QUESTIONS& ANSWERSABOUT THE OZONE LAYER
2022 UpdateWorld Meteorological Organization
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United Nations Environment Programme
Ozone Secretariat
P.O. Box 30552
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National Oceanic and Atmospheric Administration
14th Street and Constitution Avenue NWHerbert C. Hoover Building, Room 5128
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B-1049 Bruxelles
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Published April 2023
ISBN: 978-9914-733-98-3
Ross J. Salawitch (Lead Author), Laura A. McBride, Chelsea R. Thompson, Eric L. Fleming, Richard L. McKenzie, Karen H.
Rosenlof, Sarah J. Doherty, David W. Fahey, Twenty Questions and Answers About the Ozone Layer: 2022 Update, Scientific
Assessment of Ozone Depletion: 2022, 75 pp., World Meteorological Organization, Geneva, Switzerland, 2023.
This report is available on the internet at the following locations:
https://ozone.unep.org/science/assessment/sap
https://www.csl.noaa.gov/assessments/ozone/2022
Note: Figures from this report are in the public domain and may be used without permission. Attribution to this document is encouraged.
Cover image: Artistic rendering of the Aura spacecraft in Earth orbit.
Aura, launched in 2004 by the U.S. National Aeronautics and Space
Administration (NASA), carries four instruments: the Microwave Limb
Sounder (MLS), High Resolution Dynamics Limb Sounder (HIRDLS), Tropospheric Emission Spectrometer (TES) and Ozone Monitoring Instrument(OMI). The instrument suite has provided observations of stratosphericand tropospheric ozone abundances, as well as a wide range of otherchemical and physical properties of the atmosphere. The observations
have been used extensively to guide the scientific community’s understanding of the effects of human activity on Earth’s protective ozone layer.
Of particular importance have been the MLS observations of the chemical conditions in the ozone layer over Antarctica, including halogen
species (ClO and HCl) and nitric acid (HNO3 ) (see Fig. Q7-3). Image
credit: NASA.
Cover design by Chelsea R. Thompson

The Aura spacecraft in aclean-room, prior to launch.
Photo credit: NASA

Scientific Assessment of Ozone Depletion: 2022

TWENTY QUESTIONS & ANSWERSABOUT THE OZONE LAYER
2022 Update
Lead Author
Ross J. Salawitch
Coauthors
Laura A. McBride
Chelsea R. Thompson
Eric L. Fleming
Richard L. McKenzie
Karen H. Rosenlof
Sarah J. Doherty
David W. Fahey

World Meteorological Organization
United Nations Environment Programme
National Oceanic and Atmospheric Administration
National Aeronautics and Space Administration
European Commission

PREFACE
This document, Twenty Questions and Answers About the Ozone Layer: 2022 Update, is a component of the Scientific
Assessment of Ozone Depletion: 2022 report. The report is prepared quadrennially by the Scientific Assessment Panel(SAP) of the Montreal Protocol on Substances that Deplete the Ozone Layer*. The 2022 edition of the Twenty Questions
document is the fifth update of the original edition that appeared in the 2002 Assessment Report. The motivation behind
this scientific publication is to tell the story of ozone depletion, ozone-depleting substances and the success of the Montreal Protocol. The questions and answers format divides the narrative into topics that can be read and studied individually by the intended audience of specialists and non-specialists. The topics range from the most basic (e.g., What is ozone?)
to more recent developments (e.g., the Kigali Amendment to the Montreal Protocol). Each question begins with a shortanswer followed by a longer, more comprehensive answer. Figures enhance the narrative by illustrating key concepts
and results. This document is principally based on scientific results presented in the 2022 and earlier Assessment Reports
and has been extensively reviewed by scientists and non-specialists to ensure quality and readability.
We hope that you find this Twenty Questions and Answers edition of value in communicating the scientific basis of ozone
depletion and the success of the Montreal Protocol in protecting the ozone layer and climate.
David W. Fahey, Paul A. Newman, John A. Pyle, and Bonfils Safari
Co-Chairs of the Scientific Assessment Panel* https://ozone.unep.org/science/assessment/sap

CONTENTS
INTRODUCTION

1

SECTION I: OZONE IN OUR ATMOSPHERE

4

Q1.
Q2.
Q3.
Q4.

What is ozone, how is it formed, and where is it in our atmosphere?
Why do we care about atmospheric ozone?
How is total ozone distributed over the globe?
How is ozone measured in the atmosphere?

SECTION II: THE OZONE DEPLETION PROCESS
Q5.
Q6.
Q7.
Q8.
Q9.

How do emissions of halogen source gases lead to stratospheric ozone depletion?
What emissions from human activities lead to stratospheric ozone depletion?
What are the reactive halogen gases that destroy stratospheric ozone?
What are the chlorine and bromine reactions that destroy stratospheric ozone?
Why has an “ozone hole” appeared over Antarctica when ozone-depleting substances arepresent throughout the stratosphere?

SECTION III: STRATOSPHERIC OZONE DEPLETION
Q10.
Q11.
Q12.
Q13.

How severe is the depletion of the Antarctic ozone layer?
Is there depletion of the Arctic ozone layer?
How large is the depletion of the ozone layer outside of polar regions?
Do changes in the Sun and volcanic eruptions affect the ozone layer?

SECTION IV: CONTROLLING OZONE-DEPLETING SUBSTANCES
Q14.
Q15.

Are there controls on the production of ozone-depleting substances?
Has the Montreal Protocol been successful in reducing ozone-depleting substances in
the atmosphere?

SECTION V: IMPLICATIONS OF OZONE DEPLETION AND THE MONTREAL PROTOCOL
Q16.
Q17.
Q18.
Q19.

Does depletion of the ozone layer increase ground-level ultraviolet radiation?
Is depletion of the ozone layer the principal cause of global climate change?
Are Montreal Protocol controls of ozone-depleting substances also helping protect Earth’s climate?
How has the protection of climate by the Montreal Protocol expanded beyond the regulation ofozone-depleting substances?

SECTION VI: STRATOSPHERIC OZONE IN THE FUTURE
Q20. How is stratospheric ozone expected to change in the coming decades?

4
7
9
11

13
13
16
20
24
26

29
29
34
38
41

44
44
47

51
51
54
59
62

66
66

TERMINOLOGY72

CHEMICAL FORMULAE AND NOMENCLATURE

74

AUTHORS AND CONTRIBUTORS

75

INTRODUCTION
Ozone is present only in small amounts in the atmosphere. Nevertheless, ozone is vital to human well-being
as well as agricultural and ecosystem sustainability. Most of Earth’s ozone resides in the stratosphere, thelayer of the atmosphere that is more than 10 kilometers (6 miles) above the surface. About 90% of atmospheric ozone is contained in the stratospheric “ozone layer”, which shields Earth’s surface from harmful
ultraviolet radiation emitted by the Sun.

In the mid-1970s scientists discovered that some human-produced
chemicals could lead to depletion of the stratospheric ozone layer. The resulting increase in ultraviolet radiation at Earth’s surface
would increase incidents of skin cancer and eye cataracts, suppress the immune systems of humans, and also adversely affectagriculture as well as terrestrial and oceanic ecosystems.
Following the discovery of this environmental issue, researcherssought a better understanding of this threat to the ozone layer.
Monitoring stations showed that the abundances of gases that areozone-depleting substances (ODSs)1, such as chlorofluorocarbons(CFCs), were steadily increasing in the atmosphere. These trends
were linked to growing production and use of CFCs and other
ODSs for spray can propellants, refrigeration and air conditioning,
foam blowing, industrial cleaning, and other applications. 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 already occurring and to predict how much more might occur in the future.
By the mid-1980s observations of the ozone layer showed that
depletion was indeed occurring. The most severe ozone loss,
unexpected at the time of discovery, was found to be recurring
each springtime over Antarctica. The loss in this region is commonly called the “ozone hole” because the ozone depletion is solarge and localized. A thinning of the ozone layer also has been
observed over other regions of the globe, such as the Arctic andnorthern and southern midlatitudes.
The work of many scientists throughout the world has built a broadand solid scientific understanding of the ozone-depletion process.
With this foundation, we know that ozone depletion has been occurring and we understand why. Most importantly, we know that if
the most potent ODSs were to continue to be emitted and increasein the atmosphere, the result would be ever greater depletion of
the ozone layer.
In 1985, the world’s governments adopted the Vienna Convention
for the Protection of the Ozone Layer in response to the prospectof increasing ozone depletion. The Vienna Convention provided aframework through which nations agreed to take appropriate mea-sures to protect human health and the environment from activities
that harm the ozone layer, including cooperation on systematic
observations, research and exchange of information. In 1987, thisframework led to the Montreal Protocol on Substances that Deplete
the Ozone Layer (the Montreal Protocol), an international treaty
designed to control the production and consumption of CFCs andother ODSs. As a result of the broad compliance with the Montreal
Protocol and subsequent amendments and adjustments as well as
industry’s development and deployment of “ozone-friendly” substitutes to replace CFCs, the total global accumulation of ODSs in
the atmosphere has begun to decrease.
The replacement of CFCs has occurred in two phases: first via theuse of hydrochlorofluorocarbons (HCFCs) that cause considerably
less damage to the ozone layer compared to CFCs, and second by
the introduction of hydrofluorocarbons (HFCs) that do not depleteozone. In response, global ozone depletion has stabilized, and initial signs of recovery of the ozone layer are being observed. With
continued compliance, substantial recovery of the ozone layer is
expected 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. The
Montreal Protocol has also decreased the human drivers of global
warming, because many CFCs and HFCs are potent greenhousegases (GHGs).
The amendment and adjustment process is a vitally important aspect of the Montreal Protocol, allowing the protocol to evolve andaddress emerging issues as our scientific understanding matures.
The Protocol was amended or adjusted between 1990 and 2007
at meetings held in London, Copenhagen, Vienna, Beijing, andMontreal (see Q14). The most recent amendment was formulated
at the Meeting of the Parties of the Montreal Protocol held in Kigali,
Rwanda during October 2016. The Kigali Amendment phases
down future global production and consumption of some HFCs toprotect future climate, an important new milestone for the Montreal Protocol (see Q19). The Kigali Amendment was motivatedby projections of substantial increases in the global use of HFCs in
the coming decades. The control of HFCs under this amendment
marks the first time the Montreal Protocol has adopted controls
solely for the protection of climate.

Here and throughout, the term ozone-depleting substances (ODSs) refers to gases containing either chlorine or bromine that are released to the atmosphere as aresult of human activity and are controlled under Annexes A, B, C, or E of the Montreal Protocol.
1

Introduction

1

Introduction | Twenty Questions | 2022 Update
The protection of the ozone layer and climate under the Montreal
Protocol is a story of notable achievements: discovery, understanding, decisions, actions, and verification. It is a success 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, andthe Montreal Protocol is illustrated in Figure Q0-1.

success of the Montreal Protocol, the possible future of the ozone
layer, and the protection against climate change now provided by
the Kigali Amendment. Computer model projections show that
GHGs such as carbon dioxide (CO2 ), methane (CH4 ), and nitrous
oxide (N2O) will have a growing influence on global ozone in the
coming decades, and in some cases may exceed the influence ofODSs on ozone by the middle of this century, given the expected
future decline in the atmospheric abundance of ODSs.

To help communicate the broad understanding of the Montreal
Protocol, ODSs, and ozone depletion, as well as the relationship ofthese topics to GHGs and global warming, this component of the
Scientific Assessment of Ozone Depletion: 2022 report describes
the state of this science with 20 illustrated questions and answers.
The questions and answers address the nature of atmosphericozone, the chemicals that cause ozone depletion, how global andpolar ozone depletion occur, the extent of ozone depletion, theFor each question, a brief answer is first given in highlighted text;
an expanded answer then follows. The answers are based on theinformation presented in the 2022 and earlier Assessment reports
as well as other international scientific assessments. These reports
and the answers provided here were prepared and reviewed by alarge number of international scientists who are experts in different
research fields related to the science of stratospheric ozone andclimate2.

2

See Authors and Contributors at the end of this book.

Figure Q0-1. Stratospheric ozone depletion and policy milestones. This timeline highlights milestones related to the history of ozone
depletion. Events represent the occurrence of important scientific findings, the completion of international scientific assessments, andmilestones of the Montreal Protocol.
Scientific Milestones. A worldwide network of ground-based ozone measurement stations was initiated in 1957, as part of the International Geophysical Year. The number of atmospheric observations of ozone, CFCs, and other ODSs have increased substantially since
the early 1970s. The discovery of the Antarctic ozone hole in 1985 led to a major community effort that established, within a few years,
that this phenomenon was caused by the human release of various chlorine and bromine compounds. The total global emission of thesecompounds peaked in 1987, the year the Montreal Protocol was signed. The Nobel Prize in Chemistry in 1995 was awarded for researchthat identified, in 1974, the threat posed by CFCs to the global ozone layer. The graph at the bottom shows the history and near future ofannual total emissions of ozone-depleting substances (ODSs) combined with natural emissions of halogen source gases. The emissions,when weighted by their potential to destroy ozone, peaked in the late 1980s after several decades of steady increases. Between the late1980s and the present, emissions have decreased substantially as a result of the Montreal Protocol and its subsequent amendments andadjustments coming into force (see Q14). The total stratospheric halogen content peaked in the late 1990s, followed by a slow, steady decline (see Q15). In the mid-2010s scientists documented clear evidence that the rise in the abundance of upper stratospheric ozone, which
had begun in the late 1990s, was caused by the decline in the stratospheric chlorine loading driven by the Montreal Protocol. At the endof 2022, the stratospheric halogen content was 18% less than the peak value (see Q15).
International Scientific Assessments. The provisions of the Montreal Protocol and its amendments and adjustments have depended on information embodied in international scientific assessments of ozone depletion that have been produced periodically since 1989 under the
auspices of UNEP and WMO. These assessments incorporate new knowledge from ongoing observations, modeling studies, and analyses
into a report designed to reflect the latest scientific understanding of how human activity affects the ozone layer.
International Policy Milestones. The Montreal Protocol, which is built on the framework established by the Vienna Convention, was signed
on 16 September 1987. Under the Protocol, January 2010 marked the end of allowable production of CFCs and halons, with a few verysmall exemptions. In January 2013, a production and consumption freeze on HCFCs went into effect for all nations. In October 2016 theKigali Amendment brought the future production of HFCs under the auspices of the Montreal Protocol (see Q19).
(A megatonne = 1 million (106) metric tons = 1 billion (109) kilograms.)(Formally the term halogen refers to the elements fluorine, chlorine, bromine, iodine, and astatine that are in group 7A of the periodic table. Here andthroughout, unless otherwise specified, halogen refers to chlorine and bromine, since source gases containing either of these two halogens constitute theODSs controlled by the Montreal Protocol.)EESC: Equivalent effective stratospheric chlorine | IPCC: Intergovernmental Panel on Climate Change | ODS: Ozone-depleting substance| TEAP: Technology and Economic Assessment Panel of the Montreal Protocol | UNEP: United Nations Environment Programme | WMO:World Meteorological Organization
2

Introduction

2022 Update | Twenty Questions | Introduction

Stratospheric Ozone Depletion Science and Policy Milestones
Scientific Milestones
WMO World Ozone
Data Centre – 1960

Global Ozone
Observing
System – 1957

Backscatter Ultraviolet
satellite instrument
launched – 1970

Ozone hole
reported over
Antarctica – 1985

Peak global emissionsof ODSs – 1987

First surface
observations of
CFCs – 1973

Nobel Prize in Chemistry to Cruzten,
Start of continuous
Molina, and Rowland – 1995
global ozone satellite
record – 1978
Laboratory, field, and modeling studies
CFC ozone-depletion
show chlorine and bromine deplete polar
hypothesis – 1974
stratospheric ozone – 1985–1989

International Scientific
Assessments

WMO Atmospheric
Ozone – 1985

WMO The Stratosphere 1981:
Theory and Measurements

Stratospheric abundance of
EESC in midlatitudes down
18% from peak level – 2022

2014 Assessment reports clear
evidence that upper stratosphericozone increased as a result of the
ODS phaseout

UNEP/WMO Scientific Assessments of Ozone Depletion :1989 1991 19941998

WMO Report of the International
Ozone Trends Panel – 1988

2002

2006

2010

2014

2018

2022

IPCC/TEAP Special Report on
Safeguarding the Ozone Layer and
the Global Climate System – 2005

Developed countries eliminatedMontreal Protocol Amendments :90% of their HCFC production andLondon – 1990
consumption; developing countries
Copenhagen – 1992
freeze their HCFC production and
Global production
consumption – 2013of CFCs and halons
Vienna – 1995
Vienna Convention forends – 2010
Montreal – 1997
the Protection of theKigali Amendment
HCFC accelerated phaseout
Ozone Layer – 1985
Beijing – 1999
enters into force – 2019
(Montreal-2007 Adjustment)International Policy
Milestones

Global consumption of ODSs
reduced by 98% – 1986 to 2008

UNEP World Plan of Action
on the Ozone Layer – 1977

Megatonnes CFC-11-equivalent per year

Peak stratosphericabundance of chlorine
and bromine (EESC) in
midlatitudes – 1998

Montreal Protocol
universal ratification – 2009

Montreal Protocol on
Substances that Deplete the
Ozone Layer – 1987

Kigali Amendment on
the control of production
and consumption of some
HFCs adopted – 2016

1.5
Emissions of halogen source gases (ODSs + natural sources)(Weighted by Ozone Depletion Potential (ODP))
Based on atmospheric observations or industry data
Future projections

1.0

0.5
Natural emissions

0.0

1960

1970

1980

1990

2000

2010

2020

Introduction

3

Q1

What is ozone, how is it formed, andwhere is it in our atmosphere?

Ozone is a gas that is naturally present in our atmosphere. Each ozone molecule contains three atoms ofoxygen and is denoted chemically as O3 . Ozone is found primarily in two regions of the atmosphere. About
10% of Earth’s ozone is in the troposphere, which extends from the surface to about 10–15 kilometers (6–9
miles) altitude. About 90% of Earth’s ozone resides in the stratosphere, the region of the atmosphere between the top of the troposphere and about 50 kilometers (31 miles) altitude. The part of the stratospherewith the highest amount of ozone is commonly referred to as the “ozone layer”. Throughout the atmosphere,ozone is formed in multistep chemical processes that are initiated by sunlight. In the stratosphere, the process begins with an oxygen molecule (O2 ) being broken apart by ultraviolet radiation from the Sun. In the
troposphere, ozone is formed by a different set of chemical reactions that involve naturally occurring gases
as well as those from sources of air pollution.

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 inlaboratory experiments in the mid-1800s. Ozone’s presence inthe atmosphere was later discovered using chemical and opticalmeasurement 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 inconcentrated amounts. Electrical discharges are generally used to
produce ozone for industrial processes such as air and water purification and bleaching of textiles and food products.
Ozone location. Most ozone (about 90%) is found in the stratosphere, which begins about 10–15 kilometers (km) above Earth’s
surface and extends up to about 50 km altitude. The stratosphericregion with the highest concentration of ozone, between about 15
and 35 km altitude, is commonly known as the “ozone layer” (seeFigure Q1-2). The stratospheric ozone layer extends over the entire globe, with some variation in its altitude and thickness. Mostof 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. Tropospheric air is the “air we breathe”
and, as such, excess ozone in the troposphere has harmful consequences (see Q2).
Ozone abundance. Ozone molecules constitute a small fraction of
the gas molecules in the atmosphere. Most air molecules are either
oxygen (O2 ) or nitrogen (N2 ). In the stratosphere, near the peakconcentration of the ozone layer, there are typically a few thousandozone molecules for every billion air molecules (1 billion = 1,000
million). In the troposphere near Earth’s surface, ozone is even lessabundant, with a typical range of 20 to 100 ozone molecules foreach billion air molecules. The highest ozone values near the surface occur in air that is polluted by human activities. Throughout
this document the word “abundance” refers to the concentrationor amount of an atmospheric gas or some other physical quantity.
As an illustration of the low relative abundance of ozone in our
4Section I: Ozone in our atmosphere

atmosphere, one can imagine bringing all the ozone moleculesin the troposphere and stratosphere down to Earth’s surface andforming a layer of pure ozone that extends over the entire globe.
The resulting layer would have an average thickness of about three
millimeters (0.12 inches), which scientists would report as 300
Dobson Units (see Q3). Nonetheless, this extremely small fractionof the atmosphere plays a vital role in protecting life on Earth (seeQ2).

Ozone and Oxygen

Oxygen atom (O)Oxygen molecule (O2)

Ozone molecule (O3)
Figure Q1-1. Ozone and oxygen. A molecule of ozone (O3)contains three oxygen atoms (O) bound together. Oxygen
molecules (O2), which constitute about 21% of the gases inEarth’s atmosphere, contain two oxygen atoms bound together.

2022 Update | Twenty Questions

Q1

Ozone in the Atmosphere
35
20
30

Altitude (kilometers)

15

Ozone Layer
20

10
15

Altitude (miles)

Stratospheric
Ozone

25

Figure Q1-2. Ozone in the atmosphere.
Ozone is present throughout the troposphere and stratosphere. This profileshows schematically how ozone changes
with altitude in the tropics. Most ozone
resides in the stratospheric “ozone layer.”
The vertical extent or thickness of this layer varies from region to region and withseason over the globe (see Q3). Increases
in ozone near the surface are a result of air
pollutants released by human activities.
The direct exposure to ozone is harmful
to humans and other forms of life.

10

5

0

Largest increases
from air pollutants

Tropospheric
Ozone

Ozone abundance

Stratospheric ozone. Stratospheric ozone is formed naturally bychemical reactions involving solar ultraviolet radiation (sunlight)
and oxygen molecules, which make up about 21% of the atmosphere. In the first step, solar ultraviolet radiation breaks apart one
oxygen molecule (O2) to produce two oxygen atoms (2 O) (seeFigure Q1-3). In the second step, each of these highly reactive
oxygen atoms combines with an oxygen molecule to produce anozone molecule (O3). These reactions occur continually wheneversolar 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 sunlightand a wide variety of natural and human-produced chemicals inthe stratosphere. In each reaction, an ozone molecule is lost andother chemical compounds are produced. Important reactive gases that destroy ozone are hydrogen and nitrogen oxides and thosecontaining chlorine and bromine (see Q7). Some stratosphericozone is regularly transported down into the troposphere and canoccasionally influence ozone amounts at Earth’s surface.

50

Tropospheric ozone. Near Earth’s surface, ozone is produced bychemical reactions involving gases emitted into the atmospherefrom both natural sources and human activities. Ozone productionin the troposphere primarily occurs by reactions of hydrocarbon
and nitrogen oxide gases, and all require sunlight for completion.
Fossil fuel combustion and deforestation are the primary sourcesof pollutant gases that lead to production of tropospheric ozone.
As in the stratosphere, ozone in the troposphere is destroyed bynaturally 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 thoseof soils and plants.

Balance of chemical processes. Ozone abundances in thestratosphere and troposphere are determined by the balance between chemical processes that produce and destroy ozone. Thebalance is determined by the amounts of reactive gases and how
the rate or effectiveness of the various reactions varies with sunlightintensity, location in the atmosphere, temperature, and other factors. As atmospheric conditions change to favor ozone-producing
Section I: Ozone in our atmosphere5

Q1

Twenty Questions | 2022 Update

reactions in a certain location, ozone abundances increase. Similarly, if conditions change to favor other reactions that destroyozone, abundances decrease. The balance of production and loss
reactions, combined with atmospheric air motions that transportand mix air with different ozone abundances, determines the glob-

al distribution of ozone on timescales of days to many months (seealso Q3). Global stratospheric ozone decreased from the 1970s
to the late 1990s (see Q12 and Q13) because the amounts of reactive gases containing chlorine and bromine in the stratosphere
increased due to human activities (see Q6 and Q15).

Stratospheric Ozone Production
STEP 1

STEP 2O
Ultraviolet
Sunlight

O2

O3

O2

O3

O2
O
OVERALL REACTION :3 O2SUNLIGHT

2 O3

Figure Q1-3. Stratospheric ozone production. Ozone is produced naturally in the stratosphere by a two-step reaction process. In the firststep, solar ultraviolet radiation (sunlight) breaks apart an oxygen molecule to form two separate oxygen atoms. In the second step, eachoxygen atom collides with another oxygen molecule and forms an ozone molecule in a binding reaction. In the overall process, three oxygen molecules plus sunlight react to form two ozone molecules.

6

Section I: Ozone in our atmosphereQ2

Why do we care aboutatmospheric ozone?

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 this gas is harmful tohumans, plants, and animals.

Ozone in the stratosphere (Good ozone). Stratospheric ozone isconsidered good for humans and other life forms because ozone
absorbs ultraviolet (UV) radiation from the Sun (see Figure Q2-1).
The Sun emits UV radiation that scientists categorize into three
wavelength ranges: UV-C (100 to 280 nanometer (nm) wavelengths); UV-B (280 to 315 nm), and UV-A (315 to 400 nm). Theenergy of solar UV radiation, which cannot be seen by the humaneye, is higher at shorter wavelengths. Exposure to high energy
UV-C radiation is particularly dangerous to all life forms. Fortunately, UV-C radiation is entirely absorbed within the ozone layer. Most
UV-B radiation emitted by the Sun is absorbed by the ozone layer;
the rest reaches Earth’s surface. In humans, increased exposure
to UV-B radiation raises the risks of skin cancer and cataracts, andsuppresses the immune system. Exposure to UV-B radiation before
adulthood and cumulative exposure are both important health risk
factors. Excessive UV-B exposure also can damage terrestrial plantlife, including agricultural crops, single-celled organisms, andaquatic ecosystems. Low energy UV radiation, UV-A, which is notabsorbed significantly by the ozone layer, causes premature agingof the skin.
Protecting stratospheric ozone. In the mid-1970s, it was discovered that gases containing chlorine and bromine atoms releasedby human activities could cause stratospheric ozone depletion (seeQ5 and Q6). These gases, referred to as halogen source gases, andalso as ozone-depleting substances (ODSs), chemically release
their chlorine and bromine atoms after they reach the stratosphere.

UV Protection by the Stratospheric Ozone Layer

U V- CU V- BU V- A

15–400 nm
15 nm 3–3082
nm100–280

OZONE LAYER

Figure Q2-1. UV protection by the ozone layer. The
ozone layer is located in the stratosphere and surrounds the entire Earth. The Sun emits ultraviolet (UV)radiation that reaches the top of the ozone layer, whichscientists classify according to wavelength. Solar UV-C
radiation (wavelength range 100 to 280 nanometer(nm)) is extremely damaging to humans and other life
forms; UV-C radiation is entirely absorbed within theozone layer. Solar UV-B radiation (280 to 315 nm) is
only partially absorbed and, as a result, humans andother life forms are exposed to some UV-B radiation.
Excessive exposure to UV-B radiation increases therisks of skin cancer, cataracts, and a suppressed immune system for humans and also damages terrestrial plant life, single-cell organisms, as well as aquatic
ecosystems. UV-A (315 to 400 nm), visible light, andother wavelengths of solar radiation are only weakly
absorbed by the ozone layer. Exposure to UV-A is associated with premature aging of the skin and someskin cancers. Depletion of the ozone layer primarily increases the amount of UV-B radiation that reaches thesurface (Q16). Avoiding ozone depletion that would
increase human exposure to UV-B radiation is a principal objective of the Montreal Protocol.
(The unit “nanometer” (nm) is a common measure of the
wavelength of light; 1 nm equals one billionth of a meter(=10 –9 m).)Section I: Ozone in our atmosphere7

Q2

Twenty Questions | 2022 Update

Ozone depletion increases surface UV-B radiation above naturally
occurring amounts. International efforts have been successful inprotecting the ozone layer through controls on the production andconsumption of ODSs (see Q14 and Q15).
Ozone in the troposphere (Bad ozone). Ozone near Earth’s surface in excess of natural amounts is considered bad ozone (seeFigure Q1-2). Surface ozone in excess of natural levels is formedby reactions involving air pollutants emitted from human activities,
such as nitrogen oxides (NOx ), carbon monoxide (CO), and various hydrocarbons (gases containing hydrogen, carbon, and often
oxygen). Exposure to ozone at concentrations above natural levels is harmful to humans, plants, and other living systems because
ozone reacts strongly to destroy or alter molecules that constitutebiological tissue. Enhanced surface ozone caused by air pollution
reduces crop yields and forest growth. In humans, exposure tohigh levels of ozone can reduce lung capacity; cause chest pains,throat irritation, and coughing; and worsen pre-existing healthconditions related to the heart and lungs. In addition, increases intropospheric ozone lead to a warming of Earth’s surface because
ozone is a greenhouse gas (GHG) (see Q17). The negative effects
of excess tropospheric ozone contrast sharply with the protection
from harmful UV radiation afforded by preserving the natural abundance of stratospheric ozone.

8

Section I: Ozone in our atmosphere

Reducing tropospheric ozone. Limiting the emission of certain
common pollutants reduces the production of excess ozone near
Earth’s surface, where ozone can affect humans, plants, and animals. Major sources of pollutants include large cities, where fossil
fuel consumption for transportation, heating, and industrial activities is concentrated, power plants that rely on coal, oil, or naturalgas, as well as deforestation, wildfires, and the burning of savannah for agriculture. Many programs around the globe have beensuccessful in reducing or limiting the emission of pollutants thatcause production of excess ozone near Earth’s surface.
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 componentof the clean atmosphere. Natural emissions from the biosphere,
mainly from trees, participate in chemical reactions that produce
ozone. Atmospheric ozone plays important ecological roles beyond absorbing UV radiation. For example, ozone initiates thechemical removal of many pollutants as well as some GHGs, suchas methane (CH4 ). In addition, the absorption by ozone of solar UV
radiation as well as visible and infrared radiation is a natural sourceof heat in the stratosphere, causing temperatures to increase withaltitude. Stratospheric temperatures affect the balance of ozone
production and destruction processes (see Q1) and air motions
that redistribute ozone throughout the stratosphere (see Q3).

Q3

How is total ozone distributed overthe globe?

The distribution of total ozone over Earth varies with geographic location and on daily and seasonal timescales. These variations are caused by large-scale movements of stratospheric and tropospheric air and thechemical production and destruction of ozone. Total ozone is generally lowest at the equator and highest
in midlatitude and polar regions.

Total ozone. The total column of ozone at any location on the
globe is defined as the sum of all the ozone in the atmosphere directly above that location. Most ozone resides in the stratosphericozone layer and a small percentage (about 5 to 10%) is distributed
throughout the troposphere (see Q1). Total column ozone values
are usually reported in Dobson units denoted as “DU.” Typical values vary between 200 and 500 DU over the globe, with a global
average abundance of about 300 DU (see Figure Q3-1). The ozone
molecules required for total ozone to be 300 DU would form a layer of pure ozone gas at Earth’s surface having a thickness of only
3 millimeters (0.12 inches) (see Q1), which is about the height ofa stack of 2 common coins, if these molecules could be isolated
and compressed. It is remarkable that a layer of pure ozone only
3 millimeters thick protects life on Earth’s surface from most of theharmful UV radiation emitted by the Sun (see Q2).

ern Hemisphere as shown in Figure Q3-1, with increasing values in
Arctic total ozone during winter, a clear maximum in spring, and decreasing values from summer to autumn. In the Antarctic, however,a pronounced minimum in total ozone is observed during spring.
The minimum is known as the “ozone hole”, which is caused by
the widespread chemical depletion of ozone in spring by pollutants known as ozone-depleting substances (see Q5 and Q10). In
the late 1970s, before the ozone hole appeared each year, muchhigher ozone values than those currently observed were found in
the Antarctic spring (see Q10). Currently, the lowest values of totalozone across the globe and all seasons are found during early
spring in the Antarctic, as shown in Figure Q3-1 for 2021. After
spring, these low values disappear from total ozone maps as polar
air mixes with lower-latitude air containing much higher amountsof ozone.

Global distribution. Total ozone varies strongly with latitude overthe globe, with the largest values occurring at middle and high latitudes during most of the year (see Figure Q3-1). This distribution is
the result of the large-scale circulation of air in the stratosphere thatslowly transports ozone-rich air from high altitudes in the tropics,
where ozone production from solar ultraviolet radiation is largest,
toward the poles. Ozone accumulates at middle and high latitudes,
increasing the vertical extent of the ozone layer and, at the sametime, total ozone. The total column of ozone is generally smallest
in the tropics for all seasons. An exception since the mid-1980s is
the region of low values of ozone over Antarctica during spring in
the Southern Hemisphere, a phenomenon known as the Antarctic
ozone hole (dark blue, Figure Q3-1; also see Q10 and Q11).

In the tropics, the change in total ozone through the progressionof the seasons is much smaller than at higher latitudes. This feature
is present because seasonal changes in both sunlight and ozone
transport are much smaller in the tropics compared to higher latitudes.

