Twenty Questions andAnswers About the Ozone
Layer: 2014 Update
Scientific Assessment of Ozone Depletion: 2014

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

This publication, Twenty Questions and Answers About the Ozone Layer:
2014 Update, is a component of the 2014 report of the Montreal Protocol
Scientific Assessment Panel, Scientific Assessment of Ozone Depletion:
2014. This document is an update of Twenty Questions and Answers About
the Ozone Layer: 2010 (Fahey, D.W., and M.I. Hegglin) and reflects thescientific information presented in the 2014 Assessment Report.
It is a continuation of the updates of the first document with this title
scope: Twenty Questions and Answers About the Ozone Layer: 2002
(Fahey, D.W.).
Lead Author
Michaela I. Hegglin
University of Reading, UK
Coauthors
David W. Fahey
NOAA ESRL Chemical Sciences
Division, USA
Mack McFarland
Dupont Chemicals &Fluoroproducts, USA
Stephen A. Montzka
NOAA ESRL Global Monitoring
Division, USA
Eric R. Nash
NASA Goddard Space Flight
Center, USA
Contributors

Stefan Reimann
Swiss Federal Laboratories forMaterials Science and Technology,
Switzerland

John P. Burrows
University of Bremen, Germany

Matthew Rigby
University of Bristol, UK

Martin Dameris
DLR, Institut für Physik der
Atmosphäre, Germany

Michelle L. Santee
Jet Propulsion Laboratory, California
Institute of Technology, USA
Guus J.M. Velders
RIVM (National Institute for Public
Health and the Environment), The
Netherlands
Darryn W. Waugh
Johns Hopkins University, USA
Mark Weber
University of Bremen, Germany

Pablo O. Canziani
Pontificia Universidad Católica
– CONICET, Argentina

The authors would like toacknowledge in addition the many
contributors to previous versionsof this document.

Lucy J. Carpenter
University of York, UK

Reviewers

John S. Daniel
NOAA ESRL Chemical Sciences
Division, USA
Nada Derek
Commonwealth Scientific andIndustrial Research Organisation,
Australia
Vitali E. Fioletov
Environment Canada, CanadaPeter von der Gathen
Alfred Wegener Institute, Germany
Bradley D. HallNOAA ESRL Global Monitoring
Division, USA
Bryan Johnson
NOAA ESRL Global Monitoring
Division, USA
Kathy Lantz
NOAA ESRL Global Monitoring
Division / CIRES, USA
Zachary Lawrence
New Mexico Institute of Mining
and Technology, USA
Gloria L. Manney
NWRA / New Mexico Instituteof Mining and Technology, USA
Olaf Morgenstern
National Institute of Water
and Atmospheric Research,
New ZealandPaul A. Newman
NASA Goddard Space Flight
Center, USA

Special Recognition
Malcolm K.W. Ko
NASA Langley Research Center,
USA

Christine A. Ennis
CIRES / NOAA ESRL Chemical
Sciences Division, USA
Lucien Froidevaux
Jet Propulsion Laboratory, California
Institute of Technology, USA
Nathan P. Gillett
Environment Canada, Canada
Sophie Godin-Beekmann
Centre National de la Recherche
Scientifique, France
Marco Gonzalez
Sistemas de Energia, Costa Rica
Neil R.P. Harris
University of Cambridge, UK
Kenneth W. Jucks
NASA Headquarters, USA
David J. Karoly
Melbourne University, Australia

Michelle L. Santee
Jet Propulsion Laboratory, California
Institute of Technology, USA

Lambert J.M. Kuijpers
Technical University Eindhoven,
Netherlands

Ross J. Salawitch
University of Maryland, USA

Michael J. Kurylo
Universities Space Research
Association / Goddard Earth
Sciences, Technology, andResearch, USA

Rajendra Shende
TERRE Policy Centre, India
Theodore G. Shepherd
University of Reading, UK
Donald J. Wuebbles
University of Illinois, USA
In addition, this document was
discussed by other scientists who
attended the Panel Review Meetingfor the 2014 Assessment report in
Les Diablerets in June 2014:
Ayité-Lô Ajavon
North-South Environment, Togo
Stephen O. Andersen
Institute for Governance &Sustainable Development, USA
Julie M. Arblaster
Australian Bureau of Meteorology,
Australia
Pieter Aucamp
Ptersa, South Africa
Gufran Beig
Indian Institute of TropicalMeteorology, India
Geir O. Braathen
World Meteorological Organization,
Switzerland

2

Martyn P. Chipperfield
University of Leeds, UK

Section II: The ozone depletion process

Ulrike Langematz
Freie Universität Berlin, Germany
Martin R. Manning
New Zealand Climate Change
Research Institute, New Zealand
Amanda C. Maycock
University of Cambridge, UK
Pauline Midgley
University of Bern, Switzerland
Rolf Müller
Forschungszentrum Jülich,
Germany
Hiroaki Naoe
Japan Meteorological Agency,
Japan
Steven Pawson
NASA Goddard Space Flight
Center, USA
Thomas PeterETH Zurich, Switzerland
David Plummer
Environment Canada, Canada
Lorenzo Polvani
Columbia University, USA

Ronald G. Prinn
Massachusetts Institute ofTechnology, USA
John A. Pyle
National Centre for Atmospheric
Science, UK and University ofCambridge
A.R. Ravishankara
Colorado State University / NOAAESRL Chemical Sciences Division,
USA
Jose M. Rodriguez
NASA Goddard Space Flight
Center, USA
Megumi Seki
UNEP Ozone Secretariat, Kenya
Susan Solomon
Massachusetts Institute ofTechnology, USA
Wolfgang Steinbrecht
Deutscher Wetterdienst
Hohenpeissenberg, Germany
William T. Sturges
University of East Anglia, UK
David W.J. Thompson
Colorado State University, USA
Timothy J. Wallington
Ford Motor Company, USA
Ray F. Weiss
University of California San Diego /Scripps, USA
Shigeo Yoden
Kyoto University, Japan
Lingxi Zhou
Chinese Meteorological
Administration, China
Coordinating Editor
Christine A. Ennis
CIRES / NOAA ESRL Chemical
Sciences Division, USA
Technical Management Assistance
Debra A. Dailey-Fisher
NOAA ESRL Chemical Sciences
Division, USA
Graphic Art Design
Dennis Dickerson
Respond Grafiks, USA
Publication Layout and Design
Nayna Morjaria
Design & Print Studio,
University of Reading, UK

Twenty Questions andAnswers About the Ozone
Layer: 2014 Update
Scientific Assessment of Ozone Depletion: 2014

This document answers some of the most commonly
asked questions about the ozone layer.
Lead Author
Michaela I. Hegglin

Coauthors
David W. Fahey
Mack McFarland
Stephen A. Montzka
Eric R. Nash

Contributors
Pablo O. Canziani, Lucy J. Carpenter, John S. Daniel, Nada Derek,
Vitali E. Fioletov, Peter von der Gathen, Bradley D. Hall, Bryan
Johnson, Kathy Lantz, Zachary Lawrence, Gloria L. Manney, Olaf
Morgenstern, Paul A. Newman, Stefan Reimann, Matthew Rigby,
Michelle L. Santee, Guus J.M. Velders, Darryn W. Waugh, Mark Weber

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

Michaela I. Hegglin (Lead Author), David W. Fahey, Mack McFarland,
Stephen A. Montzka, and Eric R. Nash, Twenty Questions and Answers
About the Ozone Layer: 2014 Update, Scientific Assessment of Ozone
Depletion: 2014, 84 pp., World Meteorological Organization, Geneva,
Switzerland, 2015.
Published in January 2015
ISBN: 978-9966-076-02-1
2

Section I: Ozone in our atmosphere

Contents
Introduction 5
Section I: Ozone in our atmosphere
Q1: What is ozone and where is it in the atmosphere? 8
Q2: How is ozone formed in the atmosphere? 10
Q3: Why do we care about atmospheric ozone? 12
Q4: How is total ozone distributed over the globe? 14
Q5: How is ozone measured in the atmosphere? 17

Section II: The ozone depletion process
Q6:	How do emissions of halogen source gases lead to stratospheric ozone
depletion? 19
Q7:

What emissions from human activities lead to ozone depletion? 22

Q8: What are the reactive halogen gases that destroy stratospheric ozone? 27
Q9: What are the chlorine and bromine reactions that destroy stratospheric ozone? 31
Q10:	Why has an “ozone hole” appeared over Antarctica when ozone-depleting
substances are present throughout the stratosphere? 33

Section III: Stratospheric ozone depletion
Q11: How severe is the depletion of the Antarctic ozone layer? 37
Q12: Is there depletion of the Arctic ozone layer? 42
Q13: How large is the depletion of the global ozone layer? 46
Q14: Do changes in the Sun and volcanic eruptions affect the ozone layer? 49

Section IV: Controlling ozone-depleting substances
Q15: Are there controls on the production of ozone-depleting substances? 52
Q16:	Has the Montreal Protocol been successful in reducing ozone-depleting
substances in the atmosphere? 55

Section V: I mplications of ozone depletionand the Montreal Protocol
Q17: Does depletion of the ozone layer increase ground-level ultraviolet radiation? 59
Q18: Is depletion of the ozone layer the principal cause of global climate change? 62
Q19:	Have Montreal Protocol controls of ozone-depleting substances also protected
Earth’s climate? 68

Section VI: Stratospheric ozone in the future
Q20: How is ozone expected to change in the coming decades? 74

Additional Topics
Global Ozone Network 18
Understanding Stratospheric Ozone Depletion 21
The Discovery of the Antarctic Ozone Hole 36
The 2002 Antarctic Ozone Hole 41
Initial Signs of Ozone Recovery 48
The Antarctic Ozone Hole and Southern Hemisphere Surface Climate 67
3

Introduction
Ozone is present only in small amounts in Earth’s atmosphere.
Nevertheless, it is vital to human well-being and ecosystem
health. Most ozone resides in the upper part of the atmosphere.
This region, called the stratosphere, is more than 10 kilometers(6 miles) above Earth’s surface. There, about 90% of atmosphericozone is contained in the “ozone layer,” which shields us fromharmful ultraviolet radiation from the Sun.
In the mid-1970s, it was discovered that some human-produced chemicals could
lead to depletion of the ozone layer. The resulting increase in ultraviolet radiation at
Earth’s surface would likely increase the incidences of skin cancer and eye cataracts,
and also adversely affect plants, crops, and ocean plankton.
Following the discovery of this environmental issue, researchers sought a better
understanding of this threat to the ozone layer. Monitoring stations showed that theabundances of the ozone-depleting substances (ODSs) were steadily increasing in theatmosphere. These trends were linked to growing production and use of chemicals
like chlorofluorocarbons (CFCs) for refrigeration and air conditioning, foam blowing,
and industrial cleaning. Measurements in the laboratory and in the atmospherecharacterized the chemical reactions that were involved in ozone destruction.
Computer models of the atmosphere employing this information were used tosimulate how much ozone depletion was occurring and to predict how much moremight occur in the future.
Observations of the ozone layer showed that depletion was indeed occurring. The
most severe and most surprising ozone loss was discovered to be recurring in
springtime over Antarctica. The loss in this region is commonly called the “ozonehole” because the ozone depletion is so large and localized. A thinning of the ozone
layer also has been observed over other regions of the globe, such as the Arctic andnorthern and southern midlatitudes.
The work of many scientists throughout the world has provided a basis for building
a broad and solid scientific understanding of the ozone depletion process. With thisunderstanding, we know that ozone depletion is indeed occurring and why. Most
important, we know that if the most potent ODSs were to continue to be emitted andincrease in the atmosphere, the result would be more depletion of the ozone layer.
In response to the prospect of increasing ozone depletion, the governments of theworld crafted the 1987 United Nations Montreal Protocol as an international means
to address this global issue. As a result of the broad compliance with the Protocol andits Amendments and adjustments and, of great significance, industry’s developmentof “ozone-friendly” substitutes for the now-controlled chemicals, the total global
accumulation of ODSs has slowed and begun to decrease. In response, global ozone
depletion is no longer increasing, and initial signs of recovery of the ozone layer havebeen identified. With continued compliance, we can expect substantial recoveryof the ozone layer by the middle of the 21st century. The day the Montreal Protocol
was agreed upon, 16 September, is now celebrated as the International Day for the
Preservation of the Ozone Layer.
This is a story of notable achievements: discovery, understanding, decisions,actions, and verification. It is a story written by many: scientists, technologists,
economists, legal experts, and policymakers, in which continuous dialogue has beena key ingredient. A timeline of milestones related to the science of stratosphericozone depletion, international scientific assessments, and the Montreal Protocol is
illustrated in Figure Q0-1.

Introduction

5

To help maintain a broad understanding of the relationship between ozone
depletion, ODSs, and the Montreal Protocol, this component of the Scientific
Assessment of Ozone Depletion: 2014 presents 20 questions and answers about
the often-complex science of ozone depletion. Most questions and answers
are updates of those presented in previous Ozone Assessments, and a few
have been expanded to address newly emerging issues. The questions address
the nature of atmospheric ozone, the chemicals that cause ozone depletion,
how global and polar ozone depletion occur, the extent of ozone depletion, thesuccess of the Montreal Protocol, and the possible future of the ozone layer.
Computer models project that the influence on global ozone of greenhouse gases
and changes in climate will grow significantly in the coming decades and exceed
the importance of ODSs in most atmospheric regions by the end of this century.
Ozone and climate are indirectly linked because both ODSs and their substitutes
are greenhouse gases and contribute to climate change. The Antarctic ozone
hole has recently been shown also to have had a direct effect on Southern
Hemisphere surface climate during summer. A brief answer to each question is
first given in blue; an expanded answer then follows. The answers are based on
the information presented in the 2014 and earlier Assessment reports as wellas other international scientific assessments. These reports and the answers
provided here were prepared and reviewed by a large number of international
scientists who are experts in different research fields related to the science ofstratospheric ozone and climate1.

1

See inside front cover.

Figure Q0-1. Stratospheric ozone depletion milestones.
This timeline highlights milestones related to the history of ozone
depletion. Events represent the occurrence of important scientificfindings, the completion of international scientific assessments,and highlights of the Montreal Protocol. The graph shows thehistory and near future of annual total emissions of ozonedepleting substances (ODSs) combined with natural emissions of
halogen source gases. ODSs are halogen source gases controlled
under the Montreal Protocol. The emissions, when weighted bytheir potential to destroy ozone, peaked in the late 1980s after
several decades of steady increases (see Q19). Between the late1980s and the present, emissions have decreased substantiallyas a result of the Montreal Protocol and its subsequent
Amendments and adjustments coming into force (see Q15). The
Protocol began with the Vienna Convention for the Protectionof the Ozone Layer in 1985. The provisions of the Protocol andits Amendments and adjustments decisions have depended on
information embodied in international scientific assessments ofozone depletion that have been produced periodically since 1989
under the auspices of UNEP and WMO. Atmospheric observationsof ozone, CFCs, and other ODSs have increased substantially

6

Introduction

since the early 1970s. For example, the SAGE and TOMS satellite
instruments have provided essential global views of stratosphericozone for several decades. The Nobel Prize in Chemistry in 1995
was awarded for research that identified the threat to ozone
posed by CFCs and that described key reactive processes in thestratosphere. By 2008, stratospheric chlorine abundances in thestratosphere were 10% lower than their peak values reached in
the late 1990s and were continuing to decrease. January 2010
marked the end of global production of CFCs and halons under
the Protocol, and January 2013 the universal control of ODS
production and consumption. In the early 2010s, the first signsof ozone recovery were detected.
(A megatonne = 1 billion (109) kilograms.)
EESC: Equivalent effective stratospheric chlorine
ODS: Ozone-depleting substance
WMO: World Meteorological Organization
IPCC: Intergovernmental Panel on Climate Change
TEAP: Technology and Economic Assessment Panel of theMontreal Protocol
UNEP: United Nations Environment ProgrammeMilestones
theHistory
History
Stratospheric
Ozone
Depletion
Milestones ininthe
of ofStratospheric
Ozone
Depletion
Peak global emissionsof ozone-depleting
substances–1987
Ozone hole reported
over Antarctica–1985

Scientific Milestones
WMO World Ozone
Data Centre–1961

Global surface observationsof CFCs reported–1973

Stratospheric Aerosol &Gas Experiment I (SAGE-I)
instrument launched–1979

CFC ozone-depletion
hypothesis–1974

Total Ozone Mapping
Spectrometer (TOMS)
instrument launched–1978

CFCs observed in thestratosphere & reported as
potent greenhouse gases–1975

International Scientific
Assessments

WMO Atmospheric
Ozone–1985
WMO World Plan of Action
on the Ozone Layer–1977
WMO The Stratosphere 1981:
Theory and Measurements
WMO Report of the International
Ozone Trends Panel–1988

Montreal Protocol
Milestones

Upper stratospheric ozone shown to increase
as a result of the ODS phase-out–2011–2014
Peak stratospheric abundance of
EESC in midlatitudes–1997
Laboratory, field, and modeling studies
show chlorine and bromine deplete
polar stratospheric ozone–1985-1989

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

Montreal Protocol
Amendments

Vienna Convention for theProtection of the Ozone Layer–1985(Megatonnes CFC-11-equivalent per year)

Stratospheric abundance of EESC
in midlatitudes down 10%
from 1997 peak–2008

UNEP/WMO Scientific Assessments of Ozone Depletion
1989 1991 1994 1998 2002 2006 2010 2014

London–1990
Copenhagen–1992
Vienna–1995
Montreal–1997
Beijing–1999
Montreal Protocol on Substancesthat Deplete the Ozone Layer–1987

2.0

1995 Nobel Prize in Chemistry
to Crutzen, Molina and Rowland

Global productionof CFCs andhalons ends–2010
HCFC accelerated phaseout(Montreal–2007 Adjustment)

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

All ODSs under
production andconsumption
control–2013

Montreal Protocol
universal ratification(196 nations)–2009

Emissions of halogen source gases (ODSs + natural sources)(Weighted by Ozone Depletion Potential (ODP))

1.5

Based on atmosphericobservations or industry data
Future projections

1.0

0.5
Natural emissions
0
1960

1970

1980

1990

2000

2010

2020

Year

Introduction7

1

What is ozone and whereis it in the atmosphere?
Ozone is a gas that is naturally present in our atmosphere.
Each ozone molecule contains three atoms of oxygen and isdenoted chemically as O3. Ozone is found primarily in two regionsof the atmosphere. About 10% of atmospheric ozone is in the
troposphere, the region closest to Earth (from the surface to about
10–16 kilometers (6–10 miles)). The remaining ozone (about 90%)
resides in the stratosphere between the top of the troposphere andabout 50 kilometers (31 miles) altitude. The large amount of ozone
in the stratosphere is often referred to as the “ozone layer.”Ozone is a gas that is naturally present in our atmosphere. Ozone has the chemical
formula O3 because an ozone molecule contains three oxygen atoms (see FigureQ1-1). Ozone was discovered in laboratory experiments in the mid-1800s. Ozone’s
presence in the 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 detectedeven at very low amounts. Ozone reacts rapidly with many chemical compounds
and is explosive in concentrated amounts. Electrical discharges are generally used
to produce ozone for industrial processes such as air and water purification andbleaching of textiles and food products.
Ozone location. Most ozone (about 90%) is found in the stratosphere, which begins
about 10–16 kilometers (6–10 miles) above Earth’s surface and extends up to about
50 kilometers (31 miles) altitude. The stratospheric region with the highest ozone
concentration is commonly known as the “ozone layer” (see Figure Q1-2). The ozone
layer extends over the entire globe with some variation in altitude and thickness.
Most of the remaining ozone, about 10%, is found in the troposphere, which is thelowest region of the atmosphere, between Earth’s surface and the stratosphere.
Ozone abundance. Ozone molecules have a low relative abundance in theatmosphere. In the stratosphere near the peak concentration of the ozone layer,
there are typically a few thousand ozone molecules for every billion air molecules(1 billion = 1,000 million). Most air molecules are either oxygen (O2) or nitrogen (N2).
In the troposphere near Earth’s surface, ozone is even less abundant, with a typicalrange of 20 to 100 ozone molecules for each billion air molecules. The highest
surface values result when ozone is formed in air polluted by human activities.

