Paleoclimatology

Paleoclimatology (in British spelling, palaeoclimatology) is the study of changes in climate taken on the scale of the entire history of Earth. It uses a variety of proxy methods from the Earth and life sciences to obtain data previously preserved within things such as rocks, sediments, ice sheets, tree rings, corals, shells, and microfossils. It then uses the records to determine the past states of the Earth's various climate regions and its atmospheric system. Studies of past changes in the environment and biodiversity often reflect on the current situation, specifically the impact of climate on mass extinctions and biotic recovery.[1]

History

The scientific study field of paleoclimate began to form in the early 19th century, when discoveries about glaciations and natural changes in Earth's past climate helped to understand the greenhouse effect.The first observations which had a real scientific basis were probably those by John Hardcastle in New Zealand, in the 1880s. He noted that the loess deposits at Timaru in the South Island recorded changes in climate; he called the loess a 'climate register'.[2]

Reconstructing ancient climates

All palaeotemps
Palaeotemperature graphs compressed together
Sauerstoffgehalt-1000mj
The oxygen content in the atmosphere over the last billion years

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

Ice

Mountain glaciers and the polar ice caps/ice sheets provide much data in paleoclimatology. 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.

  • 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. The trapped air has proven a tremendously valuable source for direct measurement of the composition of air from the time the ice was formed.
  • Layering can be observed because of seasonal pauses in ice accumulation and can be used to establish chronology, associating specific depths of the core with ranges of time.
  • Changes in the layering thickness can be used to determine changes in precipitation or temperature.
  • Oxygen-18 quantity changes (δ18O) in ice layers represent changes in average ocean surface temperature. Water molecules containing the heavier O-18 evaporate at a higher temperature than water molecules containing the normal Oxygen-16 isotope. The ratio of O-18 to O-16 will be higher as temperature increases. It also depends on other factors such as the water's salinity and the volume of water locked up in ice sheets. Various cycles in those isotope ratios have been detected.
  • Pollen has been observed in the ice cores and can be used to understand which plants were present as the layer formed. Pollen is produced in abundance and its distribution is typically well understood. A pollen count for a specific layer can be produced by observing the total amount of pollen categorized by type (shape) in a controlled sample of that layer. Changes in plant frequency over time can be plotted through statistical analysis of pollen counts in the core. Knowing which plants were present leads to an understanding of precipitation and temperature, and types of fauna present. Palynology includes the study of pollen for these purposes.
  • Volcanic ash is contained in some layers, and can be used to establish the time of the layer's formation. Each volcanic event distributed ash with a unique set of properties (shape and color of particles, chemical signature). Establishing the ash's source will establish a range of time to associate with layer of ice.

Dendroclimatology

Climatic information can be obtained through an understanding of changes in tree growth. Generally, trees respond to changes in climatic variables by speeding up or slowing down growth, which in turn is generally reflected by a greater or lesser thickness in growth rings. Different species, however, respond to changes in climatic variables in different ways. A tree-ring record is established by compiling information from many living trees in a specific area.

Older intact wood that has escaped decay can extend the time covered by the record by matching the ring depth changes to contemporary specimens. By using that method, some areas have tree-ring records dating back a few thousand years. Older wood not connected to a contemporary record can be dated generally with radiocarbon techniques. A tree-ring record can be used to produce information regarding precipitation, temperature, hydrology, and fire corresponding to a particular area.

Sedimentary content

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

  • 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 Foraminifera 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.
Core+Repository+core samples2
Sea floor core sample labelled to identify the exact spot on the sea floor where the sample was taken. Sediments from nearby locations can show significant differences in chemical and biological composition.
Sedimentary facies
On a longer time scale, the rock record may show signs of sea level rise and fall, and 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, they include major shifts in climate.

Sclerochronology

Corals (see also sclerochronology)
Coral "rings" are similar to tree rings except that they respond to different things, such as the water temperature, freshwater influx, pH changes, and wave action. From there, certain equipment can be used to derive the sea surface temperature and water salinity from the past few centuries. The δ18O of coralline red algae provides a useful proxy of the combined sea surface temperature and sea surface salinity at high latitudes and the tropics, where many traditional techniques are limited.[3][4]

Landscapes and landforms

Within climatic geomorphology one approach is to study relict landforms to infer ancient climates.[5] Being often concerned about past climates climatic geomorphology is considered sometimes to be a theme of historical geology.[6] Climatic geomorphology is of limited use to study recent (Quaternary, Holocene) large climate changes since there are seldom discernible in the geomorphological record.[7]

Time Scale and Limitations

A multinational consortium, the European Project for Ice Coring in Antarctica (EPICA), has drilled an ice core in Dome C on the East Antarctic ice sheet and retrieved ice from roughly 800,000 years ago.[8] The international ice core community has, under the auspices of International Partnerships in Ice Core Sciences (IPICS), defined a priority project to obtain the oldest possible ice core record from Antarctica, an ice core record reaching back to or towards 1.5 million years ago.[9] The deep marine record, the source of most isotopic data, exists only 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. Resolution and confidence in the data decrease over time.

