16.3: Prehistoric Climate Change

Over Earth's history, the climate has changed a lot. For example, during the Mesozoic Era, the age of dinosaurs, the climate was much warmer and carbon dioxide was abundant in the atmosphere. However, throughout the Cenozoic Era (65 million years ago to today), the climate has been gradually cooling. This section summarizes some of these major past climate changes.

Figure \(\PageIndex{1}\): Maximum extent of Laurentide Ice Sheet.

Past Glaciations

Through geologic history, the climate has changed slowly over millions of years. Before the most recent Pliocene-Quaternary glaciation, there were three other major glaciations [20]. The oldest, known as the Huronian, occurred toward the end of the Archean Eon-early Proterozoic Eon (~2.5 billion years ago). The major event of that time, the Great Oxygenation Event, is most commonly associated with the cause of that glaciation. The increased oxygen is thought to have reacted with the potent greenhouse gas methane, causing cooling [21].

The end of the Proterozoic Eon (about 700 million years ago) had other glaciations, part of the Snowball Earth hypothesis [22]. Glacial evidence has been interpreted in widespread rock sequences globally and even has been linked to low-latitude glaciation [23]. Limestone rock (usually formed in tropical marine environments) and glacial deposits (usually formed in cold climates) are often found together from this time in regions all around the world. In Utah, Antelope Island in the Great Salt Lake has interbedded limestone and glacial deposits (diamictites) interpreted to be formed by continental glaciation [24].

The idea of the controversial Snowball Earth hypothesis is that a runaway albedo effect (ice and snow reflecting solar radiation) might cause the complete freezing of land and ocean surfaces and a collapse of biological activity. The ice-covered Earth would only melt from greenhouse heating when carbon dioxide from volcanoes reached high concentrations, due to the inability for carbon dioxide to enter the then-frozen ocean. Some studies estimated carbon dioxide was 350 times higher than today’s concentrations [22]. The complete freezing [25] and the extent of the freezing [26] has come into question because biological activity did survive. A competing hypothesis is the Slushball Earth hypothesis in which some regions of the equatorial ocean remained open. Differing scientific conclusions about the stability of Earth’s magnetic poles, impacts on ancient rock evidence from subsequent metamorphism, and alternate interpretations of existing evidence keep the idea of Snowball Earth controversial.

Glaciation also occurred in the Paleozoic Era, notably the Andean-Saharan glaciation in the late Ordovician, about 440–460 million years ago, which coincided with a major extinction event, and the Karoo Ice Age during the Pennsylvanian Period, 323 to 300 million years ago. This glaciation was one of the evidences cited by Wegener for his Continental Drift hypothesis as his proposed Pangea drifted into south polar latitudes. This also was caused by an increase of oxygen and a subsequent drop in carbon dioxide, most likely produced by the evolution and rise of land plants [27].

During the Cenozoic Era (the last 65 million years), the climate started out warm and gradually cooled to today. This warm time is called the Paleocene-Eocene Thermal Maximum and Antarctica and Greenland were ice-free during this time. Since the Eocene, tectonic events during the Cenozoic caused persistent and significant planetary cooling. For example, the collision of the Indian Plate with the Asian Plate created the Himalaya Mountains, increasing weathering and erosion rates. An increased rate of weathering of silicate minerals, especially feldspar, consumes carbon dioxide from the atmosphere and therefore reduces the greenhouse effect, resulting in long-term cooling [28].

Figure \(\PageIndex{2}\): Global average surface temperature over the past 70 million years.

At about 40 million years ago, the narrow gap between the South American Plate and the Antarctica Plate widened, resulting in the opening of the Drake Passage. This allowed for the unrestricted west-to-east flow of water around Antarctica, the Antarctic Circumpolar Current, which effectively isolated the southern ocean from the warmer waters of the Pacific, Atlantic, and Indian Oceans. The region cooled significantly, and by 35 million years ago (Oligocene) glaciers had started to form on Antarctica [29].

Figure \(\PageIndex{3}\): The Antarctic Circumpolar Current (ACC).

