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20 Questions and Answers

Here you can find answers to many of your questions about ozone and the ozone layer.

Ozone is present only in small amounts in Earth’s atmosphere. Nevertheless, it is vital to human well-being and ecosystem health. Most ozone resides in the upper part of the atmosphere. This region, called the stratosphere, is more than 10 kilometers (6 miles) above Earth’s surface. There, about 90% of atmospheric ozone is contained in the “ozone layer,” which shields us from harmful ultraviolet radiation from the Sun.

In the mid-1970s, it was discovered that some human-produced chemicals could lead to depletion of the ozone layer. The resulting increase in ultraviolet radiation at Earth’s surface would likely increase the incidences of skin cancer and eye cataracts, and also adversely affect plants, crops, and ocean plankton.

Following the discovery of this environmental issue, researchers sought a better understanding of this threat to the ozone layer. Monitoring stations showed that the abundances of the ozone-depleting substances (ODSs) were steadily increasing in the atmosphere. These trends were linked to growing production and use of chemicals like chloro uorocarbons (CFCs) for refrigeration and air conditioning, foam blowing, and industrial cleaning. Measurements in the laboratory and in the atmosphere characterized the chemical reactions that were involved in ozone destruction. Computer models of the atmosphere employing this information were used to simulate how much ozone depletion was occurring and to predict how much more might occur in the future.

Observations of the ozone layer showed that depletion was indeed occurring. The most severe and most surprising ozone loss was discovered to be recurring in springtime over Antarctica. The loss in this region is commonly called the “ozone hole” because the ozone depletion is so large and localized. A thinning of the ozone layer also has been observed over other regions of the globe, such as the Arctic and northern and southern midlatitudes.

The work of many scientists throughout the world has provided a basis for building a broad and solid scienti c understanding of the ozone depletion process. With this understanding, we know that ozone depletion is indeed occurring and why. Most important, we know that if the most potent ODSs were to continue to be emitted and increase in the atmosphere, the result would be more depletion of the ozone layer.

In response to the prospect of increasing ozone depletion, the governments of the world crafted the 1987 United Nations Montreal Protocol as an international means to address this global issue. As a result of the broad compliance with the Protocol and its Amendments and adjustments and, of great signi cance, industry’s development of “ozone-friendly” substitutes for the now-controlled chemicals, the total global accumulation of ODSs has slowed and begun to decrease. In response, global ozone depletion is no longer increasing, and initial signs of recovery of the ozone layer have been identi ed. With continued compliance, we can expect substantial recovery of the ozone layer by the middle of the 21st century. The day the Montreal Protocol was agreed upon, 16 September, is now celebrated as the International Day for the Preservation of the Ozone Layer.

This is a story of notable achievements: discovery, understanding, decisions, actions, and veri cation. It is a story written by many: scientists, technologists, economists, legal experts, and policymakers, in which continuous dialogue has been a key ingredient. A timeline of milestones related to the science of stratospheric ozone depletion, international scientific assessments, and the Montreal Protocol is illustrated in Figure Q0-1.

Figure Q0-1 Stratospheric ozone depletion milestones

These answers are based on the information presented in the 2014 and earlier Assessment reports as well as other international scientific assessments. These reports and the answers provided here were prepared and reviewed by a large number of international scientists who are experts in different research elds related to the science of stratospheric ozone and climate.

Source: https://www.esrl.noaa.gov/csd/assessments/ozone/2014/twentyquestions2014update.pdf

Ozone is a gas that is naturally present in our atmosphere. Each ozone molecule contains three atoms of oxygen and is denoted chemically as O3. Ozone is found primarily in two regions of the atmosphere. About 10% of atmospheric ozone is in the troposphere, the region closest to Earth (from the surface to about 10–16 kilometers (6–10 miles)). The remaining ozone (about 90%) resides in the stratosphere between the top of the troposphere and about 50 kilometers (31 miles) altitude. The large amount of ozone in the stratosphere is often referred to as the “ozone layer.”

Ozone is a gas that is naturally present in our atmosphere. Ozone has the chemical formula O3 because an ozone molecule contains three oxygen atoms (see Figure Q1-1). Ozone was discovered in laboratory experiments in the mid-1800s. Ozone’s presence in the atmosphere was later discovered using chemical and optical measurement methods. The word ozone is derived from the Greek word óζειν (ozein), meaning “to smell.” Ozone has a pungent odor that allows it to be detected even at very low amounts. Ozone reacts rapidly with many chemical compounds and is explosive in concentrated amounts. Electrical discharges are generally used to produce ozone for industrial processes such as air and water puri cation and bleaching of textiles and food products.

Figure Q1-1 Ozone and oxygen

Ozone location. Most ozone (about 90%) is found in the stratosphere, which begins about 10–16 kilometers (6–10 miles) above Earth’s surface and extends up to about 50 kilometers (31 miles) altitude. The stratospheric region with the highest ozone concentration is commonly known as the “ozone layer” (see Figure Q1-2). The ozone layer extends over the entire globe with some variation in altitude and thickness. Most of the remaining ozone, about 10%, is found in the troposphere, which is the lowest region of the atmosphere, between Earth’s surface and the stratosphere.

Ozone abundance. Ozone molecules have a low relative abundance in the atmosphere. In the stratosphere near the peak concentration of the ozone layer, there are typically a few thousand ozone molecules for every billion air molecules (1 billion = 1,000 million). Most air molecules are either oxygen (O2) or nitrogen (N2). In the troposphere near Earth’s surface, ozone is even less abundant, with a typical range of 20 to 100 ozone molecules for each billion air molecules. The highest surface values result when ozone is formed in air polluted by human activities.

Figure Q1-2 Ozone in the atmosphere

As an illustration of the low relative abundance of ozone in our atmosphere, one can imagine bringing all the ozone molecules in the troposphere and stratosphere down to Earth’s surface and forming a layer of pure ozone that extends over the entire globe. The resulting layer would have an average thickness of about three millimeters (about one-tenth inch) (see Q4). Nonetheless, this extremely small fraction of the atmosphere plays a vital role in protecting life on Earth (see Q3).

Ozone is formed throughout the atmosphere in multistep chemical processes that require sunlight. In the stratosphere, the process begins with an oxygen molecule (O2) being broken apart by ultraviolet radiation from the Sun. In the lower atmosphere (troposphere), ozone is formed by a different set of chemical reactions that involve naturally occurring gases and those from pollution sources.

Stratospheric ozone. Stratospheric ozone is formed naturally by chemical reactions involving solar ultraviolet radiation (sunlight) and oxygen molecules, which make up 21% of the atmosphere. In the rst step, solar ultraviolet radiation breaks apart one oxygen molecule (O2) to produce two oxygen atoms (2 O) (see Figure Q2-1). In the second step, each of these highly reactive atoms combines with an oxygen molecule to produce an ozone molecule (O3). These reactions occur continually whenever solar ultraviolet radiation is present in the stratosphere. As a result, the largest ozone production occurs in the tropical stratosphere.

The production of stratospheric ozone is balanced by its destruction in chemical reactions. Ozone reacts continually with sunlight and a wide variety of natural and human-produced chemicals in the stratosphere. In each reaction, an ozone molecule is lost and other chemical compounds are produced. Important reactive gases that destroy ozone are hydrogen and nitrogen oxides and those containing chlorine and bromine (see Q8).

Some stratospheric ozone is regularly transported down into the troposphere and can occasionally in uence ozone amounts at Earth’s surface, particularly in remote, unpolluted regions of the globe.

Tropospheric ozone. Near Earth’s surface, ozone is produced by chemical reactions involving gases from natural and pollution sources. Ozone production reactions primarily involve hydrocarbon and nitrogen oxide gases, as well as ozone itself, and all 

require sunlight for completion. Fossil fuel combustion is a primary source of pollutant gases that lead to tropospheric ozone production. The production of ozone near the surface does not signi cantly contribute to the abundance of stratospheric ozone.
The amount of surface ozone is too small in comparison and the transport of surface air to the stratosphere is not effective enough. As in the stratosphere, ozone in the troposphere is destroyed by naturally occurring chemical reactions and by reactions involving human-produced chemicals. Tropospheric ozone can also be destroyed when ozone reacts with a variety of surfaces, such as those of soils and plants.

Figure Q2-1 Stratospheric ozone production

Balance of chemical processes. Ozone abundances in the stratosphere and troposphere are determined by the balance between chemical processes that produce and destroy ozone. The balance is determined by the amounts of reactive gases and how the rate or effectiveness of the various reactions varies with sunlight intensity, location in the atmosphere, temperature, and other factors. As atmospheric conditions change to favor ozone-producing reactions in a certain location, ozone abundances increase. Similarly, if conditions change to favor other reactions that destroy ozone, abundances decrease. The balance of production and loss reactions, combined with atmospheric air motions that transport and mix air with different ozone abundances, determines the global distribution of ozone on timescales of days to many months (see also Q4). Global ozone has decreased during the past several decades (see Q13) because the amounts of reactive gases containing chlorine and bromine have increased in the stratosphere due to human activities.

Ozone in the stratosphere absorbs a large part of the Sun’s biologically harmful ultraviolet radiation. Stratospheric ozone is considered “good” ozone because of this beneficial role. In contrast, ozone formed at Earth’s surface in excess of natural amounts is considered “bad” ozone because it is harmful to humans, plants, and animals. Natural ozone near the surface and in the lower atmosphere plays an important beneficial role in chemically removing pollutants from the atmosphere.

Good ozone. Stratospheric ozone is considered good for humans and other life forms because it absorbs ultraviolet-B (UV-B) radiation from the Sun (see Figure Q3-1). If not absorbed, UV-B radiation would reach Earth’s surface in amounts that are harmful to a variety of life forms. In humans, increased exposure to UV-B radiation increases the risks of skin cancer, cataracts, and a suppressed immune system. UV-B radiation exposure before adulthood and cumulative exposure are both important health risk factors. Excessive UV-B exposure also can damage terrestrial plant life, single-cell organisms, and aquatic ecosystems. Other UV radiation, UV-A, which is not absorbed signi cantly by ozone, causes premature aging of the skin.

Protecting good ozone. In the mid-1970s, it was discovered that gases containing chlorine and bromine atoms released by human activities could cause stratospheric ozone depletion (see Q6 and Q7). These gases, referred to as halogen source gases, and as ozone-depleting substances (ODSs), chemically release their chlorine and bromine atoms after they reach the stratosphere. Ozone depletion increases surface UV-B radiation above naturally occurring amounts. International efforts have been successful in protecting the ozone layer through controls on ODS production and consumption (see Q15 and Q16).

Bad ozone. Ozone near Earth’s surface in excess of natural amounts is considered bad ozone. It is formed by reactions involving human-made pollutant gases. Increasing surface ozone above natural levels is harmful to humans, plants, and other living systems because ozone reacts strongly to destroy or alter many biological molecules. High ozone exposure caused by air pollution reduces crop yields and forest growth. In humans, exposure to high levels of ozone can reduce lung capacity; cause chest pains, throat irritation, and coughing; and worsen pre-existing health conditions related to the heart and lungs. In addition, increases in tropospheric ozone lead to a warming of Earth’s surface because ozone is a greenhouse gas (see Q18). The negative effects of excess tropospheric ozone contrast sharply with the protection from harmful UV-B radiation afforded by an abundance of stratospheric ozone.

Reducing bad ozone. Limiting the emission of certain common pollutants reduces the production of excess ozone in the air surrounding humans, plants, and animals. Natural emissions from the biosphere, mainly from trees, also participate in reactions that produce ozone. Major sources of pollutants include large cities where fossil fuel consumption and industrial activities are greatest. Many programs around the globe have already been successful in reducing or limiting the emission of pollutants that cause production of excess ozone near Earth’s surface.

Figure Q3-1 UV protection by the ozone layer

Natural ozone. In the absence of human activities, ozone would still be present near Earth’s surface and throughout the troposphere and stratosphere because ozone is a natural component of the clean atmosphere. Ozone plays important roles in the atmosphere beyond absorbing UV radiation. For example, ozone initiates the chemical removal of many pollutants, such as carbon monoxide (CO) and nitrogen oxides (NOx), as well as some greenhouse gases, such as methane (CH4). In addition, the absorption of UV-B radiation by ozone is a natural source of heat in the stratosphere, causing temperatures to increase with altitude. Stratospheric temperatures affect the balance of ozone production and destruction processes (see Q2) and air motions that redistribute ozone throughout the stratosphere (see Q4).

The distribution of total ozone over Earth varies with location on timescales that range from daily to seasonal. The variations are caused by large-scale movements of stratospheric air and the chemical production and destruction of ozone. Total ozone is generally lowest at the equator and highest in polar regions.

Total ozone. Total ozone at any location on the globe is de ned as the sum of all the ozone in the atmosphere directly above that location. Most ozone resides in
the stratospheric ozone layer and a small percentage (about 10%) is distributed throughout the troposphere (see Q1). Total ozone values are often reported in Dobson units denoted as “DU.” Typical values vary between 200 and 500 DU over the globe (see Figure Q4-1). The ozone molecules required for total ozone to be 500 DU around the globe would form a layer of pure ozone gas at Earth’s surface having a thickness of only 5 millimeters (0.2 inches) (see Q1).

Global distribution. Total ozone varies strongly with latitude over the globe, with the largest values occurring at middle and high latitudes during all seasons (see Figure Q4-1). This is the result of the large-scale circulation of air in the stratosphere that slowly transports ozone from the tropics, where ozone production from solar ultraviolet radiation is highest, toward the poles. Ozone accumulates at middle and high latitudes, increasing the thickness (or vertical extent) of the ozone layer and, at the same time, total ozone. In contrast, the values of total ozone are the lowest in the tropics in all seasons (except for those within the Antarctic ozone hole) because the thickness of the ozone layer is smallest there.

Seasonal distribution. Total ozone also varies with season (see Figure Q4-1), using two-week averages of ozone taken from 2009 satellite observations. March and September plots represent the early spring and fall seasons in the Northern and Southern Hemispheres. June and December plots similarly represent the early summer and winter seasons. Total ozone shows a maximum at high latitudes during spring as a result of increased transport of ozone from its source region in the tropics toward the polar regions during late fall and winter. This ozone transport is much weaker during the summer and early fall periods and is weaker overall in the Southern Hemisphere. An important feature of seasonal ozone changes is the natural chemical destruction that occurs when daylight is continuous in the summer polar stratosphere, which causes total ozone to decrease gradually toward its lowest values in early fall.

This natural seasonal cycle can be observed clearly in the Northern Hemisphere (see Figure Q4-1), with increasing values in Arctic total ozone during winter, a clear maximum in spring, and decreasing values from summer to fall. In the Antarctic, however, a pronounced minimum in total ozone is observed during spring. The minimum is a consequence of the “ozone hole,” which describes the widespread chemical destruction of ozone by ozone-depleting substances (see Q6 and Q11) in spring. In the late 1970s, before the ozone hole appeared each year, much higher ozone values were found in the Antarctic spring (see Q11). Now, the lowest values of total ozone across the globe and all seasons are found every late winter and spring in the Antarctic (see Figure Q4-1). After spring, these low values disappear from total ozone maps as polar air mixes with lower-latitude air containing much higher ozone values.

In the tropics, the total ozone changes through the progression of the seasons are much smaller than in the polar regions. This is because seasonal changes in both sunlight and ozone transport are smaller in the tropics than those in the polar regions.

Figure Q4-1 Total ozone

Natural variations. Total ozone varies strongly with latitude and longitude as seen within seasonal plots (see Figure Q4-1). These patterns, which change on daily to weekly timescales, come about for two reasons. First, natural air motions mix and blend air between regions of the stratosphere that have high ozone values and those that have low ozone values. Tropospheric weather systems can temporarily change the thickness of the ozone layer in a region, and thereby change total ozone. The geographical variation in these air motions in turn causes variations in the distribution of total ozone.

Second, ozone variations occur as a result of changes in the balance of chemical production and loss processes as air moves to and from different locations over the globe. This balance, for example, is very sensitive to the amount of sunlight in a region.

There is a good understanding of how chemistry and air motions work together to cause the observed large-scale features in total ozone (see Figure Q4-1). Ozone changes are routinely monitored by a large group of investigators using satellite, airborne, and ground-based instruments. The continued analysis of these observations provides an important basis to quantify the contribution of human activities to ozone depletion.

The amount of ozone in the atmosphere is measured by instruments on the ground and carried aloft on balloons, aircraft, and satellites. Some instruments measure ozone locally by continuously drawing air samples into a small detection chamber. Other instruments measure ozone remotely over long distances by using ozone’s unique optical absorption or emission properties.

The abundance of ozone in the atmosphere is measured by a variety of techniques (see Figure Q5-1). The techniques make use of ozone’s unique optical and chemical properties. There are two principal categories of measurement techniques: local and remote. Ozone measurements by these techniques have been essential in monitoring changes in the ozone layer and in developing our understanding of the processes that control ozone abundances.

Local measurements. Local measurements of atmospheric ozone abundance are those that require air to be drawn directly into an instrument. Once inside an instrument’s detection chamber, ozone is measured by its absorption of ultraviolet (UV) light or by the electrical current or light produced in a chemical reaction involving ozone. The last approach is used in the construction of “ozonesondes,” which are lightweight, ozone-measuring modules suitable for launching on small balloons. The balloons ascend far enough in the atmosphere to measure ozone in the stratospheric ozone layer. Ozonesondes are launched regularly at many locations around the world. Local ozone-measuring instruments using optical or chemical detection schemes are also used routinely on research aircraft to measure the distribution of ozone in the troposphere and lower stratosphere. High-altitude research aircraft can reach the ozone layer at most locations over the globe and can reach highest into the layer at high latitudes. Ozone measurements are also being made routinely on some commercial aircraft flights.