Notable climate events in Earth history

Knowledge of precise climatic events decreases as the record goes back in time, but some notable climate events are known:

History of the atmosphere

Earliest atmosphere

The first atmosphere would have consisted of gases in the solar nebula, primarily hydrogen. In addition, there would probably have been simple hydrides such as those now found in gas giants like Jupiter and Saturn, notably water vapor, methane, and ammonia. As the solar nebula dissipated, the gases would have escaped, partly driven off by the solar wind.[10]

Second atmosphere

The next atmosphere, consisting largely of nitrogen, carbon dioxide, and inert gases, was produced by outgassing from volcanism, supplemented by gases produced during the late heavy bombardment of Earth by huge asteroids.[10] A major part of carbon dioxide emissions were soon dissolved in water and built up carbonate sediments.

Water-related sediments have been found dating from as early as 3.8 billion years ago.[11] About 3.4 billion years ago, nitrogen was the major part of the then stable "second atmosphere". An influence of life has to be taken into account rather soon in the history of the atmosphere because hints of early life forms have been dated to as early as 3.5 billion years ago.[12] The fact that it is not perfectly in line with the 30% lower solar radiance (compared to today) of the early Sun has been described as the "faint young Sun paradox".

The geological record, however, shows a continually relatively warm surface during the complete early temperature record of Earth with the exception of one cold glacial phase about 2.4 billion years ago. In the late Archaean eon, an oxygen-containing atmosphere began to develop, apparently from photosynthesizing cyanobacteria (see Great Oxygenation Event) which have been found as stromatolite fossils from 2.7 billion years ago. The early basic carbon isotopy (isotope ratio proportions) was very much in line with what is found today, suggesting that the fundamental features of the carbon cycle were established as early as 4 billion years ago.

Third atmosphere

The constant rearrangement of continents by plate tectonics influences the long-term evolution of the atmosphere by transferring carbon dioxide to and from large continental carbonate stores. Free oxygen did not exist in the atmosphere until about 2.4 billion years ago, during the Great Oxygenation Event, and its appearance is indicated by the end of the banded iron formations. Until then, any oxygen produced by photosynthesis was consumed by oxidation of reduced materials, notably iron. Molecules of free oxygen did not start to accumulate in the atmosphere until the rate of production of oxygen began to exceed the availability of reducing materials. That point was a shift from a reducing atmosphere to an oxidizing atmosphere. O2 showed major variations until reaching a steady state of more than 15% by the end of the Precambrian.[13] The following time span was the Phanerozoic eon, during which oxygen-breathing metazoan life forms began to appear.

The amount of oxygen in the atmosphere has fluctuated over the last 600 million years, reaching a peak of 35%[14] during the Carboniferous period, significantly higher than today's 21%. Two main processes govern changes in the atmosphere: plants use carbon dioxide from the atmosphere, releasing oxygen and the breakdown of pyrite and volcanic eruptions release sulfur into the atmosphere, which oxidizes and hence reduces the amount of oxygen in the atmosphere. However, volcanic eruptions also release carbon dioxide, which plants can convert to oxygen. The exact cause of the variation of the amount of oxygen in the atmosphere is not known. Periods with much oxygen in the atmosphere are associated with rapid development of animals. Today's atmosphere contains 21% oxygen, which is high enough for rapid development of animals.[15]

Climate during geological ages

GlaciationsinEarthExistancelicenced annotated
Timeline of glaciations, shown in blue

Precambrian climate

The climate of the late Precambrian showed some major glaciation events spreading over much of the earth. At this time the continents were bunched up in the Rodinia supercontinent. Massive deposits of tillites and anomalous isotopic signatures are found, which gave rise to the Snowball Earth hypothesis. As the Proterozoic Eon drew to a close, the Earth started to warm up. By the dawn of the Cambrian and the Phanerozoic, life forms were abundant in the Cambrian explosion with average global temperatures of about 22 °C.

