Paleoclimatology: Earth Systems Change Through Time
What is Climate?
Climate is the average of the weather for a region over time. There is a common confusion that arises when people use the word climate. Often, even if they think they are discussing climate, the conversation will begin with a discussion of the current weather. It might go something like this:
“It was really hot today! Must be global warming…”
And, indeed, it might be. However, what they are actually describing is the weather. Weather is the summation of current atmospheric conditions at any given moment. And, to measure it, we use things like thermometers, barometers, anemometers, and other devices. This information is collected and is used to develop models that can be used for predictive purposes. From these, you get your 1-10 day forecasts. These forecasts of the weather are entirely based on probabilities. A 20% chance of rain just means that, under current weather conditions (or those predicted), it tends to rain 20% of the time. These conditions are a part of a region’s climate story, but they are still defined as weather. While they use recent historical data, they are intended to describe the condition of the atmosphere at a moment in time…perhaps three days from now.
Climate, on the other hand, is the average of these daily conditions over time. At its most basic and familiar level, climate can be described based upon seasons. Some places on Earth experience four of them (winter through autumn), others experience two (rainy and dry). These short term climatic conditions are based upon a large array of variables. These include latitude, topography, proximity to the sea, time of year, and so forth. Looking at changes in climate over the course of the year at a given location on the planet provides a short-term picture unique to that location. Climate can also be described on a global average basis. Under these conditions, we can follow long-term changes in temperature and associated variables (greenhouse gases, ice loss, evapotranspiration, etc.) that provide a broader picture of global climate and how it is changing.
All of this is still be based upon instrumental records on the Earth’s surface as well as satellite imagery and data from space. But what about the years before humans could measure these things? Is there any way of figuring out climatic conditions in the time before scientific measurement or human existence?
Paleoclimatology: Historical Climate
Paleoclimatology is the study of the Earth’s climate before instrumental records were available. The demarcation between paleoclimatology and climatology is analogous to paleobiology and biology. Both fields study climate, but one is very much focused on our current climatic situation. Paleoclimatologists must tease out of the Earth the story of its climatic past. And, that past is full of evidence of change.
While we tend to focus on measuring our weather and climate using instrumentation on the ground and in space, the Earth has kept a record of its past climate in a wide variety of ways. In this chapter, we will explore a wide array of these kinds of paleoclimate evidence, including things like tree rings, layers of ice, and pollen. There are many more examples. Some of these record great detail over shorter periods of time while some record more coarse detail over much greater spans of time. In reality, this means that the Earth’s data stores are really no different than the data stores we create in instrumentation. Measurement of climate is a matter of time and resolution. The usefulness of a particular data set is limited by time and resolution. So, like studies of modern climate, paleoclimatologists ask questions and pursue investigations that are guided by data limited by time and resolution.
There are some important assumptions accepted when studying paleoclimate. Like in all of the geosciences, these assumptions are uniformitarian. That is, we assume past processes generally operated the same way as in the present. That does not mean changes in the Earth’s systems that affect climate always happen in the same way, but just that the processes that force or drive the changes do.
A key assumption is that the Earth’s systems behaved throughout its history as they do today. And, we accept that the interactions among the Earth’s spheres also have remained consistent. An event that forces movement within one system will have effects in other systems. These can lead to amplifying feedbacks that further upset the system’s dynamic equilibrium and can lead to tipping points and a new set of system circumstances. Balancing feedbacks also still serve to work against such processes, as a natural means for the systems to right themselves and return to the former equilibrium.
The study of paleoclimate has lead to a wide variety of insights into our modern climate and how it works. It has helped to refine the models we use to study current problems associated with anthropogenic (human caused) warming. Studies of paleoclimate have also led to deeper knowledge of how our solar system works. While Galileo and others discovered sunspots, their periodicity and effect on climate at different scales and magnitudes would not be known if it were not for paleoclimatic research. Likewise, we would perhaps not know that the Earth, like other planets, sees gradual changes in its orbital parameters that affect its climate. Knowledge about these exospheric impacts informs our current research, not only because it helps us understand how our climate should be responding under only natural influences, but also because we can eliminate them as mechanisms behind the current changes we are causing. Paleoclimate research is critically important to understanding our future.
Intrinsic and Extrinsic Forcing Mechanisms
The Earth system is a part of the solar system. Extrinsic climate change agents must come from outside of the Earth system and result from the solar environment, or the exosphere. The solar system is powered by mostly by our Sun and also contains a variety of other objects, including planets, dwarf planets, and chunks of rock or ice hurtling through space in the form of asteroids or comets.Extrinsic forcing agents typically have a global effect, but in some cases effects can be more limited, such as the impact of a Tunguska-like asteroid versus a Chicxulub-like asteroid. The Tunguska asteroid impact, which happened over Siberia in 1908, had a wide variety of local effects, whereas the Chicxulub asteroid impact at 65 Ma devastated the planet’s biosphere, killing off the dinosaurs and eventually giving mammals the evolutionary reins. The devastation of this impact led to massive global climate effects that contributed mightily to the extinction.
Likewise, intrinsic climate change agents come from within the Earth system. Overall, these forcing mechanisms can tend to be more limited in their impact, occurring on more local or regional scales rather than global. The Earth system consists of the atmosphere, hydrosphere, geosphere, cryosphere, and biosphere. Massive perturbation in any one of these could have climatic effects. From the geosphere, volcanic eruptions provide an excellent example of scale-dependent intrinsic climate forcing. In April 1815, Mt. Tambora, a stratovolcano in Indonesia, erupted. This eruption was very large and led to three years of significant climate change, particularly for the northern Hemisphere where Europe recorded the three coldest years in centuries and saw major crop failure and famine.This calamity pales in comparison, however, to the major flood basalt eruptions of Earth’s past. The Siberian Trap eruptions that took place at the Permian/Triassic boundary caused the largest mass extinction in the planet’s history. This large igneous province in northern Siberia covered 7 million km2 in lava and lasted for perhaps hundreds of thousands of years. CO2 emissions from the lava led to runaway global warming and oceanic stagnation. Almost everything on Earth died.
Extrinsic Change Agents
There are some key extrinsic change agents that affect Earth’s climate from our exosphere. Below is a brief explanation of some of them. Ultimately, extrinsic factors do one of two things. They either adjust the amount of incoming solar radiation, like turning a handle on a spigot, or they cause direct perturbations to the Earth system. Most notably, changes in the hydrologic cycle and carbon cycle are good examples of systems that can be heavily affected by bolide impacts, for example. For this chapter, we will focus on three specific extrinsic change agents. These are Milankovitch Cycles, bolide impacts, and changes in solar energy output.