Ozone
OxygenOzoneandand Oxygen
Oxygenatom (O)
Figure Q1-1 Ozone and oxygen.
A molecule of ozone (O3) contains threeoxygen atoms (O) bound together. Oxygen
molecules (O2), which constitute 21% of
the gases in Earth’s atmosphere, containtwo oxygen atoms bound together.

8

Section I: Ozone in our atmosphere

Oxygenmolecule (O2)Ozone
molecule (O3)

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

Ozoneinin thethe Atmosphere
Ozone
Atmosphere

35
20Stratospheric
ozone

25

Ozone Layer

15

20
10

15
10
5
0

Ozone
increasesfrom pollution

Troposphericozone

Ozone abundance (mPa)

Altitude (miles)

Figure Q1-2 Ozone in the atmosphere.
Ozone is present throughout thetroposphere 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 thislayer varies from region to region and withseason over the globe (see Q4). Increases
in ozone occur near the surface as aresult of pollution from human activities.

Altitude (kilometers)

30

5

0

Section I: Ozone in our atmosphere92How is ozone formed in the atmosphere?
Ozone is formed throughout the atmosphere in multistep
chemical processes that require sunlight. In the stratosphere, theprocess begins with an oxygen molecule (O2) being broken apartby ultraviolet radiation from the Sun. In the lower atmosphere(troposphere), ozone is formed by a different set of chemical
reactions that involve naturally occurring gases and those from
pollution sources.
Stratospheric ozone. Stratospheric ozone is formed naturally by chemical reactions
involving solar ultraviolet radiation (sunlight) and oxygen molecules, which make up
21% of the atmosphere. In the first step, solar ultraviolet radiation breaks apart one
oxygen molecule (O2) to produce two oxygen atoms (2 O) (see Figure Q2-1). In thesecond step, each of these highly reactive atoms combines with an oxygen molecule
to produce an ozone molecule (O3). These reactions occur continually whenever solar
ultraviolet radiation is present in the stratosphere. As a result, the largest ozone
production occurs in the tropical stratosphere.
The production of stratospheric ozone is balanced by its destruction in chemical
reactions. Ozone reacts continually with sunlight and a wide variety of natural andhuman-produced chemicals in the stratosphere. In each reaction, an ozone molecule
is lost and other chemical compounds are produced. Important reactive gases that
destroy ozone are hydrogen and nitrogen oxides and those containing chlorine andbromine (see Q8).
Some stratospheric ozone is regularly transported down into the troposphere andcan occasionally influence ozone amounts at Earth’s surface, particularly in remote,unpolluted regions of the globe.
Tropospheric ozone. Near Earth’s surface, ozone is produced by chemical reactions
involving gases from natural and pollution sources. Ozone production reactions
primarily involve hydrocarbon and nitrogen oxide gases, as well as ozone itself, and all

Stratospheric
Ozone
Production
Stratospheric
Ozone
Production
Step

1

Figure Q2-1. Stratospheric ozone production.
Ozone is produced naturally in the stratosphere by a two-stepreaction process. In the first step, solar ultraviolet radiation(sunlight) breaks apart an oxygen molecule to form two separate
oxygen atoms. In the second step, each atom collides withanother oxygen molecule and forms an ozone molecule in abinding reaction. In the overall process, three oxygen molecules
plus sunlight react to form two ozone molecules.

10

Section I: Ozone in our atmosphere

Ultraviolet
Sunlight

Step

2

Overall reaction: 3 O2 sunlight 2 O3require sunlight for completion. Fossil fuel combustion is a primary source of pollutant
gases that lead to tropospheric ozone production. The production of ozone near thesurface does not significantly contribute to the abundance of stratospheric ozone.
The amount of surface ozone is too small in comparison and the transport of surface
air to the stratosphere is not effective enough. As in the stratosphere, ozone in the
troposphere is destroyed by naturally occurring chemical reactions and by reactions
involving human-produced chemicals. Tropospheric ozone can also be destroyed whenozone reacts with a variety of surfaces, such as those of soils and plants.
Balance of chemical processes. Ozone abundances in the stratosphere andtroposphere are determined by the balance between chemical processes that
produce and destroy ozone. The balance is determined by the amounts of reactive
gases and how the rate or effectiveness of the various reactions varies with sunlightintensity, location in the atmosphere, temperature, and other factors. As atmosphericconditions change to favor ozone-producing reactions in a certain location, ozone
abundances increase. Similarly, if conditions change to favor other reactions that
destroy ozone, abundances decrease. The balance of production and loss reactions,combined with atmospheric air motions that transport and mix air with different
ozone abundances, determines the global distribution of ozone on timescales of days
to many months (see also Q4). Global ozone has decreased during the past several
decades (see Q13) because the amounts of reactive gases containing chlorine andbromine have increased in the stratosphere due to human activities.

Section I: Ozone in our atmosphere113

Why do we care about atmospheric ozone?
Ozone in the stratosphere absorbs a large part of the Sun’sbiologically 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 naturalamounts is considered “bad” ozone because it is harmful tohumans, plants, and animals. Natural ozone near the surface
and in the lower atmosphere plays an important beneficial rolein chemically removing pollutants from the atmosphere.
Good ozone. Stratospheric ozone is considered good for humans and other life forms
because it absorbs ultraviolet-B (UV-B) radiation from the Sun (see Figure Q3-1). If notabsorbed, UV-B radiation would reach Earth’s surface in amounts that are harmful
to a variety of life forms. In humans, increased exposure to UV-B radiation increases
the risks of skin cancer, cataracts, and a suppressed immune system. UV-B radiation
exposure before adulthood and cumulative exposure are both important health risk
factors. Excessive UV-B exposure also can damage terrestrial plant life, single-cellorganisms, and aquatic ecosystems. Other UV radiation, UV-A, which is not absorbed
significantly by ozone, causes premature aging of the skin.
Protecting good ozone. In the mid-1970s, it was discovered that gases containing
chlorine and bromine atoms released by human activities could cause stratosphericozone depletion (see Q6 and Q7). These gases, referred to as halogen source gases,
and as ozone-depleting substances (ODSs), chemically release their chlorine andbromine atoms after they reach the stratosphere. Ozone depletion increases surface
UV-B radiation above naturally occurring amounts. International efforts have beensuccessful in protecting the ozone layer through controls on ODS production andconsumption (see Q15 and Q16).
Bad ozone. Ozone near Earth’s surface in excess of natural amounts is consideredbad ozone. It is formed by reactions involving human-made pollutant gases.
Increasing surface ozone above natural levels is harmful to humans, plants, andother living systems because ozone reacts strongly to destroy or alter many biologicalmolecules. High ozone exposure caused by air pollution reduces crop yields andforest growth. In humans, exposure to high 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 in tropospheric ozone
lead to a warming of Earth’s surface because ozone is a greenhouse gas (see Q18).
The negative effects of excess tropospheric ozone contrast sharply with the protection
from harmful UV-B radiation afforded by an abundance of stratospheric ozone.
Reducing bad ozone. Limiting the emission of certain common pollutants reduces
the production of excess ozone in the air surrounding humans, plants, and animals.
Natural emissions from the biosphere, mainly from trees, also participate in reactions
that produce ozone. Major sources of pollutants include large cities where fossil fuelconsumption and industrial activities are greatest. Many programs around the globe
have already been successful 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 isa natural component of the clean atmosphere. Ozone plays important roles in theatmosphere beyond absorbing UV radiation. For example, ozone initiates the chemical
removal of many pollutants, such as carbon monoxide (CO) and nitrogen oxides

12

Section I: Ozone in our atmosphere(NOx), as well as some greenhouse gases, such as methane (CH4). In addition, theabsorption of UV-B radiation by ozone is a natural source of heat in the stratosphere,causing temperatures to increase with altitude. Stratospheric temperatures affect thebalance of ozone production and destruction processes (see Q2) and air motions that
redistribute ozone throughout the stratosphere (see Q4).

UVUVProtection
byby
thetheOzone
Layer
Protection
Ozone
Layer

er
Lay
e
on
Oz
Figure Q3-1. UV protection by the ozone layer.
The ozone layer resides in the stratosphere and surrounds
the entire Earth. UV-B radiation (280- to 315-nanometer (nm)wavelength) from the Sun is strongly absorbed in this layer. As
a result, the amount of UV-B reaching Earth’s surface is greatly
reduced. UV-A (315- to 400-nm wavelength), visible light, andother solar radiation are not strongly absorbed by the ozone
layer. Human exposure to UV-B radiation increases the risks ofskin cancer, cataracts, and a suppressed immune system. UV-B
radiation exposure can also damage terrestrial plant life, singlecell organisms, and aquatic ecosystems.

r th
Ea

Section I: Ozone in our atmosphere13

4

How is total ozone
distributed over the globe?
The distribution of total ozone over Earth varies with location ontimescales that range from daily to seasonal. The variations arecaused by large-scale movements of stratospheric air and thechemical production and destruction of ozone. Total ozone is
generally lowest at the equator and highest in polar regions.
Total ozone. Total ozone at any location on the globe is defined as the sum of all
the ozone in the atmosphere directly above that location. Most ozone resides inthe stratospheric ozone layer and a small percentage (about 10%) is distributed
throughout the troposphere (see Q1). Total ozone values are often reported in Dobson
units denoted as “DU.” Typical values vary between 200 and 500 DU over the globe(see Figure Q4-1). The ozone molecules required for total ozone to be 500 DU aroundthe globe would form a layer of pure ozone gas at Earth’s surface having a thicknessof only 5 millimeters (0.2 inches) (see Q1).
Global distribution. Total ozone varies strongly with latitude over the globe, withthe largest values occurring at middle and high latitudes during all seasons (seeFigure Q4-1). This is the result of the large-scale circulation of air in the stratospherethat slowly transports ozone from the tropics, where ozone production from solar
ultraviolet radiation is highest, toward the poles. Ozone accumulates at middle andhigh latitudes, increasing the thickness (or vertical extent) of the ozone layer and, atthe same time, total ozone. In contrast, the values of total ozone are the lowest in thetropics in all seasons (except for those within the Antarctic ozone hole) because thethickness of the ozone layer is smallest there.
Seasonal distribution. Total ozone also varies with season, as is shown in Figure
Q4-1 using two-week averages of ozone taken from 2009 satellite observations.
March and September plots represent the early spring and fall seasons in the
Northern and Southern Hemispheres. June and December plots similarly represent
the early summer and winter seasons. Total ozone shows a maximum at highlatitudes during spring as a result of increased transport of ozone from its source
region in the tropics toward the polar regions during late fall and winter. This ozone
transport is much weaker during the summer and early fall periods and is weakeroverall in the Southern Hemisphere. An important feature of seasonal ozone changes
is the natural chemical destruction that occurs when daylight is continuous in thesummer polar stratosphere, which causes total ozone to decrease gradually toward
its lowest values in early fall.
This natural seasonal cycle can be observed clearly in the Northern Hemisphere asshown in Figure Q4-1, with increasing values in Arctic total ozone during winter, aclear maximum in spring, and decreasing values from summer to fall. In the Antarctic,
however, a pronounced minimum in total ozone is observed during spring. The
minimum is a consequence of the “ozone hole,” which describes the widespreadchemical destruction of ozone by ozone-depleting substances (see Q6 and Q11) in
spring. In the late 1970s, before the ozone hole appeared each year, much higher
ozone values were found in the Antarctic spring (see Q11). Now, the lowest values of
total ozone across the globe and all seasons are found every late winter and spring in
the Antarctic as shown in Figure Q4-1. After spring, these low values disappear fromtotal ozone maps as polar air mixes with lower-latitude air containing much higher
ozone values.
In the tropics, the total ozone changes through the progression of the seasons aremuch smaller than in the polar regions. This is because seasonal changes in bothsunlight and ozone transport are smaller in the tropics than those in the polar regions.

14

Section I: Ozone in our atmosphereGlobalSatellite
Ozoneinin2009
2009
GlobalSatelliteMaps
Mapsofof Total
Total Ozone
Early spring

16 – 30 March

Early winter

Early summer

Early fall

16 – 30 December

16 – 30 June

Early fall

Early summer

Early winter

Total ozone(Dobson units)

600
500
400
300

16 – 30 September

200
100

Early spring
Figure Q4-1. Total ozone.
Total ozone at any location on the globe is defined as the sum of all the ozone moleculesin the atmosphere directly above that location. Total ozone varies with latitude, longitude,
and season, with the largest values at high latitudes and the lowest values in tropicalregions. The variations are demonstrated here with two-week averages of total ozone in2009 as measured with a satellite instrument. Total ozone shows little variation in thetropics (20°N–20°S latitudes) over all seasons. Total ozone outside the tropics varies
more strongly with time on a daily to seasonal basis as ozone-rich air is moved from thetropics and accumulates at higher latitudes. The low total ozone values over Antarctica
in September constitute the “ozone hole” in 2009. Since the 1980s, the ozone hole inlate winter and early spring represents the lowest values of total ozone that occur over allseasons and latitudes (see Q11).

Section I: Ozone in our atmosphere154Natural variations. Total ozone varies strongly with latitude and longitude as seen
within the seasonal plots in Figure Q4-1. These patterns, which change on daily toweekly timescales, come about for two reasons. First, natural air motions mix andblend air between regions of the stratosphere that have high ozone values andthose that have low ozone values. Tropospheric weather systems can temporarilychange the thickness of the ozone layer in a region, and thereby change total ozone.
The geographical variation in these air motions in turn causes variations in the
distribution of total ozone.
Second, ozone variations occur as a result of changes in the balance of chemical
production and loss processes as air moves to and from different locations over the
globe. This balance, for example, is very sensitive to the amount of sunlight in a region.
There is a good understanding of how chemistry and air motions work together tocause the observed large-scale features in total ozone, such as those seen in Figure
Q4-1. Ozone changes are routinely monitored by a large group of investigators using
satellite, airborne, and ground-based instruments. The continued analysis of theseobservations provides an important basis to quantify the contribution of humanactivities to ozone depletion.

16

Section I: Ozone in our atmosphere5

How is ozone measured
in the atmosphere?
The amount of ozone in the atmosphere is measured byinstruments on the ground and carried aloft on balloons, aircraft,and satellites. Some instruments measure ozone locally bycontinuously drawing air samples into a small detection chamber.
Other instruments measure ozone remotely over long distancesby using ozone’s unique optical absorption or emission properties.
The abundance of ozone in the atmosphere is measured by a variety of techniques(see Figure Q5-1). The techniques make use of ozone’s unique optical and chemical
properties. There are two principal categories of measurement techniques: local andremote. Ozone measurements by these techniques have been essential in monitoringchanges in the ozone layer and in developing our understanding of the processes that
control ozone abundances.
Local measurements. Local measurements of atmospheric ozone abundance
are those that require air to be drawn directly into an instrument. Once inside an
instrument’s detection chamber, ozone is measured by its absorption of ultraviolet(UV) light or by the electrical current or light produced in a chemical reaction involving
ozone. The last approach is used in the construction of “ozonesondes,” which arelightweight, ozone-measuring modules suitable for launching on small balloons.
The balloons ascend far enough in the atmosphere to measure ozone in thestratospheric ozone layer. Ozonesondes are launched regularly at many locations
around the world. Local ozone-measuring instruments using optical or chemical
detection schemes are also used routinely on research aircraft to measure the
distribution of ozone in the troposphere and lower stratosphere. High-altituderesearch aircraft can reach the ozone layer at most locations over the globe and canreach highest into the layer at high latitudes. Ozone measurements are also being
made routinely on some commercial aircraft flights.

Measuring
MeasuringOzone
Ozoneininthe
theAtmosphere
AtmosphereHigh-altitude
aircraft
Satellites
Low-altitude
aircraft

Figure Q5-1. Ozone measurements.
Ozone is measured throughout the atmosphere with instrumentson the ground, aircraft, high-altitude balloons, and satellites. Some
instruments measure ozone locally in sampled air and othersmeasure ozone remotely some distance away from the instrument.
Instruments use optical techniques, with the Sun and lasers as lightsources; detect the thermal emissions from ozone; or use chemical
reactions that are unique to ozone. At many locations over the globe,
regular measurements are made to monitor total ozone amounts
and their variations over time.

Laser
beams
BalloonsondesGround-based
systems

Section I: Ozone in our atmosphere175Remote measurements. Remote measurements of total ozone amounts and thealtitude distributions of ozone are obtained by detecting ozone at large distancesfrom the instrument. Most remote measurements of ozone rely on its uniqueabsorption of UV radiation. Sources of UV radiation that can be used are the Sun,lasers, and starlight. For example, satellites use the absorption of solar UV radiationby the atmosphere or the absorption of sunlight scattered from the surface of Earth
to measure ozone over nearly the entire globe on a daily basis. Lidar instruments,
which measure backscattered laser light, are routinely deployed at ground sites andon research aircraft to detect ozone over a distance of many kilometers along thelaser light path. A network of ground-based detectors measures ozone by detecting
small changes in the amount of the Sun’s UV radiation that reaches Earth’s surface.
Other instruments measure ozone using its absorption of infrared or visible radiation
or its emission of microwave or infrared radiation at different altitudes in theatmosphere, obtaining information on the vertical distribution of ozone. Emission
measurements have the advantage of providing remote ozone measurements atnight, which is particularly valuable for sampling polar regions during winter, when
there is continuous darkness.

Global Ozone Network
The first instrument for routinely monitoring total ozone was developed by Gordon
M.B. Dobson in the United Kingdom in the 1920s. The instrument, called a Féry
spectrometer, made its measurements by examining the wavelength spectrum ofsolar ultraviolet radiation (sunlight) using a photographic plate. A small networkof instruments distributed around Europe allowed Dobson to make importantdiscoveries about how total ozone varies with location and time. In the 1930s a new
instrument was developed by Dobson, now called a Dobson spectrophotometer,
which precisely measures the intensity of sunlight at two ultraviolet wavelengths: onethat is strongly absorbed by ozone and one that is weakly absorbed. The differencein light intensity at the two wavelengths provides a measure of total ozone above theinstrument location.
A global network of total-ozone observing stations was established in 1957 as part of
the International Geophysical Year. Today, there are about 100 sites located around theworld ranging from South Pole, Antarctica (90°S), to Ellesmere Island, Canada (83°N),
that routinely measure total ozone. The accuracy of these observations is maintainedby regular instrument calibrations and intercomparisons. Data from the network havebeen essential for understanding the effects of chlorofluorocarbons (CFCs) and otherozone-depleting substances on the global ozone layer, starting before the launch ofspace-based ozone-measuring instruments and continuing to the present day. Groundbased instruments with excellent long-term stability and accuracy are now routinely
used to help calibrate space-based observations of total ozone.
Pioneering scientists have traditionally been honored by having units of measure
named after them. Accordingly, the unit of measure for total ozone is called the“Dobson unit” (see Q4).

18

Section I: Ozone in our atmosphere6

How do emissions of halogen source gases
lead to stratospheric ozone depletion?
The initial step in the depletion of stratospheric ozone by humanactivities is the emission, at Earth’s surface, of gases containing chlorine
and bromine. Most of these gases accumulate in the lower atmospherebecause they are unreactive and do not dissolve readily in rain orsnow. Natural air motions transport these accumulated gases to thestratosphere, where they are converted to more reactive gases. Someof these gases then participate in reactions that destroy ozone. Finally,
when air returns to the lower atmosphere, these reactive chlorine andbromine gases are removed from Earth’s atmosphere by rain and snow.
Emission, accumulation, and transport. The principal steps in stratospheric ozone
depletion caused by human activities are shown in Figure Q6-1. The process begins
with the emission, at Earth’s surface, of source gases containing the halogens
chlorine and bromine (see Q7). The halogen source gases, often referred to asozone-depleting substances (ODSs), include manufactured chemicals released tothe atmosphere in a variety of applications, such as refrigeration, air conditioning,
and foam blowing. Chlorofluorocarbons (CFCs) are an important example of chlorinecontaining gases. Emitted source gases accumulate in the lower atmosphere(troposphere) and are transported to the stratosphere by natural air motions. The
accumulation occurs because most source gases are highly unreactive in the lower
atmosphere. Small amounts of these gases dissolve in ocean waters. The low
reactivity of these manufactured halogenated gases is one property that makes themwell suited for specialized applications such as refrigeration.
Some halogen gases are emitted in substantial quantities from natural sources (seeQ7). These emissions also accumulate in the troposphere, are transported to thestratosphere, and participate in ozone destruction reactions. These naturally emitted
gases are part of the natural balance of ozone production and destruction that
predates the large release of manufactured halogenated gases.
Conversion, reaction, and removal. Halogen source gases do not react directly withozone. Once in the stratosphere, halogen source gases are chemically converted
to reactive halogen gases by ultraviolet radiation from the Sun (see Q8). The rate ofconversion is related to the atmospheric lifetime of a gas (see Q7). Gases with longer
lifetimes have slower conversion rates and survive longer in the atmosphere after
emission. Lifetimes of the principal ODSs vary from 1 to 100 years (see Q7). Emitted
gas molecules with atmospheric lifetimes greater than a few years circulate between
the troposphere and stratosphere multiple times, on average, before conversion occurs.
The reactive gases formed from halogen source gases react chemically to destroyozone in the stratosphere (see Q9). The average depletion of total ozone attributed
to reactive gases is smallest in the tropics and largest at high latitudes (see Q13). In
polar regions, surface reactions that occur at low temperatures on polar stratosphericclouds (PSCs) greatly increase the abundance of the most reactive chlorine gas,chlorine monoxide (ClO) (see Q10). This results in substantial ozone destruction inpolar regions in late winter and early spring (see Q11 and Q12).
After a few years, air in the stratosphere returns to the troposphere, bringing along
reactive halogen gases. These reactive halogen gases are then removed from theatmosphere by rain and other precipitation or deposited on Earth’s land or ocean
surfaces. This removal brings to an end the destruction of ozone by chlorine andbromine atoms that were first released to the atmosphere as components of halogen
source gas molecules.

Section II: The ozone depletion process

19

6
Tropospheric conversion. Halogen source gases with short lifetimes (less than 1year) undergo significant chemical conversion in the troposphere, producing reactive
halogen gases and other compounds. Source gas molecules that are not converted
are transported to the stratosphere. Only small portions of reactive halogen gases
produced in the troposphere are transported to the stratosphere because most areremoved by precipitation. Important examples of halogen gases that undergo sometropospheric removal are the hydrochlorofluorocarbons (HCFCs), methyl bromide(CH3Br) and methyl chloride (CH3Cl), and gases containing iodine (see Q7).

Principal
PrincipalSteps
Stepsinin thethe Depletion
DepletionofStratosphericof Stratospheric Ozone
Ozone
Emissions
Halogen source gases are emitted at Earth's
surface by human activities and natural processes.

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

Transport
Halogen source gases are transported to thestratosphere by air motions.

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

Chemical reaction
Reactive halogen gases cause chemical depletionof stratospheric ozone over the globe.

Figure Q6-1. Principal steps in stratospheric ozone depletion.
The stratospheric ozone depletion process begins with theemission of halogen source gases by human activities andnatural processes. Those emitted by human activities are also
called ozone-depleting substances (ODSs). Subsequent steps
are accumulation, transport, conversion, chemical reaction,and removal. Ozone depletion by halogen source gases occurs
globally. Large seasonal ozone losses occur in polar regions
as a result of reactions involving polar stratospheric clouds(PSCs). Ozone depletion ends when reactive halogen gases are
removed by rain and snow in the troposphere and deposited on
Earth’s surface.

20

Section II: The ozone depletion process

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 halogen gases returns
to the troposphere where the gases are removed bymoisture in clouds and rain.

Understanding Stratospheric Ozone Depletion
Our understanding of stratospheric ozone depletion has been obtained througha combination of laboratory studies, computer models, and atmosphericobservations. The wide variety of chemical reactions that occur in the stratospherehave been discovered and studied in laboratory studies. Chemical reactions
between two gases follow well-defined physical rules. Some of these reactions
occur on the surfaces of polar stratospheric clouds (PSCs) formed in the winterstratosphere. Reactions have been studied that involve many different molecules
containing chlorine, bromine, fluorine, and iodine and other atmosphericconstituents such as carbon, oxygen, nitrogen, and hydrogen. These studies haveshown that several reactions involving chlorine and bromine directly or indirectly
destroy ozone in the stratosphere.
Computer models have been used to examine the combined effect of the large
group of known reactions that occur in the stratosphere. These models simulate
the stratosphere by including representative chemical abundances, winds, airtemperatures, and the daily and seasonal changes in sunlight. These analyses
show that under certain conditions chlorine and bromine react in catalytic
cycles in which one chlorine or bromine atom destroys many thousands ofozone molecules. Models are also used to simulate ozone amounts observed inprevious years as a strong test of our understanding of atmospheric processes
and to evaluate the importance of new reactions found in laboratory studies. The
responses of ozone to possible future changes in the abundances of trace gases,
temperatures, and other atmospheric parameters have been extensively exploredwith specialized computer models (see Q20).
Atmospheric observations have shown what gases are present in different regionsof the stratosphere and how their abundances vary. Gas and particle abundances
have been monitored over time periods spanning a daily cycle to decades.
Observations show that halogen source gases and reactive halogen gases arepresent in the stratosphere at the amounts required to cause observed ozone
depletion. Ozone and chlorine monoxide (ClO), for example, have been observed
extensively with a variety of instruments. ClO is a highly reactive gas that is
involved in catalytic ozone destruction cycles throughout the stratosphere (see Q9).
Instruments on the ground and on satellites, balloons, and aircraft now routinely
detect ozone and ClO remotely using optical and microwave signals. High-altitude
aircraft and balloon instruments are also used to detect both gases locally in thestratosphere (see Q5). The observations of ozone and reactive gases made in past
decades are used extensively in comparisons with computer models in order toincrease confidence in our understanding of stratospheric ozone depletion.

Section II: The ozone depletion process21

7What emissions from humanactivities lead to ozone depletion?
Certain industrial processes and consumer products resultin the emission of ozone-depleting substances (ODSs) to theatmosphere. ODSs are manufactured halogen source gases thatare controlled worldwide by the Montreal Protocol. These gases
bring chlorine and bromine atoms to the stratosphere, where they
destroy ozone in chemical reactions. Important examples are thechlorofluorocarbons (CFCs), once used in almost all refrigeration
and air conditioning systems, and the halons, which were used infire extinguishers. Current ODS abundances in the atmosphere are
known directly from air sample measurements.
Halogen source gases versus ODSs. Those halogen source gases emitted by humanactivities and controlled by the Montreal Protocol are referred to as ODSs within
the Montreal Protocol, by the media, and in the scientific literature. The Montreal
Protocol now controls the global production and consumption of all ODSs (see Q15).
Halogen source gases that have only natural sources are not classified as ODSs.
The contributions of ODSs and natural halogen source gases to the total amountof chlorine and bromine entering the stratosphere are shown in Figure Q7-1. Totalamounts of chlorine peaked in 1993, and of bromine in 1998. The difference in thetiming of the peaks is explained by the phase-out schedules, how closely emissions
are tied to production, and the lifetimes of the various source gases. Also shown are
the contributions to total chlorine and bromine in 2012, highlighting the reductions
achieved due to the Montreal Protocol controls.
Ozone-depleting substances (ODSs). ODSs are manufactured for specific industrial
uses or consumer products, most of which result in the eventual emission ofthese gases to the atmosphere. Total ODS emissions increased substantially fromthe middle to the late 20th century, reached a peak in the late 1980s, and are
now in decline (see Figure Q0-1). A large fraction of the emitted ODSs reach thestratosphere, where they are converted to reactive gases containing chlorine andbromine that lead to ozone depletion. ODSs containing only carbon, chlorine, andfluorine are called chlorofluorocarbons, usually abbreviated as CFCs. The principal
CFCs are CFC-11 (CCl3F), CFC-12 (CCl2F2), and CFC-113 (CCl2FCClF2). CFCs, alongwith carbon tetrachloride (CCl4) and methyl chloroform (CH3CCl3), historically havebeen the most important chlorine-containing halogen source gases emitted byhuman activities. These and other chlorine-containing ODSs have been used inmany applications, including refrigeration, air conditioning, foam blowing, aerosol
propellants, and cleaning of metals and electronic components. As a result of theMontreal Protocol controls, the abundances of most of these chlorine source gases
have decreased between 1993 and 2012 (see Figure Q7-1). While the concentrationof CFC-12 in 2012 was 2.4% higher than it was in 1993, it peaked in 2000 andhas decreased since then (see Figure Q16-1). As substitute gases for CFCs,
hydrochlorofluorocarbons (HCFCs) increased substantially between 1993 and 2012(+157.5%). With new restrictions on global production in place since 2013, HCFCs are
expected to reach peak values between 2020 and 2030.
Another category of ODSs contains bromine. The most important of these gases
are the halons and methyl bromide (CH3Br). Halons are halocarbon gases originallydeveloped to extinguish fires. Halons were widely used to protect large computer
installations, military hardware, and commercial aircraft engines. As a consequence,
halons are often released directly into the atmosphere upon use. Halon-1211 andhalon-1301 are the most abundant halons emitted by human activities. Methyl

22

Section II: The ozone depletion processbromide is used primarily as a fumigant for pest control in agriculture and disinfectionof export shipping goods, but also has significant natural sources. As a result of theMontreal Protocol, the abundances of methyl bromide from human sources havesubstantially decreased between 1998 and 2012 (−61.8%; see Figure Q7-1). Halon1211 reached peak values in 2005 and has been decreasing since then. Halon-1301
abundances, on the other hand, have increased by +19.9% since 1998 and are
expected to increase further (although at a slower rate) because of continued small
releases and a long atmospheric lifetime.
Natural sources of chlorine and bromine. There are a few halogen source gases
present in the stratosphere that have large natural sources. These include methyl
chloride (CH3Cl) and methyl bromide (CH3Br), both of which are emitted by oceanicand terrestrial ecosystems. Natural sources of these two gases contributed about
17% to the total chlorine in the stratosphere in 2012 and about 30% of the total

HalogenSourceGases
GasesEntering
Enteringthe
theStratosphere
Halogen Source
Stratosphere3658 pptv 3300 pptv

1000

0

CFC-113+157.5%HCFCs-19.4%

Carbon tetrachloride-12.0%

CFC-11

-1.8%

1993

Total

22.1 pptv 20.1 pptv

CFC-12

Methyl chloride

2012

Chlorine source gases
Figure Q7-1. Changes in halogen source gases entering
the stratosphere.
A variety of halogen source gases emitted from natural sources
and by human activities transport chlorine and bromine into thestratosphere. Ozone-depleting substances (ODSs) are the subsetof these gases emitted by human activities that are controlled
by the Montreal Protocol. These partitioned columns show thesources and abundances of chlorine- and bromine-containing
gases entering the stratosphere in 1993 and 1998, when their
total amounts peaked, respectively, and in 2012. The overall
reductions in the total amounts of chlorine and bromine and thechanges observed for each source gas are also indicated. The
approximate amounts are derived from tropospheric observationsof each gas. Note the large difference in the vertical scales: total
chlorine entering the stratosphere is 150 times more abundant
than total bromine. Human activities are the largest source ofchlorine reaching the stratosphere and the CFCs are the most

20

15

-9%

ODSs
Human-activity sources-10.3%+2.4%

Naturalsources

Methyl chloroform - 96.0%
Other gases +17.2%

50-15.1%

Other gases+19.9%

Halon-1301+1.8%

Halon-1211- 61.8%10Natural sources

2000

Human-activity sources

3000

Ozone-depleting substances (ODSs)

Total chlorine amount (parts per trillion)-10%

Total bromine amount (parts per trillion)

Total

Methyl bromide

+0%

+0%

1998

Very short-lived gases

2012

Bromine source gases
abundant chlorine-containing gases. Methyl chloride is theprimary natural source of chlorine. Largest decreases between
1993 and 2012 are seen in CFC-11, carbon tetrachloride,and methyl chloroform. HCFCs, which are substitute gases for
CFCs and also controlled under the Montreal Protocol, show
small but still strongly increasing abundances. For bromine
entering the stratosphere, halons and methyl bromide are thelargest contributors. Methyl bromide has an additional, muchlarger, natural source. Natural sources provide a much larger
contribution to total bromine entering the stratosphere than
to total chlorine. Largest decreases between 1998 and 2012
are seen in methyl bromide. Halon-1211 and halon-1301 show
increasing abundances. (The unit “parts per trillion” is used here
as a measure of the relative abundance of a gas in air: 1 part per
trillion equals the presence of one molecule of a gas per trillion(=1012) total air molecules.)Section II: The ozone depletion process23

7
bromine (Figure Q7-1). Very short-lived source gases containing bromine, such as
bromoform (CHBr3), are also released to the atmosphere primarily from biologicalactivity in the oceans. Only a fraction of these emissions reaches the stratosphere,
because these gases are rapidly removed in the lower atmosphere. The contributionof these very short-lived gases to the total amount of bromine in the stratosphereis estimated to be about 25% in 2012, but this number has a large uncertainty.
The amounts of chlorine and bromine in the stratosphere from natural sources
are believed to be fairly constant and, therefore, cannot be the cause of the ozone
depletion observed since the 1980s.
Other human sources of chlorine and bromine. Other chlorine- and brominecontaining 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 andindustrial water supplies (including swimming pools). Most of these gases are very
short-lived and only a small fraction of their emissions reaches the stratosphere. The
contribution of very short-lived chlorinated gases from natural and human sources tototal stratospheric chlorine is only about 3% and is included in Other gases in Figure
Q7-1. The Montreal Protocol does not control the production and consumption ofvery short-lived substances (VSLS), although the concentrations of some (notably
dichloromethane, CH2Cl2) have increased substantially in recent years.
Lifetimes and emissions. After emission, halogen source gases are either naturally
removed from the atmosphere or undergo chemical conversion in the troposphereor stratosphere. The time to remove or convert about 60% of a gas is often called
its atmospheric lifetime. Lifetimes vary from less than 1 year to 100 years for the
principal chlorine- and bromine-containing gases (see Table Q7-1). The long-lived
gases are converted to other gases primarily in the stratosphere and essentially allof their original halogen content becomes available to participate in the destructionof stratospheric ozone. Gases with short lifetimes (e.g., the HCFCs, methyl bromide,
methyl chloride, and the very short-lived gases) are effectively converted to othergases in the troposphere, which are then removed by rain. Therefore, only a fraction of
their halogen content potentially contributes to ozone depletion in the stratosphere.
The amount of an emitted gas that is present in the atmosphere represents abalance between its emission rate and removal rate. Emission rates and atmospheric
lifetimes vary greatly for the source gases, as indicated in Table Q7-1. For example,
the atmospheric abundances of most of the principal CFCs and halons havedecreased since 1990 in response to smaller emission rates, while those of theleading substitute gases, the hydrochlorofluorocarbons (HCFCs), continue to increase
under the provisions of the Montreal Protocol (see Q16). In the coming decades,
the emissions and atmospheric abundances of all controlled gases are expected todecrease under these provisions.
Ozone Depletion Potential (ODP). Halogen source gases are compared in theireffectiveness to destroy stratospheric ozone using the ODP, as listed in Table Q7-1
(see Q18). A gas with a larger ODP destroys more ozone over its atmospheric lifetime.
The ODP is calculated relative to CFC-11, which has an ODP defined to be 1. The
calculations, which require the use of atmospheric computer models, use as thebasis of comparison the ozone depletion from an equal mass of each gas emitted tothe atmosphere. Halon-1211 and halon-1301 have ODPs significantly larger than thatof CFC-11 and most other chlorinated gases because bromine is much more effective(about 60 times) on a per-atom basis than chlorine in chemical reactions that destroyozone. The gases with small ODP values generally have short atmospheric lifetimes orcontain fewer chlorine and bromine atoms.

24

Section II: The ozone depletion process

Table Q7-1.
Atmospheric lifetimes, global emissions, Ozone Depletion Potentials, and Global Warming
Potentials of some halogen source gases and HFC substitute gases.