Phanerozoic climate

Phanerozoic Climate Change
500 million years of climate change

Major drivers for the preindustrial ages have been variations of the sun, volcanic ashes and exhalations, relative movements of the earth towards the sun, and tectonically induced effects as for major sea currents, watersheds, and ocean oscillations. In the early Phanerozoic, increased atmospheric carbon dioxide concentrations have been linked to driving or amplifying increased global temperatures.[16] Royer et al. 2004[17] found a climate sensitivity for the rest of the Phanerozoic which was calculated to be similar to today's modern range of values.

The difference in global mean temperatures between a fully glacial Earth and an 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 requirement for the development of large scale ice sheets seems to be the arrangement of continental land masses at or near the poles. The constant rearrangement of continents by plate tectonics can also shape long-term climate evolution. However, the presence or absence of land masses at the poles is not sufficient to guarantee glaciations or exclude polar ice caps. 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.

The relatively warm local minimum between Jurassic and Cretaceous goes along with an increase of subduction and mid-ocean ridge volcanism [18] due to the breakup of the Pangea supercontinent.

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 climate changes due to sudden collapses of natural methane clathrate reservoirs in the oceans.[19]

A similar, single event of induced severe climate change after a meteorite impact has been proposed as reason for the Cretaceous–Paleogene extinction event. Other major thresholds are the Permian-Triassic, and Ordovician-Silurian extinction events with various reasons suggested.

Quaternary climate

"EDC TempCO2Dust"
Ice core data for the past 800,000 years (x-axis values represent "age before 1950", so today's date is on the left side of the graph and older time on the right). Blue curve is temperature,[20] red curve is atmospheric CO2 concentrations,[21] and brown curve is dust fluxes.[22][23] Note length of glacial-interglacial cycles averages ~100,000 years.

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).

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.

Holocene Temperature Variations
Holocene Temperature Variations

The graph on the left shows the temperature change over the past 12,000 years, from various sources. The thick black curve is an average.

Climate forcings

Radiative-forcings
Radiative forcings, IPCC (2007)

Climate forcing is the difference between radiant energy (sunlight) received by the Earth and the outgoing longwave radiation back to space. Radiative forcing is quantified based on the CO2 amount in the tropopause, in units of watts per square meter to the Earth's surface.[24] Dependent on the radiative balance of incoming and outgoing energy, the Earth either warms up or cools down. Earth radiative balance originates from changes in solar insolation and the concentrations of greenhouse gases and aerosols. Climate change may be due to internal processes in Earth sphere's and/or following external forcings.[25]

Internal processes and forcings

The Earth's climate system involves the atmosphere, biosphere, cryosphere, hydrosphere, and lithosphere,[26] and the sum of these processes from Earth's spheres is what affects the climate. Greenhouse gasses act as the internal forcing of the climate system. Particular interests in climate science and paleoclimatology focus on the study of Earth climate sensitivity, in response to the sum of forcings.

Examples:

External forcings

  • The Milankovitch cycles determine Earth distance and position to the Sun. The solar insolation is the total amount of solar radiation received by Earth.
  • Volcanic eruptions are considered an external forcing.[27]
  • Human changes of the composition of the atmosphere or land use.[27]

Mechanisms

On timescales of millions of years, the uplift of mountain ranges and subsequent weathering processes of rocks and soils and the subduction of tectonic plates, are an important part of the carbon cycle.[28][29][30] The weathering sequesters CO2, by the reaction of minerals with chemicals (especially silicate weathering with CO2) and thereby removing CO2 from the atmosphere and reducing the radiative forcing. The opposite effect is volcanism, responsible for the natural greenhouse effect, by emitting CO2 into the atmosphere, thus affecting glaciation (Ice Age) cycles. James Hansen suggested that humans emit CO2 10,000 times faster than natural processes have done in the past.[31]

Ice sheet dynamics and continental positions (and linked vegetation changes) have been important factors in the long term evolution of the earth's climate.[32] There is also a close correlation between CO2 and temperature, where CO2 has a strong control over global temperatures in Earth history.[33]