Born in the rural Serbian village of Dalj in 1879, Milutin Milankovitch would make his mark on the study of climate in a big way. Through his research on insolation variation during Earth’s seasons, he discovered a mathematical theory of climate that helped predict changes in Earth’s climate due to changes in its orbit . His theory stated that as the Earth travels in its orbit around the sun, three cycles of variability will have an effect on changes in its climate. These are orbital eccentricity, axial obliquity, and axial precession.
The longest of these cycles is eccentricity, lasting from 90,000 to 100,000 years per cycle. As this cycle progresses, the Earth’s orbit stretches from more circular to more elliptical. The shift occurs because, while the Earth orbits the Sun, the giant planets Jupiter and Saturn over time exert gravitational forces on the Earth, which causes the Earth to shift over this period. As Johannes Kepler discovered well before Milankovitch, no planetary orbits are perfectly circular; all are elliptical (Kepler’s First Law).
However, Milankovitch discovered that the shape of our orbit changes enough to have a significant effect on our climate. Currently, our orbit is in the phase of this cycle where it is more circular, leading to only a 6% increase in incoming solar radiation from January to July. In January, the Earth is at its closest approach to the Sun for the year (called perihelion) and in July, it is at its furthest approach (aphelion). When the orbit is more highly elliptical, the difference in incoming solar radiation at perihelion can be 20% to 30% greater than it is today. This would mean a much different climate than we have today. And, as you can see in the figure below, it is this 100,000-year cycle that governs cold and hot periods (glacials and interglacials) during the Pleistocene epoch.
While the Earth’s orbital path is changing, there are also at least two cycles of obliquity that will occur at about 41,000 years each. During one of these cycles, the Earth’s tilt ranges from a minimum of 22.1° to a maximum of 24.5°. Currently, our orbital tilt is 23.5°, meaning that we are in the middle of one of these cycles rather than at one end or the other. 10,700 years ago, the planet was last at its maximum tilt. In 9,800 years, it will reach its minimum tilt. As the tilt of the Earth is the main factor that determines the behavior of our seasons, changes in this variable can also have significant effects on global climate. Greater tilt leads to more severe seasons while lesser tilt leads to milder changes between seasons. We think the greatest impact on climate occurs at times of greater tilt, as ice builds up and causes major changes in insolation albedo, leading to an amplifying feedback toward cooler climate. Higher reflectivity of the Earth’s surface reflects more of the sun’s radiation back to space. When this happens, ice sheets expand. As ice sheets continue to expand, even more sunlight is reflected back to space, cooling the planet.
Finally, precession is the shortest of the Milankovitch cycles. Over the course of about 26,000 years, the Earth’s axis (the imaginary line defining our geographic north and south poles) wobbles. As the axis does wobbles, it marks out a kind of circle on the celestial sphere over the course of that time. Currently, the axis is pointing toward the star Polaris, the North Star, making it the only star that stays in a “constant” position in the night sky. This is why it can be used for navigation. This has not always been the case over longer periods, however. At times, there is no north star, much like there is currently no “south star.” At other times, other stars have that title, as the star Thuban did in 2000 BC when ancient Egyptians might used it to navigate through the desert. In terms of climate, the changes in precession alter the dates for aphelion and perihelion during the year, which affects how seasonal insolation affects the northern and southern hemisphere. As such, the effect of precession in global climate is significant, but more difficult to tease out of data as it is not nearly as powerful as obliquity and eccentricity.
Asteroid impacts are both exciting and terrifying. We hope that, should we ever be in danger of a massive bolide collision, NASA or some other space agency will see it coming. If they do, perhaps we will get a chance to do something about it. Until then – such imaginings make for excellent disaster B-movie plots!
However, these are very real hazards. Certainly, small impacts like the Tunguska event mentioned earlier are not likely to be seen by any NEA (Near Earth Asteroid) observatory, like the Center for NEO Studies, run by NASA through the Jet Propulsion Laboratory at Caltech. According to their data, there have been 822 recorded fireballs between April 15, 1988 and March 4th, 2020:
Most of these are small and cause only localized disturbances. To better quantify risks associated with impacts, the IAU (International Astronomical Union) in 1999 created the Torino Scale. This metric communicates to the public the hazard level of a particular NEA that is being watched by specialists. Displayed using colors from white (no hazard) to red (certain collisions), it is simple and conveys important information on the type of resulting hazards associated at that level. Fortunately for Earth, none of the currently monitored NEA objects are listed as greater than 0 on the Torino Scale.
Asteroid impacts are important extrinsic climate change forcing mechanisms because of the effect they convey on the atmosphere, hydrosphere, geosphere, and biosphere. Impactors heat the atmosphere on their way in and then, as their ejecta return to Earth, heat the global atmosphere more evenly, possibly to temperatures as high as several thousand degrees. But this is a short-term effect of the impact. Post-impact, dust, ash, and other materials may block incoming solar radiation for many years to come, creating a long-term cooling effect. These are just two of the climate effects. As the oceans acidify from dissolved gases and rock debris, climate-regulating currents may be upended. New tectonic or volcanic activity from the geosphere can also contribute to climatic change.
In the time periods where we study paleoclimate, none of these warning agencies existed. In the Earth’s history, we know of many large impacts that have occurred, many of which have left lasting scars on the surface. One of the most photographed is of Manicouagan Crater in northern Quebec, which was created by a bolide that hit that part of Canada 215.5 Ma (during the Triassic period). The impactor is estimated to have been about 5 km in diameter (85km diameter crater), very close to the size of the Chicxulub impactor that caused the extinction of the dinosaurs at the K/Pg boundary (6km impactor and 150km crater).
The Earth Impact Database world map. Note that the colors indicate a rough time period of impact. These are also known impacts. There are very likely many, many more yet to be discovered. Hovering over a circle brings up more information about it.
Changes in solar insolation over time
Over the course of a year, the fraction of the energy Earth receives from the Sun is measured as insolation (the energy received), and varies by about 0.1%. While this fraction is small, it actually represents quite a bit of energy. Changes in short-term solar insolation are of great interest to climate modelers. Such changes result over time spans longer than a year due to the 11-year sunspot cycle, where the Sun’s surface experiences a cycle running from cooler to hotter due to magnetic changes. These can affect the planet’s climate over these time spans, though the evidence is clear that current climate change cannot be explained by this. To study changes in solar output, NASA in 2003 launched the Solar Radiation and Climate Experiment mission (SORCE).
The Sun itself changes over very long time spans also. Since its early days, it has been steadily warming and will continue to do so. The warming trend is significant, but does pale in comparison to the increases over time in the star’s radius and luminosity. These two variables are related, as greater radius leads to greater surface area which leads to greater luminosity. In the figure to the left, it is predicted that the Sun will continue to grow larger. One misleading piece of data here is temperature. While the Sun’s temperature may not increase as dramatically as its radius, as that radius increases, the surface of the Sun and its very hot atmosphere will inch ever closer to Earth. This effect on Earth will be very significant, very likely making within a billion years or less most life on Earth as we know it no longer possible. Ultimately, whether at this point the Earth is pushed outward be the Sun’s lower mass or pulled inward by drag and enveloped is a matter of some lively scientific debate.