Figure Q5-1 Ozone measurements

Remote measurements. Remote measurements of total ozone amounts and the altitude distributions of ozone are obtained by detecting ozone at large distances from the instrument. Most remote measurements of ozone rely on its unique absorption of UV radiation. Sources of UV radiation that can be used are the Sun, lasers, and starlight. For example, satellites use the absorption of solar UV radiation by the atmosphere or the absorption of sunlight scattered from the surface of Earth to measure ozone over nearly the entire globe on a daily basis. Lidar instruments, which measure backscattered laser light, are routinely deployed at ground sites and on research aircraft to detect ozone over a distance of many kilometers along the laser light path. A network of ground-based detectors measures ozone by detecting small changes in the amount of the Sun’s UV radiation that reaches Earth’s surface. Other instruments measure ozone using its absorption of infrared or visible radiation or its emission of microwave or infrared radiation at different altitudes in the atmosphere, obtaining information on the vertical distribution of ozone. Emission measurements have the advantage of providing remote ozone measurements at night, which is particularly valuable for sampling polar regions during winter, when there is continuous darkness.

Global Ozone Network

The first instrument for routinely monitoring total ozone was developed by Gordon M.B. Dobson in the United Kingdom in the 1920s. The instrument, called a Féry spectrometer, made its measurements by examining the wavelength spectrum of solar ultraviolet radiation (sunlight) using a photographic plate. A small network of instruments distributed around Europe allowed Dobson to make important discoveries about how total ozone varies with location and time. In the 1930s a new instrument was developed by Dobson, now called a Dobson spectrophotometer, which precisely measures the intensity of sunlight at two ultraviolet wavelengths: one that is strongly absorbed by ozone and one that is weakly absorbed. The difference in light intensity at the two wavelengths provides a measure of total ozone above the instrument location.

A global network of total-ozone observing stations was established in 1957 as part of the International Geophysical Year. Today, there are about 100 sites located around the world ranging from South Pole, Antarctica (90°S), to Ellesmere Island, Canada (83°N), that routinely measure total ozone. The accuracy of these observations is maintained by regular instrument calibrations and intercomparisons. Data from the network have been essential for understanding the effects of chlorofluorocarbons (CFCs) and other ozone-depleting substances on the global ozone layer, starting before the launch of space-based ozone-measuring instruments and continuing to the present day. Ground-based instruments with excellent long-term stability and accuracy are now routinely used to help calibrate space-based observations of total ozone.

Pioneering scientists have traditionally been honored by having units of measure named after them. Accordingly, the unit of measure for total ozone is called the “Dobson unit” (see Q4).

The initial step in the depletion of stratospheric ozone by human activities is the emission, at Earth’s surface, of gases containing chlorine and bromine. Most of these gases accumulate in the lower atmosphere because they are unreactive and do not dissolve readily in rain or snow. Natural air motions transport these accumulated gases to the stratosphere, where they are converted to more reactive gases. Some of these gases then participate in reactions that destroy ozone. Finally, when air returns to the lower atmosphere, these reactive chlorine and bromine gases are removed from Earth’s atmosphere by rain and snow.

Emission, accumulation, and transport. The principal steps in stratospheric ozone depletion caused by human activities are shown in Figure Q6-1. The process begins with the emission, at Earth’s surface, of source gases containing the halogens chlorine and bromine (see Q7). The halogen source gases, often referred to as ozone-depleting substances (ODSs), include manufactured chemicals released to the atmosphere in a variety of applications, such as refrigeration, air conditioning, and foam blowing. Chlorofluorocarbons (CFCs) are an important example of chlorine-containing gases. Emitted source gases accumulate in the lower atmosphere (troposphere) and are transported to the stratosphere by natural air motions. The accumulation occurs because most source gases are highly unreactive in the lower atmosphere. Small amounts of these gases dissolve in ocean waters. The low reactivity of these manufactured halogenated gases is one property that makes them well suited for specialized applications such as refrigeration.

Some halogen gases are emitted in substantial quantities from natural sources (see Q7). These emissions also accumulate in the troposphere, are transported to the stratosphere, and participate in ozone destruction reactions. These naturally emitted gases are part of the natural balance of ozone production and destruction that predates the large release of manufactured halogenated gases.

Figure Q6-1 Principle steps in stratospheric ozone depletion

Conversion, reaction, and removal. Halogen source gases do not react directly with ozone. Once in the stratosphere, halogen source gases are chemically converted
to reactive halogen gases by ultraviolet radiation from the Sun (see Q8). The rate of conversion is related to the atmospheric lifetime of a gas (see Q7). Gases with longer lifetimes have slower conversion rates and survive longer in the atmosphere after emission. Lifetimes of the principal ODSs vary from 1 to 100 years (see Q7). Emitted gas molecules with atmospheric lifetimes greater than a few years circulate between the troposphere and stratosphere multiple times, on average, before conversion occurs.

The reactive gases formed from halogen source gases react chemically to destroy ozone in the stratosphere (see Q9). The average depletion of total ozone attributed
to reactive gases is smallest in the tropics and largest at high latitudes (see Q13). In polar regions, surface reactions that occur at low temperatures on polar stratospheric clouds (PSCs) greatly increase the abundance of the most reactive chlorine gas, chlorine monoxide (ClO) (see Q10). This results in substantial ozone destruction in polar regions in late winter and early spring (see Q11 and Q12).

After a few years, air in the stratosphere returns to the troposphere, bringing along reactive halogen gases. These reactive halogen gases are then removed from the atmosphere by rain and other precipitation or deposited on Earth’s land or ocean surfaces. This removal brings to an end the destruction of ozone by chlorine and bromine atoms that were first released to the atmosphere as components of halogen source gas molecules.

Tropospheric conversion. Halogen source gases with short lifetimes (less than 1 year) undergo signi cant chemical conversion in the troposphere, producing reactive halogen gases and other compounds. Source gas molecules that are not converted are transported to the stratosphere. Only small portions of reactive halogen gases produced in the troposphere are transported to the stratosphere because most are removed by precipitation. Important examples of halogen gases that undergo some tropospheric removal are the hydrochlorofluorocarbons (HCFCs), methyl bromide (CH3Br) and methyl chloride (CH3Cl), and gases containing iodine (see Q7).

Understanding Stratospheric Ozone Depletion

Our understanding of stratospheric ozone depletion has been obtained through a combination of laboratory studies, computer models, and atmospheric observations. The wide variety of chemical reactions that occur in the stratosphere have been discovered and studied in laboratory studies. Chemical reactions between two gases follow well-defined physical rules. Some of these reactions occur on the surfaces of polar stratospheric clouds (PSCs) formed in the winter stratosphere. Reactions have been studied that involve many different molecules containing chlorine, bromine, uorine, and iodine and other atmospheric constituents such as carbon, oxygen, nitrogen, and hydrogen. These studies have shown that several reactions involving chlorine and bromine directly or indirectly destroy ozone in the stratosphere.

Computer models have been used to examine the combined effect of the large group of known reactions that occur in the stratosphere. These models simulate the stratosphere by including representative chemical abundances, winds, air temperatures, and the daily and seasonal changes in sunlight. These analyses show that under certain conditions chlorine and bromine react in catalytic cycles in which one chlorine or bromine atom destroys many thousands of ozone molecules. Models are also used to simulate ozone amounts observed in previous years as a strong test of our understanding of atmospheric processes and to evaluate the importance of new reactions found in laboratory studies. The responses of ozone to possible future changes in the abundances of trace gases, temperatures, and other atmospheric parameters have been extensively explored with specialized computer models (see Q20).

Atmospheric observations have shown what gases are present in different regions of the stratosphere and how their abundances vary. Gas and particle abundances have been monitored over time periods spanning a daily cycle to decades. Observations show that halogen source gases and reactive halogen gases are present in the stratosphere at the amounts required to cause observed ozone depletion. Ozone and chlorine monoxide (ClO), for example, have been observed extensively with a variety of instruments. ClO is a highly reactive gas that is involved in catalytic ozone destruction cycles throughout the stratosphere (see Q9). Instruments on the ground and on satellites, balloons, and aircraft now routinely detect ozone and ClO remotely using optical and microwave signals. High-altitude aircraft and balloon instruments are also used to detect both gases locally in the stratosphere (see Q5). The observations of ozone and reactive gases made in past decades are used extensively in comparisons with computer models in order to increase con dence in our understanding of stratospheric ozone depletion.

Certain industrial processes and consumer products result in the emission of ozone-depleting substances (ODSs) to the atmosphere. ODSs are manufactured halogen source gases that are controlled worldwide by the Montreal Protocol. These gases bring chlorine and bromine atoms to the stratosphere, where they destroy ozone in chemical reactions. Important examples are the chlorofluorocarbons (CFCs), once used in almost all refrigeration and air conditioning systems, and the halons, which were used in re extinguishers. Current ODS abundances in the atmosphere are known directly from air sample measurements.

Halogen source gases versus ODSs. Those halogen source gases emitted by human activities and controlled by the Montreal Protocol are referred to as ODSs within
the Montreal Protocol, by the media, and in the scientific literature. The Montreal Protocol now controls the global production and consumption of all ODSs (see Q15). Halogen source gases that have only natural sources are not classified as ODSs. The contributions of ODSs and natural halogen source gases to the total amount of chlorine and bromine entering the stratosphere are shown in Figure Q7-1. Total amounts of chlorine peaked in 1993, and of bromine in 1998. The difference in the timing of the peaks is explained by the phase-out schedules, how closely emissions are tied to production, and the lifetimes of the various source gases. Also shown are the contributions to total chlorine and bromine in 2012, highlighting the reductions achieved due to the Montreal Protocol controls.

Ozone-depleting substances (ODSs). ODSs are manufactured for speci c industrial uses or consumer products, most of which result in the eventual emission of these gases to the atmosphere. Total ODS emissions increased substantially from the middle to the late 20th century, reached a peak in the late 1980s, and are now in decline (see Figure Q0-1). A large fraction of the emitted ODSs reach the stratosphere, where they are converted to reactive gases containing chlorine and bromine that lead to ozone depletion. ODSs containing only carbon, chlorine, and uorine are called chloro uorocarbons, usually abbreviated as CFCs. The principal CFCs are CFC-11 (CCl3F), CFC-12 (CCl2F2), and CFC-113 (CCl2FCClF2). CFCs, along with carbon tetrachloride (CCl4) and methyl chloroform (CH3CCl3), historically have been the most important chlorine-containing halogen source gases emitted by human activities. These and other chlorine-containing ODSs have been used in many applications, including refrigeration, air conditioning, foam blowing, aerosol propellants, and cleaning of metals and electronic components. As a result of the Montreal Protocol controls, the abundances of most of these chlorine source gases have decreased between 1993 and 2012 (see Figure Q7-1). While the concentration of CFC-12 in 2012 was 2.4% higher than it was in 1993, it peaked in 2000 and has decreased since then (see Figure Q16-1). As substitute gases for CFCs, hydrochlorofluorocarbons (HCFCs) increased substantially between 1993 and 2012 (+157.5%). With new restrictions on global production in place since 2013, HCFCs are expected to reach peak values between 2020 and 2030.

Another category of ODSs contains bromine. The most important of these gases are the halons and methyl bromide (CH3Br). Halons are halocarbon gases originally developed to extinguish fires. Halons were widely used to protect large computer installations, military hardware, and commercial aircraft engines. As a consequence, halons are often released directly into the atmosphere upon use. Halon-1211 and halon-1301 are the most abundant halons emitted by human activities. Methyl bromide is used primarily as a fumigant for pest control in agriculture and disinfection of export shipping goods, but also has signi cant natural sources. As a result of the Montreal Protocol, the abundances of methyl bromide from human sources have substantially decreased between 1998 and 2012 (−61.8%; see Figure Q7-1). Halon-1211 reached peak values in 2005 and has been decreasing since then. Halon-1301 abundances, on the other hand, have increased by +19.9% since 1998 and are expected to increase further (although at a slower rate) because of continued small releases and a long atmospheric lifetime.

Natural sources of chlorine and bromine. There are a few halogen source gases present in the stratosphere that have large natural sources. These include methyl chloride (CH3Cl) and methyl bromide (CH3Br), both of which are emitted by oceanic and terrestrial ecosystems. Natural sources of these two gases contributed about 17% to the total chlorine in the stratosphere in 2012 and about 30% of the total bromine (Figure Q7-1). Very short-lived source gases containing bromine, such as bromoform (CHBr3), are also released to the atmosphere primarily from biological activity in the oceans. Only a fraction of these emissions reaches the stratosphere, because these gases are rapidly removed in the lower atmosphere. The contribution of these very short-lived gases to the total amount of bromine in the stratosphere is estimated to be about 25% in 2012, but this number has a large uncertainty. The amounts of chlorine and bromine in the stratosphere from natural sources are believed to be fairly constant and, therefore, cannot be the cause of the ozone depletion observed since the 1980s.

Figure Q7-1 Changes in halogen source gases entering the stratosphere

Other human sources of chlorine and bromine. Other chlorine- and bromine- containing gases are released to the atmosphere from human activities. Common examples are the use of chlorine-containing solvents and industrial chemicals, and the use of chlorine gases in paper production and disinfection of potable and industrial water supplies (including swimming pools). Most of these gases are very short-lived and only a small fraction of their emissions reaches the stratosphere. The contribution of very short-lived chlorinated gases from natural and human sources to total stratospheric chlorine is only about 3% and is included in Other gases in Figure Q7-1. The Montreal Protocol does not control the production and consumption of very short-lived substances (VSLS), although the concentrations of some (notably dichloromethane, CH2Cl2) have increased substantially in recent years.

Lifetimes and emissions. After emission, halogen source gases are either naturally removed from the atmosphere or undergo chemical conversion in the troposphere
or stratosphere. The time to remove or convert about 60% of a gas is often called its atmospheric lifetime. Lifetimes vary from less than 1 year to 100 years for the principal chlorine- and bromine-containing gases (see Table Q7-1). The long-lived gases are converted to other gases primarily in the stratosphere and essentially all of their original halogen content becomes available to participate in the destruction of stratospheric ozone. Gases with short lifetimes (e.g., the HCFCs, methyl bromide, methyl chloride, and the very short-lived gases) are effectively converted to other gases in the troposphere, which are then removed by rain. Therefore, only a fraction of their halogen content potentially contributes to ozone depletion in the stratosphere.

The amount of an emitted gas that is present in the atmosphere represents a balance between its emission rate and removal rate. Emission rates and atmospheric lifetimes vary greatly for the source gases, as indicated in Table Q7-1. For example, the atmospheric abundances of most of the principal CFCs and halons have decreased since 1990 in response to smaller emission rates, while those of the leading substitute gases, the hydrochlorofluorocarbons (HCFCs), continue to increase under the provisions of the Montreal Protocol (see Q16). In the coming decades, the emissions and atmospheric abundances of all controlled gases are expected to decrease under these provisions.

Ozone Depletion Potential (ODP). Halogen source gases are compared in their effectiveness to destroy stratospheric ozone using the ODP, as listed in Table Q7-1 (see Q18). A gas with a larger ODP destroys more ozone over its atmospheric lifetime. The ODP is calculated relative to CFC-11, which has an ODP de ned to be 1. The calculations, which require the use of atmospheric computer models, use as the basis of comparison the ozone depletion from an equal mass of each gas emitted to the atmosphere. Halon-1211 and halon-1301 have ODPs signi cantly larger than that of CFC-11 and most other chlorinated gases because bromine is much more effective (about 60 times) on a per-atom basis than chlorine in chemical reactions that destroy ozone. The gases with small ODP values generally have short atmospheric lifetimes or contain fewer chlorine and bromine atoms.

Table Q7-1

Fluorine and iodine. Fluorine and iodine are also halogens. Many of the source gases in Figure Q7-1 also contain fluorine in addition to chlorine or bromine. After the source gases undergo conversion in the stratosphere (see Q6), the uorine content of these gases is left in chemical forms that do not cause ozone depletion. As a consequence, halogen source gases that contain uorine and no other halogens are not classified as ODSs. An important example of these are the hydrofluorocarbons (HFCs), which are included in Table Q7-1 because they are common ODS substitute gases. HFCs have ODPs of zero and are also strong greenhouse gases (see Q18). Iodine is a component of several gases that are naturally emitted from the oceans and some human activities. Although iodine can participate in ozone destruction reactions, iodine-containing source gases all have very short lifetimes. Currently there is no evidence that any of these gases reach the stratosphere in significant amounts.

Other non-halogen gases. Other non-halogen gases that in uence stratospheric ozone abundances have also increased in the stratosphere as a result of emissions from human activities. Important examples are methane (CH4) and nitrous oxide (N2O), which react in the stratosphere to form water vapor and reactive hydrogen, and nitrogen oxides, respectively. These reactive products participate in the destruction of stratospheric ozone (see Q2). Increasing abundances of N2O and CH4, as well as CO2, are expected to signi cantly affect future stratospheric ozone through combined effects on temperature, winds, and chemistry (see Q20). Although all of these gases are part of the Kyoto Protocol (see Q15) because they are climate gases, they are not classi ed as ODSs under the Montreal Protocol. Although past emissions of ODSs still dominate global ozone depletion today, the current emissions of N2O from human activities will destroy more stratospheric ozone in the future than the current emissions of any ODS.

Emissions from human activities and natural processes represent a large source of chlorine- and bromine-containing gases that enter the stratosphere. When exposed to ultraviolet radiation from the Sun, these halogen source gases are converted to more reactive gases containing chlorine and bromine. Some reactive gases act as chemical reservoirs that convert to form the most reactive gases, namely chlorine monoxide (ClO) and bromine monoxide (BrO). The most reactive gases participate in catalytic reactions that ef ciently destroy ozone. Most volcanoes emit some reactive halogen gases that readily dissolve in water and are usually washed out of the atmosphere before they can reach the stratosphere.