See also

References

Notes

  1. ^ Sahney, S. & Benton, M.J. (2008). "Recovery from the most profound mass extinction of all time" (PDF). Proceedings of the Royal Society B: Biological Sciences. 275 (1636): 759–65. doi:10.1098/rspb.2007.1370. PMC 2596898. PMID 18198148.
  2. ^ Hardcastle,J. 1890. On the Timaru loess as a climate register. Transactions and Proceedings of the New Zealand Institute 23, 324-332. also Loess Letter 71, www.loessletter.msu.edu
  3. ^ Halfar, J.; Steneck, R.S.; Joachimski, M.; Kronz, A.; Wanamaker, A.D. (2008). "Coralline red algae as high-resolution climate recorders". Geology. 36 (6): 463. Bibcode:2008Geo....36..463H. doi:10.1130/G24635A.1.
  4. ^ Cobb, K.; Charles, C. D.; Cheng, H; Edwards, R. L. (2003). "El Nino/Southern Oscillation and tropical Pacific climate during the past millennium". Nature. 424 (6946): 271–6. Bibcode:2003Natur.424..271C. doi:10.1038/nature01779. PMID 12867972.
  5. ^ Gutiérrez, Mateo; Gutiérrez, Francisco (2013). "Climatic Geomorphology". Treatise on Geomorphology. 13. pp. 115–131.
  6. ^ Gutiérrez, Mateo, ed. (2005). "Chapter 1 Climatic geomorphology". Developments in Earth Surface Processes. 8. pp. 3–32. doi:10.1016/S0928-2025(05)80051-3. ISBN 978-0-444-51794-4.
  7. ^ Goudie, A.S. (2004). "Climatic geomorphology". In Goudie, A.S. (ed.). Encyclopedia of Geomorphology. pp. 162–164.
  8. ^ Jouzel, Jean; Masson-Delmotte, V.; Cattani, O.; Dreyfus, G.; Falourd, S.; Hoffmann, G.; Minster, B.; Nouet, J.; et al. (10 August 2007). "Orbital and Millennial Antarctic Climate Variability over the Past 800,000 Years". Science. 317 (5839): 793–796. Bibcode:2007Sci...317..793J. doi:10.1126/science.1141038. PMID 17615306.
  9. ^ "Page 1 1 International Partnerships in Ice Core Sciences (IPICS) The oldest ice core: A 1.5 million year record of climate and greenhouse gases from Antarctica". Retrieved 22 September 2011.
  10. ^ a b Zahnle, K.; Schaefer, L.; Fegley, B. (2010). "Earth's Earliest Atmospheres". Cold Spring Harbor Perspectives in Biology. 2 (10): a004895. doi:10.1101/cshperspect.a004895. PMC 2944365. PMID 20573713.
  11. ^ B. Windley: The Evolving Continents. Wiley Press, New York 1984
  12. ^ J. Schopf: Earth's Earliest Biosphere: Its Origin and Evolution. Princeton University Press, Princeton, N.J., 1983
  13. ^ Christopher R. Scotese, Back to Earth History: Summary Chart for the Precambrian, Paleomar Project
  14. ^ Beerling, David (2007). The emerald planet: how plants changed Earth's history. Oxford University press. p. 47. ISBN 9780192806024.
  15. ^ Peter Ward:[1] Out of Thin Air: Dinosaurs, Birds, and Earth's Ancient Atmosphere
  16. ^ Rosemarie E. Came, John M. Eiler, Jan Veizer, Karem Azmy, Uwe Brand & Christopher R. Weidman; Eiler; Veizer; Azmy; Brand; Weidman (September 2007). "CO
    2
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    CS1 maint: Multiple names: authors list (link)
  17. ^ Royer, Dana L. and Robert A. Berner, Isabel P. Montañez, Neil J. Tabor, David J. Beerling (July 2004). "CO2 as a primary driver of Phanerozoic climate". GSA Today. 14 (3): 4–10. doi:10.1130/1052-5173(2004)014<4:CAAPDO>2.0.CO;2.CS1 maint: Multiple names: authors list (link)
  18. ^ Douwe G. Van Der Meer, Richard E. Zeebe, Douwe J. J. van Hinsbergen, Appy Sluijs, Wim Spakman, and Trond H. Torsvik (February 2014). "Plate tectonic controls on atmospheric CO2 levels since the Triassic". PNAS. 111 (12): 4380–4385. Bibcode:2014PNAS..111.4380V. doi:10.1073/pnas.1315657111. PMC 3970481. PMID 24616495.CS1 maint: Multiple names: authors list (link)