Lunar effects (stabilization, etc.)
What about forcing factors outside of the Sun and bolide impacts? Certainly, while there may be some climatic effects caused by phenomena such as supernova explosions and the passing of our solar system through dense interstellar clouds, these effects are tiny, if they exist.
What about the Moon?
The Earth’s natural satellite (as that is really the proper term to use), the Moon, has two demonstrable effects, though small, on our climate. Both of these effects are related to the tidal gravitational forces that the Moon exerts on the Earth. When we think of tidal forces, we usually think of shorelines. This is correct, of course, but the effect on the shorelines is really a gravitational effect, as the Moon is pulling and releasing the water in the ocean as the Earth rotates and as a 29.5 day lunar cycle passes. As you may have learned, there are one to two diurnal high and low tides per day. During a lunar month, there are also times of greatest tidal range that occur during the New and Full phases of the Moon, the spring tides. The tidal range is the difference between the elevation of high and low tide. There are also what are called neap tides, when the tidal range is lowest during a lunar cycle, during the 1st and 3rd quarter phases (half moons).
These tidal effects actually affect climate on a small scale, by affecting the amount of precipitation that occurs. It does this not by pulling on Earth’s water, however, but by exerting its gravitational pull on the atmosphere. When the Moon is overhead, a full moon in particular, air pressure is actually higher. This slightly depresses the chance for precipitation. The effect is tiny, but it does exist.
Perhaps a larger lunar impact is related to the Earth’s obliquity cycle. Joutel and Robutel (1993) were able to demonstrate the Moon’s gravity stabilizes our planet’s tilt. In essence, it keeps our obliquity cycle stable, which prevents much more massive long term climatic changes. This also helps stabilize our seasons, which the Earth’s tilt is critical in defining. In times when the tilt angle could be nearly 0° or much greater than 24.5° maximum our planet currently has, the seasons would be very different. In a situation where there is zero to little tilt angle, there would be no seasonality on the planet, or very little. In a situation where the tilt might reach greater than 30°, the difference between summer and winter would be quite harsh in temperate latitudes and nearly impossible to bear in polar regions. As such, our Moon works to stabilize our climate.
The leading idea on the formation of the Moon is that it resulted from a massive impact, probably of an object about the size of Mars. In its early years then, the Moon was much, much closer to the Earth. Over time, it has gradually been moving farther and farther away. However, this also means that its gravitational tidal forces have also been lessening over time. One notable effect of this is the gradual change in the length of a day on Earth and, consequently, the number of days in a solar year. This is born out by fossil mollusks and corals from 80 Ma (and other time periods) that show that a year was 372 days at the time.
There are also potentially effects over time on other orbital parameters that affect climate. The affect of our natural satellite on our climate is still an active and exciting area of research.
Intrinsic Change Agents
Within the Earth system, there are a wide variety of factors that can affect the Earth’s climate. These intrinsic factors all have one thing in common. In one way or another, they disrupt, or place in a state of disequilibrium, one or more biogeochemical cycles. There are many biogeochemical cycles and all of them interact with more than one of the Earth’s subsystems, or spheres. Examples include the carbon cycle, nitrogen cycle, hydrologic cycle, phosphorus cycle, etc.
The Earth’s climate is regulated by the greenhouse effect, or the retention of re-radiated longwave radiation through the absorption of this energy by greenhouse gases in our atmosphere. Earth is not the only planet that has a greenhouse effect. Venus has this also, in spades. It is what leads to surface temperatures on Venus exceeding 400ºC (~800ºF) in some instances. Mars has very little atmosphere and though what it does have it almost entirely carbon dioxide, it cannot retain much of the Sun’s energy. Mars lacks a magnetic field and thus cannot hold onto its atmosphere. One other location in our solar system, Titan, a very large and gassy natural satellite of Saturn, has a moderate greenhouse effect, but due to the high concentration of methane in its atmosphere.
On Earth, the greenhouse effect is moderate, but absolutely critical for life. This is due in part to its distance from the Sun, but also because of its very different mix of atmospheric gases due to the presence of life. Photosynthesis has altered our atmosphere so that it contains 21% oxygen on average. 80% of the gas in our atmosphere is nitrogen, which also has biogenic sources. Life on Earth has, thus, reduced the carbon dioxide so prevalent in the atmospheres of Venus and Mars (from volcanism) to a mere 0.04% of the total gas content. As of this writing today, this amounts to a carbon dioxide concentration of about 416ppm. Common greenhouse gases include carbon dioxide, methane (another carbon compound), oxides of nitrogen (NOx), HCFCs (anthropogenic chlorofluorocarbons), and water vapor.
Ultimately, all of the elements that make up compounds in our atmosphere share a primordial source from the Earth’s interior. Biogeochemical cycling, aided by plate tectonics, makes these elements accessible. Biogeochemical cycles are cycles driven not only by geological processes (geo), but also biological processes (bio). As such, Earth’s versions of these cycles are unique.
Geological Sources of Important Elements for Life
|Elemental Nutrient||Primordial Source on Earth||Important Life Processes|
|Carbon||Cosmogenetic, Mantle, Volcanism||Basic element of organic chemistry, important nutrient for photosynthesis of glucose|
|Hydrogen||Cosmogenetic, Mantle, Volcanism||Basic element of organic chemistry, important nutrient for photosynthesis of glucose|
|Nitrogen||Cosmogenetic ammonia, Earth’s atmosphere||Important for the creation of amino acids, enzymes|
|Oxygen||Cosmogenetic, Volcanism, Carbon Dioxide dissociation, Photodissociation in Atmosphere of Water||By-product of photosynthesis, Necessary for respiration and associated oxidation|
|Phosphorus||Phosphides in Earth’s core, Inorganic minerals such as apatite and fluorite, ocean-floor sediments||“Spine” of DNA molecule, important micronutrient that helps store energy in cells via ATP|
|Sulfur||Cosmogenetic, Stored in Earth’s interior, Volcanism||Allows for the synthesis of a greater variety of amino acids, Important nutrient for chemosynthesis|
Changes in greenhouse gas concentrations, particularly carbon dioxide, have led to changes in the Earth’s climate today, but also in the geologic past. Today, we monitor carbon dioxide levels at NOAA’s Global Monitoring Laboratory atop Mauna Loa, an extinct volcano in Hawaii. You can view carbon dioxide concentrations as measured instrumentally and then compared to data taken from ice cores from Antarctica (a source of proxy data for carbon dioxide) in the video below (Retrieved May, 2020):
Milkankovitch Forcing, Ice Cores, and Carbon Dioxide
Similarly, methane, another greenhouse gas, has seen a rapid increase in atmospheric concentration due to human activities since monitoring began. A nearly identical trend to carbon dioxide and methane can be seen in nitrous oxide gases. While nitrous oxide gases are also a major component of the nitrogen cycle, they are intimately intertwined within the carbon cycle because so many of the sources of carbon and nitrous oxides are the same processes. These include biomass burning, fossil fuel combustion, etc.