At around 15 million years ago, subduction-related volcanism between Central and South America created the Isthmus of Panama that connected North and South America. This prevented water from flowing between the Pacific and Atlantic Oceans and reduced heat transfer from the tropics to the poles. This reduced heat transfer and created a cooler Antarctica and larger Antarctic glaciers. The expansion of that ice sheet (on land and water) increased Earth’s reflectivity of solar radiation. This enhanced the albedo effect, creating a positive feedback loop of further cooling: more reflective glacial ice, more cooling, more ice, and so on [30; 31].

By 5 million years ago (Pliocene Epoch), ice sheets had started to grow in North America and northern Europe. The most intense part of the current glaciation is the last 1 million years of the Pleistocene Epoch. The Pleistocene has significant temperature variations (through a range of almost 10°C or 18°F) on time scales of 40,000 to 100,000 years, and corresponding expansion and contraction of ice sheets. These variations are attributed to subtle changes in Earth’s orbital parameters called Milankovitch cycles [32; 33]. Over the past million years, the glaciation cycles have been approximately every 100,000 years [34] with many glacial advances in the last 2 million years (Lisiecki and Raymo, 2005) [35].

Figure \(\PageIndex{4}\): A Pliocene‐Pleistocene stack of 57 globally distributed benthic δ18O records, as a proxy for temperature. Higher values indicate cooling, and lower values indicate warming. The overall trend illustrates the glaciation cycles. The x-axis represents time in thousands of years (ka) so 200 means 200,000 years. (Source: Lisiecki and Raymo, 2005)

During an ice age, periods of warming climate are called interglacials; during interglacials, very brief periods of even warmer climate are called interstadials. These warming upticks are related to Earth’s climate variations, like Milankovitch cycles, which are changes to the Earth’s orbit that can fluctuate climate. In the last 500,000 years, there have been 5 or 6 interglacials, with the most recent belonging to our current time, the Holocene.

Two of the more recent climate swings demonstrate the complexity of the changes: the Younger Dryas and the Holocene Climatic Optimum. These events are more recent and yet have conflicting information. The Younger Dryas cooling is widely recognized in the Northern Hemisphere [36], though the timing of the event (about 12,000 years ago) does not appear to be equal everywhere [37]. It also is difficult to find in the Southern Hemisphere [38]. The Holocene Climatic Optimum is the warming around 6,000 years ago [39], though it was not universally warmer, and probably not as warm as current warming [40], and not at the same time everywhere [41].

Proxy Indicators of Past Climates

How do we know about past climates? Geologists use proxy indicators to understand past climate. A proxy indicator is a biological, chemical, or physical signature preserved in the rock, sediment, or ice record that acts like a “fingerprint” of something in the past [42]. Thus they are an indirect indicator of something like climate. For ancient glaciations from the Proterozoic and Paleozoic, there are rock formations of glacial sediments such as the diamictite (or tillite) of the Mineral Fork Formation in Utah. This dark rock has many fine-grained components plus some large out-sized clasts like a modern glacial till [43; 44].

Deep-sea sediment is an indirect indicator of climate change during the Cenozoic Era, about the last 65 million years. Researchers from the Ocean Drilling Program, an international research collaboration, collect deep-sea sediment cores that record continuous sediment accumulation. The sediment provides detailed chemical records of stable carbon and oxygen isotopes obtained from deep-sea benthic foraminifera shells that accumulated on the ocean floor over millions of years. Oxygen isotopes are a proxy indicator of deep-sea temperatures and continental ice volume [45].

Sediment Cores – Stable Oxygen Isotope

Oxygen isotopes are an indicator of past climate. The two main stable oxygen isotopes are ¹⁶O and ¹⁸O. They both occur in water (H₂O) and in the calcium carbonate (CaCO₃) shells of foraminifera as the oxygen component of both of those molecules. The most abundant and lighter isotope is ¹⁶O. Since it is lighter, it evaporates more easily from the ocean’s surface as water vapor, which later turns to clouds and precipitation on the ocean and land. This evaporation is enhanced in warmer sea water and slightly increases the concentration of ¹⁸O in the surface seawater from which the plankton derives the carbonate for its shells. Thus the ratio of ¹⁶O and ¹⁸O in the fossilized shells in seafloor sediment is a proxy indicator of the temperature and evaporation of seawater.