  19. ^ Frieling, Joost; Svensen, Henrik H.; Planke, Sverre; Cramwinckel, Margot J.; Selnes, Haavard; Sluijs, Appy (25 October 2016). "Thermogenic methane release as a cause for the long duration of the PETM". Proceedings of the National Academy of Sciences. 113 (43): 12059–12064. Bibcode:2016PNAS..11312059F. doi:10.1073/pnas.1603348113. ISSN 0027-8424. PMC 5087067. PMID 27790990.
  20. ^ Jouzel, J.; Masson-Delmotte, V.; Cattani, O.; Dreyfus, G.; Falourd, S.; Hoffmann, G.; Minster, B.; Nouet, J.; Barnola, J. M. (10 August 2007). "Orbital and Millennial Antarctic Climate Variability over the Past 800,000 Years". Science. 317 (5839): 793–796. Bibcode:2007Sci...317..793J. doi:10.1126/science.1141038. ISSN 0036-8075. PMID 17615306.
  21. ^ Lüthi, Dieter; Le Floch, Martine; Bereiter, Bernhard; Blunier, Thomas; Barnola, Jean-Marc; Siegenthaler, Urs; Raynaud, Dominique; Jouzel, Jean; Fischer, Hubertus (15 May 2008). "High-resolution carbon dioxide concentration record 650,000–800,000 years before present". Nature. 453 (7193): 379–382. Bibcode:2008Natur.453..379L. doi:10.1038/nature06949. ISSN 0028-0836. PMID 18480821.
  22. ^ Lambert, F.; Delmonte, B.; Petit, J. R.; Bigler, M.; Kaufmann, P. R.; Hutterli, M. A.; Stocker, T. F.; Ruth, U.; Steffensen, J. P. (3 April 2008). "Dust-climate couplings over the past 800,000 years from the EPICA Dome C ice core". Nature. 452 (7187): 616–619. Bibcode:2008Natur.452..616L. doi:10.1038/nature06763. ISSN 0028-0836. PMID 18385736.
  23. ^ Lambert, F.; Bigler, M.; Steffensen, J. P.; Hutterli, M.; Fischer, H. (2012). "Centennial mineral dust variability in high-resolution ice core data from Dome C, Antarctica". Climate of the Past. 8 (2): 609–623. Bibcode:2012CliPa...8..609L. doi:10.5194/cp-8-609-2012.
  24. ^ IPCC (2007). "Concept of Radiative Forcing". IPCC.
  25. ^ IPCC (2007). "What are Climate Change and Climate Variability?". IPCC.
  26. ^ "Glossary, Climate system". NASA.
  27. ^ a b "Annex III: Glossary" (PDF). IPCC AR5. Climate change may be due to natural internal processes or external forcings, such as modulations of the solar cycles, volcanic eruptions, and persistent anthropogenic changes in the composition of the atmosphere or in land use.
  28. ^ Caldeira, Ken (18 June 1992). "Enhanced Cenozoic chemical weathering and the subduction of pelagic carbonate". Nature. 357 (6379): 578–581. Bibcode:1992Natur.357..578C. doi:10.1038/357578a0.
  29. ^ Cin-Ty Aeolus Lee, Douglas M. Morton, Mark G. Little, Ronald Kistler, Ulyana N. Horodyskyj, William P. Leeman, and Arnaud Agranier (28 January 2008). "Regulating continent growth and composition by chemical weathering". PNAS. 105 (13): 4981–4986. Bibcode:2008PNAS..105.4981L. doi:10.1073/pnas.0711143105. PMC 2278177. PMID 18362343.CS1 maint: Uses authors parameter (link)
  30. ^ van der Meer, Douwe (25 March 2014). "Plate tectonic controls on Atmospheric CO2 since the Triassic". PNAS. 111 (12): 4380–4385. Bibcode:2014PNAS..111.4380V. doi:10.1073/pnas.1315657111. PMC 3970481. PMID 24616495.
  31. ^ James Hansen (2009). "The 8 Minute Epoch 65 million Years with James Hansen". University of Oregon.
  32. ^ ROYER, D. L.; PAGANI, M.; BEERLING, D. J. (1 July 2012). "Geobiological constraints on Earth system sensitivity to CO2 during the Cretaceous and Cenozoic". Geobiology. 10 (4): 298–310. doi:10.1111/j.1472-4669.2012.00320.x. PMID 22353368.
  33. ^ Royer, Dana L. (1 December 2006). "CO2-forced climate thresholds during the Phanerozoic". Geochimica et Cosmochimica Acta. 70 (23): 5665–5675. Bibcode:2006GeCoA..70.5665R. doi:10.1016/j.gca.2005.11.031.

Bibliography

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  • Cronin, Thomas N. (2010). Paleoclimates: understanding climate change past and present. New York: Columbia University Press. ISBN 978-0-231-14494-0.
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  • Margulis, Lynn; Sagan, Dorion (1986). Origins of sex: three billion years of genetic recombination. The Bio-origins series. New Haven: Yale University Press. ISBN 978-0-300-03340-3.
  • Gould, Stephen Jay (1989). Wonderful life, the story of the Burgess Shale. New York: W.W. Norton. ISBN 978-0-393-02705-1.
  • Crowley, Thomas J.; North, Gerald R. (1996). Paleoclimatology. Oxford monographs on geology and geophysics. 18. Oxford: Clarendon Press. ISBN 978-0-19-510533-9.
  • The Climates of the Geological Past. (Die Klimate der geologischen Vorzeit). 1924, Wladimir Köppen, Alfred Wegener
  • Karl-Heinz Ludwig (2006). Eine kurze Geschichte des Klimas. Von der Entstehung der Erde bis heute, (A short history of climate, From the evolution of earth till today) Herbst, ISBN 3-406-54746-X
  • William F. Ruddimann (2001). Earth's Climate — Past and Future. Palgrave Macmillan. ISBN 978-0-7167-3741-4.
  • B. Windley (1984). The Evolving Continents. New York: Wiley Press.
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External links

  • Quintana, Favia et al., 2018 ″Multiproxy response to climate- and human-driven changes in a remote lake of southern Patagonia (Laguna Las Vizcachas, Argentina) during the last 1.6 kyr″, Boletín de la Sociedad Geológica Mexicana, Mexico, VOL. 70 NO. 1 P. 173 ‒ 186 https://dx.doi.org/10.18268/BSGM2018v70n1a10
Atlantic (period)

The Atlantic in palaeoclimatology was the warmest and moistest Blytt-Sernander period, pollen zone and chronozone of Holocene northern Europe. The climate was generally warmer than today. It was preceded by the Boreal, with a climate similar to today’s, and was followed by the Subboreal, a transition to the modern. Because it was the warmest period of the Holocene, the Atlantic is often referenced more directly as the Holocene climatic optimum, or just climatic optimum.

Blytt–Sernander system

The Blytt-Sernander classification, or sequence, is a series of north European climatic periods or phases based on the study of Danish peat bogs by Axel Blytt (1876) and Rutger Sernander (1908). The classification was incorporated into a sequence of pollen zones later defined by Lennart von Post, one of the founders of palynology.

Climate

Climate is the statistics of weather over long periods of time. It is measured by assessing the patterns of variation in temperature, humidity, atmospheric pressure, wind, precipitation, atmospheric particle count and other meteorological variables in a given region over long periods of time. Climate differs from weather, in that weather only describes the short-term conditions of these variables in a given region.

A region's climate is generated by the climate system, which has five components: atmosphere, hydrosphere, cryosphere, lithosphere, and biosphere.The climate of a location is affected by its latitude, terrain, and altitude, as well as nearby water bodies and their currents. Climates can be classified according to the average and the typical ranges of different variables, most commonly temperature and precipitation. The most commonly used classification scheme was the Köppen climate classification. The Thornthwaite system, in use since 1948, incorporates evapotranspiration along with temperature and precipitation information and is used in studying biological diversity and how climate change affects it. The Bergeron and Spatial Synoptic Classification systems focus on the origin of air masses that define the climate of a region.

Paleoclimatology is the study of ancient climates. Since direct observations of climate are not available before the 19th century, paleoclimates are inferred from proxy variables that include non-biotic evidence such as sediments found in lake beds and ice cores, and biotic evidence such as tree rings and coral. Climate models are mathematical models of past, present and future climates. Climate change may occur over long and short timescales from a variety of factors; recent warming is discussed in global warming. Global warming results in redistributions. For example, "a 3°C change in mean annual temperature corresponds to a shift in isotherms of approximately 300–400 km in latitude (in the temperate zone) or 500 m in elevation. Therefore, species are expected to move upwards in elevation or towards the poles in latitude in response to shifting climate zones".

Climatic geomorphology

Climatic geomorphology is the study of the role of climate in shaping landforms and the earth-surface processes. An approach used in climatic geomorphology is to study relict landforms to infer ancient climates. Being often concerned about past climates climatic geomorphology considered sometimes to be an aspect of historical geology. Since landscape features in one region might have evolved under climates different from those of the present, studying climatically disparate regions might help understand present-day landscapes. For example, Julius Büdel studied both cold-climate processes in Svalbard and weathering processes in tropical India to understand the origin of the relief of Central Europe, which he argued was a palimpsest of landforms formed at different times and under different climates.