Carbon cycle gases are the real keys to understanding our changes in climate today. Human activities from the burning of fossil fuels to the production of fossil fuels, to deforestation, and other perturbations to the carbon cycle that lead to changes in carbon dioxide and methane concentrations are having a profound effect on our modern climate. The image below depicts natural and anthropogenic fluxes in the carbon cycle and their magnitude for a ten year period from 2000-2009. Note the magnitude of anthropogenic fluxes (red arrows and numbers).
Knowing the source of anthropogenic greenhouse gas emissions is one thing, pinning the increase and resultant climate effects is another. Apart from our ability to use carbon isotopic ratios to identify greenhouse gases as the course of new atmospheric carbon dioxide, stations monitoring atmospheric oxygen levels are seeing decreases across the globe as atmospheric oxygen is used in the combustion process of fossil fuels. Cape Grim, Tasmania, provides an excellent example.
The challenge of maintaining balance within the current carbon cycle and resulting climate is really no different today than in the past. The main difference today is the speed of changes. In Earth’s past, changes in climate and the carbon cycle were more often much slower. To illustrate this, it is useful to think of the carbon cycle as having a slow component and a fast component, as illustrated below.
Today, human activity is pulling carbon out of reservoirs (storage) that have taken thousands to millions of years to accumulate (slow carbon cycle) and then adding these stored carbon to the fast carbon cycle. The speed at which the magnitude of these anthropogenic fluxes are changing is leading to rapid change in the Earth system, leading to a state of disequilibrium. At some point, tipping points will be reached. Such tipping points include releases of methane from Arctic permafrost, ocean acidification with the potential to trigger a marine mass extinction, etc.
Similar changes to the carbon cycle, or other biogeochemical cycles, also occurred in the past. Below, a few forcing mechanisms will be discussed that have led to changes in these systems, leading to large climate change, mass extinction, and other major geologic events.
Flood basalt eruptions
At various times throughout Earth’s past, there have been major eruptions that have lasted for extended periods of time and that have affected large geographic areas. Called Large Igneous Provinces, or LIPs, these massive volcanic eruptions had a very unique character, given the massive volumes of erupted lava, the very extensive nature of them,
and their long duration. All of them share similar environmental effects also, most of which are related to climate. Some LIP events exhibited greater effects in one area than others. These include 1) global warming, 2) oceanic anoxia, 3) release of methane from gas hydrates, 4) oceanic calcification crises, 5) a period of global cooling, and 6) a significant extinction event. LIP eruptions may result from the pooling of upwelling plumes of warm rock from deep within the mantle, likely from the core/mantle boundary, sourced from large low shear velocity provinces (LLSPVs) that are probably partially molten.
The extinction connection is very significant. Within the last 258 million years, there have been five major LIP eruptive events. These include the Emeishan LIP (258 Ma) that correlates to the end-Guadalupian extinction within the Permian Period, the Siberian Traps (250 Ma) that likely caused the end-Permian “Great Dying” extinction, the Central Atlantic Magmatic Province (CAMP, 200 Ma) that likely caused the Triassic mass extinction, the Karoo-Ferrar Traps (180 Ma) that caused what is called either the Toarcian Turnover or the Toarcian Extinction during the Jurassic Period, and the Deccan Traps (65 Ma) that coincide with the extinction of the dinosaurs and the end of the Mesozoic Era.
Such extinctions likely begin with the massive releases of carbon dioxide from the eruptions. These releases warm the atmosphere through the greenhouse effect, as is being done today by humans (though at a faster rate). Following this, the ocean waters warm also. Warmer waters hold lower concentrations of dissolved oxygen and, because the warmer air disrupts the ocean circulation system (at times disrupted also because of the positions of the continents), the water is able to hold very little oxygen. This is the source of oceanic anoxia.
Also like today, the increased carbon dioxide in the atmosphere translates to increased carbon dioxide in the oceans, leading to the “calcification crises” due to ocean acidification. Animals that secrete calcite are competing with dissolved carbon dioxide for the available bicarbonate. Warming waters also releases accumulations of methane gas hydrates, which are common in offshore environments. This added methane acts as an amplifying feedback to warming, ramping up temperatures even more, thus amplifying also the warming of ocean waters and so on.
Like other terrestrial planets such as Venus and Mars, carbon dioxide was a very prevalent gas in our planet’s early atmosphere. Over time, this changed. Earth’s earliest atmosphere was composed of gases such as helium that did not stay put. Rather, Earth’s mass is not great enough to retain helium and so it floated off into space, as it does today when you release it from a balloon. The planet’s magnetic field had also not developed yet, so there was no way to prevent the solar wind from blowing away whatever atmosphere existed. The fact that Earth’s liquid outer core is so substantial (and fluid enough to convect, generating a magnetic geodynamo) is surely one of the planet’s “tricks for success” that have led to its stability over the geological long term.
Eventually, the formation of a magnetic field and lithosphere led to an atmosphere composed of accumulated volcanic gases, notably carbon dioxide. While it is hard to imagine what the climate of this early Earth was like, it was likely very hot. Life began, and started to evolve. It is likely that the first microbial organisms got their energy chemosynthetically, through the breakdown of compounds like hydrogen sulfide, which is common at deep sea volcanic vents. Genetic evidence also suggests that both Bacteria and Archaea had their origins in a very hot environment: the least derived members of both groups are hyperthermophilic extremophiles. In our solar system, the closest we get to examples of such an atmosphere are on Venus and Mars.
Earth’s final atmosphere, the one we have today, began to form around 3.8 Ga with the innovation and evolution of photosynthesis. Most free oxygen produced through this process was absorbed and retained by the oceans, which kept atmospheric levels low. However, once saturated, oxygen was able to effervesce (exsolve) into the atmosphere and accumulate there. The presence of this oxygen would eventually lead to massive cooling events, as the prior carbon dioxide-rich equilibrium was upset. Several “Snowball Earth” events would follow throughout the remaining Precambrian time.