Halogen-containing gases present in the stratosphere can be divided into two groups: halogen source gases and reactive halogen gases (Figure Q8-1). The source gases, which include ozone-depleting substances (ODSs), are emitted at Earth’s surface by natural processes and by human activities (see Q7). Once in the stratosphere, the halogen source gases chemically convert at different rates to form the reactive halogen gases. The conversion occurs in the stratosphere for most gases instead of the troposphere because solar ultraviolet radiation (a component of sunlight) is more intense in the stratosphere. Reactive gases containing the halogens chlorine and bromine lead to the chemical destruction of stratospheric ozone.

Figure Q8-1 Conversion of halogen source gases

Reactive halogen gases. The chemical conversion of halogen source gases, which involves solar ultraviolet radiation and other chemical reactions, produces a number of reactive halogen gases. These reactive gases contain all of the chlorine and bromine atoms originally present in the source gases. The most important reactive chlorine- and bromine-containing gases that form in the stratosphere are shown in Figure Q8-1. Throughout the stratosphere, the most abundant are typically hydrogen chloride (HCl) and chlorine nitrate (ClONO2). These two gases are considered important reservoir gases because, while they don’t react directly with ozone, they can be converted to the most reactive forms that do chemically destroy ozone. The most reactive forms are chlorine monoxide (ClO) and bromine monoxide (BrO), and chlorine and bromine atoms (Cl and Br). A large fraction of available reactive bromine is generally in the form of BrO, whereas usually only a small fraction of reactive chlorine is in the form of ClO. The special conditions that occur in the polar regions in winter cause the reservoir gases ClONO2 and HCl to undergo nearly complete conversion to ClO in reactions on polar stratospheric clouds (PSCs) (see Q10).

Reactive chlorine at midlatitudes. Reactive chlorine gases have been observed extensively in the stratosphere using both local and remote measurement techniques. The measurements from space displayed in Figure Q8-2 are representative of how the amounts of chlorine-containing gases change between the surface and the upper stratosphere at mid- to high latitudes. Available chlorine (see red line in Figure Q8-2) is the sum of chlorine contained in halogen source gases and in the reactive gases (e.g., CFC-11, ClONO2, ClO). Available chlorine is constant to within about 10% from the surface to above 50 kilometers (31 miles) altitude. In the troposphere, available chlorine is contained almost entirely in the source gases described in Figure Q7-1. At higher altitudes, the source gases become a smaller fraction of available chlorine as they are converted to the reactive chlorine gases. At the highest altitudes, available chlorine is all in the form of the reactive chlorine gases.

Figure Q8-2 Reactive chlorine gas observations

In the altitude range of the ozone layer at midlatitudes, as shown in Figure Q8-2, the reservoir gases HCl and ClONO2 account for most of the available chlorine. ClO, the most reactive gas in ozone depletion, is a small fraction of available chlorine. The low abundance of ClO limits the amount of ozone destruction that occurs outside of polar regions.

Reactive chlorine in polar regions. Reactive chlorine gases in polar regions undergo large changes between the fall and late winter. Meteorological and chemical conditions in both polar regions are now routinely observed from space in all seasons. Fall and winter conditions over the Antarctic are contrasted in Figure Q8-3 using seasonal observations made near the center of the ozone layer (about 18 km (11.3 miles) altitude) (see Figure Q12-3).

In fall (May), ozone values are high over the entire Antarctic continent and beyond. Temperatures are mid-range, HCl and nitric acid (HNO3) are high, and ClO is very low. High HCl indicates that substantial conversion of halogen source gases has occurred in the stratosphere. In the past decades, the reservoir gases HCl and ClONO2 have increased substantially in the stratosphere following increased emissions of halogen source gases. HNO3 is an abundant, naturally occurring stratospheric compound that plays major roles in stratospheric ozone chemistry by both moderating ozone destruction and condensing to form polar stratospheric clouds (PSCs). Low ClO indicates that little conversion of the reservoir gases occurs in the fall, thereby limiting catalytic ozone destruction.

By late winter (September), ozone amounts re ect substantial depletion at 18-km altitude over an area larger than the Antarctic continent. Antarctic ozone holes arise
from similar chemical destruction throughout much of the altitude range of the ozone layer (see altitude profile in Figure Q12-3). The meteorological and chemical conditions in late winter are very different from those found in fall: very low temperatures, very low HCl and HNO3, and very high ClO. Low stratospheric temperatures are characteristic of winter, when solar heating is reduced. Low HCl and high ClO re ect the conversion of the reactive halogen reservoir compounds, HCl and ClONO2, to the most reactive chlorine form, ClO. This conversion occurs selectively in winter on PSCs, which form at the very low temperatures (see Q10). Low HNO3 is indicative of its condensation to form PSCs, some of which subsequently descend to lower altitudes through gravitational settling. High ClO abundances generally cause ozone depletion to continue in the Antarctic region until mid-October (spring), when the lowest ozone values usually are observed (see Q11). As temperatures rise at the end of the winter, PSC formation is halted, ClO is converted back into the reservoir species HCl and ClONO2 (see Q10), and ozone destruction is curtailed.

Similar though less dramatic changes in meteorological and chemical conditions are also observed between fall and late winter in the Arctic, where ozone depletion is less severe than in the Antarctic.

Reactive bromine observations. Fewer measurements are available for reactive bromine gases in the lower stratosphere than for reactive chlorine, in part because of the lower abundance of bromine. The most widely observed bromine gas is bromine monoxide (BrO), which can be observed from space. Estimates of reactive bromine abundances in the stratosphere are larger than expected from the conversion of the halons and methyl bromide source gases, suggesting that the contribution of the very short-lived bromine- containing gases to reactive bromine must also be signi cant (see Q7).

Other sources. Other emission sources of reactive halogen gases exist that are associated with natural processes and human activities. Most emissions become trapped in the troposphere when dissolved in water, and ultimately are returned to Earth’s surface before they can reach the stratosphere. Volcanoes are an important episodic source of reactive halogen gases (e.g., HCl) that do not reach the stratosphere in appreciable quantities (see Q14). Other examples include reactive chlorine produced by evaporation of ocean spray. Sea salt products dissolve in water and are removed in the troposphere. Solid rocket motors, such as those used on the Space Shuttle, release reactive chlorine gases directly into the troposphere and stratosphere. The quantities emitted by the Space Shuttle were small in comparison with halogen emissions from other human activities.

Figure Q8-3 Chemical conditions in the ozone layer over Antarctica

Reactive gases containing chlorine and bromine destroy stratospheric ozone in “catalytic” cycles made up of two or more separate reactions. As a result, a single chlorine or bromine atom can destroy many thousands of ozone molecules before it leaves the stratosphere. In this way, a small amount of reactive chlorine or bromine has a large impact on the ozone layer. A special situation develops in polar regions in the late winter/early spring season, where large enhancements in the abundance of the most reactive gas, chlorine monoxide, lead to severe ozone depletion.

Stratospheric ozone is destroyed by reactions involving reactive halogen gases, which are produced in the chemical conversion of halogen source gases (see Figure Q8-1). The most reactive of these gases are chlorine monoxide (ClO), bromine monoxide (BrO), and chlorine and bromine atoms (Cl and Br). These gases participate in three principal reaction cycles that destroy ozone.

Cycle 1. Ozone destruction Cycle 1 is illustrated in Figure Q9-1. The cycle is made up of two basic reactions: Cl + O3 and ClO + O. The net result of Cycle 1 is to convert one ozone molecule and one oxygen atom into two oxygen molecules. In each cycle, chlorine acts as a catalyst because ClO and Cl react and are re-formed. In this way, one Cl atom participates in many cycles, destroying many ozone molecules. For typical stratospheric conditions at middle or low latitudes, a single chlorine atom can destroy hundreds of ozone molecules before it happens to react with another gas, breaking the catalytic cycle. During the total time of its stay in the stratosphere, a chlorine atom can thus destroy up to tens of thousands of ozone molecules.

Figure Q9-1 Ozone destruction Cycle 1

Polar Cycles 2 and 3. The abundance of ClO is greatly increased in polar regions during winter as a result of reactions on the surfaces of polar stratospheric clouds (PSCs) (see Q8 and Q10). Cycles 2 and 3 (see Figure Q9-2) become the dominant reaction mechanisms for polar ozone loss because of the high abundances of ClO and the relatively low abundance of atomic oxygen (which limits the rate of ozone loss by Cycle 1). Cycle 2 begins with the self-reaction of ClO. Cycle 3, which begins with the reaction of ClO with BrO, has two reaction pathways to produce either Cl and Br or BrCl. The net result of both cycles is to destroy two ozone molecules and create three oxygen molecules. Cycles 2 and 3 account for most of the ozone loss observed in the Arctic and Antarctic stratospheres in the late winter/early spring season (see Q11 and Q12). At high ClO abundances, the rate of polar ozone destruction can reach 2 to 3% per day in late winter/early spring.

Sunlight requirement. Sunlight is required to complete and maintain Cycles 1 through 3. Cycle 1 requires ultraviolet radiation (a component of sunlight) that is strong enough to break apart molecular oxygen into atomic oxygen. Cycle 1 is most important in the stratosphere at tropical and middle latitudes, where solar ultraviolet radiation is most intense.

Cycles 2 and 3 require visible sunlight to complete the reaction cycles and to maintain ClO abundances. In the continuous darkness of winter in the polar stratosphere, reaction Cycles 2 and 3 cannot occur. It is only in late winter/early spring when sunlight returns to the polar regions that these cycles are active. Therefore, the greatest destruction of ozone occurs in the partially to fully sunlit periods after midwinter in the polar stratosphere. The visible sunlight needed in Cycles 2 and 3 is not sufficient to form ozone because this process requires solar ultraviolet radiation (see Q1). In the late winter/early spring stratosphere, solar ultraviolet radiation is weak at low Sun angles. As a result, ozone destruction by Cycles 2 and 3 in the sunlit winter stratosphere greatly exceeds ozone production.

Figure Q9-2 Polar ozone destruction Cycles 2 and 3

Other reactions. Global ozone abundances are controlled by many reactions that both produce and destroy ozone (see Q2). Chlorine and bromine catalytic reactions are but one group of ozone destruction reactions. Reactive hydrogen and reactive nitrogen gases, for example, are involved in other catalytic ozone-destruction cycles that also occur in the stratosphere. These reactions occur naturally in the stratosphere and their importance has not been as strongly in uenced by human activities as have reactions involving halogens. The sources of reactive hydrogen and nitrogen gases are the naturally occurring gases methane (CH4) and nitrous oxide (N2O), respectively. The importance of reactive hydrogen and nitrogen gases relative to reactive halogen gases is expected to increase in the future, because the atmospheric abundances of the reactive halogen gases are decreasing as a result of the Montreal Protocol, while CH4 and N2O abundances (which are not controlled under the Montreal Protocol) are projected to increase substantially due to human activities.

Ozone-depleting substances are present throughout the stratospheric ozone layer because they are transported great distances by atmospheric air motions. The severe depletion of the Antarctic ozone layer known as the “ozone hole” occurs because of the special meteorological and chemical conditions that exist there and nowhere else on the globe. The very low winter temperatures in the Antarctic stratosphere cause polar stratospheric clouds (PSCs) to form. Special reactions that occur on PSCs, combined with the isolation of polar stratospheric air in the polar vortex, allow chlorine and bromine reactions to produce the ozone hole in Antarctic springtime.

The severe depletion of stratospheric ozone in late winter and early spring in the Antarctic is known as the “ozone hole” (see Q11). The ozone hole appears over Antarctica because meteorological and chemical conditions unique to this region increase the effectiveness of ozone destruction by reactive halogen gases (see Q8). In addition to an abundance of these reactive gases, the formation of the Antarctic ozone hole requires temperatures low enough to form polar stratospheric clouds (PSCs), isolation from air in other stratospheric regions, and sunlight.

Figure Q10-1 Arctic and Antarctic temperatures

Distribution of halogen gases. Halogen source gases that are emitted at Earth’s surface and have lifetimes longer than about 1 year are present in comparable abundances throughout the stratosphere in both hemispheres, even though most of the emissions occur in the Northern Hemisphere. The abundances are comparable because most long-lived source gases have no signi cant natural removal processes in the lower atmosphere, and because winds and convection redistribute and mix air efficiently throughout the troposphere on the timescale of weeks to months. Halogen gases (in the form of source gases and some reactive products) enter the stratosphere primarily from the tropical upper troposphere. Stratospheric air motions then transport these gases upward and toward the pole in both hemispheres.

Low polar temperatures. The severe ozone destruction that leads to the ozone hole requires low temperatures to be present over a range of stratospheric altitudes, over large geographical regions, and for extended time periods. Low temperatures are important because they allow liquid and solid PSCs to form. Reactions on the surfaces of these PSCs initiate a remarkable increase in the most reactive chlorine gas, chlorine monoxide (ClO) (see below and Q8). Stratospheric temperatures are lowest in the polar regions in winter. In the Antarctic winter, minimum daily temperatures are generally much lower and less variable than those in the Arctic winter (see Figure Q10-1). Antarctic temperatures also remain below PSC formation temperatures for much longer periods during winter. These and other meteorological differences occur because of the differences in the distribution of land, ocean, and mountains between the hemispheres at middle and high latitudes. Winter temperatures are low enough for PSCs to form somewhere in the Antarctic for nearly the entire winter (about 5 months), but only for limited periods (10–60 days) in the Arctic in most winters.

Isolated conditions. Stratospheric air in the polar regions is relatively isolated for long periods in the winter months. The isolation is provided by strong winds that encircle the poles during winter, forming a polar vortex, which prevents substantial transport and mixing of air into or out of the polar stratosphere. This circulation strengthens in winter as stratospheric temperatures decrease. Since winter temperatures are lower in the Southern than in the Northern Hemisphere polar stratosphere, the isolation of air in the vortex is much more effective in the Antarctic than in the Arctic. Once temperatures drop low enough, PSCs form within the polar vortex and induce chemical changes that are preserved by the isolation for many weeks to months.

Polar stratospheric clouds (PSCs). Reactions on the surfaces of liquid and solid PSCs can substantially increase the relative abundances of the most reactive chlorine gases. These reactions convert the reservoir forms of reactive chlorine gases, chlorine nitrate (ClONO2) and hydrogen chloride (HCl), to the most reactive form, ClO (see Figure Q8-1). ClO increases from a small fraction of available reactive chlorine to comprise nearly all chlorine that is available. With increased ClO, the catalytic cycles involving ClO and BrO become active in the chemical destruction of ozone whenever sunlight is available (see Q9).

Different types of liquid and solid PSC particles form when stratospheric temperatures fall below about −78°C (−108°F) in polar regions (see Figure Q10-1). As a result, PSCs are often found over large areas of the winter polar regions and over signi cant altitude ranges, with larger regions and for longer time periods in the Antarctic than in the Arctic. The most common type of PSC forms from nitric acid (HNO3) and water condensing on pre-existing liquid sulfuric acid-containing particles. Some of these particles freeze to form solid particles. At even lower temperatures (−85°C or −121°F), water condenses to form ice particles. PSC particles grow large enough and are numerous enough that cloud-like features can be observed from the ground under certain conditions, particularly when the Sun is near the horizon (see Figure Q10-2). PSCs are often found near mountain ranges in polar regions because the motion of air over the mountains can cause localized cooling in the stratosphere, which increases condensation of water and HNO3.

Figure Q10-2 Polar stratospheric clouds

When average temperatures begin increasing by late winter, PSCs form less frequently and conversion reactions on their surfaces produce less ClO. Without continued ClO production, ClO amounts decrease and other chemical reactions re-form the reservoirs gases, ClONO2 and HCl. When temperatures rise above PSC formation thresholds, which typically occurs by late January to early February in the Arctic or by mid-October in the Antarctic, the most intense period of ozone depletion ends.

Nitric acid and water removal. Once formed, PSC particles fall to lower altitudes because of gravity. The largest particles can descend several kilometers or more in the stratosphere within a few days during the low-temperature winter/spring period. Because PSCs often contain a significant fraction of available HNO3, their descent removes HNO3 from regions of the ozone layer. This process is called denitrification of the stratosphere. Because HNO3 is a source for nitrogen oxides (NOx) in the stratosphere, denitri cation removes the NOx available for converting the highly reactive chlorine gas ClO back into the reservoir gas ClONO2. As a result, ClO remains chemically active for a longer period, thereby increasing chemical ozone destruction. Significant denitrification occurs each winter in the Antarctic but only in occasional winters in the Arctic, because PSC formation temperatures must be sustained over an extensive altitude region and time period to lead to denitri cation (see Figure Q10-1).

Ice particles form at temperatures that are a few degrees lower than those required for PSC formation from HNO3. If ice particles grow large enough, they can fall several kilometers due to gravity. As a result, a signi cant fraction of water vapor can be removed from regions of the ozone layer over the course of a winter. This process is called dehydration of the stratosphere. Because of the very low temperatures required to form ice, dehydration is common in Antarctic but rare in Arctic winters. The removal of water vapor does not directly affect the catalytic reactions that destroy ozone. Dehydration indirectly affects ozone destruction by suppressing PSC formation later in winter, which reduces ClO production in PSC reactions.

Discovering the role of PSCs. Ground-based observations of PSCs were available many years before the role of PSCs in polar ozone destruction was recognized. The geographical and altitude extent of PSCs in both polar regions was not known fully until PSCs were observed by a satellite instrument in the late 1970s. The role of PSC particles in converting reactive chlorine gases to ClO was not understood until after the discovery of the Antarctic ozone hole in 1985. Our understanding of the chemical role of PSC particles developed from laboratory studies of their surface reactivity, computer modeling studies of polar stratospheric chemistry, and measurements that directly sampled particles and reactive chlorine gases, such as ClO, in the polar stratosphere.

The Discovery of the Antarctic Ozone Hole

The first decreases in Antarctic total ozone were observed in the early 1980s over research stations located on the Antarctic continent. The measurements were made with ground-based Dobson spectrophotometers (see box in Q5). The observations showed unusually low total ozone during the late winter/early spring months of September, October, and November. Total ozone was lower in these months compared with previous observations made as early as 1957. The early published reports came from the Japan Meteorological Agency and the British Antarctic Survey. The results became widely known to the world after three scientists from the British Antarctic Survey published their observations in the scientific journal Nature in 1985 and speculated that CFCs were the cause. Soon after, satellite measurements confirmed the spring ozone depletion and further showed that in each late winter/early spring season starting in the early 1980s, the depletion extended over a large region centered near the South Pole. The term “ozone hole” came about as a description of the very low total ozone values that encircled the Antarctic continent in satellite images for many weeks (see Q11). Currently, the formation and severity of the Antarctic ozone hole are documented each year by a combination of satellite, ground-based, and balloon observations of ozone.

Very early Antarctic ozone measurements. The rst total ozone measurements made in Antarctica with Dobson spectrophotometers occurred in the 1950s following extensive measurements in the Northern Hemisphere and Arctic region. Total ozone values found in spring were around 300 Dobson units (DU), surprisingly lower than in the Arctic spring, because the assumption then was that the two polar regions would have similar values. We now know that these Antarctic values were not anomalous; indeed they are similar to those observed there in the 1970s before the ozone hole appeared (see Figure Q11-3). We also now know that Antarctic total ozone values, in comparison with Arctic values, are systematically lower in the early spring because the polar vortex is much stronger and, therefore, much more effective in reducing the transport of ozone-rich air from midlatitudes to the pole (compare Figures Q11-3 and Q12-2).

In 1958, measurements of total ozone were made at the Dumont d’Urville station (66.7°S, 140°E) in Antarctica using a photographic plate method to analyze
solar ultraviolet radiation after it passed through the ozone layer. The reported measurements were anomalously low, reaching 110–120 DU in September and October. These values are similar to minimum ozone hole values now routinely observed over Antarctica in the same months (see Figure Q11-2). Some have speculated that these limited observations provide evidence that an ozone hole existed before ODS emissions were large enough to cause the depletion. However, analyses of the more extensive Dobson spectrophotometer measurements made at several other Antarctic locations in 1958 did not con rm the low total ozone values. These measurements indicate that the photographic plate determinations were not a reliable source of total ozone values at the Dumont d’Urville station in 1958.

Severe depletion of the Antarctic ozone layer was first reported in the mid-1980s. Antarctic ozone depletion is seasonal, occurring primarily in late winter and early spring (August–November). Peak depletion occurs in early October when ozone is often completely destroyed over a range of altitudes, thereby reducing total ozone by as much as two-thirds at some locations. This severe depletion creates the “ozone hole” apparent in images of Antarctic total ozone made using satellite observations. In most years the maximum area of the ozone hole far exceeds the size of the Antarctic continent.

The severe depletion of Antarctic ozone, known as the “ozone hole,” was first reported in the mid-1980s. The depletion is attributable to chemical destruction by reactive halogen gases that increased in the stratosphere in the latter half of the 20th century (see Q16). Conditions in the Antarctic winter and early spring stratosphere enhance ozone depletion because of (1) the long periods of extremely low temperatures, which cause polar stratospheric clouds (PSCs) to form; (2) the large abundance of reactive halogen gases produced in reactions on PSCs; and (3) the isolation of stratospheric air, which allows time for chemical destruction processes to occur. The long-term changes and severity of Antarctic ozone depletion can be seen using satellite observations of total ozone and ozone altitude profiles.

Figure Q11-1 Antarctic ozone hole

Antarctic ozone hole. The most widely used images of Antarctic ozone depletion are derived from measurements of total ozone made with satellite instruments. A map of  Antarctic early spring measurements shows a large region centered near the South Pole in which total ozone is highly depleted (see Figure Q11-1). This region has come to be called the “ozone hole” because of the near-circular contours of low ozone values in the maps. The area of the ozone hole is de ned here as the geographical region within the 220-Dobson unit (DU) contour in total ozone maps (see white line in Figure Q11-1) averaged between 21–30 September for each year. The area reached a maximum of 27 million square kilometers (about 10 million square miles) in 2006, which is nearly twice the area of the Antarctic continent (see Figure Q11-2). Minimum values of total ozone inside the ozone hole averaged in late September to mid-October are near 110 DU, which is two-thirds below normal springtime values of about 350 DU (see Figures Q11-3 and Q12-1). Low total ozone inside the ozone hole contrasts strongly with the distribution of much larger values outside the ozone hole. This common feature can be seen in Figure Q11-1, where a crescent-shaped region with values around 400 DU surrounds a large portion of the ozone hole in September 2013.

Altitude profiles of Antarctic ozone. The low total ozone values within the ozone hole are caused by nearly complete removal of ozone in the lower stratosphere. Balloonborne instruments (see Q5) demonstrate that this depletion occurs within the ozone layer, the altitude region that normally contains the highest ozone abundances. At geographic locations with the lowest total ozone values, balloon measurements show that the chemical destruction of ozone is complete over an altitude region of several kilometers. For example, in the ozone pro le over South Pole, Antarctica, on 9 October 2006 (see red line in left panel of Figure Q12-3), ozone abundances are essentially zero over the altitude region of 14 to 21 kilometers. The lowest winter temperatures and highest reactive chlorine (ClO) abundances occur in this altitude region (see Figure Q8-3). The differences in the average South Pole ozone pro les between the decades 1962–1971 and 1990–2013 in Figure Q12-3 show how reactive halogen gases have dramatically altered the ozone layer. In the 1960s, a normal ozone layer is clearly evident in the October average pro le, with a peak near 16 kilometers altitude. In the 1990–2013 average pro le, a broad minimum centered near 16 kilometers now occurs, with ozone values reduced by up to 90% of normal values.

Figure Q11-2 Antarctic ozone hole features

Long-term total ozone changes. The signi cant springtime ozone depletion represented by the ozone hole has only been observed since the early 1980s even though low winter temperatures and isolated conditions occur each year in the Antarctic stratosphere. Prior to 1980, the amounts of reactive halogen gases in the stratosphere were insuf cient to cause a signi cant signature of depletion in the face of strong interannual variations in stratospheric transport. Satellite observations can be used in multiple ways to examine how ozone depletion has changed in the Antarctic region over the last 30–40 years:

  • First, ozone hole areas displayed in Figure Q11-2 show that depletion has increased since 1980 to become fairly stable in the 1990s, 2000s, and even into the early 2010s near a value of 23 million square kilometers. The exception is the unexpected low depletion in 2002, which is explained in the box at the end of this Question. The ozone-hole area is de ned here as the geographical area inside the 220-DU contour on total ozone maps (see Figure Q11-1).

  • Second, minimum Antarctic ozone amounts displayed in Figure Q11-2 show that the severity of the depletion increased beginning around 1980 along with the ozone hole area. Fairly constant minimum values near 110 DU were observed in the 1990s, 2000s, and early 2010s, with 2002 being an exception.

  • Third, total ozone maps over the Antarctic and surrounding regions show how the ozone hole has developed over time (see Figure Q11-3). October averages of total ozone con rm the absence of an ozone hole in the 1970s and its persistent occurrence in the late 2000s and early 2010s.

  • Fourth, values of total ozone poleward of 63°S averaged for each October show how total ozone has changed when averaged over the ozone hole and its surrounding areas (see Figure Q12-1). The values decreased strongly starting from those found in the 1970s and now are approximately 30% less than in pre-ozone- hole years (1970–1982). The average values show a larger year-to-year variability than found for the other ozone measures noted above because the average includes areas outside the ozone hole, where the distribution of total ozone varies in response to meteorological conditions. Increased year-to-year variability is evident over the past decade. The maps in Q11-3 show also how the maximum in total ozone surrounding the ozone hole each year has diminished over the last three decades, adding to the decreases noted in Figure Q12-1.

  •  

Figure Q11-3 Antarctic total ozone

Disappearance of the ozone hole in spring. The severe depletion of Antarctic ozone occurs in the late winter/early spring season. In spring, temperatures in the polar lower stratosphere increase (see Figure Q10-1), stopping the formation of PSCs and production of ClO and, consequently, the most effective chemical cycles that destroy ozone. The polar vortex breaks down, ending the wintertime isolation of high-latitude air and increasing the exchange of air between the Antarctic stratosphere and lower latitudes. This allows substantial amounts of ozone-rich air to be transported poleward, where it displaces or mixes with air depleted in ozone. The midlatitude air also contains higher abundances of nitrogen oxide gases (NOx), which help convert the most reactive chlorine gases (ClO) back into the chlorine reservoir gas ClONO2 (see Q6). As a result of these large-scale transport and mixing processes, the ozone hole disappears by December.

Figure Q11-4 Unusual 2002 ozone hole

The 2002 Antarctic Ozone Hole

The 2002 Antarctic ozone hole showed features that looked surprising at the time (see Figure Q11-4). It had much less ozone depletion as measured by the area of the ozone hole or minimum total-ozone amounts in comparison with the 2001 ozone hole. The 2002 values now stand out clearly in the year-to-year changes in these quantities displayed in Figure Q11-2. There were no forecasts of an ozone hole with unusual features in 2002 because the chemical and meteorological conditions required to deplete ozone, namely low temperatures and available reactive halogen gases, were present that year and did not differ substantially from previous years. The ozone
hole initially formed as expected in August and early September 2002. Later, during the last week of September, an unexpected and surprisingly strong meteorological event occurred that dramatically reshaped the ozone hole into two separate depleted regions. As a result of this disturbance, the combined area of these two regions in late September and early October was signi cantly less than that observed for the previous or subsequent ozone holes.

The unexpected meteorological in uence in 2002 resulted from speci c atmospheric air motions that sometimes occur in polar regions. Meteorological analyses of the Antarctic stratosphere show that it was warmed by very strong, large-scale weather systems that originated in the lower atmosphere (troposphere) at midlatitudes in late September. At that time, Antarctic temperatures are generally very low (see Q10) and ozone destruction rates are near their peak values. The in uence of these tropospheric systems extended poleward and upward into the stratosphere, disturbing the normal circumpolar wind ow (polar vortex) and warming the lower stratosphere where ozone depletion was ongoing. Higher temperatures reduced the rate of ozone depletion and led to the higher minimum values observed for total ozone in Figure Q11-2. The higher-than-normal impact of these weather disturbances during the critical time period for ozone loss reduced the total loss of ozone in 2002.

The strong influence of the 2002 warming event is unique in the many decades of Antarctic meteorological observations. Another warming event occurred in 1988 causing smaller changes in the ozone hole features in Figure Q11-2. Large warming events are dif cult to predict because of the complex conditions leading to their formation.

In 2003 through 2009, ozone hole features returned to values observed from the mid-1990s to 2001 (see Figure Q11-2). The high ozone depletion found since the mid-1990s, with the exception of 2002, is expected to be typical of coming years. A significant, sustained reduction of Antarctic ozone depletion, leading to full recovery of total ozone, requires comparable, sustained reductions of ODSs in the stratosphere. Even with the source gas reductions already underway (see Q16), the return of Antarctic total ozone to 1980 values is not expected to occur before 2050.

Yes, significant depletion of the Arctic ozone layer now occurs in most years in the late winter/early spring period (January–March). However, the maximum depletion is less severe than that observed in the Antarctic and with large year-to-year differences as a consequence of the highly variable meteorological conditions found in the Arctic polar stratosphere. Even the most severe Arctic ozone depletion does not lead to total ozone amounts as low as those seen in the Antarctic, because Arctic ozone abundances during early winter before the onset of ozone depletion are much higher than those in the Antarctic. A large and recurrent “ozone hole,” as found in the Antarctic stratosphere, does not occur in the Arctic.

Signi cant depletion of ozone has been observed in the Arctic stratosphere in recent decades. The depletion is attributable to chemical destruction by reactive halogen gases, which increased in the stratosphere in the latter half of the 20th century (see Q16). Arctic depletion also occurs in the late winter/early spring period (January– March), but over a somewhat shorter period than in the Antarctic (July–October). Similar to the Antarctic (see Q11), Arctic depletion occurs because of (1) the periods of extremely low temperatures, which cause polar stratospheric clouds (PSCs) to form; (2) the large abundance of reactive halogen gases produced in reactions on PSCs; and (3) the isolation of polar stratospheric air, which allows time for chemical destruction processes to occur.

Arctic ozone depletion is much less than that observed each Antarctic winter/spring season. Large and recurrent ozone holes as found in the Antarctic stratosphere
do not occur in the Arctic. Stratospheric ozone abundances during early winter, before the onset of ozone depletion, are naturally higher in the Arctic than in the Antarctic because transport of ozone from its source region in the tropics to higher latitudes is more vigorous in the Northern Hemisphere. Depletion is limited because, in comparison to Antarctic conditions, Arctic average temperatures are always signi cantly higher (see Figure Q10-1) and the isolation of polar stratospheric air is less effective. Temperature and other meteorological differences occur because northern polar latitudes have more land and mountainous regions than southern polar latitudes (compare Figures Q11-3 and Q12-2). In a few Arctic winters, for example, PSCs did not form because temperatures never reached low enough values. These differences cause the extent and timing of Arctic ozone depletion to vary considerably from year to year. Depletion in some winter/spring seasons occurs over many weeks, in others only for brief early or late periods, and in some not at all.

Figure Q12-1 Average total ozone in polar regions

Long-term total ozone changes. Satellite observations can be used in two important ways to examine the average total ozone abundances in the Arctic region for the last 30–40 years and to contrast these values with Antarctic abundances:

  • First, total ozone poleward of 63°N averaged for each March shows quantitatively how total ozone has changed in the Arctic (see Figure Q12-1). The seasonal poleward and downward transport of ozone-rich air is naturally stronger in the Northern Hemisphere. As a result, total ozone values at the beginning of each winter season in the Arctic are considerably higher than those in the Antarctic. Before depletion sets in, normal Arctic values are close to 450 DU and Antarctic values close to 330 DU. Decreases from pre-ozone-hole average values (1970– 1982) were observed in the Arctic by the mid-1980s, when larger changes were already occurring in the Antarctic. The decreases in total ozone are generally much smaller than those found in the Antarctic and lead to total ozone values that are on average around 10–15% below pre-ozone-hole values. Maximum decreases in Arctic total ozone of about 30% observed in March 1997 and 2011 are the most comparable to Antarctic depletion. In both these Arctic winters, meteorological conditions kept ozone transport to high latitudes below average values, and in 2011 persistently low temperatures facilitated severe chemical ozone depletion by ODSs. Overall, Arctic total ozone values tend to show larger year-to-year variability than those in the Antarctic. Ozone differences from the 1970–1982 average value are due to a combination of chemical destruction by ODSs and meteorological (natural) variations. In the last two decades, these two aspects have contributed about equally to observed ozone changes. The amount of chemical destruction depends in large part on stratospheric temperatures. Meteorological conditions determine how well Arctic stratospheric air is isolated from ozone-rich air at lower latitudes and in uence the extent of low temperatures.
  • Second, total ozone maps over the Arctic and surrounding regions (see Figure Q12-2) show year-to-year changes in March total ozone. In the 1970s, total ozone values were near 450 DU when averaged over the Arctic region in March. Beginning in the 1990s and continuing into the early 2010s, values above 450 DU were increasingly absent from the March average maps. A comparison of the maps in the 1970s and early 2010s, for example, shows a striking reduction of total ozone throughout the Arctic region. The large geographical extent of low total ozone in the maps of March 1997 and March 2011 represent exceptional events in the Arctic observational record of the last three decades as noted above in the discussion of Figure Q12-1.

Altitude pro les of Arctic ozone. Arctic ozone is measured using a variety of instruments (see Q5), as in the Antarctic, to document daily to seasonal changes within the ozone layer. Spring Arctic and Antarctic balloonborne measurements are contrasted in Figure Q12-3. Arctic pro les were obtained from the Ny-Ålesund research station at 79°N. For 1989–2014, the March average reveals a substantial ozone layer, contrasting sharply with the severely depleted Antarctic ozone layer in the October average over a similar time period. This further demonstrates how higher stratospheric temperatures and meteorological variability have protected the Arctic ozone layer from the greater ozone losses that occur in the Antarctic, despite similar reactive halogen abundances in the two regions.

The separate Arctic pro les shown for 29 March 1996 and 1 April 2011 are two of the most severely depleted in the 25-year long record from Ny-Ålesund. Although signi cant, the depletion during both events is smaller in comparison to that routinely observed in the Antarctic, such as in the pro le from 9 October 2006. The near- complete depletion of ozone over many kilometers in altitude and over areas almost as large as North America, as is now common in the Antarctic stratosphere, has never been observed in the Arctic.

Figure Q12-3 Vertical distribution of Arctic and Antarctic ozone

Restoring ozone in spring. As in the Antarctic, ozone depletion in the Arctic is con ned to the late winter/early spring season. In spring, temperatures in the polar lower stratosphere increase (see Figure Q10-1), halting PSC formation and production of ClO, and thus the most effective chemical cycles that destroy ozone. The breakdown of the polar vortex ends the wintertime isolation of high-latitude air, increasing exchange of air between the Arctic stratosphere and lower latitudes. This allows more ozone-rich air to be transported poleward, where it displaces or mixes with air in which ozone may have been depleted. As a result of these large- scale transport and mixing processes, any ozone depletion typically disappears by April or earlier.

Depletion of the global ozone layer increased gradually in the 1980s and reached a maximum of about 5% in the early 1990s. The depletion has lessened since then and now is about 3% averaged over the globe. The average depletion exceeds the natural year-to-year variations of global total ozone. The ozone loss is very small near the equator and increases with latitude toward the poles. The larger polar depletion is attributed to the late winter/ early spring ozone destruction that occurs there each year.

Global total ozone has decreased beginning in the 1980s (see Figure Q13-1). The decreases have occurred in the stratospheric ozone layer where most ozone resides (see Figure Q1-2). In the early 1990s, the depletion of global total ozone reached a maximum of about 5% below the 1964–1980 average. The depletion has lessened since then
and during the early 2010s has averaged to about 3% below the 1964–1980 average. The observations shown in Figure Q13-1 have been smoothed to remove regular ozone changes that are due to natural seasonal and solar effects (see Q14). The depleted amounts are larger than the remaining natural variations in global total ozone amounts.

The observed global ozone depletion in the last three decades is attributable to increases in reactive halogen gases in the stratosphere. The lowest global total ozone values since 1980 have occurred in the years following the eruption of Mt. Pinatubo in 1991, which temporarily increased the number of sulfuric acid-containing particles throughout the stratosphere. These particles signi cantly increased the effectiveness of reactive halogen gases in destroying ozone (see Q14) and, thereby, increased global ozone depletion by 1–2% for several years following the eruption.

Polar regions. Observed total ozone depletion varies signi cantly with latitude on the globe (see Figure Q13-1). The largest reductions occur at high southern latitudes as a result of the severe ozone loss over Antarctica each late winter/early spring period. The next largest losses are observed in the high latitudes of the Northern Hemisphere, caused in part by winter losses over the Arctic. Although the depletion in polar regions is larger than at lower latitudes, the in uence of polar regions on global ozone is limited by their small geographical area. Latitudes poleward of 60° account for only about 13% of Earth’s surface.

Midlatitude regions. Ozone depletion is also observed at the midlatitudes spanning the region between equatorial and polar latitudes. In comparison with the 1964–1980 averages, total ozone averaged for 2008–2012 is about 3.5% lower in northern midlatitudes (35°N–60°N) and about 6% lower at southern midlatitudes (35°S–60°S). Midlatitude depletion has two contributing factors. First, ozone- depleted air over both polar regions is dispersed away from the poles during and after each winter/spring period, thereby reducing average ozone outside polar regions. Second, chemical destruction occurring at midlatitudes contributes to observed depletion in these regions. This contribution is much smaller than in polar regions because the amounts of reactive halogen gases are lower and a seasonal increase of the most reactive halogen gases, such as the increase in ClO in the polar winter lower stratosphere (see Figure Q8-3), does not occur in midlatitude regions.

Figure Q13-1 Global total ozone changes

Tropical region. Total ozone in the tropics (20°N–20°S latitude) has been only weakly affected by chemical ozone depletion. In the tropical lower stratosphere, air has been transported from the lower atmosphere (troposphere) within the previous 18 months. As a result, the conversion of ozone-depleting substances (ODSs) to reactive halogen gases is still very small. With little reactive halogen available, total ozone depletion in this region is also very small. In addition, ozone production is high because average solar ultraviolet radiation is highest in the tropics. In contrast, stratospheric air in polar regions has been in the stratosphere for an average of 4 to 7 years, allowing time for signi cant conversion of ODSs to reactive halogen gases. The systematic differences in the age of stratospheric air are well understood and are a consequence of the large-scale atmospheric transport: air enters the stratosphere in the tropics, moves poleward in both hemispheres, and then descends and ultimately returns to the troposphere in the middle to high latitudes.

Initial Signs of Ozone Recovery

The Montreal Protocol, strengthened by its Amendments and adjustments, has successfully controlled the production and consumption of ODSs that act to destroy the ozone layer (Q15). As a result, atmospheric abundances of ODSs have peaked and are now decreasing (Q7 and Q16). By 2012, equivalent effective stratospheric chlorine (EESC; the total chlorine and bromine abundances in the stratosphere) had declined by 15% at midlatitudes from peak values of around 15 years ago. This raises the question, is global ozone increasing in response to the observed EESC decreases?

Identifying an ozone increase is not easy, because ODS levels are not the only factor that determines global ozone levels. For example, the global ozone minimum was observed half a decade before the EESC maximum was reached. The difference in the timing resulted from the strong global ozone response to enhanced stratospheric aerosol loading after the Mount Pinatubo eruption in 1991, which led to increased ozone depletion for several years. Observed global ozone increases through the 1990s were therefore a result of the steady removal of the aerosol from the stratosphere, and not a sign of decreasing ODSs (see Q14). Another factor complicating the identi cation of ozone recovery in different regions of the atmosphere is the year-to-year variations of the stratospheric circulation. These variations lead to ozone variability in most regions of the atmosphere that is currently still larger than the signal expected from the observed EESC decreases. Finally, greenhouse gas increases (such as carbon dioxide, CO2) affect ozone by decreasing stratospheric temperatures (which slows down ozone depletion rates) and by strengthening the stratospheric circulation (which enhances the transport of ozone from the tropics to higher latitudes). It is therefore dif cult to attribute observed ozone changes to these different factors.

Observations now show a clear 5% increase of ozone in the upper stratosphere (42 km) over the 2000-2013 period. Model simulations that allow for separation of the different factors suggest that about half of this increase results from a cooling in this region due to CO2 increases, while the other half results from EESC decreases. Also, total column ozone declined over most of the globe during the 1980s and early 1990s (by about 2.5% averaged over 60°S to 60°N). It has remained relatively unchanged since 2000, with indications of a small increase in total column ozone in recent years. Models suggest that this small increase is likely due to EESC decreases. These ndings based on both models and observations suggest that there are initial signs of ozone recovery.

Because of their long lifetime, the impact on stratospheric ozone of the most prominent ODSs (CFC-11 and CFC-12) will continue for many decades after emissions have ceased. Assuming continued compliance with the Montreal Protocol, EESC will continue to decline over the coming decades and will return to pre-1980 levels around midcentury. With the exception of the tropics (see Q20), climate change is expected to accelerate the return of the ozone layer to pre-1980 levels. However, as long as ODS levels remain elevated in the atmosphere, the possibility of extreme low-ozone events due to volcanic eruptions or cold winter conditions persists into the second half of the 21st century.

Yes, factors such as changes in solar radiation, as well as the formation of stratospheric particles after volcanic eruptions, do influence the ozone layer. However, neither factor can explain the average decreases observed in global total ozone over the last three decades. If large volcanic eruptions occur in the coming decades while ODS levels remain high, ozone depletion will increase for several years afterward.

Changes in solar radiation and increases in stratospheric particles from volcanic eruptions both affect the abundance of stratospheric ozone. Global total ozone in
the early 1990s had decreased by about 5% when compared to pre-1980 values, and is now about 3% below pre-1980 values (see Q13). The depletion is primarily attributed to increases in reactive halogen gases, with additional depletion in the early 1990s associated with the volcanic eruption of Mt. Pinatubo. Reactive halogen gases are represented by changes in equivalent effective stratospheric chlorine (EESC). EESC values account for stratospheric chlorine and bromine abundances and their different effectiveness in destroying ozone (see de nition in Q16). A comparison of the smooth year-to-year changes in ozone and EESC shows that the quantities are inversely related to each other, with ozone rst decreasing while EESC increased (see Figure Q14-1). After the mid-1990s, the annual changes in both quantities remained constant and then slowly reversed sign. Changes in solar output and volcanic activity do not show such smooth long-term changes, as discussed below, and therefore cannot be the cause of long-term global ozone depletion.

Total ozone and solar changes. The formation of stratospheric ozone is initiated by ultraviolet (UV) radiation coming from the Sun (see Figure Q2-1). As a result, an increase in the Sun’s radiation output increases the amount of ozone in Earth’s atmosphere. The Sun’s radiation output and sunspot number vary over the well- documented 11-year solar cycle. Observations over several solar cycles since the 1960s show that global total ozone levels vary by 1 to 2% between the maximum and minimum of a typical cycle. Changes in incoming solar radiation at a wavelength of 10.7 cm are often used as a surrogate for changes in solar output at UV wavelengths. The long-term changes in the 10.7-cm output in Figure Q14-1 clearly show alternating periods of maximum and minimum values in total solar output separated by about 5 to 6 years. If changes in solar output were the cause of global ozone depletion, a gradually decreasing output would have been observed around 1980 or earlier, slowing sharply in the mid-1990s. Since such a decrease was not observed, nor is expected based on longer-term solar observations, the long-term decreases in global ozone cannot result from changes in solar output alone. Most analyses presented in this and previous international scienti c assessments quantitatively account for the in uence of the 11-year solar cycle on long-term variations in ozone.

Total ozone and past volcanoes. Explosive volcanic eruptions inject sulfur gases directly into the stratosphere, causing new sulfate particles to be formed. The particles initially form in the stratosphere downwind of the volcano and then spread throughout the hemisphere or globally as air is transported by stratospheric winds. One method of detecting the presence of volcanic particles in the stratosphere uses observations of the transmission of solar radiation through the atmosphere (see Figure Q14-1). When large amounts of new particles are formed in the stratosphere over an extensive region, solar transmission is measurably reduced. The eruptions of Mt. Agung (1963), El Chichón (1982), and Mt. Pinatubo (1991) are the most recent examples of sulfur injections that temporarily reduced solar transmission.

Laboratory measurements and stratospheric observations have shown that chemical reactions on the surfaces of volcanically produced particles can increase ozone destruction by increasing the amounts of the highly reactive chlorine gas, chlorine monoxide (ClO). The ozone response depends on the total abundance of EESC after the eruption (see Q16). At times of relatively low EESC, such as the early 1960s, ozone is not very sensitive to stratospheric injection of volcanic sulfate particles. At times of higher EESC amounts, such as from 1980 to the present, global ozone is expected to decrease signi cantly following large explosive eruptions. The most recent large eruption was that of Mt. Pinatubo, which resulted in as much as a 10-fold increase in the number of particles available for surface reactions. Both El Chichón and Mt. Pinatubo increased global ozone depletion for a few years (see Figure Q14-1). EESC was too low for ozone depletion to occur after the Mt. Agung eruption in 1963. The effect on ozone diminishes during the years following an eruption as volcanic particles are gradually removed from the stratosphere by gravitational settling and large-scale atmospheric transport. As particles are removed, solar transmission is restored. Based on the short residence time of volcanic particles in the stratosphere, the two large eruptions in the past three decades cannot account directly for the continuous long-term decreases in global total ozone observed over the same period.

Reactive chlorine from volcanoes. Explosive volcanic plumes generally contain large quantities of reactive chlorine in the form of hydrogen chloride (HCl). HCl is a reactive halogen gas that can be converted to ClO, which rapidly destroys ozone (see Figure Q8-3). The plumes also contain a considerable amount of water vapor, which forms rainwater and ice in the rising fresh plume. Rainwater and ice ef ciently scavenge and remove HCl while it is still in the lower atmosphere (troposphere). As a result, most of the HCl in explosive volcanic plumes does not enter the stratosphere. After recent explosive eruptions, observations of HCl in the stratosphere have con rmed that increases are small compared with the total amount of chlorine in the stratosphere from other sources.

Figure Q14-1 Solar changes and volcanoes

Antarctic volcanoes. Volcanoes on the Antarctic continent are of special interest due to their proximity to the Antarctic ozone hole. An explosive eruption could in principle inject volcanic aerosol and small amounts of HCl directly into the stratosphere over Antarctica, which could lead to ozone depletion. However, to be a possible cause of the annually recurring ozone hole beginning in the early 1980s, explosive Antarctic eruptions would need to have occurred at least every few years to maintain volcanic emissions in the stratosphere. This is not the case. Only the Mt. Erebus volcano is currently active in Antarctica. No explosive eruptions of Mt. Erebus or any other Antarctic volcano have occurred since 1980. Therefore, explosive volcanic eruptions in the last three decades have not caused the Antarctic ozone hole and, as noted above, have not been suf cient to cause the long-term depletion of global total ozone.

Total ozone and future volcanoes. Observations and atmospheric models indicate that the record-low ozone levels observed in 1992–1993 resulted from the large number of particles produced by the Mt. Pinatubo eruption, combined with the relatively large amounts of EESC present in the stratosphere in the early 1990s. If the Mt. Pinatubo eruption had occurred before 1980, changes to global ozone would have been much smaller than observed in 1992–1993 because EESC values were much lower. EESC values will remain substantial in the early decades of the 21st century even as ODSs decline globally, with 1980 values not reached until about 2050 (see Figures Q16-1 and Q20-2). Large volcanic eruptions in the intervening years will cause more ozone depletion. If an explosive eruption larger than Mt. Pinatubo were to occur, peak ozone losses could be larger than previously observed and substantial ozone losses could persist for longer time periods. As halogen gas abundances gradually decline, the effect of volcanic eruptions on ozone will lessen.

Yes, the production and consumption of ozone-depleting substances are controlled under a 1987 international agreement known as the “Montreal Protocol on Substances that Deplete the Ozone Layer” and by its subsequent Amendments and adjustments. The Protocol, now rati ed by all 196 United Nations members, establishes legally binding controls on national production and consumption of ozone-depleting substances (ODSs). Production and consumption of all principal ODSs by developed and developing nations will be almost completely phased out before the middle of the 21st century.

Montreal Protocol. In 1985, a treaty called the Vienna Convention for the Protection of the Ozone Layer was signed by 20 nations in Vienna. The signing nations agreed to take appropriate measures to protect the ozone layer from human activities. The Vienna Convention was a framework agreement that supported research, exchange of information, and future protocols. In response to growing concern, the Montreal Protocol on Substances that Deplete the Ozone Layer was signed in 1987 and, following suf cient country rati cation, entered into force in 1989. The Protocol has been successful in establishing legally binding controls for developed and developing nations on the production and consumption of halogen source gases known to cause ozone depletion. Halogen source gases containing chlorine and bromine controlled under the Montreal Protocol are referred to as ozone-depleting substances (ODSs). National consumption of an ODS is de ned as the amount that production and imports of the substance exceed its export to other nations. The Protocol provisions are structured for developed countries to act rst and for developing countries to follow with some nancial assistance. In 2009, the Montreal Protocol became the rst international treaty to achieve universal rati cation by all 196 United Nations members.

Amendments and adjustments. As the scientific basis of ozone depletion became more certain after 1987 and substitutes and alternatives became available to replace ODSs, the Montreal Protocol was strengthened with Amendments and adjustments. Each Amendment is named after the city in which the meeting of the Parties to the Montreal Protocol took place and by the year of the meeting. The timeline in Figure Q0-1 shows some of the major decisions that have been adopted in the last three decades. These decisions have put additional ODSs under control, accelerated the timing of existing control measures, and prescribed phase-out dates for the production and consumption of certain gases. The initial Protocol called for a 50% reduction in chloro uorocarbon (CFC) production and a freeze on halon production. The 1990 London Amendment called for a phase-out of the production and consumption of the most damaging ODSs in developed nations by 2000 and in developing nations by 2010. The 1992 Copenhagen Amendment accelerated the phase-out date to 1996 in developed nations. Further controls on ODSs were agreed upon in later meetings in Vienna (1995), Montreal (1997, 2007), and Beijing (1999).

Figure Q15-1 Effect of the Montreal Protocol

 

 

Influence of the Montreal Protocol. Montreal Protocol controls are based on several factors that are considered separately for each ODS. The factors include (1) the effectiveness in depleting ozone in comparison with other substances (see Ozone Depletion Potential, ODP, in Q18), (2) the availability of suitable substitutes for domestic and industrial use, and (3) the potential impact of controls on developing nations. The in uence of Montreal Protocol provisions on stratospheric ODS abundances can be demonstrated with long-term changes in equivalent effective stratospheric chlorine (EESC). Calculations of EESC combine the amounts of chlorine and bromine present in surface air to form a measure of the potential for ozone destruction in a particular stratospheric region on an annual basis (see de nition in Q16). The long-term changes in EESC at midlatitudes are shown in Figure Q15-1 for several cases:

  • No Protocol. Without the Montreal Protocol the production, use, and emissions of CFCs and other ozone-depleting substances is expected to have increased after 1987 with an annual growth rate of about 3% (business-as-usual scenario). As a result, EESC is projected to have increased nearly 10-fold by the mid-2050s compared with the 1980 value. Computer models of the atmosphere indicate that such high EESC values would have at least doubled global total ozone depletion between 1990 and 2010 and increased it far beyond that by midcentury. As a result, harmful UV-B radiation would have increased substantially at Earth’s surface, causing damage to ecosystem health and a global rise in skin cancer and cataract cases (see Q17).
  • Montreal Protocol provisions. International compliance with only the 1987 provisions of the Montreal Protocol and the later 1990 London Amendment would have substantially slowed the projected growth of EESC. Not until the 1992 Copenhagen Amendments and adjustments did the Protocol projections show a decrease in future EESC values. The provisions became more stringent with the Amendments and adjustments adopted in Beijing in 1999 and Montreal in 1997 and 2007. Now, with full compliance to the Protocol, most ODSs will be almost completely phased out, with some exemptions for critical uses (see Q16). Global EESC is slowly decaying from its peak value in the late 1990s and is expected to reach 1980 values in the mid-21st century. The success of the Montreal Protocol to date is demonstrated by the decline in ODP-weighted emissions of ODSs shown in Figure Q0-1. Total emissions peaked in 1987 at values about 10-fold higher than natural emissions. Between 1987 and 2014, ODS emissions from human activities have decreased by almost 80%.
  • Zero emissions. EESC values in the coming decades will be in uenced by (1) the slow natural removal of ODSs still present in the atmosphere, (2) emissions from continued production and use of ODSs, and (3) emissions from currently existing banks containing a variety of ODSs. ODS banks are associated with applications that involve long-term containment of halogen gases. Examples are CFCs in refrigeration equipment and insulating foams, and halons in re- ghting equipment. New emissions are projected based on continued production and consumption of ODSs, particularly in developing nations, under existing Protocol provisions. The zero-emissions case demonstrates the EESC values that would occur if it were possible to set all ODS emissions to zero beginning in 2014. This would eliminate the contributions from new production and bank emissions. Signi cant differences from the Montreal 2007 projections are evident in the rst decades following 2014 because the phase-out of all ODS production under the Protocol is not yet complete and bank emissions are substantial. Zero emissions would bring forward the return of EESC to 1980 levels by about 13 years.