Seasonal distribution. Total ozone also varies with season,
as shown in Figure Q3-1 using two-week averages of ozone taken from satellite observations acquired in 2021. March and September plots represent the early spring and autumn seasons in
the Northern and Southern Hemispheres, respectively. June andDecember plots similarly represent the early summer and winter seasons. During spring, total ozone exhibits maximums atlatitudes poleward of about 45° N in the Northern Hemisphereand between 45° and 60° S in the Southern Hemisphere. These
spring maximums are a result of increased transport of ozone fromits source region in the tropics toward high latitudes during lateautumn and winter. This poleward ozone transport is much weaker
during the summer and early autumn periods and is weaker overall
in the Southern Hemisphere.
This natural seasonal cycle can be observed clearly in the North-

The abundance of ozone is larger at midlatitudes in the Northern
Hemisphere (NH) than the Southern Hemisphere (SH), for all four
seasons in the respective hemispheres. The thinner ozone layer at
SH midlatitudes compared to NH midlatitudes is due to several
factors: differences in the large-scale circulation of the two hemispheres that preceded the development of the ozone hole as wellas larger abundances of tropospheric ozone in the NH comparedto the SH that is caused by more pollution in the more heavily populated NH. Dilution of ozone-depleted air from the Antarctic ozone
hole region starting in the 1980s further increases the hemisphericdifference in total ozone. This hemispheric total ozone difference
results in higher levels of UV light reaching the surface in the SH
compared to the NH (see Q16).

Natural variations. Total ozone varies strongly with latitude andlongitude, as seen within the seasonal plots in Figure Q3-1. These
patterns come about for two reasons. First, atmospheric winds
transport air between regions of the stratosphere that have highozone values and those that have low ozone values. Tropospheric weather systems can temporarily alter the vertical extent of theozone layer in a region, and thereby change total ozone. The regular nature of these air motions, in some cases associated withgeographical features (oceans and mountains), in turn causes recurring patterns in the distribution of total ozone. Second, ozone
Section I: Ozone in our atmosphere9Q3

Twenty Questions | 2022 Update

variations occur as a result of changes in the balance of chemical
production and loss processes. This balance is very sensitive to theamount of solar UV radiation (see Q2) reaching the various parts of
the atmosphere.
There is a good understanding of how chemistry and air motions
work together to cause the observed large-scale features in totalozone, such as those seen in Figure Q3-1. Ozone changes areroutinely monitored by a large group of scientists using satellite,
airborne, and ground-based instruments. The continued analyses of these observations provide an important long-term basis toquantify the contribution of human activities to ozone depletion.

Global Satellite Maps of Total Ozone in 2021
Spring16 – 30 March

Summer

WinterAutumn16 – 30 June

16 – 30 December

Total Ozone (Dobson units)

Autumn600

Summer

Winter500
400

16 – 30 September

300
200
100

SpringFigure Q3-1. Total ozone. Total column ozone at any location on the globe is defined as the sum of all the ozone molecules in the atmosphere directly above that location. Total ozone varies with latitude, longitude, and season, with the largest values at high latitudes andthe lowest values usually found in tropical regions. The variations are demonstrated here with two-week averages of total ozone in 2021
as measured with a satellite instrument. Total ozone shows little variation in the tropics (20°N–20°S latitudes) over all seasons. Total ozone
outside of the tropics varies more strongly with time on a daily to seasonal basis as ozone-rich air is moved from the tropics and accumulates at higher latitudes, with more ozone being transported in winter. The low total ozone values over Antarctica shown here in September
constitute the “ozone hole” in 2021. Since the mid-1980s, the ozone hole in late winter/early spring represents the lowest values of totalozone that occur over all seasons and latitudes (see Q10).

10

Section I: Ozone in our atmosphereQ4

How is ozone measured in theatmosphere?

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

The abundance of ozone in the atmosphere is measured by avariety of techniques (see Figure Q4-1). The techniques makeuse of ozone’s unique optical and chemical properties. There are
two principal categories of measurement techniques: local andremote. 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.

Measuring Ozone in the Atmosphere

Satellites

Local measurements. Local measurements of the atmospheric abundance of ozone are those that require air to be drawn
directly into an instrument. Once inside an instrument’s detection
chamber, the amount of ozone is determined by measuring theabsorption of ultraviolet (UV) radiation or by the electrical current
or light produced in a chemical reaction involving ozone. The latter approach is used in “ozonesondes”, which are lightweight,
ozone-measuring modules suitable for launching on small balloons. The balloons ascend up to altitudes of about 32 to 35 kilometers (km), high enough to measure ozone in the stratosphericozone layer. Ozonesondes are launched regularly at many locations around the world. Local ozone-measuring instruments
using optical or chemical detection schemes are also used on
research aircraft to measure the distribution of ozone in the troposphere and lower stratosphere (up to altitudes of about 20 km).
High-altitude research aircraft can reach the ozone layer at most
locations over the globe and can reach furthest into the layer athigh latitudes. Ozone measurements are also being made routinely on some commercial aircraft flights. Local measurementsof the abundance of ozone at the surface are obtained at many
thousands of sites over the globe, which provide hourly data critical for assessing and improving air quality throughout the world.

Remote measurements. Remote measurements of total ozone
amounts and the altitude distributions of ozone are obtained bydetecting 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 sunlight (and reflected sunlight from the moon), lasers, and starlight.
For example, satellite instruments 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 theentire globe on a daily basis. Lidar instruments, which measurebackscattered 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

High-altitude
aircraft

Ozonesondes
Low-altitude
aircraft

Ground-based
systems

Figure Q4-1. Ozone measurements. Ozone is measured throughout
the atmosphere with instruments on the ground, aircraft, high-altitude balloons, and satellites. Some instruments measure ozone locally in sampled air and others measure ozone remotely a considerable
distance away from the instrument. Instruments use optical techniques with either the Sun (direct, reflected, or scattered radiation) or
lasers (green lines) as light sources; detect the thermal emissions fromozone and other atmospheric molecules (not shown); or use chemical
reactions that are unique to ozone (ozonesonde). At many locations
over the globe, regular measurements are made to monitor ozone
amounts and their variations over time.
Section I: Ozone in our atmosphere

11

Q4

Twenty Questions | 2022 Update

instruments measures ozone by detecting small changes in theamount of the Sun’s UV radiation that reaches Earth’s surface. Other instruments measure ozone using either its absorption of infrared, visible, or ultraviolet radiation or its emission of microwave or
infrared radiation at different altitudes in the atmosphere, thereby

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 samplingpolar 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 (UV) 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 UV wavelengths: onethat is strongly absorbed by ozone and one that is weakly absorbed. The difference in light intensity at the two wavelengths provides
a measure of the total amount of 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 hundreds of sites located around the world ranging from South Pole, Antarctica (90°S) to Ellesmere Island, Canada (83°N) thatroutinely measure total ozone. The Brewer spectrophotometer was introduced into the global network starting in 1982. Whereas
the original Dobson instrument measures atmospheric ozone based on observations of UV light at only two wavelengths, modern
Dobson instruments as well as Brewer instruments use observations from multiple pairs of wavelengths. The accuracy of these observations is maintained by regular instrument calibrations and intercomparisons. At many of the stations, observations of total ozone
are augmented by measurements of the vertical distribution of ozone obtained either from twilight measurements with Dobson or
Brewer instruments, by routine launches of ozonesondes, or with the use of lidar instruments. Numerous stations also quantify theatmospheric abundances of a wide variety of related compounds, taking advantage of the unique optical properties of atmosphericgases.
Data from the network have been essential for understanding the effects of chlorofluorocarbons and other ozone-depleting substances on the global ozone layer, starting before the launch of space-based ozone-measuring instruments and continuing to thepresent day. Ground-based instruments with excellent long-term stability and accuracy are now routinely used to help calibrate
space-based observations of total ozone as well as numerous other gases and physical quantities involved in ozone chemistry.
Keeping with the tradition of naming units after pioneering scientists, the unit of measure for total ozone is called the “Dobson unit”(see Q3).

12

Section I: Ozone in our atmosphereQ5

How do emissions of halogen source gases
lead to stratospheric ozone depletion?

The initial step in the depletion of stratospheric ozone by human activities is the emission, at Earth’s
surface, of gases that contain chlorine and bromine and have long atmospheric lifetimes. Most of these gases
accumulate in the lower atmosphere because they are relatively unreactive and do not dissolve readily inrain or snow. Natural air motions eventually transport these accumulated gases to the stratosphere, where
they are converted to more reactive gases. Some of these gases then participate in reactions that destroyozone. Finally, when stratospheric large-scale circulation patterns return this air to the lower atmosphere,
these reactive chlorine and bromine gases are removed from Earth’s atmosphere by rain and snow.

The principal steps in stratospheric ozone depletion caused by human activities are shown in Figure Q5-1.

greater than a few decades circulate between the troposphere andstratosphere multiple times, on average, before conversion occurs.

Emission, accumulation, and transport. The process begins

The reactive gases formed from halogen source gases react chemically to destroy ozone in the stratosphere (see Q8). The average
depletion of total ozone attributed to reactive gases is smallest in the tropics and largest at high latitudes (see Q12). In polar
regions, reactions that occur on the surface of polar stratosphericclouds, which exist only at low temperatures, greatly increase theabundance of the most important reactive chlorine gas, chlorinemonoxide (ClO) (see Q9). This process results in substantial ozone
destruction in polar regions in late winter/early spring (see Q10
and Q11).

with the emission, at Earth’s surface, of long-lived source gases containing the halogens chlorine and bromine (see Q6). The
halogen source gases, often referred to as ozone-depleting
substances (ODSs), include manufactured chemicals released to theatmosphere when used in a variety of applications, such as refrigeration, air conditioning, and foam blowing. Chlorofluorocarbons(CFCs) are an important example of chlorine-containing sourcegases. Emitted source gases accumulate in the lower atmosphere(troposphere) and are slowly transported to the stratosphere by
natural air motions. The accumulation occurs because most sourcegases are highly unreactive in the lower atmosphere. Furthermore,
only a small amount of halogen source gases dissolve in ocean
waters. The low reactivity of these manufactured halogenatedgases within the lower atmosphere is one property that made themwell suited for specialized applications such as refrigeration.
Some halogen gases are emitted in substantial quantities fromnatural sources (see Q6). These emissions also accumulate in the
troposphere, are transported to the stratosphere, and participate
in ozone destruction reactions. These naturally emitted gases arepart of the natural balance of ozone production and destructionthat predates the large release of manufactured halogenated gases and the associated observed 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 and reservoir
halogen gases by the absorption of ultraviolet radiation from theSun (see Q7). The rate of conversion is related to the atmospheric
lifetime of a gas (see Q6). Gases with longer lifetimes have slower
conversion rates and survive longer in the atmosphere after emission. Lifetimes of the principal ODSs vary from about 1 to 100 years(see Table Q6-1). Emitted gas molecules with atmospheric lifetimes

Air in the stratosphere is generally isolated from the troposphere.
A small portion of stratospheric air returns to the troposphere every
day, bringing along reactive and reservoir halogen gases. On average, it takes several years for air throughout the global stratosphereto return to the troposphere. Reactive halogen gases that are transported back to the troposphere are removed from the atmosphereby rain and other precipitation or deposited by wind onto Earth’s
land or ocean surfaces. These removal processes bring to an end
the destruction of ozone by chlorine and bromine atoms that werefirst released to the atmosphere as components of halogen sourcegas molecules.

Tropospheric conversion. Halogen source gases with short lifetimes (less than 1 year) undergo significant chemical conversion in
the troposphere, producing reactive and reservoir halogen gases.
Source gas molecules that are not converted are transported to thestratosphere. Only small portions of reactive and reservoir 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, prior to transport to the stratosphere, are the hydrochlorofluorocarbons (HCFCs), methyl bromide (CH3Br), methyl chloride(CH3Cl), and gases containing iodine (see Q6).

Section II: The ozone depletion process

13

Q5

Twenty Questions | 2022 Update

Principal Steps in the Depletionof Stratospheric Ozone

1

Halogen source gases are emitted at Earth's
surface by human activities and natural processes.

2

Halogen source gases accumulate in theatmosphere and are globally distributed throughout
the lower atmosphere by winds and other air motions.

3

Halogen source gases are transported to thestratosphere by air motions.

4

Emissions

Accumulation

Transport

Conversion
Most halogen source gases are converted in thestratosphere to reactive and reservoir halogen
gases in chemical reactions involving ultraviolet
radiation from the Sun.

Chemical reaction

5
6

14

Reactive halogen gases cause chemical depletionof stratospheric ozone over the globe.
Low-temperature surface reactions on polar
stratospheric clouds (PSCs) significantly increase
reactive halogen gases and thereby cause severe
ozone loss in polar regions in late winter and early
spring.

RemovalAir containing reactive and reservoir halogen gases
returns to the troposphere where the gases are
removed by moisture in clouds and rain.

Section II: The ozone depletion process

Figure Q5-1. Principal steps in stratospheric ozone
depletion. The stratospheric ozone depletion process
begins with the emission of halogen source gases byhuman activities and natural processes. These compounds have at least one carbon and one halogen atom,
causing them to be chemically stable and leading tocommon use of the term halocarbon, an abbreviation for
halogen and carbon. Many halocarbon gases emitted
by human activities are also ozone-depleting substances(ODSs); all ODSs contain at least one chlorine or bromine
atom (see Q7). Most of these compounds undergo little
or no chemical loss within the troposphere, the lowest
region of the atmosphere, and accumulate until transported to the stratosphere. Subsequent steps are theconversion of ODSs to reactive and reservoir halogen
gases and chemical reactions that remove ozone (seeQ8). Ozone depletion by halogen gases occurs throughout the globe (see Q12), with the largest losses occurringin the polar regions during late winter and early spring(see Q9 to Q11). Ozone depletion ends when reactive
and reservoir halogen gases are removed by rain andsnow in the troposphere and deposited on Earth’s surface.

2022 Update | Twenty Questions

Q5

Understanding Stratospheric Ozone Depletion
Our understanding of stratospheric ozone depletion has been obtained through a combination of laboratory studies, computermodels, and atmospheric observations. The wide variety of chemical reactions that occur in the stratosphere have been discovered
and investigated in laboratory studies. Chemical reactions between two gases follow well-defined physical rules. Some of thesereactions occur on the surfaces of polar stratospheric clouds formed in the winter stratosphere. Reactions have been studied thatinvolve many different molecules containing chlorine, bromine, fluorine, and iodine and other atmospheric constituents such ascarbon, 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 representation of chemical abundances, winds, air temperatures, andthe 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 importanceof new reactions found in laboratory studies. The response of ozone to possible future changes in the abundances of trace gases,
temperatures, and other atmospheric parameters has 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 with respect to time and location. Gas and particle abundances have been monitored over time periods spanning a daily cycle
to decades. Observations show that reactive halogen gases are present in the stratosphere at the amounts required to cause observed ozone depletion (see Q7). 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 (seeQ8). Instruments on the ground and on satellites, balloons, and aircraft now routinely measure the abundance of ozone and ClO
remotely using optical and microwave signals. High-altitude aircraft and balloon instruments are also used to measure both gases
locally in the stratosphere (see Q4). Observations of ozone and reactive gases made in past decades are used extensively in comparisons with computer models to increase confidence in our understanding of stratospheric ozone depletion.

Section II: The ozone depletion process

15

Q6

What emissions from human activities
lead to stratospheric ozone depletion?

Certain industrial processes and consumer products result in the emission of ozone-depleting substances(ODSs) to the atmosphere. Principal ODSs are manufactured halogen source gases that are now controlledworldwide 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 asfire extinguishing agents. Current ODS abundances in the atmosphere are known directly from air sample
measurements.

Halogen source gases versus ozone-depleting substances(ODSs). Halogen source gases that are emitted by human activities and controlled by the Montreal Protocol are generally referred
to as ODSs. The Montreal Protocol controls the global production
and consumption of ODSs (see Q14). Halogen source gases suchas methyl chloride (CH3Cl) that have predominantly natural sources are not classified as ODSs. The contributions of various ODSs
and natural halogen source gases to the total amount of chlorine
and bromine entering the stratosphere are shown in Figure Q6-1.
Total chlorine and total bromine entering the stratosphere peaked in 1993 and 1999, respectively. The difference in the timing ofthese peaks is a result of various phaseout schedules specified by
the Montreal Protocol and its amendments and adjustments, different atmospheric lifetimes of halogen source gases, and the timedelays between production and emission of the numerous sourcegases. Also shown are the contributions to total chlorine andbromine in 2020, highlighting the reductions of 11% and 15%,
respectively, achieved by the controls of the Montreal Protocol.

Ozone-depleting substances (ODSs). The principal ODSs are
manufactured for specific industrial uses or consumer products,
most of which result in the eventual emission of these gases to theatmosphere. 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). Because of their long atmospheric lifetimes, a large fraction of the emitted ODSs reach thestratosphere, where they are converted to reactive and reservoir
gases containing chlorine and bromine that lead to ozone depletion.
ODSs containing only chlorine, fluorine, and carbon are called
chlorofluorocarbons, usually abbreviated as CFCs. The principal CFCs are CFC-11 (CCl3F), CFC-12 (CCl2F2 ), and CFC-113(CCl2FCClF2 ). CFCs, along with carbon tetrachloride (CCl4 ) andmethyl chloroform (CH3CCl3 ), historically have been the mostimportant chlorine-containing halogen source gases emitted byhuman activities. These and other chlorine-containing ODSs havebeen used in many applications, including refrigeration, air conditioning, foam blowing, spray can propellants, and cleaning of metals and electronic components. As a result of the Montreal Protocol
controls, the abundances of these chlorine source gases have de16

Section II: The ozone depletion processcreased since 1993 (see Figure Q6-1). The abundances of CFC-11
and CFC-12 in 2020 were 16% and 2.8% lower than their values in
1993, respectively.
The class of compounds known as hydrochlorofluorocarbons(HCFCs) contain hydrogen, in addition to chlorine, fluorine, andcarbon. HCFC-22 (CHF2Cl), developed in the 1930s, has beenused as a refrigerant, primarily in residential air conditioners, sincethe 1940s. As detailed below, HCFCs are less harmful to the ozone
layer compared to CFCs. In the 1990s, the use of HCFC-22 expanded and other HCFCs were developed as substitutes for CFCs.
Consequently, the chlorine content of HCFCs entering the stratosphere increased by 185% between 1993 and 2020 (see FigureQ6-1). With restrictions on production starting in 1996, and globally in place since 2013, the atmospheric abundances of HCFCs
are expected to peak between 2023 and 2030 (see Figures Q0-1
and Q15-1). Classes of compounds known as hydrofluorocarbons(HFCs) and hydrofluoroolefins (HFOs) constitute the replacement
for many applications of HCFCs.
Another category of ODSs contains bromine. The most important of these gases are the halons and methyl bromide (CH3Br).
Halons are a group of industrial compounds that contain at least
one bromine and one carbon atom; halons may or may not containa chlorine atom. Halons were originally developed to extinguishfires and were widely used to protect large computer installations,military hardware, and commercial aircraft engines. Consequently,
halons are often released directly into the atmosphere upon use
or testing of these fire suppression systems. The most abundanthalons emitted by human activities are halon-1211 (CBrClF2 ) andhalon-1301 (CBrF3 ). Methyl bromide is used primarily as a fumigant
for pest control in agriculture and disinfection of export shipping
goods, and also has significant natural sources.
As a result of the controls of the Montreal Protocol, the contribution to the atmospheric abundance of methyl bromide from humanactivities decreased by 71% between 1999 and 2020 (seeFigure Q6-1). The concentration of halon-1211 peaked in 2005 andhas been decreasing ever since, reaching an abundance in 2020
that was 22% below that measured in 1999. The abundance of
halon-1301, on the other hand, increased by 19% since 1999
and is expected to slowly decline into the next decade because

2022 Update | Twenty Questions

Q6

Halogen Source Gases Entering the Stratosphere

2500
2000
1500
1000
500
0

+185%
–27%
–16%

–2.8%

Naturalsources

–0.4%

1993

CFC-11

CFC-12

Methylchloride

2020

Chlorine source gases

20

15

22.1 ppt

0–25%

Other halons+19%

Halon-1301–22%

Halon-1211

–71%

10

5

18.9 ppt–15%

ODSs from human activity

3000

–11%
+63%
–16%

Very short-lived
gases
Methyl chloroform,–99%
Other gases, +11%
CFC-113
HCFCs
Carbontetrachloride

nochangeNatural sources

3500

Total

3240 ppt

Total bromine amount (parts per trillion)

3660 ppt

Ozone-depleting substances (ODSs)
from human activity

Total chlorine amount (parts per trillion)

Total

no
change1999

Methyl bromide
human, natural

Very short-lived
gases

2020

Bromine source gases

Figure Q6-1. Changes in halogen source gases entering the stratosphere. A variety of halogen source gases emitted by human activities and natural processes transport chlorine and bromine into the stratosphere. Ozone-depleting substances (ODSs) are the subset ofthese gases emitted by human activities that are controlled by the Montreal Protocol. These partitioned columns show the abundancesof chlorine- and bromine-containing gases entering the stratosphere in 1993 and 1999, when their total amounts peaked, respectively,
and in 2020. The overall reductions in the total amounts of chlorine and bromine entering the stratosphere and the changes observedfor each source gas are also indicated. The amounts are derived from tropospheric observations of each gas. Note the large difference inthe vertical scales: total chlorine entering the stratosphere is about 150 times greater than total bromine. Both, however, are importantbecause bromine is about 60 times more effective on a per-atom basis than chlorine at destroying ozone. Human activities are the largest
source of chlorine reaching the stratosphere and CFCs are the most abundant chlorine-containing gases. Methyl chloride is the primary natural source of chlorine. The largest decreases between 1993 and 2020 are seen in methyl chloroform, carbon tetrachloride, and
CFC-11. The abundance of HCFCs, which are substitute gases for CFCs and also controlled under the Montreal Protocol, have risen substantially since 1993 and have approached expected peak atmospheric abundances (see Figure Q15-1). The abundance of chlorine-containing very short-lived gases entering the stratosphere has risen substantially since 1993; these compounds originate primarily from human activity, undergo chemical loss within the troposphere, and are not controlled by the Montreal Protocol. Halons and methyl bromide
are the largest contributors to bromine entering the stratosphere. The largest decrease between 1999 and 2020 is seen in the abundanceof methyl bromide attributed to human activities, because of the success of the Montreal Protocol. Halon-1301 is the only brominated ODS
showing an increased abundance relative to 1999. Methyl bromide also has a natural source, which is now substantially greater than thehuman source due to the success of the Montreal Protocol. Natural sources, which make a much larger fractional contribution to bromine
entering the stratosphere than occurs for chlorine, have remained fairly constant in the recent past.
(The unit “parts per trillion” is used here as a measure of the relative abundance of a substance in dry air: 1 part per trillion equals the presence of one molecule of a gas per trillion (=1012) total air molecules.)of continued small releases and a long atmospheric lifetime (seeFigure Q15-1). In 2020, the bromine content of other halons (mainly
halon-1202 and halon-2402) was 25% below the amount present
in 1999.

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. In addition, very short-lived source gases(defined as compounds with atmospheric lifetimes typically lessthan 0.5 year) containing bromine such as bromoform (CHBr3 ) anddibromomethane (CH2Br2 ) are also released to the atmosphere,
primarily from biological activity in the oceans. Only a fraction of
the emissions of very short-lived source gases reaches the stratosphere because these gases are efficiently removed in the lower
atmosphere. Volcanoes provide an episodic source of reactive halogen gases that sometimes reach the stratosphere in appreciable
quantities.
Other natural sources of halogens include reactive chlorine andbromine produced by evaporation of ocean spray. However, these
Section II: The ozone depletion process

17

Q6

Twenty Questions | 2022 Update

reactive chemicals play no role in stratospheric ozone depletion
because they readily dissolve in water and are removed in the troposphere.

ozone depletion in the stratosphere. Methyl chloride, despite its
large source, constituted only about 17% (540 ppt) of the halogen
source gases entering the stratosphere in 2020 (see Figure Q6-1).

In 2020, natural sources contributed about 17% of total stratospheric chlorine and about 56% of total stratospheric bromine(see Figure Q6-1). The amount of chlorine and bromine entering
the stratosphere from natural sources is known to be fairly constantover time and, therefore, cannot be the cause of the ozone depletion observed since the 1980s.

The amount of an emitted gas that is present in the atmosphere
represents a balance between its emission and removal rates. A
wide range of current emission rates and atmospheric lifetimes
are derived for the various source gases (see Table Q6-1). 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 important substitute gases, the HCFCs, continue to slowly increase under the provisions of the Montreal Protocol (see Q15). In the past few years, the rate of the increase of theatmospheric abundance of HCFCs has declined. In the coming
decades, the emissions and atmospheric abundances of all controlled ODSs are expected to decrease under these provisions.

Other human activities that are sources of chlorine andbromine gases. 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 andonly a small fraction of their emissions reaches the stratosphere.
The contribution of very short-lived chlorinated gases from natural sources and human activities to total stratospheric chlorine was
63% larger in 2020 than in 1993, and now contributes about 4%(130 ppt) of the total chlorine entering the stratosphere (see FigureQ6-1). The Montreal Protocol does not control the production andconsumption of very short-lived chlorine source gases, although
the atmospheric abundances of some (notably dichloromethane,
CH2Cl2 ) have increased substantially in recent years. Solid rocket
engines, such as those used to propel payloads into orbit, release
reactive chlorine gases directly into the troposphere and stratosphere. The quantities of chlorine emitted globally by rockets iscurrently small in comparison with halogen emissions from otherhuman activities.

Lifetimes and emissions. Estimates of global emissions in2020 for a selected set of halogen source gases are given in
Table Q6-1. These emissions occur from continued production of
HCFCs and HFCs as well as the release of gases from banks. Emission from banks refers to the atmospheric release of halocarbons
from existing equipment, chemical stockpiles, foams, and otherproducts. In 2020 the global emission of the refrigerant HCFC-22
constituted the largest annual release, by mass, of a halocarbon
from human activities. Release in 2020 of HFC-134a (CH2FCF3 ),another refrigerant, was second largest. The emission of methyl
chloride (CH3Cl) is primarily from natural sources such as the ocean
biosphere, terrestrial plants, salt marshes and fungi. The humansource of methyl chloride is small relative to the total natural source(see Q15).
After emission, halogen source gases are either removed from theatmosphere or undergo chemical conversion in the troposphere,
stratosphere, or mesosphere. The time to remove or convert about
63% of a gas is often called its atmospheric lifetime. Lifetimes varyfrom less than 1 year to 100 years for the principal chlorine- andbromine-containing gases (see Table Q6-1). The long-lived gases
are converted to other gases primarily in the stratosphere and essentially all of their original halogen content becomes available toparticipate in the destruction of stratospheric ozone. Conversely,
gases with short lifetimes such as methyl bromide, methyl chloride,and some HCFCs are converted to other gases in the troposphere,
which are then removed from the atmosphere by rain and snow.
Therefore, only a fraction of their halogen content contributes to18

Section II: The ozone depletion process

Ozone Depletion Potential (ODP). The effectiveness of halogen source gases at destroying stratospheric ozone is given by the
ODP (see Table Q6-1 and Q17). A gas with a larger ODP destroys
more stratospheric ozone than a gas with a smaller ODP. The calculation of ODP requires the use of computer models that simulate
atmospheric ozone and is found relative to CFC-11, which has an
ODP defined to be 1. The ODP of a gas is based upon a comparison of the amount of ozone depletion caused by the continuous
emission to the atmosphere of a certain mass of that gas, relative
to the amount of ozone depletion following emission of the same
mass of CFC-11. Halogen source gases controlled by the Montreal
Protocol have a wide range of ODPs. Halon-1211 and halon-1301
have ODPs significantly larger than that of CFC-11 and most otherchlorinated 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 smaller values of ODP generally
have shorter atmospheric lifetimes or contain fewer chlorine andbromine atoms compared to gases with larger ODPs.

HFCs & other fluorine-containing gases. Many of the sourcegases in Figure Q6-1 also contain fluorine, another halogen, in addition to chlorine or bromine. After the source gases undergo conversion in the stratosphere (see Q5), the fluorine content of thesegases is left in chemical forms that do not cause ozone depletion.
As a consequence, halogen source gases that contain fluorine andno other halogens are not classified as ODSs. An important example of these are the HFCs, which are included in Table Q6-1 because they are common ODS substitute gases. HFCs do not contain chlorine or bromine and, consequently, all HFCs have an ODPof zero.
Many HFCs are strong greenhouse gases, as quantified by a metrictermed the Global Warming Potential (GWP) (see Q17). The Kigali
Amendment to the Montreal Protocol now controls the production
and consumption of HFCs (see Q19), especially those HFCs withhigh GWPs. As a result, industry has transitioned in part to production and use of a subset of HFCs with very low GWPs known ashydrofluoroolefins (HFOs), which are also composed of hydrogen,fluorine, and carbon atoms. Here, the “O” stands for olefin, a term
used by chemists to refer to the double carbon bond of these compounds that results in small tropospheric lifetimes and GWPs for
HFOs. One such HFO, HFO-1234yf (CF3CFCH2), has a GWP of lessthan 1 due to its 12 day lifetime.

2022 Update | Twenty Questions
Iodine containing gases. Iodine is a component of several gases

Q6

Q20). Important examples are methane (CH4 ), which reacts in thestratosphere to form water vapor and reactive hydrogen, and nitrous oxide (N2O), which reacts in the stratosphere to form nitrogen oxides. These reactive products participate in the destructionof stratospheric ozone. Increased levels of atmospheric carbon
dioxide (CO2 ) alter stratospheric temperature and winds, which
also affect the abundance of stratospheric ozone. Should futureatmospheric abundances of CO2 , CH4 , and N2O increase significantly relative to present-day values, these increases will affectfuture levels of stratospheric ozone through combined effects on
temperature, winds, and chemistry (see Figure Q20-2). Efforts areunderway to reduce the emissions of these gases under the Paris
Agreement of the United Nations Framework Convention on Climate Change because they cause surface warming (see Q18 andQ19). Although past emissions of ODSs still dominate global ozone
depletion today, future emissions of N2O from human activities are
expected to become relatively more important for ozone depletion as the atmospheric abundances of ODSs decline (see Q20).

that are naturally emitted from the oceans and from some humanactivities. Research on the importance of iodine for stratosphericozone is being conducted, in part, because trifluoroiodomethane(CF3I) is a possible replacement for halons in fire extinguishers andalso because CF3I has been proposed as an ingredient of lowGWP refrigerant blends. Although iodine can participate in ozone
destruction reactions, iodine-containing source gases all have very
short lifetimes, with most of the removal occurring in the lower atmosphere within a few days. Since the last assessment, there has
been an upward revision to the upper limit on the amount of iodine
reaching the stratosphere, which is now estimated to be about 1
ppt. The importance for stratospheric ozone of very short-lived
iodine containing source gases, including a possible enhancementof polar ozone depletion, remains an active area of investigation.

Other non-halogen gases. Other non-halogen gases that influence stratospheric ozone abundances have also increased in thestratosphere as a result of emissions from human activities (seeTable Q6-1. Atmospheric lifetimes, global emissions, Ozone Deletion Potentials, and Global Warming Potentials of some halogen sourcegases and HFC substitute gases.
Gas

Atmospheric Lifetime(years)Global Emissions in2020 (kt/yr) a

Ozone Depletion
Potential (ODP) b

Global Warming
Potential (GWP) b

CFC-11 (CCl3F)

52

Carbon tetrachloride (CCl4 )30

36 – 58

16410

27 – 60

0.87

CFC-113 (CCl2FCClF2 )

2150

93

1 – 130.82

6530

CFC-12 (CCl2F2 )102

3 – 480.75

12,500

Halogen Source Gases
Chlorine Gases

Methyl chloroform (CH3CCl3)

5.0

1–3

0.12

164

HCFC-141b (CH3CCl2F)

8.8

48 – 67

0.102

808

HCFC-142b (CH3CClF2 )17

15 – 23

0.057

2190

HCFC-22 (CHF2Cl)

12

284 – 403

0.038

1910

Methyl chloride (CH3Cl)

0.9

3759 – 5677

0.015

6

72

1–2

17

7430

Bromine Gases
Halon-1301 (CBrF3)
Halon-1211 (CBrClF2 )

16

1–5

7.1

1990

Methyl bromide (CH3Br)

0.8

111 – 154

0.57

2

Hydrofluorocarbons (HFCs)
HFC-23 (CHF3)

228

16 – 180

14,700

HFC-143a (CH3CF3)

52

27 – 33

05900

HFC-125 (CHF2CF3)

31

78 – 9803820

HFC-134a (CH2FCF3)

14

216 – 275

01470

HFC-32 (CH2F2 )5.3

56 – 770749

HFC-152a (CH3CHF2 )1.5

41 – 63

0

153

HFO-1234yf (CF3CFCH2 )0.03

not available

0

less than 1aIncludes both human activities (production and banks) and natural sources. Emissions are in units of kilotonnes per year (1 kilotonne = 1000 metric tons
= 1 gigagram = 109 grams). These emission estimates are based on analysis of atmospheric observations. The range of values for each emission estimate
reflects the uncertainty in estimating emissions from atmospheric observations.

b

100-year GWP. ODPs and GWPs are discussed in Q17. Values are calculated for emissions of an equal mass of each gas. ODPs given here reflect current
scientific values and in some cases differ from those used in the Montreal Protocol.

Section II: The ozone depletion process

19

Q7

What are the reactive halogen gasesthat destroy stratospheric ozone?

Chlorine and bromine containing halogen source gases that enter the stratosphere arise from both humanactivities and natural processes (Q6). When exposed to ultraviolet radiation from the Sun, these halogen
source gases are converted to other gases that also contain chlorine and bromine. Some of the gases act aschemical reservoirs, which can then be converted into ClO and BrO, the two most important reactive gasesthat participate in catalytic reactions that destroy ozone.