This last atmosphere is unique to Earth. No other planet in our solar system has an atmosphere like our planet. The biosphere’s ability to alter its environment led to not just local changes, but global change. As evidenced by climate change today, we know this is still possible, even by a single species. Massive releases of a photosynthetic waste product, oxygen, would force major changes in the hydrosphere, followed by the atmosphere. Eventually, these would also be reflected in the geosphere as banded iron formations formed in the oceans and later by massive iron-rich red beds in continental environments. The iron so prevalent in the early Earth’s geology was oxidized (chemically weathered) by all of this free oxygen.
As mountains rise, they are also weathered and eroded. At different periods in Earth’s history, there was more orogenic activity (mountain-building) going on than at other times. Orogenies occur at plate boundaries, the largest of which occur at convergent plate boundaries. As two continents plow into one another, as exemplified by India converging with Asia about 45 Ma, they form large mountain chains along massive collisional belts. In the case of the Himalayas, the largest mountains above sea level exist there rise to above 8,800m (~29,000 ft) of elevation, nearly to the stratosphere.
As mountains rise up, they are subject to more and more chemical weathering. This is because the troposphere’s carbon dioxide combines with the clouds and precipitation occurring as a part of the hydrologic cycle, forming a weak carbonic acid. This acid over time reacts with the feldspar minerals so common in igneous and metamorphic rocks that form the cores of such mountain ranges, allowing them to be weathered downward. Mundane as it seems, this process not only leads to salty oceans, produces a great deal of free bicarbonate and calcite for shell-making organisms, but also pulls carbon dioxide out of the atmosphere, leading to a cooler climate. As you can see in the image above, there is a chemical equation that describes this process:
2(CO2) + 3(H2O) + CaSiO3 → Ca2+ + 2(HCO3-) + H4SiO4
Thus, the rise of mountains leads to global cooling. It also leads to the expansion of carbonate environments.
An excellent example of the power of silicate weathering lies in its role in moderating the climate system. If tectonics tunes the climate over millions of years, weathering of silicates is a key process that serves to reduce or amplify warming or cooling, depending upon the initial inputs. In the image of feedback loops to the left, times of fast seafloor spreading (that also drive subduction and orogeny) lead to carbon dioxide input from volcanism which warms the climate. As resulting mountains intrude into the atmosphere, chemical weathering increases, carbon dioxide is removed, and warming is reduced. Likewise, when seafloor spreading slows, the reverse series of processes occurs, eventually moderates the resulting cooling trend.
By the end of the Pliocene Epoch, the initial conditions were very much like the top loop. There was a great deal of global tectonic activity and seafloor spreading that had been going on for nearly 200 Ma since Pangaea broke up. This led to accumulations of large concentrations of carbon dioxide in the atmosphere and, with the exception of some smaller orogenic events such as those in the American west (Sevier Orogeny, Antler Orogeny), there were no really major orogenic uplifts until about 45 Ma, when India begin to slam into Asia, as mentioned above. Prior to this, the Earth reached a maximum warm period, referred to as the Paleocene-Eocene Thermal Maximum (PETM).
Measuring Paleoclimatic Change: Proxy Records
By definition, paleoclimate data had no one there to collect it. So, how do we know what we know? As it turns out, the Earth has its own methods for recording its temperature (among other variables). As we get better as speaking the language of Rock, we learn to tease out of the stratigraphic record evidence of changes in temperature. Not all of these changes are recorded in rock, however. Some of them are recorded in ice. Others are recorded in wood. Still others are recorded in ocean or lake sediments. In many cases then, it’s not rock, but sediment or other carbon-rich materials where we seek data. Collectively, we call these “proxy records”. This simply means that they are records that stand in the place of the instrumentation we usually use to measure things like temperature.
Temporal Scale and Proxy Records
Before we begin our brief overview of a variety of paleoclimate proxies, it is important to note that not all proxy records are created the same. Some record climate records over just the last few hundred to thousand years while others go back millions of years.
In the figure to the right, we can see the time limitations of historical and instrumental records. Beyond these time horizons, about 3-400 years, we refer to paleoclimate. Proxy records like tree rings can take our understanding of temperatures back to as much as 10,000 years ago and ice cores close to 800,000 years before present. Coral reef data extends to about the same time horizon as ice. From there, sediment records can give us climatic data going back as much as several tens of millions of years. At this point, the rock record and all of the data it contains become critical. For fossil shellfish, foraminifera, and other carbonate creatures, oxygen isotope records are very useful. Carbon isotope ratios extracted from black shales hundreds of millions of years old grant us insights into climate. There are numerous examples of proxy records.
You can find, explore, download, and freely use paleoclimate proxy data of great variety by using the Paleoclimatology Data Map maintained by the National Centers for Environmental Information.
Let’s explore some key examples of proxy data sets commonly used by paleoclimatologists today.
Biological Proxy Records
Plant Fossils (Moderate to Long Timescales)
Plant fossils provide proxy records that vary in timeframes. Typically, macrofossils of plants, or remains large enough to be visible without a microscope, are the focus of such work. Using our knowledge of modern flora as they related to climate, including factors such as needle length, leaf shape, and more, we can use these macrofossils to describe ancient terrestrial environments. Some of these plant fossils are found in rock, other times they are found as a part of packrat middens. Fossil plants have more to say about really ancient environments while packrat middens might only describe environments of a few hundred to thousand years ago, and even then they are only applicable in arid regions.
When you think of plant-free harsh environments on Earth today, Antarctica might come to mind. But, a wealth of plant and animal fossils have been found there, dating back hundreds of millions of years to the Permian Period. At that time, Antarctica was a part of a very large land mass geologists call Gondwana. It was humid, what we might describe today as tropical. Antarctic land was carpeted with forests containing a wide diversity of plants. One of the well known fossil plant examples from this region is Glossopteris, woody plants with tongue-shaped leaves arranged in thick mats. They may have even been deciduous.
Plants with leaves can be used to provide quantitative measurements of climate. Using uniformitarian principles, researchers evaluating modern environments have identified a linear relationship between the percentage of leaves in a location that have a smooth, “Entire” (toothless, non-lobate) margin and temperature. This relationship in modern species can be applied to assemblages of fossil leaves from the same time period to help determine what the temperature was like in that area.
Leaf margins were used successfully in this way for determining what the Eocene climate was like in the Geodetic Hills of Axel-Heiberg island in northern Canada.
Pollen (Short to Long Timescales)
The record of pollen is directly linked to the existence of seeds. The earliest known plants that produced seeds were seed ferns. The oldest fossil seed ferns hearken to the late Devonian Period. These ranged from the size of small trees to shrubs. Since that time, over the subsequent 300 million years, the record of pollen has been accumulating. Pollen is by nature airborne. Whether transmitted through the air by insects or other organisms or blown by the wind, it can be carried great distances and, as such, was a fabulous evolutionary innovation. This adaptation persists today in all seed-bearing plants. You plant them in your garden, walk through them in the forest, sneeze because of them, and pick them for a lover or a friend.