Dry Andes

The Dry Andes (Spanish: Andes áridos) is a climatic and glaciological subregion of the Andes. Together with the Wet Andes it is one of the two subregions of the Argentine and Chilean Andes. The Dry Andes runs from the Atacama Desert in northern Chile and Northwest Argentina south to a latitude of 35°S in Chile. In Argentina the Dry Andes reaches 40°S due to the leeward effect of the Andes. According to Luis Lliboutry

the Dry Andes can be defined by the distribution of penitentes. The southernmost well developed penitentes are found on Lanín Volcano.

Geohazard

A geohazard is a geological state that may lead to widespread damage or risk. Geohazards are geological and environmental conditions and involve long-term or short-term geological processes. Geohazards can be relatively small features, but they can also attain huge dimensions (e.g., submarine or surface landslide) and affect local and regional socio-economy to a large extent (e.g., tsunamis).

Geologic temperature record

The Geologic temperature record are changes in Earth's environment as determined from geologic evidence on multi-million to billion (109) year time scales. The study of past temperatures provides an important paleoenvironmental insight because it is a component of the climate and oceanography of the time.

Late Pleistocene

The Late Pleistocene is a geochronological age of the Pleistocene Epoch and is associated with Upper Pleistocene (or Tarantian) stage rocks. The beginning of the stage is defined by the base of the Eemian interglacial phase before the final glacial episode of the Pleistocene 126,000 ± 5,000 years ago. Its end is defined at the end of the Younger Dryas, some 11,700 years ago. The age represents the end of the Pleistocene epoch and is followed by the Holocene epoch.

Much of the Late Pleistocene age was dominated by glaciations, such as the Wisconsin glaciation in North America and the Weichselian glaciation and Würm glaciation in Eurasia). Many megafauna became extinct during this age, a trend that continued into the Holocene. The Late Pleistocene contains the Upper Paleolithic stage of human development, including the out-of-Africa migration and dispersal of anatomically modern humans and the extinction of the last remaining archaic human species.

Lonnie Thompson

Lonnie Thompson (born July 1, 1948), is an American paleoclimatologist and Distinguished University Professor in the School of Earth Sciences at The Ohio State University. He has achieved global recognition for his drilling and analysis of ice cores from mountain glaciers and ice caps in the tropical and sub-tropical regions of the world. He and his wife, Ellen Mosley-Thompson, run the ice core paleoclimatology research group at the Byrd Polar Research Center.

Mediterranean Basin

In biogeography, the Mediterranean Basin (also known as the Mediterranean region or sometimes Mediterranea) is the region of lands around the Mediterranean Sea that have a Mediterranean climate, with mild, rainy winters and hot, dry summers, which supports characteristic Mediterranean forests, woodlands, and scrub vegetation.

Oldest Dryas

The Oldest Dryas was a climatic period, which occurred during the coldest stadial after the Weichselian glaciation in north Europe. In the Alps, the Oldest Dryas corresponds to the Gschnitz stadial of the Würm glaciation. The three “Dryas” periods (younger, older, oldest) are named for a marker species, Dryas octopetala, detected in core samples of glacial ice and peat bogs. The Oldest Dryas corresponds to pollen zone Ia.

Palaeogeography, Palaeoclimatology, Palaeoecology

Palaeogeography, Palaeoclimatology, Palaeoecology ("Palaeo3") is a peer-reviewed scientific journal publishing multidisciplinary studies and comprehensive reviews in the field of palaeoenvironmental geology. The journal is edited by D. J. Bottjer, T. Corrège, A. P. Kershaw, and F. Surlyk. It was established in 1965 and is currently published by Elsevier.

Paleosol

In the geosciences, paleosol (palaeosol in Great Britain and Australia) can have two meanings. The first meaning, common in geology and paleontology, refers to a former soil preserved by burial underneath either sediments (alluvium or loess) or volcanic deposits (volcanic ash), which in the case of older deposits have lithified into rock. In Quaternary geology, sedimentology, paleoclimatology, and geology in general, it is the typical and accepted practice to use the term "paleosol" to designate such "fossil soils" found buried within either sedimentary or volcanic deposits exposed in all continents as illustrated by Rettallack (2001), Kraus (1999), and other published papers and books.

In soil science, paleosols are soils formed long periods ago that have no relationship in their chemical and physical characteristics to the present-day climate or vegetation.