Halogen-containing gases present in the stratosphere can be divided into two groups: halogen source gases as well as reactive
and reservoir halogen gases (see Figure Q7-1). The source gases,
which include ozone-depleting substances (ODSs), are emitted at
Earth’s surface by natural processes and by human activities (seeQ6) and are chemically inert in the lower atmosphere. Once in thestratosphere, the halogen source gases chemically convert at different rates to form the reactive and reservoir halogen gases. The
conversion occurs in the stratosphere instead of the tropospherebecause solar ultraviolet (UV) radiation (a component of sunlight)
is needed for the breakup of these compounds, and solar UV radiation is more intense in the stratosphere than the troposphere(see Q2). Reactive gases containing the halogens chlorine andbromine participate in a series of chemical reactions that remove
stratospheric ozone (see Q8).

Reactive and reservoir halogen gases. The chemical conversion of halogen source gases, which involves solar ultraviolet radiation and other chemical reactions, produces a number of reactive
and reservoir halogen gases. These reactive and reservoir gases
contain all of the chlorine and bromine atoms originally present in
the source gases. The chlorine content of all of the reactive and reservoir gases is termed available chlorine, while the bromine content of similar gases is termed available bromine.

Stratospheric Halogen Gases
Halogen Source Gases
Ozone-Depleting Substances (ODSs)
CFCs (CFC-11, CFC-12, CFC-113)
HCFCs
Halon-1211 (CBrClF2)
Halon-1301 (CBrF3)
Methyl bromide (CH3Br)
Methyl chloride (CH3Cl)
Methyl chloroform (CH3CCl3)
Carbon tetrachloride (CCl4)
Very Short-Lived (VSL) Substances
Chloroform (CHCl3)
Dichloromethane (CH2Cl2)
Bromoform (CHBr3)
Dibromomethane (CH2Br2)
Methyl iodide (CH3I)

Reactive and Reservoir
Halogen Gases
Chemicalconversion
Solar ultraviolet (UV)radiation andchemical reactions

Hydrogen bromide (HBr)
Hydrogen chloride (HCl)
Chlorine nitrate (ClONO2)
Bromine nitrate (BrONO2)
Chlorine monoxide (ClO)
Bromine monoxide (BrO)
Chlorine atoms (Cl)Bromine atoms (Br)}

Largestreservoirs}

Reactive

Figure Q7-1. Conversion of halogen source gases. Halogen source gases containing chlorine and bromine are chemically converted toreactive and reservoir halogen gases, primarily in the stratosphere. The conversion requires solar ultraviolet radiation and a few chemical
reactions. The very short-lived substances undergo partial loss in the troposphere, so a smaller fraction of these gases reach the stratosphere compared to the other source gases. The gases produced by this chemical conversion can be grouped into the reservoir gases,
which do not directly destroy ozone, and the reactive gases, which participate in ozone destruction cycles (see Q8). A principal reactive
gas, ClO, is formed from the reservoir gases HCl and ClONO2 by reactions that occur on the surfaces of liquid and solid polar stratosphericclouds (PSCs) (see Q9).
20 Section II: The ozone depletion process

Q72022 Update | Twenty Questions
The most important reactive and reservoir chlorine and bromine
containing gases that form in the stratosphere are shown in Figure
Q7-1. Throughout the stratosphere, the most abundant are typically hydrogen chloride (HCl) and chlorine nitrate (ClONO2 ). These
two gases are considered reservoir gases because, while they do
not react directly with ozone, they can be converted to the mostreactive forms that do chemically destroy ozone. The halogens
most reactive with ozone are chlorine monoxide (ClO) and bromine monoxide (BrO) molecules, as well as chlorine and bromine(Cl and Br) atoms. A large fraction of available bromine is generally
in the form of BrO, whereas usually only a small fraction of available
chlorine is in the form of ClO. The unusually cold conditions that
occur in the polar regions during winter cause the reservoir gases
HCl and ClONO2 to undergo nearly complete conversion to ClO
and related reactive gases. This conversion occurs through chemical reactions that take place on the surface or within polar stratospheric cloud (PSC) particles (see Q9).

to within about 10% from the surface to above 50 km (31 miles) altitude. In the troposphere, total chlorine is contained almost entirelyin the source gases described in Figure Q6-1. At higher altitudes,
the source gases become a smaller fraction of total chlorine as they
are converted to the reactive and reservoir chlorine gases. At thehighest altitudes, total chlorine is all in the form of reactive and reservoir chlorine gases.
In the altitude range of the ozone layer at midlatitudes, as shownin Figure Q7-2, the reservoir gases HCl and ClONO2 account formost of the available chlorine. The abundance of ClO, the most important reactive gas in ozone depletion, is a small fraction of total
chlorine. The abundance of ClO peaks in the upper stratosphereabout 40 km (24.9 miles) above the surface. In this region of theatmosphere the abundance of ozone reached a minimum in thelate 1990s, at about the time the abundance of ClO in the upper
stratosphere maximized. In the lower and middle stratosphere (altitudes below about 30 km, or 18.6 miles above the surface), thelow abundance of ClO tends to limit the amount of ozone destruction that occurs outside of polar regions.

Chlorine at midlatitudes. Reactive and reservoir chlorine gases
have been observed extensively in the stratosphere using both
local and remote measurement techniques, including observations from satellite instruments. The measurements from space
displayed in Figure Q7-2 are representative of how the amountsof chlorine-containing gases change between the surface and theupper stratosphere at middle to high latitudes. Total chlorine (seered line in Figure Q7-2) is the sum of chlorine contained in halogen
source gases (e.g., CFC-11, CFC-12) and in the reservoir and reactive gases (e.g., HCl, ClONO2, and ClO). Total chlorine is constant

Chlorine in polar regions. Chlorine gases in polar regions undergo large changes between autumn and late winter. Meteorological and chemical conditions in both polar regions are now
routinely observed from space in all seasons. Maps of autumn andlate winter conditions at an altitude of 18 km (11.2 miles), near thecenter of the ozone layer (see Figure Q11-3) over the Antarctic arecontrasted in Figure Q7-3. These observations document dramatic differences in chemistry and temperature for these two seasons.

Measurements of Chlorine Gases
Annual mean 2006 (30°–70°N)
50

30

Hydrogenchloride (HCl)25

ClONO2

Total
chlorine

20

20

15

Altitude (miles)

30

Other
chlorinegases

Ozone Layer
Stratosphere

ClO

40

10
Chlorine source gases(CFCs, HCFCs, CCl4, CH3Cl, etc.)

10(The unit “parts per trillion” is defined in the caption of Figure
Q6-1.)

0

0

1000

2000

3000

Troposphere

Altitude (kilometers)

Figure Q7-2. Chlorine gas observations. The abundances of chlorine source gases as well as reactive andreservoir chlorine gases as measured in 2006, averaged
over 30° to 70°N latitude, are displayed as a function ofaltitude. In the troposphere (below about 12 km), all of
the measured chlorine is contained in the source gases.
In the stratosphere, the total chlorine content of reactive
and reservoir gases (termed available chlorine) increases
with altitude as the amount of chlorine source gases declines. This transition is a consequence of chemical reactions initiated by solar ultraviolet radiation that convert
source gases to available chlorine (see Figure Q7-1). The
principal reactive and reservoir chlorine gases formed are
HCl, ClONO2, and ClO. Adding up the source gases withavailable chlorine gives “Total chlorine”, which is nearly
constant with altitude throughout the atmosphere. In
the midlatitude ozone layer (15–35 km), chlorine sourcegases are still present and HCl and ClONO2 constitute themost abundant forms of available chlorine.

5

0
4000

Chlorine (parts per trillion)
Section II: The ozone depletion process

21

Q7

Twenty Questions | 2022 Update

Ozone values are high over the entire Antarctic continent during
autumn in the Southern Hemisphere. 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 tothis reservoir gas has occurred in the stratosphere. In the 1980s
and early 1990s, the abundances of the reservoir gases HCl and
ClONO2 increased substantially in the stratosphere following increased emissions of halogen source gases. HNO3 is an abundant,
primarily naturally-occurring stratospheric compound that plays amajor role in stratospheric ozone chemistry by both moderatingozone destruction and condensing to form polar stratosphericclouds (PSCs), thereby enabling conversion of chlorine reservoirs
gases to ozone-destroying forms (see Q9). The low abundance of
ClO indicates that little conversion of the reservoir to reactive gases
occurs in autumn, thereby limiting chemical ozone destruction.

ozone depletion to continue in the Antarctic region until midOctober (spring), when the lowest ozone values usually are observed (see Q10). 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 Q9), and ozone destruction is curtailed.

By late winter (September), a remarkable change in the composition of the Antarctic stratosphere has taken place. Low amounts ofozone reflect substantial depletion at 18 km altitude over an arealarger 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 Q11-3). The
meteorological and chemical conditions in late winter, characterized by very low temperatures, very low HCl and HNO3, and very
high ClO, are distinctly different from those found in autumn. Low
stratospheric temperatures occur during winter, when solar heating is reduced. Low HCl and high ClO reflect the conversion of
the halogen reservoir compounds, HCl and ClONO2, to the mostimportant reactive form of chlorine, ClO. This conversion occurs
selectively in winter on PSCs, which form at very low temperatures(see Q9). Low HNO3 is indicative of its condensation to form these
PSCs, some of which subsequently fall to lower altitudes throughgravitational settling. High abundances of ClO generally cause

Bromine observations. Fewer measurements are available forreactive and reservoir bromine gases in the lower stratospherethan for chlorine gases. This difference arises in part because of
the lower abundance of bromine, which makes quantification of
its atmospheric abundance more challenging. The most widely
observed bromine gas is BrO, which can be observed from space.
Estimates of the concentration of available bromine in the stratosphere are higher than expected from the decomposition of
halons and methyl bromide, the most important bromine
source gases that are produced by human activities. Thisdifference was the first direct evidence that very short-lived (VSL)
bromine-containing source gases reach the stratosphere. Subsequently, direct observations of VSL source gases have confirmed their importance. In 2020, slightly more than one-quarterof the total stratospheric bromine is supplied by these naturally
occurring, VSL source gases (see Q6).

Similar changes in meteorological and chemical conditions are also
observed between autumn and late winter for some years in the
Arctic, leading to substantial ozone loss. In spring of 2020, Arctic
ozone reached exceptionally low values. A very stable, cold, andlong-lived stratospheric Arctic vortex enabled halogen-catalyzed
chemical ozone loss that exceeded the previous record-breaking
loss observed in spring 2011 (see Q11). Substantial chemical loss ofArctic ozone will continue to occur in cold winters/springs, as longas the concentrations of ODSs are well above natural levels.

Figure Q7-3. Chemical conditions in the ozone layer over Antarctica. Observations of the chemical conditions in the Antarctic region
highlight the changes associated with the formation of the ozone hole. Satellite instruments routinely monitor ozone, reactive and reservoir chlorine gases, and temperatures in the global stratosphere. Satellite observations are shown here for autumn (May) and late winter(September) seasons in the Antarctic region, for a narrow altitude region near 18 km (11.2 miles) within the ozone layer (see Figure Q11-3).
Ozone over Antarctica has naturally high values in autumn, before the onset of ozone destruction reactions that drive widespread depletion. The high ozone is accompanied by moderate temperatures, large values of the reservoir gases HCl and HNO3, and very low amountsof reactive ClO. When the abundance of ClO is low, significant ozone destruction from halogens does not occur. Chemical conditions arequite different in late winter when ozone undergoes severe depletion. Temperatures are much lower, HCl has been converted to ClO (themost important reactive chlorine gas), and HNO3 has been removed by the gravitational settling of polar stratospheric cloud particles. The
abundance of ClO closely surrounding the South Pole is low in September because formation of ClO requires sunlight, which is still gradually returning to the most southerly latitudes. The high values of ClO in late winter cover an extensive area that at times exceeds that of the
Antarctic continent and can last for several months, leading to efficient destruction of ozone in sunlit regions in late winter/early spring.
Ozone typically reaches its minimum values in early to mid-October (see Q11). Note that the first and last colors in the color bar represent
values outside the indicated range of values.
(The unit “parts per billion,” abbreviated “ppb,” is used here as a measure of the relative abundance of a substance in dry air: 1 part per billion equals thepresence of one molecule of a gas per billion (=109) total air molecules (compare to ppt in Figure Q6-1).)

22 Section II: The ozone depletion process2022 Update | Twenty Questions

Q7

Chemical Conditions Observed in the Ozone Layer Over Antarctica
Normal ozone amounts in autumn (1 May 2021) at 18 km altitude

Temperature

Nitric acid (HNO3)

Hydrogen chloride (HCl)

Chlorine monoxide (ClO)

Ozone

Large ozone depletion in late winter (15 September 2021) at 18 km altitude

Temperature

Nitric acid (HNO3)

Hydrogen chloride (HCl)

Chlorine monoxide (ClO)

Ozone

Temperatures and chemical abundances
Low
750
– 95°
0.0
0.0
0.0

High
Ozone
Temperature
HNO3
HCl
ClO

3600 ppb– 50° C12.0 ppb
2.2 ppb
2.2 ppb

Section II: The ozone depletion process 23

Q8What are the chlorine and bromine
reactions that destroy stratosphericozone?

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 chlorineor 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 important 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 Q7-1). The most reactive of thesegases are chlorine monoxide (ClO), bromine monoxide (BrO), andchlorine and bromine atoms (Cl and Br). These gases participate inthree principal reaction cycles that destroy ozone.

Cycle 1. Ozone destruction Cycle 1 is illustrated in Figure Q8-1.
The cycle is made up of two basic reactions: ClO + O and Cl + O3.
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 reformed. In

this way, one Cl atom participates in many cycles, destroying manyozone molecules. For typical stratospheric conditions at middle
or low latitudes, a single chlorine atom can destroy thousandsof ozone molecules before it happens to react with another gas,
breaking the catalytic cycle. During the total time of its stay in thestratosphere, a chlorine atom can thus destroy many thousands ofozone molecules.

Polar Cycles 2 and 3. The abundance of ClO is greatly increasedin polar regions during late winter and early spring, relative to other seasons, as a result of reactions on the surfaces of polar stratospheric clouds (see Q7 and Q9). Cycles 2 and 3 (see Figure Q8-2)

Ozone Destruction Cycle 1 : Upper Stratosphere
Cl
O2

O3
ClO + OREACTION

Chlorine
catalytic
cycle

Cl + O3
REACTION

Ozone
destruction

O2
O

ClO
ClO + OCl + O3
NET : O + O3

24 Section II: The ozone depletion process

Cl + O2
ClO + O2
2 O2

Figure Q8-1. Ozone destruction Cycle 1. The destruction of ozone in Cycle 1 involves two separate chemical
reactions. The cycle can be considered to begin with either ClO or Cl.
When starting with ClO, the first reaction is ClO with O to form Cl and O2.
Then, Cl reacts with O3 and reforms
ClO, consuming O3 in the process andforming another O2. The net or overall reaction is that of atomic oxygen(O) with ozone (O3 ), forming two oxygen molecules (O2 ). The cycle thenbegins again with another reactionof ClO with O. Chlorine is considereda catalyst for ozone destruction because Cl and ClO are reformed eachtime the reaction cycle is completed,and hence available for further destruction of ozone. Atomic oxygen
is formed when solar ultraviolet (UV)radiation reacts with O3 and O2 molecules (see Figure Q1-3). Cycle 1 ismost important in the stratosphere attropical and middle latitudes, where
solar UV radiation is most intense.

2022 Update | Twenty Questions

Q8

Ozone Destruction Cycles
2 and 3 : Polar Regions
CYCLE 2 :ClO + ClO
(ClO)2 + sunlightClOO
2(Cl + O3(ClO)2
ClOO + Cl
Cl + O2
ClO + O2)

NET: 2O3

3O2

ClO + BrO

Cl + Br + O2

CYCLE 3 :+ BrO(BrClClO+ sunlightorCl + O3
Br + O3

NET: 2O3(

BrCl + O2
Cl + Br
ClO + O2
BrO + O2

Figure Q8-2. Polar ozone destruction Cycles 2 and 3. Significant destruction of ozone occurs during late winter and early spring in the polar
regions when abundances of ClO reach large values. In this case, thecycles initiated by the reaction of ClO with another ClO (Cycle 2) or thereaction of ClO with BrO (Cycle 3) efficiently destroy ozone. The net reaction in both cases is two ozone (O3 ) molecules forming three oxygen(O2 ) molecules. The reaction of ClO with BrO has two pathways to form
the Cl and Br product gases that lead to loss of ozone. The destructionof ozone by Cycles 2 and 3 is catalytic, as illustrated for Cycle 1 in Figure Q8-1, because chlorine and bromine gases react and are re-formed
each time the reaction cycle is completed. Sunlight is required to complete each cycle and to help form and maintain elevated abundances of
ClO. During polar night and other periods of darkness, ozone cannot
be destroyed by these reactions.

3O2

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 that produce either Cl and Br or BrCl. The net result ofboth cycles is to destroy two ozone molecules and create threeoxygen molecules. Cycles 2 and 3 account for most of the ozone
loss observed in the stratosphere over the Arctic and Antarctic
regions in the late winter/early spring season (see Q10 and Q11).
At high ClO abundances, the rate of polar ozone destruction canreach 2 to 3% per day.

Sunlight requirement. Sunlight is required to complete andmaintain these reaction cycles. Cycle 1 requires ultraviolet (UV)radiation (a component of sunlight) that is strong enough to breakapart molecular oxygen into atomic oxygen. Cycle 1 is most important in the stratosphere at altitudes above about 30 km (18.6
miles), where solar UV-C radiation (100 to 280 nanometer (nm)wavelengths) is most intense (see Figure Q2-1).
Cycles 2 and 3 also require sunlight. In the continuous darknessof winter in the polar stratosphere, reaction Cycles 2 and 3 cannot
occur. Sunlight is needed to break apart (ClO)2 and BrCl, resulting
in abundances of ClO and BrO large enough to drive rapid loss ofozone by Cycles 2 and 3. These cycles are most active when sunlight returns to the polar regions in late winter/early spring. Therefore, the greatest destruction of ozone occurs in the partially to fully
sunlit periods after midwinter in the polar stratosphere.

Sunlight in the UV-A (315 to 400 nm wavelengths) and visible (400
to 700 nm wavelengths) parts of the spectrum needed in Cycles 2
and 3 is not sufficient to form ozone because this process requires
more energetic solar UV-C solar radiation (see Q1 and Q2). In
the late winter/early spring, only UV-A and visible solar radiation
is present in the polar stratosphere, due to low Sun angles. As aresult, the rate of ozone destruction by Cycles 2 and 3 in the sunlit
polar stratosphere during springtime greatly exceeds the rate ofozone production.

Other reactions. Global abundances of ozone are controlled
by many other reactions (see Q1). Reactive hydrogen and reactive nitrogen gases, for example, are involved in catalytic ozonedestruction cycles, similar to those described above, that also takeplace in the stratosphere. Reactive hydrogen is supplied by thestratospheric decomposition of water (H2O) and methane (CH4 ).
Methane emissions result from both natural sources and a wide
variety of human activities. The abundance of stratospheric H2O iscontrolled by the temperature of the upper tropical troposphere aswell as the decomposition of stratospheric CH4. Reactive nitrogenis supplied by the stratospheric decomposition of nitrous oxide(N2O), also emitted by natural sources and human activities. The
importance of reactive hydrogen and nitrogen gases in ozone
depletion relative to reactive halogen gases is expected to increasein the future because the atmospheric abundances of the reactive
halogen gases are decreasing as a result of the Montreal Protocol,
while abundances of CH4 and N2O are projected to increase due
to various human activities (see Q20).

Section II: The ozone depletion process 25

Q9

Why has an “ozone hole” appeared over
Antarctica when ozone-depleting substances
are present throughout the stratosphere?

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 particular 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. Specific chemical reactions that occur on PSCs, combined with theisolation of polar stratospheric air inside the polar vortex, allow chlorine and bromine reactions to produce
the ozone hole over Antarctica in springtime.

The severe depletion of stratospheric ozone in late winter and early spring in the Antarctic is known as the “ozone hole” (see Q10).
The ozone hole appears over Antarctica because meteorologicaland chemical conditions unique to this region increase the effectiveness of ozone destruction by reactive halogen gases (see Q7
and Q8). The formation of the Antarctic ozone hole requires the
combination of temperatures low enough to form polar strato-

spheric clouds (PSCs), isolation of polar vortex air from air in otherstratospheric regions, sunlight, and sufficient amounts of available
chlorine (see Q8).

Distribution of halogen gases. Halogen source gases that are
emitted at Earth’s surface and have lifetimes longer than about 1
year (see Table Q6-1) are present in comparable amounts through-Minimum Air Temperaturesin the Polar Stratosphere
Arctic winter
Nov

Dec

Jan

Feb

–55

Average of daily minimum values
Arctic, 1978/1979 to 2020/2021
Antarctic, 1979 to 2021

–60

Range of daily minimum values
PSC formation temperature

March

April

50° to 90° Latitude

–60
–70
–80

–65
–90
–70
–100

–75

–110

–80

Arctic

–120

–85

–130

–90

Antarctic
–95

–140
May

June

July

Aug

Antarctic winter26 Section II: The ozone depletion process

Sept

Oct

Temperature (°F)

Temperature (°C)–50

Figure Q9-1. Arctic and Antarctic temperatures.
Air temperatures in the lower stratosphere of bothpolar regions reach minimum values in the lower
stratosphere in their respective winter seasons. Average daily minimum values over Antarctica are as
low as −92°C in July and August in a typical year.
Over the Arctic, average minimum values are near
−80°C in late December and January. Polar stratospheric clouds (PSCs) are formed in the ozone layer when winter minimum temperatures fall below
their formation threshold of about −78°C. This occurs on average for 1 to 4 months over the Arctic
and about 5 months over Antarctica each year (seeheavy orange and magenta lines). Chemical reactions within and on the surfaces of liquid and solid PSC particles cause the highly reactive chlorine
gas ClO to be formed, which catalytically destroysozone (see Q8). The range of winter minimum temperatures found in the Arctic is much greater thanthat in the Antarctic. In some years, PSC formation
temperatures are not reached in the Arctic, andsignificant ozone depletion does not occur. In contrast, PSC formation temperatures are always present for many months somewhere in the Antarctic,
and severe ozone depletion occurs each winter/spring season (see Q10).

2022 Update | Twenty Questions
out the stratosphere in both hemispheres, even though most of theemissions occur in the Northern Hemisphere. The stratosphericabundances are comparable in both hemispheres because mostlong-lived source gases have no significant natural removal processes in the lower atmosphere, and because winds and convection redistribute and mix air efficiently throughout the troposphereon the timescale of weeks to months. Halogen source gases enter
the stratosphere primarily from the tropical upper troposphere.
Stratospheric air motions then transport these gases upward andtoward 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 geographicalregions, and for an extended period of time. Low temperatures
are important because they allow liquid and solid PSC particles toform. Chemical reactions within and on the surfaces of these PSC
particles initiate a remarkable increase in the most important reactive chlorine gas, chlorine monoxide (ClO) (see below as well as Q7
and Q8). Air is usually too warm to enable the formation of clouds
in the stratosphere. Only the polar regions during winter have temperatures low enough for stratospheric clouds to form, because
air cools due to lack of sunlight. In the Antarctic winter, minimum
daily temperatures are generally much lower and less variable
than those in the Arctic winter (see Figure Q9-1). Antarctic temperatures also remain below PSC formation temperatures for much
longer periods during winter. These and other meteorological differences occur because of variations between the hemispheres inthe distributions of land, ocean, and mountains at middle and highlatitudes. As a consequence, winter temperatures are low enoughfor PSCs to form somewhere in the Antarctic for nearly the entire
winter (about 5 months), and only for limited periods (about 1 to 4
months) in the Arctic for most winters.

Q9

most important reactive form, ClO (see Figure Q7-3). The abundance of ClO increases from a small fraction of available chlorine
to comprise more than half of all available chlorine (see Q7). With
increased ClO, the catalytic cycles involving ClO and BrO become
active in the chemical destruction of ozone whenever sunlight isavailable (see Q8).
Different types of liquid and solid PSC particles form when stratospheric temperatures fall below about −78°C (−108°F) in polar
regions (see Figure Q9-1). As a result, PSCs are often found overlarge areas of the winter polar regions and over extensive altitude
ranges in both hemispheres, with substantially larger regions andfor longer time periods in the Antarctic than in the Arctic. The mostcommon type of PSC forms from nitric acid (HNO3 ) and water

Arctic Polar Stratospheric Clouds (PSCs)

Isolated conditions. Stratospheric air in the polar regions is relatively isolated for long periods in the winter months. This isolation results from strong winds that encircle the poles during winter,
forming a polar vortex, which prevents substantial transport andmixing of air into or out of the polar stratosphere. This circulation
strengthens in winter as stratospheric temperatures decrease. The
polar vortex circulation tends to be stronger in the Southern Hemisphere (SH) than in the Northern Hemisphere (NH), because northern latitudes have more mountainous regions and adjacent areasof ocean and land with contrasting temperatures than is present at
southern latitudes. This situation leads to more meteorological disturbances in the circulation of the NH, which increase the mixing
in of air from lower latitudes toward the pole, warming the Arctic
stratosphere. Since winter temperatures are therefore lower in theSH than in the NH polar stratosphere, the isolation of air in the polar 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 such as an increase in theabundance of ClO (see Q8). These changes persist for many weeks
to months due to the isolation of stratospheric air in the Antarctic.

Polar stratospheric clouds (PSCs). Chemical reactions within
and on the surfaces of liquid and solid PSC particles can substantially increase the relative abundances of reactive chlorine gases.
These reactions convert the reservoir forms of chlorine gases,
hydrogen chloride (HCl) and chlorine nitrate (ClONO2 ), to the

Figure Q9-2. Polar stratospheric clouds. This photograph of
an Arctic polar stratospheric cloud (PSC) was taken in Kiruna,
Sweden (67°N), on 27 January 2000. PSCs form in the ozone
layer during winters in the Arctic and Antarctic, wherever sufficiently low temperatures occur (see Figure Q9-1). The particles
grow from the condensation of water, nitric acid (HNO3 ), andsulfuric acid (H2SO4 ). The clouds often can be seen with the human eye when the Sun is near the horizon. Reactions within andon PSC surfaces lead to the formation of the highly reactive gas
chlorine monoxide (ClO), which is very effective in the chemical
destruction of ozone (see Q7 and Q8).
Section II: The ozone depletion process 27

Q9 Twenty Questions | 2022 Update
condensing on pre-existing liquid sulfuric acid (H2SO4 )-containing
particles. Some of these particles freeze to form solid particles. At
even lower temperatures (−85°C or −121°F), water condenses toform ice particles. PSC particles grow large enough and are numerous enough that cloud-like features can be observed from theground under certain conditions, particularly when the Sun is near
the horizon (see Figure Q9-2). PSCs are often found near mountain ranges in polar regions because the motion of air over themountains can cause localized cooling in the stratosphere, which
increases condensation of water and HNO3.
When average temperatures begin increasing in late winter, PSCs
form less frequently, which slows down the conversion of chlorine
from reservoir to reactive forms throughout the polar region. Without continued production, the abundance of ClO decreases asother chemical reactions re-form the reservoir gases, ClONO2 andHCl. When temperatures rise above PSC formation thresholds,
usually sometime between late January and early March in the Arctic and by mid-October in the Antarctic (see Figure Q9-1), the mostintense period of ozone depletion ends.

Nitric acid and water removal. Once formed, the largest PSC
particles fall to lower altitudes because of gravity. The largest particles can descend several kilometers in the stratosphere within a few
days during the low-temperature winter/spring period. Because
PSCs often contain a significant fraction of available HNO3, theirdescent removes HNO3 from regions of the ozone layer. This process is called denitrification of the stratosphere. Because HNO3 is asource for nitrogen oxides (NOx ) in the stratosphere, denitrification
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 increasingchemical ozone destruction. Significant denitrification occurs each
winter in the Antarctic and only for occasional winters in the Arctic,
because PSC formation temperatures must be sustained over an
extensive altitude region and time period to lead to denitrification(see Figure Q9-1).
Ice particles form at temperatures that are a few degrees lower
than those required for PSC formation from HNO3. If these water
ice particles grow large enough, their gravitational settling can remove a significant fraction of water vapor from regions of the ozone
layer over the course of a winter. This process is called dehydrationof the stratosphere. Because of the very low temperatures required
to form ice, dehydration is common in the Antarctic and rare in the
Arctic. 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 the production of ClO by reactions on PSCs.

Discovering the role of PSCs. Ground-based observations of
PSCs were available many decades before the role of PSCs in polarozone destruction was recognized. The geographical and altitudinal extent of PSCs in both polar regions was not fully known until
PSCs were observed by satellite instruments starting in the late1970s. The role of PSC particles in converting reservoir chlorine
gases to ClO was not understood until after the discovery of the
Antarctic ozone hole in 1985. Our understanding of the chemicalrole of PSC particles developed from laboratory studies of their surface reactivity, computer modeling studies of polar stratosphericchemistry, and measurements that sampled particles and reactive
chlorine gases, such as ClO, in the polar stratosphere.

The Discovery of the Antarctic Ozone HoleThe 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 Q4) installed as part of the effort
to increase observations of Earth’s atmosphere during the International Geophysical Year that began in 1957 (see Figure Q0-1).
The observations showed unusually low total ozone during the austral late winter/early spring months of September, October, andNovember. Total ozone was lower in these months in the early 1980s 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 prestigious scientificjournal Nature in 1985. They suggested that rising abundances of atmospheric CFCs were the cause of the steady decline in totalozone over the Halley Bay research station (76°S) observed during successive Octobers starting in the early 1970s. Soon after, satellite measurements confirmed the spring ozone depletion and further showed that for each late winter/early spring season startingin the early 1980s, the depletion of ozone extended over a large region centered near the South Pole. The term “ozone hole” cameabout as a description of the very low values of total ozone, apparent in satellite images, that extend over the Antarctic continent
for many weeks each October (spring in the Southern Hemisphere) (see Q10). 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 first 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
observed in the Antarctic spring were found to be around 300 Dobson units (DU), lower than those in the Arctic spring. The Antarctic
values were surprising because the assumption at the time was that the two polar regions would have similar values. We now knowthat these 1950s Antarctic values were not anomalous; in fact, similar values were observed near the South Pole in the early 1970s,
before the ozone hole appeared (see Figure Q10-3). Antarctic total ozone values in early spring are systematically lower than thosein the Arctic early spring because the Southern Hemisphere polar vortex is much stronger and colder and, therefore, much more
effective in reducing the transport of ozone-rich air from midlatitudes to the pole (compare Figures Q10-3 and Q11-2).

28 Section II: The ozone depletion process

Q10

How severe is the depletion of
the Antarctic ozone layer?

Severe depletion of the Antarctic ozone layer was first reported in the mid-1980s. Antarctic ozone depletionis 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 stratospheric altitudes, thereby
reducing total ozone by as much as two-thirds at some locations. This severe depletion creates the “ozonehole” apparent in images of Antarctic total ozone acquired using satellite instruments. 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 “ozonehole”, was first reported in the mid-1980s (see box in Q9). The
depletion is attributable to chemical destruction by reactive halogen gases (see Q7 and Q8), which increased everywhere in thestratosphere in the latter half of the 20th century (see Q15). Conditions in the Antarctic winter and early spring stratosphere enhanceozone depletion because of (1) the long periods of extremely low
temperatures during polar night, which cause polar stratosphericclouds (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 after the return of sunlight. The severity of Antarctic
ozone depletion as well as long-term changes can be seen using
satellite observations of total ozone and profiles of ozone versus
altitude.

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 springmeasurements shows a large region centered near the South
Pole in which total ozone is highly depleted (see Figure Q10-1).

Antarctic Ozone Hole

Figure Q10-1. Antarctic ozone hole. Total ozone values asmeasured by a satellite instrument are shown for high southernlatitudes averaged over 21 to 30 September 2021. The dark
blue and purple regions over the Antarctic continent show thesevere ozone depletion or “ozone hole” now found every austral spring. Minimum values of total ozone inside the ozone hole
are close to 150 Dobson units (DU), compared with Antarctic
springtime values of about 350 DU observed in the early 1970s(see Figure Q10-3). The area of the ozone hole is usually definedas the geographical region within the 220-DU contour (thickwhite line) on total ozone maps. The maximum values of totalozone in the Southern Hemisphere in late winter/early springare generally located in a crescent-shaped region (better seen in
Figure Q10-3) that surrounds and is isolated from the ozone hole,due to stratospheric winds at the boundary of the polar vortex.
In late spring or early summer (November–December), these
atmospheric winds weaken and the ozone hole disappears dueto the cessation of the ozone-depleting chemical reactions andthe transport of ozone-enriched air masses towards the pole.

21 – 30 September 2021
100

200

300

400

500

600

Total Ozone (Dobson units)
Section III: Stratospheric ozone depletion 29

Q10 Twenty Questions | 2022 Update

Area of depletion(million square kilometers)

Antarctic Ozone Depletion
30

Area equivalent
to the size of:Africa25

North
America

20

South
America
Antarctica
EuropeAustralia

15
10

Average daily area of O3 < 220 DU(21–30 September)

5(Dobson units)

Minimum total ozone

0
Average daily minimum(21 September–16 October)

200
150
100
1980

1990

2000

2010

Year

This region has come to be called the “ozone hole” because of
the near-circular contours of low ozone values in the maps. The
reported area of the ozone hole for a given year is defined here
as the geographical region within the 220-Dobson unit (DU) contour in total ozone maps (see white line in Figure Q10-1) averaged
between 21–30 September. The area reached a maximum of 28
million square km (about 11 million square miles) in 2006, which
is more than twice the area of the Antarctic continent (see Figure
Q10-2). Minimum values of total ozone inside the ozone hole averaged in late September to mid-October are near 120 DU, which isonly one-third of the springtime values of about 350 DU observedin the early 1970s (see Figures Q10-3 and Q11-1). Low total ozone
inside the ozone hole contrasts strongly with the distribution ofmuch larger values outside the ozone hole. This common feature
can be seen in Figure Q10-1, where a broad geographic region
with values of total ozone around 350 DU surrounds the ozone
hole in September 2021, revealing the edge of the polar vortex
that acts as a barrier to the transport of ozone-rich midlatitude air
into the polar region (see Q9).

Altitude profiles of Antarctic ozone. The low total ozone values within the ozone hole are caused by nearly complete removalof ozone in the lower stratosphere. Balloon-borne ozone instruments (see Q4) demonstrate that this depletion occurs within theozone layer, the altitude region that normally contains the highest
abundances of ozone. At geographic locations with the lowest
total ozone values, ozonesonde measurements show that thechemical destruction of ozone has often been complete over an altitude region of up to several kilometers. For example, in the ozone
profile over South Pole, Antarctica on 10 October 2020 (see red
30 Section III: Stratospheric ozone depletion

2020

Figure Q10-2. Antarctic ozone holefeatures. Long-term changes are shownfor key aspects of the Antarctic ozone
hole: the area enclosed by the 220-DU
contour on maps of total ozone (upper
panel) and the minimum total ozone
amount measured over Antarctica (lower panel). The values are from satellite
observations and averaged for each year
at a time near the peak of ozone depletion, as defined by the dates shown ineach panel. The areas of continents areincluded for reference in the upper panel. The magnitude of Antarctic ozone depletion gradually increased beginning in
1980. In the past two and a half decades
the depletion reached steady year-toyear values, except for the unusuallysmall amount of depletion in 2002 and2019 (see Figure Q10-4 and followingbox). The magnitude of Antarctic ozone
depletion will steadily decline as ODSs
are removed from the atmosphere (seeFigure Q15-1). The return of Antarctic total ozone to 1980 values is expected tooccur around the mid-2060s (see Q20).

line in left panel of Figure Q11-3), ozone abundances are essentially zero over the altitude region from 14 to 20 km. The lowest
winter temperatures and highest abundances of reactive chlorine(ClO) occur in this altitude region (see Figure Q7-3). The differences in the South Pole ozone profiles averaged over 1967–1971
and over 1990–2021 in Figure Q11-3 show how reactive halogen
gases have dramatically altered the ozone layer. For the 1967–1971
time period, a normal ozone layer is clearly evident in the Octoberaverage profile, with a peak near 16 km altitude. In the 1990–2021
average profile, a broad minimum centered near 16 km is observed, with ozone values reduced at some altitudes by up to 90%
relative to normal values.

Long-term total ozone changes. Prior to 1960, the amount ofreactive halogen gases in the stratosphere was insufficient to causesignificant chemical loss of Antarctic ozone. Ground-based observations show that the steady decline in total ozone over the Halley
Bay research station (76°S) (see box in Q9) during each October
first became apparent in the early 1970s.
Satellite observations reveal that in 1979, total ozone during October near the South Pole was slightly lower than found at other highsoutherly latitudes (see Figure Q10-3). Computer model simulations indicate that Antarctic ozone depletion actually began in theearly 1960s. Until the early 1980s, depletion of total ozone was notlarge enough to result in minimum values falling below the 220DU threshold that is now commonly used to denote the boundaryof the ozone hole (see Figure Q10-1). Starting in the mid-1980s, a
region of total ozone well below 220 DU centered over the South
Pole became apparent in satellite maps of October total ozone(see Figure Q10-3). Observations of total ozone from satellite

2022 Update | Twenty Questions

Q10

Antarctic Total Ozone
October monthly averages

1971

1979

1984

1985

1989

1994

1999

2004

2009

2015

2018

2021

100

200

300

400

500

600

Total Ozone (Dobson units)
Figure Q10-3. Antarctic total ozone. Long-term changes in Antarctic total ozone are demonstrated with this series of total ozone maps derived from satellite observations. Each map is an average during October, the month of maximum ozone depletion over Antarctica. In the
1970s an ozone hole, as defined by a significant region with total ozone values less than 220 DU (thick white line), was not observed. The
ozone hole initially appeared in the early 1980s and increased in size until the early 1990s. A large ozone hole has occurred each year since
the mid-1980s as shown in Figure Q10-2. Maps from the mid-2010s show the large extent (about 25 million square km) of recent ozone
holes. The largest values of total ozone in the Southern Hemisphere during October are found in a crescent-shaped region outside of theozone hole. The satellite data show that these maximum values and their geographic extent have significantly diminished since the 1970s.

Section III: Stratospheric ozone depletion31Q10 Twenty Questions | 2022 Update
instruments can be used in multiple ways to examine how ozone
depletion has changed in the Antarctic region over the past 50
years, including:• First, ozone hole areas displayed in Figure Q10-2 show thatthe area of depletion increased after 1980, then became fairly stable in the 1990s, 2000s, and into the mid-2010s, often
reaching an area of 25 million square km (about the size ofNorth America). Exceptions are the unexpectedly low areas ofdepletion observed in 2002 and 2019, which are explained in
the box at the end of this Question.
• Second, minimum Antarctic ozone amounts displayed in Figure Q10-2 show that the severity of the depletion increased beginning around 1980 along with the rise in the ozone hole area.
Fairly constant minimum values of total ozone, near 110 DU,
were observed in the 1990s and 2000s, with the exceptions of2002 and 2019. There is some indication of an increase in theminimum value of total ozone since the early 2010s. However,in 2020 and 2021 unusually cold conditions produced large
and long-lasting ozone holes, whereas in 2019 a weather disturbance resulted in a small, shallow ozone hole (see box). The
ozone holes of 2019, 2020, and 2021 demonstrate the importance of year-to-year variability in ozone hole conditions, whichchallenges the identification of recovery due to declining levels
of ODSs.
• Third, total ozone maps over the Antarctic and surrounding regions (Figure Q10-3) show how the ozone hole has developed
over time. October averages show the absence of an ozone
hole in the 1970s, the extent of the ozone hole around the timeof its discovery in 1985 (see box in Q9), followed by the progression throughout the 1990s, the 2010s, and into the early2020s.
• Fourth, values of total ozone averaged poleward of 63°S
for each October (Figure Q11-1) show how total ozone haschanged in the Antarctic region. After decreasing steeply in the

32 Section III: Stratospheric ozone depletion

early years of the ozone hole, polar-cap averages of total ozone
are now approximately 30% smaller than those in pre-ozone
hole years (1970–1982). Increased year-to-year variability in
Antarctic-region ozone is evident since 2000.

Disappearance of the Antarctic ozone hole. Severe depletion of Antarctic ozone occurs each spring. In mid-October, temperatures in the polar lower stratosphere begin to increase (seeFigure Q9-1), eventually rising to a level that prevents the formationof PSCs and production of ClO. Consequently, the most effective
chemical cycles that destroy ozone are curtailed (see Q8). Typically,
the polar vortex breaks down during late November or early
December, ending the isolation of high-latitude air and increasing
the exchange of air between the Antarctic stratosphere and lower
latitudes. This exchange allows substantial amounts of ozone-rich
air to be transported poleward, where it displaces or mixes with
air depleted in ozone. As a result of these large-scale transportand mixing processes, the ozone hole typically disappears by midDecember.
Long-term recovery of the Antarctic ozone hole. There are
emerging indications that the size and maximum ozone depletion(severity) of the Antarctic ozone hole has diminished since 2000.
The signature of recovery is clearest during September, which is
early spring in the Southern Hemisphere. Although accounting forthe effect of natural variability on the size and depth of the ozone
hole is challenging, the weight of evidence suggests that the
decline in the amount of reactive halogen gases in the stratospheresince 2000 has made a substantial contribution to the observed
reductions in the size and depth of the ozone hole during September. The reduction of Antarctic ozone depletion leading to fullrecovery of total ozone back to the value observed in the early
1980s requires continued, sustained reductions of ozone-depleting substances in the stratosphere. Even with the halogen sourcegas reductions already underway (see Q15), the return of Antarctic
total ozone to 1980 values is not expected to occur until the mid2060s (see Q20).

2022 Update | Twenty Questions

Q10

Unusual 2019 Antarctic Ozone Hole17 September 2018100

17 September 2019
200

300

400

17 September 2020
500

600

Total Ozone (Dobson units)
Figure Q10-4. Unusual 2019 ozone hole. Views from space of the Antarctic ozone hole are shown for 17 September for three consecutive
years: 2018, 2019, and 2020. The ozone holes in 2018 and 2020 are considered typical of those observed since the late 1990s. An initiallycircular hole in 2019 was transformed by the significant transport of air from midlatitudes into the region of the ozone hole starting in lateAugust. This unusual event, initiated by meteorological disturbances in the midlatitude troposphere, transported high amounts of ozone
from midlatitudes into the region of the ozone hole and led to higher temperatures that reduced the rate of ozone depletion in 2019 compared to most other years. Consequently, the region of low total ozone in 2019 was displaced from the normal circumpolar position, andthe amount of ozone depletion in 2019 as measured by the area of the ozone hole or minimum total ozone amounts was much less than ineither 2018 or 2020 (see Figure Q10-2).

The 2019 Antarctic Ozone HoleThe 2019 Antarctic ozone hole exhibited much less ozone depletion as measured by the area of the ozone hole or minimum totalozone amounts in comparison with ozone holes in 2018 and 2020. The 2019 values, along with those observed in 2002, stand out
clearly in the year-to-year changes in these quantities displayed in Figure Q10-2. Both of these anomalies are due to unusually warm
meteorological conditions.
The ozone hole initially formed as expected in August 2019. At the end of August, an unusual weather pattern occurred that resulted
in the transport of air from midlatitudes into the region of the ozone hole. As a result of this meteorological disturbance, the regionof low total ozone was displaced from the normal circular position centered over the South Pole during September 2019 (see Figure
Q10-4). Air with anomalously large amounts of ozone was transported from midlatitudes into the Antarctic stratosphere during thefirst two weeks of September 2019, resulting in the smallest ozone hole in the 21st century and one of the smallest ever observed in
September since 1984, when the ozone hole was discovered (see Figure Q10-2).
The unexpected meteorological influence in 2019 resulted from specific atmospheric air motions that sometimes occur in polar regions. Meteorological analyses show that the 2019 disturbance was initiated by strong, large-scale weather systems that originatedin the lower atmosphere (troposphere) at midlatitudes earlier in the year. This tropospheric weather system extended poleward andupward into the stratosphere, disturbing the normal circumpolar wind (polar vortex), transporting air with high levels of ozone into
the vortex, and warming the lower stratosphere where ozone depletion was ongoing. The impact of this weather disturbance duringlate winter/early spring, when ozone loss processes are normally most effective, resulted in less Antarctic ozone depletion in 2019.
A similar stratospheric warming event occurred in 2002, resulting in ozone hole area and minimum total ozone amount values thatare comparable to those observed in 2019 (see Figure Q10-2). Another warming event in 1988 caused somewhat smaller changes
in the ozone hole features than those observed in 2002 and 2019. Large warming events are difficult to predict because of thecomplex conditions leading to their formation. The size and maximum ozone depletion (depth) of the ozone hole in 2020 and 2021
returned to values similar to those observed since the late 1990s (see Figures Q10-2 and Q10-3).

Section III: Stratospheric ozone depletion 33

Q11

Is there depletion of the Arctic
ozone layer?

Yes, significant depletion of the Arctic ozone layer now occurs in most years in the late winter and early
spring period (January– March). However, Arctic ozone depletion is less severe than that observed in the
Antarctic and exhibits larger year-to-year differences as a consequence of the more variable meteorological conditions found in the Arctic polar stratosphere. Even the most severe Arctic ozone depletion does notlead 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 larger than those in the Antarctic. Consequently,
an extensive and recurrent “ozone hole”, as found in the Antarctic stratosphere, does not occur in the Arctic.

Significant depletion of ozone has been observed in the Arctic
stratosphere in recent decades. The depletion is attributable tochemical destruction by reactive halogen gases (see Q8), which
increased in the stratosphere in the latter half of the 20th century (see Q15). Arctic depletion also occurs in the late winter/early
spring period (January–March), however over a somewhat shorter
period than in the Antarctic (July–October). Similar to the Antarctic(see Q10), Arctic ozone depletion occurs because of (1) periods ofvery low temperatures, which lead to the formation of polar stratospheric clouds (PSCs); (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 after the return of sunlight. However, these conditions occur less often in the Arctic than in the Antarctic.

sive 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 transportof ozone from its source region in the tropics to higher latitudes ismore vigorous in the Northern Hemisphere. Furthermore, ozone
depletion is limited because, in comparison to Antarctic conditions, average temperatures in the Arctic stratosphere are always
significantly higher (see Figure Q9-1) and the isolation of polar
stratospheric air is less effective (see Q9). These differences occur
because northern latitudes have a greater prevalence of mountainous regions and contrasting areas of ocean and land than southernlatitudes (compare Figures Q10-3 and Q11-2), which creates more
meteorological disturbances that warm the Arctic stratosphere(see box in Q10). Consequently, the extent and timing of Arctic
ozone depletion varies considerably from year to year.

Total ozone observed in the Arctic during spring (see Figure
Q11-1), even for years with significant ozone depletion, is higher
than is commonly observed in the Antarctic during spring. Exten-

Average Total Ozone in Polar Regions
500
Arctic region in March (63°– 90°N)

Total ozone (DU)

450
400
350
Antarctic region in October (90°– 63°S)

300
250
200

1970 –1982 observed
springtime average(over 63° – 90° )

1970

1980

1990

Year

34 Section III: Stratospheric ozone depletion

2000

2010

2020

Figure Q11-1. Average total ozone in polar regions. Long-term changes in total ozone measured with satellite instruments are shown for the
Antarctic and Arctic. The measurements are averaged over latitudes poleward of 63° for March (inthe Arctic) and for October (in the Antarctic) anddisplayed as Dobson units (DU) (see box in Q4).
The reference values (black lines) are averagesof springtime total ozone from observations acquired between 1970 and 1982. After 1982, significant ozone depletion is found in most years in
the Arctic and all years in the Antarctic. The largest
average depletions have occurred in the Antarctic
since 1990. The ozone changes arise from a combination of chemical destruction and natural variations in meteorological conditions that influence
the year-to-year values of ozone, particularly in
the Arctic. The year-to-year variation in Antarctic
total ozone has increased since 2000.

2022 Update | Twenty Questions

Q11

Arctic Total Ozone
March monthly averages

1971

1979

1984

1985

1990

1997

2001

2006

2011

2014

2020

2021

100

200

300

400

500

600

Total Ozone (Dobson units)
Figure Q11-2. Arctic total ozone. Long-term changes in Arctic total ozone are evident in this series of total ozone maps derived from satellite observations. Each map is an average during March, the month when some ozone depletion is usually observed in the Arctic. The years
chosen are shown to highlight the strong year-to-year variability of Arctic ozone as well as the tendency towards lower values over time,
during the half-century long record of satellite data. In the 1970s and 1980s, the Arctic region had normal ozone values in March, with values of 450 DU and above (red colors). Low values of total ozone on the scale of the Antarctic ozone hole do not occur in the Arctic. Instead,in late winter and early spring ozone depletion has eroded the normal high values of total ozone. For most years starting in the 1990s, thegeographic extent of ozone values of 450 DU and above is reduced in comparison with the 1970s. The large regions of low total ozone in
March of 1997, 2011, and 2020 (blue colors) are unusual in the Arctic record (see Figure Q11-1). Meteorological conditions during thesethree winters led to below-average stratospheric temperatures and strong polar vortices, conditions favorable to strong ozone depletionby reactive halogen gases.
Section III: Stratospheric ozone depletion 35

Q11

Twenty Questions | 2022 UpdateLong-term total ozone changes. Two important ways in which

Arctic ozone depletion in some winter/spring seasons occurs overmany weeks, in others only for brief periods. These differences
are controlled by the stability of and temperature within the Arctic vortex circulation system. Meteorological conditions affect thestrength of the circulation of air in the Arctic stratosphere and theamount of ozone transported to the polar region. When the Arctic
vortex is weak and disrupted, the polar region has larger amountsof ozone since temperatures are higher than normal and moreozone is transported by the winds to high latitudes. In contrast, astrong and stable vortex leads to lower-than-normal amounts of total ozone in the Arctic due to both more severe chemical loss andweaker poleward transport of ozone.

satellite observations can be used are to examine the average totalozone abundances in the Arctic region for the last half century andto contrast these values with Antarctic abundances.
• First, total ozone averaged poleward of 63°N for each Marchshows how total ozone has changed in the Arctic (see Figure
Q11-1). The seasonal poleward and downward transport ofozone-rich air is naturally stronger in the Northern Hemisphere.
As a result, before ozone depletion begins normal Arctic values
are close to 450 DU while Antarctic values are close to 330
DU. Changes in total ozone from the 1970–1982 average value

Polar Ozone Depletion
30

Antarctic ozone

Arctic ozone

South Pole (90°S)

Ny-Ålesund (79°N)

25

Altitude (km)

20

10
15

Altitude (miles)

15

10
October averages:
1967–1971 (280 DU)
1990–2021 (150 DU)

5March averages:
1967–1971 (no data available)
1990–2021 (382 DU)

5

4 April 2011 (259 DU)
28 March 2020 (273 DU)

10 October 2020 (104 DU)
South Pole elevation

0

0

5

10

15

0

5

10

15

0

Ozone abundance (mPa)
Figure Q11-3. Vertical distribution of Antarctic and Arctic ozone. At high latitudes during winter and spring, most stratospheric ozone
resides between about 10 and 30 km (6 to 19 miles) above Earth’s surface. Long-term observations of the ozone layer with balloon-borneozone instruments also allow vertical profiles of ozone during winter to be compared between the Antarctic and Arctic regions. In the
Antarctic at the South Pole (left panel), a normal ozone layer was observed to be present between 1967 and 1971. As shown here, ozone
was almost completely destroyed between 14 and 20 km (9 to 12 miles) over the South Pole on 10 October 2020. Average October valuesof ozone in the last three decades (1990–2021) are 90% lower than pre-1971 values at the peak altitude of the ozone layer (16 km). For
the Arctic, the average March profile for 1990–2021 as measured over the Ny-Ålesund research station is shown (right panel). No datafrom Ny-Ålesund are available for the 1967–1971 period. Some profiles reveal significant depletion, as shown here for 4 April 2011 and 28
March 2020. In such years, winter minimum temperatures are generally lower than normal, allowing for formation of PSCs and increases
in ClO over longer periods of time. The number in parentheses for each profile is the total column ozone value in Dobson units (DU) (seebox in Q4). Ozone abundances are shown here as the pressure of ozone at each altitude using the unit “milli-Pascals” (mPa) (100 million
mPa = atmospheric sea-level pressure).
36 Section III: Stratospheric ozone depletion

2022 Update | Twenty Questions(horizontal line in Figure Q11-1) are due to a combination ofchemical destruction by halogens and meteorological (natural)variations. In the last quarter century, these two aspects havecontributed about equally to the observed year-to-year variations of ozone in the Arctic. Decreases from pre-ozone-holeaverage values (1970–1982) were observed in the Arctic by themid-1980s, when larger changes were already occurring in the
Antarctic. The decreases in total ozone in the Arctic are generally much smaller than those found in the Antarctic and leadto total ozone values that are typically about 10 to 20% below
normal. Maximum decreases in total ozone of about 30%
observed in March 1997, 2011, and 2020 for considerable
regions of the Arctic (see Figure Q11-2) are the most comparable to Antarctic depletion. In all three Arctic winters, meteorological conditions inhibited transport of ozone-rich air tohigh latitudes, and in 2011 and 2020 persistently low temperatures facilitated severe chemical depletion of ozone by reactive
halogens (see Q8).
• Second, total ozone maps over the Arctic and surrounding
regions (see Figure Q11-2) show year-to-year changes in totalozone during March. In the 1970s, total ozone values were near
450 DU when averaged over the Arctic region in March. Beginning in the 1990s and continuing until present, values above
450 DU have been increasingly absent from the March averagemaps. A comparison of the maps in the 1970s and 2020, for example, shows a striking reduction of total ozone throughout the
Arctic region. The large geographical extent of low total ozone
in the March maps for years 1997, 2011, and 2020 represent
exceptional events in the Arctic observational record of the last
five decades, as noted above in the discussion of Figure Q11-1.
The large-scale differences between Arctic ozone distributions
observed in March 2011, 2014, 2020, and 2021 are a prime example of the influence of meteorology in driving year-to-year
variations in Arctic ozone depletion. Despite the decline in theconcentration of reactive gases in the polar stratosphere beginning around year 2000, there is little evidence of recovery
of Arctic ozone. Detection of ozone recovery for the Arctic is
made difficult by the smaller amount of depletion comparedto the Antarctic, combined with larger natural variability in the
year-to-year changes in the amount of Arctic ozone.

Altitude profiles of Arctic ozone. As in the Antarctic, Arctic
ozone is measured using a variety of instruments (see Q4). These
measurements document daily to seasonal changes within theozone layer. Spring Arctic and Antarctic balloon-borne ozone
measurements are shown in Figure Q11-3. Arctic profiles were ob-

Q11tained from the Ny-Ålesund research station at 79°N. For 1990–
2021, the March average reveals the presence of a substantial
ozone layer, contrasting sharply with the severely depleted Antarctic ozone layer in the October average over the same time period. This contrast further demonstrates how higher stratospherictemperatures, more variable meteorology, and larger amounts ofozone present in mid-winter have protected the Arctic ozone layerfrom having extended altitude regions of near-zero ozone as regularly occurs in the Antarctic, despite there being similar abundances of halogen gases (see Q7) in the two regions.
The Arctic ozone profiles shown in Figure Q11-3 for 4 April 2011
and 28 March 2020 are two of the most severely depleted in the
30-year record from Ny-Ålesund. Arctic ozone reached exceptionally low values in the spring of 2020 due to the presence of a very
stable, cold, and long-lived polar vortex that led to lower-than-normal poleward transport of ozone and enabled extensive halogencatalyzed chemical loss of ozone, which exceeded the previous
record-breaking loss in spring 2011 (see Figure Q11-1). Chemicalloss averaged over the Arctic vortex for March 2020 is estimated to
be about 120 DU, nearly equal to the difference between the totalozone values associated with the 28 March 2020 profile and the
1991–2021 average profile shown in Figure Q11-3. Chemical lossof Arctic ozone in spring 2020 is similar in magnitude to the loss
that regularly occurs over Antarctica.
Substantial chemical loss of Arctic ozone will continue to occur
during cold winters and springs, as long as the concentrations of
ODSs are well above natural levels. Although significant, the Arctic
ozone depletion during both 2011 and 2020 resulted in amountsof ozone between 15 and 20 km altitude that are substantially larger than routinely observed in the Antarctic during October, suchas in the profile from 10 October 2020 shown in Figure Q11-3. In
the Antarctic stratosphere, near-complete depletion of ozone overmany kilometers in altitude and over areas almost as large as North
America is a common occurrence (see Figure Q10-2).

Restoring ozone in spring. As in the Antarctic, ozone depletionin the Arctic is largest in the late winter/early spring season. In
spring, temperatures in the polar lower stratosphere increase (seeFigure Q9-1), halting the formation of PSCs, the production of ClO,
and the chemical cycles that destroy ozone. The breakdown of thepolar vortex ends the isolation of air in the high-latitude region,
allowing 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,
the signature of large-scale and extensive ozone depletion at highnorthern latitudes typically disappears by April or earlier.

Section III: Stratospheric ozone depletion 37

Q12

How large is the depletion of theozone layer outside of polar regions?

The abundance of total ozone between 60°S and 60°N is now about 2–3% below the amount present during
1964–1980. The abundance of total ozone in this region declined steadily throughout the 1980s due to theincreases in reactive halogen gases in the stratosphere resulting from human activities. In the early 1990s,
due to additional ozone loss that followed the 1991 eruption of Mount Pinatubo, total ozone between 60°S
to 60°N was depleted by 5% relative to the 1964–1980 average, the maximum depletion observed overthe past six decades. In both hemispheres, total ozone depletion is small near the equator and increases
toward the poles. The larger depletion at higher latitudes, particularly in the Southern Hemisphere, is driven
in part by the late winter/early spring destruction of polar ozone, which influences ozone at lower latitudes
following the breakdown of the polar vortex in each hemisphere.

Decreases in total ozone averaged over 60°S to 60°N, termed
here global ozone, first became apparent in the 1980s (see Figure
Q12-1) due to the rise in stratospheric halogens that result fromhuman activities (see Figure Q15-1). Most of the depletion has
occurred in the stratospheric ozone layer, where most ozone resides (see Figure Q1-2). Following the eruption of Mount Pinatubo
in June 1991, global ozone in the early 1990s reached a minimum
value that was 5% lower than the 1964–1980 average. This massive volcanic eruption dramatically increased the number of sulfuric acid-containing particles throughout the stratosphere. These
particles significantly increased the effectiveness of reactive halogen gases in destroying ozone (see Q13) and, thereby, decreased
global ozone by an additional 2% relative to its long-term trend forseveral years following the eruption. Since the mid-1990s, global
ozone has gradually increased, due to the recovery of the ozone
layer from this volcanically-induced perturbation as well as theslow, steady decline in stratospheric halogens driven by the Montreal Protocol (see Q14). Since 2010, global ozone has been about
2 to 3% below the 1964–1980 average.

Polar regions. Observed total ozone depletion varies significantly with latitude across the globe. The largest reductions occur at high southern latitudes as a result of the severe ozone loss
over Antarctica each year during winter/spring (see Q9 and Q10).
The next largest depletion is observed in the high latitudes of theNorthern Hemisphere, caused in part by winter losses over theArctic in some years (see Q11). Since ozone loss in polar regions is
described extensively in the answers to other questions, the focus
below is on the 60°S–60°N region, where the vast majority of theworld’s population resides.

Midlatitude regions. Ozone depletion is also observed at midlatitudes. Present-day (2017–2020 average) total ozone at midlatitudes of the Southern Hemisphere (SH) (35°S–60°S) is about
5% below the 1964–1980 average, whereas total ozone at midlatitudes of the Northern Hemisphere (NH) (35°N–60°N) is about
4% below the 1964–1980 average (see Figure Q12-1). Midlatitude ozone depletion has two contributing factors. First, ozonedepleted air over both polar regions is dispersed away from the
38 Section III: Stratospheric ozone depletion

poles during and after each winter/spring period, thereby reducing ozone at midlatitudes. Second, chemical destruction of ozone
occurs at midlatitudes, contributing to the observed reductionsin these regions. Ozone depletion at midlatitudes is much smaller than in polar regions (see Figure Q20-1) because the amountof reactive and reservoir halogen gases is lower and the dramatic seasonal increase of ClO, the most important reactive halogen
gas (see Figure Q7-3), does not occur at midlatitudes. The slightly larger depletion of ozone at SH midlatitudes, compared to NH
midlatitudes, is caused by the dispersion of air from the Antarctic
ozone hole.

Tropical region. Total ozone in the tropics (20°S–20°N latitude)
has been only weakly affected by chemical depletion. Present-day
total ozone in the tropics is about 1% below the 1964–1980 average. In the tropical lower stratosphere, air is transported from thelower atmosphere (troposphere) over about an 18-month period.
As a result, the fraction of ozone-depleting substances (ODSs)(halogen source gases) that is converted to reactive and reservoir
halogen gases (see Figure Q7-1) is small. With little reactive andreservoir halogen gases available, total ozone depletion in thisregion is also small. In addition, net ozone production occurs in
the tropics because of high average amounts of solar ultraviolet
radiation. In contrast, stratospheric air in polar regions has been in
the stratosphere for about 4 to 7 years, allowing time for significant conversion of ODSs to reactive and reservoir halogen gases(see Figure Q5-1). These systematic differences in the compositionof stratospheric air are a consequence of large-scale atmospherictransport: 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.
Ozone recovery. The Montreal Protocol, strengthened by its
amendments and adjustments, has successfully controlled the production and consumption of ODSs, which act to destroy the ozone
layer (see Q14). As a result, a quantity termed equivalent effective
stratospheric chlorine (EESC; the total chlorine and bromine abundances in the stratosphere) peaked in the late 1990s and is now
decreasing (see Q6 and Q15). An increase in upper stratosphericQ122022 Update | Twenty Questions
Global and Regional Total Ozone Changes
Northern Midlatitudes

Global60°S – 60°N

290
280
1970

1980

1990

Year

2000

2010

0
–2%
–4%
2020

Ozone (DU)

Ozone (DU)

Observed changes relative to the 1964 –1980 average
350
340

Ozone (DU)

–4%
–6%
1970

1980

1990

Year

2000

2010

2020

Tropics

270

20°S – 20°N

260
1970

1980

1990

Year

2000

2010

Ozone (DU)

330

60°S – 35°S

320
310
300

1970

1980

1990

Year

2000

2010

2%
0
–2%

2020

Southern Midlatitudes
Three-year running average
Uncertainty range
1964–1980 average

0
–2%

330
320

2%

35°N – 60°N

2%
0
–2%
–4%
–6%

2020

Figure Q12-1. Global total ozone changes. Ground-based and satellite observations show depletion of global total ozone beginningin the 1980s. The various panels compare the difference between annual averages of total ozone averaged over 60°S–60°N (“Global”),
35°N–60°N (“Northern Midlatitudes”), 20°S–20°N (“Tropics”), and 35°S–60°S (“Southern Midlatitudes”) to the amount of ozone thatwas present in these regions during the period 1964–1980. Seasonal effects have been removed from the observational data set. A 1964–
1980 baseline is used because large amounts of ozone depletion had not occurred during these years, as evident in the data record.
Global ozone decreased between 1980 and 1990. The depletion worsened for a few years after 1991 due to the effect of volcanic aerosols
from the eruption of Mount Pinatubo (see Q13). Since 2010, global ozone has been about 2 to 3% below the 1964–1980 average, ozone
in the northern and southern midlatitudes has been about 2 to 5% and 4 to 7% below the 1964–1980 average, respectively, and ozone in
the Tropics is about 0 to 2% below the 1964–1980 average.

ozone coincident with the decline in EESC is well documented,
constituting an important initial sign of the recovery of the ozone
layer. However, ozone in the upper stratosphere makes only a small
contribution to total ozone.
The depletion of total ozone over 1979 to 1995 varies as a functionof latitude, with the largest losses occurring at the highest latitudes
in both hemispheres (see Figure Q12-2). Over this time period
EESC nearly doubled. The pattern of ozone loss as a function of latitude is caused by two factors: 1) the tendency for larger levels of reactive and reservoir halogen gases to be present at higher latitudes
due to the pattern of large-scale atmospheric transport along withthe time required for significant conversion of ODSs to reactive andreservoir halogen gases as described above (see Q5 and Q7) and2) greater influence of the dispersal of severely ozone-depleted air
from the Antarctic (Q10) and Arctic (Q11) vortices at the end of theirrespective winters at the higher latitudes.

Over the time period 1996 to 2020, there has been a rise in totalozone of 0.4 % per decade averaged over 60°S–60°N (see Figures Q12-1 and Q12-2) that is consistent with the decline in theabundance of ODSs (see Q14 and Q15). The data indicate mostof this increase in total ozone occurs between 35°– 60° of theSouthern Hemisphere. For the tropics (20°S–20°N), trends intotal ozone over 1996 to 2020 are small and not statisticallysignificant. The abundance of EESC at midlatitudes declined byabout 15% from 1996 to 2020. The smaller decline of EESC overthis 24-year time period compared to the doubling of EESC overthe shorter 1979 to 1995 time period is due to the long lifetime
for atmospheric removal of ODSs (see Table Q6-1), combined withthe rapid growth in the emissions of ODSs in the 1970s and 1980s(see Figure Q0-1). The small increase in ozone shown in Figure
Q12-2 for 1996 to 2020 is consistent with the scientific community’s current understanding of the processes that control the abundance of atmospheric ozone.
Section III: Stratospheric ozone depletion 39

Q12 Twenty Questions | 2022 Update
Total Ozone Trends

Total column ozone trend(% per decade)

Trend in total column ozone due to halogen chemistry
1996 – 2020

2%
0
–2%
–4%

1979 – 1995

–6%
–8%

–60°

–30°

0°

Latitude30°

Identifying an ozone increase that is attributable to the observed
decrease in the amount of ODSs is challenging because halogen levels are not the only factor that determines the abundanceof stratospheric ozone. Total ozone declined over most of the
globe (60°S–60°N) during the 1980s and early 1990s, reachinga minimum in 1993 due to the combined effects of ODSs andthe eruption of Mount Pinatubo in 1991 (see Figures Q12-1 andQ13-1). This global ozone minimum was observed half a decade
before the EESC maximum was reached due to the strong global ozone response to enhanced amounts of stratospheric sulfate
aerosol after the volcanic eruption of Mount Pinatubo, which led toincreased ozone depletion for several years. Observed global ozone
increases in the mid-1990s were caused by the steady removal of
volcanic aerosol from the stratosphere by natural processes, which
occurred at the time EESC was approaching its maximum (see Q13).
Another factor complicating the identification of ozone recovery
in different regions of the atmosphere is the year-to-year variationof the stratospheric circulation. In most regions of the atmosphere,
these variations lead to ozone variability that is currently still larger
than the increases in ozone expected from the observed decrease
in EESC. Finally, increases in greenhouse gases (GHGs) such as

40 Section III: Stratospheric ozone depletion

Figure Q12-2. The change in total ozone due to ODSs as
a function of latitude, from 60°S to 60°N, for two time
periods, 1979–1995 and 1996–2020, expressed aspercent change per decade. The trends are based on a
statistical model analysis of measurements of total ozone
from numerous ground-based and satellite instruments.
The model analysis includes the effects on total ozone
from ozone-depleting substances (ODSs), stratosphericaerosols, solar ultraviolet (UV) radiation emitted by theSun, and stratospheric air motions. The figure isolates theresulting chemically-driven change in total ozone due
to trends in ODSs only for two time periods (solid lines),along with the total uncertainty (shaded regions).

60°

carbon dioxide (CO2 ), which warm the lower atmosphere, affectozone by decreasing stratospheric temperatures and by strengthening the stratospheric circulation. A colder stratosphere slows
down the rate of ozone loss reactions (outside of polar regions),
while a stronger circulation enhances the transport of ozone fromthe tropics to middle and high latitudes.
The impact on stratospheric ozone from accumulated emissionsof most ODSs will continue for several decades because of thelong atmospheric lifetime of these gases. Assuming compliancewith the Montreal Protocol, EESC will continue to decline over the
coming decades and will return to pre-1980 levels around 2066
(see Figure Q15-1). Increases in the abundance of most GHGs are
expected to accelerate this EESC-driven return of the global ozone
layer to pre-1980 levels. However, not all ozone-depleting gases
are halogen compounds. Nitrous oxide (N2O), a GHG with sources attributed to natural processes and various human activities, isprojected to increase in the future, which will result in additionalozone depletion (see Q20). Finally, as long as atmospheric abundances of ODSs remain elevated, the possibility of substantial
reductions in total ozone following major volcanic eruptions (seeQ13) will persist.

Q13

Do changes in the Sun and volcanic
eruptions affect the ozone layer?

Yes, factors such as changes in solar radiation and the formation of stratospheric aerosol particles after
explosive volcanic eruptions do influence the ozone layer. Global ozone abundances vary by 1–2% between
the maximum and minimum of the 11-year solar cycle. The abundance of global ozone decreased by about
2% for a few years after the June 1991 eruption of Mount Pinatubo, due to volcanic enhancement of stratospheric sulfate aerosols. However, neither factor can explain the observed decrease in global total ozone
or the severe ozone depletion observed in polar regions over the past half century. The primary influence on
long-term changes in total global ozone is the abundance of stratospheric halogens.

Changes in solar radiation and increases in stratospheric aerosols(small particles) from volcanic eruptions both affect the abundanceof stratospheric ozone. Global total ozone in the early 1990s decreased by about 5% when compared to the 1964–1980 average,
and is now about 2 to 3% below this value (see Q12). The longterm depletion of ozone is primarily attributed to increases in
halogen source gases, with additional depletion in the early 1990s
associated with the volcanic eruption of Mount Pinatubo. Equivalent effective stratospheric chlorine (EESC) (see definition in Q15)
is often used as a measure of the potential of reactive and reservoir
halogen gases to deplete ozone. Comparisons of the long-termchanges in solar radiation, stratospheric volcanic aerosol, and
EESC are useful in evaluating the contribution of these factors tolong-term changes in total ozone.

Total ozone and solar changes. The formation of stratosphericozone is initiated by ultraviolet (UV) radiation emitted by the Sun(see Q1). As a result, an increase in the Sun’s UV radiation output
increases the amount of ozone in Earth’s atmosphere. Since the1960s, ground-based and satellite instruments have recorded variations in the total energy emitted by the Sun, which is well correlated with changes in solar UV radiation. The Sun’s radiation output
varies over the well-documented 11-year solar cycle, as shown in
Figure Q13-1 in the quantity labeled incoming solar radiation. The
long-term solar record exhibits alternating maximum and minimum
values of total output, with maximum values separated by about 11
years. Global total ozone is relatively high compared to surrounding years during times of solar maxima and is relatively low duringsolar minima due to the sensitivity of ozone production to UV radiation, which increases during solar maxima. Analysis of measurements of ozone and incoming solar radiation shown in Figure
Q13-1 shows that ozone levels vary by 1 to 2% between the maximum and minimum of a typical solar cycle. In addition to this 11year variation, the total ozone record exhibits a long-term downward trend from the early 1980s to the early 2000s. If a decline inincoming solar radiation were the primary cause of the long-term
decline in global total ozone, then the solar radiation would exhibit a similar long-term decrease. Instead, incoming solar radiation
varies about a stable baseline over the modern instrument record.
This comparison demonstrates that the observed long-term de-cline in global total ozone does not result from changes in the Sun’s
UV radiation output.

Total ozone and past volcanoes. Explosive volcanic eruptions
inject sulfur gases directly into the stratosphere, causing new sulfate aerosol particles to be produced. These particles initially form
downwind of the volcano and then disperse over large regions,as air is transported by stratospheric winds. The largest impact
on global ozone usually takes place after explosive volcanic eruptions in the tropics, because the stratospheric circulation efficiently
spreads tropical volcanic plumes to both hemispheres. A principal
method of detecting the presence of volcanic particles in the stratosphere is to measure the transmission of solar radiation through thestratosphere to the ground, which is termed stratospheric aerosol
optical depth (SAOD). When large amounts of new particles form
over an extensive region of the stratosphere, SAOD increases andsolar transmission is measurably reduced. Figure Q13-1 shows thelong-term record of SAOD averaged over the entire stratosphere,
based on measurements from ground-based and satellite instruments. Large increases in SAOD (reductions in solar transmission)
are apparent after the explosive eruptions of Mount Agung (1963),
Volcán de Fuego (1974), El Chichón (1982), and Mount Pinatubo(1991), all of which occurred in the tropics. Reduced transmission of solar radiation persists for a few years after each of these
eruptions, until the stratospheric circulation and gravitational
settling bring the volcanic sulfate aerosol particles back to thetroposphere, where they are removed by precipitation.
Volcanic aerosol is primarily composed of sulfur compounds (sulfate). Chemical reactions on the surface of sulfate aerosol particles
destroy stratospheric ozone by increasing the abundance of chlorine monoxide (ClO), a highly reactive chlorine gas (see Q7). The
extent of ozone depletion depends on both the amount of sulfate
aerosol produced following the eruption and the value of EESC (seeQ15). Global ozone decreased for a few years following the eruptions of Mount Agung, Volcán de Fuego, El Chichón, and Mount
Pinatubo. The ozone reduction from the eruption of Mount Pinatubo stands out in the global ozone record because it occurred at atime when EESC was near its peak and the perturbation to stratospheric sulfate aerosol was especially large (see Figures Q12-1 andQ13-1). Analysis of ozone observations shows that global total
Section III: Stratospheric ozone depletion

41

Q13 Twenty Questions | 2022 Update
Effects on Global Ozone of
Human and Natural Factors

ozone declined by about 2% following the eruption of Mount
Pinatubo in June 1991, and that this effect persisted for 2 to 3 years
after the eruption. At times of relatively low EESC, such as the early 1960s, total ozone is not as sensitive to a volcanically induced
increase in stratospheric aerosol as during current times, when values of EESC are much higher than background levels.
If changes in the abundance of volcanic aerosol in the stratosphere
were the primary cause of the long-term decline in global totalozone, then the record of stratospheric aerosol optical depth
42 Section III: Stratospheric ozone depletion

Change from 1964 – 1980
average

Ozone(Dobson units)

Stratosphericaerosol opticaldepth(W/m2)

Total solar irradiance

EESC(parts per trillion)

Figure Q13-1. The effects on ozone
of EESC, solar changes, and volcanic eruptions. A comparison of thelong-term variation in EESC, total solar
radiation, and a measure of the abun2000 Stratospheric halogens
dance of stratospheric sulfate particles
with global total ozone provides abasis to evaluate the primary influenc1000
es on ozone over the past half century. The top panel shows the record
Natural sourcesof equivalent effective stratospheric0chlorine (EESC) for the midlatitude,
lower stratosphere (about 19 km altiIncoming solar radiationtude). EESC represents the potential
1362
for stratospheric ozone depletion byhalogens, which result primarily fromhuman activities (see Q15). The sec1361
ond panel shows the total energy ofincoming solar radiation, with peaks
and valleys defining the maxima and
1360
0.2
minima of the 11-year solar cycle. FolMajor volcanic eruptions:
Stratospheric aerosols
lowing the explosive volcanic erupMt. Agung
Fuego El Chichón Mt. Pinatubo(8°S)(14°N)(17°N)(15°N)
tions marked on the third panel, thenumber of sulfur-containing particles
0.1
Increasingin the stratosphere exhibits a dramatstratosphericic rise. These particles decrease theaerosols
transmission of solar radiation through0
the stratosphere, which is recordedby an increase in the quantity termed
2%
Global total ozone changes
295
stratospheric aerosol optical depth.
(60°S – 60°N)
The bottom panel shows differences in0
global total ozone, averaged over 60°S
290
to 60°N latitude, from the 1964–1980
average value. Global total ozone is 1–2%285
to 2% higher at solar maxima than solar minima due to enhanced formationof ozone by solar ultraviolet radiation
280
3-month averages–4%(see Q1). Reactions on sulfate particles
Change due to EESC
enhance the abundance of highly reactive chlorine compounds, increasing
1960
1970
2020
1980
1990
2000
2010
the depletion of stratospheric ozone
Yearafter major volcanic eruptions. The
maximum depletion of ozone occurs in mid-1993, after the eruption of Mount Pinatubo. The bottom panel also shows the change in global
total ozone attributed to EESC, found using an analysis that considers the effects on ozone of numerous natural and human-related factors.
The long-term changes in ozone are consistent with the variation of EESC: prior to the mid-1990s EESC steadily rose while global ozone declined, and since the late 1990s EESC has declined and global ozone has risen. This figure illustrates that the primary influence on changes
in global total ozone over the past half century is the abundance of stratospheric halogens.

(marker of volcanic sulfate particles) would exhibit a slow, gradual rise. Instead, stratospheric volcanic aerosol has been quite
low since 1995, a period of time over which global total ozone
has been about 2 to 3% below the 1964–1980 value. The data
record shown in Figure Q13-1 provides evidence that the long-term
decrease in global total ozone relative to the 1964–1980 average
does not result from changes in volcanic aerosol.

Total ozone and EESC. Values of EESC are derived from surface
observations of ozone-depleting substances (ODSs) and repre-

2022 Update | Twenty Questions
sent the potential for ozone depletion from halogens at particular
times and locations of the stratosphere (see Q15). The EESC record
for the midlatitude, lower stratosphere rose well above the natural background level in the 1980s, peaked in 1998, and in 2022
was 18% below the peak value. The bottom panel of Figure Q13-1
compares the observed long-term record of global total ozone(magenta line) to the variation in ozone attributed to the changes
in EESC (blue dashed line). This attribution curve is computed by astatistical model that considers the effects on ozone of EESC, stratospheric sulfur containing particles, variations in the total energy ofincoming solar radiation, as well as a few factors related to changes in stratospheric circulation. The observed record of global totalozone follows the same general tendencies of the EESC attribution
curve over the past half century, providing strong evidence
that changes in stratospheric halogens in response to humanactivities are the primary factor responsible for the long-term variation of ozone depletion. Further evidence linking ODSs and longterm variations in total column ozone is provided by the climatechemistry model simulations highlighted in Q20.

Halogen gases from volcanic eruptions. Explosive volcanic
eruptions have the potential to inject halogens directly into thestratosphere, in the form of gases such as hydrogen chloride (HCl),bromine monoxide (BrO), and iodine monoxide (IO). Although
HCl does not react directly with ozone, stratospheric injections of
HCl and other chlorine-containing gases following explosive volcanic eruptions can lead, through chemical reactions, to elevatedamounts of reactive chlorine monoxide (ClO) that destroys ozone(see Figure Q7-3). Eruption plumes also contain a considerable
amount of water vapor, which forms rainwater and ice in the rising
fresh plume. Rainwater and ice efficiently scavenge and remove
HCl while the plume is still in the lower atmosphere (troposphere).
Most of the HCl in the explosive plume of Mount Pinatubo did notenter the stratosphere because of this scavenging by precipitation. The amount of injected halogens depends on the chemicalcomposition of the magma, conditions of the eruption such as its

Q13

explosivity and the local meteorology. Recent analyses of several
historic, extremely large volcanic eruptions show the potential forquite large ozone loss from the stratospheric injection of halogens.
A volcanic eruption of this nature has not occurred during the time
period of the modern observational record.

Antarctic volcanoes. Volcanoes on the Antarctic continent are of
special interest due to their close proximity to the Antarctic ozone
hole. An explosive eruption could in principle inject volcanic aerosol or halogens directly into the stratosphere over Antarctica andcontribute to ozone depletion. To be a possible cause of the annually recurring ozone hole beginning in the early 1980s, explosive
eruptions of Antarctic volcanoes large enough to inject material
into the stratosphere would need to have occurred at least every
few years. This is not the case. Mount Erebus and Deception Island
are the only two currently active volcanoes in Antarctica. No explosive eruptions of these volcanoes, or any other Antarctic volcano, have occurred since 1980. Explosive volcanic eruptions in thelast three decades have not caused the Antarctic ozone hole and,as noted above, have not been sufficient to cause the long-term
depletion of global total ozone.
Total ozone and future volcanoes. The abundance of EESC willremain high for much of the 21st century due to the long atmospheric lifetime of ODSs (see Figure Q15-1). With its slow decline,
EESC will remain above the 1960 value throughout this century.
Consequently, throughout the rest of this century, increases in theabundance of stratospheric sulfate aerosol particles caused bylarge volcanic eruptions similar to Mount Pinatubo have the potential to reduce global total ozone values for a few years. The ozone
layer will be most vulnerable to such an eruption until midcentury,
since EESC is projected to return to the 1980 value around 2066.
Following an explosive eruption much larger than Mount Pinatubo,
or an eruption that injects halogens into the stratosphere, peakozone losses could both be greater than previously observed andpersist for longer periods of time.

Section III: Stratospheric ozone depletion 43

Q14

Are there controls on the productionof ozone-depleting substances?

Yes, the production and consumption of ozone-depleting substances (ODSs) are controlled under a 1987
international agreement known as the “Montreal Protocol on Substances that Deplete the Ozone Layer”
and its subsequent amendments and adjustments. The Protocol, now ratified by 198 parties, establishes
legally binding controls on the global production and consumption of ODSs. Production and consumptionof controlled ODSs by developed and developing nations will be almost completely phased out by 2030.

The Vienna Convention and the Montreal Protocol. In 1985,
a treaty called the Vienna Convention for the Protection of the
Ozone Layer was signed by 26 nations. The signing nations agreed
to take appropriate measures to protect the ozone layer fromhuman 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 ratification, entered into force in 1989. The Protocolhas been successful in establishing and enforcing 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 brominethat are emitted by human activities and are controlled under theMontreal Protocol are referred to as ozone-depleting substances(ODSs). National consumption of an ODS is defined as production plus imports of the controlled substance, minus exports ofthe substance. Inadvertent emissions of ODSs used as feedstock(raw material used to synthesize other substances through a process of chemical transformation) are not controlled by the Montreal
Protocol. Nations are, however, required to report imports, exports, and production of ODSs used as feedstocks. The Protocol
provisions are structured for developed countries to act first and fordeveloping countries to follow with financial assistance, technology transfer, and the sharing of knowledge for mitigation of emissions of ODSs as well as the destruction of ODSs within equipment
or products. In 2009, the Montreal Protocol became the firstmultilateral environmental agreement to achieve universal ratification.
Amendments and Adjustments. As the scientific basis ofozone 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 yearof the meeting. The timeline in Figure Q0-1 shows some of the
major decisions that have been adopted. These decisions brought
additional ODSs under control, accelerated the timing of existing
control measures, or prescribed phaseout dates for the production
and consumption of certain gases. The initial Protocol measures
were a 50% reduction in chlorofluorocarbon (CFC) production andconsumption by 1998 and a freeze on halon production and con44 Section IV: Controlling ozone-depleting substances

sumption. The 1990 London Amendment called for a phaseout ofthe production and consumption of the most damaging ODSs indeveloped nations by 2000 and in developing nations by 2010.
The 1992 Copenhagen Amendment accelerated the phaseout
date for CFCs, halons, carbon tetrachloride, and methyl chloroform to 1996 in developed nations and also initiated controls on
future production and consumption of hydrochlorofluorocarbons(HCFCs) in developed nations. Further controls on ODSs wereagreed upon in later meetings in Vienna (1995), Montreal (1997,2007), and Beijing (1999). The latest development is the 2016 Kigali
Amendment (see Q19), which expanded the Montreal Protocol
to control production and consumption of hydrofluorocarbons(HFCs) with high global warming potentials (GWPs) (see Table
Q6-1). As explained below, HFCs are greenhouses gases (GHGs)
which warm climate in the lower atmosphere and do not cause
ozone depletion.

Influence of the Montreal Protocol. Montreal Protocol phasedown schedules for each group of ODSs are based on several
factors including (1) the effectiveness in depleting stratosphericozone in comparison with other substances (see Ozone Depletion
Potential, ODP, in Q17), (2) the availability of suitable substitutes,
and (3) the potential impact of controls on developing nations.
The influence of Montreal Protocol provisions on stratospheric
ODS abundances can be demonstrated with long-term changes inequivalent effective stratospheric chlorine (EESC).
Calculations of EESC combine the amounts of chlorine andbromine present in air near Earth’s surface to form an estimate of
the potential for ozone destruction in a particular stratosphericregion on an annual basis (see definition in Q15). EESC values in the
coming decades will be influenced by (1) the slow natural removalof ODSs present in the atmosphere, (2) emissions from continued
production and use of ODSs, and (3) emissions from existing ODS
banks. The phrase ODS banks refers to long-term containment of
ODSs in various applications. Examples are CFCs in refrigeration
and air-conditioning equipment as well as insulating foams, andhalons in fire-extinguishing equipment. Annual emissions areprojected based on release from existing banks and any new production and consumption of ODSs. The long-term changes in EESC
at midlatitudes are shown in Figure Q14-1, based upon estimatesof future emissions of ODSs that were published in year 2007 forthe following cases:

2022 Update | Twenty Questions• No Protocol. In a world-avoided scenario without the Montreal Protocol, the production, use, and emissions of CFCs andother ODSs continue to increase unabated after 1987. This No
Protocol scenario is illustrated using an annual growth rate of
3% for the emissions of all ODSs. As a result, EESC increases nearly 10-fold by the mid-2050s compared with the 1980
value. Computer models of the atmosphere show that EESC
under the No Protocol scenario dramatically increases global
total ozone depletion between 1990 and 2020 relative to what
actually occurred, and increases ozone depletion much moreby midcentury. As a result, harmful UV-B radiation would haveroughly doubled at Earth’s surface by the middle of the 21st
century, causing damage to ecosystem health, and a global
rise in skin cancer and cataract cases (see Q16). Since ODSs are
powerful GHGs, the climate forcing from ODSs increases substantially without the Montreal Protocol (see Q18).

Q14• Montreal Protocol provisions. International compliance with
only the 1987 provisions of the Montreal Protocol and the later
1990 London Amendment substantially slowed the projected
growth of EESC relative to the world-avoided (No Protocol)
scenario. The projections show a decrease in future EESC values for the first time with the 1992 Copenhagen Amendments
and Adjustments. The provisions became more stringent withthe amendments and adjustments adopted in Beijing in 1999
and Montreal in 1997 and 2007. Now, with full compliance tothe Protocol, ODSs will ultimately be phased out, with someexemptions for critical uses (see Q15). Global EESC is slowly
decaying from its peak value in 1998 and is expected to reach1980 values around 2066. 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

Figure Q14-1. Effect of the Montreal Protocol. The Montreal Protocol protects the ozone
layer through control of the production andLong-term changes in equivalent effective
consumption of ozone-depleting substances
stratospheric chlorine (EESC)(ODSs) by the world’s nations. Equivalent ef6000
fective stratospheric chlorine (EESC) is a quantity that represents the abundance of halogens
available for ozone depletion in the stratosphere. Values of EESC are based upon either
5000
analysis of surface observations of ODSs (seeMontreal
Q15) or projections of future abundances of1987
ODSs. Projections of EESC for the midlatitude
No Protocol
lower stratosphere (about 19 km altitude) are
4000
shown separately for: no Protocol provisions;the provisions of the original 1987 Montreal
LondonProtocol and some of its subsequent amend1990
ments and adjustments as projected in in 2007;
3000
the current trajectory for EESC (red solid curve);
and for future EESC assuming zero emissionsof ODSs starting in 2023 (black dashed curve).
The city names and years indicate where andwhen changes to the original 1987 Protocol
2000
Current trajectory
provisions were agreed upon (see Figure Q0-1).
Without the Protocol, EESC values would have
Copenhagen
1992
increased significantly in the 21st century,
Montreal
2007
leading to large amounts of ozone depletion
1000
throughout the world, including over highly
Zero emissions
in 2023
populated regions. Only with the CopenhaNatural sources
gen (1992) and subsequent Amendments andAdjustments do projected EESC values show a
0
1950
2000
2050
2100
long-term decrease. The EESC values from obYear
servations shown in Figure Q15-1 lie above theMontreal 2007 projection due mainly to the useof larger emissions of ODSs from banks as well as increased emissions of ODSs from feedstocks used to manufacture various other chemicals than was considered in the original Montreal 2007 projection. Upward revision of the lifetime of CFC-11 and a few other ODSs since2007, as well as unreported emissions of CFC-11, make small contributions to the slower decline in EESC than projected in 2007. If future
emissions of ODSs from feedstocks, banks, and quarantine and pre-shipment applications of CH3Br currently allowed under the Montreal
Protocol could all be eliminated starting in 2023 (black dashed line), EESC would fall below the trajectory of full compliance following theMontreal 2007 Adjustment, as had been projected in 2007 (green line). The contributions from very short-lived gases (see Figure Q6-1),
which have been minor to this point, are not included in any of these EESC time series.

EESC (parts per trillion)

Effect of the Montreal Protocol(The unit “parts per trillion” is defined in the caption of Figure Q6-1.)

Section IV: Controlling ozone-depleting substances 45

Q14 Twenty Questions | 2022 Updatefrom methyl chloride and methyl bromide (see Q15). Between
1987 and 2022, the emissions of ODSs from human activities
decreased by almost 80%.
Figure Q14-1 also shows long-term changes in EESC for two additional cases from the 2022 Scientific Assessment of Ozone Depletion report:• Current trajectory. The current trajectory of EESC shown in
Figure Q14-1 is based on observed abundances of ODSs from2007 to 2021 and projected future abundances, assuming
compliance with the Montreal Protocol. The current trajectoryfor EESC lies above the Montreal 2007 curve for two reasons:increased emissions of ODSs from banks as well as from feedstocks used for production of other chemicals, both relative tothe assumptions of the original Montreal 2007 projection. The
upward revision of the lifetime of CFC-11 and a few other ODSs
since 2007, as well as unreported emissions of CFC-11, makesmall contributions to the slower decline of EESC than was projected in 2007.
• Zero emissions. The zero emissions scenario demonstrates
the reduction in EESC that occurs if emissions of all ODSs are
set to zero beginning in 2023. This assumption eliminates theemissions from new production, feedstocks, and banks. Significant differences from the current trajectory are evident in thefirst decades following 2023 because the phaseout of all ODS
production under the Protocol will not be complete in 2023
and continued bank emissions are substantial. In the zeroemissions scenario, EESC returns to the 1980 value about adecade earlier than currently projected (solid red and dashedblack lines, Figure Q14-1).

HCFC substitute gases. The Montreal Protocol provides for theuse of hydrochlorofluorocarbons (HCFCs) as transitional, shortterm 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,
fluorine, and carbon. Per amount emitted to the atmosphere,
HCFCs are 88 to 98% less effective than CFC-11 in depleting stratospheric ozone because their chemical removal occurs primarily
in the troposphere (see ODPs in Table Q6-1). The dominance of
tropospheric removal for HCFCs prevents most of the halogen
content this group of ODSs from reaching the stratosphere. In contrast, CFCs and some other ODSs release all their halogen content
in the stratosphere because they are chemically inert in the troposphere (see Q5).

46 Section IV: Controlling ozone-depleting substances

Under the provisions of the Montreal Protocol, developed anddeveloping countries may continue to produce and import HCFCs
over the rest of this decade before they are ultimately phased out.
In the 2007 Adjustment to the Protocol, the phaseout of HCFCs
was accelerated so that nearly all production ceased by 2020 for
developed countries and ceases by 2030 for developing countries, about a decade earlier than in previous provisions. In making
this decision, the Parties reduced the contribution of HCFC emissions to both long-term ozone depletion and future climate forcing(see Q17 and Q18).

HFC substitute gases. Hydrofluorocarbons (HFCs) are transitional substitute compounds for CFCs, HCFCs, and other ODSs. HFCs
contain hydrogen, fluorine, and carbon. HFCs do not contributeto ozone depletion because they contain no chlorine or bromine.
However, most HFCs and ODSs are also GHGs with long atmosphere lifetimes, so they contribute to human-induced climatechange (see Q18 and Q19). Under the auspices of the United Nations Framework Convention on Climate Change (UNFCCC), HFCs
are included in the basket of GHGs for which regular reporting of
national annual emissions are required. The Paris Agreement of the
UNFCCC is an international accord designed to reduce the emissions of GHGs in order to limit global warming to well below 2.0°C
relative to the start of the Industrial Era and pursue efforts to limit
global warming to 1.5°C warming. Future growth in the production and consumption of HFCs with high GWPs is now limited by
the 2016 Kigali Amendment to the Montreal Protocol (see Q19).
Very short-lived (VSL) chlorine source gases. VSL halogenated source gases, defined as compounds with atmospheric lifetimes typically shorter than 0.5 years, are primarily converted to reactive halogen gases in the lower atmosphere (troposphere). Atmospheric release of most very short-lived chlorine
source gases, such as dichloromethane (CH2Cl2, which is also
known as methylene chloride) and chloroform (CHCl3 ), results
primarily from human activities. This class of compounds is notcontrolled by the Montreal Protocol and therefore they are notincluded in the estimates of EESC shown in Figure Q14-1. The
atmospheric abundance of VSL chlorine source gases has increased substantially since the early 1990s and these gases presently contribute about 4% (130 ppt) to the total chlorine entering the stratosphere (see Figure Q6-1). Furthermore, the fractionof VSL gases that reach the stratosphere varies as a function of
the location where the gases are emitted into the atmosphere.
Should this class of compounds ever come under the control of theMontreal Protocol, future actions would be effective at lowering
atmospheric abundances rather quickly, because these compounds are removed from the atmosphere within a few years.

Q15

Has the Montreal Protocol beensuccessful in reducing ozone-depleting
substances in the atmosphere?

Yes, as a result of the Montreal Protocol, the overall abundance of ozone-depleting substances (ODSs) in theatmosphere has been decreasing for the past two decades. 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, HCFC-22 (CHF2Cl) and HCFC-141b (CH3CCl2F), will begin
to decrease this decade if compliance with the Protocol continues. However, it is only after midcentury thatthe abundance of ODSs is expected to fall to values that were present before the Antarctic ozone hole was
first observed in the early 1980s, due to the long atmospheric lifetime of these gases.

The Montreal Protocol and its amendments and adjustments havebeen very successful in reducing the atmospheric abundances ofozone-depleting substances (ODSs). ODSs are halogen sourcegases released by human activities whose production and consumption are now controlled by the Montreal Protocol (see Q14).
The success of the Montreal Protocol controls is documented by(1) observed changes and future projections of the atmosphericabundances of the principal ODSs and (2) the long-term decreasein equivalent effective stratospheric chlorine (EESC).

Individual ODS reductions. The reduction in the atmosphericabundance of an ODS in response to controls on production andconsumption depends principally on how rapidly this ODS is used
and released to the atmosphere after being produced, as well as
the atmospheric lifetime of the ODS (see Table Q6-1). For example, the abundances of ODSs with short lifetimes, such as methyl
chloroform, respond quickly to emission reductions. In contrast,
the abundances of ODSs with long lifetimes such as CFC-11 andCFC-12 respond slowly to emission reductions. Estimates of longterm changes in the atmospheric abundances of ODSs are based
upon: (1) their measured abundances in air trapped for years within accumulated snow in polar regions, (2) observed atmosphericabundances using ground-based measurements, (3) projections of
future abundances based on estimated future demand and compliance with Montreal Protocol provisions for the production andconsumption of ODSs, and (4) emissions from ODS banks. The
term bank refers to the total amounts of ODSs contained in existing
equipment, chemical stockpiles, foams, and other products thathave not yet been released to the atmosphere. The destruction of
ODSs in banks prevents the eventual release of these compounds
into the atmosphere. The long-term changes of the atmosphericabundances of individual ODSs and the natural chlorine andbromine source gases, methyl chloride (CH3Cl) and methyl
bromide (CH3Br), assuming compliance with the Montreal Protocol,
are shown in Figure Q15-1. Key aspects of families of ODSs shownin this figure are:
• CFCs. Chlorofluorocarbons (CFCs) include some of the mostdestructive chlorine-containing ODSs. CFC-11 and CFC-12,

with large Ozone Depletion Potentials (ODPs) of 1 and 0.75,
are the most abundant CFCs in the atmosphere owing to large
historical emissions and long atmospheric lifetimes of about 50
and 100 years, respectively (see Table Q6-1). Under the Montreal Protocol, allowed production and consumption of CFCs
ended in January 1996 for developed countries and in January2010 for developing countries. As a consequence, the atmospheric abundances of CFC-11 and CFC-113 peaked in 1994
and 1996, respectively, and have been declining for more than
two decades. In contrast, the abundance of CFC-12 peaked in
2002 and has been slowly decreasing, owing to its longer lifetime (about 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, except for the use of certain CFCs
(mainly CFC-113 and CFC-114) as feedstocks and some limited
exempted uses, along with some continuing emissions frombanks, CFC abundances are projected to decline steadily
throughout this century. From 2012 to 2018, the annual decline in CFC-11 slowed measurably compared to the expected
decline, due to unreported production outside the provisionsof the Montreal Protocol. In 2019 and 2020, global emissionsof CFC-11 declined in a manner indicative of the elimination ofmost of these unreported emissions.
• Halons. Halons are currently the most important brominecontaining ODSs. Halon-1211 and halon-1301 account fora significant fraction of bromine from all ODSs (see FigureQ6-1). Under the Montreal Protocol, production and consumption of halons for controlled uses ended in January 1994 fordeveloped countries and in January 2010 for developing
countries, with some essential use exemptions for both developed and developing countries. Atmospheric abundances of
halon-1211 show significant decreases since peak concentrations were measured in the mid-2000s. Halon-2402 abundances have been decreasing slowly for the past two decades while
those of halon-1301 have nearly peaked and are expected to
decline in coming decades. The slow decline in the emission of
halon-1301 is likely due to substantial banks in fire-extinguishing
Section IV: Controlling ozone-depleting substances 47

Q15 Twenty Questions | 2022 Update
Past and Projected Atmospheric Abundances of Halogen Source Gases
600
CFCs
Halons
HCFCs
CCl4
CH3CCl3

EESC (parts per trillion)

Mid-latitudes

1500
1980 level

1000

500
CH3Br (natural + human sources)
CH3Cl (natural sources)

Atmospheric abundance (parts per trillion)

0140

Methyl chloroform(CH3CCl3)

Historical estimates
Direct air
observations
Future projections

120
100

500

300

Carbon tetrachloride(CCl4)

60

CFC-11
HCFC-22200
100
CFC-113
0
4

Halon-1211
Halon-1301

3
2
1

80

CFC-12400

Atmospheric abundance (parts per trillion)

2000

Equivalent effective stratospheric chlorine

Halon-2402

0Methyl chloride (CH3Cl)

500

40
HCFC-141b

8

400

20

Methyl bromide (CH3Br)

HCFC-142b01960 1980 2000 2020 2040 2060 2080 2100

Year

300

10

6
1960 1980 2000 2020 2040 2060 2080 2100

Year

Figure Q15-1. Halogen source gas changes and contributions to EESC. The surface abundances of individual gases shown here wereobtained using a combination of direct atmospheric measurements, estimates of historical abundances, and future projections of abundances assuming compliance with the Montreal Protocol. The gases are all ODSs except for methyl chloride, which is primarily emitted
from natural processes. Therefore methyl chloride is not controlled by the Montreal Protocol and its abundance is projected to be constantin the future. The past increases of the principal CFCs shown in the figure, along with those of carbon tetrachloride, methyl chloroform,
halon-1211, and halon-2402, have slowed and reversed in the last three decades. The abundances of most HCFCs, which are used astransitional substances to replace CFCs, will likely peak during this decade before production and consumption are completely phased
out. The abundance of halon-1301 has nearly peaked and is expected to slowly decline in coming decades because of continued small
releases and a long atmospheric lifetime. Future decreases in methyl bromide are expected to be modest, since most of the current sourceis from natural processes and the industrial production of methyl bromide is now much smaller than occurred in the 1990s. Values of EESC
shown in the upper left panel are for the midlatitude lower stratosphere (about 19 km altitude), based on observations and projections (seeQ14-1). The colored shaded regions depict contributions to EESC from CH3Cl, CH3Br, CH3CCl3, CCl4, HCFCs, halons, and CFCs and the
top of the magenta shaded region is the sum of all contributions, or EESC. The value of EESC rose rapidly during the latter half of the priorcentury, peaking in the late 1990s. Since the late 1990s, the trend in EESC has reversed due to the effectiveness of the Montreal Protocol
in reducing the production and consumption of ODSs. In the midlatitude stratosphere EESC is projected to return to its 1980 value around2066. In polar regions EESC is projected to return to 1980 values about two decades later. International compliance with the provisionsof the Montreal Protocol is required to ensure that EESC will continue to decrease as projected (see Q14). Over 2012 to 2018, the decline
in CFC-11 slowed measurably compared to the expected decline, due to unreported production. In 2019 and 2020, CFC-11 declined in amanner indicative of the elimination of most of this unreported production. The contributions from very short-lived gases (see Figure Q6-1)
are not included in this EESC time series.
(The unit “parts per trillion” is defined in the caption of Figure Q6-1.)

48 Section IV: Controlling ozone-depleting substances

2022 Update | Twenty Questions
and other equipment that gradually release this compoundto the atmosphere years after production. The abundance of
halon-1301 is expected to remain high well into the 21st century
because of its long lifetime (72 years) and continued release.
Atmospheric emissions of halons continue to occur due to theuse of halon-containing equipment in various fire-fighting applications.
• Carbon tetrachloride. Production and consumption ofcarbon tetrachloride (CCl4 ) for controlled uses in developed
countries was phased out in 1996 and that in developing countries in 2010, with some essential use exemptions. As a result,atmospheric abundances of carbon tetrachloride have been
decreasing for two decades. The decline is considerably lessrapid than expected, suggesting that actual emissions are larger than the emissions derived from the reported consumption.
Carbon tetrachloride that is used as raw material (feedstock) tomake other chemicals is exempted when calculating the controlled levels of production and consumption under the Montreal Protocol, and some residual emissions do occur. Current
understanding of global sources suggests emissions of carbon
tetrachloride are presently dominated by inadvertent production and subsequent release during the chemical manufacturing processes of other compounds, as well as release fromlandfills and contaminated soils.
• Methyl chloroform. The largest reduction to date in the abundance of an ODS (99% from its peak value) has been observed
for methyl chloroform (CH3CCl3 ). Production and consumptionof methyl chloroform in developed countries ended in January
1996 and that in developing countries ended in January 2015,
with limited essential use exemptions. Atmospheric abundances responded rapidly to the reduced emissions starting in the
mid-1990s because methyl chloroform has a short atmospheric
lifetime of about 5 years. Methyl chloroform was used mainly asa solvent and had typically been emitted soon after production.
This compound is now approaching complete removal fromthe atmosphere due to the success of the Montreal Protocol.

Q15chloride is not controlled under the Montreal Protocol. The
abundance of CH3Cl in the atmosphere has remained fairly constant throughout the last 40 years (see Figure Q15-1). Current
sources of methyl chloride from human activities are thought tobe small relative to its natural source, and to be dominated by
the combustion of coal and chemical manufacturing.
In contrast, methyl bromide is controlled under the Montreal
Protocol. Methyl bromide was primarily used as a fumigant.
Nearly all developed country production and consumption of
methyl bromide ended in January 2005 and that in developing
countries ended in January 2015. The Protocol currently provides limited exemptions for methyl bromide production anduse as a fumigant in agriculture as well as for quarantine andpre-shipment applications. Atmospheric abundances of methyl bromide declined rapidly in response to the reduced emissions starting in 1999, because its atmospheric lifetime is lessthan 1 year (see Figure Q15-1). Future projections show only
small changes in methyl bromide abundances based on theassumptions of unchanged contributions from natural sources
and small continued critical use exemptions. An important uncertainty in these projections is the future amount that will be
produced and emitted under Montreal Protocol critical use,quarantine, and pre-shipment exemptions.

Equivalent effective stratospheric chlorine (EESC). Important measures of the success of the Montreal Protocol are thepast and projected changes in the values of equivalent effective
stratospheric chlorine, which was introduced in Figures Q13-1
and Q14-1. EESC is one measure of the potential for ozone depletion in the stratosphere that can be calculated from atmospheric surface abundances of ODSs and natural chlorine andbromine gases. The calculation considers CFCs, HCFCs, methyl
chloroform, carbon tetrachloride, halons, as well as methyl
chloride and methyl bromide. For both past and future EESC values, the required atmospheric abundances are derived from measurements, historical estimates, or future projections based oncompliance with the provisions of the Montreal Protocol.

• HCFC substitute gases. The Montreal Protocol allows for theuse of hydrochlorofluorocarbons (HCFCs) as short-term, transitional substitutes for CFCs and in other specific applications.
As a result, the atmospheric abundances of HCFC-22, HCFC141b, and HCFC-142b continue to grow in response to continued production, mainly in the developing world. HCFCs posea lesser threat to the ozone layer than CFCs, because HCFCs
have lower ODP values (less than about 0.1; see Table Q6-1).
The 2007 Montreal Adjustment to the Protocol accelerated thephaseout of HCFCs by a decade for both developed countries(2020) and developing countries (2030) (see Q14). Future projections indicate that the principal HCFCs will all reach peakvalues between 2023 and 2030, and steadily decrease thereafter. The response of atmospheric abundances to decreasing
emissions (due to gradual releases from existing banks such as
insulating foams) will be relatively rapid because of the short atmospheric lifetimes of HCFCs (less than 17 years).

EESC is derived from the amount of chlorine and bromine 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. Althoughchlorine is much more abundant in the stratosphere than bromine(about 150-fold) (see Figure Q6-1), bromine atoms are about 60
times more efficient than chlorine atoms in chemically destroying
ozone in the lower stratosphere. The term effective indicates that
only the estimated fractions of ODSs that have been converted
to reactive and reservoir halogen gases, for a particular region of
the stratosphere at a specified time, are included in the computed value of EESC value (see Q5 and Q7). Long-term changes of
EESC generally depend on the altitude and latitude region in thestratosphere under consideration. The EESC curve shown in Figure
Q15-1 is for the midlatitude, lower stratosphere (about 19 km altitude).

• Methyl chloride and methyl bromide. Both methyl chloride(CH3Cl) and methyl bromide (CH3Br) are distinct among halogen source gases because substantial fractions of their emissions are associated with natural processes (see Q6). Methyl

ry up until the 1990s, EESC values steadily increased (see Figure
Q15-1), causing global ozone depletion. As a result of the Montreal Protocol regulations, the long-term increase in EESC slowed,

Long-term changes in EESC. In the latter half of the 20th centu-

Section IV: Controlling ozone-depleting substances 49

Q15 Twenty Questions | 2022 Updatevalues reached a peak in 1999, and EESC then began to decrease. By 2022, EESC at midlatitudes had declined by about
18% from the peak value. The initial decrease came primarily
from the substantial, rapid reductions in the atmospheric abundance of methyl chloroform, which has a lifetime of only 5 years.
The decrease is continuing with declining abundances of CFCs,

50 Section IV: Controlling ozone-depleting substancescarbon tetrachloride, and halon-1211. Decreases depend on natural processes that gradually decompose and remove halogencontaining gases from the global atmosphere (see Q5). Reductionof EESC to 1980 values or lower will require four more decades because the most abundant ODS gases now in the atmosphere have
lifetimes ranging from 10 to 100 years (see Table Q6-1).

Q16

Does depletion of the ozone layer
increase ground-level ultraviolet
radiation?

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 increasedin large geographic regions in response to ozone depletion.

Depletion of stratospheric ozone leads to an increase in solar ultraviolet radiation at Earth’s surface. The increase occurs primarily inthe ultraviolet-B (UV-B) component of the Sun’s radiation. UV-B isdefined 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 changes in total ozone (see Q3).
Exposure to UV-B radiation can harm humans, other life forms, andmaterials (see Q2). Most of the effects of sunlight on the humanbody are caused by UV-B radiation. A principal effect is sunburn,which first appears as reddening of the skin, also called erythema.
Excess exposure to UV-B radiation can lead to skin cancer. Erythemal radiation is regularly reported to the public in many countries
in the form of the UV Index (UVI), which is proportional to the erythemally weighted UV radiation at Earth’s surface. The UVI ranges
from zero at night to more than 20 at noon for high elevations in
the tropics.

Surface UV-B radiation. The amount of UV-B radiation reaching
Earth’s surface at a particular location depends in large part on total
column ozone (see Q3) in the atmosphere at that location. Ozone
molecules in the stratosphere and in the troposphere absorb UV-B
radiation, thereby significantly reducing the amount that reaches Earth’s surface (see Q2). If conditions occur that reduce theabundance of ozone molecules somewhere in the troposphere or
stratosphere, total ozone is reduced and the amount of UV-B radiation reaching Earth’s surface is increased proportionately.
Additional causes of UV changes. The actual amount of UV-B
radiation reaching Earth’s surface at a specific location and timedepends on a number of other factors, in addition to the amountof total ozone. The primary additional factor is the elevation of the
Sun in the sky, which changes at any location with daily and seasonal cycles. Other factors include the altitude of the location, local
cloudiness, the amount of ice or snow cover, and the amounts ofatmospheric 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.
The seasonal change in the Earth–Sun distance is also a significant
factor affecting the amount of UV-B radiation reaching the surface.
Measurements indicate that both increases and decreases in UV
radiation at certain locations have resulted from variations in one

or more of these factors. Estimating the impact of changes in thesefactors is complex. For example, an increase in cloud cover usually
results in a reduction of UV radiation below the clouds and at thesame time could increase UV radiation at a location in the mountains above the clouds. Conversely, if clouds don’t sufficiently
block the direct beam from the Sun, reflections from cloud particles can result in an increase in the amount of UV-B reaching thesurface.

Biologically-weighted UV and the UV Index (UVI). The effectof ozone on the amount of biologically relevant UV that arrives at the
Earth’s surface is governed by the wavelength dependence of thebiological action involved, which typically — as in the case for skindamage — increases towards shorter (higher energy) UV-B wavelengths. Effects are commonly reported in terms of the UVI. All else
being equal, the UVI rises as the abundance of total ozone declines.
This inverse relationship between total ozone and the UVI measured
at surface locations at stations throughout the world is shown in
Figure Q16-1. The observations show that a 50% decline in totalozone is associated with a factor of two increase in the UVI.
The UVI is used internationally to increase public awareness about
the detrimental effects of UV on human health and to guide theneed for personal protective measures. Largest values of the UVI
occur in the tropics, where the midday Sun has the highest elevation throughout the year and total ozone values tend to be low(see Figure Q3-1). At all latitudes, the UVI is larger in mountainous
areas (due to less overhead air to scatter or absorb the radiation)and over snow- or ice-covered regions (due to increased surface
reflectivity). Values of the UVI greater than 10 are considered‘extreme’: under this circumstance, damage to sensitive fair skincan occur within 15 minutes of exposure.
The expected variation of the UVI with respect to total ozone isshown by the line marked “Model” in Figure Q16-1, which is inexcellent agreement with the observations. The sensitivity of UV-B
to ozone change is similar to that for the UVI, whereas other biological weightings exhibit different sensitivities. For example, thesensitivity for damage to DNA as a function of total ozone is about
twice as large as that for the UVI and UV-B, while the sensitivity forthe production of Vitamin D in the skin is intermediate between
those for DNA-damage and erythema.

Long-term surface UV changes. Changes in surface UV since
Section V: Implications of ozone depletion and the Montreal Protocol

51

Q16 Twenty Questions | 2022 Update
The maximum daily UVI varies dramatically with location andseason due largely to its strong dependence on solar elevation
angle. Ground-based measurements since the early 1990s
from the Palmer research station on the Antarctic Peninsula, thecity of San Diego in southern California, and at Point Barrownear Utqiaġvik in northern Alaska enable a direct comparison
between the UVI at polar and lower latitudes. The data showthat the large geographical differences in the historical UVI
have been dramatically affected by ozone depletion, especially at the Antarctic site, where the increases exceed historical
geographic differences (see Figure Q16-2).

Relationship Between Changes in
Total Column Ozone and UV Index

UV Index Change

x 2.5

47° N
44° N
41° N
20° N
45° S
90° S

x2

For San Diego and Point Barrow, the daily maximum UVI is largest during summer, when the midday Sun is closest to being
overhead. For the Antarctic site, the daily maximum UVI now
peaks in spring, the season of lowest total ozone due to theozone hole (see Q10). The daily maximum UVI decreases significantly after mid-December due to the seasonal recovery
of total ozone, following the break-up of the ozone hole (see
Q10).

Model

x 1.5

x1

–60

–50

–40

–30

–20

–10

0

% Ozone Change
Figure Q16-1. Relationship between changes in total column ozone
and the UVI measured at six sites distributed throughout the world.
All data are for cloud-free conditions and are acquired at the same solar elevation for each site during the 1990s. The observed variations
in ozone are large. For example, at the South Pole the ozone amountis sometimes reduced by more than 50% from its maximum. The corresponding UVI for that lowest ozone amount is a factor of 2.5 larger
than for the highest ozone amount. At lower latitudes, ozone changes are smaller, but increases in the UVI can still exceed a factor of 1.5.
UVI data for all sites show a similar ozone dependence, which agrees
well with the dependence predicted using a model of radiative transfer through the atmosphere, as shown by the smooth curve labeled“Model”.

the onset of ozone depletion are not well documented. Estimatedchanges derived from satellite measurements of ozone are incomplete because the backscattered radiation used by the satellites
does not fully penetrate the lowermost portion of the atmosphere.
Consequently, effects on UV radiation at Earth’s surface due to interactions of sunlight with atmospheric aerosols and clouds must be
estimated using computer models. Few ground-based measurements of UV radiation suitable for long term trend analysis were available prior to the early 1990s, during the period of most pronounced
ozone depletion (see Q12). The start of the UV instrument record isfurther complicated by the volcanic eruption of Mount Pinatubo in
1991 (see Q13). Stratospheric aerosols from this eruption contributed
to widespread ozone loss and also directly blocked solar radiation,
including UV radiation, for more than a year.
52 Section V: Implications of ozone depletion and the Montreal Protocol

Prior to the development of the Antarctic ozone hole, the UVI
was always much higher at San Diego (32°N) than at Palmer
Station (64°S). Measurements at Palmer Station demonstrate
the dramatic effect of Antarctic ozone depletion. There, estimates of the UVI for the years 1970 to 1976, a period before
the appearance of the ozone hole, are compared with measurements for the period 1990–2020 when Antarctic ozone
depletion increased the UVI throughout spring and into summer (orange shading). The development of the ozone hole led
to large enhancements of the UVI that persist for many months,
with the greatest increases occurring during spring.
The maximum UVI at Palmer Station is now larger by about afactor of 2.5 compared with the pre-ozone-hole period. The
highest UVIs observed in spring now exceed those measured
in spring and early summer in San Diego, despite San Diego’smuch lower latitude. The large levels of UV radiation reaching
the surface due to the Antarctic ozone hole have had an adverse effect on the microscopic plants and animals at the baseof the food chain in the high-latitude marine environment.
At San Diego, measurements of the UVI since 1992 and UVI
values reconstructed based on pre-ozone depletion values arealmost indistinguishable. This small change is consistent withthe small variation in total ozone observed at subtropical latitudes (see Q12) and with the finding that the maximum daily
UVI has remained essentially constant at these latitudes overabout the past 20 years. At the Arctic site near Point Barrow,
the UVI has increased by approximately 20% since the 1970s.

Modeling UV changes. Despite the short data record,
trends in the UVI measured at unpolluted sites since the early
1990s agree well with trends calculated from changes in totalozone. These measurements show that the change in the UVIfrom 1996 (around the time that stratospheric chlorine peaked, see Figure Q15-1) to 2020 have been small.
Model simulations of ozone in a world without the Montreal
Protocol show that large increases in the UVI due to ozone depletion occur during the period 1996 to 2020. In these simulations, at mid-southern latitudes the UVI in summer increases by2022 Update | Twenty Questions

Q16

Seasonal Changes in the UV Index
Maximum UV index on a given day of the year
Northern Hemisphere month
J

F

M

A

M

J

J

A

S

O

N

D

16

Maximum UV Index

14
12
San Diego,
California (32°N)

10Palmer,Antarctica(64°S)

8
6
4
2
0

Point Barrow, Alaska (71°N)

J

A

S

O

N

D

J

F

M

A

M

Southern Hemisphere monthPre-ozone hole

Since early 1990s

1970 – 19761970 – 19761970 – 1976

1990 – 20201992 – 2008
1991 – 2016

(calculated)

Palmer, Antarctica (64°S)
San Diego, California (32°N)
Point Barrow, Alaska (71°N)(measured)

about 20%, and in springtime the UVI in Antarctica doubles without the Montreal Protocol. These illustrative estimates serve as a
powerful testament to the benefit of the Montreal Protocol’s role inprotecting human health and the environment.

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 radiation exposure formany people. Increased UV exposure has adverse health effects,
primarily associated with eye and skin disorders. UV radiation is arecognized risk factor for eye cataracts. For the skin, the most common threat is skin cancer. Over the past decades, the incidence ofseveral types of skin tumors has risen significantly among peopleof all skin types. On the other hand, an important human healthbenefit of UV-B radiation exposure is the production of vitamin D,
which plays a significant role in bone metabolism and the immune
system. Human exposure to solar UV-B radiation requires a carefulbalance to maintain adequate levels of vitamin D, while minimizing
the risks of skin and eye disorders.

3

J

Figure Q16-2. Long-term changes in the UV
Index (UVI), which is a measure of the erythemal radiation that reaches the surface. The
figure shows the daily maximum UVI at threelocations: San Diego, California; Point Barrownear Utqiaġvik in northern Alaska; and Palmer
Station in Antarctica. The daily maximum UVI
is shown for two time periods: a more recent
period corresponding to the availability ofmeasurements starting in the early 1990s (solid lines, as indicated) and a pre-ozone holetime period of 1970–1976, computed using
simulated total ozone values at the three locations for this time period. The shaded region
for each measurement site represents the difference between the pre-ozone hole UVI andthe observed UVI, starting with early 1990s
data. The UVI observations from Palmer, Antarctica demonstrate the importance of ozone
depletion on the amount of erythemal radiation reaching the surface. For 1990–2020,
springtime values of the daily maximum UVI at
Palmer equaled or exceeded those measured
in spring at San Diego, which is located at amuch lower latitude and, as a consequence,experienced much larger values of the daily
maximum UVI than Palmer prior to the development of the ozone hole. The three thin vertical dotted lines show the dates of the springequinox, the summer solstice, and the autumn
equinox.

Skin cancer in humans typically occurs long after exposure to UV
radiation that causes sunburn. Even under the current provisions of
the Montreal Protocol and its amendments and adjustments, projections of additional skin cancer cases associated with ozone depletion are largest in the first half of the 21st century. This projection
represents a significant global health issue. Since recovery to 1980
values of total ozone averaged over 60°S to 60°N is projected tooccur around the middle of this century (see Figure Q20-1), ozone
depletion will continue to contribute to adverse human health effects over the coming decades.
In addition to detrimental effects on human health, increases in UV
radiation reaching the surface also impact air quality, aquatic andterrestrial plants and ecosystems, biogeochemical cycling, andoutdoor materials. The impacts of UV radiation are discussed ingreater detail in reports by the Environmental Effects Assessment
Panel (EEAP) of the Montreal Protocol on Substances that Deplete
the Ozone Layer3.

https://ozone.unep.org/science/assessment/eeap

Section V: Implications of ozone depletion and the Montreal Protocol 53

Q17

Is depletion of the ozone layer the
principal cause of global climate change?

No, ozone depletion is not the principal cause of global climate change. Ozone depletion and global
climate change are linked because both ozone-depleting substances and their substitutes are greenhouse gases. The best estimate is that stratospheric ozone depletion has led to a small amount of surface
cooling. Conversely, increases in tropospheric ozone and other greenhouse gases lead to surface warming.
The effect on global climate change from stratospheric ozone depletion is small compared to the warming from the greenhouse gases responsible for observed global climate change. Since the early 1980s, theAntarctic ozone hole has contributed to changes in Southern Hemisphere surface climate through effects
on the atmospheric circulation.

While stratospheric ozone depletion is not the principal cause ofclimate 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 (or halogen source gases) andtheir substitutes; and other leading greenhouse gases.

Greenhouse gases and the radiative forcing of climate.
The warming of Earth’s surface and troposphere by the Sun is enhanced by the presence of greenhouse gases (GHGs). The natural abundances of GHGs in Earth’s atmosphere absorb outgoing
infrared radiation, trapping heat in the atmosphere and warming
the surface. The most important natural GHG is water vapor. Without this natural greenhouse effect, Earth’s surface would be muchcolder than current conditions. Human activities have led to significant increases in the atmospheric abundances of a number of longlived and short-lived GHGs since 1750, the start of the Industrial
Era, leading to warming of Earth’s surface and associated climatechanges. This group includes carbon dioxide (CO2 ), methane(CH4 ), nitrous oxide (N2O), tropospheric ozone, and halocarbons.
Ozone-depleting substances (ODSs) and their substitutes make upa large fraction of the halocarbons in today’s atmosphere. Increases in the abundances of these gases from human activities cause
more outgoing infrared radiation to be absorbed and reemitted
back to the surface, further warming the atmosphere and surface.
This change in Earth’s energy 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 expressedusing units of watts per square meter (W/m2). The potential forclimate change rises as this radiative forcing increases.
A summary of radiative forcings of climate in 2019 resulting fromthe increases in the principal long-lived and short-lived GHGs
from human activities since 1750 is shown in Figure Q17-1. Positive forcings lead to warming and negative forcings lead to cooling of Earth’s surface. Climate forcings also lead to other changes,
for example reductions in glacier and sea-ice extent, variations inprecipitation patterns, and more extreme weather events. International climate assessments conclude that much of the observed
54 Section V: Implications of ozone depletion and the Montreal Protocol

surface warming and changes in other climate parameters overthe last several decades are due to increases in the atmosphericabundances of CO2 and other GHGs, which result from a variety ofhuman activities.

Carbon dioxide, methane, and nitrous oxide. All three ofthese GHGs have both human and natural sources. The accumulation of CO2 since 1750 represents the largest climate forcing
caused by 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 CO2 exceeded 416 parts per million(ppm) in 2021, which is about 50% larger than the abundance of
CO2 present in 1750. Carbon dioxide is considered a long-lived
gas, since a significant fraction remains in the atmosphere 100 to1000 years after emission.
Methane is a short-lived climate gas (atmospheric lifetime of about
12 years). Sources related to human activities include livestock,fossil fuel extraction and use, rice agriculture, and landfills. Natural
sources include wetlands, termites, and oceans. The global mean
atmospheric abundance of CH4 has more than doubled since
1750.
Nitrous oxide is a long-lived climate gas (atmospheric lifetime ofabout 109 years). The largest source related to human activities isagriculture, especially the use of fertilizer. Microbial processes insoils that are part of natural biogeochemical cycles represent thelargest natural source. In the stratosphere, nitrous oxide is the principal source of reactive nitrogen species that participate in ozone
destruction cycles (see Q8). The global mean atmospheric abundance of nitrous oxide has increased by about 22% since 1750.

Halocarbons. Halocarbons in the atmosphere contribute to bothozone depletion and climate change. The halocarbons considered in Figures Q17-1 and Q17-2 are gases containing chlorine,
bromine, or fluorine atoms that are either controlled under theMontreal Protocol or are GHGs that fall under the auspices of
the United Nations Framework Convention on Climate Change(UNFCCC). Historically, ODSs were the only halocarbons controlled under the Montreal Protocol. In 2016, the Kigali Amend-

2022 Update | Twenty Questions

Q17

Greenhouse Gases and Climate Change
Climate forcing due to changes in greenhouse gases caused byhuman activities between 1750 and 2019
Carbon dioxide (CO2)Methane (CH4)
Halocarbons
Nitrous oxide (N2O)
Tropospheric ozone (O3)Stratospheric ozone (O3)
–1.5

–1.0

–0.5

0

0.5

1.0

1.5

2.0

2.5

2

Radiative forcing (W/m )Surface coolingSurface warming

Figure Q17-1. Radiative forcing of greenhouse gases and ozone depletion. Human activities since the start of the Industrial Era (around1750) have caused increases in the atmospheric abundance of greenhouse gases (GHGs). Rising levels of GHGs lead to an increase in theradiative forcing (RF) of climate by trapping infrared radiation released by Earth’s surface. Here, values of RF are for the time period from1750 to 2019 and are expressed in units of watts per square meter (W/m2); black whiskers on each bar show uncertainties. Positive valuesof RF (shown in red) contribute to global warming and negative values (shown in blue) contribute to climate cooling. The largest positive
RFs are due to carbon dioxide (CO2 ), methane (CH4 ), tropospheric ozone (O3 ), halocarbons, and nitrous oxide (N2O). Halocarbons include
all ozone-depleting substances, hydrofluorocarbons, and a few other gases (see Figure Q17-2). The RF due to ozone is shown as the separate response to changes in ozone within two layers of the atmosphere: the troposphere and stratosphere. Tropospheric ozone increases
result from the emission of air pollutants and lead to a positive RF whereas stratospheric ozone depletion leads to a small RF that could be
negative or positive.

ment to the Montreal Protocol established controls on the future production and consumption of certain hydrofluorocarbon(HFC) gases. Perfluorocarbons (PFCs) and sulfur hexafluoride(SF6 ) are in the UNFCCC group of GHGs that now fall under theParis Agreement. Perfluorocarbons are compounds that contain only carbon and fluorine atoms, such as carbon tetrafluoride(CF4 ) and perfluoroethane (C2F6 ). Technically, SF6 is not a halocarbon since it lacks carbon. However, the environmental effectsof SF6 are commonly examined with those of halocarbon gases
since all of these compounds contain at least one halogen atom.

The large contribution of the CFCs to radiative forcing has been
gradually decreasing following the decline in their atmosphericabundance and is expected to further decrease (see Figure Q15-1).
Based on their long lifetimes, CFCs will still make a significant contribution, and most likely the largest contribution from ODSs, tothe radiative forcing by halocarbons at the end of this century. Even
with adherence to the provisions of the Kigali Amendment to theMontreal Protocol, the radiative forcing from HFCs is projected toincrease for another two to three decades before starting to slowly
decline (see Figure Q19-2).

In 2019, the halocarbon contribution to the radiative forcing (RF)of climate was 0.41 W/m2, which is the fourth largest GHG forcing following carbon dioxide, methane, and tropospheric ozone(see Figure Q17-1). The contributions of individual halocarbon gases controlled by the Montreal Protocol are highlighted in Figure
Q17-2. Within the halocarbons, CFC-12, CFC-11, and CFC-113
combined contribute the largest percentage (67%) to radiative
forcing in 2019. The intermediate-term ODS substitutes, hydrochlorofluorocarbons (HCFCs), make the next largest contribution(15%). The long-term ODS substitutes HFCs plus PFCs and SF6 contribute 13%. Finally, CCl4 and CH3CCl3 contribute an additional 3%
and the minor CFCs and halons contribute 2% to radiative forcing
in 2019.

Stratospheric and tropospheric ozone. Ozone in both thestratosphere and the troposphere absorbs infrared radiation emitted from Earth’s surface, trapping heat in the atmosphere. Ozone
also significantly absorbs solar ultraviolet (UV) 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 a variety of humanactivities has led to increases in global tropospheric ozone (seeQ2), causing a positive radiative forcing (warming) estimated to be+0.47 W/m2 over the 1750–2019 time period, with a range of uncertainty spanning +0.24 to +0.70 W/m2 (see Figure Q17-1). The
large uncertainty in the climate forcing due to release of air pollutants reflects our limited knowledge of changes in the abundanceof tropospheric ozone between 1750 and the mid-1950s as well
Section V: Implications of ozone depletion and the Montreal Protocol 55

Q17 Twenty Questions | 2022 Update
Radiative Forcing of Climate by Halocarbons
Climate forcing due to increases in controlled gases containing chlorine,
bromine, and fluorine from human activities between 1750 and 2019
CFC-12,
CFC-11,
CFC-113

Major CFCs
Sum of CFC-115, CFC-13,
CFC-114, Halon-1211, Halon-1301

Minor CFCs and Halons
Other ODSs:
Carbon tetrachloride
Methyl chloroform

CCl4,
CH3CCl3
HCFC-22,
HCFC-142b,
HCFC-141b
HFCs,
PFCs,
SF6

Hydrochlorofluorocarbons (HCFCs)
Hydrofluorocarbons (HFCs)
Perfluorocarbons (PFCs)
Sulfur hexafluoride (SF6)Montreal Protocol gases
Other halogenated climate gases

Total halocarbon climate forcing
0

0.1

0.2

Radiative forcing (W/m )
2

0.3
0.4
Increasing surface warming

0.5

Figure Q17-2. Halocarbons and radiative forcing of climate. Halocarbon gases in the atmosphere represent an important contribution
to the radiative forcing (RF) of climate since the start of the Industrial Era (see Figure Q17-1). Halocarbons are gases containing chlorine,
bromine, or fluorine atoms, with at least one carbon atom, that contribute to RF by trapping infrared radiation released by Earth’s surface.
The rise in RF between 1750 and 2019 is shown for all halocarbons controlled either under the Montreal Protocol (red) or included in the
Paris Agreement (blue) along with the RF due to the rise in SF6. Note that while SF6 is technically not a halocarbon because it lacks any
carbon atoms, it is an important halogen-containing gas in the atmosphere. Separate contributions to RF of each gas or group of gases are
based upon atmospheric abundance histories and the radiative efficiency specific to each compound. The gases listed in the right-handlabels begin with the largest contribution in each group and proceed in descending order, except for the entry for minor CFCs and halons,
which are shown as one total value. The individual RF terms add together to form the bottom bar, representing the total RF due to controlled halocarbons, PFCs, and SF6. The RFs of CFC-11 and CFC-12, the largest halocarbon contributors, are decreasing and will continue todecline as CFCs are gradually removed from the atmosphere (see Figure Q15-1). In contrast, the total RF of HCFCs, the intermediate-term
ODS substitute gases, is projected to grow for about another one to two decades before decreasing. HFCs are the long-term ODS substitute gases. With the October 2016 Kigali Amendment, the Montreal Protocol now controls future production and consumption of HFCs
with high Global Warming Potentials. As a result, nearly all of the RF due to halogen-containing GHGs is now controlled by the Montreal
Protocol (bottom bar). The future RF of climate due to HFCs is expected to peak in about two decades under the provisions of the Kigali
Amendment (see Q19).

as the difficulty in modeling the complex chemical processes that
control the production of tropospheric ozone.
On the other hand, rising abundances of ODSs in the atmospheresince the middle of the 20th century have led to decreases in stratospheric ozone, causing a negative radiative forcing of –0.02 W/m2
(cooling) over the 1750–2019 time period, with a range of uncertainty spanning – 0.15 to +0.11 W/m2 (see Figure Q17-1). The signof the radiative forcing due to stratospheric ozone depletion is uncertain because this quantity is the difference between two termsof comparable magnitude, each of which has an associated uncertainty. The first term represents the trapping by ozone of outgoing
infrared radiation released by the surface and lower atmosphere:
this is a cooling term because less ozone results in less trapping ofheat. The second term represents the absorption of solar UV radiation by ozone: this is a warming term because less ozone resultsin greater penetration of solar UV radiation into the lower atmo56 Section V: Implications of ozone depletion and the Montreal Protocol

sphere (troposphere). This radiative forcing due to stratosphericozone depletion will diminish in the coming decades, as ODSs aregradually removed from the atmosphere.
It is clear that stratospheric ozone depletion is not a principal causeof present-day global warming. First, the climate forcing fromozone depletion is small. Second, the total radiative forcing of climate from other GHGs such as carbon dioxide, methane, halocarbons, and nitrous oxide is large and positive, leading to warming(see Figure Q17-1). The total forcing from these other GHGs is theprincipal cause of the observed warming of Earth’s surface.

Ozone Depletion Potentials and Global Warming Potentials. A useful way of comparing the influence of individual emissions of halocarbons on ozone depletion and climate change is tocompare Ozone Depletion Potentials (ODPs) and Global Warming
Potentials (GWPs). The ODP and GWP are the effectiveness of anemission of a gas in causing ozone depletion and climate forcing,2022 Update | Twenty Questions

Q17

Evaluation of Selected Ozone-Depleting Substances and Substitute Gases

Chlorine gases

CFC-11(1)

Carbon tetrachloride(0.87)

CFC-113(0.82)

CFC-12(0.75)

HCFC-22 (0.04)
Bromine gases

Ozone-depleting substances

Relative importance of equal mass emissions for ozone depletion and climate change

Halon-1301
Halon-1211
Methyl bromide(2)(0.57)

Hydrofluorocarbons (HFCs)

Substitute Gases

HFC-23 (0)HFC-143a (0)HFC-125 (0)HFC-134a (0)HFC-32 (0)HFC-152a (0)(153)

HFO-1234yf (0)

(<1)

0

4

8

12

16

Ozone Depletion Potential (ODP)
Increasing ozone depletion

0

5000

10000

15000

Global Warming Potential (GWP, 100-yr)
Increasing surface warming

Figure Q17-3. ODPs and GWPs. The environmental impacts of ozone-depleting substances and their substitutes are commonly comparedbased upon their Ozone Depletion Potentials (ODPs) and Global Warming Potentials (GWPs) (see Table Q6-1). The ODPs and GWPs represent the magnitude of ozone depletion and climate forcing, respectively, of a given mass of gas emitted to the atmosphere, relative tothat of CFC-11 (for ODP) or CO2 (for GWP). Therefore, the ODP of CFC-11 and the GWP of CO2 are assigned reference values of 1. The GWPs
shown here are evaluated for a 100-year time interval after emission. The CFCs, halons, and HCFCs are ozone-depleting substances (ODSs)
since they contain either chlorine or bromine (see Q6). HFCs, used as ODS substitutes, do not destroy ozone (ODPs equal zero) since they
are mixtures of only hydrogen, fluorine, and carbon atoms. The ODPs of halons far exceed those of the CFCs, since all halons contain bromine. The GWPs of these gases span a wide range of values, from less than 1 (HFO-1234yf) to 14,700 (HFC-23).

respectively, relative to a reference gas (see Table Q6-1). The principal halocarbon gases are contrasted with each other in Figure
Q17-3. The ODP of CFC-11 and the GWP of carbon dioxide areassigned reference values of 1. The CFCs and carbon tetrachloride all have ODPs near 1, indicating comparable effectiveness incausing ozone depletion per mass emitted. The principal halons
have ODPs greater than 7, making them the most effective ozonedepleting substances per mass emitted. All HFCs have ODPs ofzero since they contain no chlorine and bromine, and therefore do
not directly cause ozone depletion (see Q6).
All halocarbons have non-zero GWPs and, therefore, contributeto the RF of climate. The GWP does not correspond strongly with

the ODP of a gas because these quantities depend on different
chemical and physical properties of the molecule. For example,
while HFC-143a does not destroy ozone (ODP equals zero), eachkilogram emitted is about 6000 times more effective than a kilogram of carbon dioxide in causing climate forcing. When HFCs are
released to the atmosphere, their contribution to climate forcing
depends on their GWPs, which vary over a wide range (less than
1 to 15,000).
Montreal Protocol regulations have led to reductions in CFC emissions and increases in HCFC emissions (see Q15). As a result of
these actions, the total RF from ODSs stopped increasing and is
now slowly decreasing because HCFCs have lower GWPs than
Section V: Implications of ozone depletion and the Montreal Protocol 57

Q17 Twenty Questions | 2022 Update
CFCs (see Q18). However, the overall RF of all halocarbons is slowly increasing because of growing contributions from non-ODS
gases (HFCs, PFCs, and SF6 ). The growth in the HFC contribution
will be limited by the provisions of the 2016 Kigali Amendment(see Q19). It is important to note that despite having a GWP that issmall in comparison to many other halocarbons and other greenhouse gases, carbon dioxide is the most important greenhousegas produced by human activities because its emissions are large,
its atmospheric lifetime is long, and its atmospheric abundanceis far greater than those of all other greenhouse gases associatedwith human activities.

The Antarctic ozone hole and Southern Hemisphereclimate. While stratospheric ozone depletion is not the principal
cause of global climate change, the reoccurring Antarctic ozone
hole has contributed to observed changes in climate parameters in
the atmosphere and oceans of the Southern Hemisphere since theearly 1980s. These research findings are explained in more detailin the box below.

The Antarctic Ozone Hole and Southern Hemisphere Surface Climate
Links between stratospheric ozone depletion and changes in surface climate were first 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 has occurred every spring since the early 1980s,
was shown to contribute to observed changes in Southern Hemisphere (SH) surface climate during summer due to the effect of theozone hole on atmospheric circulation.
The severe springtime depletion of ozone over the Antarctic leads to a strong cooling of the polar lower stratosphere persisting into
early summer in the SH. This cooling increases the temperature contrast between the tropics and the polar region and strengthens
stratospheric winds. As a result, in the SH there has been 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 stream (which is associated withweather systems). There is evidence from both models and observations that subtropical and midlatitude summer precipitation
patterns in the SH have been affected by these changes. Model studies indicate that even though long-lived greenhouse gases thatcause climate change exacerbate this shift in the summertime tropospheric circulation in the Southern Hemisphere, ozone depletionhas been the dominant contributor to the observed changes since the early 1980s. Paleoclimate reconstructions suggest the current
state of these climate features caused by the annually recurring ozone hole is unprecedented over the past 600 years.
Initial signs of the recovery of the ozone hole, particularly during September, are emerging (see Q10). During the 21st century, as
the ozone hole further recovers due to the decline of stratospheric halogens, the ozone-depletion related climate impacts discussed
above will lessen (see Q20). The influence of the ozone hole on SH circulation trends during summer has recently stabilized, due tothe initial recovery of the area and depth of the ozone hole (see Q10). The Southern Hemisphere surface climate response to ozone
depletion in other seasons is weaker than the summer response.

58 Section V: Implications of ozone depletion and the Montreal Protocol

Q18

Are Montreal Protocol controls ofozone-depleting substances also
helping protect Earth’s climate?

Yes. Many ozone-depleting substances (ODSs) are also potent greenhouse gases that contribute to global
warming when they accumulate in the atmosphere. Montreal Protocol controls have led to a substantial
reduction in the emissions of ODSs over the last two decades. These reductions, while protecting the ozone
layer, have the additional benefit of reducing the human contribution to climate change. Without Montreal
Protocol controls, the global warming due to ODSs could now be nearly three times the present value. Withthe 2016 Kigali Amendment to the Montreal Protocol, climate protection was extended to include controls
on HFCs, which do not deplete ozone but contribute to global warming (see Q19).

The success of the Montreal Protocol in controlling the production and consumption of ozone-depleting substances (ODSs) has
protected the ozone layer (see Q14). The resulting reductions inemissions and atmospheric abundances of ODSs also decreased
the human influence on climate because all ODSs are greenhouse gases (see Q17). By protecting both ozone and climate, theMontreal Protocol has provided a dual benefit to society and
Earth’s ecosystems. As shown in Figure Q18-1 and described below, the dual benefit of the Montreal Protocol is highlighted by considering long-term baseline and world-avoided scenarios of ODS
emissions, Ozone Depletion Potentials (ODPs), Global Warming
Potentials (GWPs), equivalent effective stratospheric chlorine(EESC), and the radiative forcing (RF) of climate.

Baseline ODS scenario. The baseline scenario refers to actual
past ODS emissions of the principal halogen source gases and projected emissions for the years 2021 to 2025. The baseline scenario
is labeled “from observed ODS abundances” in Figure Q18-1
since, for 1960–2020, the emissions are based upon analysis ofobserved abundances of the principal ODS gases at Earth’s surface(see Figure Q15-1). This scenario also includes emissions of thenaturally occurring halogen source gases methyl chloride (CH3Cl)
and methyl bromide (CH3Br). For this scenario the peak emission of
ODSs occurs in the late 1980s (see Figure Q0-1).
For all of the emission scenarios shown in Figure Q18-1, the annual emissions of each gas are added together after being weighted (multiplied) by their corresponding Ozone Depletion Potential(ODP) (upper left) or Global Warming Potential (GWP) (upper right)(see Q17 and Table Q6-1). The ODP and GWP of a given gas quantify how effective the gas is at destroying ozone (ODP) or warming
climate (GWP) for the emission of a certain mass of the gas, relative
to the effect on ozone or warming of the emission of the same massof CFC-11 (for ODP) or CO2 (for GWP). In both cases, the reference
gases (CFC-11 and CO2 ) are assigned a value of 1, and the ODP and
GWP for all other gases are scaled accordingly (see Table Q6-1
and Q17). For example, 1 kg of halon-1211 emissions is expressedas 7.1 kg of CFC-11-equivalent emissions because the ODP of
halon-1211 is 7.1. Similarly, the GWP-weighted sum is expressed as
CO2 -equivalent emissions because CO2 is the reference gas, withan assigned GWP of 1. Likewise, 1 kg of carbon tetrachloride emissions is considered to be 2150 kg of CO2 -equivalent emissions
because the GWP of carbon tetrachloride is 2150. GWP-100 values are shown here and throughout reflecting a choice of a time
horizon of 100 years.

World-avoided ODS scenario. The baseline scenario of ODS
emissions can be contrasted with a scenario of ODS emissions thatthe world has avoided by successfully implementing the Montreal
Protocol (see Figure Q18-1). The world-avoided scenario is derived
by assuming that, from 1987 onwards, emissions of ODSs increaseat a rate of 3% per year. This growth rate is consistent with thestrong market for ODSs in the late 1980s, which included a wide
variety of current and potential applications and had the potential
for substantial growth in developing countries.
CO2 emission scenario. Long-term emissions of CO2 are also

shown in the upper right panel of Figure Q18-1. Atmospheric CO2 isthe principal greenhouse gas emitted by human activities. The CO2
emission curve represents global emissions from the sum of eachnation’s reported emissions from the combustion of coal, oil, natural gas, as well as the fuels used by the world’s ships and airplanes,
cement manufacturing, and the release of CO2 due to global deforestation.

ODP-weighted emissions. The ODP-weighted emission scenario based upon observed ODS abundances is one measure of how
the overall threat to stratospheric ozone from ODSs has changed
over time (see Figure Q18-1, upper left panel). Since most ODSs
remain in the atmosphere for years (see “Atmospheric lifetime” column in Table Q6-1), when ODP-weighted emissions rise this means
there will be an increase in ozone destruction for many futureyears. Conversely, when emissions decline, less ozone will be destroyed in future years than if emissions had remained high. Annual
ODP-weighted emissions increased substantially between 1960
and 1987, the year the Montreal Protocol was signed (see FigureQ0-1). After 1987, annual ODP-weighted emissions began a longand steady decline to present-day values. The decline in emissions
is expected to continue, causing the atmospheric abundances of
all individual ODSs to eventually decrease (see Figure Q15-1). The
Section V: Implications of ozone depletion and the Montreal Protocol 59

Q18 Twenty Questions | 2022 Update
The Montreal Protocol Protection of Ozone and Climate
from Global Emissions of Ozone-Depleting Substances (ODSs)

5
4

Emissions

Emissions(megatonnes CFC-11-equivalent per year)

6

Emissions weighted by Global
Warming Potentials (GWPs)

3
Montreal Protocol(1987)

2

Ozone Layer
Protection

1
0
1960

1980

Year

2000

(gigatonnes CO2-equivalent per year)

Emissions weighted by Ozone
Depletion Potentials (ODPs)
40

30

20

Equivalent effective
stratospheric chlorine (EESC)

8000

ClimateProtection

10

0
1960

2020

Montreal Protocol(1987)

1980

Year

2000

2020

Radiative forcing of climate
2.0

7000

Radiative forcing

5000
Ozone Layer
Protection

4000
3000

Montreal Protocol(1987)

2000(watts per square meter)

EESC(parts per trillion)

6000
1.5
ClimateProtection

1.0
Montreal Protocol(1987)

0.5

1000
0
1960

1980

Year

2000

from observed ODS abundances

2020

0.0
1960

World-avoided scenario

1980

Year

2000

2020

CO2 from human activities

Figure Q18-1. Montreal Protocol protection of ozone and climate. The Montreal Protocol has protected the ozone layer and also reduced
the potential for climate change, because ozone-depleting substances (ODSs) are greenhouse gases (GHGs). The baseline ODS scenario(orange line) includes actual emissions of all principal gases weighted either by their Ozone Depletion Potentials (ODPs) (upper left) or
Global Warming Potentials (GWPs) (100-yr timeline) (upper right) coupled with projected emissions for years 2021 to 2025. With theseweightings, emissions are expressed as CFC-11-equivalent or CO2-equivalent mass per year. The lower panels show equivalent effective
stratospheric chlorine (EESC) (see Figure Q15-1) and total radiative forcing (RF) of climate (see Figure Q17-2), as derived from the observed
abundances of ODSs as well as projected abundances for 2021 to 2025. The world-avoided emission scenario (purple line) assumes a 3%
per year growth in the emission of ODSs beyond 1987 values, consistent with the assumption for the No Protocol scenario shown in Figure
Q14-1. The emission and RF of atmospheric CO2 (magenta line) are shown for reference on the right panels. The magnitude of the dual
benefit of the Montreal Protocol has steadily increased since 1987, as shown by the differences between the world-avoided scenario andthe observed ODS abundance scenario (purple shaded region) in each panel.
(A megatonne = 1 million (106) metric tons = 1 billion (109) kilograms. A gigatonne = 1 billion (109) metric tons = 1 trillion (1012) kilograms. A CFC-11-equivalentemission of an ODS is an emission amount that results in the same ozone depletion as the release of the same mass of CFC-11; a CO2-equivalent emission ofa non-CO2 GHG is an emission amount that results in the same RF of climate over a 100-year time interval as the release of the same mass of CO2. Emissionsof CO2 used in this figure are from the Global Carbon Project.

60 Section V: Implications of ozone depletion and the Montreal Protocol

2022 Update | Twenty Questions
reductions in ODP-weighted emissions relative to the peak value in1987 represent lower limits of the annual emissions avoided by theMontreal Protocol, which are a measure of its increasing successover time in protecting the ozone layer.
The upper limits of annual reductions in ODP-weighted emissions
are derived from the world-avoided scenario. The difference between the world-avoided emission scenario and the baseline scenario (purple shaded region in Figure Q18-1, upper left panel) represents an estimate of the ozone layer protection provided by theMontreal Protocol.

GWP-weighted emissions. The GWP-weighted emission scenario based upon observed ODS abundances is a measure of how
the overall threat to Earth’s climate from ODSs has changed overtime (see Figure Q18-1, upper right panel). As GWP-weighted emissions rise, the RF of climate in the future due to the accumulationof ODSs in the atmosphere also 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 adecrease afterwards. This similarity follows from the dominant rolethat both CFC-11 and CFC-12 play in ozone depletion and climate
forcing from ODSs. The difference between the world-avoided
emission scenario and the baseline scenario (purple shaded regionin Figure Q18-1, upper right panel) represents an estimate of theclimate protection provided by the Montreal Protocol.
Annual GWP-weighted emissions of ODSs were a large percentage (about 20– 40%) of global emissions of CO2 between 1960
and 1987. Thereafter, this percentage has steadily decreased andwas 2–3% of global CO2 emissions in 2022. This past trend stands
in sharp contrast to the world-avoided scenario, in which emissionsof ODSs rise to more than 50% of CO2 emissions in 2022. Anotherway to understand the climate benefit of the Montreal Protocol is tocompare the height of the purple shaded region in 2022 to the risein the emissions of CO2 since 1987, as shown in Figure Q18-1 (upper
right panel). These two quantities are nearly equal in magnitude,
demonstrating that since 1987 the Montreal Protocol has avoided
an increase in GWP-weighted emissions of ODSs that nearly equals
the increase in global emissions of CO2 over this same time period.

EESC scenarios. The EESC scenarios in Figure Q18-1 (lower left
panel) provide a measure of the year-to-year potential of the atmospheric abundances of ODSs to destroy stratospheric ozone.
Two scenarios are shown: the baseline that uses observed abundances of ODSs (with a projection to 2025) and the world-avoided scenario described above. The derivation of EESC from ODS
atmospheric abundances is discussed in Q15 and the same EESC
baseline scenario is shown in Figures Q13-1, Q14-1 (red curve),
and Q15-1 for different time intervals. When ODS- weighted emissions declined after 1987, EESC did not decrease in a proportionalmanner because of the long atmospheric lifetimes of the principal
ODSs (see Table Q6-1). As shown in Figure Q18-1, EESC reached its

Q18

peak value nearly a decade after the peak in ODP-weighted emissions, and by 2022 the decrease in EESC from its peak value wasonly about 18%, compared to the 80% decrease in ODP-weighted
emissions achieved by 2022. Conversely, had the emissions of
ODSs followed the world-avoided scenario, EESC would be morethan twice the value in today’s stratosphere. In this case, computer
simulations show global total ozone values in 2020 are about 17%
lower than the 1964–1980 average. Even larger depletions occur
in subsequent years. The Montreal Protocol and its amendments
and adjustments have provided vitally important protection to theglobal ozone layer and climate.

Radiative forcing (RF) scenarios. The RF of climate scenarios in
Figure Q18-1 (lower right panel) provide a measure of the year- toyear contribution to climate change from the atmospheric abundances of ODSs. The RF of an ODS is equal to the net increase in its
atmospheric abundance since 1750 multiplied by its radiative efficiency, which quantifies how effective a given ODS molecule is atretaining infrared radiation. The RF of ODSs up to the present is calculated using observed atmospheric abundances. The RF due toODSs increases smoothly from 1960 onward, peaks in 2010, anddecreases very gradually in subsequent years. The decline of RF ofclimate in response to ODS emission reductions is slow becauseof the high abundances of the two principal contributing gases,
CFC-11 and CFC-12, and their long atmospheric lifetimes of about
50 and 100 years, respectively.

Increasing the benefits of the Montreal Protocol. The benefitof the Montreal Protocol for protection of climate was expanded
in 2016 through the Kigali Amendment, which placed controlson the production and consumption of some hydrofluorocarbons(HFCs) (see Q19). HFC compounds do not contain chlorine or bromine, and therefore do not deplete ozone. Many HFC gases havea high radiative efficiency and a long atmosphere lifetime, whichleads to significant global warming (see Figure Q19-2). The ozone
layer and climate benefits of the Montreal Protocol could be further
increased by expanded capture and destruction of halons, chlorofluorocarbons (CFCs), and hydrochlorofluorocarbons (HCFCs) inbanks, by avoiding emissions in continued use of ODSs as feedstock for the production of other chemicals, and by eliminating
future emissions of halogen source gases not controlled by theMontreal Protocol, such as dichloromethane (CH2Cl2 ). Banks arelargely associated with ODSs contained in refrigeration, air conditioning, fire protection equipment, insulating foams, and stockpiles for servicing long-term applications. Atmospheric release of
ODSs from existing banks is projected to contribute more to ozone
depletion in the coming decades than the limited production andconsumption of ODSs (HCFCs and CH3Br) allowed by the Montreal Protocol after 2023. If all available options were implemented to
avoid future atmospheric release of ODSs starting in 2023, the return of EESC to 1980 values would be advanced by about a decadefor both the midlatitude (see Fig Q14-1) and polar stratosphere.

Section V: Implications of ozone depletion and the Montreal Protocol

61

Q19

How has the protection of climate
by the Montreal Protocol expanded
beyond the regulation of ozonedepleting substances?

At the 28th Meeting of the Parties to the Montreal Protocol held in Kigali, Rwanda, in October 2016, theMontreal Protocol was amended to control the production and consumption of hydrofluorocarbons (HFCs).
The Montreal Protocol phaseout of chlorofluorocarbons (CFCs) led to the temporary use of hydrochlorofluorocarbons (HCFCs). The subsequent phaseout of HCFCs led to expanded long-term use of HFCs, because
HFCs pose no direct threat to the ozone layer. However, HFCs are greenhouse gases and therefore contribute to climate change. Limiting the production and consumption of those HFCs with high global warming
potentials is projected to avoid 0.3 to 0.5°C of global warming over this century. The Kigali Amendment
marks the first time the Montreal Protocol has adopted regulations solely for the protection of climate.

The control of ozone-depleting substances (ODSs) by the Montreal Protocol provides the dual benefit of protecting Earth’s ozone
layer and global climate (see Q18). The widespread global use ofhydrofluorocarbons (HFCs) and their projected future growth in
the coming decades has been recognized by the Montreal Protocol as a potentially significant contribution to climate change fromhuman activities. In response, the Kigali Amendment was adopted
to control production and consumption of HFCs with high Global Warming Potentials (GWPs) (see Q17). Full compliance with theprovisions of the Kigali Amendment will significantly enhance theclimate-protection benefit of the Montreal Protocol.

Hydrofluorocarbons (HFCs). HFCs are replacement compounds
for ODSs that were chosen because they contain no chlorine or
bromine that cause ozone depletion. HFCs are widely used in
the residential air-conditioning and refrigeration sectors and asfoam-blowing agents, spray-can propellants, and feedstocks forthe production of other chemicals. These uses are growing as theglobal phaseout of hydrochlorofluorocarbons (HCFCs), the early
replacement compounds, nears completion. The GWPs of HFCs
vary over a wide range due to differences in their physical and radiative properties (see Table Q6-1 and Figure Q17-3). For example,
the GWP of HFC-134a (primarily used in air conditioning and refrigeration) is 1470, which means that after release to the atmosphere,each kilogram of HFC-134a is 1470 times more effective than akilogram of CO2 in increasing climate forcing over a century-longtime period. In contrast, the GWP of HFO-1234yf, a substitute for
HFC-134a, is less than 1.
HFC-23. HFC-23 is considered separately in the Kigali Amendment because this gas is primarily produced as an unwanted byproduct in the manufacture of HCFC-22 and HFCs. The global
warming potential of HFC-23 is quite large (14,700), in part due
to its long atmospheric lifetime of 228 years. Although many methods exist to chemically destroy HFC-23 at production facilities, thiscompound continues to be released to the atmosphere. For example, the atmospheric abundance of HFC-23 increased by 44%
62 Section V: Implications of ozone depletion and the Montreal Protocol

between 2009 and 2019. In 2019, the radiative forcing of HFC-23
was 0.006 W/m2, which is approximately 15% of the total forcing
from all HFCs. The Kigali Amendment phases down, in conjunction with the other HFCs, unwanted by-production of HFC-23, but
provides no specific control measures for emissions of HFC-23.
Instead, the amendment directs nations to destroy HFC-23 to theextent practicable in order to avoid future emissions and the associated increased climate forcing.

Climate implications of HFC use. The total global emission of
HFCs expressed in terms of CO2-equivalent emissions has grownsteadily since 2000, equaling about 1 gigatonne CO2-equivalent
per year in 2020 (see Figure Q19-1). The primary emissions of
HFCs are of HFC-134a as well as HFC-143a, HFC-125 and HFC32, which are widely used in blended refrigerants such as R404A(52% HFC-143a, 44% HFC-125, and 4% HFC-134a) and R410A(50% HFC-32, 50% HFC-125). Recent growth in the consumption(and emissions) of HFCs is due in part to replacing HCFCs thatare being phased out under the Montreal Protocol with HFCs. In
2019, the atmospheric abundances of HFCs contributed about
10% of climate forcing from all halocarbon compounds (see Figure
Q17-2) and less than 1% of the total climate forcing from all otherlong-lived greenhouse gases (see Figure Q17-1). Projections basedon current production and consumption patterns and future economic growth indicate that, without the Kigali Amendment, HFCemissions could have reached around 5 gigatonnes CO2-equivalent
per year by 2050 and nearly double that value by 2100 (see Figure
Q19-1). This projected emission value for 2050 is about one half ofthe peak in CO2-equivalent emissions of ODSs in 1987 (see Figure
Q18-1). Thus, in the absence of the Kigali Amendment, the projected growth in HFC emissions in the coming decades offsets a significant amount of the climate protection gained from reductions in
ODS emissions under the Montreal Protocol.

Kigali Amendment. The future of HFC emissions was changed by
the Montreal Protocol with the adoption of the Kigali Amendment
in 2016. The amendment requires a phasedown of the global pro-

2022 Update | Twenty Questions

Q19

Projected Emissions of Hydrofluorocarbons (HFCs)

Emissions(gigatonnes CO2 –equivalent per year)

10

Without the Kigali Amendment

10

8

8

6

6

4

4

2

2

0
2000 2020

2040

2060

Year

2080 2100

With the Kigali Amendment02000 2020

Low-GWP Alternatives
Other HFCs
HFC-32 (GWP = 705)
HFC-134a (GWP = 1360)
HFC-143a (GWP = 5080)
HFC-125 (GWP = 3450)

2040

2060

Year

2080 2100

Figure Q19-1. HFC emissions and the Kigali Amendment. The Kigali Amendment to the Montreal Protocol limits the production andconsumption of a group of hydrofluorocarbons (HFCs) with high global warming potentials (GWPs). HFCs are considered replacement
compounds for ozone-depleting substances (ODSs) because HFCs lack chlorine and bromine, and therefore pose no direct threat to theozone layer. Avoiding substantial emissions of high-GWP HFCs through the implementation of the Kigali Amendment will increase futureclimate protection. HFCs have a wide range of GWPs given their different physical and chemical properties (see Table Q6-1 and Figure
Q17-3). The panels show emissions of widely used high-GWP HFCs. The emissions are weighted by the 100-yr GWP of each compound;with this weighting, emissions are expressed as CO2-equivalent mass per year. In the left panel, emissions are based upon an analysis of
atmospheric observations up to 2013 and projections to 2100 that represent an upper range to future global emissions in the absence ofthe Kigali Amendment and national regulations. The right panel shows GWP-weighted emissions based on atmospheric observations up
to 2020 as well as projections to 2100, again assuming international compliance with the provisions of the Kigali Amendment. The projections in the right panel include a category termed low-GWP alternatives that is comprised of refrigerant compounds that have GWPs much
lower than the refrigerants they replace. Low-GWP alternatives include a subset of HFCs known as hydrofluoroolefins (HFOs), which arealso composed only of hydrogen, fluorine and carbon atoms. The chemical structure of HFOs results in these compounds being more reactive in the lower atmosphere (troposphere) than other HFCs and, consequently, HFOs have shorter lifetimes after atmospheric release (seeTable Q6-1). As a result, emissions of HFOs cause substantially lower radiative forcing than emissions of the same mass of high-GWP HFCs.
(A gigatonne = 1 billion (109) metric tons = 1 trillion (1012) kilograms. A CO2 -equivalent emission of a non-CO2 GHG is an emission amount that results in thesame RF of climate over a 100-year time interval as a release of the same mass of CO2.)

duction and consumption of high-GWP HFCs by more than 80% (inCO2 -equivalent) from the baseline level over the next 30 years. The
phasedown schedule accommodates the concerns and interestsof developed and developing countries, including those with highambient temperatures that are likely going to have future increased
demand for the use of air-conditioners. The Kigali Amendment
entered into force on 1 January 2019. Figure Q19-1 shows how theamendment provisions dramatically reduce projected emissionsof HFCs in the coming decades. The emissions of HFCs that are
avoided by 2100 total about 420 gigatonnes CO2 -equivalent,
which is more than 10 years of present-day annual emissions of CO2
due to human activities.

Expanding climate protection. The Kigali Amendment substantially expands the protection of climate afforded by the Montreal
Protocol (see Q18). With full implementation of the amendment,

annual global emissions of HFCs reach their peak value before
2040 (see Figure Q19-1). Without the amendment, yearly emissions are projected to increase until market saturation is reachedin the second half of the century, at a value of about 10 gigatonnes
CO2 -equivalent per year, nearly five times more than the emission
peak under the amendment. Furthermore, as shown in Figure
Q19-2, the long-term radiative forcing of climate, which is proportional to atmospheric abundances, is substantially reduced. Without the amendment, projected radiative forcing from HFCs increases throughout this century, reaching a value of about 0.6 W/m2 in
2100. In this scenario, radiative forcing due to HFCs by the end of
the century exceeds that of nitrous oxide and rivals that of methane. With the amendment, the radiative forcing of climate by HFCs
reaches a peak value before 2050 and gradually decreases to about
0.07 W/m2 in 2100. The ranges of radiative forcing values for methSection V: Implications of ozone depletion and the Montreal Protocol 63

Q19 Twenty Questions | 2022 Update

Radiative forcing of HFCs

Emission of HFCs

8
6
4
2
0
2000 2020 2040 2060 2080 2100

Year

1.0
0.8
0.6
0.4
0.2
0.0
2000 2020 2040 2060 2080 2100

Year

Without the Kigali Amendment
With the Kigali Amendment

Temperature increase due to HFCs
Temperature change (°C)

10

Radiative forcing (W m-2)

HFC Emissions(gigatonnes CO2-equivalent per year)

Climate Benefit of the Kigali Amendment
0.6
0.5
0.4
0.3
0.2
0.1
0.0
2000 2020 2040 2060 2080 2100

Year

Methane range in 2100
Nitrous oxide range in 2100

Figure Q19-2. Kigali Amendment Climate Protection. The successful implementation of the Kigali Amendment will enhance the protection
to Earth’s climate afforded by the Montreal Protocol. The panels display the CO2-equivalent emissions (left), radiative forcing (middle) andsurface temperatures (right) for HFC emission scenarios without (blue shaded regions) and with (gold lines) the implementation of the Kigali
Amendment and national regulations. Historical emissions of HFCs are derived from atmospheric observations. Emissions for latter years
are based upon projections of production and consumption patterns as well as future economic growth. All emissions are weighted by the
100-yr GWP of each compound (CO2-equivalent emissions). Emissions of HFC-23 are excluded. The emission projections without the Kigali
Amendment and national regulations are based on lower and upper ranges of projected HFC consumption. The increases in global meansurface temperature from HFC emissions are shown beginning in year 2000. For comparison, the radiative forcing and surface temperature
increases are shown for methane (CH4 ) and nitrous oxide (N2O) in the middle and right panel margins, respectively, for year 2100 based on
accumulated emissions since 1750. Compliance with the Kigali Amendment has the potential to avoid a 0.3 to 0.5°C rise in global surface
temperature over this century due to restrictions on the future emission of high-GWP HFCs.
(A gigatonne = 1 billion (109) metric tons = 1 trillion (1012) kilograms. The end of century values for CH4 and N2O are based upon the Shared Socioeconomic
Pathway (SSP) SSP1-2.6 (lower limit) and SSP3-7.0 (upper limit) scenarios.)

ane and nitrous oxide in 2100 as shown in Figure Q19-2 far exceed
the 0.07 W/m2 forcing due to HFCs under the Kigali Amendment.
The benefit of avoiding HFC radiative forcing over many decades
as a result of the Kigali Amendment provisions can be expressedas an avoided increase in globally averaged surface temperature.
The increase in temperature by the year 2100 due to future atmospheric growth of HFCs without the Kigali Amendment and national regulations is projected to be between 0.3 and 0.5°C (seeFigure Q19-2). In contrast, the temperature increase is projected
to be about 0.06°C with full implementation of the amendment,
which is significantly less, for example, than the warming expected from projected abundances of methane and nitrous oxide in
2100. Currently, global warming due to all emissions from humanactivities is about 1.2°C since 1750, the start of the Industrial Era.
The goal of the United Nations Framework Convention on Climate
Change Paris Agreement is to limit global warming to well below
2.0°C since the start of the Industrial Era and to pursue efforts tolimit global warming to 1.5°C. The temperature increase of 0.3 to
0.5°C avoided by the Kigali Amendment contributes substantiallyto the achievability of this goal.

64 Section V: Implications of ozone depletion and the Montreal Protocol

Low-GWP substances. The Kigali Amendment encourages theuse of low-GWP substances or other alternatives to replace highGWP HFCs in the coming decades (see Table Q6-1 and Figure
Q17-3). Other alternatives include propane, ammonia, and otherclimate-friendly technologies. The low-GWP substances include
a subset of HFCs known as hydrofluoroolefins (HFOs), which arealso composed only of hydrogen, fluorine and carbon atoms. The
chemical structure of HFOs includes a double carbon bond, causing these compounds to be more reactive in the lower atmosphere(troposphere) than other HFCs. Consequently, HFOs have very
short atmospheric lifetimes. One such compound, HFO-1234yf,
has a lifetime of only 12 days, in contrast to HFC-23, HFC-143a, andHFC-134a with lifetimes of 228, 52, and 14 years, respectively (seeTable Q6-1). The short atmospheric lifetimes of HFOs lead to very
low GWPs. As a result, the emission of an HFO results in substantially lower climate forcing than the forcing caused by emission of
the same mass of high-GWP HFCs (see Figure Q19-1).
The projections of emissions under the Kigali Amendment include
a group of compounds labeled Low-GWP Alternatives in Figure
Q19-1. These compounds are expected to cover the application
demand from sectors in which the use of high-GWP HFCs is phased2022 Update | Twenty Questions
down. Even with the emissions of a large mass of these low-GWP
alternatives, the future projected contribution to climate change
is much less than contributions from projected emissions of highGWP HFCs in the absence of the Kigali Amendment.

Other environmental consequences of HFC use. The atmospheric abundance of trifluoroacetic acid (TFA, chemical formulaC2HF3O2 ) is expected to increase in the coming decades due to future emissions of HFCs (including HFOs), HCFCs, and related compounds. When these compounds breakdown in the atmosphere,
they produce TFA, a persistent, long-lived chemical with potentially harmful effects on animals, plants, and humans. The current concentration of TFA in rainwater and ocean water is generally very far
below toxicity limits. Potential future environmental impacts of TFA
are a subject of active research.

The Future. The phasedown of HFCs under the Kigali Amendment sets a path in which HFCs play a very limited role in future

Q19

climate forcing. Achieving the maximum climate protection fromthe implementation of the amendment requires that compounds
replacing high-GWP HFCs have much smaller or negligible GWPs.
Technological developments related to low-GWP replacement
substances or other alternatives along with improved refrigeration and air conditioning equipment will help achieve this maximum protection. The release of greenhouse gases in generating
electricity for powering refrigeration and air conditioning equipment contributes to the indirect climate forcing from this sector.
Improvements in the energy efficiency of equipment in this sector
during the transition to low-GWP alternative refrigerants could
potentially double the direct climate benefits of the Kigali Amendment. The combination of low-GWP replacement compounds,
energy efficiency improvements, and the growth in renewable energy sources has great potential to minimize the direct and indirectcontributions to climate forcing from global refrigeration and air
conditioning applications.

Section V: Implications of ozone depletion and the Montreal Protocol 65

Q20

How is stratospheric ozone expected
to change in the coming decades?

Recovery of the global ozone layer from the effects of ozone-depleting substances (ODSs) is expectedaround the middle of the 21st century, assuming global compliance with the Montreal Protocol. Recovery
will occur as ODSs and reactive halogen gas abundances in the stratosphere decrease in the coming decades. In addition to responding to ODSs, ozone abundances will increasingly be influenced by climatechange. 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.

The expected recovery of the global and polar ozone layer is a direct consequence of the success of the Montreal Protocol in reducing the global production and consumption of ODSs. Currently,
the atmospheric abundances of most major ODSs and the associated annual values of equivalent effective stratospheric chlorine(EESC) are in decline (see Q15). In contrast to the diminishing roleof ODSs, changes in climate are expected to have an increasing
influence on future levels of total ozone. Climate change is driven
by the projected growth in the abundance of greenhouse gases(GHGs), primarily carbon dioxide (CO2), methane (CH4), and nitrous oxide (N2O). Increasing amounts of GHGs will lead to changes in temperature, chemistry, and the circulation of the stratosphere, all of which affect ozone. Chemistry-climate models canbe used to project how ozone is expected to respond to changes
in ODSs and climate in particular geographic regions during therecovery period. Sporadic events such as major volcanic eruptions
and wildfires, or deliberate actions such as the injection of aerosols
into the stratosphere to mitigate global warming may also influence future ozone levels.

Using chemistry-climate models. Chemistry-climate models(CCMs) are complex computer programs that simulate temperature, winds, radiation, and the chemical composition of the atmosphere, including stratospheric ozone. The simulations of ozone
are conducted as a function of time, altitude, and geographic location. Projections of total ozone presented here are based on the
results from a group of chemistry-climate models that account for
the influences of changes in ODSs and GHGs. These models show
how changes in ozone are expected to vary across geographic regions by evaluating the complex interactions of the processes that
control ozone and climate involving radiation, chemistry, and transport of chemicals by stratospheric winds. Model inputs include
historical and projected concentrations of ODSs, GHGs (including CO2, CH4, and N2O), air pollutant gases, as well as incoming
solar radiation. The results from chemistry-climate model simulations are used to identify particular processes that are importantfor future abundances of ozone. For example, model projections
driven by increases in the atmospheric abundance of GHGs overthe coming decades show a strengthening in the global-scaleatmospheric circulation that brings air from the troposphere into
66 Section VI: Stratospheric ozone in the future

the stratosphere in the tropics, moves stratospheric air poleward
into both hemispheres, and then returns air to the troposphere at
middle to high latitudes. These circulation changes will significantly
alter the global distribution of ozone and the atmospheric lifetimesof ODSs and other long-lived gases (see Q6). Also, while Earth’s
surface is expected to continue to warm in response to positive radiative forcing of climate from GHGs (see Q17), the stratosphere is
expected to continue to cool. A colder upper stratosphere leads
to increases in ozone because lower temperatures slow down thegas-phase reactions responsible for ozone loss. Conversely, colder
conditions at lower altitudes in the polar stratosphere have the potential to exacerbate ozone loss, especially if the presence of polar
stratospheric clouds (PSCs) increases.
Methane and nitrous oxide are both involved in the chemistry that
determines levels of stratospheric ozone. The main effect of higher
amounts of CH4 is to increase ozone, while that of rising N2O is todecrease ozone. The stratospheric decomposition of CH4 leads tomore reactive hydrogen gases that produce ozone in the lowest
parts of the stratosphere and increases the conversion of reactive
chlorine into its reservoir gas HCl (see Q7). The decomposition of
CH4 also leads to larger abundances of H2O, which cool the upper stratosphere, slowing down ozone loss reactions. Conversely,
the decomposition of N2O produces reactive nitrogen gases that
destroy ozone (see Q8). The effect of higher amounts of N2O on
the depletion of the ozone layer becomes increasingly importantas halogen levels decline.

Simulating recent ozone changes. Comparisons of model
results with observations help confirm the causes of ozone depletion and increase confidence in model projections of future ozone
amounts. Two important measures are the globally averaged total
column ozone outside of polar regions (see Q3) and total ozone in
the Antarctic during October (the month of peak ozone depletion).
These observations are compared to simulations from a group ofchemistry-climate models in Figure Q20-1. Both time series ofozone show substantial depletion since 1980. The average model
values of ozone follow the observed general decline in both regions, suggesting that the main processes involved in ozone depletion are well represented by these models.

2022 Update | Twenty Questions Q20

Changes in Global and Antarctic Ozone
Observations and model projections

Total ozone (DU)

60°S – 60°N
300

1980 level

290

280

350

Total ozone (DU)

Global total ozone

Observations
Historical simulation
Simulated uncertainties

90°S – 70°S

300

Antarctic total ozone (October)1980 level

250
200

Future Projection Scenarios:
Low climate forcing (SSP1-2.6)
Medium climate forcing (SSP2-4.5)
High climate forcing (SSP3-7.0)

150
1960

1980

2000

2020

2040

2060

2080

2100

Year
Figure Q20-1. Simulations of ozone depletion. Chemistry-climate model (CCM) simulations that account for changes in ozone-depleting
substances (ODSs) and greenhouse gases (GHGs) are widely used to assess past ozone changes and project future values of ozone. Agreement between model results and observations increases confidence in the model projections and our understanding of the processes
leading to ozone depletion. Observed values of total ozone (in Dobson units, DU) annually averaged over 60°S–60°N latitude (top panel)and total ozone values over 70°S–90°S during October (bottom panel) decreased during the latter decades of the prior century (magentalines). Future ODS projections assume compliance with the Montreal Protocol, and future abundances of GHGs follow either the SSP1-2.6(low climate forcing), SSP2-4.5 (medium climate forcing), or SSP3-7.0 (high climate forcing) scenario for 2015–2100. The simulation uncertainties for modeled ozone represent the standard deviation about the multi-model mean (MMM), either added to the SSP3-7.0 MMM(upper limit) or subtracted from the SSP1-2.6 MMM (lower limit). As ODS abundances decrease in this century, CCMs project global totalozone to increase steadily and exceed the 1980 value by mid-century for the medium climate forcing scenario. Antarctic ozone during
October is projected to exceed the 1980 values around mid-century for all three scenarios.

There are significant year-to-year variations in global and Antarctic
ozone that are not captured by these simulations. The differences between the observed and modeled values of ozone are due
to factors such as interannual meteorological variability that are
not well represented in these simulations. Over the time period
1996–2020, observed global ozone has exhibited considerable
variability and risen slightly, by less than 1% (see Q12). Antarctic
ozone during October has exhibited even more year-to-year variability, but there are emerging indications that the ozone hole has
diminished in size and depth (maximum amount of ozone deple-tion) since 2000, particularly during September when the effectsof meteorological variability are smaller (see Q10).

Long-term total ozone projections. Total ozone changes derived from chemistry-climate models are shown for 1960 to 2100 in
Figure Q20-1. These simulations use the projected abundances of
ODSs given in the 2018 Scientific Assessment of Ozone Depletion
report. A computation of EESC (see Q15) based on this 2018 projection of ODSs shows a return to the 1980 level in years 2061 and
2077 for the midlatitude and polar stratospheres, respectively.
Section VI: Stratospheric ozone in the future 67

Q20 Twenty Questions | 2022 Update
The model simulations shown in Figure Q20-1 use projections ofCO2, CH4, and N2O termed the Shared Socioeconomic Pathways(SSPs). Each SSP provides an estimate of future abundances of
GHGs based on projected emissions, which are constructed using
various assumptions of population growth and economic development, as well as technological innovation and political decisions
related to the environment. Figure Q20-1 shows the evolution ofannually averaged ozone outside the polar regions (60°S–60°N)(top) and ozone during October at 70°S–90°S (bottom), for a highclimate forcing scenario (SSP3-7.0; purple line), a medium scenario(SSP2-4.5; dark orange line), and a low scenario (SSP1-2.6; blueline). These three lines represent the multi-model mean (MMM)
projection from numerous CCMs.
These simulations show that the future recovery of the ozone layer
outside of the polar regions will be governed mostly by GHGs, assuming continued adherence to the Montreal Protocol. The wide
range of possible future levels of CO2, CH4, and N2O is an important
limitation to providing accurate future projections of ozone globally and for the ozone hole (Figure Q20-1), as well as in other geographic regions. For 60°S–60°N, total ozone recovers to the 1980
level more rapidly under the high climate forcing scenario because
large future increases in both CO2 and CH4 tend to increase ozone.
Under the low climate forcing scenario, the MMM of the CCMs
projects that ozone over 60°S–60°N may not reach the 1980 levelby the end of this century. In this low climate forcing simulation, future declines in total ozone driven by rising N2O outweigh the small
future ozone increases caused by CO2 and CH4. For the MMM of
the medium climate forcing scenario (dark orange line in Figure
Q20-1), total ozone over 60°S–60°N is projected to return to
the 1980 level around year 2040.
The one-standard deviation range of computed values of ozone
from these CCMs for all three scenarios is shown by the orange
shaded region in Figure Q20-1 (see caption), which depicts
the uncertainty in the model projections of ozone for both regions. There is a considerable range of model projections for total ozone at the end of the century for 60°S–60°N, because of
sensitivity to the actual future abundance of GHGs for the three
scenarios considered here, as well as variations among CCMs
regarding how the stratospheric circulation will actually respondto a specific GHG scenario. The CCM simulations in Figure
Q20-1 show that for the Antarctic ozone hole (October,
70°S–90°S) the future evolution of total ozone is largely governed
by ODSs, with modest sensitivity to GHGs. Here, total ozone isprojected to return to the 1980 level in 2066 for both the low andmedium climate forcing scenarios (blue and dark orange lines,
Figure Q20-1). Under the high climate forcing scenario, a more
rapid recovery of October total ozone for 70°S–90°S is projected, with a return to the 1980 level at mid-century. Because of the
impact of GHGs, these recovery years for the return of Antarctic
total ozone to the 1980 level precede 2078, which is the year EESC
is projected to return to its 1980 level for the polar stratosphere.
The multi-model mean of CCM simulations conducted for Northern Hemisphere (NH) midlatitudes (35°N–60°N) and Southern
Hemisphere (SH) midlatitudes (35°S–60°S) (results not shown)project a return of total ozone to the 1980 level around years 2035
and 2045, respectively, for the medium climate forcing scenario.
The faster return of total ozone to the 1980 level forecast for NH
68 Section VI: Stratospheric ozone in the future

midlatitudes under this scenario is caused by a greater sensitivity ofozone to future abundances of GHGs in the NH compared to theSH, as explained below. The return to the 1980 level of total ozone
for both NH and SH midlatitudes occurs considerably sooner thanyear 2061, when EESC is projected to return to the 1980 level.
Finally, the MMM of CCM simulations for the tropics (20°S–20°N)(results not shown) under the medium climate forcing scenario indicates that total ozone will remain below the 1980 level until theend of this century.
Figure Q20-2 illustrates, for six geographic regions, the combined effect on ozone of future changes in ODSs, the GHGs CO2,
CH4, and N2O (white line marked “Sum of all gases”) as well as theindividual effect on ozone of each GHG and all ODSs. These effects on ozone are determined by varying the abundance of eachquantity individually, while holding the abundances of the others
constant at their level in 1960. All of the results shown in Figure
Q20-2 are for the medium climate forcing SSP2-4.5 scenario andrely on simulations from a single CCM. As such, the timing of therecovery of ozone to the 1980 level differs slightly from the values
stated above, which are based on the MMM of simulations fromnumerous CCMs. In Figure Q20-2 the change of ozone is shown
relative to the 1960 level to more clearly display the results prior
to 1980.
The breakdown of individual drivers of future ozone for these six
selected regions shown in Figure Q20-2 is an illustrative exampleof the complexity of projecting the future recovery of stratospheric ozone, as the sensitivity to various factors differs markedly between regions.
Future changes in total ozone for the various regions are describedwith respect to a return to the level in year 1960 (see figure caption):• Antarctic. Total ozone changes are largest in the Antarctic region in springtime (October). Chemistry-climate models showthat ODSs are the predominant factor in Antarctic ozone depletion in the past and in the coming decades. The future increaseof CO2 and CH4 in this scenario acts to increase ozone, while
the future increase in N2O decreases ozone. All of these impacts on ozone are smaller than the present impact of ODSs,
and the effects of CO2, CH4, and N2O nearly cancel one another later in the century. As a result, changes in total ozone when
all forcings are considered (white line) mostly follow the change
driven by ODSs (dark pink shaded region). Although meteorological variability in the Antarctic in late winter/early springwhen ozone depletion occurs causes a substantial range in theobservations and model projections (see Figure Q20-1), Antarctic total ozone is projected to remain below the 1960 level
throughout the rest of this century.
• Arctic. Total ozone changes in the Arctic during spring (March)
are considerably smaller than in the Antarctic (see Q11). After
midcentury, Arctic total ozone increases to values above those
expected from future reductions in ODSs alone because of
the strengthening of atmospheric circulation and enhanced
stratospheric cooling associated with increases in GHGs suchas CO2. For this medium climate forcing scenario, Arctic spring
total ozone is projected to exceed the 1960 level after around2065.

2022 Update | Twenty Questions Q20
Total Ozone Changes in Response to
Ozone-Depleting Substances and Greenhouse Gases

0

450
420

–10%
70°N – 90°N

390

320

10%

290

0

260

60°S – 60°N

–10%

Ozone (DU)480

Northern Midlatitudes

Global10%

Ozone (DU)

Ozone (DU)

Arctic (Spring)1980 2020 2060 2100

370

Ozone (DU)

350

0

290

10%

260

0

230

20°S – 20°N –10%

1980 2020 2060 2100

320

Year

–10%

290

–20%

260
–30%

230

–40%

200

Southern Midlatitudes
Change in ozone (in Dobson units) induced
by individual greenhouse gases and ODSs
CH4

ODSs

CO2

N2O

Sum of all gases

90°S – 70°S –50%

170
1980

2020

2060

Ozone (DU)

Ozone (DU)

Antarctic (Spring)
10%

1980 2020 2060 2100

Tropics

Year

380

35°N – 60°N –10%

Year

Year

1980 2020 2060 2100

0

340
310

10%

350

10%

320

0

290

60°S – 35°S –10%

1980 2020 2060 2100

Year

2100

Year
Figure Q20-2. Total ozone changes in response to greenhouse gases and ODSs. Results from model simulations of a single CCM show
the sensitivity of total ozone to individual changes in the abundances of the greenhouse gases CO2, CH4, and N2O as well as to ODSs starting in 1960. These total ozone changes (in Dobson units, DU) are shown for spring in the polar regions (March for the Arctic and October
for the Antarctic) and for the annual average for the Tropics, for both the northern and southern midlatitudes, and for the global averageoutside of the polar regions (“Global”). The solid shaded regions illustrate how each particular gas (or family of gases for ODSs) affects
total ozone, with negative contributions shown below the 1960 level of ozone and positive contributions shown above. These simulations
use CO2, CH4, and N2O from the medium climate forcing (SSP2-4.5) scenario of Figure Q20-1. The abundances of the greenhouse gases
and ODSs are varied individually while the others are kept constant at their level in 1960 to compute the individual constituent’s contribution to prior and projected total ozone. Increasing CO2 results in lower temperatures throughout the stratosphere. Lower temperatures
slow down the rate of most ozone destruction chemical reactions, particularly those that regulate the abundance of ozone outside ofpolar regions, leading to an increase in ozone. Increasing CO2 also enhances the transport of ozone from the tropics to middle and highlatitudes by stratospheric winds. Increasing CH4 leads to changes in stratospheric chemistry that augment the increase in ozone driven bystratospheric cooling. Conversely, the expected future increase in N2O results in significant decreases in ozone due to chemical effects.
As the atmospheric abundance of ODSs declines towards the end of this century, ozone reductions due to increasing N2O may become
more prominent. The white “Sum of all gases” line on each panel represents the combined effect on ozone of CO2, CH4, N2O, and ODSs.
(These projections use abundances for CO2, CH4, N2O from the SSP2-4.5 scenario and ODSs from the baseline (A1) projection of the 2018 Scientific Assessment of Ozone Depletion report.)• Northern and Southern Midlatitudes. Changes in the annual averages of total ozone in midlatitudes are much smaller than
the springtime losses in polar regions. In the northern midlatitudes, the all-forcings simulation (white line) indicates a returnof total ozone to the 1960 value around 2045, even though
EESC remains above its 1960 value until the end of the century.

In the southern midlatitudes, total ozone returns to the 1960
level around 2085, about four decades later than projected
for the Northern Hemisphere. The maximum ozone depletion
near year 2000 is much larger for the Southern Hemisphere,
and total ozone more closely follows the depletion driven
by ODSs. This behavior reflects the influence of the Antarctic
Section VI: Stratospheric ozone in the future 69

Q20 Twenty Questions | 2022 Update
ozone hole on the southern midlatitudes, driven by the transport of ozone-depleted air following the breakup of the polar
vortex in late spring (see Q10). The larger value of midlatitude
total ozone from mid-century to the end of this century when
all forcings are considered (white lines) compared to ozone
depletion due only to ODSs (dark pink shaded region) reflects
the influence of changes in stratospheric circulation and upper
stratospheric temperatures by CO2 and CH4 for this medium climate forcing scenario, especially in the NH.
• Tropics. Total ozone changes in the tropics are smaller than in
any other region. Ozone is less sensitive to ODSs in the tropicalstratosphere because of the dominant roles of production andtransport in controlling ozone and the low amounts of reactive
halogens available in this region (see Q12). Total ozone gradually increases until around 2060, and then remains fairly constant and slightly below the 1960 level until the end of the century. The near-constant value of tropical ozone during the latterpart of the century is primarily due to climate-change-induced
strengthening of the stratospheric circulation, which leads toenhanced transport of ozone out of the tropics and into midlatitudes. This circulation change also influences the Arctic andmidlatitude regions, as noted above.
• The globe. The annual average of global (60°S to 60°N) total ozone is projected to return to the 1960 level around 2075,
even though EESC remains above its 1960 value until the endof the century. Chemistry-climate model analysis suggests thatthe early return of total ozone relative to EESC, as well as theend-of-century rise, are primarily a result of upper stratospheric cooling and strengthening of the stratospheric circulation
caused by increasing GHGs. Towards the latter part of this century, increased abundances of N2O may lead to more ozone depletion (dark blue shaded region) than ODSs (dark pink shaded
region).

Future ultraviolet radiation. Projections of long-term changes
in total ozone can be used to estimate long-term changes in solar
ultraviolet (UV) radiation reaching Earth’s surface (see Q16). The
UV-B component of ultraviolet radiation (see Q2) decreases as total ozone increases. Based on the ozone increases in the chemistry-climate model projections, clear-sky (cloud-free) UV-B radiation
is expected to be below the 1960 value by the end of the century
across much of the globe in regions that are not currently affectedby high levels of tropospheric air pollution. Clear-sky UV-B radiation values are expected to remain high in Antarctica, where totalozone is expected to remain below its 1960 value even at the endof this century.

Volcanoes, wildfires, and climate intervention. There are several factors not included in chemistry-climate model projections

70 Section VI: Stratospheric ozone in the futureshown above that have the potential to affect future amounts of total ozone. For example, explosive volcanic eruptions have temporarily reduced global total ozone in the past (see Q13) by enhancing the stratospheric sulfate aerosol layer. Large volcanic eruptions
that occur while EESC values remain high are expected to reduce
total ozone for a few years. Such volcanic eruptions are an additional source of uncertainty that is not included in the ozone projections in Figures Q20-1 and Q20-2. Extreme wildfires can generate
deep thunderstorms, termed pyrocumulonimbus (pyroCb) clouds,
that inject particles from biomass burning into the stratosphere.
The biomass burning particles consist of organic carbon, inorganic components and a significant fraction of black carbon thatenhance the absorption of solar radiation. These pyroCb particles
likely impact the ozone layer and climate in a manner that differs
substantially from the impact of sulfate aerosols.
Numerous climate intervention (also called geoengineering) methods have been proposed to reduce climate forcing from humanactivities. A widely discussed method is the intentional enhancement of sulfate aerosols in the stratosphere from direct injectionsof sulfur-containing substances, known as Stratospheric Aerosol
Injection (SAI). With sufficient enhancement, the added aerosols
would cool the Earth’s surface through increased reflection ofsunlight back to space, similar to the effect observed after volcanic eruptions that inject sulfate into the stratosphere (see Figure
Q13-1). While SAI could reduce some of the impacts of global
warming, it cannot restore past climatic conditions and may cause
unintended side effects, including changes in the concentrationof stratospheric ozone, a delay in the recovery of the ozone hole,and changes to the atmospheric circulation. Aerosol heating of thelowermost stratosphere by SAI using sulfates could result in further
residual impacts, including changes in regional surface temperature and precipitation patterns. How much aerosol injection would
be needed and the duration of that injection would depend on thedesired climate outcomes and the trajectory of greenhouse gas
emissions and decarbonization efforts. The sign of stratosphericozone changes will depend on the details of the SAI implementation as well as the concentrations of EESC and N2O at the time of
injection. Changes in stratospheric water vapor due to variations intemperature and circulation resulting from SAI may also play a rolein governing how stratospheric ozone would be affected by SAI.
Studies are currently being conducted to determine materials for
SAI other than sulfate that have chemical and radiative properties
that would reduce the effect of SAI on ozone and stratosphericheating. These studies are just beginning; significant laboratory
and modeling work on materials other than sulfate is needed before their effects on ozone and stratospheric transport can be reliably estimated.

Terminology | Twenty Questions | 2022 Update

TERMINOLOGY
CCM

chemistry-climate model

CFC

chlorofluorocarbon, a group of industrial compounds that contains at least one chlorine, fluorine, and carbon atom

CFC-11-equivalenta unit for the measure of the mass of emission of an ODS, equal to the product of the actual mass emission of the
ODS times its ODP

CO2-equivalenta unit for the measure of the mass of emission of a GHG, equal to the product of the actual mass emission of the
GHG times its GWP

DU

Dobson unit, a measure of total column ozone; 1 DU = 2.687 ×1016 molecules/cm2

EEAP

Environmental Effects Assessment Panel of the Montreal Protocol

EESC

equivalent effective stratospheric chlorine, a measure of the total amount of reactive and reservoir chlorine andbromine gases in the stratosphere that is available to deplete stratospheric ozone

GHG

greenhouse gas

gigatonne

1 billion (109) metric tons = 1 trillion (1012) kilograms

GWP

global warming potential, a measure of the effectiveness of the emission of a gas to cause an increase in the radiative forcing of climate, relative to the radiative forcing caused by the emission of the same mass of CO2; all GWPs
used here are for a 100-year time interval

halogenelements fluorine, chlorine, bromine, iodine, and astatine that form group 7A of the periodic tablehalon

a group of industrial compounds that contain at least one bromine and carbon atom; may or may not contain achlorine atom

HCFChydrochlorofluorocarbon, a group of industrial compounds that contain at least one hydrogen, chlorine, fluorine,and carbon atom

HFChydrofluorocarbon, a group of industrial compounds that contain at least one hydrogen, fluorine, and carbon atomand no chlorine or bromine atoms

HFO

hydrofluoroolefin, a group of industrial compounds that contain at least one hydrogen, fluorine, and carbon atomand no chlorine or bromine atoms, and also include a double carbon bond that causes these gases to be morereactive in the troposphere than other HFCs

IPCC

Intergovernmental Panel on Climate Changekilotonne

1000 metric tons = 1 million (106) kilograms

megatonne

1 million (106) metric tons = 1 billion (109) kilograms

MMM

multi-model meanmPa

pascal (Pa) is a unit of pressure; 1 mPa (millipascal) equals 0.001 Pa.

nm

nanometer, one billionth of a meter (10–9 m)

ODP

ozone depletion potential, a measure of the effectiveness of the emission of a gas to deplete the ozone layer, relative to the ozone depletion caused by the emission of the same mass of CFC-11

ODS

ozone-depleting substanceozone layer

the region in the stratosphere with the highest concentration of ozone, between about 15 and 35 km altitudeppb

parts per billion; 1 part per billion equals the presence of one molecule of a gas per billion (109) total air molecules

ppm

parts per million; 1 part per million equals the presence of one molecule of a gas per million (10 6) total air molecules

ppt

parts per trillion; 1 part per trillion equals the presence of one molecule of a gas per trillion (1012) total air molecules

PFC

perfluorocarbon, a group of industrial compounds that contain only carbon and fluorine atoms

PSC

polar stratospheric cloudRF

radiative forcing of climate

72 Terminology

2022 Update | Twenty Questions | Terminology

SAI

stratospheric aerosol injection

SAOD

stratospheric aerosol optical depthsource gas

a halogen-containing chemical with a sufficiently long atmospheric lifetime that its emissions contribute to
the stratospheric abundance of chlorine or bromine molecules, which participate in chemical reactions that
deplete ozone

SAP

Scientific Assessment Panel of the Montreal Protocol

SSP

Shared Socioeconomic Pathway

stratospherelayer of the atmosphere above the troposphere that extends up to around 50 km altitude, and that includes
the ozone layer

TEAP

Technology and Economic Assessment Panel of the Montreal Protocol

TFA

trifluoroacetic acid, a decomposition product of HFCs, HCFCs, and HFOs

troposphere

lower layer of the atmosphere that extends from the surface up to about 10–15 km (6 –9 miles) altitude

UNEP

United Nations Environment Programme

UNFCCC

United Nations Framework Convention on Climate Change

UVultraviolet radiation

UV-A

ultraviolet radiation between wavelengths of 315 and 400 nm

UV-B

ultraviolet radiation between wavelengths of 280 and 315 nm

UV-C

ultraviolet radiation between wavelengths of 100 and 280 nm

UVI

ultraviolet radiation index

VSLvery short-lived

WMO

World Meteorological Organization

Terminology 73

Chemical Formulae | Twenty Questions | 2022 Update

CHEMICAL FORMULAE AND NOMENCLATURE
Chlorine Compounds
Cl
ClO(ClO)2
ClONO2HCl
CCl4
CH2Cl2CH3CCl3
CH3Cl
CCl3F
CCl2F2
CCl2FCClF2
CHF2ClCH3CCl2F
CH3CClF2

atomic chlorine
chlorine monoxide
chlorine monoxide dimer, chemical
structure ClOOCl
chlorine nitratehydrogen chloride (hydrochloric acid)carbon tetrachloride
dichloromethane
methyl chloroform
methyl chloride
CFC-11
CFC-12
CFC-113
HCFC-22
HCFC-141b
HCFC-142b

Bromine Compounds
Br
BrO
BrCl
CH2Br2
CHBr3
CH3Br
CBrClF2
CBrF3
CBrF2CBrF2

Other Gases

Other Halogens
CHF3
CH2F2
CHF2CF3
CH2FCF3
CH3CF3
CH3CHF2
CF3CFCH2
CF4
C2F6
CF3I
C2HF3O2
IO
SF6

HFC-23
HFC-32
HFC-125
HFC-134a
HFC-143a
HFC-152a
HFO-1234yf
carbon tetrafluoride
perfluoroethane
trifluoroiodomethane
trifluoroacetic acidiodine monoxide
sulfur hexafluoride

74 Chemical formulae and nomenclature

atomic bromine
bromine monoxide
bromine monochloride
dibromomethane
bromoform
methyl bromide
halon-1211
halon-1301
halon-2402

CH4
CO
CO2
H
H2O
HNO3
H2SO4
N2
N2O
NOx
O
O2
O3

methanecarbon monoxidecarbon dioxide
atomic hydrogenwater vapor
nitric acidsulfuric acidmolecular nitrogen
nitrous oxide
nitrogen oxides
atomic oxygen
molecular oxygenozone

2022 Update | Twenty Questions | Acknowledgements

AUTHORS AND CONTRIBUTORS
Co-Chairs of the Scientific Assessment Panel (SAP) of the Montreal Protocol
and Assessment Co-Chairs
David W. Fahey
Paul A. Newman
John A. Pyle
Bonfils Safari

NOAA Chemical Sciences Laboratory
NASA Goddard Space Flight Center
University of Cambridge and the National Centre for Atmospheric Science (NCAS)
University of Rwanda, College of Science and Technology

USA
USA
UK
RwandaLead Author
Ross J. Salawitch

University of Maryland, College Park

USA

Authors
Laura A. McBride
Chelsea R. Thompson
Eric L. Fleming
Richard L. McKenzie
Karen H. Rosenlof
Sarah J. Doherty
David W. Fahey

Albright College
NOAA Chemical Sciences Laboratory
Science Systems and Applications, Inc. (SSAI) at NASA Goddard Space Flight Center
National Institute of Water and Atmospheric Research (NIWA)NOAA Chemical Sciences Laboratory
University of Colorado, Cooperative Institute for Research in Environmental Sciences (CIRES)
at NOAA Chemical Sciences Laboratory
NOAA Chemical Sciences Laboratory

USA
USA
USA
New Zealand
USA
USA
USA

Contributing Authors
Simon Alexander
Germar Bernhard
Lucy Carpenter
Gabriel Chiodo
Robert Damadeo
John S. DanielSean Davis
Vitali Fioletov
Hella Garny
Brad Hall
Birgit Hassler
James Haywood
Lei HuYue Jia
Bryan Johnson
James Keeble
Johannes C. Laube
Eric Nash
Stefan Reimann
Matt Rigby
Michelle L. Santee
Susan Tegtmeier
Simon Tilmes
Walter Tribett
Guus J. M. Velders
Peter von der Gathen
Krzysztof Wargan
Mark Weber
Paul Young

Australian Antarctic Division
Biospherical Instruments, Inc.
University of YorkETH Zürich, Institute for Atmospheric and Climate ScienceNASA Langley Research Center
NOAA Chemical Sciences Laboratory
NOAA Chemical Sciences Laboratory
Environment and Climate Change Canada
Deutsches Zentrum für Luft und Raumfahrt (DLR), Institut für Physik der Atmosphäre (IPA)
NOAA Global Monitoring Laboratory
Deutsches Zentrum für Luft und Raumfahrt (DLR), Institut für Physik der Atmosphäre (IPA)
University of Exeter and Met Office Hadley Centre
University of Colorado, Cooperative Institute for Research in Environmental Sciences (CIRES)
at NOAA Global Monitoring Laboratory
University of Colorado, Cooperative Institute for Research in Environmental Sciences (CIRES)
at NOAA Chemical Sciences Laboratory
NOAA Global Monitoring Laboratory
University of Cambridge, Department of Chemistry
Forschungszentrum Jülich, Institute for Energy and Climate Research: Stratosphere (IEK-7)
Science Systems and Applications, Inc. (SSAI) at NASA Goddard Space Flight Center
Swiss Federal Laboratories for Materials Science and Technology (Empa)
University of Bristol, School of Chemistry
NASA Jet Propulsion Laboratory, California Institute of Technology
University of Saskatchewan, Institute of Space and Atmospheric Studies
National Center for Atmospheric Research (NCAR), Atmospheric Chemistry Observations & Modeling
University of Maryland, College Park
National Institute for Public Health and the Environment (RIVM) & Utrecht University
Alfred Wegener Institute, Helmholtz Centre for Polar and Marine Research
Science Systems and Applications Inc. (SSAI) at NASA Goddard Space Flight Center
Universität Bremen, Institute of Environmental Physics
Lancaster University

Australia
USA
UK
Switzerland
USA
USA
USA
CanadaGermanyUSA
Germany
UK
USA
USA
USA
UK
Germany
USA
Switzerland
UK
USA
CanadaUSA
USA
Netherlands
GermanyUSA
Germany
UK

Authors and contributors 75

The Twenty Questions and Answers About the Ozone Layer: 2022 Update is a componentof the Scientific Assessment of Ozone Depletion: 2022 report. The report is prepared
quadrennially by the Scientific Assessment Panel (SAP) of the Montreal Protocol on
Substances that Deplete the Ozone Layer. The motivation behind this publication is totell the story of ozone depletion, ozone-depleting substances, and the success of theMontreal Protocol. This 2022 Update coincides with the 35th anniversary of the signingof the Montreal Protocol in 1987.

Lead Author
Ross J. Salawitch
Coauthors
Laura A. McBride
Chelsea R. Thompson
Eric L. Fleming
Richard L. McKenzie
Karen H. Rosenlof
Sarah J. Doherty
David W. Fahey

World Meteorological Organization
United Nations Environment Programme
National Oceanic and Atmospheric Administration
National Aeronautics and Space Administration
European Commission