Pollen provides a record of the vegetation of a region. Much of the record we recover and use as proxies for paleoclimate comes from lake sediments. Drifting airborne pollen becomes stuck when it touches water and eventually settles to the bed of lakes. From the emerging vegetation profile, painstakingly amassed via sometimes tedious work at microscopes, a detailed view of the climate of a region at the time and place may emerge. Pollen records run from recent times back several hundred million years. So, they are useful paleoclimate proxies over their entire time span.
Fire History (Short Timescales)
Fire history data tends to record more recent paleoclimatic data. It is also more sporadic, recording not continuously but in seasonal moments of intense burning. The data is also multiproxy, meaning that it comes in more than one form. Some fire data may come from tree rings. Other fire history data may come from lake sediments. There is even a global charcoal database!
Fire histories give us insights into ancient climate and how it has changed in a region over time. They also contribute to modern wildfire knowledge and contribute to the creation of forest management models. These system models use the history of fires over time as one of many variables to help understand feedbacks and variables related to everything as diverse as drought frequency to the volume of tinder on the forest floor.
Faunal Data (Short to Moderate Timescales)
Faunal data is very similar to plant data, in that the types of faunal remains found in a location provide excellent insights into the climate of the area at that time. Faunal remains come in various forms. There is fossil faunal data, of course. This can include everything from marine shellfish to terrestrial mastodon remains.
One significant project that has relied heavily on faunal data has been the characterization of the changes in the Sahara Desert region of north Africa during the Holocene Epoch. Analysis of the organisms from this vast region has produced evidence of past humid environments, rather than today’s arid landscape. Animals of various types used waterways that existed then as migration routes. Some of these animals were aquatic. These waterways consisted of linked lakes, rivers, and inland deltas. These environments existed during the last interglacial of the Pleistocene Epoch and also during the early portion of the Holocene.
The faunal evidence also includes human data. The migrations of humans were influenced by these waterways.
Physical Proxy Records
Sediments – Oceanic and Lake (Short to Long Timescales)
Sediments accumulate in basins after being eroded from highlands. As this occurs, they bring with them loads of evidence of their former lives and their journey to this new destination. Before becoming rock, these sediments contain a wealth of climatic clues, whether entombed in oceanic sediments or lacustrine.
Oceanic sediments are analyzed not only for their physical attributes, but also for chemical variables. These include isotopic data and trace metal analyses. They also contain a wealth of microfossil information. The ocean sediments themselves can give information on ocean health just by their color. Darker colored sediments may contain few shells and could provide evidence for a period of little to no oxygen (dysoxic to anoxic) . Light-colored sediments likely contain lots of shells and suggest a more healthy, oxygen-rich ocean. Changes in ocean oxygenation are clues for past acidification events that are associated with periods of climate change, often associated with atmospheric changes such as a rise in carbon dioxide.
Lake sediments are used for similar studies, but also to analyze moisture profiles in a region. Because freshwater lakes do have so much depth variation that is directly related to seasonal and other hydrological changes, lake level is a useful measure of moisture in a region. As such, it becomes possible to describe the climate based upon this moisture. Is it humid or arid?
A great tool for exploring this is NOAA’s “Lake Level Viewer” for the North American Great Lakes. Here, you can explore a wide variety of variables related to these lakes. In particular, predicted changes in shoreline can be explored using simple tools.
Apart from lake sediments, it is possible to find imagery of changes in lakes over time that also provide insights. While not specifically paleoclimatic, such imagery is very useful for documenting climatic change. Lake Mead, created by the Hoover Dam in southern Nevada, is a classic example of a lake experiencing arid conditions and, over time and through overuse, is drying out, leaving a very visible bathtub ring record.
Tree Rings (Short Timescales)
You were likely exposed to tree rings in your younger years. As children, exposure to tree rings is an introduction to a hidden world, one where curious information abounds at your fingertips. It is perhaps one of the most tangible and familiar proxy records. Major sources of tree ring data include the Southwest Paleoclimate CLIMAS portal, the Colorado River Basin Tree Ring Analysis, and the Living Blended Drought Prediction (LBDP).
Rings record a good deal of information. Some rings are thin, some wide. All rings vary in thickness around the center of the tree. Why? Most rings have two different types of wood – so called “early wood” and “late wood.” This couplet of two materials make up an annual growth ring. The early wood is produced in spring and feeds off of the starches stored in last summer’s late wood. The thickness of early wood is a function of the quality of the late
wood from the year before, put on during the summer. Late wood is produced from the sugars created through photosynthesis in the leaves and then stored for the cycle to begin anew the next year. Late wood is also structurally stronger and more dense than early wood. Overall, most trees produce all of their new growth during about an eight week window from spring to early summer, until the hot weather arrives.
The thickness of a ring is a function of a variety of variables. Two of these are less important to the dendroclimatologist, the name of a specialist who studies climate through tree rings. Rings reflect the growth of the tree and its structure. A single ring will vary in thickness as you follow it up the tree and the tree narrows. It will also have a slightly variable structure depending upon how the tree grows–does it lean, split, or otherwise have diseased growths, etc. The one variable that a climatologist is interested in, moisture, has to be teased out using sophisticated statistical programs on a computer.
Tree rings are most often extracted from living trees or structural beams from buildings. This is done using an increment borer, which allows a small straw-like sample to be extracted from its core. This is mounted, polished, and the rings are meticulously measured along the length of the core. This data is input to a program, and compared to a database of other trees. All of this is meant to remove growth variables, leaving behind a record of moisture. From this, it is also possible to extract calendar dates for growth, particularly from buildings. This data is critical in understanding modern and paleodroughts. As a proxy record, tree rings only take researchers back several thousand years. This is enough for Quaternary geoscientists, who study the most recent geologic period, and archaeologists, who are interested in the material records of humanity, to learn how changes in climate have affected environments and human cultures. Today, governments use this and other data to monitor drought conditions. Other data are also useful, such as wood chemistry and isotopic fractionation.
Loess and Eolian Dust
Dust is something you breathe in daily. On the U.S. Gulf coast and in the Caribbean regions, some of this dust is blown in all the way from the Sahara desert. Over 100 tons of dust from the Sahara is blown westward annually. When silty sediment is windblown, it is called loess.
Dust is being blown all over the place, as long as the particles are small enough to be blown and wind energy great enough to keep the particles suspended. As such, ice cores are one of the most important repositories for the dust record, as much of it makes it over polar regions to find a resting place on ice caps.
What does it mean? Dust is sourced from dry land. Wet sediments do not blow away as easily as dry. So, from a paleoclimate perspective, large influxes of dust are an indication that some regions were experiencing periods of intense drought. This dust can be chemically analyzed too. Carbon dates, percentage of organic matter, and other measures can provide a wider variety of information about the dust source and depositional location.