HCFC substitute gases. The Montreal Protocol provides for the use of hydrochloro uorocarbons (HCFCs) as transitional, short-term substitute compounds for ODSs with higher ODPs, such as CFC-12. HCFCs are used for refrigeration, in making insulating foams, and as solvents, all of which were primary uses of CFCs. HCFCs are generally more reactive in the troposphere than other ODSs because they contain hydrogen (H) in addition to chlorine, uorine, and carbon. HCFCs are 88 to 98% less effective than CFC-12 in depleting stratospheric ozone because their chemical removal occurs primarily in the troposphere (see ODPs in Table Q7-1). This removal protects stratospheric ozone from most of the halogen content of HCFC emissions. In contrast, CFCs and some other ODSs release all of their halogen content in the stratosphere because they are chemically inert in the troposphere (see Q6).

Under the provisions of the Montreal Protocol, developed and developing countries may continue to use HCFCs as ODS substitutes in the coming decades before they are ultimately phased out. In the most recent adjustment to the Protocol (Montreal 2007), the phase-out of HCFCs for all Parties was accelerated so that it will be complete in 2030, a decade earlier than in previous provisions. In adopting this decision, the Parties reduced the contribution of HCFC emissions to both long-term ozone depletion and future climate forcing (see Q18 and Q19).

HFC substitute gases. Hydro uorocarbons (HFCs) are used as substitute compounds for CFCs, HCFCs, and other ODSs. HFCs contain hydrogen, uorine, and carbon. HFCs do not contribute to ozone depletion because they contain no chlorine, bromine, or iodine. As a consequence, HFCs are not ODSs and are not subject to the provisions of the Montreal Protocol. HFCs and all ODSs are radiatively active gases that contribute to human-induced climate change based on their accumulation in the atmosphere (see Q18). HFCs are included in the basket of gases controlled by the Kyoto Protocol of the United Nations Framework Convention on Climate Change (UNFCCC). The Kyoto Protocol is an international treaty designed to protect climate by controlling emissions of carbon dioxide (CO2), methane (CH4), nitrous oxide (N2O), HFCs, per uorocarbons (PFCs), and sulfur hexa uoride (SF6). HFC emissions are expected to grow substantially in coming decades as the demand for their use as substitute gases and in new applications increases in developed and developing countries. The climate change contribution from future HFC emissions would be minimized if HFCs with very low Global Warming Potentials (GWPs) (less than 100) were chosen to meet the growing demand (see Q18 and Q19).

Yes, as a result of the Montreal Protocol, the overall abundance of ozone-depleting substances (ODSs) in the atmosphere has been decreasing for more than a decade. If the nations of the world continue to comply with the provisions of the Montreal Protocol, the decrease will continue throughout the 21st century. Those gases that are still increasing in the atmosphere, such as halon-1301 and HCFCs, will begin to decrease in the coming decades if compliance with the Protocol continues. However, it is only after midcentury that the effective abundance of ODSs is expected to fall to values that were present before the Antarctic ozone hole was rst observed in the early 1980s.

The Montreal Protocol and its Amendments and adjustments have been very successful in reducing the atmospheric abundance of ozone-depleting substances (ODSs). ODSs are halogen source gases released by human activities. Under the Protocol, the production and consumption of individual ODSs are now controlled in all 196 nations that are Parties to the Protocol (see Q15). The success of the Montreal Protocol controls is documented by (1) observed changes and future projections of the atmospheric abundances of the principal ODSs and (2) long-term changes in equivalent effective stratospheric chlorine (EESC).

Individual ODS reductions. The reduction in the atmospheric abundance of an ODS in response to controls on production and consumption depends principally
on two factors: (1) how rapidly an ODS is used and released to the atmosphere after being produced and (2) the lifetime for the removal of the ODS from the atmosphere (see Table Q7-1). For example, the abundances of ODSs with short lifetimes respond quickly to emission reductions. Long-term changes in ODS abundances are constructed from (1) estimates of historical emissions and bank quantities using industry reports, (2) abundances measured in air trapped for years in accumulated snow ( rn) in polar regions, (3) observed atmospheric abundances using ground-based measurements, and (4) projections of future abundances based on compliance with Montreal Protocol provisions and patterns of use in developed and developing countries. The results for individual ODSs and the natural chlorine and bromine source gases, methyl chloride (CH3Cl) and methyl bromide (CH3Br), are shown in Figure Q16-1 and described as follows:

  • CFCs. Chlorofluorocarbons (CFCs) include some of the most destructive chlorine- containing ODSs. CFC-11 and CFC-12, each with an Ozone Depletion Potential (ODP) near 1, are the most abundant ODSs in the atmosphere owing to large historical emissions and long atmospheric lifetimes (50–100 years). Production and consumption of CFCs in developed countries ended in 1996 and that in developing countries ended by January 2010. As a consequence, CFC-11 and CFC-113 abundances have peaked in the atmosphere and have been declining for more than a decade. In contrast, CFC-12 abundances have only recently shown a decrease, owing to its longer lifetime (100 years) and continuing emissions from CFC-12 banks, namely, refrigeration and air conditioning equipment and thermal insulating foams. With no further global production of the principal CFCs allowed except for limited exempted uses and with some continuing emissions from banks, CFC abundances are projected to decline steadily throughout this century.
  • Halons. Halons are the most destructive bromine-containing ODSs. The most abundant in the atmosphere, halon-1211 and halon-1301, have abundances about 100 times less than CFC-11 and CFC-12 and account for a signi cant fraction of bromine from all ODSs (see Figure Q7-1). Production and consumption of halons in developed countries ended in 1994 and that in developing countries ended by January 2010. Atmospheric abundances of halon-1211 show signi cant decreases since peak concentrations were measured in the early 2000s. Halon- 2402 abundances have been decreasing slowly for a few years while those of halon-1301 continue to increase. The increase is likely due to substantial banks in re-extinguishing and other equipment that gradually release halon-1301 to the atmosphere. The abundance of halon-1301 is expected to remain high well into the 21st century because of its long lifetime (65 years) and continued release.
  • Methyl chloroform. The largest reduction to date in the abundance of an ODS (96% from its peak value) has been observed for methyl chloroform. Production and consumption of methyl chloroform in developed countries ended in January 1996 and that in developing countries is scheduled to end by January 2015. Atmospheric abundances responded rapidly to the reduced emissions starting in the mid-1990s because methyl chloroform has a short atmospheric lifetime (about 5 years). Methyl chloroform is used mainly as a solvent and has no signi cant long-term storage following production. It will approach complete removal from the atmosphere after the phase-out in developing countries is complete.
  • HCFC substitute gases. The Montreal Protocol allows for the use of hydrochlorofluorocarbons (HCFCs) as short-term, transitional substitutes for CFCs and in other speci c applications. As a result, the abundances of HCFC-22, HCFC- 141b, and HCFC-142b continue to grow in the atmosphere in response to greater application demand. HCFCs pose a lesser threat to the ozone layer than CFCs because they have low ODP values (less than 0.12). The most recent adjustment to the Montreal Protocol (Montreal 2007) accelerated the phase-out of HCFCs by a decade for both developed countries (2020) and developing countries (2030) (see Q15). Even with the accelerated phase-out, future projections show HCFC abundances that continue to increase, reach peak values around 2030, and steadily decrease thereafter. The response of atmospheric abundances to decreasing emissions will be relatively rapid because of the short atmospheric lifetimes of HCFCs (less than 20 years).
  • Carbon tetrachloride. Carbon tetrachloride has been phased out in both developed countries (January 1996) and developing countries (January 2010). As a result, atmospheric abundances of carbon tetrachloride have been decreasing for two decades. The decline is somewhat less rapid than expected, suggesting that actual emissions are larger than reported or the atmospheric lifetime is greater than current estimates. Production of carbon tetrachloride for use as raw material (feedstock) to make other chemicals is exempted under the Protocol because emission to the atmosphere from this use is believed to be minimal.
  • Methyl chloride and methyl bromide. Both methyl chloride and methyl bromide are distinct among halogen source gases because substantial fractions of their emissions are associated with natural processes (see Q7). Methyl chloride is not an ODS under the Montreal Protocol because it is not manufactured or used in appreciable quantities. Its abundance in the atmosphere has remained fairly constant throughout the last 60 years and will remain constant if the balance of total sources and loss processes remains unchanged. Methyl bromide is controlled under the Protocol because it is manufactured for use as a fumigant. Developed country production and consumption of methyl bromide ended in January 2005 and that in developing countries is scheduled to end by January 2015. The Protocol currently provides exemptions for some methyl bromide production and use as an agricultural and pre-shipment fumigant. Atmospheric abundances of methyl bromide responded rapidly to the reduced emissions starting in 1999 because its atmospheric removal lifetime is less than 1 year. Future projections show only small changes in methyl bromide abundances based on the assumptions of unchanged contributions from natural sources and small use in developing countries. An important uncertainty in these projections is the future amounts that will be produced and emitted under Montreal Protocol use exemptions.

Figure Q16-1 Halogen source gas changes

Equivalent effective stratospheric chlorine (EESC). Important measures of the success of the Montreal Protocol are the past and projected changes in the values of equivalent effective stratospheric chlorine (EESC), which was introduced in Figures Q14-1 and Q15-1. EESC is designed as one measure of the potential for ozone depletion in the stratosphere that can be calculated from atmospheric surface abundances of ODSs and natural chlorine and bromine gases. All gases used in the calculation are shown in Figure Q7-1. For both past and future EESC values, the required atmospheric abundances are derived from measurements, historical estimates, or future projections based on compliance with Montreal Protocol provisions.

EESC is expressed as a hypothetical amount of chlorine available in the stratosphere to deplete ozone. The term equivalent indicates that bromine gases, scaled by their greater per-atom effectiveness in depleting ozone, are included in EESC. The term effective indicates that only the estimated fraction of ODSs that are currently in the form of reactive halogen gases in the stratosphere is included in an EESC value (see Q8). Although chlorine is much more abundant in the stratosphere than bromine (about 150-fold) (see Figure Q7-1), bromine atoms are about 60 times more effective than chlorine atoms in chemically destroying ozone in the lower stratosphere. EESC values generally depend on the year and latitude region in the stratosphere being considered.

Another quantitative measure of the potential for ozone depletion in the stratosphere that is highly related to EESC is effective stratospheric chlorine (ESC). ESC, as de ned in Q20, is calculated with chemistry-climate models rather than inferred from tropospheric ODS observations.

Long-term changes in EESC. In the latter half of the 20th century up until the 1990s, EESC values steadily increased (see Figure Q16-1), causing global ozone depletion. As a result of the Montreal Protocol regulations, the long-term increase in EESC slowed, values reached a peak, and EESC began to decrease in the late 1990s. By 2012, EESC at midlatitudes had declined by about 15% from its peak value in 1997. The initial decrease came primarily from the substantial, rapid reductions in emissions of methyl chloroform, which has a lifetime of only 5 years. The decrease is continuing with declining emissions of CFCs and other long-lived ODSs. Decreasing EESC means that the potential for stratospheric ozone depletion is now lessening each year as a result of the Montreal Protocol. Decreases in EESC are projected to continue throughout the 21st century if all nations continue to comply with the provisions of the Protocol. The decrease will continue because as emissions are reduced, natural processes continue to gradually remove halogen-containing gases from the global atmosphere. Reduction of EESC to 1980 values or lower will require several more decades because the most abundant ODS gases now in the atmosphere have lifetimes ranging from 10 to 100 years.

Yes, ultraviolet radiation at Earth’s surface increases as the amount of overhead total ozone decreases, because ozone absorbs ultraviolet radiation from the Sun. Measurements by ground-based instruments and estimates made using satellite data provide evidence that surface ultraviolet radiation has increased in large geographic regions in response to ozone depletion.

The depletion of stratospheric ozone leads to an increase in solar ultraviolet radiation at Earth’s surface. The increase occurs primarily in the ultraviolet-B (UV-B) component of the Sun’s radiation. UV-B is de ned as radiation in the wavelength range of 280 to 315 nanometers, which is invisible to the human eye. Long-term changes in UV-B radiation reaching the surface have been measured directly and can be estimated from total ozone changes.

UV-B radiation can harm humans, other life forms, and materials (see Q3). Most of the effects of sunlight on the human body are caused by UV-B exposure. A principal effect is skin erythema, which leads to sunburning. Excess exposure may lead to skin cancers. Erythemal radiation is regularly reported to the public in many countries in the form of the “UV Index.” The long-term changes in surface UV-B radiation are important to study because of its potential harmful effects and relationship to ozone depletion.

Surface UV-B radiation. The amount of UV-B radiation reaching Earth’s surface at a particular location depends in large part on total ozone at that location. Ozone molecules in the stratosphere and in the troposphere absorb UV-B radiation, thereby signi cantly reducing the amount that reaches Earth’s surface (see Q3). If conditions occur that reduce the abundance of ozone molecules somewhere in the troposphere or stratosphere, total ozone is reduced and the amount of UV-B radiation reaching Earth’s surface below is increased proportionately. This relationship between total ozone and surface UV-B radiation has been con rmed at a variety of locations with direct measurements of both quantities.

Additional causes of UV changes. The actual amount of UV-B radiation reaching Earth’s surface at a specific location and time depends on a number of factors in addition to total ozone. The primary additional factor is the position of the Sun in the sky, which changes with daily and seasonal cycles. Other factors include local cloudiness, the altitude of the location, the amount of ice or snow cover, and the amounts of atmospheric particles (aerosols) in the atmosphere above the location. Changes in clouds and aerosols are partially related to air pollution and greenhouse gas emissions from human activities. Measurements indicate that both increases and decreases in UV radiation at certain locations have resulted from changes in one or more of these factors. Estimating the impact of changes in these factors is complex. For example, an increase in cloud cover usually results in a reduction of UV radiation below the clouds and could at the same time increase UV radiation at a location in the mountains above the clouds.

Figure 17-1 Long-term changes in surface erythemal radiation

Long-term surface UV changes. Long-term changes in UV-B radiation have been estimated from measurements made with special UV monitoring instruments at several surface locations since about 1990. For example, as a consequence of Antarctic ozone depletion, the average UV-B measured at the South Pole during spring between 1991 and 2006 was 55–85% larger than estimated for the years 1963–1980. In addition, satellite observations of ozone changes have been used to estimate changes in surface UV-B radiation that have occurred over the past three decades. With satellite observations, the UV-B radiation changes can be separately attributed to changes in ozone and clouds. The results show that erythemal radiation has increased by an average of 3% between 1979 and 2008 over a wide range of latitudes outside the tropics (see lower panel of Figure Q17-1). The largest percentage increases have occurred at high polar latitudes in both hemispheres, where the largest annual decreases in total ozone are observed (see Figure Q13-1). Over this time period the UV increases due to ozone depletion are partially offset by changes in cloudiness, primarily in the high latitudes of the Southern Hemisphere (see top panel in Figure Q17-1). Without changes in cloudiness, the increases in erythemal radiation at these latitudes would have reached a maximum close to 9%. The smallest changes in erythemal UV have been in the tropics, where long-term changes in total ozone are smallest (see Q13). These ndings have not changed substantially between 2008 and 2014.

UV Index changes. The UV Index is a measure of the erythemal radiation that occurs at a particular surface location and time. The index is used internationally to increase public awareness about the detrimental effects of UV on human health and to guide the need for personal protective measures. The maximum daily UV Index varies with location and season, as shown for three sites in Figure Q17-2. The UV Index increases when moving from high to low latitudes and is highest in summer when the midday Sun is closest to overhead. For example, UV Index values in San Diego, California, at 32°N, are generally higher than those in Barrow, Alaska, at 71°N. At all latitudes, UV Index values increase in mountainous regions and over snow- or ice-covered regions. The UV Index is zero during periods of continuous darkness in winter at high-latitude locations.

The UV Index over Antarctica has increased dramatically due to ozone depletion, as illustrated in Figure Q17-2. Normal index values for Palmer, Antarctica, at 64°S in spring were estimated from satellite measurements made during the period 1978–1980, before the appearance of the ozone hole over Antarctica. In the period from 1990 to 2006, the severe and persistent ozone depletion that occurred in late winter and early spring over Antarctica increased the average UV Index well above normal values for several months. Now, the spring UV Index measured in Palmer, Antarctica, routinely equals or exceeds spring and summer values measured in San Diego, California, which is located at a much lower latitude (32°N).

Figure Q17-2 Long-term changes in the UV index

UV changes and human health. Over the past several decades, depletion of the stratospheric ozone layer together with societal changes in lifestyle have increased UV-B radiation exposure for many people. Increased exposure has adverse health effects, primarily associated with eye and skin disorders. UV radiation is a recognized risk factor for some types of eye cataracts. For the skin, the most common threat is skin cancer. Over the past decades, the incidence of several types of skin tumors has risen signi cantly among people of all skin types. Skin cancer in humans occurs long after exposure to sunburning UV. With current Montreal Protocol provisions, projections of additional skin cancer cases associated with ozone depletion are largest in the early to middle decades of the 21st century and represent a signi cant global health issue. Since it is projected that the recovery of the ozone layer to 1980 levels will not occur until the mid-21st century, extensive and unprotected UV-B exposure over the next decades will continue to have adverse health effects. An important human health bene t of UV-B radiation exposure is the production of vitamin D, which plays a signi cant role in bone metabolism and the immune system. Human exposure to solar UV-B radiation requires a careful balance to maintain adequate vitamin D levels while minimizing long-term risks of skin and eye disorders now and well into the future.

No, ozone depletion itself is not the principal cause of global climate change. Changes in ozone and climate are directly linked because ozone absorbs solar radiation and is also a greenhouse gas. Stratospheric ozone depletion leads to surface cooling, while the observed increases in tropospheric ozone and other greenhouse gases lead to surface warming. The cooling from ozone depletion is small compared to the warming from the greenhouse gases responsible for observed global climate change. Ozone depletion and global climate change are indirectly linked because both ozone-depleting substances and their substitutes are greenhouse gases. The Antarctic ozone hole has contributed to changes in Southern Hemisphere surface climate through indirect effects on the atmospheric circulation.

While stratospheric ozone depletion is not the principal cause of climate change, aspects of ozone depletion and climate change are closely linked. Both processes involve gases released to the atmosphere by human activities. The links are best understood by examining the contribution to climate change of the gases involved: ozone; ozone-depleting substances (ODSs) (or halogen source gases) and their substitutes; and other leading greenhouse gases.

Greenhouse gases and the radiative forcing of climate. The warming of Earth by the Sun is enhanced by the presence of natural greenhouse gases, of which water vapor is an important example. The natural abundances of these gases in Earth’s atmosphere change the balance between incoming solar radiation and outgoing infrared radiation, warming the atmosphere and surface. Without this natural greenhouse effect, Earth’s surface would be much colder. Human activities in the Industrial Era have led to long-term increases in the atmospheric abundances of a number of long-lived and short-lived greenhouse gases. This group includes carbon dioxide (CO2), methane (CH4), nitrous oxide (N2O), tropospheric ozone, and halocarbons. ODSs and their substitutes make up a large fraction of the halocarbons in today’s atmosphere. Increases in the abundances of these gases from human activities cause more outgoing radiation to be absorbed, which further warms the atmosphere and surface. This change in the Earth’s radiative balance caused by human activities is called a radiative forcing of climate or, more simply, a climate forcing. The magnitude of this energy imbalance is usually evaluated at the top of the troposphere (tropopause) and is expressed using units of watts per square meter (W/m2). The potential for climate change increases as this radiative forcing increases.

Figure Q18-1 Radiative forcing of greenhouse gases

A summary of radiative forcings in 2011 resulting from the increases in the principal long-lived and short-lived greenhouse gases during the Industrial Era is shown in Figure Q18-1. All forcings shown relate to human activities. Positive forcings generally lead to warming and negative forcings lead to cooling of Earth’s surface. Climate forcings also lead to other changes, such as in precipitation patterns and extreme weather events. International climate assessments conclude that much of the observed surface warming and changes in other climate parameters over the last decades is due to increases in the abundances of carbon dioxide and other greenhouse gases caused by human activities.

Carbon dioxide, methane, and nitrous oxide. All three of these greenhouse gases have both human and natural sources. The accumulation of carbon dioxide during the Industrial Era represents the largest climate forcing related to human activities. Carbon dioxide concentrations continue to increase in the atmosphere primarily as the result of burning fossil fuels (coal, oil, and natural gas) for energy and transportation, as well as from cement manufacturing. The global mean atmospheric abundance of carbon dioxide is now more than 40% above what it was in preindustrial times (before 1750), approaching values of 400 parts per million by volume (ppmv). Carbon dioxide is considered a long-lived gas, since a signi cant fraction remains in the atmosphere 100–1000 years after emission.

Methane is a short-lived climate gas (atmospheric lifetime of about 10 years). Human sources include livestock, fossil fuel extraction and use, rice agriculture, and land lls. Natural sources include wetlands, termites, and oceans. Methane has increased by about 150% since preindustrial times.

Nitrous oxide is a long-lived climate gas (atmospheric lifetime of about 110 years). The largest human source is agricultural activities, especially related to fertilization. Microbial processes in soils that are part of natural biogeochemical cycles represent the largest natural source. In the stratosphere, nitrous oxide is the principal source of reactive nitrogen species, which participate in ozone destruction cycles (see Q2 and Q7). Nitrous oxide has increased by about 20% since preindustrial times.

Figure Q18-2 Halocarbons and radiative forcing of climate

Halocarbons. Halocarbons in the atmosphere contribute to both ozone depletion and climate change. As used here, halocarbons represent those gases containing chlorine, bromine, or uorine atoms that are now controlled substances under the Montreal Protocol or the Kyoto Protocol. ODSs are the halocarbons controlled under the Montreal Protocol. HFC substitute gases, per uorocarbons (PFCs), and sulfur hexa uoride (SF6) are controlled under the Kyoto Protocol. In 2011, the halocarbon contribution to climate forcing was 0.36 W/m2, which is the fourth largest forcing following carbon dioxide, methane, and tropospheric ozone (see Figure Q18-1). The contributions of individual halocarbon gases are highlighted in Figure Q18-2. Within the halocarbons, CFCs contribute the largest percentage (73%) to the 2011 climate forcing. HCFCs, the intermediate-term ODS substitutes, make the next largest contribution (14%). The atmospheric abundance of HFCs, the longer-term ODS substitutes, contributes 5% to the 2011 halocarbon climate forcing.

The large contribution of the CFCs has been gradually decreasing following the decline in their atmospheric abundance and is expected to further decrease (see Figure Q16-1). Based on their long lifetimes, CFCs will still make a signi cant contribution, and most likely the largest ODS contribution, to halocarbon climate forcing at the end of the 21st century. Halocarbons controlled under the Kyoto Protocol (HFCs, PFCs, and SF6) represent about 8% of halocarbon climate forcing in 2011. With the projected growth of HFC production and consumption in developing nations, this percentage contribution is expected to continue to increase in the coming decades (see Figure Q19-2).

Stratospheric and tropospheric ozone. Ozone in both the stratosphere and the troposphere absorbs infrared radiation emitted from Earth’s surface, trapping heat in the atmosphere. Stratospheric ozone also signi cantly absorbs solar radiation. As a result, increases or decreases in stratospheric or tropospheric ozone induce a climate forcing and, therefore, represent direct links between ozone and climate. Air pollution from human activities has led to increases in global tropospheric ozone, causing a positive radiative forcing (warming) since preindustrial times (see Figure Q18-1). The large uncertainty in tropospheric ozone forcing re ects the difficulty in quantifying tropospheric ozone trends and in modeling the complex production and loss processes that control its abundance. On the other hand, rising ODS abundances in the atmosphere since the middle of the 20th century have led to decreases in stratospheric ozone and in the amounts of stratospheric ozone transported into the troposphere, causing a negative radiative forcing (cooling) since preindustrial times. The negative radiative forcing resulting from stratospheric ozone depletion will diminish in the coming decades as ODSs are gradually removed from the atmosphere.

Stratospheric ozone depletion cannot be a principal cause of present-day global climate change for two reasons: rst, the climate forcing from ODS-induced ozone depletion acts to cool Earth’s surface. Second, the total forcing from other long-lived and short-lived gases in Figure Q18-1 is positive (leading to warming). The total forcing from these other greenhouse gases (carbon dioxide, methane, halocarbons, and nitrous oxide) is the principal cause of observed climate change.

Ozone Depletion Potentials and Global Warming Potentials. An important way of comparing the in uence of individual emissions of halocarbons on ozone depletion and climate change is to use Ozone Depletion Potentials (ODPs) and Global Warming Potentials (GWPs). The ODP and GWP are the effectiveness of an emission of a gas in causing ozone depletion and climate forcing, respectively, relative to a reference gas (see Table Q7-1). The principal halocarbon gases are contrasted with each other in Figure Q18-3. The ODP of CFC-11 and the GWP of carbon dioxide are assigned reference values of 1. The CFCs and carbon tetrachloride all have ODPs near 1, indicating comparable effectiveness in causing ozone depletion per mass emitted. The principal halons have ODPs greater than 7, making them the most effective ozone-depleting substances per mass emitted. HFCs have ODPs of zero since they cause no ozone depletion (see Q7).

All halocarbons have non-zero GWPs and, therefore, contribute to climate forcing. The GWP does not correspond strongly with the ODP of a gas because these quantities depend on different chemical and physical properties. For example, while HFC-134a does not destroy ozone (ODP equal 0), each gram emitted is 1,360 times more effective than a gram of carbon dioxide in causing climate forcing. The future selection of speci c HFCs as ODS substitutes or for use in new global applications will have important consequences for climate forcing. When these HFCs are eventually released to the atmosphere, their contribution to climate forcing will depend on their GWPs, which could vary over a wide range (2 to 13,000).

Figure Q18-3 ODPs and GWPs

Montreal Protocol regulations have led to reductions in CFC emissions and increases in HCFC emissions (see Q16). As a result of these actions, the total radiative forcing from ODSs stopped increasing and is now slowly decreasing (see Q19). Overall halocarbon radiative forcing, however, is slowly increasing because of growing contributions from HFCs, PFCs, and SF6. It is important to note that, despite having a GWP that is small in comparison to many other halocarbons and other greenhouse gases, carbon dioxide is the most important greenhouse gas related to human activities because its emissions are large and its atmospheric abundance is far greater than the abundances of other emitted gases.

The Antarctic ozone hole and Southern Hemisphere climate. While stratospheric ozone depletion is not the principal cause of global climate change, the reoccurring Antarctic ozone hole has indirectly contributed to observed changes in Southern Hemisphere summertime surface climate through circulation changes. These new research ndings are explained in more detail below.

The Antarctic Ozone Hole and Southern Hemisphere Surface Climate

Links between stratospheric ozone depletion and changes in surface climate were rst found in research studies in the early 2000s, based on both observations and models. While increasing greenhouse gases (such as carbon dioxide, methane, and nitrous oxide) are the primary drivers of global climate change, the Antarctic ozone hole, which reoccurs every spring since the early 1980s, was shown to contribute indirectly to observed changes in Southern Hemisphere surface climate during summer due to its effects on atmospheric circulation.

The severe springtime ozone depletion over the Antarctic leads to a strong cooling of the polar lower stratosphere persisting into early summer. This increases the temperature contrast between the tropics and the polar region, which strengthens stratospheric winds. For reasons that are not well understood but are consistently reproduced in models, this leads to a poleward shift of tropospheric circulation features including the tropical Hadley cell (which determines the location of the subtropical dry zones) and the midlatitude jet (which is associated with weather systems). There is some evidence from both models and observations that subtropical and midlatitude summer precipitation patterns have been affected by these changes. The observed wind changes over the Southern Ocean have also likely driven signi cant changes in ocean currents. Model studies suggest that long-lived greenhouse gases that cause climate change exacerbate this shift in summertime Southern Hemisphere tropospheric circulation, but that ozone depletion has been the dominant contributor to the changes over the last few decades.

During the 21st century, ozone recovery is expected to lead to the reversal of the above climate impacts driven by the ozone hole. This reversal will result in some cancellation of future Southern Hemisphere circulation changes driven by greenhouse gas increases. The extent of such a cancellation depends on the greenhouse gas emissions assumed in future climate projections. The Southern Hemisphere surface climate response to ozone depletion in other seasons is weaker. No such links between ozone depletion and regional climate change have been observed for the Northern Hemisphere.

Yes. Many ozone-depleting substances are also potent greenhouse gases that contribute to climate forcing when they accumulate in the atmosphere. Montreal Protocol controls have led to a substantial reduction in the emissions of ozone-depleting substances (ODSs) over the last two decades. These reductions, while protecting the ozone layer, have the additional bene t of reducing the human contribution to climate change. Without Montreal Protocol controls, the contribution to climate forcing from annual ODS emissions could now be 10-fold larger than its present value, which would be a signi cant fraction of the climate forcing from current carbon dioxide (CO2) emissions. Increases in ODS substitute gases, which are also greenhouse gases, could offset much of this climate bene t by substantially contributing to human- induced climate forcing in the coming decades.

The success of the Montreal Protocol in controlling the production and consumption of ozone-depleting substances (ODSs) has protected the ozone layer (see Q15). The resulting reductions in atmospheric abundances of ODSs also reduced the human in uence on climate because all ODSs are greenhouse gases (see Q18). By protecting both ozone and climate, the Montreal Protocol has provided a dual bene t to society and Earth’s ecosystems. In the following, the dual bene t of the Montreal Protocol is highlighted by considering long-term baseline and world-avoided scenarios of ODS emissions that use Ozone Depletion Potentials (ODPs), Global Warming Potentials (GWPs), equivalent effective stratospheric chlorine (EESC), and the radiative forcing of climate. However, the climate bene t of the Montreal Protocol may be signi cantly offset by increasing future use of ODS substitute gases such as HFCs.

Baseline ODS scenarios. The baseline scenarios of past and future ODS emissions presented here include the emissions of principal halogen source gases assuming compliance with the Montreal Protocol provisions. They are constructed from (1) historical annual production and consumption of individual ODSs reported to the Montreal Protocol, (2) projected annual production and consumption of ODSs for future years based on provisions of the Protocol, (3) estimates of ODS banks, (4) atmospheric observations of ODSs and some naturally occurring halogen source gases, such as methyl chloride (CH3Cl), and (5) weighting factors related to ozone depletion and climate change.

In forming two of the baseline scenarios shown in Figure Q19-1 (upper panels), the emissions of each gas are added together after being weighted (multiplied) by the Ozone Depletion Potential (ODP) or the Global Warming Potential (GWP) of the respective gas (see Q18 and Table Q7-1). In the ODP-weighted scenario, the sum of emissions is expressed as CFC-11-equivalent emissions because CFC-11 is the reference gas, with an assigned ODP value of 1. For example, in the sum, 1 kg of halon-1211 emissions is added as 6.9 kg of CFC-11-equivalent emissions because the ODP of halon-1211 is 6.9. Similarly, the GWP-weighted sum is expressed as CO2-equivalent emissions because CO2 is the reference gas, with an assigned GWP of 1. For example, in the sum, 1 kg of carbon tetrachloride emissions is added as 1730 kg of CO2-equivalent emissions because the GWP of carbon tetrachloride is 1730.

Figure Q19-1 Montreal Protocol protection of ozone and climate

World-avoided ODS scenarios. The baseline scenario of ODS emissions can be contrasted with a scenario of ODS emissions that the world has avoided by agreeing to the Montreal Protocol (see Figure Q19-1). These world-avoided emissions are estimated by assuming that emissions of ODSs in the baseline scenario increase beyond 1987 values with a 2 or 3% annual growth rate. These growth rates are consistent with the strong market for ODSs in the late 1980s that included a wide variety of current and potential applications and that had potential for substantial new growth in developing countries.

CO2 emission scenarios. Long-term CO2 emission scenarios are also shown for comparison, as derived from past and projected CO2 emissions, because CO2 is the principal greenhouse gas related to human activities. The projected CO2 emissions have high and low scenarios that are derived using different basic assumptions about future economies, technical progress, and societal decisions.

ODP-weighted emissions scenarios. The ODP-weighted emissions in the ODS baseline scenario are a measure of the overall threat to stratospheric ozone from ODSs (see Figure Q19-1, upper left panel). When ODP-weighted emissions increase (decrease) in a given year, more (less) ozone will be destroyed in future years. ODP-weighted emissions increased substantially in the baseline scenario between 1960 and 1987, the year the Montreal Protocol was signed (see Figure Q19-1 and Q0-1). After 1987, ODP-weighted emissions began a long and steady decline to present-day values. The decline in emissions is expected to continue, causing the atmospheric abundances of individual ODSs to decrease (see Figure Q16-1). The reduction in ODP-weighted emissions from the 1987 value is a conservative measure of the annual emissions avoided by the Montreal Protocol since 1987 and, hence, of the success of the Montreal Protocol in protecting the ozone layer.

Annual ODP-weighted emissions in the world-avoided scenario are about double the 1987 values by 2020. The annual differences between the world-avoided emissions and the baseline scenario (blue shaded region in Figure Q19-1) provide reasonable upper limits to the ODP-weighted emissions avoided by the Montreal Protocol each year since 1987.

GWP-weighted emissions scenarios. The GWP-weighted emissions in the ODS baseline scenario are a measure of the overall threat to climate from ODSs (see Figure Q19-1, upper right panel). As ODS emissions accumulate in the atmosphere, their climate forcing contribution increases. The long-term changes in the GWP-weighted scenario are very similar to those in the ODP-weighted scenario. Both show an increase before 1987 and decrease afterwards. The similarity follows from the predominant role that CFC-11 and CFC-12 emissions play in ozone depletion and climate forcing from ODSs. The reduction in GWP-weighted emissions since 1987 is a conservative measure of the substantial success of the Montreal Protocol in reducing the potential for climate change from human activities. The annual differences since 1987 between the world-avoided emissions and the baseline scenario (blue shaded region in Figure Q19-1) provide reasonable upper limits to the GWP-weighted emissions avoided by the Montreal Protocol each year since 1987.

The climate protection calculated using differences between world-avoided emissions and the baseline scenario has two offsetting effects. The rst is the additional ozone depletion that would be caused by world-avoided ODS emissions. Ozone depletion offsets ODS climate forcing because a greenhouse gas (ozone) is being removed from the atmosphere in response to ODS emissions (see Q18). The second effect is the increase in emissions of HFC substitute gases that occurred in response to ODS reductions from Montreal Protocol controls. More HFCs in the atmosphere have the potential to offset much of the gain in climate protection from ODS reductions (see Figure 19-2) because HFCs are also greenhouse gases (see Q18).

Figure Q19-2 Implications of HFC usage for climate

The combined magnitude of these offsets in 2010, for example, is about 30% of the difference between the baseline and world-avoided scenarios. The resulting
net GWP-weighted emission reduction in 2010 is about 9.7–12.5 gigatonnes CO2-equivalent per year. In contrast, the annual emissions reduction target adopted by the Kyoto Protocol during its rst commitment period (2008–2012) is estimated as 2 gigatonnes CO2-equivalent per year (see Figure Q19-1). The reductions are expected to result from controlling the Kyoto Protocol basket of gases that includes HFCs and does not include ODSs (see Q18). As a result, the upper limit for the net reduction in annual GWP-weighted emissions achieved by the Montreal Protocol in 2010 is 5- to 6-fold larger than the Kyoto Protocol target.

Annual GWP-weighted emissions of ODSs were a large percentage (about 20–40%) of CO2-baseline emissions between 1960 and 1987 (see Figure Q19-1). Thereafter, this percentage has steadily decreased and is projected to reach 2–3% by 2020. This projection stands in sharp contrast to the world-avoided scenario, in which the percentage increases to 40–75% of CO2-baseline emissions by 2020.

EESC scenarios. The EESC scenario in Figure Q19-1 (lower left panel) provides a measure of the year-to-year potential of the atmospheric abundances of ODSs
to destroy stratospheric ozone. Changes in historical and projected atmospheric emissions of ODSs cause changes in their atmospheric abundances. The derivation of EESC from ODS atmospheric abundances is discussed in Q16 and similar EESC baseline scenarios are shown in Figures Q14-1, Q15-1, and Q16-1 for different time intervals. An increase in ODP-weighted emissions always leads to some increase in EESC in the years following the emissions. When ODS-weighted emissions decreased after 1987, EESC did not proportionally decrease because of the long atmospheric lifetimes of the principal ODSs. In Figure Q19-1, for example, EESC reached its peak nearly a decade after the peak in ODP-weighted emissions, and by 2010 the decrease in EESC from its peak value was only about 10%, compared to the 70% decrease in ODP-weighted emissions achieved by 2010.

Radiative forcing of climate. The radiative forcing derived for the ODS baseline scenario in Figure Q19-1 (lower right panel) provides a measure of the year-to-year contribution to climate forcing from atmospheric ODS abundances. The radiative forcing of an ODS is proportional to its radiative ef ciency (that is, the radiative forcing per abundance change) and the net increase in its atmospheric abundance during the Industrial Era. Increases in abundance up to the present are derived from atmospheric observations. Future abundances rely on projected emissions and atmospheric lifetimes of each gas. In Figure Q19-1, radiative forcing due to ODSs increases smoothly from 1960 onward, peaks in 2003, and decreases very gradually in subsequent years. Radiative forcing responds to ODS emission reductions in a manner similar to EESC, with the current slow decline attributable to the two principal contributing gases, CFC-11 and CFC-12, and their long atmospheric lifetimes (50–100 years).

The differences in ODS climate forcing between the world-avoided and baseline scenarios are offset by additional ozone depletion and HFC emissions in a manner similar to that noted above for differences in GWP-weighted emissions. After accounting for these two offsets, the climate forcing due to ODSs in the world-avoided scenario is approximately 70% higher than that in the baseline scenario in 2010 and approximately 30% of that due to CO2.

The considerable contributions that ODSs could have made to climate forcing, if not controlled by the Montreal Protocol, attests to their potency as greenhouse gases. ODSs had negligible atmospheric abundances 50–60 years ago and, as a group, represent chlorine amounts that currently are about 100,000 times less abundant in the atmosphere than CO2.

Implications of HFC usage for climate. HFCs are currently principal substitute gases for ODS uses. Many of these HFCs are also potent greenhouse gases. In
the coming decades, the projected growth in HFC emissions has the potential to offset much of the climate bene t gained from the Montreal Protocol controls, depending on emissions scenarios. An average future projection scenario based on current production and consumption patterns and projected economic growth suggests that increasing HFC demands could result in GWP-weighted emissions of around 5 gigatonnes CO2-equivalent per year by 2050 (see Figure Q19-2), primarily in developing nations. The 2050 value is comparable to half of the peak in GWP-weighted ODS emissions in 1987 (see Figure Q19-1). The resulting radiative forcing from these projected emissions would reach levels of the early 2000s and contribute almost 10% of that due to CO2 to the total anthropogenic climate forcing by 2050.

This projection assumes that HFC application demand in developing countries would be met using the same suite of HFCs currently used in developed countries. If future HFC demand is met instead with lower-GWP substances, the 2050 projection would be substantially reduced.

Increasing the bene ts of the Montreal Protocol into the future. Fewer control options are available to increase the dual bene t of the Montreal Protocol beyond 2020 because the most effective and abundant ODSs have already been phased out under Montreal Protocol provisions (see Q16). The most recent Montreal Protocol action increased ozone and climate protection by accelerating the phase-out of HCFCs (Montreal in 2007) (see Q15). This provision is expected to reduce total GWP-weighted emissions of HCFCs by about 50% between 2010 and 2050, corresponding to about 18 gigatonnes of CO2-equivalent emissions.

Ozone and climate protection could be further enhanced with Montreal Protocol provisions that increase the effectiveness of capturing and destroying ODSs contained in halon, CFC, and HCFC banks, namely, those ODSs currently being used in refrigeration, air conditioning, and re protection equipment, or stockpiled for servicing long-term applications. Current banks are projected to contribute more to ozone depletion in the coming decades than the limited future production of ODSs allowed by the Montreal Protocol provisions.

The available options, such as cutting the remaining production of HCFCs, halting emissions of CCl4 and quarantine or pre-shipment CH3Br, as well as destroying halon, CFC, and HCFC banks, would together advance the return of the ozone layer to 1980 levels by 11 years.

Substantial recovery of the ozone layer from the effects of ozone- depleting substances (ODSs) is expected near the middle of the 21st century, assuming global compliance with the Montreal Protocol. Recovery will occur as ODSs and reactive halogen gases in the stratosphere decrease in the coming decades. In addition to responding to ODSs, future ozone amounts will increasingly be in uenced by expected changes in climate. The impacts of future climate change on the ozone layer will vary between the tropics, midlatitudes, and polar regions, and strongly depend on future emissions of carbon dioxide (CO2), methane (CH4), and nitrous oxide (N2O). During the long recovery period, large volcanic eruptions could temporarily reduce global ozone amounts for several years.

Substantial recovery from the depletion of global and polar ozone caused by ODSs is expected in the later decades of this century. The recovery follows from the success of the Montreal Protocol in reducing the global production and consumption of ODSs. Now, the atmospheric abundances of most major ODSs and the associated values of equivalent effective stratospheric chlorine (EESC) are in decline (see Q16). In contrast to the diminishing role of ODSs, changes in climate are expected to have an increasing in uence on future ozone. Climate change is driven by the projected growth in greenhouse gas abundances, primarily carbon dioxide (CO2), methane (CH4), and nitrous oxide (N2O), which will lead to changes in temperatures, chemistry, and the circulation of the stratosphere, all of which affect ozone. Chemistry-climate models can be used to project how ozone is expected to respond to changes in ODSs and climate in different geographical regions during the recovery period. Global events, such as major volcanic eruptions and geoengineering actions, may also in uence future ozone levels.

Figure Q20-1 Simulations of ozone depletion

Using chemistry-climate models. Projections of total ozone presented here are based on the results from a group of chemistry-climate models that take into account the in uences of changes in ODSs and climate. These models show how changes in ozone are expected to vary across global regions by evaluating the complex interactions of the processes that control ozone and climate involving radiation, chemistry, and transport. Required model inputs include, for example, historical and projected concentrations of ODSs, CO2, CH4, and N2O, and changes in sea surface temperatures. The results from chemistry-climate model simulations are used to identify aspects of the models that are particularly important for future ozone abundances. For example, model projections for the coming decades show a strengthening in the atmospheric circulation that brings air to the stratosphere in the tropics, moves air poleward into both hemispheres, and then returns it to the troposphere at middle to high latitudes. These circulation changes will signi cantly alter the global distribution of ozone and the atmospheric lifetimes of ODSs and other long-lived gases. Also, while Earth’s surface is expected to continue to warm in response to positive radiative forcing from greenhouse gases, the stratosphere is expected to continue to cool. A colder upper stratosphere leads to increases in ozone because lower temperatures decrease the effectiveness of gas-phase ozone-loss reactions. Finally, CH4 and N2O are both involved in the chemistry that determines stratospheric ozone levels.

Simulating recent ozone changes. Comparisons of model results with observations help con rm the causes of ozone depletion and increase con dence in model projections of future ozone amounts. Two important measures of ozone, global total ozone and minimum total ozone values in the Antarctic (available since 1980), are compared to a group of chemistry-climate model simulations in Figure Q20-1. Both ozone measures show substantial depletion since 1980. The average model value follows the observed decline in both ozone measures, indicating that the main processes involved in ozone depletion are reasonably well represented by the models. Some differences between the simulations and observations can be explained by unusual meteorological conditions, volcanic eruptions, changes in solar activity, or other natural in uences, which are not all fully accounted for in the various models used here. In Figure Q20-1, global total ozone has increased in the last 20 years, while Antarctic minimum ozone values have been relatively constant. The global total ozone increases cannot be attributed solely to reductions in ODSs that began in the 1990s, because total ozone has also recovered from additional depletion caused by the 1991 Mt. Pinatubo eruption (see Q14).

Equivalent stratospheric chlorine (ESC) projections. Equivalent stratospheric chlorine (ESC) values are also projected in chemistry-climate models to represent how the potential for reactive halogen gases (see Q8) to destroy ozone varies over time. ESC is similar to equivalent effective stratospheric chlorine (EESC) (see Q16) in that both represent a weighted sum of reactive chlorine and bromine gases in the stratosphere. The weighting in each accounts for the greater effectiveness of bromine atoms in destroying ozone. ESC can be calculated more precisely and comprehensively than EESC because the chemistry-climate models can derive ESC as a function of altitude, latitude, longitude, and time. The ESC calculations are based on the history and projections of ODS surface abundances and the chemical and transport processes that control (1) the conversion of ODSs to reactive halogen gases, (2) the distribution of reactive halogen gases in the global stratosphere, and (3) their ultimate removal from the stratosphere.

Figure Q20-2 Long-term changes in ozone and equivalent stratospheric chlorine (ESC)

The long-term changes in ESC in the global and regional analyses shown in Figure Q20-2 are highly similar. In all regions, changes in ESC from 1960 values increase smoothly with time, reach a peak near the end of the 20th century, and decrease gradually until the end of the 21st century. Values at the end of the 21st century are approaching those in 1960, indicating that ODSs are largely removed from the stratosphere by that time. ESC returns to 1980 values several decades sooner than to 1960 values because of the slow rate of decline.

Peak ESC values around the year 2000 are highest in polar regions and lowest in the tropics. In the tropics, stratospheric air has only recently been transported from the troposphere, with the result that only a small fraction of ODSs has undergone conversion to reactive halogen gases (see Q8). In polar regions, the fraction is much larger because stratospheric air requires several years on average to journey from its entry point in the tropics to the polar lower stratosphere. During this time, a much larger fraction of ODSs undergoes conversion to reactive halogen gases.

Long-term total ozone projections. Total ozone changes derived from chemistry- climate models, as referenced to 1960 values, are shown in Figure Q20-2. The range of values from the group of models is included in the gure as one measure of the uncertainty in the model projections. Total ozone changes in the different regions are described as follows:

  • Antarctic. Total ozone changes are largest in the Antarctic region in springtime (October). Chemistry-climate models show that ODSs are the predominant factor in Antarctic ozone depletion in the past and in the coming decades. Changes in climate parameters have a smaller role. As a result, total ozone changes mostly mirror changes in ESC: as ESC increases, ozone proportionately decreases; as ESC decreases, ozone proportionately increases. Antarctic total ozone is projected to return to 1980 levels after midcentury and later than in any other region. Meteorological variability in the Antarctic in late winter/early spring when ozone depletion occurs causes a large range in the model projections.
  • Arctic. Total ozone changes in the Arctic region in springtime (March) are considerably smaller than in the Antarctic. In contrast to the Antarctic, ozone changes do not closely mirror changes in ESC. After midcentury, Arctic total ozone increases to values above those that would be expected from ESC reductions alone because of the strengthening of the atmospheric circulation and the enhanced stratospheric cooling associated with increases in CO2. By 2100 Arctic total ozone is projected to be well above both 1960 and 1980 values. The large range in projections compared to other nonpolar regions is due to greater meteorological variability, as noted for the Antarctic. Arctic total ozone is projected to return to 1980 levels between 2020 and 2035, which is two to three decades before ESC returns to 1980 levels, and to continue to increase until the end of the century.
  • Northern and southern midlatitudes. The annual averages of total ozone changes in midlatitudes are much smaller than the springtime losses in polar regions. Both midlatitude regions resemble the Arctic, with total ozone returning to 1960 and 1980 values much sooner than ESC. In the northern midlatitudes, the models predict a return of total ozone to 1960 values in 2030, whereas ESC requires the full century to return close to 1960 values. Total ozone in southern midlatitudes changes in a manner similar to that in the northern midlatitudes. Differences are that in southern midlatitudes, the 1960 return date for total ozone is somewhat later (2055) and the maximum ozone depletion observed near 2000 is greater. Both aspects re ect the in uence of Antarctic polar ozone depletion, with stratospheric air low in ozone being transported to southern midlatitudes in spring after the breakup of the polar vortex (see Q11). The more rapid return of total ozone in both regions compared with ESC again is the result of the in uence of climate-induced changes in transport and upper stratospheric temperatures in the model projections. The model range is signi cantly less in midlatitude annual averages than in the polar springtime averages. After returning to either 1960 or 1980 values, total ozone continues to increase in both hemispheres and signi cantly exceeds those values by 2100.
  • Tropics. Total ozone changes in the tropics are smaller than in any other region. Ozone is less sensitive to ESC in the tropical stratosphere because of the dominant roles of production and transport in controlling ozone and the low amounts of ESC available in this region. In contrast to other regions, chemistry-climate models project total ozone to remain below 1960 values throughout the 21st century. Total ozone gradually returns to 1980 values in 2040, peaks around 2060, and decreases again until the end of the century. The different total ozone behavior in the tropics is due to the changing balance between ozone increases in the upper stratosphere and decreases in the lower stratosphere. The ozone increases in the upper stratosphere occur from declining ODS abundances and decreases in temperatures caused by increasing greenhouse gases, primarily CO2. In the lower stratosphere, ozone is reduced because the strengthening of the stratospheric circulation reduces the time for ozone production before air leaves the tropics. These circulation changes also in uence the Arctic and midlatitude regions as noted above.
  • The globe. The annual averages of global total ozone are projected to return to 1960 levels around the middle of the century (2040–2080) while ESC returns to 1960 values near the century’s end. The return dates for 1980 values are substantially earlier. Chemistry-climate model analysis suggests that this early return of total ozone is primarily a result of upper stratospheric cooling. The larger total ozone changes shown for polar regions in springtime do not signi cantly in uence global total ozone values because these values are annual averages and polar regions are a small geographical fraction of the globe.

Future ultraviolet radiation. Projections of long-term changes in total ozone can be used to estimate long-term changes in solar ultraviolet radiation reaching Earth’s surface (see Q17). The UV-B component of ultraviolet radiation increases as total ozone decreases. Based on the global total ozone projections, clear-sky UV-B radiation is expected to decrease below 1960 values by the end of the century because ozone is above 1960 values in the later decades. The latitude regions expected to experience increased ultraviolet radiation levels are the Antarctic and the tropics, where total ozone remains lower than 1960 values until the end of the century. Large increases or decreases in surface erythemal radiation away from 1960 values are expected to lead to adverse effects on human and ecosystem health.

Greenhouse gas emission scenarios and impact on future ozone. The climate impacts on future ozone discussed above have been derived using a particular greenhouse gas emission scenario. Emission scenarios specify the amounts of greenhouse gases emitted into the atmosphere in the future and are constructed based on socio-economic assumptions on future population growth, economical developments, technology innovation, and political decisions. Future emissions are highly uncertain and therefore various emission scenarios are designed to cover a wide range of possible future outcomes. Chemistry-climate model simulations are run using the different emission scenarios to determine the range of impacts climate change may have on the ozone layer. Such simulations show that the evolution of ozone and the resulting surface UV levels in different regions of the world are strongly dependent on the mixture and relative abundances of greenhouse gases that will be used.

Figure Q20-3 Global total ozone changes in response to increasing greenhouse gases

Simulations with a simpli ed model using increases in CO2, N2O, and CH4 separately illustrate the reasons for these dependencies (see Figure Q20-3). The simulation using increases in CO2 only shows that the main effect of CO2 is to increase global ozone in the future. The ozone increases are explained mainly by decreases in stratospheric temperatures. As detailed above, lower temperatures in the upper stratosphere slow down the chemical gas-phase reactions that lead to ozone depletion. Note that circulation changes only play a minor role for global mean ozone since decreases in total ozone in the tropics and increases at higher latitudes largely cancel. Simulations using increases in CH4 and N2O separately show that the main effect of CH4 is to increase ozone, while that of N2O is to decrease ozone in the future. The reason for this behavior is that N2O and CH4 are source gases of reactive gases (NOx and HOx) that strongly affect stratospheric ozone chemistry. The reactive gases produced from N2O lead to stratospheric ozone depletion, while those from CH4 lead to the production of stratospheric ozone.

Volcanoes and geoengineering. Other factors not included in chemistry-climate models can potentially affect future total ozone amounts. Explosive volcanic eruptions have temporarily reduced global total ozone in the past (see Q14). Similar eruptions, especially in the early decades of this century when ESC values are highest, are also expected to reduce total ozone for a few years. Volcanic eruptions are an additional source of uncertainty not included in the ozone projections in Figures Q20-2 and Q20-3.

Several geoengineering methods have been proposed to reduce climate forcing from human activities. A widely discussed proposal is the enhancement of sulfate aerosols in the stratosphere from direct injections of sulfur or sulfuric acid. The expected response is a cooling of the climate system from increased aerosol scattering of sunlight, similar to that observed after some explosive volcanic eruptions. The required injections, if sustained over many years, are likely to have unintended consequences such as reductions in total ozone amounts and changes in stratospheric temperatures and circulation.