Outline the properties of water that are essential for life as we know it on Earth. Describe the hydrologic cycle including the important reservoirs for water on Earth and processes reservoirs (or pools). Define characteristics of soils and the factors that determine what type of soil develops in a given Discuss the layers in Earth's atmosphere and distinguish between climate and weather in Ecology. Illustrate the important factors in Earth's energy balance. Communicate the causes and climate consequences of Atmospheric and Oceanic circulation. Delineate the natural causes of climate change on Earth. Summarize some important changes in Earth's climate through geologic time.
2.1: Properties of Water 2.2: Water (Hydrologic) Cycle 2.3: Soils 2.4: The Atmosphere 2.5: Earth's Energy Balance 2.6: Atmospheric and Oceanic Circulation 2.7: What Makes the Climate Change 2.8: Past Climate Change
Summary This chapter describes the physical environment of the planet Earth and the importance of the physical environment for the field of Ecology. Sections 2.1 and 2.2 discuss important properties of water and they hydrologic cycle. Section 2.3 characterizes soils, their importance for ecology, and how they form. Earth's atmosphere, energy balance and atmospheric and oceanic circulation are detailed in sections 2.4, 2.5, and 2.6. Finally Sections 2.7 and 2.8 illustrate the natural causes of climate change on earth and some of the important past changes in Earth's climate. For a description of the human impact on current and future global climate, see the chapter on "Human Impact on Global Climate". 2: The Physical Environment is shared under a CC BY-NC-SA license and was authored, remixed, and/or curated by LibreTexts.
2.1: Properties of Water Do you ever wonder why scientists spend time looking for water on other planets? It is because water is essential to life, as we understand it; even minute traces of it on another planet can indicate that life could or did exist on that planet. Water is one of the more abundant molecules in living cells and the one most critical to life as we know it. Approximately 60-70 percent of your body is made up of water. Without it, life simply would not exist. Water Is a Polar Molecule The hydrogen and oxygen atoms within water molecules form polar covalent bonds. The shared electrons spend more time associated with the oxygen atom than they do with hydrogen atoms. There is no overall charge to a water molecule, but there is a slight positive charge on each hydrogen atom and a slight negative charge on the oxygen atom. Because of these charges, the slightly positive hydrogen atoms repel each other and form the unique shape seen in (Figure ).
: Hydrogen bonds form between slightly positive (+) and slightly negative (-) charges of polar covalent molecules, such as water.
Each water molecule attracts other water molecules because of the positive and negative charges in the different parts of the molecule allowing the formation of hydrogen bonds. Water also attracts and form hydrogen bonds with other polar molecules (such as sugars). When a substance readily forms hydrogen bonds with water, it can dissolve in water and is referred to as hydrophilic ("water-loving"). Hydrogen bonds are not readily formed with nonpolar substances like oils and fats. These nonpolar compounds are hydrophobic ("water-fearing") and will not dissolve in water.
Water Stabilizes Temperature Temperature is a measure of the motion (kinetic energy) of molecules. As the motion increases, energy is higher and thus temperature is higher. Water absorbs a great deal of energy before its temperature rises. Increased energy disrupts the hydrogen bonds between water molecules. The hydrogen bonds in water allow it to absorb and release heat energy more slowly than many other substances (giving water a high specific heat). Because these bonds can be created and disrupted rapidly, water absorbs an increase in energy and temperature changes only minimally. This means that water moderates temperature changes within organisms and in their environments. As energy input continues, the balance between hydrogen-bond formation and destruction swings toward the destruction side. More bonds are broken than are formed. This process results in the release of individual water molecules at the surface of the liquid (such as a body of water, the leaves of a plant, or the skin of an organism) in a process called evaporation. Evaporation of sweat, which is 90 percent water, allows for cooling of an organism, because breaking hydrogen bonds requires an input of energy and takes heat away from the body. Conversely, as molecular motion decreases and temperatures drop, less energy is present to break the hydrogen bonds between water molecules. These bonds remain intact and begin to form a rigid, lattice-like structure (e.g., ice) (Figure a). When frozen, ice is less dense than liquid water (the molecules are farther apart). This means that ice floats on the surface of a body of water (Figure b). In lakes, ponds, and oceans, ice will form on the surface of the water, creating an insulating barrier to protect the animal and plant life beneath from freezing in the water. If this did not happen, plants and animals living in water would freeze in a block of ice and could not move freely, making life in cold temperatures difficult or impossible.
Figure : (a) The lattice structure of ice makes it less dense than the freely flowing molecules of liquid water. Ice's lower density enables it to (b) float on water (credit a: modification of work by Jane Whitney; credit b: modification of work by Carlos Ponte). CONCEPT IN ACTION
Click here to see a 3-D animation of the structure of an ice lattice. (credit: image created by Jane Whitney using Visual Molecular Dynamics (VMD) software (Humphrey et al, 1996) Water Is an Excellent Solvent Because water is a polar molecule, with slight positive and negative charges, ionic compounds and polar molecules can readily dissolve in it. Water is, therefore, what is referred to as a solvent--a substance capable of dissolving another substance. The charged particles will form hydrogen bonds with a surrounding layer of water molecules. This is referred to as a sphere of hydration or hydration shells and serves to keep the particles separated or dispersed in the water. In the case of table salt (NaCl) mixed in water (Figure ), the sodium and chloride ions separate, or dissociate, in the water, and spheres of hydration are formed around the ions. A positively charged sodium ion is surrounded by the partially negative charges of oxygen atoms in water molecules. A negatively charged chloride ion is surrounded by the partially positive charges of hydrogen atoms in water molecules. The polarity of the water molecule makes it an effective solvent and is important in its many roles in living systems (i.e. cell transport, secretion of waste products, nutrient absorbtion).
Figure : When table salt (NaCl) is mixed in water, spheres of hydration form around the ions.
Water Is Cohesive Have you ever filled up a glass of water to the very top and then slowly added a few more drops? Before it overflows, the water actually forms a dome-like shape above the rim of the glass. This water can stay above the glass because of the property of cohesion. In cohesion, water molecules are attracted to each other (because of hydrogen bonding), keeping the molecules together at the liquid-air (gas) interface. Cohesion gives rise to surface tension, the capacity of a substance to withstand rupture when placed under tension or stress. When you drop a small scrap of paper onto a droplet of water, the paper floats on top of the water droplet, although the object is denser (heavier) than the water. This occurs because of the surface tension that is created by the water molecules. Cohesion and surface tension keep the water molecules intact and the item floating on the top. It is even possible to "float" a steel needle on top of a glass of water if you place it gently, without breaking the surface tension (Figure ).
Figure : The weight of a needle on top of water pulls the surface tension downward; at the same time, the surface tension of the water is pulling it up, suspending the needle on the surface of the water and keeping it from sinking. Notice the indentation in the water around the needle (credit: Cory Zanker).
These cohesive forces are also related to the water's property of adhesion, or the attraction between water molecules and other molecules. This is observed when water "climbs" up a straw placed in a glass of water. You will notice that the water appears to be higher on the sides of the straw than in the middle. This is because the water molecules are attracted to the straw and therefore adhere to it. Cohesive and adhesive forces are important for sustaining life. For example, because of these forces, water can flow up from the roots to the tops of plants for use in photosynthesis. Many aquatic organisms such as water striders or snails can travel along the surface of the water because of surface tension created by cohesive forces. CONCEPT IN ACTION To learn more about water, visit the U.S. Geological Survey Water Science for Schools: All About Water! website. pH, Acids, Bases, and Buffers The pH of a solution is a measure of its acidity or alkalinity. You have probably used litmus paper, paper that has been treated with a natural water-soluble dye so it can be used as a pH indicator, to test how much acid or base (alkalinity) exists in a solution. You might have even used some to make sure the water in an outdoor swimming pool is properly treated. In both cases, this pH test measures the amount of hydrogen ions that exists in a given solution. High concentrations of hydrogen ions yield a low pH (or acidic), whereas low levels of hydrogen ions result in a high pH (or alkaline/basic). The overall concentration of hydrogen ions is inversely related to its pH and can be measured on the pH scale (Figure ). Therefore, the more hydrogen ions present, the lower the pH; conversely, the fewer hydrogen ions, the higher the pH. The pH scale ranges from 0 to 14. A change of one unit on the pH scale represents a change in the concentration of hydrogen ions by a factor of 10, a change in two units represents a change in the concentration of hydrogen ions by a factor of 100. Thus, small changes in pH represent large changes in the concentrations of hydrogen ions. Pure water is neutral. It is neither acidic nor alkaline, and has a pH of 7.0. Anything below 7.0 (ranging from 0.0 to 6.9) is acidic, and anything above 7.0 (from 7.1 to 14.0) is alkaline. The blood in your veins is slightly alkaline (pH = 7.4). The environment in your stomach is highly acidic (pH = 1 to 2). Orange juice is mildly acidic (pH = approximately 3.5), whereas baking soda is basic (pH = 9.0).
Figure Stevens).
: The pH scale measures the amount of hydrogen ions (H+) in a substance (credit: modification of work by Edward
Acids are substances that provide hydrogen ions (H+) and lower pH, whereas bases provide hydroxide ions (OH-) and raise pH. The stronger the acid, the more readily it donates H+. For example, hydrochloric acid and lemon juice are very acidic and readily give up H+ when added to water. Conversely, bases are those substances that readily donate OH-. The OH- ions combine with H+ to produce water, which raises a substance's pH. Sodium hydroxide and many household cleaners are very alkaline and give up OH- rapidly when placed in water, thereby raising the pH.
Most cells in our bodies operate within a very narrow window of the pH scale, typically ranging only from 7.2 to 7.6. If the pH of the body is outside of this range, the respiratory system malfunctions, as do other organs in the body. Cells no longer function properly, and proteins will break down. Deviation outside of the pH range can induce coma or even cause death.
So how is it that we can ingest or inhale acidic or basic substances and not die? Buffers are the key. Buffers readily absorb excess H+ or OH-, keeping the pH of the body carefully maintained in the aforementioned narrow range. Carbon dioxide is part of a prominent buffer system in the human body; it keeps the pH within the proper range. This buffer system involves carbonic acid (H2CO3) and bicarbonate (HCO3-) anion. If too much H+ enters the body, bicarbonate will combine with the H+ to create carbonic acid and limit the decrease in pH. Likewise, if too much OH- is introduced into the system, carbonic acid will rapidly dissociate into bicarbonate and H+ ions. The H+ ions can combine with the OH- ions, limiting the increase in pH. While carbonic acid is an important product in this reaction, its presence is fleeting because the carbonic acid is released from the body as carbon dioxide gas each time we breathe. Without this buffer system, the pH in our bodies would fluctuate too much and we would fail to survive.
The same buffer system also works in the ocean where CO2 reacts with seawater to form carbonic acid (H2CO3), which dissociates into bicarbonate (HCO3-) and carbonate (CO32-) ions:
The sum of these three carbon species is called dissolved inorganic carbon (
equilibrium between the species depends on the pH. In the current ocean, pH is about 8.1, which leads to about 86.5 % of DIC
being in the form of bicarbonate ions, 13.0 % in the form of carbonate ions, and only 0.5 % in the form of aqueous CO2 (Fig. \ (\PageIndex{6}; Zeebe and Wolf-Gladrow, 2001).
Figure : Ratios of carbonate species concentrations as a function of the pH. Currently average seawater has a pH of about 8.1. Therefore most carbon in the ocean is in the form of bicarbonate. Addition of anthropogenic CO2 decreases the pH. "Bjerrum Plot" is available in the public domain.
Dissociation of carbonic acid into bicarbonate (baking soda) produces a hydrogen ion H+, which increases the pH of the water. Most hydrogen ions, however, re-combine with carbonate ions to form additional bicarbonate ions. Nevertheless, adding CO2 to seawater increases its hydrogen ion concentration (decreases its pH) and decreases the carbonate ion concentration. This process is called ocean acidification. Summary Water has many properties that are critical to maintaining life. It is a polar molecule, allowing for the formation of hydrogen bonds, which allow ions and other polar molecules to dissolve in water. Therefore, water is an excellent solvent. The hydrogen bonds between water molecules give water the ability to hold heat better than many other substances. As the temperature rises, the hydrogen bonds between water continually break and reform, allowing for the overall temperature to remain stable, although increased energy is added to the system. Water's cohesive forces allow for the property of surface tension. All of these unique properties of water are important in the chemistry and physics of living organisms. The pH of a solution is a measure of the concentration of hydrogen ions in the solution. A solution with a high number of hydrogen ions is acidic and has a low pH value. A solution with a high number of hydroxide ions is basic and has a high pH value. The pH scale ranges from 0 to 14, with a pH of 7 being neutral. Buffers are solutions that moderate pH changes when an acid or base is added to the buffer system. Buffers are important in biological systems because of their ability to maintain constant pH conditions. Sources Humphrey, W., Dalke, A., & Schulten, K. (1996). VMD--Visual Molecular Dynamics. Journal of Molecular Graphics, 14(1), 33- 38. http://www.ks.uiuc.edu/Research/vmd/ Zeebe, R. E., & Wolf-Gladrow, D. A. (2001). CO in seawater: Equilibrium, kinetics, isotopes. Elsevier. Contributors and Attributions Modified by Kyle Whittinghill (University of Pittsburgh) from the following sources: Samantha Fowler (Clayton State University), Rebecca Roush (Sandhills Community College), James Wise (Hampton University). Original content by OpenStax (CC BY 4.0; Access for free at https://cnx.org/contents/b3c1e1d2-83...4e119a8aafbdd). 2.1: The Building Blocks of Molecules https://bio.libretexts.org/Bookshelv...fig-ch02_01_06 1.4: Carbon by Andreas Schmittner is licensed CC BY-NC 4.0. 2.1: Properties of Water is shared under a CC BY-NC-SA license and was authored, remixed, and/or curated by LibreTexts. 2.2: Water by OpenStax is licensed CC BY 4.0.
2.2: Water (Hydrologic) Cycle Water is essential for all living processes. The human body is more than one-half water and human cells are more than 70 percent water. Water is also the only substance that occurs naturally on earth in three forms: solid, liquid and gas. Of the stores of water on Earth, 97.5% is salt water (Figure ). Of the remaining water, 99% is locked as underground water or ice. Ice caps and glaciers are the largest reservoirs of fresh water but this water is inconveniently located, mostly in Antarctica and Greenland. Shallow groundwater is the largest reservoir of usable fresh water. Less than one percent of fresh water is present in lakes and rivers, the most heavily used water resource. Many organisms are dependent on this small amount of surface fresh water supply, a lack of which can have important effects on ecosystem dynamics. If all of world's water was shrunk to the size of 1 gallon, then the total amount of fresh water would be about 1/3 cup, and the amount of readily usable fresh water would be 2 tablespoons.
Figure : Earth's Water Reservoirs -- Bar chart Distribution of Earth's water including total global water, fresh water, and surface water and other fresh water and Pie chart Water usable by humans and sources of usable water. Source: United States
Geographical Survey Igor Skiklomanov's chapter "World fresh water resources" in Peter H. Gleick (editor), 1993, Water in Crisis: A Guide to the World's Fresh Water Resources.
The hydrosphere is the area of Earth where water movement and water storage occurs. Water reservoirs are the locations where
water is stored. (Note that this term can also refer to artificial lakes created by dams.) Water is found as a liquid on the surface
(rivers, lakes, oceans) and beneath the surface (groundwater), as ice (polar ice caps and glaciers), and as water vapor in the
illustrates the average time that an individual water molecule may spend in the Earth's major water
reservoirs. Residence time is a measure of the average time an individual water molecule stays in a particular reservoir.
Figure : Average residence time that water remains in each reservoir. Image from OpenStax is licensed under CC-BY.
The Water Cycle The various processes that occur during the cycling of water are illustrated in Figure evaporation and sublimation transpiration condensation and precipitation infiltration and subsurface water flow (groundwater discharge) surface runoff and snowmelt streamflow
Figure : Water from the land and oceans enters the atmosphere by evaporation or sublimation, where it condenses into clouds and falls as rain or snow. Precipitated water may enter freshwater bodies or infiltrate the soil. The cycle is complete when surface or groundwater reenters the ocean (credit: modification of work by John M. Evans and Howard Perlman, USGS).
The water cycle is driven by the Sun's energy as it warms the oceans and other surface waters. This leads to evaporation (water to water vapor) of liquid surface water and sublimation (ice to water vapor) of frozen water, thus moving large amounts of water into the atmosphere as water vapor. As the water vapor rises in the atmosphere, it cools and condenses. Condensation is the process in which water vapor changes to tiny droplets of liquid water. Over time, this water vapor condenses into clouds as liquid or frozen droplets and eventually leads to precipitation (rain or snow), which returns water to Earth's surface. Most precipitation falls into the ocean. Some frozen precipitation becomes part of ice caps and glaciers. These masses of ice can store frozen water for hundreds of years or longer. Rain reaching Earth's surface may evaporate again, flow over the surface, or percolate into the ground. Most easily observed is surface runoff: the flow of fresh water either from rain or melting ice. Runoff can make its way through streams and lakes to the oceans or flow directly to the oceans themselves. In most natural terrestrial environments rain encounters vegetation before it reaches the soil surface. A significant percentage of water evaporates immediately from the surfaces of plants. What is left reaches the soil and begins to move down. Surface runoff will occur only if the soil becomes saturated with water in a heavy rainfall. Infiltration is the process through which water sinks into the ground and is determined by the soil or rock type through which water moves. Soil texture (particle size) is an important control on water movement. Most water in the soil will be taken up by plant roots. Water enters the vascular system of the plant through the roots and evaporates, or transpires, through the stomata of the leaves. The plant will use some of this water for its own metabolism, and some of that will find its way into animals that eat the plants, but much of it will be lost back to the atmosphere through a process known as evapotranspiration. Water in the soil that is not taken up by a plant and that does not evaporate is able to percolate into the subsoil and bedrock. Here it forms groundwater.
Groundwater is a significant reservoir of fresh water. It exists in the pore spaces between particles in sand and gravel, or in the fissures in rocks. Shallow groundwater flows slowly through these pore spaces and fissures and eventually finds its way to a stream or lake where it becomes a part of the surface water again. Streams do not flow because they are replenished from rainwater directly; they flow because there is a constant inflow from groundwater below. Some groundwater is found very deep in the bedrock and can persist there for millennia. Most groundwater reservoirs, or aquifers, are the source of drinking or irrigation water drawn up through wells. In many cases these aquifers are being depleted faster than they are being replenished by water percolating down from above. An important part of the water cycle is how water varies in salinity, which is the abundance of dissolved ions in water. Freshwater (such as lakes, rivers, and near-surface groundwater) has a relatively low salinity. The saltwater in the oceans is highly saline, with about 35,000 mg of dissolved ions per liter of seawater. Evaporation is a distillation process that produces nearly pure water with almost no dissolved ions. As water vaporizes, it leaves the dissolved ions in the original liquid phase. Eventually, condensation forms clouds and sometimes precipitation. After rainwater falls onto land, it dissolves minerals in rock and soil, which increases its salinity. Rain and surface runoff are major ways in which minerals, including phosphorus and sulfur, are cycled from land to water. Precipitation Patterns Rain and surface runoff are major ways in which minerals, including carbon, nitrogen, phosphorus, and sulfur, are cycled from land to water. More precipitation falls near the equator, and landmasses there are characterized by a tropical rainforest climate (Figure ). Less precipitation tends to fall near 20-30° north and south latitude, where the world's largest deserts are located. These rainfall and climate patterns are related to global wind circulation cells. The intense sunlight at the equator heats air, causing it to rise and cool, which decreases the ability of the air mass to hold water vapor and results in frequent rainstorms. Around 30° north and south latitude, descending air conditions produce warmer air, which increases its ability to hold water vapor and results in dry conditions. Both the dry air conditions and the warm temperatures of these latitude belts favor evaporation. Global precipitation and climate patterns are also affected by the size of continents, major ocean currents, and mountains.
Figure : World Rainfall Map -- The false-color map above shows the amount of rain that falls around the world. Areas of high rainfall include Central and South America, western Africa, and Southeast Asia. Since these areas receive so much rainfall, they are where most of the world's rainforests grow. Areas with very little rainfall usually turn into deserts. The desert areas include North Africa, the Middle East, western North America, and Central Asia. Source: United States Geological Survey Earth Forum,
Houston Museum Natural Science. In the United States, the 100th Meridian roughly marks the boundary between the humid and arid parts of the country (figure ). Irrigation is required to grow crops west of the 100th Meridian. In the West, surface water is stored in reservoirs (artificial lakes) and mountain snowpacks and strategically released through a system of canals during times of high use.
Figure : Distribution of precipitation in the United States. The 100th Meridian is approximately where the average precipitation transitions from relatively wet to dry. Eastern states are mostly green, indicating greater precipitation than most
western states, which are mainly orange or red (indicating low precipitation). Source: U.S. Geological Survey. Further "Reading" The steps of the water cycle are also explained in the video below.
For more information on the water cycle you might want to watch this water cycle video from USGS. Contributors and Attributions Modified by Kyle Whittinghill and Melissa Ha from the following sources Samantha Fowler (Clayton State University), Rebecca Roush (Sandhills Community College), James Wise (Hampton University). Original content by OpenStax (CC BY 4.0; Access for free at https://cnx.org/contents/b3c1e1d2-83...4e119a8aafbdd). 20.2: Biogeochemical Cycles by OpenStax, is licensed CC BY from Biology 2e by OpenStax (CC-BY). Access for free at openstax.org. 6.6: Water Cycle by CK-12: Biology Concepts, is licensed CC BY-NC Biogeochemical Cycles and the Flow of Energy in the Earth System and Water Cycle and Fresh Water Supply from Sustainability: A Comprehensive Foundation by Tom Theis and Jonathan Tomkin, Editors (licensed under CC-BY). Download for free at CNX. Water Cycle and Fresh Water Supply, Water Supply Problems and Solutions, and Biogeochemical Cycles from Environmental Biology by Matthew R. Fisher (licensed under CC-BY) The Water Cycle, Water Use and Distribution, and Groundwater from An Introduction to Geology by Chris Johnson et al. (licensed under CC-BY-NC-SA) Essentials of Environmental Science by Kamala Dorsner is licensed under CC BY 4.0. 2.2: Water (Hydrologic) Cycle is shared under a CC BY-NC-SA license and was authored, remixed, and/or curated by LibreTexts. 20.2: Biogeochemical Cycles by OpenStax is licensed CC BY 4.0.
2.3: Soils Soil is the outer loose layer that covers the surface of Earth. Soil quality is a major determinant, along with climate, of plant distribution and growth. Soil quality depends not only on the chemical composition and texture of the soil, but also the topography (regional surface features) and the presence of living organisms. In agriculture, the history of the soil, such as the cultivating practices and previous crops, modify the characteristics and fertility of that soil. Soil develops very slowly over long periods of time, and its formation results from natural and environmental forces acting on mineral, rock, and organic compounds. Soils can be divided into two groups: organic soils are those that are formed from sedimentation and primarily composed of organic matter, while those that are formed from the weathering of rocks and are primarily composed of inorganic material are called mineral soils. Mineral soils are predominant in terrestrial ecosystems, where soils may be covered by water for part of the year or exposed to the atmosphere. Soil is Important Soil is important to our society primarily because it provides the foundation of agriculture and forestry. Soil plays a key role in plant growth. Beneficial aspects to plants include providing physical support, water, heat, nutrients, and oxygen. Mineral nutrients from the soil can dissolve in water and then become available to plants. Soil plays a role in nearly all biogeochemical cycles on the Earth's surface. Global cycling of key elements such as carbon (C), nitrogen (N), phosphorous (P), and sulfur (S) all pass through soil. In the hydrologic (water) cycle, soil helps to mediate infiltration (percolating) from the surface into the groundwater. Microorganisms living in soil can also be important components of biogeochemical cycles through the action of decomposition and other processes such as nitrogen fixation. Through their roots, plants absorb water and minerals (e.g., nitrates, phosphates, potassium, copper, zinc). Of the mineral nutrients absorbed from the soil, macronutrients, including nitrogen (N), potassium (K), calcium (Ca), magnesium (Mg), phosphorus (P), sulfur (S), and silicon (Si) are needed by plants in significant quantities. Micronutrients are essential elements that are needed only in small quantities, but can still be limiting to plant growth since these nutrients are not so abundant in nature. Micronutrients include chlorine (Cl), iron (Fe), boron (B), manganese (Mn), sodium (Na), zinc (Zn), copper (Cu), nickel (Ni), and molybdenum (Mo). There are some other elements that tend to aid plant growth but are not absolutely essential. Micronutrients and macronutrients are desirable in particular concentrations and can be detrimental to plant growth when concentrations in soil solution are either too low (limiting) or too high (toxicity). Mineral nutrients are useful to plants only if they are in an extractable form in soil solutions, such as a dissolved ion rather than in solid mineral. Many nutrients move through the soil and into the root system as a result of concentration gradients, moving by diffusion from high to low concentrations across semi-permeable membranes in root cells (passive transport). However, some nutrients are selectively absorbed by active transport across the root membranes, enabling concentrations to become higher inside the plant than in the soil. Although many aspects of soil are beneficial to plants, excessively high levels of trace metals (either naturally occurring or added by humans) or applied herbicides can be toxic to some plants (Figure ).
: These native plants are affected by herbicides that were used to kill nearby weeds. Image by Matt Lavin is available under CC-BY-SA 2.0.
An important factor affecting soil fertility is soil pH (the negative log of the hydrogen ion concentration). Soil pH is a measure of the acidity or alkalinity of the soil solution. On the pH scale (0 to 14) a value of seven represents a neutral solution; a value less than seven represents an acidic solution and a value greater than seven represents an alkaline solution. Soil pH affects the health of microorganisms in the soil and controls the availability of nutrients in the soil solution. Strongly acidic soils (less than 5.5) hinder the growth of bacteria that decompose organic matter in the soil. This results in a buildup of organic matter that has yet to be decomposed, which leaves important nutrients such as nitrogen in forms that are unusable by plants. Soil pH also affects the solubility of nutrient-bearing minerals. This is important because the nutrients must be dissolved in solution for plants to assimilate them through their roots. Most minerals are more soluble in slightly acidic soils than in neutral or slightly alkaline soils. Strongly acid soils (pH four to five), though, can result in high concentrations of aluminum, iron and manganese in the soil solution, which may inhibit the growth of some plants.
Several factors determine soil pH. Organic material in soil decreases pH to an extent, but it also acts as a buffer, limiting changes in pH. Climate is also important, with high amounts of rainfall lowering pH and increasing leaching (draining of nutrients and other dissolved ions with percolating water). Some types of parent material, such as those high in silicon, decrease pH, while others, such as limestone increase pH.
Soil consists of organic matter (about 5%), inorganic mineral matter (40-45% of soil volume), water (about 25%) and air (about 25%) (Figure ). The amount of each of the four major components of soil depends on the amount of vegetation, soil compaction, and water present in the soil.
: The four major components of soil are shown: inorganic minerals, organic matter, water, and air. Image is from OERTX and is available in the public domain.
The organic material consists of dead organisms in various stages of decomposition. It is dark-colored because it contains humus, partially decayed matter containing organic acids. Humus enriches the soil with nutrients, gives the soil a loose texture that holds water, and allows air to diffuse through it. Oxygen is important for plant roots and many inhabitants of the soil. The organic component of soil serves as a cementing agent, returns nutrients to the plant, allows soil to store moisture, makes soil tillable for farming, and provides energy for soil microorganisms. Most soil microorganisms--bacteria, algae, or fungi--are dormant in dry soil, but become active once moisture is available.
The inorganic material of soil consists of rock, slowly broken down into smaller particles that vary in size. Soil particles that are 100 m to 2 mm in diameter are sand. (A micrometer, m, 10-6 m, or a millionth of a meter.) Soil particles between 2 and 100 m are called silt, and even smaller particles, less than 2 m in diameter, are called clay.
Soil should ideally contain 50 percent solid material and 50 percent pore space (Figure ). Pore space refers to the gaps in between soil particles. The larger the soil particles, the larger the pore spaces. Water can quickly pass through large pore spaces, so soils high in sand drain easily. Smaller soil particles have more surface area relative to volume and produce narrow pore spaces. Water clings to these surfaces, and soils high in clay thus retain water. (Clay is also negatively charged, which attracts water.) About one-half of the pore space should contain water, and the other half should contain air.
Figure : Soil particles and the pore spaces in between them. The black shading represents water, which fills the pore spaces in the lower portion of the soil and clings to narrow portions of pore space in the top portion of the soil. Image from "Forest
physiography; physiography of the United States and principles of soils in relation to forestry" is available in the public domain. Soil texture is based on percentages of sand, silt, and clay (Figure ). Soils that have a high percentage of one particle size are named after that particle (a clay soil has a high percentage of clay). Other soils have a mixture of two particle sizes and very little of the third size. For example, silty clay has roughly 50% clay and 50% silt while sandy clay has 50-60% sand and 35-50% clay. Some soils have no dominant particle size and contain a mixture of sand, silt, and humus. These soils are called loams, and they are optimal for agriculture. A medium loam has roughly 40% sand, 40% silt, and 20% clay. Larger particles (sand) facilitate drainage, and small particles (clay) facilitate water retention, so loam soils both have good drainage and can remain moist. Soils that deviate slightly from a medium loam include loamy sand, sandy loam, sandy clay loam, clay loam, silty clay loam, and silty loam.
Figure : A soil texture triangle is used to determine the soil texture based on the percentages of sand, silt, and clay. The left axis represents the percentage of clay. The right axis represents the percentage of silt. The bottom axis represents the percentage of sand. Most soils that contain more than 40% of clay particles are classified as clays. Exceptions are silty clays, which contain 4060% clay and 40-60% silt. Another exception are sandy clays, which contain 35-50% clay and 50-70% sand. Clay loams and silty
clay loams contain 25-40% clay. Clay loams contain 20-60% silt, and silty clay loams contain 60-75% silt. Sandy clay loams contain 20-35% clay and 50-80% sand. Medium loams contain 5-25% clay, 20-50% silt, and 30-55% sand. Soils with more than 50% silt and no more than 25% clay are either silty loams or silts. Silts have 80-100% silt and no more than 15% clay. Sand have 90-100% sand. Loamy sands have 75-90% sand and no more than 15% clay. A soil with a composition that does not fit into the already described categories would be a sandy loam. To determine the soil texture, first find the percentage of clay and trace a line horizontally to the right. At the same time, find the percentage of silt and trace a line diagonally (down and to the left, parallel to the clay axis). The two lines will meet at the correct soil type. For example, traced lines from 20% clay and 40% silt intersect at the medium loam category. To confirm, trace a diagonal line (down and to the right, parallel to the sit axis) from this point to the sand axis. This line cross the axis at 40%. In summary, a line can be traced from each axis that is parallel to axis that is counterclockwise of it. "Soil Composition" by Richard Wheeler is available under CC-BY-SA 3.0. Soils can be divided into two groups based on how they form. Organic soils are those that are formed from sedimentation and often contain more than 30% organic matter. They form when organic matter, such as leaf litter, is deposited more quickly than it can be decomposed (Figure ). Mineral soils are formed from the weathering of rocks, typically contain no more than 30% organic matter, and are primarily composed of inorganic material. Weathering occurs when biological, physical, and chemical processes, such as erosion, leaching, or high temperatures, break down rocks.
: Decomposition of organic material occurs very slowly in this acidic bog. The organic material accumulates, which is characteristic of organic soils. "Sphagnum Bog" by William L. Farr is available under CC-BY-SA 4.0.
Soil has Horizons Soil distribution is not uniform because its formation results in the production of layers; together, the vertical section of a soil is called the soil profile. Within the soil profile, soil scientists define zones called horizons. A horizon is a soil layer with distinct physical and chemical properties that differ from those of other layers. The soil profile has four distinct layers: 1) O horizon; 2) A horizon; 3) B horizon and 4) C horizon (Figure ).
Figure : This image shows the different horizons, or layers, in soil. Vegetation grows at the surface. The top, deep brown layer is the O (organic) horizon. Below this is the gray A horizon (topsoil, or surface). Next is the light brown B horizon (subsoil). The C horizon (substratum, or soil base) is tan and includes larger pieces of rocks. The gray bedrock is the deepest layer. The E (eluviated) horizon that is sometimes found between the A and B horizons is absent here. "Graphic of Soil Profile" by Wilsonbiggs is licensed
under CC BY-SA 4.0. Upper horizons (labeled as the A and O horizons) are richer in organic material and so are important in plant growth, while deeper layers (such as the B and C horizons) retain more of the original features of the bedrock below. Some soils may have additional layers (like the E horizon, Figure ), or lack one of these layers. The thickness of the layers is also variable, and depends on the factors that influence soil formation. In general, immature soils may have O, A, and C horizons, whereas mature soils may display all of these, plus additional layers.
Figure ): Photograph shows a soil profile from South Dakota revealing three horizons. The A horizon (topsoil) is dark brown and extend about 1 foot. Below it is the white E (eluviated) horizon, which is about 6 inches thick. The B horizon (subsoil, labeled Bt) is the medium brown and the lowest visible layer The yellow arrows symbolize translocation of fine clays to the B horizon. The
scale is in feet. Source: University of Idaho and modified by D. Grimley.
O horizon The very top of the O horizon consists of partially decayed organic debris like leaves (Figure color because of humus.
A horizon The A horizon (topsoil) consists of a mixture of organic material (from decomposition) with inorganic products of chemical weathering of rocks, and it is therefore the beginning of true mineral soil (Figure ). In this area, rainwater percolates through the soil and carries materials from the surface. The A horizon may be only 5 cm (2 in.), or it may over a meter. For instance, river deltas like the Mississippi River delta have deep layers of topsoil. Microbial processes occur in the top soil, and this horizon supports plant growth. Many organisms, such as earthworms and insects live among the plant roots in this horizon.
B horizon The B horizon (subsoil) consists of small particles that have moved downward, resulting in a dense layer in the soil (Figure ). In some soils, the B horizon contains nodules or a layer of calcium carbonate. The subsoil is usually lighter in color than topsoil and often contains an accumulation of minerals.
C horizon The C horizon (soil base), includes the parent material, the organic and inorganic substances from which soils form (Figure ). Weathering parent material represents the first steps in the chemical breakdown of rock into soil. Often the weathered parent material is underlain by the parent material itself, although in some places it has been carried from another location by wind, water, or glaciers. Beneath the C horizon lies bedrock. The chemical nature of the parent material, whether granite, limestone, or sandstone, for example, has a great influence on the fertility of the soil derived from it.
Several Factors Affect Soil Formation and Composition The fundamental factors that affect soil genesis can be categorized into five elements: climate, organisms, topography, parent material, and time. One could say that the relief, climate, and organisms dictate the local soil environment and act together to cause weathering and mixing of the soil parent material over time.
The role of climate in soil development includes aspects of temperature and precipitation. Soils in very cold areas with permafrost conditions (such as the arctic tundra) tend to be shallow and weakly developed due to the short growing season. In warm, tropical climates, soils tend to be thicker (but lacking in organic matter), with extensive leaching and mineral alteration. In such climates, organic matter decomposition and chemical weathering occur at an accelerated rate. The presence of moisture and nutrients from weathering will also promote biological activity: a key component of a quality soil.
Ancient soils, sometimes buried and preserved in the subsurface, are referred to as paleosols (Figure and environmental conditions.
Figure : Modern versus Buried Soil Profiles: a buried soil profile, or paleosol (above geologist `s head), represents soil development during the last interglacial period, while a modern soil profile (Alfisol) occurs near the land surface. Source: D.
Grimley. Organisms The presence of living organisms in the soil (soil biota) greatly affects soil formation and structure. A diversity of animals are found in the soil such as nematodes, spiders, insects, centipedes, millipedes, pillbugs, slugs, and earthworms (Figure ). The soil also contains microorganisms like bacteria, archaea, fungi, and "protists". Animals and microorganisms can produce pores and crevices, and plant roots can penetrate into crevices to produce more fragmentation. Additionally, leaves and other material that fall from plants decompose and contribute to soil composition. Microorganisms not only decompose organic matter, but contribute to other processes in nutrient cycles, such as nitrogen fixation. Many soil animals also contribute decomposition either directly as detritivores or indirectly through the soil food web.
Figure : Illustration of animals commonly found in the soil. "Soil Fauna" by Arne Henriks is licensed under CC BY 2.0.
Parent Material Mineral soils form directly from the weathering of bedrock, the solid rock that lies beneath the soil, and therefore, they have a similar composition to the original rock. Other soils form in materials that came from elsewhere, such as sand and glacial drift. Materials located in the depth of the soil are relatively unchanged compared with the deposited material. Sediments in rivers may have different characteristics, depending on whether the stream moves quickly or slowly. A fast-moving river could have sediments of rocks and sand, whereas a slow-moving river could have fine-textured material, such as clay. The type of parent material may also affect the rapidity of soil development. Parent materials that are highly weatherable (such as volcanic ash) will transform more quickly into highly developed soils, whereas parent materials that are quartz-rich rocks (and more difficult to break down) will take longer to develop into soils. Parent materials also provide nutrients to plants and can affect soil internal drainage. Topography Regional surface features (familiarly called "the lay of the land") can have a major influence on the characteristics and fertility of a soil. Topography affects water runoff, which strips away parent material and affects plant growth. Soils on steep slopes are more prone to erosion and may be thinner than soils that are on relatively level ground. Infiltration, the percolating of water through the soil, is limited in steep soils. The local topography can have important microclimatic effects. In the northern hemisphere, south-facing slopes are exposed to more direct sunlight angles and are thus warmer and drier than north-facing slopes. The cooler, moister north-facing slopes have a more dynamic plant community and thicker soils because extensive root systems stabilize the soil and reduce erosion (figure ).
: A north-facing slope (left) and south-facing slope (right) in the Mediterranean climate (chaparral) of the Santa
Monica Mountains, California. Vegetation on the north-facing slope is denser due to cool, moist conditions relative to the south-
facing slope. "Slope Effect" by Noah Elhardt is licensed under CC BY-SA 2.5.
Time Time is an important factor in soil formation because soils develop over long periods. Soil formation is a dynamic process. Materials are deposited over time, decompose, and transform into other materials that can be used by living organisms or deposited onto the surface of the soil. In general, soil profiles tend to become thicker (deeper), more developed, and more altered over time. However, the rate of change is greater for soils in youthful stages of development. The degree of soil alteration and deepening slows with time and at some point, after tens or hundreds of thousands of years, may approach an equilibrium condition where erosion and deepening (removals and additions) become balanced. Young soils (< 10,000 years old) are strongly influenced by parent material and typically develop horizons and character rapidly. Over time, as weathering processes deepen, mix, and alter the soil, the parent material becomes less recognizable as chemical, physical, and biological processes take effect. Moderate age soils (roughly 10,000 to 500,000 years old) are slowing in profile development and deepening, and may begin to approach equilibrium conditions. Old soils (>500,000 years old) have generally reached their limit as far as soil horizonation and physical structure, but may continue to alter chemically or mineralogically. Soil development is not always continual. Geologic events such as landslides, glacier advance, or the rising of shorelines can rapidly bury soils. Erosion in rivers and shorelines can cause removal or truncation of soils, and wind or flooding slowly deposit sediment that adds to the soil. Animals can mix the soil and sometimes cause soil regression, a reversal or "bump in the road" for the normal path of development, and this increases development over time.
Soils are classified into one of 12 soil orders based on soil horizons, how they form, and their chemical compositions. For example,
), which are found in temperate grasslands, have a thick topsoil rich in organic content. Aridisols, on the
other hand, are dry soils that contain calcium carbonate and are found in deserts. Each soil order is further divided into suborders.
See USDA's The Twelve Orders of Soil Taxonomy and The Twelve Soil Orders from the University of Idaho for more details.
Contributiors and Attributions Modified by Kyle Whittinghill (University of Pittsburgh) and Melissa Ha from the following sources: The Soil from General Biology by OpenStax (CC-BY) Soil Profiles and Processes and Soil-Plant Interactions from Environmental Biology by Matthew R. Fisher (licensed under CCBY) Essentials of Environmental Science by Kamala Dorsner is licensed under CC BY 4.0.
Soil from Biology by John W. Kimball (licensed under CC-BY) 2.3: Soils is shared under a CC BY-NC-SA license and was authored, remixed, and/or curated by LibreTexts. 31.2: The Soil by OpenStax is licensed CC BY 4.0.
2.4: The Atmosphere The Atmosphere The atmosphere, the gaseous layer that surrounds the earth, formed over four billion years ago and is held in place by the attractive forces of gravity. During the evolution of the solid earth, volcanic eruptions released gases into the developing atmosphere. Assuming the outgasing was similar to that of modern volcanoes, the gases released included: water vapor (H2O), carbon monoxide (CO), carbon dioxide (CO2), hydrochloric acid (HCl), methane (CH4), ammonia (NH3), nitrogen (N2) and sulfur gases. The atmosphere was reducing because there was no free oxygen. Most of the hydrogen and helium that outgassed would have eventually escaped into outer space due to the inability of the earth's gravity to hold on to their small masses. There may have also been significant contributions of volatiles from the massive meteoritic bombardments known to have occurred early in the earth's history. Water vapor in the atmosphere condensed and rained down, eventually forming lakes and oceans. The oceans provided homes for the earliest organisms which were probably similar to cyanobacteria. Oxygen was released into the atmosphere by these early organisms, and carbon became sequestered in sedimentary rocks. This led to our current oxidizing atmosphere, which is mostly comprised of nitrogen (roughly 78 percent) and oxygen (roughly 21 percent) with 0.9% argon (Ar) and 0.04% carbon dioxide (CO2). The atmosphere also contains highly variable concentrations of water vapor, which can range from only 0.01% in frigid winter air in the Arctic to 5% in warm, humid, tropical air. The atmosphere also contains several gases in trace amounts, such as helium, neon, methane and nitrous oxide. One very important trace gas is ozone (O3), which absorbs harmful UV radiation from the sun. The earth's atmosphere extends outward to about 1,000 kilometers where it transitions to interplanetary space. However, most of the mass of the atmosphere (greater than 99 percent) is located within the first 40 kilometers. On average, the total weight of the atmospheric mass exerts a pressure at sea level of around 1.0 × 105 pascals (Pa; or one atmosphere), which is equivalent to 1.0 kg per cm2. The density of the atmospheric mass is much greater close to the surface and decreases rapidly with increasing altitude. The vertical temperature profile of the atmosphere is variable and depends upon the types of radiation that affect each atmospheric layer. This, in turn, depends upon the chemical composition of that layer (mostly involving trace gases). Based on these factors, the atmosphere can be divided into four distinct layers: the troposphere, stratosphere, mesosphere, and thermosphere (Figure ). The boundaries of the layers are inexact because they may vary over time and space. Beyond the atmosphere is outer space, an immeasurably vast region where the Earth exerts no detectable chemical or thermal influences.
Figure : The layers of the atmosphere. the troposphere is the closest to the Earth's surface (0-12 km). Next, is the stratosphere (12-50 km), mesophere (50-80 km), and thermosphere (80+ km). The outermost layer (the exosphere) is not shown. Ground-level
ozone in the troposphere is a form of air pollution, but the ozone layer in the stratosphere helps filter UV rays. "Atmospheric Layers" by GFDL is licensed under CC BY-SA 3.0. The troposphere (or lower atmosphere) contains 85-90% of the atmospheric mass and extends from the surface to an altitude of 820 km. It is thinner at high latitudes, and thicker at equatorial latitudes, but also varies seasonally, at any place being thicker during the summer than in the winter. Because convective air currents (winds) are common in the troposphere (the name troposphere means "region of mixing") it is sometimes referred to as the "weather layer." It also contains some 99 percent of the total water vapor of the atmosphere. The temperature of the troposphere is warm (roughly 17º C) near the surface of the earth. This is due to the absorption of infrared radiation from the surface by water vapor and other greenhouse gases (e.g. carbon dioxide, nitrous oxide and methane) in the troposphere. The concentration of these gases decreases with altitude, and therefore, the heating effect is greatest near the surface. The temperature in the troposphere decreases at a rate of roughly 6.5º C per kilometer of altitude. The temperature at its upper boundary is very cold (roughly -60º C). Water vapor evaporated from the earth's surface condenses in the cooler upper regions of the troposphere and falls back to the surface as rain. Dust and pollutants injected into the troposphere become well mixed in the layer, but are eventually washed out by rainfall. The troposphere is therefore self cleaning. A narrow zone at the top of the troposphere is called the tropopause. It effectively separates the underlying troposphere and the overlying stratosphere. The temperature in the tropopause is relatively constant. Strong eastward winds, known as the jet stream, also occur here. The stratosphere extends from the troposphere to as high as about 50 km, depending on the season and latitude. Within the stratosphere there are few convective air currents. The temperature profile of the stratosphere is quite different from that of the troposphere. The temperature remains relatively constant up to roughly 25 kilometers and then gradually increases up to the upper boundary of the layer. The amount of water vapor in the stratosphere is very low, so it is not an important factor in the temperature regulation of the layer. Instead, it is ozone (O3) that causes the observed temperature inversion. Most of the ozone in the atmosphere is contained in a layer of the stratosphere from roughly 20 to 30 kilometers. This ozone layer absorbs solar energy in the form of ultraviolet radiation (UV), and the energy is ultimately dissipated as heat in the stratosphere. This heat leads to the rise in temperature. Stratospheric ozone is also very important for living organisms on the surface of the earth as it protects them by absorbing most of the harmful UV radiation from the sun. The upper boundary of the stratosphere is known as the stratopause, which is marked by a sudden decrease in temperature. The third layer in the earth's atmosphere is called the mesosphere. It extends from the stratopause (about 50 kilometers) to roughly 85 kilometers above the earth's surface. Because the mesosphere has negligible amounts of water vapor and ozone for generating heat, the temperature drops across this layer. It is warmed from the bottom by the stratosphere. The air is very thin in this region with a density about 1/1000 that of the surface. With increasing altitude this layer becomes increasingly dominated by lighter gases, and in the outer reaches, the remaining gases become stratified by molecular weight. The fourth layer, the thermosphere, extends outward from about 85 kilometers to about 600 kilometers. Its upper boundary is ill defined. The temperature in the thermosphere increases with altitude, up to 1500º C or more. The high temperatures are the result of absorption of intense solar radiation by the last remaining oxygen molecules. The temperature can vary substantially depending upon the level of solar activity. The lower region of the thermosphere (up to about 550 kilometers) is also known as the ionosphere. Because of the high temperatures in this region, gas particles become ionized. The ionosphere is important because it reflects radio waves from the earth's surface, allowing long-distance radio communication. The visual atmospheric phenomenon known as the northern lights also occurs in this region. The outer region of the atmosphere is known as the exosphere. The exosphere represents the final transition between the atmosphere and interplanetary space. It extends about 1000 kilometers and contains mainly helium and hydrogen. Most satellites operate in this region. Climate and Weather both Happen in the Atmosphere A common misconception about global climate change is that a specific weather event occurring in a particular region (for example, a very cool week in June in central Indiana) is evidence of global climate change. However, a cold week in June is a weather-related event and not a climate-related one. These misconceptions often arise because of confusion over the terms climate and weather. Both long-term climate patterns and short-term weather patterns have to do with conditions in the atmosphere.
Climate refers to the long-term, predictable atmospheric conditions of a specific area. The climate of a biome is characterized by having consistent temperature and annual rainfall ranges. Climate does not address the amount of rain that fell on one particular day in a biome or the colder-than-average temperatures that occurred on one day. Climate scientists often average over decades to determine climate in a given location. Climate change refers to any significant change in the measures of climate lasting for an extended period of time. In other words, climate change includes major changes in temperature, precipitation, or wind patterns, among other effects, that occur over several decades or longer. More details about climate can be found in the sections of this book on "Earth's Energy Balance," "What Makes the Climate Change?," and "Past Climate Change." In contrast, weather refers to the conditions of the atmosphere during a short period of time. Weather forecasts are usually made for 48-hour cycles. Long-range weather forecasts are available but can be unreliable. For more on how Atmospheric and Oceanic Circulation affect weather patterns across the globe, see the section of this book on "Atmospheric and Oceanic Circulation." We won't be covering short-term weather patterns in this text, but there are many good resources for learning more about meteorology including Roland Stull's book Practical Meteorology. To better understand the difference between climate and weather, imagine that you are planning an outdoor event in northern Wisconsin. You would be thinking about climate when you plan the event in the summer rather than the winter because you have long-term knowledge that any given Saturday in the months of May to August would be a better choice for an outdoor event in Wisconsin than any given Saturday in January. However, you cannot determine the specific day that the event should be held on because it is difficult to accurately predict the weather on a specific day. Climate can be considered "average" weather. Contributors and Attributions Modified by Kyle Whittinghill, Melissa Ha, and Rachel Schleiger from the following sources: Connie Rye (East Mississippi Community College), Robert Wise (University of Wisconsin, Oshkosh), Vladimir Jurukovski (Suffolk County Community College), Jean DeSaix (University of North Carolina at Chapel Hill), Jung Choi (Georgia Institute of Technology), Yael Avissar (Rhode Island College) among other contributing authors. Original content by OpenStax (CC BY 4.0; Download for free at http://cnx.org/contents/185cbf87-c72...f21b5eabd@9.87). Kyle Whittinghill (University of Pittsburgh) Contributed by Melissa Ha and Rachel Schleiger Faculty (Biological Sciences) at Yuba College & Butte College Climate and the Effects of Global Climate Change from General Biology by OpenStax (licensed under CC-BY) Climate Change from Environmental Biology by Matthew R. Fisher (licensed under CC-BY) Carbon Cycle from Biology by John W. Kimball (licensed under CC-BY) The Physical World from Environmental Science: A Canadian Perspective by Bill Freedman Essentials of Environmental Science by Kamala Dorsner is licensed under CC BY 4.0 The Atmosphere from AP Environmental Science by University of California College Prep accessed through CNX. 2.4: The Atmosphere is shared under a CC BY-NC-SA license and was authored, remixed, and/or curated by LibreTexts. 44.5: Climate and the Effects of Global Climate Change by OpenStax is licensed CC BY 4.0. 10.2: Ozone Depletion by Matthew R. Fisher is licensed CC BY 4.0. Original source: https://openoregon.pressbooks.pub/envirobiology.
2.5: Earth's Energy Balance Earth's Temperature is a Balancing Act Earth's temperature depends on the balance between energy entering and leaving the planet. When incoming energy from the sun is absorbed, Earth warms. When the sun's energy is reflected back into space, Earth avoids warming. When energy is released from Earth into space, the planet cools. Many factors, both natural and human, can cause changes in Earth's energy balance, including: Changes in the greenhouse effect, which affects the amount of heat retained by Earth's atmosphere; Variations in the sun's energy reaching Earth; Changes in the reflectivity of Earth's atmosphere and surface. The balance of incoming and outgoing heat on Earth is referred to as its energy budget. As with any budget, to maintain constant conditions the budget must be balanced so that the incoming heat equals the outgoing heat. The energy budget of Earth appears below (Figure ).
Figure : Earth's energy budget. Of all of the solar radiation reaching Earth, 30% is reflected back to space and 70% is absorbed by the Earth (47%) and atmosphere (23%). The heat absorbed by the land and oceans is exchanged with the atmosphere through conduction, radiation, and latent heat (phase change). The heat absorbed by the atmosphere is eventually radiated back into
space. Source: Paul Webb, Roger Williams University. Solar radiation from the sun is composed of mostly ultraviolet (UV), visible light, and infrared (IR) radiation (Figure ). Components of solar radiation include parts with a shorter wavelength than visible light, like ultraviolet light, and parts of the spectrum with longer wavelengths, like IR and others (Figure ). Of all of the solar energy reaching the Earth, about 30% is reflected back into space from the atmosphere, clouds, and surface of the Earth (Figure ). Another 23% of the energy is absorbed by the water vapor, clouds, and dust in the atmosphere, where it is converted into heat.
Figure : Incoming radiation absorbed, scattered, and reflected by atmospheric gases.
Just under half (47%) of the incoming solar radiation is absorbed by the land and ocean, and this energy heats up the Earth's surface. The energy absorbed by the Earth returns to the atmosphere through three processes; conduction, radiation, and latent heat (phase change) (Figure ). Conduction is the transfer of heat through direct contact between the surface and the atmosphere. Air is a relatively poor thermal conductor (which means it is a good insulator), so conduction represents only a small part of the energy transfer between the Earth and the atmosphere; equal to about 7% of the incoming solar energy. All bodies with a temperature above absolute zero (-273o C) radiate heat in the form of longwave, infrared radiation (Figure ). The warmed Earth is no exception, and about 16% of the original solar energy is radiated from the Earth to the atmosphere (Figure ). When sunlight warms a surface such as a paved surface, a patio, or deck, the warmer surface emits more thermal radiation, which is a type of IR radiation. So, there is a conversion from visible, UV, and IR to just thermal IR. This thermal IR is what we experience as heat. Some of this radiated energy will dissipate into space, but a significant amount of heat will be absorbed by the atmosphere. This is the basis for the greenhouse effect (Figure ). In the greenhouse effect, shortwave solar radiation passes through the atmosphere and reaches the Earth's surface where it gets absorbed. When the radiation is re-emitted by the Earth, it is now in the form of long wavelength, infrared radiation, which does not easily pass through the atmosphere. Instead, this infrared radiation is absorbed by the atmosphere, particularly by the greenhouse gases such as CO2, methane, and water vapor. As a result, the atmosphere heats up. Some of the infrared radiation absorbed by the atmosphere is also re-emitted back to the planet's surface, further increasing warming. Without the greenhouse effect, the average temperature on Earth would be about -18o C, which is too cold for liquid water, and therefore life as we know it could not exist!
Figure : An explanation of the greenhouse effect. "Earth's Greenhouse Effect" by US EPA is available in the public domain.
The largest pathway for energy exchange between the land or oceans and the atmosphere is latent heat transferred through phase changes; heat released or absorbed when water moves between solid, liquid, and vapor forms. Heat must be added to liquid water to make it evaporate, and when water vapor is formed, that heat is removed from the ocean and transferred to the atmosphere along with the water vapor. When water vapor condenses into rain, that heat is then returned to the oceans. The same process happens with the formation and melting of ice. Heat is absorbed by ice when it melts, and heat is released when ice forms, and these phase changes transfer heat between the oceans and the atmosphere. To complete the energy budget, the heat that is absorbed by the atmosphere either directly from solar radiation or as a result of conduction, radiation and latent heat, is eventually radiated back into space (Figure ). Greenhouse Gasses Throughout this chapter we will be talking about the role of greenhouse gases (GHGs) in controlling the climate, so it is important to understand what greenhouse gases are and how they work. As you know, the dominant gases of the atmosphere are nitrogen (as N2) and oxygen (as O2). These gas molecules have only two atoms each and are not GHGs. Some of the other important gases of the atmosphere are water vapor (H2O), carbon dioxide (CO2), and methane (CH4). All of these have more than two atoms, and they are GHGs. (For more information on the relative amounts of different gasses in the atmosphere, see Atmosphere). All molecules vibrate at various frequencies and in various ways, and some of those vibrations take place at frequencies within the range of the infrared (IR) radiation that is emitted by Earth's surface. Gases with two atoms, such as O2, can only vibrate by stretching (back and forth), and those vibrations are much faster than the IR radiation (Figure ). Gases with three or more atoms (such as CO2) vibrate by stretching as well, but they can also vibrate in other ways, such as by bending (Figure ). Those vibrations are slower and match IR radiation frequencies. When IR radiation interacts with CO2 or with one of the other GHGs, the molecular vibrations are enhanced because there is a match between the wavelength of the IR light and the vibrational frequency of the molecule. This makes the molecule vibrate more vigorously, heating the surrounding air in the process. These molecules also emit IR radiation in
all directions, some of which reaches Earth's surface. The heating caused by the more vigorous vibrations of GHGs is the greenhouse effect.
Figure : Stretching versus bending vibrations in atmospheric gases. Image by Steven Earle is licensed under CC BY.
Differential Heating of Earth's Surface If the Earth was a flat surface facing the sun, every part of that surface would receive the same amount of incoming solar radiation. However, because the Earth is a sphere, sunlight is not equally distributed over the Earth's surface, so different regions of Earth will be heated to different degrees. This differential heating of Earth's surface occurs for a number of reasons. First, because of the curvature of Earth, sunlight only falls perpendicularly to the surface at the center of the sphere (equatorial regions). At any other point on Earth, the angle between the surface and the incoming solar radiation is less than 90o. Because of this, the same amount of incoming solar radiation will be concentrated in a smaller area at the equator, but will be spread over a much larger area at the poles (Figure ). Thus the tropics receive more intense sunlight and a greater amount of heating per unit of area than the polar regions. The angle at which sunlight strikes the Earth contributes to differential heating of the surface in an additional way. At the poles, because of the angle at which the solar energy strikes the surface, more of the light will glance off of the surface and the
atmosphere and be reflected back into space. At the equator, the direct angle with which light reaches the surface results in more of the energy being absorbed by the earth rather than being reflected (bouncing off the surface). Recall that the earth is also tilted, changing which part of the earth's surface is pointed directly at the sun during different seasons (see What Makes the Climate Change for more on Earth's Tilt).
Figure : Because of the curvature of the Earth, the same amount of sunlight will be spread out over a larger area at the poles compared to the equator. The equator therefore receives more intense sunlight, and a greater amount of heat per unit of area. "Solar
Angle of Incidence on Earth" by Thebiologyprimer is available in the public domain. Finally, the poles reflect more solar energy than other parts of the Earth because the poles have a higher albedo. The albedo refers to reflectivity of a surface, expressed as the percentage of light that reflects off a given material. Lighter surfaces are more reflective than darker surfaces (which absorb more energy), and therefore have a higher albedo. Water in the oceans or on a lake is one of the darkest surfaces, reflecting less than 10% of the incident light, while clouds and snow or ice are among the brightest surfaces, reflecting 70% to 90% of the incident light (Figure ).
Figure : Typical albedo values for Earth surfaces. Image by Steven Earle is licensed under CC BY.
At the poles, the ice, snow and cloud cover create a much higher albedo, and the poles reflect more and absorb less solar energy than the lower latitudes (Figure ). Through all of these mechanisms, the poles absorb much less solar radiation than equatorial regions, which is why the poles are cold and the tropics are very warm.
Figure : Surface of Earth with Cloud Cover Removed -- The surface of the Earth with cloud cover removed. The poles and deserts are much brighter than the oceans and forests. Source: NASA Goddard Space Flight Center Image by Reto Stöckli. Courtesy
of NASA's Earth Observatory. But there is an interesting twist to this global distribution of heat. The tropical regions actually receive more radiant heat than they emit, and the poles emit more heat than they receive (Figure ). We should therefore expect that the tropics will be getting continually warmer, while the poles become increasingly cold. Yet this is not the case; so what is happening? Rather than the heat remaining isolated near the equator, about 20% of the heat from the tropics is transported to the poles before it is emitted. This large scale transport of energy moderates the climates at both extremes. The mechanisms for this heat transfer are ocean and atmospheric circulation.
: The balance between heat gain and heat loss as a function of latitude. Excess heat received near the equator is transferred towards the poles Source: National Oceanography Centre (NOC).
One of the fundamental facts of life at Earth's mid-latitudes, is that there are significant variations in the heat we receive from the
Sun during the course of the year. The difference between seasons gets more pronounced the farther north or south from the equator
we travel, and the seasons in the Southern Hemisphere are the opposite of what we find on the northern half of Earth. The seasons
are actually caused by the 23.5° tilt of Earth's axis which continues to point the same direction in the sky throughout the year
(figure ). As Earth travels around the Sun, in June the Northern Hemisphere "leans into" the Sun and is more directly
illuminated. In December, the situation is reversed: the Southern Hemisphere leans into the Sun, and the Northern Hemisphere
leans away. In September and March, Earth leans "sideways"--neither into the Sun nor away from it--so the two hemispheres are
equally favored with sunshine. When we lean into the Sun, sunlight hits us at a more direct angle and is more effective at heating
). During the summer season, the sun also spends more time above the horizon, resulting in more
Figure : We see Earth at different seasons as it circles the Sun. In June, the Northern Hemisphere "leans into" the Sun, and those in the North experience summer and have longer days. In December, during winter in the Northern Hemisphere, the Southern Hemisphere "leans into" the Sun and is illuminated more directly. In spring and autumn, the two hemispheres receive more equal
shares of sunlight. Source: 4.2: The Seasons.
: The Sun's Rays in Summer and Winter. (a) In summer, the Sun appears high in the sky and its rays hit Earth more
directly, spreading out less. (b) In winter, the Sun is low in the sky and its rays spread out over a much wider area, becoming less
effective at heating the ground. Source: 4.2: The Seasons.
Further Reading The National Aeronautical and Space Administration (NASA) Earth Observatory website has an array of climate resources. For a more in-depth discussion of Earth's energy budget, go to http://earthobservatory.nasa.gov/Features/EnergyBalance/ Contributors and Attributions This page was modified from the following sources by Kyle Whittinghill (University of Pittsburgh) Connie Rye (East Mississippi Community College), Robert Wise (University of Wisconsin, Oshkosh), Vladimir Jurukovski (Suffolk County Community College), Jean DeSaix (University of North Carolina at Chapel Hill), Jung Choi (Georgia Institute of Technology), Yael Avissar (Rhode Island College) among other contributing authors. Original content by OpenStax (CC BY 4.0; Download for free at http://cnx.org/contents/185cbf87-c72...f21b5eabd@9.87). Contributed by Melissa Ha and Rachel Schleiger Faculty (Biological Sciences) at Yuba College & Butte College Contributed by Paul Webb Professor (Biology) at Rodger Williams University Contributed by Chris Johnson, Matthew D. Affolter, Paul Inkenbrandt, & Cam Mosher Faculty (Geology) at Salt Lake Community College Contributed by Steven Earle Faculty (Geology) at Vancover Island University 5.2 Climate Processes; External and Internal Controls from Sustainability: A Comprehensive Foundation by Tom Theis and Jonathan Tomkin 19.1 What Makes Climate Change from Physical Geology by Steven Earle (licensed under CC-BY) Sourced from OpenGeology from An Introduction to Geology Section 15.2 Earth's Temperature 8.1 Earth's Heat Budget from Introduction to Oceanography by Paul Webb (licensed under CC-BY) Climate and the Effects of Global Climate Change from General Biology by OpenStax (licensed under CC-BY) Climate Change from Environmental Biology by Matthew R. Fisher (licensed under CC-BY) 4.2: The Seasons from Astronomy by OpenStax (licensed under CC-BY) 2.5: Earth's Energy Balance is shared under a CC BY-NC-SA license and was authored, remixed, and/or curated by LibreTexts. 10.4: Climate Change by Matthew R. Fisher is licensed CC BY 4.0. Original source: https://openoregon.pressbooks.pub/envirobiology. 44.5: Climate and the Effects of Global Climate Change by OpenStax is licensed CC BY 4.0.
2.6: Atmospheric and Oceanic Circulation Atmospheric Circulation Differential heating of the Earth's surface results in equatorial regions receiving more heat than the poles. As air is warmed at the equator it becomes less dense and rises, while at the poles the cold air is denser and sinks. If the Earth was non-rotating, the warm air rising at the equator would reach the upper atmosphere and begin moving horizontally towards the poles. As the air reached the poles it would cool and sink, and would move over the surface of Earth back towards the equator. This would result in one large atmospheric convection cell in each hemisphere (Figure ), with air rising at the equator and sinking at the poles, and the movement of air over the Earth's surface creating the winds. On this non-rotating Earth, the prevailing winds would thus blow from the poles towards the equator in both hemispheres (Figure ).
Figure : Hypothetical atmospheric convection cells on a non-rotating Earth. Air rises at the equator and sinks at the poles, creating a single convection cell in each hemisphere. The prevailing winds moving over the Earth's surface blow from the poles towards the equator in both hemispheres (Modified by Paul Webb from globe image by Location_of_Cape_Verde_in_the_globe.svg: Eddo derivative work: Luan fala! [CC BY-SA 3.0], via Wikimedia Commons).
The non-rotating situation in Figure ( ) is of course only hypothetical, and in reality the Earth's rotation makes this atmospheric circulation a bit more complex. The paths of the winds on a rotating Earth are deflected by the Coriolis Effect. The Coriolis Effect is a result of the fact that different latitudes on Earth rotate at different speeds. This is because every point on Earth must make a complete rotation in 24 hours, but some points must travel farther, and therefore faster, to complete the rotation in the same amount of time. In 24 hours a point on the equator must complete a rotation distance equal to the circumference of the Earth, which is about 40,000 km. A point right on the poles covers no distance in that time; it just spins. So the speed of rotation at the equator is about 1600 km/hr, while at the poles the speed is 0 km/hr. Latitudes in between rotate at intermediate speeds; approximately 1400 km/hr at 30o and 800 km/hr at 60o. As objects move over the surface of the Earth they encounter regions of varying speed, which causes their path to be deflected by the Coriolis Effect. To explain the Coriolis Effect, imagine a cannon positioned at the equator and facing north. Even though the cannon appears stationary to someone on Earth, it is in fact moving east at about 1600 km/hr due to Earth's rotation. When the cannon fires the projectile travels north towards its target; but it also continues to move to the east at 1600 km/hr, the speed it had while it was still in the cannon. As the shell moves over higher latitudes, its momentum carries it eastward faster than the speed at which the ground beneath it is rotating. For example, by 30o latitude the shell is moving east at 1600 km/hr while the ground is moving east at only 1400 km/hr. Therefore, the shell gets "ahead" of its target, and will land to the east of its intended destination. From the point of view of the cannon, the path of the projectile appears to have been deflected to the right (red arrow, Figure ). Similarly, a cannon located at 60o and facing the equator will be moving east at 800 km/hr. When its shell is fired towards the equator, the shell will be moving east at 800 km/hr, but as it approaches the equator it will be moving over land that is traveling east faster than the projectile. So the projectile gets "behind" its target, and will land to the west of its destination. But from the point of view of the cannon facing the equator, the path of the shell still appears to have been deflected to the right (green arrow, Figure ). Therefore, in the Northern Hemisphere, the apparent Coriolis deflection will always be to the right.
Figure : The Coriolis Effect. Objects moving from the equator towards the poles (red arrows) move into a region of slower rotational speed and their paths are deflected "ahead" of their point of origin. Movement from high latitudes to low latitudes (green arrows) goes from a region of low speed to a region of higher rotation speed, and there is deflection "behind" their point of origin. In the Northern Hemisphere this deflection is always to the right from the point of origin, and in the Southern Hemisphere the deflection is always to the left (Modified by Paul Webb from globe image by Location_of_Cape_Verde_in_the_globe.svg: Eddo derivative work: Luan fala! [CC BY-SA 3.0], via Wikimedia Commons).
In the Southern Hemisphere the situation is reversed (Figure ). Objects moving towards the equator from the south pole are moving from low speed to high speed, so are left behind and their path is deflected to the left. Movement from the equator towards the south pole also leads to deflection to the left. In the Southern Hemisphere, the Coriolis deflection is always to the left from the point of origin. The magnitude of the Coriolis deflection is related to the difference in rotation speed between the start and end points. Between the poles and 60o latitude, the difference in rotation speed is 800 km/hr. Between the equator and 30o latitude, the difference is only 200 km/hr (Figure ). Therefore the strength of the Coriolis Effect is stronger near the poles, and weaker at the equator. Because of the rotation of the Earth and the Coriolis Effect, rather than a single atmospheric convection cell in each hemisphere, there are three major cells per hemisphere. Warm air rising at the equator cools as it moves through the upper atmosphere, and it descends at around 30o latitude. The convection cells created by rising air at the equator and sinking air at 30o are referred to as Hadley Cells, of which there is one in each hemisphere. The cold air that descends at the poles moves over the Earth's surface towards the equator, and by about 60o latitude it begins to rise, creating a Polar Cell between 60o and 90o. Between 30o and 60o lie the Ferrel Cells, composed of sinking air at 30o and rising air at 60o (Figure ). With three convection cells in each hemisphere that rotate in alternate directions, the surface winds no longer always blow from the poles towards the equator as in the non-rotating
Earth in Figure . Instead, surface winds in both hemispheres blow towards the equator between 90o and 60o latitude, and between 0o and 30o latitude. Between 30o and 60o latitude, the surface winds blow towards the poles (Figure ).
Figure : On a rotating Earth, there are three atmospheric convection cells in each hemisphere, leading to alternating bands of surface winds (red arrows) (Modified by Paul Webb from globe image by Location_of_Cape_Verde_in_the_globe.svg: Eddo derivative work: Luan fala! [CC BY-SA 3.0], via Wikimedia Commons).
The surface winds created by the atmospheric convection cells are also influenced by the Coriolis Effect as they change latitudes. The Coriolis Effect deflects the path of the winds to the right in the Northern Hemisphere and to the left in the Southern Hemisphere. Adding this deflection leads to the pattern of prevailing winds illustrated in Figure ). Between the equator and 30o latitude are the trade winds; the northeast trade winds in the Northern Hemisphere and the southeast trade winds in the Southern Hemisphere (note that winds are named based on the direction from which they originate, not where they are going). The westerlies are the dominant winds between 30o and 60o in both hemispheres, and the polar easterlies are found between 60o and the poles.
: The atmospheric circulation cells, showing direction of winds at Earth's surface. The atmosphere is thicker at the equator than at the poles. Created by Zachary Wilson for CK-12 Foundation CC BY-NC 3.0.
In between these wind bands lie regions of high and low pressure. High pressure zones occur where air is descending, while low pressure zones indicate rising air. Along the equator the rising air creates a low pressure region called the doldrums, or the Intertropical Convergence Zone (ITCZ) (convergence zone because this is where the trade winds converge). The doldrums refer to regions of low pressure around the equator. In these areas, air is rising rather than moving horizontally, so these regions commonly encounter very light winds. The lack of wind could leave sailing ships becalmed for days or weeks at a time, which was not good for the morale of the ship's crew. At 30o latitude there are high pressure zones of descending air known as the horse latitudes, or the subtropical highs. Like the doldrums the horse latitudes are also areas with light winds, this time due to descending air, which could leave ships becalmed. The term trade winds may have originally derived from the terms for "track" or "path", but the term may have become more common during European exploration and commercialization of the New World. Mariners sailing from Europe to the New World could sail south until they reached the trade winds, which would then propel their ships across the Atlantic to the Caribbean. To return to Europe, ships could sail to the northeast until they entered the westerlies, which would then steer them back to Europe. Finally, at 60o lies another low pressure region called the polar front. It should be noted that these high and low pressure zones are not fixed in place; their latitude fluctuates depending on the season, and these fluctuations have important implications for regional climates.
In addition to their role in creating the surface winds, these high and low pressure systems also influence other climatic phenomena (Figure ). Along the equator air is rising as it is warmed by solar radiation. Warm air contains more water vapor than cold air, which is why we experience humidity during the summer and not during the winter. The water content of air roughly doubles with every 10o C increase in temperature. So the air rising at the equator is warm and full of water vapor; as it rises into the upper atmosphere it cools, and the cool air can no longer hold as much water vapor, so the water condenses and forms rain. Therefore, low pressure systems are associated with precipitation, and we see wet habitats like tropical rainforests near the equator (Figure ). The ITCZ migrates slightly with the season, following the relative motion of sunlight between the Tropic of Cancer and Tropic of Capricorn. Because land areas heat more quickly than the oceans and there are more land areas in the Northern Hemisphere, during the Northern Hemisphere's summer, the ITCZ is approximately 5 degrees north of the equator, while in the winter, it shifts back and is approximately at the equator. As the ITCZ shifts, the dominant wind belts also shift slightly north in summer and south in winter, which causes the wet and dry seasons in this area.
Figure : Major global climatic regions in relation to atmospheric convection cells. Rising air and low pressure creates rain and wet environments at 0o and 60o latitudes, while high pressure, sinking air creates drier conditions at 30o and 90o latitudes.
(Modified by Paul Webb; Map by Waitak at en.Wikipedia Later version(s) were uploaded by Splette at en.Wikipedia; Sun by Inductiveload (Own work Based on File:Nuvola_apps_kweather.svg); Raincloud by Calusarul (Own work); all [CC BY-SA 3.0], via Wikimedia Commons). After rising and producing rain near the equator, the air masses move towards 30o latitude and sink back towards Earth as part of the Hadley convection cells. This air has lost most of its moisture after producing the equatorial rains, so the sinking air is dry, resulting in arid climates near 30o latitude in both hemispheres. Many of the major desert regions on Earth are located near 30o latitude, including much of Australia, the Middle East, and the Sahara Desert of Africa (Figure ). The air also becomes compressed and heats up as it sinks, absorbing any moisture from the clouds and creating clear skies. Thus high pressure systems are associated with dry weather and clear skies. This cycle of high and low pressure regions continues with the Ferrel and Polar convection cells, leading to rain and the boreal forests at 60o latitude in the Northern Hemisphere (there are no corresponding large land masses at these latitudes in the Southern Hemisphere). At the poles, descending, dry air produces little precipitation, leading to the polar desert climate. The elevation of the land also plays a role in precipitation and climactic characteristics. As moist air moves over land and encounters mountains it rises, expands, and cools because of the declining pressure and temperature. The cool air holds less water vapor, so condensation occurs and rain falls on the windward side of the mountains (Figure ). As the air passes over the mountains to the leeward side, it is now dry air, and as it sinks the pressure increases, it heats back up, any moisture revaporizes, and it creates dry, deserts regions behind the mountains (Figure ). This phenomenon is referred to as a rain shadow, and can be found in areas such as the Tibetan Plateau and Gobi Desert behind the Himalayas, Death Valley behind the Sierra Nevada mountains, and the dry San Joaquin Valley in California.
Figure : A rain shadow. Air rising over mountains cools and condenses and forms rain, leaving dry descending air and arid conditions on the other side of the mountain. Modified by Paul Webb from Thebiologyprimer, public domain via Wikimedia Commons.
Ocean Circulation Not unexpectedly, the oceans are warmest near the equator--typically 25° to 30°C--and coldest near the poles--around 0°C (Figure ). (Sea water will remain unfrozen down to about -2°C) Variations in sea-surface temperatures (SST) are related to redistribution of water by ocean currents, as we will see below. A good example of that is the plume of warm Gulf Stream water that extends across the northern Atlantic. St. John's, Newfoundland, and Brittany in France are at about the same latitude (47.5° N), but the average SST in St. John's is a frigid 3°C, while that in Brittany is a reasonably comfortable 15°C.
Figure : The global distribution of average annual sea-surface temperatures. Source: Plumbago, licensed under CC BY-SA.
Currents in the open ocean are created by wind moving across the water and by density differences related to temperature and salinity. An overview of the main ocean currents is shown in Figure . As you can see, the northern hemisphere currents form circular patterns (gyres) that rotate clockwise, while the southern hemisphere gyres are counter-clockwise. This happens for the
same reason that the water in your northern hemisphere sink rotates in a clockwise direction as it flows down the drain; this is caused by the Coriolis effect.
Figure : Overview of the main open-ocean currents. Red arrows represent warm water moving toward colder regions. Blue arrows represent cold water moving toward warmer regions. Black arrows represent currents that don't involve significant temperature changes. Source: image "Corrientes Oceanicas" by Dr. Michael Pidwirny is available in the public domain.
Because the ocean basins are not like bathroom basins, not all ocean currents behave the way we would expect. In the North Pacific, for example, the main current flows clockwise, but there is a secondary current in the area adjacent to our coast--the Alaska Current--that flows counter-clockwise, bringing relatively warm water from California, past Oregon, Washington, and B.C. to Alaska. On Canada's eastern coast, the cold Labrador Current flows south past Newfoundland, bringing a stream of icebergs past the harbour at St. John's (Figure ). This current helps to deflect the Gulf Stream toward the northeast, ensuring that Newfoundland stays cool, and western Europe stays warm.
Figure : An iceberg floating past Exploits Island on the Labrador Current. Source: image "Newfoundland Iceberg" by Shawn is licensed under CC BY-SA 2.0.
The currents shown in Figure are all surface currents, and they only involve the upper few hundred meters of the oceans. But
there is much more going on underneath. The Gulf Stream, for example, which is warm and saline, flows past Britain and Iceland
into the Norwegian Sea (where it becomes the Norwegian Current). As it cools down, it becomes denser, and because of its high
salinity, which also contributes to its density, it starts to sink beneath the surrounding water (Figure
known as North Atlantic Deep Water (NADW), and it flows to significant depth in the Atlantic as it heads back south. Meanwhile, at the southern extreme of the Atlantic, very cold water adjacent to Antarctica also sinks to the bottom to become Antarctic Bottom Water (AABW) which flows to the north, underneath the NADW.
: A depiction of the vertical movement of water along a north-south cross-section through the Atlantic basin.
Source: Steven Earle, licensed under CC BY.
The descent of the dense NADW is just one part of a global system of seawater circulation, both at surface and at depth, as
). The water that sinks in the areas of deep water formation in the Norwegian Sea and adjacent to
Antarctica moves very slowly at depth. It eventually resurfaces in the Indian Ocean between Africa and India, and in the Pacific
: The thermohaline circulation system, also known as the Global Ocean Conveyor. Source: image "A summary of the
path of thermohaline circulation" by NASA is available in the public domain.
The thermohaline circulation is critically important to the transfer of heat on Earth. It brings warm water from the tropics to the poles, and cold water from the poles to the tropics, thus keeping polar regions from getting too cold and tropical regions from getting too hot. A reduction in the rate of thermohaline circulation would lead to colder conditions and enhanced formation of sea ice at the poles. This would start a positive feedback process that could result in significant global cooling. There is compelling evidence to indicate that there were major changes in thermohaline circulation, corresponding with climate changes, during the Pleistocene Glaciation. The movement of surface currents also plays a role in the vertical movements of deeper water, mixing the upper water column. Upwelling is the process that brings deeper water to the surface, and its major significance is that it brings nutrient-rich deep water to the nutrient-deprived surface, stimulating primary production (see section 7.3). Downwelling is where surface water is forced downwards, where it may deliver oxygen to deeper water. Downwelling leads to reduced productivity, as it extends the depth of the nutrient-limited layer. For more details on upwelling, downwelling, and ocean productivity see the section of Dr. Paul Webb's Oceanography book on Currents, Upwelling and Downwelling. There are several interactions between ocean currents, ocean temperatures, and atmospheric circulation that cause cyclic global climate variability. The most well know of which is the El Niño/Southern Oscillation (ENSO) in the Pacific Ocean [also called El Niño-La Niña Cycles] associated with a band of warm ocean water that develops in the central and east-central equatorial Pacific. For more information on El Niño and La Niña, see the section of Dr. Paul Webb's Oceanography book on El Niño and La Niña. El Niño/Southern Oscillation (ENSO) is perhaps the most important ocean-atmosphere interaction phenomenon to cause cyclic global climate variability. Contributors and Attributions This page was modified from the following sources by Kyle Whittinghill (University of Pittsburgh) Contributed by Paul Webb Professor (Biology) at Rodger Williams University Contributed by Steven Earle Faculty (Geology) at Vancover Island University 18.4 Ocean Water from Physical Geology by Steven Earle (licensed under CC-BY) 8.2 Winds and the Coriolis Effect from Introduction to Oceanography by Paul Webb (licensed under CC-BY) 8.3 Winds and Climate from Introduction to Oceanography by Paul Webb (licensed under CC-BY) 9.5: Currents, Upwelling and Downwelling from Introduction to Oceanography by Paul Webb (licensed under CC-BY) 9.6: El Niño and La Niña from Introduction to Oceanography by Paul Webb (licensed under CC-BY) 10.02: Controls of Climate from Physical Geography and Natural Disasters by Adam Dastrup (licensed under CC BY-NC-SA 4.0) 2.6: Atmospheric and Oceanic Circulation is shared under a CC BY-NC-SA license and was authored, remixed, and/or curated by LibreTexts.
2.7: What Makes the Climate Change There are two parts to climate change, the first one is known as climate forcing, which is when conditions change to give the climate a little nudge in one direction or the other. The second part of climate change, and the one that typically does most of the work, is what we call a feedback. When a climate forcing changes the climate a little, a whole series of environmental changes take place, many of which either exaggerate the initial change (positive feedbacks), or suppress the change (negative feedbacks). In this section we will be discussing primarily natural climate forcing. Natural climate forcing has been going on throughout geological time. A wide range of processes has been operating at widely different time scales, from a few years to billions of years. An example of a climate-forcing mechanism is the increase in the amount of carbon dioxide (CO2) in the atmosphere that results from our use of fossil fuels. CO2 traps heat in the atmosphere and leads to climate warming. Warming changes vegetation patterns; contributes to the melting of snow, ice, and permafrost; causes sea level to rise; reduces the solubility of CO2 in sea water; and has a number of other minor effects. Most of these changes contribute to more warming. Melting of permafrost, for example, is a strong positive feedback because frozen soil contains trapped organic matter that is converted to CO2 and methane (CH4) when the soil thaws. Both these gases accumulate in the atmosphere and add to the warming effect. On the other hand, if warming causes more vegetation growth, that vegetation should absorb CO2, thus reducing the warming effect, which would be a negative feedback. Under our current conditions--a planet that still has lots of glacial ice and permafrost--most of the feedbacks that result from a warming climate are positive feedbacks and so the climate changes that we cause get naturally amplified by natural processes. Natural Climate Forcing Life Cycle of the Sun The longest-term natural forcing variation is related to the evolution of the Sun. Like most other stars of a similar mass, our Sun is evolving. For the past 4.57 billion years, its rate of nuclear fusion has been increasing, and it is now emitting about 40% more energy (as light) than it did at the beginning of geological time (Figure ). A difference of 40% is big, so it's a little surprising that the temperature on Earth has remained at a reasonable and habitable temperature for all of this time. The mechanism for that relative climate stability has been the evolution of our atmosphere from one that was dominated by CO2, and also had significant levels of CH4--both GHGs--to one with only a few hundred parts per million of CO2 and just under 1 part per million of CH4 because, over geological time, life and its metabolic processes have evolved and changed the atmosphere.
Figure : The life cycle of our Sun and of other similar stars. Source: image "Solar Life Cycle" by Oliver Beatson is available in the public domain.
Changes occurring in the sun itself can affect the intensity of the sunlight that reaches Earth's surface. The intensity of the sunlight can cause either warming (during periods of stronger solar intensity) or cooling (during periods of weaker solar intensity). The sun follows a natural 11-year cycle of small ups and downs in intensity, but the effect of these 11-year cycles on Earth's climate is small. Milankovitch Cycles Earth's orbit around the Sun is nearly circular, but like all physical systems, it has natural oscillations (Figure ). The importance of changes in the eccentricity (shape) of the Earth's orbit, tilt of the Earth, and precession (where the Earth's axis points) to Earth's climate cycles (now known as Milankovitch Cycles) was first pointed out by Yugoslavian engineer and
mathematician Milutin Milankovitch in the early 1900s. Milankovitch recognized that although the variations in the orbital cycles did not affect the total amount of insolation (light energy from the Sun) that Earth received, it did affect where on Earth that energy was strongest. Glaciations are most sensitive to the insolation received at latitudes of around 65°, and with the current configuration of continents, it would have to be 65° north (because there is almost no land at 65° south). First, the shape of the orbit changes on a regular time scale -- close to 100,000 years -- from being close to circular to being very slightly elliptical. But the circularity of the orbit is not what matters; it is the fact that as the orbit becomes more elliptical, the position of the Sun within that ellipse becomes less central or more eccentric (Figure a). Eccentricity is important because when it is high, the Earth-Sun distance varies more from season to season than it does when eccentricity is low.
Figure : The cycles of Earth's orbit and rotation. These are also called the Milankovitch Cycles. Source: Steven Earle, licensed under CC BY.
Second, Earth rotates around an axis through the North and South Poles, and that axis is at an angle to the plane of Earth's orbit around the Sun (Figure b). The angle of tilt (also known as obliquity) varies on a time scale of 41,000 years. When the angle is at its maximum (24.5°), Earth's seasonal differences are accentuated. When the angle is at its minimum (22.1°), seasonal differences are minimized. The current hypothesis is that glaciation is favored at low seasonal differences as summers would be cooler and snow would be less likely to melt and more likely to accumulate from year to year. Third, the direction in which Earth's rotational axis points also varies, on a time scale of about 20,000 years (Figure c). This variation, known as precession, means that although the North Pole is presently pointing to the star Polaris (the pole star), in 10,000 years it will point to the star Vega. The most important aspects are whether the northern hemisphere is pointing toward the Sun at its closest or farthest approach, and how eccentric the Sun's position is in Earth's orbit. Two opposing situations are when the northern hemisphere is at its farthest distance from the Sun during summer, which means cooler summers, and when the northern hemisphere is at its closest distance to the Sun during summer, which means hotter summers. Cool summers--as opposed to cold winters--are the key factor in the accumulation of glacial ice, so the scenario where the northern hemisphere is at its farthest distance from the Sun during summer is the one that promotes glaciation. This factor is greatest when eccentricity is high (so the difference between the closest point to the Sun and furthest point from the Sun is larger). For a more detailed explanation on Mikankovich Cycles see Milankovitch. As already stated, climate feedbacks are critically important in amplifying weak climate forcings into full-blown climate changes. When Milankovitch published his hypothesis in 1924, it was widely ignored, partly because it was evident to climate scientists that the forcing produced by the orbital variations was not strong enough to drive the significant climate changes of the glacial cycles. Those scientists did not recognize the power of positive feedbacks. It was not until 1973, 15 years after Milankovitch's death, that sufficiently high-resolution data were available to show that the Pleistocene glaciations were indeed driven by the orbital cycles, and it became evident that the orbital cycles were just the forcing that initiated a range of feedback mechanisms that made the climate change.
Data for tilt, eccentricity, and precession over the past 400,000 years have been used to determine the insolation levels at 65° north, as shown in Figure . Also shown in Figure are Antarctic ice-core temperatures from the same time period. The correlation between the two is clear, and it shows up in the Antarctic record because when insolation changes lead to growth of glaciers in the northern hemisphere, southern-hemisphere temperatures are also affected.
Figure : Insolation at 65° N in July compared with Antarctic ice-core temperatures. Source: Steven Earle, licensed under CC BY. Based on data from Valerie Masson-Delmotte, "EPICA Dome C Ice Core 800KYr Deuterium Data and Temperature Estimates," WDCA Contribution Series Number : 2007 -091, NOAA/NCDC Paleoclimatology Program, Boulder CO, USA. Retrieved from: NOAA and from Berger, A. and Loutre, M.F. (1991). Insolation values for the climate of the last 10 million years. Quaternary Science Reviews, 10, 297-317.
Plate Tectonic Processes Plate tectonic processes contribute to climate forcing in several different ways, and on time scales ranging from tens of millions to hundreds of millions of years. One mechanism is related to continental position. For example, we know that Gondwana (South America + Africa + Antarctica + Australia) was positioned over the South Pole between about 450 and 250 Ma, during which time there were two major glaciations (Andean-Saharan and Karoo) affecting the South polar regions and cooling the rest of the planet at the same time. Another mechanism is related to continental collisions. For example, the collision between India and Asia, which started at around 50 Ma, resulted in massive tectonic uplift. The consequent accelerated weathering of this rugged terrain consumed CO2 from the atmosphere and contributed to gradual cooling over the remainder of the Cenozoic. Changes in continental position can also lead to changes in ocean circulation, and therefore the distribution of energy from equator to the poles. For example, the opening of the Drake Passage -- due to plate-tectonic separation of South America from Antarctica -- led to the development of the Antarctic Circumpolar Current, which isolated Antarctica from the warmer water in the rest of the ocean and thus contributed to Antarctic glaciation starting at around 35 Ma.
Figure : A map showing 15 of the Earth's tectonic plates and the approximate rates and directions of plate motions. Source: image "Tectonic Plates" by USGS is available in the public domain and was adapted by Steven Earle.
Volcanic eruptions don't just involve lava flows and exploding rock fragments; various particulates and gases are also released, including carbon dioxide, water vapor, sulfur dioxide, hydrogen sulfide, hydrogen, and carbon monoxide. Volcanic eruptions can last a few days, but the solids and gases released during an eruption can influence the climate over a period of a few years, causing short-term climate changes. Generally, volcanic eruptions cool the climate. Sulphur dioxide is an aerosol that reflects incoming solar radiation and has a net cooling effect that is short-lived (a few years in most cases, as the particulates settle out of the atmosphere within a couple of years), and doesn't typically contribute to longer-term climate change. This occurred in 1783 when volcanoes in Iceland erupted and caused the release of large volumes of sulfuric oxide. This led to haze-effect cooling, a global phenomenon that occurs when dust, ash, or other suspended particles block out sunlight and trigger lower global temperatures as a result; haze-effect cooling usually extends for one or more years. In Europe and North America, haze-effect cooling produced some of the lowest average winter temperatures on record in 1783 and 1784. Volcanic CO2 emissions can contribute to climate warming but only if a greater-than-average level of volcanism is sustained over a long time (at least tens of thousands of years). It is widely believed that the catastrophic end-Permian extinction (at 250 Ma) resulted from warming initiated by the eruption of the massive Siberian Traps over a period of at least a million years. Greenhouse Gas Concentrations Since scientists cannot go back in time to directly measure climatic variables, such as average temperature and precipitation, they must instead indirectly measure temperature. Antarctic ice cores are a key example of such evidence. These ice cores are samples of polar ice obtained by means of drills that reach thousands of meters into ice sheets or high mountain glaciers. Viewing the ice cores is like traveling backwards through time; the deeper the sample, the earlier the time period. Trapped within the ice are bubbles of air and other biological evidence that can reveal temperature and carbon dioxide data. Antarctic ice cores have been collected and analyzed to indirectly estimate the temperature of the Earth over the past 400,000 years (Figure a). The 0 °C on this graph refers to the long-term average. Temperatures that are greater than 0 °C exceed Earth's long-term average temperature.
Conversely, temperatures that are less than 0 °C are less than Earth's average temperature. This figure shows that there have been periodic cycles of increasing and decreasing temperature. Before the late 1800s, the Earth has been as much as 9 °C cooler and about 3 °C warmer. Note that the graph in Figure b shows that the atmospheric concentration of carbon dioxide has also risen and fallen in periodic cycles; note the relationship between carbon dioxide concentration and temperature. Figure b shows that carbon dioxide levels in the atmosphere have historically cycled between 180 and 300 parts per million (ppm) by volume.
Figure : Ice at the Russian Vostok station in East Antarctica was laid down over the course 420,000 years and reached a depth of over 3,000 m. By measuring the amount of CO2 trapped in the ice, scientists have determined past atmospheric CO2 concentrations. Temperatures relative to modern day were determined from the amount of deuterium (an isotope of hydrogen) present.
Changes in Ocean Currents Ocean currents are important to climate, and currents also have a tendency to oscillate. Glacial ice cores show clear evidence of changes in the Gulf Stream (and other parts of the thermohaline circulation system) that affected global climate on a time scale of about 1,500 years during the last glaciation. The east-west changes in sea-surface temperature and surface pressure in the equatorial Pacific Ocean--known as the El Niño Southern Oscillation or ENSO--varies on a much shorter time scale of between two and seven years. These variations tend to garner the attention of the public because they have significant climate implications in many parts of the world. The past 65 years of ENSO index values are shown in Figure . The strongest El Niños in recent decades were in 1983 and 1998, and those were
both very warm years from a global perspective. During a strong El Niño, the equatorial Pacific sea-surface temperatures are warmer than normal and heat the atmosphere above the ocean, which leads to warmer-than-average global temperatures.
Figure (MEI)."
: Variations in the ENSO index from 1950 to early-2019. Source: Steven Earle with NOAA ("Multivariate ENSO Index
For more information on El Niño and La Niña, see the section of Dr. Paul Webb's Oceanography book on El Niño and La Niña.
Contributors and Attributions This page was modified from the following sources by Kyle Whittinghill (University of Pittsburgh). Connie Rye (East Mississippi Community College), Robert Wise (University of Wisconsin, Oshkosh), Vladimir Jurukovski (Suffolk County Community College), Jean DeSaix (University of North Carolina at Chapel Hill), Jung Choi (Georgia Institute of Technology), Yael Avissar (Rhode Island College) among other contributing authors. Original content by OpenStax (CC BY 4.0; Download for free at http://cnx.org/contents/185cbf87-c72...f21b5eabd@9.87). Contributed by Melissa Ha and Rachel Schleiger Faculty (Biological Sciences) at Yuba College & Butte College Contributed by Paul Webb Professor (Biology) at Rodger Williams University Contributed by Chris Johnson, Matthew D. Affolter, Paul Inkenbrandt, & Cam Mosher Faculty (Geology) at Salt Lake Community College Contributed by Steven Earle Faculty (Geology) at Vancover Island University 19.1 What Makes Climate Change from Physical Geology by Steven Earle (licensed under CC-BY) Sourced from OpenGeology from An Introduction to Geology Section 15.2 Earth's Temperature 8.1 Earth's Heat Budget from Introduction to Oceanography by Paul Webb (licensed under CC-BY) Climate and the Effects of Global Climate Change from General Biology by OpenStax (licensed under CC-BY) Climate Change from Environmental Biology by Matthew R. Fisher (licensed under CC-BY) Fundamentals of Plate Tectonics from Physical Geology by Steven Earle (licensed under a Creative Commons Attribution 4.0 International License)
2.7: What Makes the Climate Change is shared under a CC BY-NC-SA license and was authored, remixed, and/or curated by LibreTexts. 44.5: Climate and the Effects of Global Climate Change by OpenStax is licensed CC BY 4.0.
2.8: Past Climate Change Over Earth's history, the climate has changed a lot. For example, during the Mesozoic Era, the Age of Dinosaurs, the climate was much warmer and carbon dioxide was abundant in the atmosphere. However, throughout the Cenozoic Era (65 Million years ago to today), the climate has been gradually cooling. This section summarizes some of these major past climate changes.
Figure : Maximum extent of Laurentide Ice Sheet.
Past Glaciations Through geologic history, the climate has changed slowly over millions of years. Before the most recent Pliocene-Quaternary glaciation (Figure ), there were three other major glaciations (Deynoux et al., 2004). The oldest, known as the Huronian, occurred toward the end of the Archean-early Proterozoic (~2.5 billion years ago). The major event of that time, the great oxygenation event, is most commonly associated with the cause of that glaciation. The increased oxygen is thought to have reacted with the potent greenhouse gas methane, causing cooling (Kopp et al., 2005). The end of the Proterozoic (about 700 million years ago) had another glaciation, known as the Snowball Earth hypothesis (Hoffman et al., 1998). Glacial evidence has been interpreted in widespread rock sequences globally and even has been linked to low-latitude glaciation (Schopf & Klein, 1992). Limestone rock (usually formed in tropical marine environments) and glacial deposits (usually formed in cold climates) are often found together from this time in regions all around the world. In Utah, Antelope Island in the Great Salt Lake has interbedded limestone and glacial deposits (diamictites) interpreted to be formed by continental glaciation (Doelling et al., 1988). The idea of the controversial Snowball Earth hypothesis is that a runaway albedo effect (ice and snow reflecting solar radiation) might cause the complete freezing of land and ocean surfaces and a collapse of biological activity. The ice-covered earth would only melt when carbon dioxide from volcanoes reached high concentrations, due to the inability for carbon dioxide to enter the then-frozen ocean. Some studies estimated carbon dioxide was 350 times higher than today's concentrations (Hoffman et al., 1998). The complete freezing and the extent of the freezing has come into question (Allen & Etienne, 2008; Eyles & Januszczak, 2004). Glaciation also occurred in the Paleozoic, most notably with the Karoo Glaciation of the Pennsylvanian (323 to 300 million years ago). This also was caused by an increase of oxygen and a subsequent drop in carbon dioxide, most likely produced by the evolution and rise of land plants (Berner, 1998).
Figure : Global average surface temperature over the past 70 million years.
During the Cenozoic Era (the last 65 million years), the climate started out warm and gradually cooled to today. This warm time is called the Paleocene-Eocene Thermal Maximum and Antarctica and Greenland were ice-free during this time (Figure ). Since the Eocene, tectonic events during the Cenozoic caused persistent and significant planetary cooling. For example, the collision of the Indian Plate with the Asian Plate created the Himalaya Mountains increasing weathering and erosion rates. An increased rate of weathering of silicate minerals, especially feldspar, consumes carbon dioxide from the atmosphere and therefore reduces the greenhouse effect, resulting in long-term cooling (Kump et al., 2000). At about 40 Ma, the narrow gap between the South American Plate and the Antarctica Plate widened, resulting in the opening of the Drake Passage. This allowed for the unrestricted west-to-east flow of water around Antarctica, the Antarctic Circumpolar Current, which effectively isolated the southern ocean from the warmer waters of the Pacific, Atlantic, and Indian Oceans (Figure ). The region cooled significantly, and by 35-million-year ago (Oligocene) glaciers had started to form on Antarctica (Lagabrielle et al., 2009).
Figure : The Antarctic Circumpolar Current.
At around 15 Ma, subduction-related volcanism between Central and South America created the Isthmus of Panama that connected North and South America. This prevented water from flowing between the Pacific and Atlantic Oceans and reduced heat transfer from the tropics to the poles. This created a cooler Antarctica and larger Antarctic glaciers. The expansion of that ice sheet (on land and water) increased Earth's reflectivity (albedo), a positive feedback loop of further cooling: more reflective glacial ice, more cooling, more ice, and so on (Kuipers Munneke et al., 2011; Curry et al., 1995). By 5 million years ago (Pliocene Epoch), ice sheets had started to grow in North America and northern Europe. The most intense part of the current glaciation is the last 1 million years of the Pleistocene Epoch. The Pleistocene has significant temperature variations (through a range of almost 10°C) on time scales of 40,000 to 100,000 years, and corresponding expansion and contraction of ice sheets. These variations are attributed to subtle changes in Earth's orbital parameters called Milankovitch cycles, which are explained in more detail in the section on natural climate forcings (Milankovitch, 1930; Roe, 2006). Over the past
million years, the glaciation cycles have been approximately every 100,000 years with many glacial advances in the last 2 million years (Figure ) (Abe-Ouchi et al., 2013; Lisiecki and Raymo, 2005).
Figure : A Pliocene-Pleistocene stack of 57 globally distributed benthic 18O records. X-axis is time in thousands of years (ka) so 200 is actually 200,000. Source: Lisiecki and Raymo, 2005.
Warmer portions of climate within an ice age are called interglacials, with brief versions called interstadials. These warming upticks are related to variations in Earth's climate like Milankovitch cycles. In the last 500,000 years, there have been 5 or 6 interglacials, with the most recent belonging to our current time, the Holocene (Figure ). Two of the more recent climate swings demonstrate the complexity of the changes: the Younger Dryas and the Holocene Climatic Optimum. These events are more recent and yet have conflicting information. The Younger Dryas cooling is widely recognized in the Northern Hemisphere, though the timing of the event (about 12,000 years ago) does not appear to be equal everywhere (Carlson, 2013; Fairbanks, 1990). It also is difficult to find in the Southern Hemisphere (Kaplan et al., 2010). The Holocene Climatic Optimum is the warming around 6,000 years ago, though it was not universally warmer, and probably not as warm as current warming, and not at the same time everywhere (Ganopolski et al., 1998; Hewitt & Mitchell, 1998; He et al., 2004). Two significant temperature anomalies, or irregularities, have occurred in the last 2000 years. These are the Medieval Climate Anomaly (or the Medieval Warm Period) and the Little Ice Age. A third temperature anomaly aligns with the Industrial Era. The Medieval Climate Anomaly occurred between 900 and 1300 AD. During this time period, many climate scientists think that slightly warmer weather conditions prevailed in many parts of the world; the higher-than-average temperature changes varied between 0.10 °C and 0.20 °C above the norm. Although 0.10 °C does not seem large enough to produce any noticeable change, it did free seas of ice. Because of this warming, the Vikings were able to colonize Greenland. The Little Ice Age was a cold period that occurred between 1550 AD and 1850 AD. During this time, a slight cooling of a little less than 1 °C was observed in North America, Europe, and possibly other areas of the Earth. This 1 °C change in global temperature is a seemingly small deviation in temperature (as was observed during the Medieval Climate Anomaly); however, it also resulted in noticeable changes. Historical accounts reveal a time of exceptionally harsh winters with much snow and frost.
Proxy Indicators of Past Climates How do we know about past climates? Scientists use proxy indicators to understand past climate. A proxy indicator is a biological, chemical, or physical signature preserved in the rock, sediment, or ice record that acts like a "fingerprint" of something in the past (Weissert, 2000). Thus they are an indirect indicator of something like climate. For ancient glaciations from the Proterozoic and Paleozoic, there are rock formations of glacial sediments such as the diamictite (or tillite) of the Mineral Fork Formation in Utah. This dark rock has many fine-grained components plus some large out-sized clasts like a modern glacial till (Ojakangas & Matsch, 1980; Christie-Blick et al., 1982). For climate changes during the Cenozoic Era (the last 65 Ma), there is a detailed chemical record from the coring of deep-sea sediments as part of the Ocean Drilling Program (Figure ). Studies of deep-sea sediment use stable carbon and oxygen isotopes obtained from the shells of deep-sea benthic foraminifera that have settled on the ocean floor over millions of years. Oxygen isotopes are a proxy indicator of deep-sea temperatures and continental ice volume (Zachos et al., 2001). Sediment Cores - Stable Oxygen Isotope
Figure : Sediment core from the Greenland continental slope. Source: Hannes Grobe.
Oxygen isotopes are an indicator of past climate. The two main stable oxygen isotopes are 16O and 18O. They both occur in water (H2O) and in the calcium carbonate (CaCO3) shells of foraminifera as the oxygen component of both of those molecules. The most abundant and lighter isotope is 16O. Since it is lighter, it evaporates more easily from the ocean's surface as water vapor, which later turns to clouds and precipitation on the ocean and land.
Figure : Antarctic temperature changes during the last few glaciations compared to global ice volume. The first two curves are based on the deuterium (heavy hydrogen) record from ice cores (EPICA Community Members 2004, Petit et al. 1999). The bottom line is ice volume based on oxygen isotopes from a composite of deep-sea sediment cores (Lisiecki and Raymo 2005).
During geologic times when the climate is cooler, more of this precipitation is locked onto land in the form of glacial ice. Consider the giant ice sheets, more than a mile thick, that covered a large part of North America during the last ice age only 14,000 years ago. During glaciation, the glaciers effectively lock away more 16O, thus the ocean water and foraminifera shells become enriched in 18O. Therefore, a ratio of 18O to 16O in calcium carbonate shells of foraminifera is an indicator of past climate. The sediment cores from the Ocean Drilling Program record a continuous accumulation of sediment (Figure ).
Sediment Cores - Boron Isotopes and Acidity Boron-isotope ratios in ancient planktonic foraminifera shells in deep-sea sediment cores have been used to estimate the pH (acidity) of the ocean over the past 60 million years. Ocean acidity is a proxy for past atmospheric CO2 concentrations. In the early Cenozoic, around 60 million years ago, CO2 concentrations were over 2,000 ppm and started falling around 55 to 40 million years ago possibly due to reduced CO2 outgassing from ocean ridges, volcanoes, and metamorphic belts and increased carbon burial due to uplift of the Himalaya Mountains. By the Miocene (about 24 million years ago), CO2 levels were below 500 ppm and by 800,000 years ago CO2 levels didn't exceed 300 ppm (Pearson & Palmer, 2000; Lüthi et al., 2008). Carbon Dioxide Concentrations in Ice Cores
Figure : 19 cm long section of ice core showing 11 annual layers with summer layers (arrowed) sandwiched between darker winter layers. Source: US Army Corps of Engineers.
For the more recent Pleistocene climate, there is a more detailed and direct chemical record from coring into the Antarctic and Greenland ice sheets (Figure ). Snow accumulates on these ice sheets and creates yearly layers. Ice cores have been extracted from ice sheets covering the last 800,000 years. Oxygen isotopes are collected from these annual layers and the ratio of 18O to 16O is used to determine temperature as discussed above. In addition, the ice traps small atmospheric gas bubbles as the snow turns to ice.
Figure CSIRO.
: Antarctic ice showing hundreds of tiny trapped air bubbles from the atmosphere thousands of years ago. Source:
Small pieces of this ice are crushed and the ancient air extracted into a mass spectrometer that can detect the chemistry of the ancient atmosphere. Carbon dioxide levels are recreated from these measurements. Over the last 800,000 years, the maximum carbon dioxide concentration during warm times was about 300 ppm and the minimum during cold stretches was about 170 ppm (Pearson & Palmer, 2000; Lüthi et al., 2008; Oak Ridge National Laboratory, 2008) (Figure ). The carbon dioxide content of earth's atmosphere is currently over 400 ppm.
Figure : Composite carbon dioxide record from the last 800,000 years based on ice core data from EPICA Dome C Ice Core.
Oceanic Microfossils Microfossils, like foraminifera, diatoms, and radiolarians, can be used to interpret past climate records. In sediment cores, different species of microfossils are found in different layers. Groups of these microfossils are called assemblages. One assemblage consists of species that lived in cooler ocean water (in glacial times) and another assemblage found at a different level in the same sediment core is made of warmer water species (Cunningham et al., 1999).
Every year a tree will grow one ring with a light section and dark section (Figure
). The rings vary in width. Since trees need a
lot of water to survive, narrower rings indicate colder and drier climates. Since some trees can be several thousands years old, we
can use their rings for regional paleoclimatic reconstructions. Further, dead trees such as those used in Puebloan ruins can be used
to extend this proxy indicator, which showed long term droughts in the region and why their villages were abandoned (Figure
: Tree rings form every year. Rings that are farther apart are from wetter years and rings that are closer together are
: Summer temperature anomalies for the past 7000 years. Source: R.M.Hantemirov.
Flowering plants produce pollen grains. Pollen is distinctive when viewed under a microscope (Figure
can be preserved in lake sediments that accumulate every year. Coring of lake sediments can reveal ancient pollen. Fossil pollen
assemblages are groups of pollen from multiple species such as spruce, pine, and oak. Through time (via the sediment cores and
radiometric age-dating techniques), the pollen assemblage will change revealing the plants that lived in the area at the time. Thus
the pollen assemblages are an indicator of past climate since different plants will prefer different climates (Webb & Thompson,
1986). For example, in the Pacific Northwest east of the Cascades, a region close to the border of grasslands and forest, a study
tracked pollen over the last 125,000 years covering the last two glaciations. As shown in the figure (Fig. 2 from reference Whitlock
and Bartlein, 1997), pollen assemblages with more pine tree pollen are found during glaciations and pollen assemblages with less
pine tree pollen are found during interglacial times (Whitlock & Bartlein, 1997).
: Scanning electron microscope image of modern pollen with false color added to distinguish plant species. Source:
Dartmouth Electron Microscope Facility, Dartmouth College.
Other Proxy Indicators Paleoclimatologists study many other phenomena to understand past climates such as human historical accounts, human instrument record from the recent past, lake sediments, cave deposits of calcium carbonates, and corals (both oxygen isotopes and growth rings). Sources Abe-Ouchi, A. et al. Insolation-driven 100,000-year glacial cycles and hysteresis of ice-sheet volume. Nature 500, 190-193 (2013). Allen, P. A. & Etienne, J. L. Sedimentary challenge to Snowball Earth. Nat. Geosci. 1, 817-825 (2008). Berner, R. A. The carbon cycle and carbon dioxide over Phanerozoic time: the role of land plants. Philos. Trans. R. Soc. Lond. B Biol. Sci. 353, 75-82 (1998). Carlson, A. E. The younger Dryas climate event. (2013). Christie-Blick, N., Ojakangas, R. W. & Matsch, C. L. Upper Precambrian (Eocambrian) Mineral Fork Tillite of Utah: A Continental Glacial and Glaciomarine Sequence: Discussion and Reply. The Geological Society of America Bulletin 93, 184-187 (1982). Cunningham, W. L., Leventer, A., Andrews, J. T., Jennings, A. E. & Licht, K. J. Late Pleistocene-Holocene marine conditions in the Ross Sea, Antarctica: evidence from the diatom record. The Holocene 9, 129-139 (1999). Curry, J. A., Schramm, J. L. & Ebert, E. E. Sea Ice-Albedo Climate Feedback Mechanism. J. Clim. 8, 240-247 (1995). Deynoux, M., Miller, J. M. G. & Domack, E. W. Earth's Glacial Record. (Cambridge University Press, 2004). Doelling, H. H. et al. Geology of Antelope Island. Davis County, Utah: Utah Geological and Mineral Survey Open-File Release 144, 82 (1988). Eyles, N. & Januszczak, N. `Zipper-rift': a tectonic model for Neoproterozoic glaciations during the breakup of Rodinia after 750 Ma. Earth-Sci. Rev. (2004). Fairbanks, R. G. The age and origin of the "Younger Dryas climate event" in Greenland ice cores. Paleoceanography 5, 937-948 (1990). Ganopolski, A., Kubatzki, C., Claussen, M., Brovkin, V., V. & Petoukhov, V., V. The influence of vegetation-atmosphere-ocean interaction on climate during the mid-holocene. Science 280, 1916-1919 (1998). He, Y. et al. Asynchronous Holocene climatic change across China. Quat. Res. 61, 52-63 (2004). Hewitt, C. D. & Mitchell, J. F. B. A fully coupled GCM simulation of the climate of the mid-Holocene. Geophys. Res. Lett. 25, 361-364 (1998). Hoffman, P. F., Kaufman, A. J., Halverson, G. P. & Schrag, D. P. A neoproterozoic snowball earth. Science 281, 1342-1346 (1998).
Kaplan, M. R. et al. Glacier retreat in New Zealand during the Younger Dryas stadial. Nature 467, 194-197 (2010). Kopp, R. E., Kirschvink, J. L., Hilburn, I. A. & Nash, C. Z. The Paleoproterozoic snowball Earth: a climate disaster triggered by the evolution of oxygenic photosynthesis. Proc. Natl. Acad. Sci. U. S. A. 102, 11131-11136 (2005). Kuipers Munneke, P. et al. A new albedo parameterization for use in climate models over the Antarctic ice sheet. J. Geophys. Res. 116, D05114 (2011). Kump, L. R., Brantley, S. L. & Arthur, M. A. Chemical Weathering, Atmospheric CO2, and Climate. Annu. Rev. Earth Planet. Sci. 28, 611-667 (2000). Lagabrielle, Y., Goddéris, Y., Donnadieu, Y., Malavieille, J. & Suarez, M. The tectonic history of Drake Passage and its possible impacts on global climate. Earth Planet. Sci. Lett. 279, 197-211 (2009). Lisiecki, L. E. & Raymo, M. E. A Pliocene-Pleistocene stack of 57 globally distributed benthic 18O records. Paleoceanography 20, (2005). Lüthi, D. et al. High-resolution carbon dioxide concentration record 650,000-800,000 years before present. Nature 453, 379-382 (2008). Milankovitch, M. Mathematische klimalehre und astronomische theorie der klimaschwankungen. (1930). Oak Ridge National Laboratory. 800,000-year Ice-Core Records of Atmospheric Carbon Dioxide (CO2). (2008). Available at: http://cdiac.ornl.gov/trends/co2/ice_core_co2.html. (Accessed: 14th September 2016) Ojakangas, R. W. & Matsch, C. L. Upper Precambrian (Eocambrian) Mineral Fork Tillite of Utah: A continental glacial and glaciomarine sequence. GSA Bulletin 91, 495-501 (1980). Pearson, P. N. & Palmer, M. R. Atmospheric carbon dioxide concentrations over the past 60 million years. Nature 406, 695-699 (2000). Roe, G. In defense of Milankovitch. Geophys. Res. Lett. 33, L24703 (2006). Schopf, J. W. & Klein, C. Late Proterozoic Low-Latitude Global Glaciation: the Snowball Earth. in The Proterozoic biosphere : a multidisciplinary study (eds. Schopf, J. W. & Klein, C.) 51-52 (Cambridge University Press, 1992). Webb, T. & Thompson, W. Is vegetation in equilibrium with climate? How to interpret late-Quaternary pollen data. Vegetatio 67, 75-91 (1986). Weissert, H. Deciphering methane's fingerprint. Nature 406, 356-357 (2000). Whitlock, C. & Bartlein, P. J. Vegetation and climate change in northwest America during the past 125 kyr. Nature 388, 57-61 (1997). Zachos, J., Pagani, M., Sloan, L., Thomas, E. & Billups, K. Trends, rhythms, and aberrations in global climate 65 Ma to present. Science 292, 686-693 (2001). Contributors and Attributions Modified by Kyle Whittinghill (University of Vermont) from the following sources: Contributed by Chris Johnson, Matthew D. Affolter, Paul Inkenbrandt, & Cam Mosher Faculty (Geology) at Salt Lake Community College Sourced from OpenGeology from An Introduction to Geology Section 15.2 Prehistoric Climate Change 2.8: Past Climate Change is shared under a CC BY-NC-SA license and was authored, remixed, and/or curated by LibreTexts.