One location where loess data has been very useful in understanding its paleoclimate is the Chinese Loess Plateau. Maher and Hu (2006) used magnetic susceptibility data and grain size measurements, for example, to create a detailed record of the southeast Asian monsoon during the
Holocene Epoch. Noteable periods of increased aridity, and dust, occurred at 12,500 BP (Before Present) and 11,500 BP, a time period also known as the “Younger Dryas”, a time of sudden cooling that occurred after the start of the Holocene. Another noteworthy period of aridity began to occur after 5,000 BP, very close to what is now referred to as the start of the Meghalayan Age of the Holocene Epoch. These were likely periods of time when major shifts in the region’s monsoons occurred.
Chemical Proxy Records
Chemical proxy records come primarily as isotopic data. You may recall from an early chemistry class that some elements exist in different forms. That is, there is a stable form and then there are versions of that element that have different numbers of neutrons, while retaining the same atomic number. Different isotopes of an element are named for their atomic mass, the total number protons and neutrons in their nucleus (for example, Oxygen-16 and Oxygen-18). Generally, in situations like oxygen and carbon, the stable versions of these elements, Carbon-12 and Oxygen-16, are found nearly everywhere. However, they do not make up 100% of the carbon or oxygen in the environment.
Let us explore oxygen, because the oxygen isotopic “thermometer” is such a critical one in paleoclimatic studies. In the environment, specifically the hydrologic cycle, there are two main forms of oxygen that exist, Oxygen-16 and Oxygen-18. There is some Oxygen-17 also, but its oxygen cousins are the big players. Oxygen-16 far outstrips Oxygen-18 in terms of prevalence, making up an average of 99.762% of all oxygen in the environment. Oxygen-18, by contrast, makes up an average of only 0.2% of all environmental oxygen. In both cases, these isotopes are taken up into organisms, clouds, precipitation, etc. and made into molecules. But, because Oxygen-16 and Oxygen-18 are treated differently than one another in these situations, depending upon the climate in a region, there can be very important variations in the ratio between these two isotopes. In some situations, such as in Arctic ice, Oxygen-18 is less common than average. In other situations, such as water vapor in clouds in the sub-tropics, Oxygen-18 would be more common than on average.
An excellent and in depth explanation of isotopic fractionation can be had from watching this video:
These variations from their average abundance provide useful metrics. In the case of some isotopes, the ratio between the two isotopes can be used as a proxy thermometer. In order to calculate the ratio and obtain a temperature, the Oxygen-16 to Oxygen-18 ratio from a sample is obtained using mass spectroscopy and then compared to a known standard. A commonly used known standard for coral reef data is SMOW (Standard Mean Ocean Water). The ratio is converted to “per mil” notation, or parts per thousand. The calculated ratio is represented using the Greek lower case letter delta, δ. So, the Oxygen-18 ratio is notated δ18O.
The name of this field of study is stable isotope thermometry. And, oxygen isotopes are not the only example that fractionates in the environment in such a useful way. An isotope of hydrogen, deuterium, is also incorporated into water. Deuterium, δ2H, is hydrogen with a neutron and is similarly as rare as Oxygen-18. It fractionates in much the same way in the environment when compared to its more common counterpart, plain old hydrogen.
Among the many other stable isotope thermometry examples are carbon isotopes. Carbon isotopes, notated as δ13C, are reflective of different environmental variables. The carbon cycle governs the fractionation of carbon isotopes. Carbon-12 is most common, while Carbon-13 and Carbon-14 less so. For thermometry, we use Carbon-13 to Carbon-12 ratios. The primary reason for this is that photosynthesis discriminates against Carbon-13 (and 14) in favor of Carbon-12. The atom is smaller and is more easily dealt with by plants. Vegetation is then a reservoir of carbon depleted in Carbon-13. When the
terrestrial biosphere is expanded, carbon with low δ13C values is sequestered. When the biosphere is more restricted (such as after extinction events), atmospheric carbon reservoirs are more enriched in δ13C. Changes in this ratio tell us something about the health of the biosphere, which can be directly related to climate. Used in conjunction with δ18O data, δ13C provides deeper insights into environmental changes associated with climate and the biosphere. One example of this is a δ13C excursion (big wiggle) that occurred right at the Paleocene-Eocene Thermal Maximum, which was discussed briefly earlier in this chapter. What caused a sudden and brief depletion of 13C? One hypothesis is that warming led to release of methane from coastal methane hydrate deposits, decreasing the δ13C ratio and contributing to a spike in warming and a small extinction episode. Methane from decomposing plant material would be much more depleted in δ13C.
There is a cave in the U.S. state of West Virginia called Buckeye Creek Cave. Buckeye Creek Cave is similar to any other cave formed in karst regions. Regions underlain by limestone that chemically weathers to form certain features, such as sinkholes, caverns, etc., are called karst. Karst is very common around the world and is a landscape that is particularly environmentally sensitive, given how easily water tends to flow through the environment. Like other caverns, Buckeye Creek Cave is located in a limestone, the Greenbriar Limestone. It contains speleothems, cave formations like stalactites and stalagmites — again like many other limestone caves.
Importantly and distinctively though, scientists have found 7,000 years of climate data in these Buckeye Creek Cave speleothems.
Caves form while below the water table. Once the water table drops, or the cave is elevated above the water table for other reasons, water typically begins dripping through joints and bedding in the bedrock above. Drips of water, over time, leave behind calcium carbonate speleothem deposits, the ones tour guides refer to variously as “cave bacon” or “gnomes” or “bridal veils.” But how does this happen?
The atmosphere contains carbon dioxide (CO2). This carbon dioxide, when combined with precipitating water, forms a weak acid(carbonic acid, H2CO3) that then infiltrates into the soil and bedrock. That soil, often the site of decomposition of organic material, can contain a good deal of carbon dioxide also, which can lower the pH of infiltrating water even further. This weak acid is enough to dissolve the limestone bedrock as it moves through, which dissociates the Ca+ ions from their former bicarbonate partner. Now an aqueous solution of carbon dioxide, bicarbonate, and Ca+ ions while in the rock, the water continues its work. That is, until it encounters the open space of a cavern. The water will run and drip as usual, but the open space allows the carbon dioxide to leave the solution and the Ca+ and bicarbonate then recombine. The entire process is called carbonation, and its responsible for those lovely karst features.
Air + Water → Carbonic Acid:
CO2 + H2O → H2CO3
Limestone + Carbonic Acid → Calcium Bicarbonate:
Ca(CO3) + H2CO3 → Ca(HCO3)2
As speleothems form, they create layers. As rain can be very seasonal, these layers can take on an annual character and be directly attributed to particular years. Because of this annual accumulation, scientists can use speleothem layers much like they do coral growth layers, tree rings, and ice cores.