Atmospheric
Lifetime (years)

Global Emissions
in 2012 (Kt/yr) a

Ozone Depletion
Potential (ODP) c

Global Warming
Potential (GWP) c

CFC-11

52

46–68

1

5160

CFC-12

102

16–64

0.73

10300

CFC-113

93

0–7

0.81

6080

Carbon tetrachloride (CCl4)

26

40–74

0.72

1730

HCFCs

1–18

400–528

0.01–0.10

800–2070

Methyl chloroform (CH3CCl3)

5

0–5

0.14

153

Methyl chloride (CH3Cl)

0.9

2707

0.015

Very short-lived Cl-containing gases

less than 0.5

b

b, d

Halon-1301

72

1.4–2

15.2

6670

Halon-1211

16

0.3–9.3

6.9

1750

Methyl bromide (CH3Br)

0.8

85

0.57

2

Very short-lived Br-containing gases(e.g., CHBr3)

less than 0.5b

HFC-134a14

144–21501360

HFC-23

228

11–14

0

12500

HFC-143a

51

20–25

0

5080

HFC-125

31

31–47

0

3450

HFC-152a

1.6

40–66

0

148

HFC-32

5.4

12–30

0

700

Gas
Halogen source gases
Chlorine gasesevery low11bless than 1

Bromine gases

260–1080

b, dvery lowbvery low

Hydrofluorocarbons (HFCs)a

Includes 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).
b
Estimates are very uncertain for most species.
c100-year GWP. ODPs and GWPs are discussed in Q18. Values are calculated foremissions of an equal mass of each gas.
d
ODP depends strongly on where the emissions occur for these gases. Emissions in thetropics have higher ODPs than emissions at midlatitudes.
e
New measurements are now available of CFC-112a, CFC-112, CFC-113a, and HCFC133. These gases are by-products of ODS-manufacturing and constitute only very small
contributions to the total of chlorine- and bromine-containing gases.

Section II: The ozone depletion process25

7Fluorine and iodine. Fluorine and iodine are also halogens. Many of the sourcegases in Figure Q7-1 also contain fluorine in addition to chlorine or bromine. After thesource gases undergo conversion in the stratosphere (see Q6), the fluorine content
of these gases is left in chemical forms that do not cause ozone depletion. As aconsequence, halogen source gases that contain fluorine and no other halogens are
not classified as ODSs. An important example of these are the hydrofluorocarbons(HFCs), which are included in Table Q7-1 because they are common ODS substitutegases. HFCs have ODPs of zero and are also strong greenhouse gases (see Q18).
Iodine is a component of several gases that are naturally emitted from the oceans
and some human activities. Although iodine can participate in ozone destruction
reactions, iodine-containing source gases all have very short lifetimes. Currently there
is no evidence that any of these gases reach the stratosphere in significant amounts.
Other non-halogen gases. Other non-halogen gases that influence stratosphericozone abundances have also increased in the stratosphere as a result of emissions
from human activities. Important examples are methane (CH4) and nitrous oxide(N2O), which react in the stratosphere to form water vapor and reactive hydrogen, andnitrogen oxides, respectively. These reactive products participate in the destructionof stratospheric ozone (see Q2). Increasing abundances of N2O and CH4, as well as
CO2, are expected to significantly affect future stratospheric ozone through combined
effects on temperature, winds, and chemistry (see Q20). Although all of these gases
are part of the Kyoto Protocol (see Q15) because they are climate gases, they are
not classified as ODSs under the Montreal Protocol. Although past emissions of
ODSs still dominate global ozone depletion today, the current emissions of N2O fromhuman activities will destroy more stratospheric ozone in the future than the current
emissions of any ODS.

26

Section II: The ozone depletion process8What are the reactive halogen gasesthat destroy stratospheric ozone?
Emissions from human activities and natural processes represent
a large source of chlorine- and bromine-containing gases that enter
the stratosphere. When exposed to ultraviolet radiation from theSun, these halogen source gases are converted to more reactive
gases containing chlorine and bromine. Some reactive gases act aschemical reservoirs that convert to form the most reactive gases,
namely chlorine monoxide (ClO) and bromine monoxide (BrO). The
most reactive gases participate in catalytic reactions that efficiently
destroy ozone. Most volcanoes emit some reactive halogen gasesthat readily dissolve in water and are usually washed out of theatmosphere before they can reach the stratosphere.
Halogen-containing gases present in the stratosphere can be divided into two groups:
halogen source gases and reactive halogen gases (Figure Q8-1). The source gases,
which include ozone-depleting substances (ODSs), are emitted at Earth’s surface
by natural processes and by human activities (see Q7). Once in the stratosphere,
the halogen source gases chemically convert at different rates to form the reactive
halogen gases. The conversion occurs in the stratosphere for most gases instead of
the troposphere because solar ultraviolet radiation (a component of sunlight) is more
intense in the stratosphere. Reactive gases containing the halogens chlorine andbromine lead to the chemical destruction of stratospheric ozone.

Stratospheric Halogen Gases
Stratospheric
Halogen Gases

ChemicalSolarultraviolet (UV)radiation andchemical
reactionsconversion

Figure Q8-1. Conversion of halogen source gases.
Halogen source gases containing chlorine and bromine are chemically converted toreactive halogen gases primarily in the stratosphere. The conversion requires solar
ultraviolet radiation and a few chemical reactions. The shorter-lived gases undergo someconversion in the troposphere. The reactive halogen gases contain all the chlorine andbromine originally present in the source gases before conversion. The reactive gases
can be grouped into the reservoir gases, which do not directly destroy ozone, and themost reactive gases, which participate in ozone destruction cycles (see Q9). A principal
reactive gas, ClO, is formed by reactions of the reservoir gases HCl and ClONO2 that
occur on the surfaces of liquid and solid polar stratospheric clouds (PSCs) (see Q10).

Section II: The ozone depletion process27

8Measurementsof
ofChlorine
Chlorine Gases
Space
Measurements
Gasesfrom
from
Space
Annual mean 2006 (30°–70°N)

50

30

ClO

25

Other
chlorine
gases

20
ClONO2

Total
available
chlorine

15

Altitude (miles)

40

Altitude (kilometers)

Figure Q8-2. Reactive chlorine gas observations.
The abundances of chlorine source gases and reactive chlorine
gases as measured from space in 2006 are displayed withaltitude for a midlatitude location. In the troposphere (below
about 10 kilometers), all chlorine is contained in the sourcegases. In the stratosphere, the abundances of reactive chlorine
gases increase with altitude as the amounts of chlorine sourcegases decline. This is a consequence of chemical reactions
initiated by solar ultraviolet radiation that convert source gases
to reactive gases (see Figure Q8-1). The principal reactive
chlorine gases formed are HCl, ClONO2, and ClO. Summing
the source gases with the reactive gases gives “Total available
chlorine,” which is nearly constant with altitude throughout thestratosphere. In the ozone layer (15-25 km), chlorine sourcegases are still present and HCl and ClONO2 are the mostabundant reactive chlorine gases. (The unit “parts per trillion”
is defined in the caption of Figure Q7-1.)

30
Hydrogenchloride (HCl)20
10
Chlorine source gases(CFCs, HCFCs, CCl4, CH3Cl, etc.)

10

0

0

1000

2000

5

3000

0
4000

Chlorine (parts per trillion)

Reactive halogen gases. The chemical conversion of halogen source gases, which
involves solar ultraviolet radiation and other chemical reactions, produces a numberof reactive halogen gases. These reactive gases contain all of the chlorine andbromine atoms originally present in the source gases. The most important reactive
chlorine- and bromine-containing gases that form in the stratosphere are shown in
Figure Q8-1. Throughout the stratosphere, the most abundant are typically hydrogenchloride (HCl) and chlorine nitrate (ClONO2). These two gases are considered
important reservoir gases because, while they don’t react directly with ozone, they
can be converted to the most reactive forms that do chemically destroy ozone. The
most reactive forms are chlorine monoxide (ClO) and bromine monoxide (BrO), andchlorine and bromine atoms (Cl and Br). A large fraction of available reactive bromine
is generally in the form of BrO, whereas usually only a small fraction of reactive
chlorine is in the form of ClO. The special conditions that occur in the polar regions
in winter cause the reservoir gases ClONO2 and HCl to undergo nearly complete
conversion to ClO in reactions on polar stratospheric clouds (PSCs) (see Q10).
Reactive chlorine at midlatitudes. Reactive chlorine gases have been observed
extensively in the stratosphere using both local and remote measurement techniques.
The measurements from space displayed in Figure Q8-2 are representative of how
the amounts of chlorine-containing gases change between the surface and the upper
stratosphere at mid- to high latitudes. Available chlorine (see red line in Figure Q8-2)
is the sum of chlorine contained in halogen source gases and in the reactive gases(e.g., CFC-11, ClONO2, ClO). Available chlorine is constant to within about 10% fromthe surface to above 50 kilometers (31 miles) altitude. In the troposphere, available
chlorine is contained almost entirely in the source gases described in Figure Q7-1. At
higher altitudes, the source gases become a smaller fraction of available chlorine asthey are converted to the reactive chlorine gases. At the highest altitudes, available
chlorine is all in the form of the reactive chlorine gases.
In the altitude range of the ozone layer at midlatitudes, as shown in Figure Q8-2, thereservoir gases HCl and ClONO2 account for most of the available chlorine. ClO, the most
28

Section II: The ozone depletion processreactive gas in ozone depletion, is a small fraction of available chlorine. The low abundanceof ClO limits the amount of ozone destruction that occurs outside of polar regions.
Reactive chlorine in polar regions. Reactive chlorine gases in polar regions undergo
large changes between the fall and late winter. Meteorological and chemical conditions
in both polar regions are now routinely observed from space in all seasons. Fall andwinter conditions over the Antarctic are contrasted in Figure Q8-3 using seasonalobservations made near the center of the ozone layer (about 18 km (11.3 miles)
altitude) (see Figure Q12-3).
In fall (May), ozone values are high over the entire Antarctic continent and beyond.
Temperatures are mid-range, HCl and nitric acid (HNO3) are high, and ClO is very low. HighHCl indicates that substantial conversion of halogen source gases has occurred in thestratosphere. In the past decades, the reservoir gases HCl and ClONO2 have increased
substantially in the stratosphere following increased emissions of halogen source gases.
HNO3 is an abundant, naturally occurring stratospheric compound that plays major roles
in stratospheric ozone chemistry by both moderating ozone destruction and condensing
to form polar stratospheric clouds (PSCs). Low ClO indicates that little conversion of thereservoir gases occurs in the fall, thereby limiting catalytic ozone destruction.
By late winter (September), ozone amounts reflect substantial depletion at 18-km
altitude over an area larger than the Antarctic continent. Antarctic ozone holes arise
from similar chemical destruction throughout much of the altitude range of the ozone
layer (see altitude profile in Figure Q12-3). The meteorological and chemical conditions
in late winter are very different from those found in fall: very low temperatures, very low
HCl and HNO3, and very high ClO. Low stratospheric temperatures are characteristic ofwinter, when solar heating is reduced. Low HCl and high ClO reflect the conversion of thereactive halogen reservoir compounds, HCl and ClONO2, to the most reactive chlorine
form, ClO. This conversion occurs selectively in winter on PSCs, which form at the very low
temperatures (see Q10). Low HNO3 is indicative of its condensation to form PSCs, someof which subsequently descend to lower altitudes through gravitational settling. High ClO
abundances generally cause ozone depletion to continue in the Antarctic region until
mid-October (spring), when the lowest ozone values usually are observed (see Q11). As
temperatures rise at the end of the winter, PSC formation is halted, ClO is converted backinto the reservoir species HCl and ClONO2 (see Q10), and ozone destruction is curtailed.
Similar though less dramatic changes in meteorological and chemical conditions arealso observed between fall and late winter in the Arctic, where ozone depletion is lesssevere than in the Antarctic.
Reactive bromine observations. Fewer measurements are available for reactive bromine
gases in the lower stratosphere than for reactive chlorine, in part because of the lower
abundance of bromine. The most widely observed bromine gas is bromine monoxide(BrO), which can be observed from space. Estimates of reactive bromine abundances inthe stratosphere are larger than expected from the conversion of the halons and methyl
bromide source gases, suggesting that the contribution of the very short-lived brominecontaining gases to reactive bromine must also be significant (see Q7).
Other sources. Other emission sources of reactive halogen gases exist that areassociated with natural processes and human activities. Most emissions becometrapped in the troposphere when dissolved in water, and ultimately are returned to
Earth’s surface before they can reach the stratosphere. Volcanoes are an important
episodic source of reactive halogen gases (e.g., HCl) that do not reach the stratospherein appreciable quantities (see Q14). Other examples include reactive chlorine producedby evaporation of ocean spray. Sea salt products dissolve in water and are removed in
the troposphere. Solid rocket motors, such as those used on the Space Shuttle, release

Section II: The ozone depletion process29

8reactive chlorine gases directly into the troposphere and stratosphere. The quantities
emitted by the Space Shuttle were small in comparison with halogen emissions fromother human activities.

Figure Q8-3. Chemical conditions in theozone layer over Antarctica.
Observations of the chemical conditions in
the Antarctic region highlight the changes
associated with severe ozone depletionthat forms the ozone hole. Satellite
instruments now routinely monitor ozone,
reactive chlorine gases, and temperaturesin the global stratosphere. Results areshown here for fall (May) and late winter(September) seasons in Antarctic winter fora narrow altitude region near 18 kilometers(11.2 miles) within the ozone layer (seeFigure Q12-3). Ozone has naturallyhigh values in fall, before the onset ofozone destruction reactions that cause
widespread depletion. The high ozone isaccompanied by moderate temperatures,
high values of HCl and HNO3, and very low
ClO amounts. When ClO is not enhanced,
significant ozone destruction from ozonedepleting substances does not occur.
Chemical conditions are quite different inlate winter when ozone undergoes severe
depletion. Temperatures are much lower,
HCl has been converted to ClO (the mostreactive chlorine gas), and HNO3 has been
removed by the gravitational settling of PSC
particles. ClO values closely surrounding
the South Pole are low in September
because ClO formation requires sunlight,
which is still gradually returning to thehighest latitudes. The high ClO values inlate winter covering an extensive areathat at times exceeds that of the Antarctic
continent can last for 1 to 2 months, andefficiently destroy ozone in sunlit regionsin late winter/early spring. Ozone typicallyreaches its minimum values in early to midOctober (see Q12). Note that the first andlast colors in the color bar represent values
outside the indicated range of values. (Theunit “parts per billion,” abbreviated “ppb,”
is used here as a measure of the relative
abundance of a gas in air: 1 part per billion
equals the presence of one molecule of agas per billion (=109) total air molecules(compare to ppt in Figure Q7-1).)30

Section II: The ozone depletion process

Chemical Conditions Observed in the Ozone Layer Over Antarctica
Chemical Conditions Observed in the Ozone Layer Over Antarctica
Normal ozone amounts in fall (1 May 2008) at 18-km altitude

Ozone

Temperature

Nitric acid (HNO3)

Hydrogen chloride (HCl)

Chlorine monoxide (ClO)

Large ozone depletion in late winter (15 September 2008) at 18-km altitude

Ozone

Low

Temperature

Nitric acid (HNO3)

Hydrogen chloride (HCl)

Chlorine monoxide (ClO)

Temperatures and chemical abundances
1800

-85°
1.0
0.2
0.1

Ozone

Temperature
HNO3
HCl
ClO

High

2800 ppb-60°C
9.0 ppb
1.7 ppb
1.1 ppb9What are the chlorine and bromine reactions
that destroy stratospheric ozone?
Reactive gases containing chlorine and bromine destroy
stratospheric ozone in “catalytic” cycles made up of two or more
separate reactions. As a result, a single chlorine or bromine atom
can destroy many thousands of ozone molecules before it leaves
the stratosphere. In this way, a small amount of reactive chlorine or
bromine has a large impact on the ozone layer. A special situation
develops in polar regions in the late winter/early spring season,
where large enhancements in the abundance of the most reactive
gas, chlorine monoxide, lead to severe ozone depletion.
Stratospheric ozone is destroyed by reactions involving reactive halogen gases, which
are produced in the chemical conversion of halogen source gases (see Figure Q8-1).
The most reactive of these gases are chlorine monoxide (ClO), bromine monoxide(BrO), and chlorine and bromine atoms (Cl and Br). These gases participate in threeprincipal reaction cycles that destroy ozone.
Cycle 1. Ozone destruction Cycle 1 is illustrated in Figure Q9-1. The cycle is madeup of two basic reactions: Cl + O3 and ClO + O. The net result of Cycle 1 is to convert
one ozone molecule and one oxygen atom into two oxygen molecules. In each cycle,
chlorine acts as a catalyst because ClO and Cl react and are re-formed. In this way,
one Cl atom participates in many cycles, destroying many ozone molecules. For typicalstratospheric conditions at middle or low latitudes, a single chlorine atom can destroy
hundreds of ozone molecules before it happens to react with another gas, breaking
the catalytic cycle. During the total time of its stay in the stratosphere, a chlorine
atom can thus destroy up to tens of thousands of ozone molecules.
Polar Cycles 2 and 3. The abundance of ClO is greatly increased in polar regions
during winter as a result of reactions on the surfaces of polar stratospheric clouds(PSCs) (see Q8 and Q10). Cycles 2 and 3 (see Figure Q9-2) become the dominantFigure Q9-1. Ozone destruction Cycle 1.
The destruction of ozone in Cycle 1
involves two separate chemical reactions.
The net or overall reaction is that of
atomic oxygen with ozone, forming two
oxygen molecules. The cycle can beconsidered to begin with either ClO or Cl.
When starting with ClO, the first reaction
is ClO with O to form Cl. Then, Cl reacts
with ozone and re-forms ClO, consuming
ozone in the process. The cycle thenbegins again with another reaction of
ClO with O. Chlorine is considered acatalyst for ozone destruction because
Cl and ClO are re-formed each time thereaction cycle is completed, and hence
available for further ozone destruction.
Atomic oxygen (O) is formed when solar
ultraviolet radiation (sunlight) reacts withozone and oxygen molecules. Cycle 1 ismost important in the stratosphere attropical and middle latitudes, where solar
ultraviolet radiation is most intense.

Ozone
OzoneDestruction
DestructionCycle
Cycle11
Oxygen molecule (O2)

Chlorine atom (Cl)Ozone (O3)Cl

ClO + O
reaction

Chlorine
catalytic
cycle

Cl + O3
reaction

Ozone
destruction

Cl
Oxygen atom (O)Chlorine monoxide (ClO)

Cl + O3
ClO + O
Net: O + O3

Oxygen molecule (O2)

ClO + O2
Cl + O2
2 O2

Section II: The ozone depletion process31

9reaction 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 withthe reaction of ClO with BrO, has two reaction pathways to produce either Cl and Br or
BrCl. The net result of both cycles is to destroy two ozone molecules and create threeoxygen molecules. Cycles 2 and 3 account for most of the ozone loss observed in the
Arctic and Antarctic stratospheres in the late winter/early spring season (see Q11 and
Q12). At high ClO abundances, the rate of polar ozone destruction can reach 2 to 3%
per day in late winter/early spring.

Ozone Destruction Cycles
in Polar Regions

Cycle 2
ClO + ClO
(ClO)2+ sunlightClOO
2(Cl + O3
Net: 2O3(ClO)2
ClOO + Cl
Cl + O2
ClO + O2 )3O2

Cycle 3or(

ClO + BrO
ClO + BrO
BrCl + sunlight
Cl + O3
Br + O3
Net: 2O3

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

3O2

Figure Q9-2. Polar ozone destruction
Cycles 2 and 3.
Significant destruction of ozone
occurs in polar regions where ClO
abundances reach large values. In
this case, the cycles initiated by thereaction of ClO with another ClO (Cycle
2) or the reaction of ClO with BrO(Cycle 3) efficiently destroy ozone.
The net reaction in both cases is twoozone molecules forming three oxygen
molecules. The reaction of ClO with BrO
has two pathways to form the Cl and Br
product gases. Ozone destruction Cycles
2 and 3 are catalytic, as illustrated for
Cycle 1 in Figure Q9-1, because chlorine
and bromine gases react and are reformed each time the reaction cycle iscompleted. Visible sunlight is required
to complete each cycle and to help
form and maintain ClO abundances.
During polar night and other periodsof darkness, ozone cannot be destroyedby these reactions.

32

Section II: The ozone depletion process

Sunlight requirement. Sunlight is required to complete and maintain Cycles 1through 3. Cycle 1 requires ultraviolet radiation (a component of sunlight) that isstrong enough to break apart molecular oxygen into atomic oxygen. Cycle 1 is mostimportant in the stratosphere at tropical and middle latitudes, where solar ultraviolet
radiation is most intense.
Cycles 2 and 3 require visible sunlight to complete the reaction cycles and to maintain
ClO abundances. In the continuous darkness of winter in the polar stratosphere,
reaction Cycles 2 and 3 cannot occur. It is only in late winter/early spring whensunlight returns to the polar regions that these cycles are active. Therefore, thegreatest destruction of ozone occurs in the partially to fully sunlit periods after
midwinter in the polar stratosphere. The visible sunlight needed in Cycles 2 and 3 is
not sufficient to form ozone because this process requires solar ultraviolet radiation(see Q1). In the late winter/early spring stratosphere, solar ultraviolet radiation is
weak at low Sun angles. As a result, ozone destruction by Cycles 2 and 3 in the sunlit
winter stratosphere greatly exceeds ozone production.
Other reactions. Global ozone abundances are controlled by many reactions
that both produce and destroy ozone (see Q2). Chlorine and bromine catalytic
reactions are but one group of ozone destruction reactions. Reactive hydrogenand reactive nitrogen gases, for example, are involved in other catalytic ozonedestruction cycles that also occur in the stratosphere. These reactions occur naturally
in the stratosphere and their importance has not been as strongly influenced byhuman activities as have reactions involving halogens. The sources of reactivehydrogen and nitrogen gases are the naturally occurring gases methane (CH4) andnitrous oxide (N2O), respectively. The importance of reactive hydrogen and nitrogengases relative to reactive halogen gases is expected to increase in the future,
because the atmospheric abundances of the reactive halogen gases are decreasing
as a result of the Montreal Protocol, while CH4 and N2O abundances (which are notcontrolled under the Montreal Protocol) are projected to increase substantially due
to human activities.

10

Why has an “ozone hole” appeared over
Antarctica when ozone-depleting substances
are present throughout the stratosphere?
Ozone-depleting substances are present throughout the stratosphericozone layer because they are transported great distances byatmospheric air motions. The severe depletion of the Antarctic ozone
layer known as the “ozone hole” occurs because of the special
meteorological and chemical conditions that exist there and nowhere
else on the globe. The very low winter temperatures in the Antarctic
stratosphere cause polar stratospheric clouds (PSCs) to form. Special
reactions that occur on PSCs, combined with the isolation of polar
stratospheric air in the polar vortex, allow chlorine and bromine
reactions to produce the ozone hole in Antarctic springtime.
The severe depletion of stratospheric ozone in late winter and early spring in the
Antarctic is known as the “ozone hole” (see Q11). The ozone hole appears over
Antarctica because meteorological and chemical conditions unique to this region
increase the effectiveness of ozone destruction by reactive halogen gases (see Q8).
In addition to an abundance of these reactive gases, the formation of the Antarctic

Minimum
Temperatures inStratosphere
Minimum
Air Air
Temperatures
inthe
thePolar
Polar
Stratosphere

-50
-55

Arctic winter
Jan
Feb

Dec

March

April
-60

Average daily values (40º to 90º latitude)
Arctic 1979/80 to 2013/14
Antarctic 1979 to 2013-70

Ranges of daily values

-60

-80

PSC formation temperature

-65
-90

-70

-100

-75

Arctic

-110

-80

-120

-85
Antarctic-90

Temperature (degrees Fahrenheit)

Nov

Temperature (degrees Celsius)

Figure Q10-1. Arctic and Antarctic
temperatures.
Air temperatures in both polar regions
reach minimum values in the lower
stratosphere in the winter season.
Average daily minimum values over
Antarctica are as low as −90°C in Julyand August in a typical year. Over theArctic, average minimum values are near
−80°C in late December and January.
Polar stratospheric clouds (PSCs) areformed in the polar ozone layer when
winter minimum temperatures fall below
the formation temperature of about−78°C. This occurs on average for 1 to 2
months over the Arctic and 5 to 6 months
over Antarctica (see heavy red and bluelines). Reactions on liquid and solid
PSC particles cause the highly reactive
chlorine gas ClO to be formed, whichcatalytically destroys ozone (see Q9). The
range of winter minimum temperatures
found in the Arctic is much greater thanin the Antarctic. In some years, PSC
formation temperatures are not reachedin the Arctic, and significant ozone
depletion does not occur. In contrast,
PSC formation temperatures are always
present for many months somewhere in
the Antarctic, and severe ozone depletion
now occurs in each winter season (seeQ11). (Note that the dashed black lines
denote the upper limits of the Antarctic
temperature range where they overlap
with the Arctic temperature range.)

-130

-95

-140
May

June

July
AugustAntarctic winter

Sept

Oct

Section II: The ozone depletion process33

10
ozone hole requires temperatures low enough to form polar stratospheric clouds(PSCs), isolation from air in other stratospheric regions, and sunlight.
Distribution of halogen gases. Halogen source gases that are emitted at Earth’s
surface and have lifetimes longer than about 1 year are present in comparable
abundances throughout the stratosphere in both hemispheres, even though most of
the emissions occur in the Northern Hemisphere. The abundances are comparable
because most long-lived source gases have no significant natural removal processes
in the lower atmosphere, and because winds and convection redistribute and mix
air efficiently throughout the troposphere on the timescale of weeks to months.
Halogen gases (in the form of source gases and some reactive products) enter the
stratosphere primarily from the tropical upper troposphere. Stratospheric air motions
then transport these gases upward and toward the pole in both hemispheres.
Low polar temperatures. The severe ozone destruction that leads to the ozone holerequires low temperatures to be present over a range of stratospheric altitudes,
over large geographical regions, and for extended time periods. Low temperatures
are important because they allow liquid and solid PSCs to form. Reactions on thesurfaces of these PSCs initiate a remarkable increase in the most reactive chlorine
gas, chlorine monoxide (ClO) (see below and Q8). Stratospheric temperatures
are lowest in the polar regions in winter. In the Antarctic winter, minimum daily
temperatures are generally much lower and less variable than those in the Arctic
winter (see Figure Q10-1). Antarctic temperatures also remain below PSC formation
temperatures for much longer periods during winter. These and other meteorologicaldifferences occur because of the differences in the distribution of land, ocean,and mountains between the hemispheres at middle and high latitudes. Wintertemperatures are low enough for PSCs to form somewhere in the Antarctic for nearly
the entire winter (about 5 months), but only for limited periods (10–60 days) in the
Arctic in most winters.
Isolated conditions. Stratospheric air in the polar regions is relatively isolated for long
periods in the winter months. The isolation is provided by strong winds that encircle
the poles during winter, forming a polar vortex, which prevents substantial transportand mixing of air into or out of the polar stratosphere. This circulation strengthensin winter as stratospheric temperatures decrease. Since winter temperatures arelower in the Southern than in the Northern Hemisphere polar stratosphere, theisolation of air in the vortex is much more effective in the Antarctic than in the Arctic.
Once temperatures drop low enough, PSCs form within the polar vortex and induce
chemical changes that are preserved by the isolation for many weeks to months.
Polar stratospheric clouds (PSCs). Reactions on the surfaces of liquid and solid
PSCs can substantially increase the relative abundances of the most reactive chlorine
gases. These reactions convert the reservoir forms of reactive chlorine gases,
chlorine nitrate (ClONO2) and hydrogen chloride (HCl), to the most reactive form, ClO(see Figure Q8-1). ClO increases from a small fraction of available reactive chlorine tocomprise nearly all chlorine that is available. With increased ClO, the catalytic cycles
involving ClO and BrO become active in the chemical destruction of ozone wheneversunlight is available (see Q9).
Different types of liquid and solid PSC particles form when stratospherictemperatures fall below about −78°C (−108°F) in polar regions (see Figure Q10-1).
As a result, PSCs are often found over large areas of the winter polar regions andover significant altitude ranges, with larger regions and for longer time periods in
the Antarctic than in the Arctic. The most common type of PSC forms from nitric acid
(HNO3) and water condensing on pre-existing liquid sulfuric acid-containing particles.
Some of these particles freeze to form solid particles. At even lower temperatures

34

Section II: The ozone depletion processArctic Polar Stratospheric
Clouds (PSCs)(−85°C or −121°F), water condenses to form ice particles. PSC particles grow large
enough and are numerous enough that cloud-like features can be observed from theground under certain conditions, particularly when the Sun is near the horizon (seeFigure Q10-2). PSCs are often found near mountain ranges in polar regions because
the motion of air over the mountains can cause localized cooling in the stratosphere,
which increases condensation of water and HNO3.
When average temperatures begin increasing by late winter, PSCs form less frequently
and conversion reactions on their surfaces produce less ClO. Without continued ClO
production, ClO amounts decrease and other chemical reactions re-form the reservoirs
gases, ClONO2 and HCl. When temperatures rise above PSC formation thresholds,
which typically occurs by late January to early February in the Arctic or by mid-October inthe Antarctic, the most intense period of ozone depletion ends.

Figure Q10-2. Polar stratosphericclouds.
This photograph of an Arctic polar
stratospheric cloud (PSC) was takenfrom the ground at Kiruna, Sweden(67°N), on 27 January 2000. PSCs formin the ozone layer during winters in
the Arctic and Antarctic stratospheres
wherever low temperatures occur (seeFigure Q10-1). The particles grow fromthe condensation of water and nitricacid (HNO3). The clouds often can beseen with the human eye when the Sun
is near the horizon. Reactions on PSCs
cause the highly reactive chlorine gas
ClO to be formed, which is very effectivein the chemical destruction of ozone(see Q9).

Nitric acid and water removal. Once formed, PSC particles fall to lower altitudes
because of gravity. The largest particles can descend several kilometers or more in
the stratosphere within a few days during the low-temperature winter/spring period.
Because PSCs often contain a significant fraction of available HNO3, their descentremoves HNO3 from regions of the ozone layer. This process is called denitrificationof the stratosphere. Because HNO3 is a source for nitrogen oxides (NOx) in thestratosphere, 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 increasing chemical ozone destruction.
Significant denitrification occurs each winter in the Antarctic but only in occasional
winters in the Arctic, because PSC formation temperatures must be sustained over an
extensive altitude region and time period to lead to denitrification (see Figure Q10-1).
Ice particles form at temperatures that are a few degrees lower than those requiredfor PSC formation from HNO3. If ice particles grow large enough, they can fall several
kilometers due to gravity. As a result, a significant fraction of water vapor can beremoved from regions of the ozone layer over the course of a winter. This process
is called dehydration of the stratosphere. Because of the very low temperatures
required to form ice, dehydration is common in Antarctic but rare in Arctic winters.
The removal of water vapor does not directly affect the catalytic reactions that destroyozone. Dehydration indirectly affects ozone destruction by suppressing PSC formation
later in winter, which reduces ClO production in PSC reactions.
Discovering the role of PSCs. Ground-based observations of PSCs were available
many years before the role of PSCs in polar ozone destruction was recognized. The
geographical and altitude extent of PSCs in both polar regions was not known fully
until PSCs were observed by a satellite instrument in the late 1970s. The role of PSC
particles in converting reactive chlorine gases to ClO was not understood until after
the discovery of the Antarctic ozone hole in 1985. Our understanding of the chemicalrole of PSC particles developed from laboratory studies of their surface reactivity,
computer modeling studies of polar stratospheric chemistry, and measurementsthat directly sampled particles and reactive chlorine gases, such as ClO, in thepolar stratosphere.

Section II: The ozone depletion process

35

10
The Discovery of the Antarctic Ozone HoleThe first decreases in Antarctic total ozone were observed in the early 1980s overresearch stations located on the Antarctic continent. The measurements were madewith ground-based Dobson spectrophotometers (see box in Q5). The observations
showed unusually low total ozone during the late winter/early spring monthsof September, October, and November. Total ozone was lower in these months
compared with previous observations made as early as 1957. The early publishedreports came from the Japan Meteorological Agency and the British Antarctic
Survey. The results became widely known to the world after three scientists fromthe British Antarctic Survey published their observations in the scientific journal
Nature in 1985 and speculated that CFCs were the cause. Soon after, satellite
measurements confirmed the spring ozone depletion and further showed that ineach late winter/early spring season starting in the early 1980s, the depletion
extended over a large region centered near the South Pole. The term “ozone hole”
came about as a description of the very low total ozone values that encircled theAntarctic continent in satellite images for many weeks (see Q11). Currently, theformation 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 found in spring were around 300 Dobson units (DU),
surprisingly lower than in the Arctic spring, because the assumption then was thatthe two polar regions would have similar values. We now know that these Antarctic
values were not anomalous; indeed they are similar to those observed there in the
1970s before the ozone hole appeared (see Figure Q11-3). We also now know that
Antarctic total ozone values, in comparison with Arctic values, are systematicallylower in the early spring because the polar vortex is much stronger and, therefore,much more effective in reducing the transport of ozone-rich air from midlatitudes
to the pole (compare Figures Q11-3 and Q12-2).
In 1958, measurements of total ozone were made at the Dumont d’Urville station(66.7°S, 140°E) in Antarctica using a photographic plate method to analyze
solar ultraviolet radiation after it passed through the ozone layer. The reported
measurements were anomalously low, reaching 110–120 DU in September andOctober. These values are similar to minimum ozone hole values now routinely
observed over Antarctica in the same months (see Figure Q11-2). Some havespeculated that these limited observations provide evidence that an ozone holeexisted before ODS emissions were large enough to cause the depletion. However,analyses of the more extensive Dobson spectrophotometer measurements made atseveral other Antarctic locations in 1958 did not confirm the low total ozone values.
These measurements indicate that the photographic plate determinations were nota reliable source of total ozone values at the Dumont d’Urville station in 1958.

36

Section II: The ozone depletion process

11

How severe is the depletionof the Antarctic ozone layer?
Severe depletion of the Antarctic ozone layer was first reported in the
mid-1980s. Antarctic ozone depletion is seasonal, occurring primarily
in late winter and early spring (August–November). Peak depletionoccurs in early October when ozone is often completely destroyed
over a range of altitudes, thereby reducing total ozone by as muchas two-thirds at some locations. This severe depletion creates the“ozone hole” apparent in images of Antarctic total ozone made using
satellite observations. In most years the maximum area of the ozone
hole far exceeds the size of the Antarctic continent.
The severe depletion of Antarctic ozone, known as the “ozone hole,” was first reported
in the mid-1980s. The depletion is attributable to chemical destruction by reactive
halogen gases that increased in the stratosphere in the latter half of the 20th century(see Q16). Conditions in the Antarctic winter and early spring stratosphere enhanceozone depletion because of (1) the long periods of extremely low temperatures, whichcause polar stratospheric clouds (PSCs) to form; (2) the large abundance of reactive
halogen gases produced in reactions on PSCs; and (3) the isolation of stratosphericair, which allows time for chemical destruction processes to occur. The long-termchanges and severity of Antarctic ozone depletion can be seen using satellite
observations of total ozone and ozone altitude profiles.
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 Ozone Hole
Antarctic
Ozone Hole

Figure Q11-1. Antarctic ozone hole.
Total ozone values are shown for high southern latitudes on
14 September 2013 as measured by a satellite instrument. The
dark blue and purple regions over the Antarctic continent show
the severe ozone depletion or “ozone hole” now found during
every spring. Minimum values of total ozone inside the ozone
hole are close to 100 Dobson units (DU) compared with normal
Antarctic springtime values of about 350 DU (see Figure Q113). The ozone hole area is usually defined as the geographicalarea within the 220-DU contour (see white line) on total ozone
maps. In late spring or early summer (November–December) theozone hole disappears in satellite images as ozone-depleted air
is displaced and mixed with ozone-rich air transported toward thepole from outside the ozone hole. Note that maximum total ozone
values in the Southern Hemisphere in this late winter period are
generally located in a crescent-shaped region surrounding theozone hole.

14 September 2013
100

200
300
400
500
Total ozone (Dobson units)

600

Section III: Stratospheric ozone depletion

37

11
Antarctic early spring measurements shows a large region centered near the South Polein which total ozone is highly depleted (see Figure Q11-1). This region has come to becalled the “ozone hole” because of the near-circular contours of low ozone values in
the maps. The area of the ozone hole is defined here as the geographical region within
the 220-Dobson unit (DU) contour in total ozone maps (see white line in Figure Q11-1)
averaged between 21–30 September for each year. The area reached a maximum of
27 million square kilometers (about 10 million square miles) in 2006, which is nearly
twice the area of the Antarctic continent (see Figure Q11-2). Minimum values of totalozone inside the ozone hole averaged in late September to mid-October are near
110 DU, which is two-thirds below normal springtime values of about 350 DU (seeFigures Q11-3 and Q12-1). Low total ozone inside the ozone hole contrasts strongly
with the distribution of much larger values outside the ozone hole. This common
feature can be seen in Figure Q11-1, where a crescent-shaped region with values
around 400 DU surrounds a large portion of the ozone hole in September 2013.
Altitude profiles of Antarctic ozone. The low total ozone values within the ozone
hole are caused by nearly complete removal of ozone in the lower stratosphere.
Balloonborne instruments (see Q5) demonstrate that this depletion occurs within theozone layer, the altitude region that normally contains the highest ozone abundances.
At geographic locations with the lowest total ozone values, balloon measurements showthat the chemical destruction of ozone is complete over an altitude region of several
kilometers. For example, in the ozone profile over South Pole, Antarctica, on 9 October
2006 (see red line in left panel of Figure Q12-3), ozone abundances are essentiallyzero over the altitude region of 14 to 21 kilometers. The lowest winter temperaturesand highest reactive chlorine (ClO) abundances occur in this altitude region (see FigureQ8-3). The differences in the average South Pole ozone profiles between the decades

38

Section III: Stratospheric ozone depletion

Antarctic Ozone Depletion

Area (million square
kilometers)

Antarctic Ozone Depletion
30

AfricaNorth
America

Average daily area
(21–30 September, 220 DU)

SouthAmerica20

Antarctica
Europe10Australia

0

Minimum total ozone(Dobson units)

Figure Q11-2. Antarctic ozone
hole features.
Long-term changes are shown for key
aspects of the Antarctic ozone hole: thearea enclosed by the 220-DU contour
on maps of total ozone (upper panel)and the minimum total ozone amountmeasured south of 40°S (lower panel).
The values are derived from satellite
observations and averaged for eachyear at a time near the peak of ozone
depletion, as defined by the dates
shown in each panel. The areas ofcontinents are included for referencein the upper panel. The magnitude ofozone depletion gradually increased
beginning in 1980. In the past two
decades, the depletion reached steadyyear-to-year values, except for theunusually low depletion in 2002 (seeFigure Q11-4 and following box). The
magnitude of Antarctic ozone depletion
will decrease as ODSs are removed fromthe atmosphere (see Figure Q16-1). The
return of Antarctic total ozone to 1980
values is not expected to occur before
2050 (see Q20).

200
150
100

Average daily minimum(21 September–16 October)

1980

1990

2000

Year

2010

2020

1962–1971 and 1990–2013 in Figure Q12-3 show how reactive halogen gases havedramatically altered the ozone layer. In the 1960s, a normal ozone layer is clearly
evident in the October average profile, with a peak near 16 kilometers altitude.
In the 1990–2013 average profile, a broad minimum centered near 16 kilometers
now occurs, with ozone values reduced by up to 90% of normal values.
Long-term total ozone changes. The significant springtime ozone depletion
represented by the ozone hole has only been observed since the early 1980s eventhough low winter temperatures and isolated conditions occur each year in the
Antarctic stratosphere. Prior to 1980, the amounts of reactive halogen gases in thestratosphere were insufficient to cause a significant signature of depletion in theface of strong interannual variations in stratospheric transport. Satellite observations
can be used in multiple ways to examine how ozone depletion has changed in the
Antarctic region over the last 30–40 years:• First, ozone hole areas displayed in Figure Q11-2 show that depletion has
increased since 1980 to become fairly stable in the 1990s, 2000s, and even into
the early 2010s near a value of 23 million square kilometers. The exception is theunexpected low depletion in 2002, which is explained in the box at the end of thisQuestion. The ozone-hole area is defined here as the geographical area inside the220-DU contour on total ozone maps (see Figure Q11-1).
• Second, minimum Antarctic ozone amounts displayed in Figure Q11-2 show thatthe severity of the depletion increased beginning around 1980 along with the ozone
hole area. Fairly constant minimum values near 110 DU were observed in the
1990s, 2000s, and early 2010s, with 2002 being an exception.
• Third, total ozone maps over the Antarctic and surrounding regions show how
the ozone hole has developed over time (see Figure Q11-3). October averages of
total ozone confirm the absence of an ozone hole in the 1970s and its persistentoccurrence in the late 2000s and early 2010s.
Figure Q11-3. Antarctic total ozone.
Long-term changes in Antarctic totalozone are demonstrated with this seriesof total ozone maps derived from satellite
observations. Each map is an averageduring October, the month of maximum
ozone depletion over Antarctica. In the
1970s, no ozone hole was observed as
defined by a significant region with totalozone values less than 220 DU (darkblue and purple colors). Starting in the
1980s, the ozone hole began to appear
and increased in size (not shown here). A
large ozone hole has occurred each year
since the early 1990s as shown in Figure
Q11-2. The maps from the late 2000s
and early 2010s show the large extent(about 25 million square kilometers) of
the most recent ozone holes. The largest
total-ozone values at high southernlatitudes are still found in a crescentshaped region outside the ozone holein October. The maps show that thesemaximum values and their extent havesignificantly diminished since the 1970s.

Antarctic
Ozone
AntarcticTotal
Total Ozone(Octobermonthly
monthly averages)(Octoberaverages)

1970

1971

1972

1979

2007

2009

2011

2013

100

200
300
400
500
Total ozone (Dobson units)

600

Section III: Stratospheric ozone depletion

39

11
Unusual 2002 Antarctic Ozone Hole

Unusual 2002 Antarctic Ozone Hole24 September 2001

24 September 2002
100

200
300
400
500
Total ozone (Dobson units)

24 September 2003
600

Figure Q11-4. Unusual 2002 ozone hole.
Views from space of the Antarctic ozone hole are shown for 24 September in theyears 2001, 2002, and 2003. The ozone holes in 2001 and 2003 are considered
typical of those observed since the early 1990s. An initially circular hole in 2002 wastransformed into two smaller, depleted regions in the days preceding 24 September.
This unusual event is attributable to an early warming of the polar stratosphere causedby meteorological disturbances that originated in the troposphere at midlatitudes.
Warmer temperatures reduced the rate of ozone depletion. As a consequence, totalozone depletion was unusually low that year in comparison with 2001 and 2003 andall other years since the early 1990s (see Figure Q11-2).

• Fourth, values of total ozone poleward of 63°S averaged for each Octobershow how total ozone has changed when averaged over the ozone hole and its
surrounding areas (see Figure Q12-1). The values decreased strongly starting fromthose found in the 1970s and now are approximately 30% less than in pre-ozonehole years (1970–1982). The average values show a larger year-to-year variability
than found for the other ozone measures noted above because the average
includes areas outside the ozone hole, where the distribution of total ozone varies
in response to meteorological conditions. Increased year-to-year variability isevident over the past decade. The maps in Q11-3 show also how the maximum
in total ozone surrounding the ozone hole each year has diminished over the last
three decades, adding to the decreases noted in Figure Q12-1.
Disappearance of the ozone hole in spring. The severe depletion of Antarctic ozone
occurs in the late winter/early spring season. In spring, temperatures in the polar
lower stratosphere increase (see Figure Q10-1), stopping the formation of PSCs andproduction of ClO and, consequently, the most effective chemical cycles that destroyozone. The polar vortex breaks down, ending the wintertime isolation of high-latitude
air and increasing the exchange of air between the Antarctic stratosphere andlower latitudes. This allows substantial amounts of ozone-rich air to be transported
poleward, where it displaces or mixes with air depleted in ozone. The midlatitude air
also contains higher abundances of nitrogen oxide gases (NOx), which help convert
the most reactive chlorine gases (ClO) back into the chlorine reservoir gas ClONO2(see Q6). As a result of these large-scale transport and mixing processes, the ozone
hole disappears by December.

40

Section III: Stratospheric ozone depletion

The 2002 Antarctic Ozone HoleThe 2002 Antarctic ozone hole showed features that looked surprising at the time(see Figure Q11-4). It had much less ozone depletion as measured by the area of theozone hole or minimum total-ozone amounts in comparison with the 2001 ozone hole.
The 2002 values now stand out clearly in the year-to-year changes in these quantities
displayed in Figure Q11-2. There were no forecasts of an ozone hole with unusual
features in 2002 because the chemical and meteorological conditions required todeplete ozone, namely low temperatures and available reactive halogen gases, were
present that year and did not differ substantially from previous years. The ozone
hole initially formed as expected in August and early September 2002. Later, during
the last week of September, an unexpected and surprisingly strong meteorologicalevent occurred that dramatically reshaped the ozone hole into two separate depleted
regions. As a result of this disturbance, the combined area of these two regions in late
September and early October was significantly less than that observed for the previous
or subsequent ozone holes.
The unexpected meteorological influence in 2002 resulted from specific atmospheric
air motions that sometimes occur in polar regions. Meteorological analyses of the
Antarctic stratosphere show that it was warmed by very strong, large-scale weather
systems that originated in the lower atmosphere (troposphere) at midlatitudes inlate September. At that time, Antarctic temperatures are generally very low (seeQ10) and ozone destruction rates are near their peak values. The influence of these
tropospheric systems extended poleward and upward into the stratosphere, disturbing
the normal circumpolar wind flow (polar vortex) and warming the lower stratosphere
where ozone depletion was ongoing. Higher temperatures reduced the rate of ozone
depletion and led to the higher minimum values observed for total ozone in Figure
Q11-2. The higher-than-normal impact of these weather disturbances during thecritical time period for ozone loss reduced the total loss of ozone in 2002.
The strong influence of the 2002 warming event is unique in the many decades ofAntarctic meteorological observations. Another warming event occurred in 1988 causing
smaller changes in the ozone hole features in Figure Q11-2. Large warming events are
difficult to predict because of the complex conditions leading to their formation.
In 2003 through 2009, ozone hole features returned to values observed from themid-1990s to 2001 (see Figure Q11-2). The high ozone depletion found since themid-1990s, with the exception of 2002, is expected to be typical of coming years.
A significant, sustained reduction of Antarctic ozone depletion, leading to fullrecovery of total ozone, requires comparable, sustained reductions of ODSs in thestratosphere. Even with the source gas reductions already underway (see Q16), thereturn of Antarctic total ozone to 1980 values is not expected to occur before 2050.

Section III: Stratospheric ozone depletion

41

12

Is there depletion of the Arctic ozone layer?
Yes, significant depletion of the Arctic ozone layer now occurs inmost years in the late winter/early spring period (January–March).
However, the maximum depletion is less severe than that observedin the Antarctic and with large year-to-year differences as aconsequence of the highly variable meteorological conditions foundin the Arctic polar stratosphere. Even the most severe Arctic ozone
depletion does not lead to total ozone amounts as low as thoseseen in the Antarctic, because Arctic ozone abundances during
early winter before the onset of ozone depletion are much higher
than those in the Antarctic. A large and recurrent “ozone hole,” asfound 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 to chemical destruction by reactive halogen
gases, which increased in the stratosphere in the latter half of the 20th century (seeQ16). Arctic depletion also occurs in the late winter/early spring period (January–
March), but over a somewhat shorter period than in the Antarctic (July–October).
Similar to the Antarctic (see Q11), Arctic depletion occurs because of (1) the periodsof extremely low temperatures, which cause polar stratospheric clouds (PSCs) toform; (2) the large abundance of reactive halogen gases produced in reactions on
PSCs; and (3) the isolation of polar stratospheric air, which allows time for chemical
destruction processes to occur.
Arctic ozone depletion is much less than that observed each Antarctic winter/spring
season. Large and recurrent ozone holes as found in the Antarctic stratosphere
do not occur in the Arctic. Stratospheric ozone abundances during early winter,
before the onset of ozone depletion, are naturally higher in the Arctic than in the
Antarctic because transport of ozone from its source region in the tropics to higher
latitudes is more vigorous in the Northern Hemisphere. Depletion is limited because,in comparison to Antarctic conditions, Arctic average temperatures are always
significantly higher (see Figure Q10-1) and the isolation of polar stratospheric air
is less effective. Temperature and other meteorological differences occur because
northern polar latitudes have more land and mountainous regions than southern
polar latitudes (compare Figures Q11-3 and Q12-2). In a few Arctic winters, forexample, PSCs did not form because temperatures never reached low enough values.
These differences cause the extent and timing of Arctic ozone depletion to varyconsiderably from year to year. Depletion in some winter/spring seasons occurs overmany weeks, in others only for brief early or late periods, and in some not at all.
Long-term total ozone changes. Satellite observations can be used in two importantways to examine the average total ozone abundances in the Arctic region for the last
30–40 years and to contrast these values with Antarctic abundances:• First, total ozone poleward of 63°N averaged for each March shows quantitatively
how total ozone has changed in the Arctic (see Figure Q12-1). The seasonalpoleward and downward transport of ozone-rich air is naturally stronger in the
Northern Hemisphere. As a result, total ozone values at the beginning of each
winter season in the Arctic are considerably higher than those in the Antarctic.
Before depletion sets in, normal Arctic values are close to 450 DU and Antarctic
values close to 330 DU. Decreases from pre-ozone-hole average values (1970–1982) were observed in the Arctic by the mid-1980s, when larger changes werealready occurring in the Antarctic. The decreases in total ozone are generally muchsmaller than those found in the Antarctic and lead to total ozone values that are

42

Section III: Stratospheric ozone depletion

AverageTotal
TotalOzone
Ozone inin Polar
Average
PolarRegions
Regions
500

Arctic region in March

Total ozone (Dobson units)

450
400
350

Antarctic region in October

300
Chemical destruction
and natural variationof ozone250

1970–1982 Average(over 63°–90°)

200
1970

1980

1990

2000

2010

Year
Figure Q12-1. Average total ozone in polar regions.
Long-term changes in average total ozone are shown for the Antarctic and Arctic,
defined by latitudes between 63° and 90°. Total ozone is measured with satellite
instruments. The reference values (red lines) are averages of springtime total ozone
values available from observations between 1970 and 1982. Each point represents
a monthly average for October in the Antarctic or March in the Arctic. 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 are the combination of chemical destruction and natural variations.
Variations in meteorological conditions influence the year-to-year changes in ozone,
particularly in the Arctic. Natural variations have also increased in the Antarctic since
2000. Essentially all of the decrease in the Antarctic and usually about 50% of thedecrease in the Arctic each year are attributable to chemical destruction by reactive
halogen gases. In the Arctic, the other 50% is attributable to natural variations in theamounts of ozone transported to polar regions before and during winter. Average totalozone values over the Arctic are naturally larger at the beginning of each winter season
because, in the preceding months, more ozone is transported poleward in the Northern
Hemisphere than in the Southern Hemisphere.

on average around 10–15% below pre-ozone-hole values. Maximum decreases in
Arctic total ozone of about 30% observed in March 1997 and 2011 are the mostcomparable to Antarctic depletion. In both these Arctic winters, meteorologicalconditions kept ozone transport to high latitudes below average values, and in 2011
persistently low temperatures facilitated severe chemical ozone depletion by ODSs.
Overall, Arctic total ozone values tend to show larger year-to-year variability thanthose in the Antarctic. Ozone differences from the 1970–1982 average value are
due to a combination of chemical destruction by ODSs and meteorological (natural)variations. In the last two decades, these two aspects have contributed about
equally to observed ozone changes. The amount of chemical destruction depends
in large part on stratospheric temperatures. Meteorological conditions determine
how well Arctic stratospheric air is isolated from ozone-rich air at lower latitudes
and influence the extent of low temperatures.

Section III: Stratospheric ozone depletion

43

12
Figure Q12-2. 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 depletionis usually observed in the Arctic. In the
1970s, the Arctic region had normalozone values in March, with values of
450 DU and above (red colors). Ozone
depletion on the scale of the Antarctic
ozone hole does not occur in the Arctic.
Instead, late winter/early spring ozone
depletion has eroded the normal highvalues of total ozone. On the maps fromthe late 2000s and early 2010s, theextent of values of 450 DU and aboveis greatly reduced in comparison withthe 1970s maps. The large regions of
low total ozone in 1997 and 2011 (bluecolors) are unusual in the Arctic record,
but not unexpected. The meteorologicalconditions led to below-average
stratospheric temperatures and a strong
polar vortex in these winters, conditions
favorable to strong ozone depletion.

Arctic Total Ozone(Marchmonthly
ArcticTotal averages)
Ozone(March monthly averages)

1971

1972

1979

1980

1997

2008

2011

2014

100

200
300
400
500
Total ozone (Dobson units)

600• Second, total ozone maps over the Arctic and surrounding regions (see Figure
Q12-2) show year-to-year changes in March total ozone. In the 1970s, total ozone
values were near 450 DU when averaged over the Arctic region in March. Beginningin the 1990s and continuing into the early 2010s, values above 450 DU wereincreasingly absent from the March average maps. A comparison of the maps in
the 1970s and early 2010s, for example, shows a striking reduction of total ozone
throughout the Arctic region. The large geographical extent of low total ozone in themaps of March 1997 and March 2011 represent exceptional events in the Arctic
observational record of the last three decades as noted above in the discussion of
Figure Q12-1.
Altitude profiles of Arctic ozone. Arctic ozone is measured using a variety of
instruments (see Q5), as in the Antarctic, to document daily to seasonal changes
within the ozone layer. Spring Arctic and Antarctic balloonborne measurements
are contrasted in Figure Q12-3. Arctic profiles were obtained from the Ny-Ålesund
research station at 79°N. For 1989–2014, the March average reveals a substantial
ozone layer, contrasting sharply with the severely depleted Antarctic ozone layer in
the October average over a similar time period. This further demonstrates how higher
stratospheric temperatures and meteorological variability have protected the Arctic
ozone layer from the greater ozone losses that occur in the Antarctic, despite similar
reactive halogen abundances in the two regions.
The separate Arctic profiles shown for 29 March 1996 and 1 April 2011 are two of
the most severely depleted in the 25-year long record from Ny-Ålesund. Althoughsignificant, the depletion during both events is smaller in comparison to that routinely
observed in the Antarctic, such as in the profile from 9 October 2006. The nearcomplete depletion of ozone over many kilometers in altitude and over areas almost

44

Section III: Stratospheric ozone depletionas large as North America, as is now common in the Antarctic stratosphere, has
never been observed in the Arctic.
Restoring ozone in spring. As in the Antarctic, ozone depletion in the Arctic isconfined to the late winter/early spring season. In spring, temperatures in thepolar lower stratosphere increase (see Figure Q10-1), halting PSC formation andproduction of ClO, and thus the most effective chemical cycles that destroy ozone.
The breakdown of the polar vortex ends the wintertime isolation of high-latitude
air, increasing exchange of air between the Arctic stratosphere and lower latitudes.
This allows more ozone-rich air to be transported poleward, where it displaces ormixes with air in which ozone may have been depleted. As a result of these largescale transport and mixing processes, any ozone depletion typically disappears
by April or earlier.

Polar
Ozone Depletion
Polar Ozone Depletion
30

Antarctic ozone

Arctic ozone

South Pole (90°S)

Ny-Ålesund (79°N)

25

15

10

Altitude (km)Altitude (miles)

20

15

10
5

March averages

October averages
51962 – 1971 (no data)
1989 – 2014 (380 DU)

1962 – 1971 (298 DU)
1990 – 2013 (153 DU)

29 March 1996 (258 DU)
1 April 2011 (260 DU)

9 October 2006 (93 DU)
0

0
0

5

10

15 0

5

10

15

Ozone abundance (mPa)

Figure Q12-3. Vertical distribution of Arctic and Antarctic ozone.
Most stratospheric ozone resides between about 10 and 30 kilometers (6 to 19 miles)
above Earth’s surface. Long-term observations of the ozone layer with balloonborne
instruments allow winter ozone altitude profiles 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 1962 and 1971. In the past two decades, as
shown here for 9 October 2006, ozone is almost completely destroyed between 14 and21 kilometers (9 to 13 miles) in the Antarctic in spring. Average October values in thelast decades (1990–2013) are 90% lower than pre-1980 values at the peak altitudeof the ozone layer (16 kilometers). In contrast, the Arctic ozone layer is still present inspring as shown by the average March profile for 1989–2014 obtained over the NyÅlesund site (right panel). No Ny-Ålesund data are available for the 1962–1971 period
before significant ODS destruction began. Some March profiles do reveal significant
depletion, as shown here for 29 March 1996 and 1 April 2011. In such years, winterminimum temperatures are generally lower than normal, allowing PSC formation for
longer periods. Depletion similar to that shown for 9 October 2006 at the South Pole
has never been observed in the Arctic. The number in parentheses for each profile isthe total ozone value in Dobson units (DU). Ozone abundances are shown here as thepressure of ozone at each altitude using the unit “milli-Pascals” (mPa) (100 million mPa= atmospheric sea-level pressure).
Section III: Stratospheric ozone depletion

45

13

How large is the depletionof the global ozone layer?
Depletion of the global ozone layer increased gradually in the
1980s and reached a maximum of about 5% in the early 1990s.
The depletion has lessened since then and now is about 3%
averaged over the globe. The average depletion exceeds the natural
year-to-year variations of global total ozone. The ozone loss is verysmall near the equator and increases with latitude toward thepoles. The larger polar depletion is attributed to the late winter/
early spring ozone destruction that occurs there each year.
Global total ozone has decreased beginning in the 1980s (see Figure Q13-1). The
decreases have occurred in the stratospheric ozone layer where most ozone resides (seeFigure Q1-2). In the early 1990s, the depletion of global total ozone reached a maximumof about 5% below the 1964–1980 average. The depletion has lessened since thenand during the early 2010s has averaged to about 3% below the 1964–1980 average.
The observations shown in Figure Q13-1 have been smoothed to remove regular ozone
changes that are due to natural seasonal and solar effects (see Q14). The depleted
amounts are larger than the remaining natural variations in global total ozone amounts.
The observed global ozone depletion in the last three decades is attributable toincreases in reactive halogen gases in the stratosphere. The lowest global total ozone
values since 1980 have occurred in the years following the eruption of Mt. Pinatubo
in 1991, which temporarily increased the number of sulfuric acid-containing particles
throughout the stratosphere. These particles significantly increased the effectivenessof reactive halogen gases in destroying ozone (see Q14) and, thereby, increased
global ozone depletion by 1–2% for several years following the eruption.
Polar regions. Observed total ozone depletion varies significantly with latitude on
the globe (see Figure Q13-1). The largest reductions occur at high southern latitudes
as a result of the severe ozone loss over Antarctica each late winter/early spring
period. The next largest losses are observed in the high latitudes of the Northern
Hemisphere, caused in part by winter losses over the Arctic. Although the depletion inpolar regions is larger than at lower latitudes, the influence of polar regions on global
ozone is limited by their small geographical area. Latitudes poleward of 60° accountfor only about 13% of Earth’s surface.
Midlatitude regions. Ozone depletion is also observed at the midlatitudes
spanning the region between equatorial and polar latitudes. In comparison with the
1964–1980 averages, total ozone averaged for 2008–2012 is about 3.5% lower innorthern midlatitudes (35°N–60°N) and about 6% lower at southern midlatitudes(35°S–60°S). Midlatitude depletion has two contributing factors. First, ozonedepleted air over both polar regions is dispersed away from the poles during and after
each winter/spring period, thereby reducing average ozone outside polar regions.
Second, chemical destruction occurring at midlatitudes contributes to observed
depletion in these regions. This contribution is much smaller than in polar regions
because the amounts of reactive halogen gases are lower and a seasonal increase of
the most reactive halogen gases, such as the increase in ClO in the polar winter lower
stratosphere (see Figure Q8-3), does not occur in midlatitude regions.
Tropical region. Total ozone in the tropics (20°N–20°S latitude) has been only weakly
affected by chemical ozone depletion. In the tropical lower stratosphere, air has beentransported from the lower atmosphere (troposphere) within the previous 18 months. As
a result, the conversion of ozone-depleting substances (ODSs) to reactive halogen gases
is still very small. With little reactive halogen available, total ozone depletion in this regionis also very small. In addition, ozone production is high because average solar ultraviolet

46

Section III: Stratospheric ozone depletionradiation is highest in the tropics. In contrast, stratospheric air in polar regions has beenin the stratosphere for an average of 4 to 7 years, allowing time for significant conversion
of ODSs to reactive halogen gases. The systematic differences in the age of stratospheric
air are well understood and are a consequence of the large-scale atmospheric transport:
air enters the stratosphere in the tropics, moves poleward in both hemispheres, and thendescends and ultimately returns to the troposphere in the middle to high latitudes.

GlobalTotal
TotalOzone
Ozone Changes
GlobalChanges

Ozone change (%)

2

0

-2

-4

-6

Annual averages
Range of observations

1970

1980

1990
Year

2000

2010

Changes between 1964–1980
and 2008–2012 averages
0-5
Ozone change (%)

Figure Q13-1. Global total ozone changes.
Satellite observations show depletion of global total ozone
beginning in the 1980s. The top panel compares annual
averages of global ozone with the average from the period 1964
to 1980 before the ozone hole appeared. Seasonal and solar
effects have been removed from the observational data set. On
average, global ozone decreased each year between 1980 and
1990. The depletion worsened for a few years after 1991 dueto the effect of volcanic aerosol from the Mt. Pinatubo eruption(see Q14). Average global ozone for 2008–2012 is about 2.5%
below the 1964–1980 average. The bottom panel shows how
the 2008–2012 depletion varies with latitude over the globe.
The largest decreases have occurred at high latitudes in bothhemispheres because of the large winter/spring depletion inpolar regions. The losses in the Southern Hemisphere are greaterthan those in the Northern Hemisphere because of the Antarctic
ozone hole. Long-term changes in the tropics are much smaller
because reactive halogen gases are less abundant in the tropicallower stratosphere than at mid or high latitudes, and ozone
production rates are greater.

Changes from 1964–1980 average-10

Average
Uncertainty range

-15
-20
-25
-30
-90

South

-60

-30
0
30
Latitude (degrees)

60

90

North

Section III: Stratospheric ozone depletion47

13
Initial Signs of Ozone Recovery
The Montreal Protocol, strengthened by its Amendments and adjustments, has
successfully controlled the production and consumption of ODSs that act to destroy
the ozone layer (Q15). As a result, atmospheric abundances of ODSs have peaked andare now decreasing (Q7 and Q16). By 2012, equivalent effective stratospheric chlorine(EESC; the total chlorine and bromine abundances in the stratosphere) had declined
by 15% at midlatitudes from peak values of around 15 years ago. This raises thequestion, is global ozone increasing in response to the observed EESC decreases?
Identifying an ozone increase is not easy, because ODS levels are not the only factor
that determines global ozone levels. For example, the global ozone minimum wasobserved half a decade before the EESC maximum was reached. The difference inthe timing resulted from the strong global ozone response to enhanced stratosphericaerosol loading after the Mount Pinatubo eruption in 1991, which led to increased
ozone depletion for several years. Observed global ozone increases through the 1990s
were therefore a result of the steady removal of the aerosol from the stratosphere, andnot a sign of decreasing ODSs (see Q14). Another factor complicating the identificationof ozone recovery in different regions of the atmosphere is the year-to-year variationsof the stratospheric circulation. These variations lead to ozone variability in mostregions of the atmosphere that is currently still larger than the signal expected fromthe observed EESC decreases. Finally, greenhouse gas increases (such as carbon
dioxide, CO2) affect ozone by decreasing stratospheric temperatures (which slows
down ozone depletion rates) and by strengthening the stratospheric circulation (which
enhances the transport of ozone from the tropics to higher latitudes). It is therefore
difficult to attribute observed ozone changes to these different factors.
Observations now show a clear 5% increase of ozone in the upper stratosphere (42km) over the 2000-2013 period. Model simulations that allow for separation of thedifferent factors suggest that about half of this increase results from a cooling in thisregion due to CO2 increases, while the other half results from EESC decreases. Also,
total column ozone declined over most of the globe during the 1980s and early 1990s(by about 2.5% averaged over 60°S to 60°N). It has remained relatively unchanged
since 2000, with indications of a small increase in total column ozone in recent
years. Models suggest that this small increase is likely due to EESC decreases. Thesefindings based on both models and observations suggest that there are initial signs ofozone recovery.
Because of their long lifetime, the impact on stratospheric ozone of the mostprominent ODSs (CFC-11 and CFC-12) will continue for many decades after emissions
have ceased. Assuming continued compliance with the Montreal Protocol, EESC willcontinue to decline over the coming decades and will return to pre-1980 levels aroundmidcentury. With the exception of the tropics (see Q20), climate change is expected toaccelerate the return of the ozone layer to pre-1980 levels. However, as long as ODS
levels remain elevated in the atmosphere, the possibility of extreme low-ozone events
due to volcanic eruptions or cold winter conditions persists into the second half of the
21st century.

48

Section III: Stratospheric ozone depletion

14

Do changes in the Sun and volcanic
eruptions affect the ozone layer?
Yes, factors such as changes in solar radiation, as well as theformation of stratospheric particles after volcanic eruptions, do
influence the ozone layer. However, neither factor can explain the
average decreases observed in global total ozone over the last
three decades. If large volcanic eruptions occur in the coming
decades while ODS levels remain high, ozone depletion will
increase for several years afterward.
Changes in solar radiation and increases in stratospheric particles from volcanic
eruptions both affect the abundance of stratospheric ozone. Global total ozone in
the early 1990s had decreased by about 5% when compared to pre-1980 values,
and is now about 3% below pre-1980 values (see Q13). The depletion is primarily
attributed to increases in reactive halogen gases, with additional depletion in theearly 1990s associated with the volcanic eruption of Mt. Pinatubo. Reactive halogen
gases are represented by changes in equivalent effective stratospheric chlorine(EESC). EESC values account for stratospheric chlorine and bromine abundances andtheir different effectiveness in destroying ozone (see definition in Q16). A comparisonof the smooth year-to-year changes in ozone and EESC shows that the quantities areinversely related to each other, with ozone first decreasing while EESC increased (seeFigure Q14-1). After the mid-1990s, the annual changes in both quantities remained
constant and then slowly reversed sign. Changes in solar output and volcanic activity
do not show such smooth long-term changes, as discussed below, and therefore
cannot be the cause of long-term global ozone depletion.
Total ozone and solar changes. The formation of stratospheric ozone is initiatedby ultraviolet (UV) radiation coming from the Sun (see Figure Q2-1). As a result, an
increase in the Sun’s radiation output increases the amount of ozone in Earth’s
atmosphere. The Sun’s radiation output and sunspot number vary over the welldocumented 11-year solar cycle. Observations over several solar cycles since the1960s show that global total ozone levels vary by 1 to 2% between the maximum andminimum of a typical cycle. Changes in incoming solar radiation at a wavelength of
10.7 cm are often used as a surrogate for changes in solar output at UV wavelengths.
The long-term changes in the 10.7-cm output in Figure Q14-1 clearly show alternating
periods of maximum and minimum values in total solar output separated by about
5 to 6 years. If changes in solar output were the cause of global ozone depletion,a gradually decreasing output would have been observed around 1980 or earlier,slowing sharply in the mid-1990s. Since such a decrease was not observed, nor is
expected based on longer-term solar observations, the long-term decreases in global
ozone cannot result from changes in solar output alone. Most analyses presented inthis and previous international scientific assessments quantitatively account for the
influence of the 11-year solar cycle on long-term variations in ozone.
Total ozone and past volcanoes. Explosive volcanic eruptions inject sulfur gases
directly into the stratosphere, causing new sulfate particles to be formed. The
particles initially form in the stratosphere downwind of the volcano and then spreadthroughout the hemisphere or globally as air is transported by stratospheric winds.
One method of detecting the presence of volcanic particles in the stratosphere uses
observations of the transmission of solar radiation through the atmosphere (seeFigure Q14-1). When large amounts of new particles are formed in the stratosphere
over an extensive region, solar transmission is measurably reduced. The eruptions of
Mt. Agung (1963), El Chichón (1982), and Mt. Pinatubo (1991) are the most recent
examples of sulfur injections that temporarily reduced solar transmission.

Section III: Stratospheric ozone depletion

49

14
Laboratory measurements and stratospheric observations have shown that chemical
reactions on the surfaces of volcanically produced particles can increase ozone
destruction by increasing the amounts of the highly reactive chlorine gas, chlorinemonoxide (ClO). The ozone response depends on the total abundance of EESC after
the eruption (see Q16). At times of relatively low EESC, such as the early 1960s,
ozone is not very sensitive to stratospheric injection of volcanic sulfate particles.
At times of higher EESC amounts, such as from 1980 to the present, global ozone
is expected to decrease significantly following large explosive eruptions. The mostrecent large eruption was that of Mt. Pinatubo, which resulted in as much as a 10-foldincrease in the number of particles available for surface reactions. Both El Chichón
and Mt. Pinatubo increased global ozone depletion for a few years (see Figure Q14-1).
EESC was too low for ozone depletion to occur after the Mt. Agung eruption in 1963.
The effect on ozone diminishes during the years following an eruption as volcanic
particles are gradually removed from the stratosphere by gravitational settling andlarge-scale atmospheric transport. As particles are removed, solar transmission isrestored. Based on the short residence time of volcanic particles in the stratosphere,
the two large eruptions in the past three decades cannot account directly for thecontinuous long-term decreases in global total ozone observed over the same period.
Reactive chlorine from volcanoes. Explosive volcanic plumes generally contain large
quantities of reactive chlorine in the form of hydrogen chloride (HCl). HCl is a reactive
halogen gas that can be converted to ClO, which rapidly destroys ozone (see FigureQ8-3). The plumes also contain a considerable amount of water vapor, which forms
rainwater and ice in the rising fresh plume. Rainwater and ice efficiently scavenge andremove HCl while it is still in the lower atmosphere (troposphere). As a result, mostof the HCl in explosive volcanic plumes does not enter the stratosphere. After recent
explosive eruptions, observations of HCl in the stratosphere have confirmed that
increases are small compared with the total amount of chlorine in the stratospherefrom other sources.
Antarctic volcanoes. Volcanoes on the Antarctic continent are of special interest due
to their proximity to the Antarctic ozone hole. An explosive eruption could in principle
inject volcanic aerosol and small amounts of HCl directly into the stratosphere over
Antarctica, which could lead to ozone depletion. However, to be a possible cause of
the annually recurring ozone hole beginning in the early 1980s, explosive Antarctic
eruptions would need to have occurred at least every few years to maintain volcanic
emissions in the stratosphere. This is not the case. Only the Mt. Erebus volcano
is currently active in Antarctica. No explosive eruptions of Mt. Erebus or any other
Antarctic volcano have occurred since 1980. Therefore, explosive volcanic eruptionsin the last 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. Observations and atmospheric models indicatethat the record-low ozone levels observed in 1992–1993 resulted from the large
number of particles produced by the Mt. Pinatubo eruption, combined with therelatively large amounts of EESC present in the stratosphere in the early 1990s. If
the Mt. Pinatubo eruption had occurred before 1980, changes to global ozone would
have been much smaller than observed in 1992–1993 because EESC values weremuch lower. EESC values will remain substantial in the early decades of the 21st
century even as ODSs decline globally, with 1980 values not reached until about
2050 (see Figures Q16-1 and Q20-2). Large volcanic eruptions in the intervening
years will cause more ozone depletion. If an explosive eruption larger than Mt.
Pinatubo were to occur, peak ozone losses could be larger than previously observed
and substantial ozone losses could persist for longer time periods. As halogen gas
abundances gradually decline, the effect of volcanic eruptions on ozone will lessen.
50

Section III: Stratospheric ozone depletion

The Solar
Cycle,
Volcanic
GlobalOzone,The Solar
Cycle,
Volcanic Eruptions,
Eruptions, Global
Ozone,
andEquivalent
Effective
Stratospheric
Chlorineand Equivalent Effective Stratospheric Chlorine
10.7– cm solar radiation(relative amounts)

250
200

Incoming solar
radiation

150
100
50

Transmission
(%)

95
90
85
80

Change from 1964–1980
average (%)

75

EESC
(relative amounts)

Figure Q14-1. Solar changes andvolcanoes.
Global total ozone values have decreased
beginning in the early 1980s. Ozone
values shown are the 3-month averages
without corrections for solar or volcanic
effects and annual averages corrected forseasonal and solar effects (from Figure
Q13-1). These long-term ozone decreases
are primarily attributable to equivalenteffective stratospheric chlorine (EESC),
which has increased over the same time
period. Since the mid-1990s, changes
in both quantities have slowed andreversed sign. Incoming solar radiation
varies on a well-recognized 11-year cycle
related to sunspot activity. The amount ofincoming solar radiation at a wavelengthof 10.7 cm is often used as a surrogate
for incoming solar radiation at UV
wavelengths that produce stratosphericozone. The 10.7-cm radiation values
clearly show the recent periods of solar
maximum and minimum. A comparisonof the solar radiation and ozone changes
strongly indicates that the cyclic changes
in solar output alone cannot account for
the long-term decrease in total ozone.
Following large volcanic eruptions,transmission of solar radiation to Earth’s
surface is reduced by the large numberof new sulfur-containing particles formedin the stratosphere. The three large
volcanic eruptions that occurred between
1960 and 2010 temporarily decreased
solar transmission as measured in
Hawaii. Volcanic particles increase ozone
depletion only for a few years before they
are removed from the stratosphere by
natural processes. As a consequence,
the two most recent volcanic eruptions
cannot be the cause of the continuous
long-term decrease found in global
total ozone.

Solar transmission
at Hawaii (20°N)

Mt. Agung(8°S)

El Chichón(17°N)

2

Mt. Pinatubo(15°N)

Major volcanic
eruptionsGlobal total ozone changes

0
-2
-4
-6

1

3-month averages(no solar or volcanic corrections)

Annual averages(seasonal and solar corrections)

Equivalent effective
stratosphericchlorine (EESC)

Natural sources
0
1960

1970

1980

Year

1990

2000

2010

Section III: Stratospheric ozone depletion

51

15

Are there controls on the productionof ozone-depleting substances?
Yes, the production and consumption of ozone-depleting
substances are controlled under a 1987 international agreement
known as the “Montreal Protocol on Substances that Deplete the
Ozone Layer” and by its subsequent Amendments and adjustments.
The Protocol, now ratified by all 196 United Nations members,
establishes legally binding controls on national production andconsumption of ozone-depleting substances (ODSs). Production
and consumption of all principal ODSs by developed and developing
nations will be almost completely phased out before the middle ofthe 21st century.
Montreal Protocol. In 1985, a treaty called the Vienna Convention for the Protectionof the Ozone Layer was signed by 20 nations in Vienna. The signing nations agreed
to take appropriate measures to protect the ozone layer from human activities. The
Vienna Convention was a framework agreement that supported research, exchange
of information, and future protocols. In response to growing concern, the Montreal
Protocol on Substances that Deplete the Ozone Layer was signed in 1987 and,following sufficient country ratification, entered into force in 1989. The Protocol has
been successful in establishing legally binding controls for developed and developing
nations on the production and consumption of halogen source gases known to cause
ozone depletion. Halogen source gases containing chlorine and bromine controlled
under the Montreal Protocol are referred to as ozone-depleting substances (ODSs).
National consumption of an ODS is defined as the amount that production andimports of the substance exceed its export to other nations. The Protocol provisions
are structured for developed countries to act first and for developing countries tofollow with some financial assistance. In 2009, the Montreal Protocol became the firstinternational treaty to achieve universal ratification by all 196 United Nations members.