Such soils form on extremely old continental cratons and as small scattered localities in outliers of ancient rock.

Proxy (climate)

In the study of past climates ("paleoclimatology"), climate proxies are preserved physical characteristics of the past that stand in for direct meteorological measurements and enable scientists to reconstruct the climatic conditions over a longer fraction of the Earth's history. Reliable global records of climate only began in the 1880s, and proxies provide the only means for scientists to determine climatic patterns before record-keeping began.

Examples of proxies include ice cores, tree rings, sub-fossil pollen, boreholes, corals, lake and ocean sediments, and carbonate speleothems. The character of deposition or rate of growth of the proxies' material has been influenced by the climatic conditions of the time in which they were laid down or grew. Chemical traces produced by climatic changes, such as quantities of particular isotopes, can be recovered from proxies. Some proxies, such as gas bubbles trapped in ice, enable traces of the ancient atmosphere to be recovered and measured directly to provide a history of fluctuations in the composition of the Earth's atmosphere. To produce the most precise results, systematic cross-verification between proxy indicators is necessary for accuracy in readings and record-keeping.Proxies can be combined to produce temperature reconstructions longer than the instrumental temperature record and can inform discussions of global warming and climate history. The distribution of proxy records, just like the instrumental record, is not at all uniform, with more records in the northern hemisphere.

Subboreal

The Subboreal is a climatic period, immediately before the present one, of the Holocene. It lasted from 3710 to 450 BCE.

Subfossil

A subfossil is a part of a dead organism that is partially, rather than fully, fossilized, as is a fossil. Partial fossilization may be present because not enough time has elapsed since the animal died for full fossilization, or because the conditions in which the remains were deposited were not optimal for fossilization.

For remains such as molluscan seashells, which frequently do not change their chemical composition over geological time, and may occasionally even retain such features as the original color markings for millions of years, the label "subfossil" is applied to shells that are understood to be thousands of years old, but are of Holocene age, and therefore are not old enough to be from the Pleistocene epoch.

Unfossilized or partially fossilized remains can include bones, exoskeletons, nests, skin imprints, or fecal deposits. Subfossils of vertebrates are often found in caves or other shelters, where the remains have been preserved for thousands of years. The main importance of these vertebrate subfossil (versus fully fossilized) remains is that they contain organic material, which can be used for radiocarbon dating or the extraction and sequencing of DNA, protein, or other biomolecules. Additionally, isotope ratios can provide information about the ecological conditions under which extinct animals lived. Subfossils are useful for studying the evolutionary history of an environment and can be important to studies in paleoclimatology.

Subfossils are also often found in depositionary environments, such as lake sediments, oceanic sediments, and soils. Once deposited, physical and chemical weathering may alter the state of preservation, and small subfossils can also be ingested by living organisms. Subfossil remains that date from the Mesozoic are exceptionally rare, are usually in an advanced state of decay, and are consequently much disputed. The vast bulk of subfossil material comes from Quaternary sediments, including many subfossilized chironomid head capsules, ostracod carapaces, diatoms, and foraminifera.

Timeline of glaciation

There have been five or six major ice ages in the history of Earth over the past 3 billion years.

The Late Cenozoic Ice Age began 34 million years ago, its latest phase being the Quaternary glaciation, in progress since 2.58 million years ago.

Within ice ages, there exist periods of more severe glacial conditions and more temperate referred to as glacial periods and interglacial periods, respectively. The Earth is currently in such an interglacial period of the Quaternary glaciation, with the last glacial period of the Quaternary having ended approximately 11,700 years ago, the current interglacial being known as the Holocene epoch.

Based on climate proxies, paleoclimatologists study the different climate states originating from glaciation.

Δ13C

In geochemistry, paleoclimatology and paleoceanography δ13C (pronounced "delta c thirteen") is an isotopic signature, a measure of the ratio of stable isotopes 13C : 12C, reported in parts per thousand (per mil, ‰).

The definition is, in per mil:

where the standard is an established reference material.

δ13C varies in time as a function of productivity, the signature of the inorganic source, organic carbon burial and vegetation type. Biological processes preferentially take up the lower mass isotope through kinetic fractionation. However some abiotic processes do the same, methane from hydrothermal vents can be depleted by up to 50%.

Δ18O

In geochemistry, paleoclimatology and paleoceanography δ18O or delta-O-18 is a measure of the ratio of stable isotopes oxygen-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-independent fractionation.

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