Buckeye Creek Cave has three speleothem formations that record 7,000 years of data. According to researchers (Hardt et al., 2010), years where calcite is enriched in δ18O represent an increase in the summer precipitation in annual totals. Critically, the data also show a major change in δ18O at 4.2 ka, which is now recognized as the start of the Meghalayan age, as documented by a speleothem from northern India. The presence of this data in both locations a globe away testifies not only to the usefulness of this data, but also to the global nature of the 4.2 ka climate event. Below is the 7,000 year record of δ18O values.
Coral and Sclerosponges
Coral reefs are perhaps one of the more endangered groups of animals on this planet. Coral are marine animals that create massive skeletal reefs, adding to them annually, that create widespread habitat for myriad organisms in otherwise nutrient-poor tropical waters. Everything about coral reefs is about symbioses. It begins with the coral animals themselves, which cannot survive without the waste products, food to the coral, produced through photosynthesis practiced by the zooxanthellae algae that live with the coral polyps. Here is how it works.
Coral polyps bring in dissolved minerals from the seawater and produce proteins. At this point, the coral polyp begins secreting calcium carbonate, using the water around it, in which the current δ18O value will be recorded. The zooxanthellae algae provide the coral with oxygen, sugars, and fats through photosynthesis. In return, the coral polyp provides shelter, and nutrients needed by the algae for growth, such as nitrogen.
The coral polyps are filter feeders. This requires them to live near the surface of the water, where there is constant movement. This entire relationship requires water temperatures in a narrow range, stable sea levels, and the presence of other animals and plants in the community that contribute in other ways.
For paleoclimate researchers, it is the annual rings that are of particular interest. Because they record the δ18O values of the seawater at that time and location, they are a window into the temperature, particularly the water temperature. Other geochemical analyses on these rings can add to the insights provided by oxygen isotopes. Since coral have been in around in one form another as far back as the Cambrian Period, they are very useful proxy records.
Ocean microplankton form the very bottom of the food web and trophic pyramid. They are the primary producers of the marine realm, the survival of everything else hinging upon their continued happiness.
Some of these microfossils, foraminiferans, secrete calcium carbonate shells, or tests. Others secrete siliceous tests, made form silicon dioxide. The fossils will use the oxygen isotopes that are around and available, as well as the hydrogen, as they secrete these tests. Because they are so small and planktonic, which by definition means they float around, they are easily moved around by ocean currents. Once they die, their bodies rain down to the seafloor, where researchers can extract them from sediment core samples.
They are excellent paleoclimate indicators in a variety of ways. First, they do include δ18O values that are useful temperature indicators of the surrounding seawater (higher δ18O = lower temperature). This characteristic is not just unique to carbonate microfossils such as foraminiferans and coccoliths, but also to siliceous microfossils, like diatoms and radiolarians.
Some species are cold-tolerant. When this is known, finding them in particular areas is a really good indicator of a colder water temperature in that region at that time of their lives and deposition of their skeletal remains. Microfossils can also be indicators of upwelling and nutrient movement, which can tell scientists about the kinds of wind and weather patterns in that area at the time.
Of all of the paleoclimate proxies, ice cores are perhaps the most glamorous. This is not just because they are much more familiar to people, but also because the wealth of information contained in an ice core is quite astounding. Ice cores are extracted from glaciers using a core drill. Ice is layered on annually, like tree rings, coral growth, and speleothems. Each year, the annual snowfall eventually compacts to firn and then to ice. While doing so, it traps dust and bubbles of gas associated with that moment in time with it. Then it sits. Conveniently for us, scientists can come back 800,000 years later and pull this data “out of the freezer.”
Ice core data includes oxygen, deuterium, carbon, and often other stable isotopes. It also includes dust, atmospheric gas concentrations (such as carbon dioxide and methane), along with other geochemical data, including lead, sulfur, and other industrial and environmental parameters. Thin layers of volcanic ash help constrain the age of the ice. It’s a rich treasure trove of information!
There are numerous ice cores that have been extracted from mountain glaciers, ice sheets, and more. Perhaps the most famous is the Vostok ice core, extracted by the Soviet Union at Vostok Station in 1987. Formerly the oldest, but still one of the most detailed cores we have, goes back 800,000 years and is referred to as Antarctic Dome C, extracted by the EPICA project (European Project for Ice Coring in Antarctica) in 1996. The oldest ice core, containing ice dating as far back as 2.7 Ma, was extracted from the Allen Hills in Antarctica. by a team led by Princeton University in 2015.
Greenland has produced some important cores that contain long climate records also. The Greenland Ice Core Project Program (GRIP) and North Greenland Ice Core Project Program (NGRIP) have extracted multiple core samples from areas of Greenland where ice thickness exceeds several thousand feet. Like the Antarctic ice cores, these cores have been critical for understanding past global conditions, including temperatures, moisture, anthropogenic emissions, and other environmental factors.
Significant Paleoclimatological Events
Many significant paleoclimatological events from Earth’s long history have been mentioned in the text above. However, should you wish to explore in more detail some of these compelling stories, below are a series of links you can pursue. All of these have in common several themes. The climate changes. It does so for a variety of reasons and at a variety of rates. As you read more about these specific events, you should see many common themes emerge, but you should also see proxy data at work helping us understand our past climate.
Great Oxygenation Event
Coffin, Millard & Duncan, Robert & Eldholm, Olav & Fitton, J. & Frey, Fred & Larsen, Hans & Mahoney, John & Saunders, Andrew & Schlich, Roland & Wallace, Paul. (2006). Large Igneous Provinces and Scientific Ocean Drilling: Status Quo and A Look Ahead. Oceanography. 19. 150-160. 10.5670/oceanog.2006.13.
Zachos, J.C. & MO, Pagani & Sloan, L.C. & Thomas, Ellen & Billups, Katharina. (2001). Trends, Rhythms, and Aberrations in Global Climate 65 Ma to Present. Science (New York, N.Y.). 292. 686-93. 10.1126/science.1059412.
- 1 Paleoclimatology: Earth Systems Change Through Time
- 1.1 What is Climate?
- 1.2 Paleoclimatology: Historical Climate
- 1.3 Geological Sources of Important Elements for Life
- 1.3.1 Measuring Paleoclimatic Change: Proxy Records
- 22.214.171.124 Temporal Scale and Proxy Records
- 126.96.36.199.1 Biological Proxy Records
- 188.8.131.52.2 Physical Proxy Records
- 184.108.40.206.3 Chemical Proxy Records
- 220.127.116.11 Temporal Scale and Proxy Records
- 1.4 Significant Paleoclimatological Events
- 1.5 Further Reading: