heavier isotopes of hydrogen and oxygen are slightly different during warmer and colder periods, changes in the average temperature of the ocean surface are reflected in slightly different ratios between those isotopes. Various cycles in those isotope ratios have been detected.

Dendroclimatology

This science retrieves climate information from tree rings. Rings from living trees of great age give data about recent centuries back to a few millennia. Older intact wood that has escaped decay can extend the time covered by identifying patterns that match rings of known age from live trees. Petrified tree rings give paleoclimatology data over a much larger stretch of time. The fossil itself is dated with radioactive dating within a wide margin of error. The rings themselves can give some information about rainfall and temperature during that epoch.

On a longer time scale, geologists must refer to the sedimentary record for data.

Sedimentary content

• Sediments, sometimes lithified to form rock, may contain remnants of preserved vegetation, animals, plankton or pollen, which may be characteristic of certain climatic zones.
• Biomarker molecules such as the alkenones may yield information about their temperature of formation.
• Chemical signatures, particularly Mg/Ca ratio of calcite in foramanifera tests, can be used to reconstruct past temperature.
• Isotopic ratios can provide further information. Specifically, the δ¹⁸O record responds to changes in temperature and ice volume, and the δ¹³C record reflects a range of factors, which are often difficult to disentangle.

Sedimentary facies

On a longer time scale, the rock record may show signs of sea level rise and fall; further, features such as "fossilised" sand dunes can be identified. Scientists can get a grasp of long term climate by studying sedimentary rock going back billions of years. The division of earth history into separate periods is largely based on visible changes in sedimentary rock layers that demarcate major changes in conditions. Often these include major shifts in climate.

Corals

Coral "rings" are similar to tree rings, except they respond to different things, such as the water temperature and wave action. From this source, certain equipment can be used to derive the sea surface temperature and water salinity from the past few centuries. The δ¹⁸O of coraline red algae provides a useful proxy of sea surface temperature at high latitudes, where many traditional techniques are limited.^[1]

Limitations

All records decrease in utility back in time. The oldest ice core taken was from the Antarctic is dates to 800, 000 years old. An international effort is currently being made in the same location to core to 1.2 million years ago. The deep marine record, our source of most isotopic data, only exists on oceanic plates, which are eventually subducted - the oldest remaining material is 200 million years old. Older sediments are also more prone to corruption by diagenesis. Consequently, our resolution and confidence in data decrease over time.

Planet's timeline

See also: Geologic time scale and History of Earth

Our knowledge of precise climatic events decreases as we go further back in time. Some notable events are noted below, with a timescale for context.

• Snowball Earth (~800Ma)
• Andean-Saharan glaciation (~450Ma)
• Permian-Triassic extinction event (251.4Ma)
• Paleocene-Eocene Thermal Maximum (Paleocene-Eocene, 55Ma)
• Younger Dryas/The Big Freeze (~11ka)
• Holocene climatic optimum (~7-3ka)
• Climate changes of 535-536 (535-536 AD)
• Medieval warm period (900-1300)
• Little ice age (1300-1800)
• Year Without a Summer (1816)

History of the atmosphere

Earliest atmosphere

The earliest atmosphere of the Earth was probably stripped away by solar winds early in the history of the planet. These gases were later replaced by an atmosphere derived from outgassing from the Earth. Sometime during the late Archean Era an oxygen atmosphere began to develop from photosynthesizing algae.

Carbon dioxide and free oxygen

Free oxygen did not exist until about 1,700 Ma (megaannum, a million years) and this can be seen with the development of the red beds and the end of the banded iron formations. This signifies a shift from a reducing atmosphere to an oxidising atmosphere. The early atmosphere and hydrosphere (up until about 2,000 Ma) were devoid of free oxygen. After photosynthesis developed, photoautotrophs began releasing O₂.

The very early atmosphere of the earth contained mostly carbon dioxide (CO₂): about 80%. This gradually dropped to about 20% by 3,500 Ma. This coincides with the development of the first bacteria about 3,500 Ma. By the time of the development of photosynthesis (2,700 Ma), CO₂ levels in the atmosphere were in the range of 15%. During the period from about 2,700 Ma to about 2,000 Ma, photosynthesis dropped the CO₂ concentrations from about 15% to about 8%. By about 2,000 Ma free O₂ was beginning to accumulate. This gradual reduction in CO₂ levels continued to about 600 Ma at which point CO₂ levels were below 1% and O₂ levels had risen to more than 15%.^[2] 600 Ma corresponds to the end of the Precambrian and the beginning of the Cambrian, the end of the Cryptozoic and the beginning of the Phanerozoic, and the beginning of oxygen-breathing metazoan life.

Precambrian climate

The climate of the late Precambrian was typically cold with glaciation spreading over much of the earth. At this time. the continents were bunched up in a supercontinent called Rodinia. Massive deposits of tillites are found and anomalous isotopic signatures are found which are consistent with the idea that the earth at this time was a massive snowball. Map of Rodinia at the end of the Precambrian after Australia and Antarctica rotated away from the southern hemisphere.

As the Proterozoic Eon drew to a close, the Earth started to warm up. By the dawn of the Cambrian and the Phanerozoic Eon, Earth was experiencing average global temperatures of about +22 °C. Hundreds of millions of years of ice were replaced with the balmy tropical seas of the Cambrian Period within which life exploded at a rate never seen before or after.

Phanerozoic Climate

500 million years of climate change

500 million years of changes in carbon dioxide concentrations

Qualitatively, the Earth's climate has varied between conditions that support large-scale continental glaciation and those which are extensively tropical and lack permanent ice caps even at the poles. The time scale for this variation is roughly 140 million years and may be related to Earth's motion into and out of galactic spiral arms (Veizer and Shaviv 2003). The difference in global mean temperatures between a fully glacial earth and ice free Earth is estimated at approximately 10 °C, though far larger changes would be observed at high latitudes and smaller ones at low latitudes. One key requirement for the development of large scale ice sheets is the arrangement of continental land masses at or near the poles. With plate tectonics constantly rearranging the continents, it can also shape long-term climate evolution. However, the presence of land masses at the poles is not sufficient to guarantee glaciations. Evidence exists of past warm periods in Earth's climate when polar land masses similar to Antarctica were home to deciduous forests rather than ice sheets.

Changes in the atmosphere may also exert an important influence over climate change. The establishment of CO₂-consuming (and oxygen-producing) photosythesizing organisms in the Precambrian led to the production of an atmosphere much like today's, though for most of this period it was much higher in CO₂ than today. Similarly, the Earth's average temperature was also frequently higher than at present, though it has been argued that over very long time scales climate is largely decoupled from carbon dioxide variations (Veizer et al. 2000). Or more specifically that changing continental configurations and mountain building probably have a larger impact on climate than carbon dioxide. Others dispute this, and suggest that the variations of temperature in response to carbon dioxide changes have been underestimated (Royer et al. 2004). However, it is clear that the preindustrial atmosphere with only 280 ppm CO₂ is not far from the lowest ever occurring since the rise of macroscopic life.

Superimposed on the long-term evolution between hot and cold climates have been many short-term fluctuations in climate similar to, and sometimes more severe than, the varying glacial and interglacial states of the present ice age. Some of the most severe fluctuations, such as the Paleocene-Eocene Thermal Maximum, may be related to rapid increases in atmospheric carbon dioxide due to the collapse of natural methane reservoirs in the oceans (see methane clathrates). Severe climate changes also seem to have occurred during the course of the Cretaceous-Tertiary, Permian-Triassic, and Ordovician-Silurian extinction events; however, it is unclear to what degree these changes caused the extinctions rather than merely responding to other processes that may have been more directly responsible for the extinctions.

Quaternary sub-era

The Quaternary sub-era includes the current climate. There has been a cycle of ice ages for the past 2.2-2.1 million years (starting before the Quaternary in the late Neogene Period).

Ice core data for the past 400,000 years. Note length of glacial cycles averages ~100,000 years. Blue curve is temperature, green curve is CO₂, and red curve is windblown glacial dust (loess). Today's date is on the left side of the graph.

Note in the graphic on the right the strong 120,000 year periodicity of the cycles, and the striking asymmetry of the curves. This asymmetry is believed to result from complex interactions of feedback mechanisms. It has been observed that ice ages deepen by progressive steps, but the recovery to interglacial conditions occurs in one big step.

Controlling Factors

Geologically short-term (<120,000 year) temperatures are believed to be driven by orbital factors (see Milankovitch cycles). The arrangements of land masses on the Earth's surface are believed to reinforce these orbital forcing effects.

Continental drift obviously affects the thermohaline circulation, which transfers heat between the equatorial regions and the poles, as does the extent of polar ice coverage.

The timing of ice ages throughout geologic history is in part controlled by the position of the continental plates on the surface of the Earth. When landmasses are concentrated near the polar regions, there is an increased chance for snow and ice to accumulate. Small changes in solar energy can tip the balance between summers in which the winter snow mass completely melts and summers in which the winter snow persists until the following winter. See the web site Paleomap Project for images of the polar landmass distributions through time.

Comparisons of plate tectonic continent reconstructions and paleoclimatic studies show that the Milankovitch cycles have the greatest effect during geologic eras when landmasses have been concentrated in polar regions, as is the case today. Today, Greenland, Antarctica, and the northern portions of Europe, Asia, and North America are situated such that a minor change in solar energy will tip the balance between year-round snow/ice preservation and complete summer melting. The presence of snow and ice is a well-understood positive feedback mechanism for climate. The Earth today is considered to be prone to ice age glaciations.

Another proposed factor in long term temperature change is the Uplift-Weathering Hypothesis, first put forward by T. C. Chamberlin in 1899 and later independently proposed in 1988 by Maureen Raymo and colleagues, where upthrusting mountain ranges expose minerals to weathering resulting in their chemical conversion to carbonates thereby removing CO₂ from the atmosphere and cooling the earth. Others have proposed similar effects due to changes in average water table levels and consequent changes in sub-surface biological activity and PH levels.

Over the very long term the energy output of the sun has gradually increased, on the order of 5% per billion (10⁹) years, and will continue to do so until it reaches the end of its current phase of stellar evolution.

See also

• Geologic temperature record
• Paleothermometry, the study of ancient temperatures
• Paleotempestology, the study of past tropical cyclone activity
• Dendroclimatology
• Historical climatology, the study of climate over human history (as opposed to earth's)
• CLIWOC, Climatological database for the world's oceans (1759-1854)
• Shen Kuo, 11th century Chinese scientist who realized the possibilities of paleoclimatology while observing ancient petrified bamboos buried underground in a northern, dry climate

References

Notes

Bibliography

External links