15.3: Geological Sources of Important Elements for Life

Table \(\PageIndex{1}\): Table of important elements for life. All of these elements are ultimately sourced from within the Earth's interior. Plate tectonic processes, driven by Earth's interior heat, drive one end of the biogeochemical cycling of these nutrients while the Sun's incoming energy drives them from above.

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 - Quiz

Exercise \(\PageIndex{1}\)

At 1:27 in the above video, the ice core record going back 800,000 years begins. As you watch, you will begin to see a pattern emerge in the carbon dioxide data. Compare this with the above section of the text on Milankovitch Cycles (Extrinsic factors). Which cycle do you see controlling the very large up and down swings in carbon dioxide?

a. Eccentricity, about 100,000 years

b. Annual, about 1 year

c. Precession, about 26,000 years

d. Obliquity, about 41,000 years

Answer

a. Eccentricity, about 100,000 years

Figure \(\PageIndex{1}\): Atmospheric methane concentrations since 1988 (Source: NOAA Global Monitoring Laboratory, May 2020).

Figure \(\PageIndex{2}\): Atmosphere nitrous oxide concentrations, Mauna Loa, Hawaii, May 2020 (Source: NOAA Global Monitoring Laboratory).

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.

Figure \(\PageIndex{3}\): Declining atmospheric oxygen levels accompany increasing carbon dioxide levels due to combustion of greenhouse gases. While these decreases in oxygen are not enough to affect human health, they provide stark evidence for the effect of the burning of fossil fuels on the Earth’s changing climate (Source: Scripps \(\ce{O2}\) Program).

Figure \(\PageIndex{4}\): The carbon cycle, including anthropogenic inputs. Pg stands for petagrams of carbon per year (same as gigatonnes). Numbers in boxes represent the total mass of carbon in the reservoir (stock). Numbers next to arrows represent fluxes, or changes in carbon concentrations from 2000-2009. Black numbers are natural fluxes. Red numbers are anthropogenic fluxes. Both carbon dioxide and methane, among other carbon compounds, are included (Source: IPCC 5th Assessment Report).

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.

Figure \(\PageIndex{5}\): The carbon cycle, broken up into its two parts, the fast and slow carbon cycles. The fast carbon cycle moves carbon through the Earth’s systems across short time periods, such as through photosynthesis and decomposition. The slow carbon cycle sees fluxes in carbon storage across much longer time spans, such as storage of carbon in limestones and fossil fuel deposits. Numbers indicate flux magnitudes and are unitless (Source: NASA Earth Observatory).

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

Figure \(\PageIndex{6}\): Large Igneous Provinces (LIPs), also known as flood basalt provinces, exist in a wide variety of locations all over the planet (Source: Williamborg after USGS via Wikimedia)

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.

Figure \(\PageIndex{7}\): The fate of subducting lithosphere (blue) and the rising of hot spots and plumes that create LIP provinces. Transient LIPs, the kind that exist and then go away, likely originate from the core/mantle boundary from LLSPVs (Large Low Shear Velocity Provinces) (Used with permission: Coffin et al., 2006).

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.

Figure \(\PageIndex{8}\): LIP eruptions superimposed upon Sepkoski (1996) extinction of marine genera over time. There is a close or direct correlation between LIP eruptions and a number of extinction events in Earth’s history. Global climate changes are the main connection (Used with permission: Coffin et al., 2006).

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.

Figure \(\PageIndex{9}\): One effect of increasing carbon dioxide levels in the atmosphere is direct increase of carbon dioxide levels in the marine realm, as carbon dioxide emitted is concentrated first in the hydrosphere. Adding carbon dioxide to the ocean leads to a lowering of pH and dissociation of \(\ce{Ca^+}\) and bicarbonate, making shell production more difficult to impossible, in extreme scenarios (Source: Brittanica Online)

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.

Figure \(\PageIndex{10}\): Burning ice – methane hydrates (Source: USGS).

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.

Figure \(\PageIndex{11}\): Locations of methane gas hydrates, recovered and inferred, today (Source: USGS).

The Biosphere

Figure \(\PageIndex{12}\): Earth’s earliest atmosphere was made up of gases outgassing from the planet’s early formation (Source: NOAA).

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.

Figure \(\PageIndex{13}\): Earth’s second atmosphere was primarily volcanic. By this time, the internal dynamo was creating the magnetic field that prevented the solar wind from flowing this atmosphere away. About 4.2 Ga (Source: NOAA).

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.

Figure \(\PageIndex{14}\): Earth’s third atmosphere becomes our modern atmosphere. It is dominated by gases resulting from plant and animal respiration, but mostly planet photosynthesis. Begins to form about 3.8 Ga (Source: NOAA).

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.

Silicate Weathering

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.

Figure \(\PageIndex{15}\): The silicate weathering cycle.

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:

\[\bf{2(\ce{CO2}) + 3(\ce{H2O}) + \ce{CaSiO3} \rightarrow \ce{Ca^{2+}} + 2(\ce{HCO_3^-}) + \ce{H4SiO4}}\nonumber\]

Thus, the rise of mountains leads to global cooling. It also leads to the expansion of carbonate environments.

Figure \(\PageIndex{16}\): Tectonics drives climate on the scale of millions of years of change. A rise in atmospheric carbon dioxide can lead to a warmer climate. When coupled with rising mountains, increased carbon dioxide leads to increased chemical weathering, carbon dioxide removal, and then reduced warming. (Source: Public Domain)

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).

Figure \(\PageIndex{17}\): Major temperature fluctuations across the Paleocene-Eocene Thermal Maximum (PETM), as recorded by oxygen and carbon isotope excursions. Global temperatures reached a maximum at about 55 Ma, after which cooling began. Much of this cooling was very likely due to the collision of India with Asia at 45 Ma, a reduction in seafloor spreading rates at 39 Ma, and the rise of the Oligocene phase of the Andean Mountain Orogeny at 28 Ma. Note how much of the first half of the Cenozoic likely saw ice free polar regions. (Figure from Zachos et al., 2001; reproduced with permission)

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

Figure \(\PageIndex{18}\): The time span of some examples of proxy records. (Sourc: NOAA).

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.

Figure \(\PageIndex{19}\): Glossopteris sp. leaves from the Permian of Australia.

Figure \(\PageIndex{20}\): Artist’s rendition of a Glossopteris tree (Public Domain).

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.

Figure \(\PageIndex{21}\): Leaf margin terms. “Entire” refers to leaves with smooth edges, no teeth, no lobes, etc. (Source: CC BY-SA 3.0; By derivative work: McSush (talk)Leaf_morphology_no_title.png: User: Debivort – Leaf_morphology_no_title.png, https://commons.wikimedia.org/w/index.php?curid=7681206)

Figure \(\PageIndex{22}\): Percent Entire-Margined species versus temperature. There is a linear relationship between the two variables that is quantifiable. This relationship can be applied to fossil assemblages to help determine paleoclimatic conditions (Source: By Jack Wolfe – USGS, U.S. Govt. Printing Office, Public Domain, https://commons.wikimedia.org/w/index.php?curid=25994245).

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)

Figure \(\PageIndex{23}\): Electron microscope image of pollen grains from various common plants. Examples here include sunflowers, hollyhocks, lilies, caster beans, morning glories, and primroses. (Source: Windows2Universe.org via Chuck Daghlian and Louisa Howard, Dartmouth Electron Microscope Facility).

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.

Figure \(\PageIndex{24}\): Reconstruction of late Pleistocene and early Holocene waterways of the Sahara, primarily from faunal data (Source: Drake et al., 2011).

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.

Figure \(\PageIndex{25}\): Ocean sediment core sample depicting, among other things, the base and top of the PETM period. Note the color change from lighter colored to darker colored during this time (Source: NSF, Scripps Institution of Oceanography).

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.

Figure \(\PageIndex{26}\): Lake Mead water level drop visualization, 2001-2015 (Source: USGS).

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).

Figure \(\PageIndex{27}\): Growth rings on a tree at the Bristol Zoo, England. Each ring contains lighter early wood and darker late wood and represent a single year of annual growth.

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

Figure \(\PageIndex{28}\): Cross section of annual rings of Pinus Taeda. Early wood is lighter in color and late wood is darker in color (Source: CC BY-SA 3.0; By Pollinator, https://commons.wikimedia.org/w/index.php?curid=1674129)

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.

Figure \(\PageIndex{29}\): An increment borer, a drill with a hollow center, is used to extract core samples from trees without minimal damage. Samples can also be taken from structural wood (Source: CC BY-SA 2.5; By Hannes Grobe/AWI – Own work, https://commons.wikimedia.org/w/index.php?curid=1135047).

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.

Figure \(\PageIndex{30}\): Dust blowing across the Atlantic Ocean between July 6-12, 2018. Data from Suomi/NPP satellite, NASA. (Source: NOAA, via climate.gov).

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.

Figure \(\PageIndex{31}\): The “Loess Hills” in the U.S. state of Iowa (Source: CC BY-SA 4.0; By BillyBlueJay – Own work, https://commons.wikimedia.org/w/index.php?curid=69537356).

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

Figure \(\PageIndex{32}\): A canyon formed in the soft loess of Linxia County, Xihe Township, China (Source: CC BY-SA 3.0; By Vmenkov – Own work, https://commons.wikimedia.org/w/index.php?curid=7853712).

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.

Figure \(\PageIndex{33}\): Central China Loess Plateau region (Source: CC BY-SA 4.0; By Maintenance script – https://editors.eol.org/eoearth/wiki/Central_China_loess_plateau_mixed_forests, https://commons.wikimedia.org/w/index.php?curid=63963818).

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.

Figure \(\PageIndex{34}\): Stable oxygen atoms, Oxygen-16 and Oxygen-18. Both have eight protons, but Oxygen-18 has two extra neutrons (white balls), making it heavier by about 12.5% (Source: NASA Earth Observatory).

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.

Figure \(\PageIndex{35}\): A version of the hydrologic cycle, focused on oxygen isotope fractionation. “Heavy”, Oxygen-18 rich water condenses over lower latitudes and rains out, leaving “lighter” Oxygen-16 enriched water to head to the poles, depleted of Oxygen-18 (Source: NASA Earth Observatory).

An excellent and in depth explanation of isotopic fractionation can be had from watching this video:

\[\delta\ce{^{18}O} = \left( \frac{\left( \frac{\ce{^{18}O}}{\ce{^{16}O}}\right)_{\text{sample}}}{\left( \frac{\ce{^{18}O}}{\ce{^{16}O}}\right)_{\text{standard}}}-1\right) \times 1000 \%_{\text{o}} \nonumber\]

Figure \(\PageIndex{36}\): To calculate the Oxygen-18 and 16 ratio of a sample, the ratio of a the sample (obtained from spectroscopy) is divided by a known standard ratio, such as SMOW (Standard Mean Ocean Water), and then converted to “per mil” notation, or parts per thousand (Source: Wikipedia).

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, \(\delta\). So, the Oxygen-18 ratio is notated \(\delta \ce{^{18}O}\).

Figure \(\PageIndex{37}\): The oxygen isotope proxy thermometer. Concentration of Oxygen-18 decreases with temperature. The coldest sites in this data have about 5% less Oxygen-18 than ocean water (Source: NASA Earth Observatory).

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, \(\delta \ce{^2H}\), 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 \(\delta \ce{^{13}C}\), 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.

Figure \(\PageIndex{38}\): Depleted ratios of Oxygen-18 and Carbon-13 at the PETM boundary. Whatever caused the warming at the PETM may have been also responsible for the release of fossil carbon from decayed plant sources (After Zachos et al., 2001).

When the terrestrial biosphere is expanded, carbon with low \(\delta \ce{^{13}C}\) values is sequestered. When the biosphere is more restricted (such as after extinction events), atmospheric carbon reservoirs are more enriched in \(\delta \ce{^{13}C}\). Changes in this ratio tell us something about the health of the biosphere, which can be directly related to climate. Used in conjunction with \(\delta \ce{^{18}O}\) data, \(\delta \ce{^{13}C}\) provides deeper insights into environmental changes associated with climate and the biosphere. One example of this is a \(\delta \ce{^{13}C}\) 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 \(\ce{^{13}C}\)? One hypothesis is that warming led to release of methane from coastal methane hydrate deposits, decreasing the \(\delta \ce{^{13}C}\) ratio and contributing to a spike in warming and a small extinction episode. Methane from decomposing plant material would be much more depleted in \(\delta \ce{^{13}C}\).

Figure \(\PageIndex{39}\): Major temperature fluctuations across the Paleocene-Eocene Thermal Maximum (PETM), as recorded by oxygen can carbon isotope excursions. Global temperatures reached a maximum at about 55 Ma, after which cooling began. Much of this cooling was very likely due to the collision of India with Asia at 45 Ma, a reduction in seafloor spreading rates at 39 Ma, and the rise of the Andes Mountains at 28 Ma. Note how much of the first half of the Cenozoic likely saw ice free polar regions (Used with permission, Zachos et al., 2001).

Speleothems

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.

Figure \(\PageIndex{40}\): Cavern interior with six common speleothems labeled (Source: CC BY-SA 2.5; By Dave Bunnell / Under Earth Images – Own work, https://commons.wikimedia.org/w/index.php?curid=22613190).

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 (\(\ce{CO2}\)). This carbon dioxide, when combined with precipitating water, forms a weak acid(carbonic acid, \(\ce{H2CO3}\)) 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 \(\ce{Ca^+}\) ions from their former bicarbonate partner. Now an aqueous solution of carbon dioxide, bicarbonate, and \(\ce{Ca^+}\) 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 \(\ce{Ca^+}\) and bicarbonate then recombine. The entire process is called carbonation, and its responsible for those lovely karst features.

\[\text{Air } + \text{ Water } \rightarrow \text{ Carbonic Acid:} \\ \bf{\ce{CO2} + \ce{H2O} \rightarrow \ce{H2CO3}} \nonumber\]

\[\text{Limestone } + \text{ Carbonic Acid } \rightarrow \text{ Calcium Bicarbonate:} \\ \bf{\ce{Ca(CO3)} + \ce{H2CO3} \rightarrow \ce{Ca(HCO3)2}}\nonumber\]

Figure \(\PageIndex{41}\): Cut face of a stalagmite from Jeita Cave, Lebanon (Source: EGU).

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 \(\delta \ce{^{18}O}\) represent an increase in the summer precipitation in annual totals. Critically, the data also show a major change in \(\delta \ce{^{18}O}\) 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 \(\delta \ce{^{18}O}\) values.

Figure \(\PageIndex{42}\): 7,000 years of \(\delta \ce{^{18}O}\) values from Buckeye Creek Cave, West Virgina. Note the change beginning around 4.2 ka and then again in the last few years. This kind of data from karst environments is a measure of moisture (Source: Hardt et al., 2010).

Coral and Sclerosponges

Figure \(\PageIndex{43}\): Coral reef outcrop, Great Barrier Reef (Source: CC BY-SA 3.0; By Toby Hudson – Own work, https://commons.wikimedia.org/w/index.php?curid=11137678).

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.

Figure \(\PageIndex{44}\): Bleached Acropora coral, with unbleached (normal) coral in the background. Coral become bleached when they lose their zooxanthellae algae, typically due to high water temperatures (Source: Public Domain).

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 \(\delta \ce{^{18}O}\) 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.

Figure \(\PageIndex{45}\): Montastrea cavernosa coral polyps. (Public Domain)

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 \(\delta \ce{^{18}O}\) 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.

Figure \(\PageIndex{46}\): Coral growth rings are produced annually. The coral polyp will produce its calcite using oxygen available to it. In cooler times, there will be less Oxygen-18 than in warmer times. Because of this, both fossil and modern coral skeletons are excellent climate indicators (Source: NASA Earth Observatory).

Oceanic microplankton

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.

Figure \(\PageIndex{47}\): Various marine microfossils, inlcluding, from upper left to lower right, diatom, ostracod, radiolarian, sponge spicule, radiolarian, two planktonic foraminiferans, and a coccolith (Source: CC BY 3.0; By Hannes Grobe/AWI – Own work, https://commons.wikimedia.org/w/index.php?curid=8831075).

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 \(\delta \ce{^{18}O}\) values that are useful temperature indicators of the surrounding seawater (higher \(\delta \ce{^{18}O}\) = 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.

Ice Cores

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!

Figure \(\PageIndex{48}\): Dust, carbon dioxide, and temperature (from d\(\ce{^{18}O}\) values) for the last 420,000 years as taken from the Vostok Ice Core (Source: CC BY-SA 3.0; By Vostok-ice-core-petit.png: NOAAderivative work: Autopilot (talk) – Vostok-ice-core-petit.png, https://commons.wikimedia.org/w/index.php?curid=10684392).

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.

Figure \(\PageIndex{49}\): The evolution of temperature after the last major glacial advance and into the Holocene Epoch. Note the period of the Younger Dryas and the 8.2 ky event. This last event defines the boundary between the Northgrippian Age and the Greenlandian Age of the Holocene Epoch (Source: CC BY-SA 4.0; By Daniel E. Platt et al., https://commons.wikimedia.org/w/index.php?curid=78989262).