Figure \(\PageIndex{5}\): Sediment cores from the Greenland continental slope (Source: Hannes Grobe)

Keep in mind, it is harder to evaporate the heavier water and easier to condense it. As evaporated water vapor drifts toward the poles and tiny droplets form clouds and precipitation, droplets of water with ¹⁸O tend to form more readily than droplets of the lighter form and precipitate out, leaving the drifting vapor depleted in ¹⁸O.

Figure \(\PageIndex{6}\): Antarctic temperature changes during the last few glaciations compared to global ice volume. The first two curves are based on the deuterium (heavy hydrogen) record from ice cores (EPICA Community Members 2004, Petit et al. 1999). The bottom line is ice volume based on oxygen isotopes from a composite of deep-sea sediment cores (Lisiecki and Raymo 2005).

During geologic times when the climate is cooler, more of this lighter precipitation is locked onto land in the form of glacial ice. Consider the giant ice sheets, more than a mile thick, that covered a large part of North America during the last ice age only 14,000 years ago. During glaciation, the glaciers effectively lock away more ¹⁶O, thus the ocean water and foraminifera shells become enriched in ¹⁸O. Therefore, a ratio of ¹⁸O to ¹⁶O (\(\delta^{18}\)O) in calcium carbonate shells of foraminifera is a proxy indicator of past climate. The sediment cores from the Ocean Drilling Program record a continuous accumulation of sediment of these fossils and the sediment provide a record of glacials, interglacials, and interstadials.

Sediment Cores – Boron Isotopes and Acidity

Ocean acidity is affected by carbonic acid and is a proxy for past atmospheric CO₂ concentrations. To estimate the ocean’s pH (acidity) over the past 60 million years, researchers collected deep-sea sediment cores and examined the ancient planktonic foraminifera shells’ boron-isotope ratios. Boron has two isotopes: ¹¹B and ¹⁰B. In aqueous compounds of boron, the relative abundance of these two isotopes is sensitive to pH (acidity), hence CO₂ concentrations. In the early Cenozoic, around 60 million years ago, CO₂ concentrations were over 2000 ppm and started falling around 55 to 40 million years ago possibly due to reduced CO₂ outgassing from ocean ridges, volcanoes, and metamorphic belts and increased carbon burial due to uplift of the Himalaya Mountains. By the Miocene (about 24 million years ago), CO₂ levels were below 500 ppm [46] and by 800,000 years ago CO₂ levels didn’t exceed 300 ppm [47].

Carbon Dioxide Concentrations in Ice Cores

For the more recent Pleistocene Epoch's climate, there is a more detailed and direct chemical record from coring into the Antarctic and Greenland ice sheets. Snow accumulates on these ice sheets and creates yearly layers. Ice cores have been extracted from ice sheets covering the last 800,000 years. Oxygen isotopes are collected from these annual layers and the ratio of ¹⁸O to ¹⁶O is used to determine temperature as discussed above. In addition, the ice traps small atmospheric gas bubbles as the snow turns to ice. Analysis of these bubbles reveals the composition of the atmosphere at these previous times.

Figure \(\PageIndex{7}\): 19 cm long section of ice core showing 11 annual layers with summer layers (arrowed) sandwiched between darker winter layers. (Source: US Army Corps of Engineers)

Small pieces of this ice are crushed and the ancient air is extracted into a mass spectrometer that can detect the chemistry of the ancient atmosphere. Carbon dioxide levels are recreated from these measurements. Over the last 800,000 years, the maximum carbon dioxide concentration during warm times was about 300 ppm (parts per million) and the minimum during cold stretches was about 170 ppm [46; 47; 48]. The carbon dioxide content of Earth’s atmosphere is currently over 410 ppm.

Figure \(\PageIndex{8}\): Antarctic ice showing hundreds of tiny trapped air bubbles from the atmosphere thousands of years ago. (Source: CSIRO)

Figure \(\PageIndex{9}\): Composite carbon dioxide record from the last 800,000 years based on ice core data from EPICA Dome C Ice Core.

Oceanic Microfossils

Microfossils, like foraminifera, diatoms, and radiolarians, can be used to interpret past climate records. In sediment cores, different species of microfossils are found in different layers. Groups of these microfossils are called assemblages. One assemblage consists of species that lived in cooler ocean water (in glacial times) and another assemblage found at a different level in the same sediment core is made of warmer water species [49]. Watch this video that shows how climate data is captured in sediment cores.

Tree Rings

Tree rings, which form every year as a tree grows, are another past climate indicator. Rings that are thicker indicate wetter years, and rings that are thinner and closer together indicate drier years. Every year a tree will grow one ring with a light section and dark section. The rings vary in width. Since trees need a lot of water to survive, narrower rings indicate colder and drier climates. Since some trees can be several thousand years old, we can use their rings for regional paleoclimatic reconstructions, for example, to reconstruct past temperature, precipitation, vegetation, streamflow, sea-surface temperature, and other climate-dependent conditions. Paleoclimatic study means relating to a distinct past geologic climate. Further, dead trees such as those used in Puebloan ruins can be used to extend this proxy indicator, which showed long term droughts in the region and why their villages were abandoned.

Figure \(\PageIndex{10}\): Tree rings form every year. Rings that are farther apart are from wetter years and rings that are closer together are from drier years.

Figure \(\PageIndex{11}\): Summer temperature anomalies for the past 7000 years from tree ring data. (Source: R.M.Hantemirov)

Pollen

Pollen is also a proxy climate indicator. Flowering plants produce pollen grains. Pollen is distinctive when viewed under a microscope. Sometimes pollen can be preserved in lake sediments that accumulate in layers every year. Coring of lake sediments can reveal ancient pollen. Fossil pollen assemblages are groups of pollen from multiple species such as spruce, pine, and oak. Through time (via the sediment cores and radiometric age-dating techniques), the pollen assemblage will change revealing the plants that lived in the area at the time. Thus the pollen assemblages are an indicator of past climate since different plants will prefer different climates [50]. For example, in the Pacific Northwest east of the Cascades, a region close to the border of grasslands and forest, a study tracked pollen over the last 125,000 years covering the last two glaciations. Pollen assemblages with more pine tree pollen are found during glaciations and pollen assemblages with less pine tree pollen are found during interglacial times [51].

Figure \(\PageIndex{12}\): Scanning electron microscope image of modern pollen with false color added to distinguish plant species. (Source: Dartmouth Electron Microscope Facility, Dartmouth College)

Other Proxy Indicators

Paleoclimatologists study many other phenomena to understand past climates such as human historical accounts, human instrument record from the recent past, lake sediments, cave deposits, and corals.

References

20. Deynoux, M., Miller, J. M. G. & Domack, E. W. Earth’s Glacial Record. (Cambridge University Press, 2004).

21. Kopp, R. E., Kirschvink, J. L., Hilburn, I. A. & Nash, C. Z. The Paleoproterozoic snowball Earth: a climate disaster triggered by the evolution of oxygenic photosynthesis. Proc. Natl. Acad. Sci. U. S. A. 102, 11131–11136 (2005).

22. Hoffman, P. F., Kaufman, A. J., Halverson, G. P. & Schrag, D. P. A neoproterozoic snowball earth. Science 281, 1342–1346 (1998).

23. Schopf, J. W. & Klein, C. Late Proterozoic Low-Latitude Global Glaciation: the Snowball Earth. in The Proterozoic biosphere : a multidisciplinary study (eds. Schopf, J. W. & Klein, C.) 51–52 (Cambridge University Press, 1992).

24. Doelling, H. H. et al. Geology of Antelope Island. Davis County, Utah: Utah Geological and Mineral Survey Open-File Release 144, 82 (1988).

25. Allen, P. A. & Etienne, J. L. Sedimentary challenge to Snowball Earth. Nat. Geosci. 1, 817–825 (2008).

26. Eyles, N. & Januszczak, N. ‘Zipper-rift’: a tectonic model for Neoproterozoic glaciations during the breakup of Rodinia after 750 Ma. Earth-Sci. Rev. (2004).

27. Berner, R. A. The carbon cycle and carbon dioxide over Phanerozoic time: the role of land plants. Philos. Trans. R. Soc. Lond. B Biol. Sci. 353, 75–82 (1998).

28. Kump, L. R., Brantley, S. L. & Arthur, M. A. Chemical Weathering, Atmospheric CO2, and Climate. Annu. Rev. Earth Planet. Sci. 28, 611–667 (2000).

29. Lagabrielle, Y., Goddéris, Y., Donnadieu, Y., Malavieille, J. & Suarez, M. The tectonic history of Drake Passage and its possible impacts on global climate. Earth Planet. Sci. Lett. 279, 197–211 (2009).

30. Kuipers Munneke, P. et al. A new albedo parameterization for use in climate models over the Antarctic ice sheet. J. Geophys. Res. 116, D05114 (2011).

31. Curry, J. A., Schramm, J. L. & Ebert, E. E. Sea Ice-Albedo Climate Feedback Mechanism. J. Clim. 8, 240–247 (1995).

32. Milankovitch, M. Mathematische klimalehre und astronomische theorie der klimaschwankungen. (1930).

33. Roe, G. In defense of Milankovitch. Geophys. Res. Lett. 33, L24703 (2006).

34. Abe-Ouchi, A. et al. Insolation-driven 100,000-year glacial cycles and hysteresis of ice-sheet volume. Nature 500, 190–193 (2013).

35. Lisiecki, L. E. & Raymo, M. E. A Pliocene-Pleistocene stack of 57 globally distributed benthic δ18O records. Paleoceanography 20, (2005).

36. Carlson, A. E. The younger Dryas climate event. (2013).

37. Fairbanks, R. G. The age and origin of the “Younger Dryas climate event” in Greenland ice cores. Paleoceanography 5, 937–948 (1990).

38. Kaplan, M. R. et al. Glacier retreat in New Zealand during the Younger Dryas stadial. Nature 467, 194–197 (2010).

39. Ganopolski, A., Kubatzki, C., Claussen, M., Brovkin, V., V. & Petoukhov, V., V. The influence of vegetation-atmosphere-ocean interaction on climate during the mid-holocene. Science 280, 1916–1919 (1998).

40. Hewitt, C. D. & Mitchell, J. F. B. A fully coupled GCM simulation of the climate of the mid-Holocene. Geophys. Res. Lett. 25, 361–364 (1998).

41. He, Y. et al. Asynchronous Holocene climatic change across China. Quat. Res. 61, 52–63 (2004).

42. Weissert, H. Deciphering methane’s fingerprint. Nature 406, 356–357 (2000).

43. Ojakangas, R. W. & Matsch, C. L. Upper Precambrian (Eocambrian) Mineral Fork Tillite of Utah: A continental glacial and glaciomarine sequence. GSA Bulletin 91, 495–501 (1980).

44. Christie-Blick, N., Ojakangas, R. W. & Matsch, C. L. Upper Precambrian (Eocambrian) Mineral Fork Tillite of Utah: A Continental Glacial and Glaciomarine Sequence: Discussion and Reply. The Geological Society of America Bulletin 93, 184–187 (1982).

45. Zachos, J., Pagani, M., Sloan, L., Thomas, E. & Billups, K. Trends, rhythms, and aberrations in global climate 65 Ma to present. Science 292, 686–693 (2001).

46. Pearson, P. N. & Palmer, M. R. Atmospheric carbon dioxide concentrations over the past 60 million years. Nature 406, 695–699 (2000).

47. Lüthi, D. et al. High-resolution carbon dioxide concentration record 650,000–800,000 years before present. Nature 453, 379–382 (2008).

48. Oak Ridge National Laboratory. 800,000-year Ice-Core Records of Atmospheric Carbon Dioxide (CO2). (2008). Available at: http://cdiac.ornl.gov/trends/co2/ice_core_co2.html. (Accessed: 14th September 2016)

49. Cunningham, W. L., Leventer, A., Andrews, J. T., Jennings, A. E. & Licht, K. J. Late Pleistocene–Holocene marine conditions in the Ross Sea, Antarctica: evidence from the diatom record. The Holocene 9, 129–139 (1999).

50. Webb, T. & Thompson, W. Is vegetation in equilibrium with climate? How to interpret late-Quaternary pollen data. Vegetatio 67, 75–91 (1986).

51. Whitlock, C. & Bartlein, P. J. Vegetation and climate change in northwest America during the past 125 kyr. Nature 388, 57–61 (1997).