Effectofofthe
the Montreal
Montreal Protocol
ProtocolEffectLong-term changes in equivalent effective
stratospheric chlorine (EESC)

Montreal
1987

EESC (relative amounts)

4
No Protocol
3

London1990
2
Beijing
1999

Montreal
2007
1

Copenhagen
1992

Zero Emissions
in 2014Natural sources

0
1960

1980

2000

2020

2040

2060

Year

52

Section IV: Controlling ozone-depleting substances

2080

2100

Figure Q15-1. Effect of the Montreal Protocol.
The objective of the Montreal Protocol is the protection of
the ozone layer through control of the global production andconsumption of ODSs. Projections of the future abundances of
ODSs expressed as equivalent effective stratospheric chlorine(EESC) values (see Q16) are shown separately for the midlatitudestratosphere for (1) no Protocol provisions, (2) the provisions of
the original 1987 Montreal Protocol and some of its subsequent
Amendments and adjustments, and (3) zero emissions of ODSs
starting in 2014. The city names and years indicate where andwhen changes to the original 1987 Protocol provisions were agreed
upon (see Figure Q0-1). EESC is a relative measure of the potential
for stratospheric ozone depletion that combines the contributionsof chlorine and bromine from ODS surface observations (see Q16).
Without the Protocol, EESC values are projected to have increased
significantly in the 21st century. Only with the Copenhagen (1992)
and subsequent Amendments and adjustments did projected
EESC values show a long-term decrease.

Amendments and adjustments. As the scientific basis of ozone depletion became
more certain after 1987 and substitutes and alternatives became available toreplace ODSs, the Montreal Protocol was strengthened with Amendments andadjustments. Each Amendment is named after the city in which the meeting of theParties to the Montreal Protocol took place and by the year of the meeting. The
timeline in Figure Q0-1 shows some of the major decisions that have been adopted
in the last three decades. These decisions have put additional ODSs under control,
accelerated the timing of existing control measures, and prescribed phase-out dates
for the production and consumption of certain gases. The initial Protocol called
for a 50% reduction in chlorofluorocarbon (CFC) production and a freeze on halon
production. The 1990 London Amendment called for a phase-out of the production
and consumption of the most damaging ODSs in developed nations by 2000 and indeveloping nations by 2010. The 1992 Copenhagen Amendment accelerated thephase-out date to 1996 in developed nations. Further controls on ODSs were agreed
upon in later meetings in Vienna (1995), Montreal (1997, 2007), and Beijing (1999).
Influence of the Montreal Protocol. Montreal Protocol controls are based on several
factors that are considered separately for each ODS. The factors include (1) the
effectiveness in depleting ozone in comparison with other substances (see Ozone
Depletion Potential, ODP, in Q18), (2) the availability of suitable substitutes fordomestic and industrial use, 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 in equivalent effective
stratospheric chlorine (EESC). Calculations of EESC combine the amounts of chlorine
and bromine present in surface air to form a measure of the potential for ozone
destruction in a particular stratospheric region on an annual basis (see definition in
Q16). The long-term changes in EESC at midlatitudes are shown in Figure Q15-1 forseveral cases:• No Protocol. Without the Montreal Protocol the production, use, and emissionsof CFCs and other ozone-depleting substances is expected to have increased
after 1987 with an annual growth rate of about 3% (business-as-usual scenario).
As a result, EESC is projected to have increased nearly 10-fold by the mid-2050s
compared with the 1980 value. Computer models of the atmosphere indicate thatsuch high EESC values would have at least doubled global total ozone depletion
between 1990 and 2010 and increased it far beyond that by midcentury. As
a result, harmful UV-B radiation would have increased substantially at Earth’s
surface, causing damage to ecosystem health and a global rise in skin cancer andcataract cases (see Q17).
• Montreal Protocol provisions. International compliance with only the 1987
provisions of the Montreal Protocol and the later 1990 London Amendment
would have substantially slowed the projected growth of EESC. Not until the 1992
Copenhagen Amendments and adjustments did the Protocol projections show adecrease in future EESC values. The provisions became more stringent with theAmendments and adjustments adopted in Beijing in 1999 and Montreal in 1997
and 2007. Now, with full compliance to the Protocol, most ODSs will be almost
completely phased out, with some exemptions for critical uses (see Q16). Global
EESC is slowly decaying from its peak value in the late 1990s and is expected toreach 1980 values in the mid-21st century. The success of the Montreal Protocol todate 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 thannatural emissions. Between 1987 and 2014, ODS emissions from human activities
have decreased by almost 80%.

Section IV: Controlling ozone-depleting substances53

15• Zero emissions. EESC values in the coming decades will be influenced by (1) theslow natural removal of ODSs still present in the atmosphere, (2) emissions fromcontinued production and use of ODSs, and (3) emissions from currently existingbanks containing a variety of ODSs. ODS banks are associated with applications
that involve long-term containment of halogen gases. Examples are CFCs inrefrigeration equipment and insulating foams, and halons in fire-fighting equipment.
New emissions are projected based on continued production and consumption of
ODSs, particularly in developing nations, under existing Protocol provisions.
The zero-emissions case demonstrates the EESC values that would occur if it werepossible to set all ODS emissions to zero beginning in 2014. This would eliminate
the contributions from new production and bank emissions. Significant differencesfrom the Montreal 2007 projections are evident in the first decades following 2014
because the phase-out of all ODS production under the Protocol is not yet complete
and bank emissions are substantial. Zero emissions would bring forward the returnof EESC to 1980 levels by about 13 years.
HCFC substitute gases. The Montreal Protocol provides for the use ofhydrochlorofluorocarbons (HCFCs) as transitional, short-term substitute compounds
for ODSs with higher ODPs, such as CFC-12. HCFCs are used for refrigeration, inmaking 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. HCFCs are
88 to 98% less effective than CFC-12 in depleting stratospheric ozone because theirchemical removal occurs primarily in the troposphere (see ODPs in Table Q7-1). This
removal protects stratospheric ozone from most of the halogen content of HCFC
emissions. In contrast, CFCs and some other ODSs release all of their halogen content
in the stratosphere because they are chemically inert in the troposphere (see Q6).
Under the provisions of the Montreal Protocol, developed and developing countries
may continue to use HCFCs as ODS substitutes in the coming decades before they
are ultimately phased out. In the most recent adjustment to the Protocol (Montreal2007), the phase-out of HCFCs for all Parties was accelerated so that it will becomplete in 2030, a decade earlier than in previous provisions. In adopting thisdecision, the Parties reduced the contribution of HCFC emissions to both long-termozone depletion and future climate forcing (see Q18 and Q19).
HFC substitute gases. Hydrofluorocarbons (HFCs) are used as substitute compounds
for CFCs, HCFCs, and other ODSs. HFCs contain hydrogen, fluorine, and carbon. HFCs
do not contribute to ozone depletion because they contain no chlorine, bromine, or
iodine. As a consequence, HFCs are not ODSs and are not subject to the provisions of
the Montreal Protocol. HFCs and all ODSs are radiatively active gases that contributeto human-induced climate change based on their accumulation in the atmosphere(see Q18). HFCs are included in the basket of gases controlled by the Kyoto Protocolof the United Nations Framework Convention on Climate Change (UNFCCC). The
Kyoto Protocol is an international treaty designed to protect climate by controllingemissions of carbon dioxide (CO2), methane (CH4), nitrous oxide (N2O), HFCs,
perfluorocarbons (PFCs), and sulfur hexafluoride (SF6). HFC emissions are expected
to grow substantially in coming decades as the demand for their use as substitutegases and in new applications increases in developed and developing countries. The
climate change contribution from future HFC emissions would be minimized if HFCs
with very low Global Warming Potentials (GWPs) (less than 100) were chosen to meetthe growing demand (see Q18 and Q19).

54

Section IV: Controlling ozone-depleting substances

16

Has the Montreal Protocol beensuccessful in reducing ozone-depleting
substances in the atmosphere?
Yes, as a result of the Montreal Protocol, the overall abundance ofozone-depleting substances (ODSs) in the atmosphere has been
decreasing for more than a decade. If the nations of the worldcontinue to comply with the provisions of the Montreal Protocol, thedecrease will continue throughout the 21st century. Those gasesthat are still increasing in the atmosphere, such as halon-1301 and
HCFCs, will begin to decrease in the coming decades if compliancewith the Protocol continues. However, it is only after midcentury thatthe effective abundance of ODSs is expected to fall to values thatwere present before the Antarctic ozone hole was first observed in
the early 1980s.
The Montreal Protocol and its Amendments and adjustments have been very
successful in reducing the atmospheric abundance of ozone-depleting substances(ODSs). ODSs are halogen source gases released by human activities. Under theProtocol, the production and consumption of individual ODSs are now controlled in all
196 nations that are Parties to the Protocol (see Q15). The success of the Montreal
Protocol controls is documented by (1) observed changes and future projections of
the atmospheric abundances of the principal ODSs and (2) long-term changes inequivalent effective stratospheric chlorine (EESC).
Individual ODS reductions. The reduction in the atmospheric abundance of an
ODS in response to controls on production and consumption depends principallyon two factors: (1) how rapidly an ODS is used and released to the atmosphere
after being produced and (2) the lifetime for the removal of the ODS from theatmosphere (see Table Q7-1). For example, the abundances of ODSs with short
lifetimes respond quickly to emission reductions. Long-term changes in ODS
abundances are constructed from (1) estimates of historical emissions and bankquantities using industry reports, (2) abundances measured in air trapped for years
in accumulated snow (firn) in polar regions, (3) observed atmospheric abundances
using ground-based measurements, and (4) projections of future abundances basedon compliance with Montreal Protocol provisions and patterns of use in developed
and developing countries. The results for individual ODSs and the natural chlorine
and bromine source gases, methyl chloride (CH3Cl) and methyl bromide (CH3Br), areshown in Figure Q16-1 and described as follows:• CFCs. Chlorofluorocarbons (CFCs) include some of the most destructive chlorinecontaining ODSs. CFC-11 and CFC-12, each with an Ozone Depletion Potential(ODP) near 1, are the most abundant ODSs in the atmosphere owing to large
historical emissions and long atmospheric lifetimes (50–100 years). Production
and consumption of CFCs in developed countries ended in 1996 and that indeveloping countries ended by January 2010. As a consequence, CFC-11 andCFC-113 abundances have peaked in the atmosphere and have been declining for
more than a decade. In contrast, CFC-12 abundances have only recently shown adecrease, owing to its longer lifetime (100 years) and continuing emissions from
CFC-12 banks, namely, refrigeration and air conditioning equipment and thermal
insulating foams. With no further global production of the principal CFCs allowed
except for limited exempted uses and with some continuing emissions from banks,
CFC abundances are projected to decline steadily throughout this century.

Section IV: Controlling ozone-depleting substances

55

16
Past
Projected
AtmosphericAbundances
Abundances ofPast
andandProjected
Atmosphericof Halogen
HalogenSource
SourceGases
Gases
600

Equivalent effective stratosphericchlorine (EESC)
Midlatitudes

1

400

1980
level

Atmospheric abundance (parts per trillion)

CH3Br (natural + human sources)

0

CH3Cl (natural sources)

Carbon tetrachloride(CH3CCl3)
140

Estimates
Observations
Future projections

120
100
80

Methyl chloroform (CCl4)

60
40

HCFC-141b

20
0
1950

HCFC-142b
2000

2050

2100

Year
Figure Q16-1. Halogen source gas changes.
The atmospheric abundances of individual ozone-depleting
substances (ODSs) at Earth’s surface were obtained using a
combination of direct atmospheric measurements, estimates ofhistorical abundance, and future projections of abundance. The
gases shown are all ODSs except for methyl chloride. The pastincreases of CFCs, along with those of carbon tetrachloride, methyl
chloroform, and halon-1211, have slowed and reversed in the lasttwo decades. HCFCs, which are used as transitional substances
to replace CFCs, will likely continue to increase in the coming
decades before a complete phase-out. Halon-1301 abundances
will also likely continue to grow in the next decade while current
halon banks are depleted. Future decreases in methyl bromide
are expected to be small. Abundances of methyl chloride, which is
not controlled under the Montreal Protocol and has large natural

56 Section IV: Controlling ozone-depleting substances

CFC-12500

CFCs
Halons
HCFCs
CCl4
CH3CCl3

Atmospheric abundance (parts per trillion)

EESC (relative amount)2300

CFC-11
HCFC-22200
CFC-113100
0Halon-1211

4

Halon-1301

3
2
1

Halon-2402

0

12

500

Methyl chloride (CH3Cl)

10

400

Methyl bromide (CH3Br)

8

2050

6
2100

300
1950

2000

Year
sources, are projected to be constant in the future. The very shortlived gases are not included here.
The rise in equivalent effective stratospheric chlorine (EESC)
values in the 20th century has slowed and reversed in the last
decade (top left panel). EESC is a measure of the potential
for ozone depletion in the stratosphere derived from surface
measurements of chlorine and bromine gases (see Q16). EESC
values are reduced as ODS emissions decrease. EESC, as shown
here for the stratosphere at midlatitudes, returns to 1980 values
around 2050. In polar regions, the return to 1980 values occurs
more than two decades later. International compliance with theMontreal Protocol provisions will ensure that EESC values willcontinue to decrease as projected (see Q15). (The unit “parts pertrillion” is defined in the caption of Figure Q7-1.)

• Halons. Halons are the most destructive bromine-containing ODSs. The mostabundant in the atmosphere, halon-1211 and halon-1301, have abundances
about 100 times less than CFC-11 and CFC-12 and account for a significant
fraction of bromine from all ODSs (see Figure Q7-1). Production and consumptionof halons in developed countries ended in 1994 and that in developing countries
ended by January 2010. Atmospheric abundances of halon-1211 show significant
decreases since peak concentrations were measured in the early 2000s. Halon2402 abundances have been decreasing slowly for a few years while those of
halon-1301 continue to increase. The increase is likely due to substantial banks
in fire-extinguishing and other equipment that gradually release halon-1301 to theatmosphere. The abundance of halon-1301 is expected to remain high well into the
21st century because of its long lifetime (65 years) and continued release.
• Methyl chloroform. The largest reduction to date in the abundance of an ODS(96% from its peak value) has been observed for methyl chloroform. Production andconsumption of methyl chloroform in developed countries ended in January 1996
and that in developing countries is scheduled to end by January 2015. Atmosphericabundances responded rapidly to the reduced emissions starting in the mid-1990s
because methyl chloroform has a short atmospheric lifetime (about 5 years). Methyl
chloroform is used mainly as a solvent and has no significant long-term storage
following production. It will approach complete removal from the atmosphere after
the phase-out in developing countries is complete.
• HCFC substitute gases. The Montreal Protocol allows for the use ofhydrochlorofluorocarbons (HCFCs) as short-term, transitional substitutes for CFCs
and in other specific applications. As a result, the abundances of HCFC-22, HCFC141b, and HCFC-142b continue to grow in the atmosphere in response to greaterapplication demand. HCFCs pose a lesser threat to the ozone layer than CFCs
because they have low ODP values (less than 0.12). The most recent adjustment
to the Montreal Protocol (Montreal 2007) accelerated the phase-out of HCFCs
by a decade for both developed countries (2020) and developing countries(2030) (see Q15). Even with the accelerated phase-out, future projections show
HCFC abundances that continue to increase, reach peak values around 2030,
and steadily decrease thereafter. The response of atmospheric abundances todecreasing emissions will be relatively rapid because of the short atmospheric
lifetimes of HCFCs (less than 20 years).
• Carbon tetrachloride. Carbon tetrachloride has been phased out in both developed
countries (January 1996) and developing countries (January 2010). As a result,atmospheric abundances of carbon tetrachloride have been decreasing for two
decades. The decline is somewhat less rapid than expected, suggesting that
actual emissions are larger than reported or the atmospheric lifetime is greaterthan current estimates. Production of carbon tetrachloride for use as raw material(feedstock) to make other chemicals is exempted under the Protocol because
emission to the atmosphere from this use is believed to be minimal.
• Methyl chloride and methyl bromide. Both methyl chloride and methyl bromide
are distinct among halogen source gases because substantial fractions of theiremissions are associated with natural processes (see Q7). Methyl chloride is not
an ODS under the Montreal Protocol because it is not manufactured or used inappreciable quantities. Its abundance in the atmosphere has remained fairly
constant throughout the last 60 years and will remain constant if the balance of
total sources and loss processes remains unchanged. Methyl bromide is controlled
under the Protocol because it is manufactured for use as a fumigant. Developed
country production and consumption of methyl bromide ended in January 2005
and that in developing countries is scheduled to end by January 2015. The Protocol

Section IV: Controlling ozone-depleting substances57

16
currently provides exemptions for some methyl bromide production and use as
an agricultural and pre-shipment fumigant. Atmospheric abundances of methyl
bromide responded rapidly to the reduced emissions starting in 1999 because its
atmospheric removal lifetime is less than 1 year. Future projections show only smallchanges in methyl bromide abundances based on the assumptions of unchanged
contributions from natural sources and small use in developing countries. An
important uncertainty in these projections is the future amounts that will be
produced and emitted under Montreal Protocol use exemptions.
Equivalent effective stratospheric chlorine (EESC). Important measures of thesuccess of the Montreal Protocol are the past and projected changes in the values of
equivalent effective stratospheric chlorine (EESC), which was introduced in Figures
Q14-1 and Q15-1. EESC is designed as one measure of the potential for ozone
depletion in the stratosphere that can be calculated from atmospheric surface
abundances of ODSs and natural chlorine and bromine gases. All gases used in
the calculation are shown in Figure Q7-1. For both past and future EESC values,
the required atmospheric abundances are derived from measurements, historicalestimates, or future projections based on compliance with Montreal Protocol provisions.
EESC is expressed as a hypothetical amount of chlorine available in the stratosphereto deplete ozone. The term equivalent indicates that bromine gases, scaled by theirgreater per-atom effectiveness in depleting ozone, are included in EESC. The term
effective indicates that only the estimated fraction of ODSs that are currently in theform of reactive halogen gases in the stratosphere is included in an EESC value (seeQ8). Although chlorine is much more abundant in the stratosphere than bromine (about
150-fold) (see Figure Q7-1), bromine atoms are about 60 times more effective thanchlorine atoms in chemically destroying ozone in the lower stratosphere. EESC values
generally depend on the year and latitude region in the stratosphere being considered.
Another quantitative measure of the potential for ozone depletion in the stratospherethat is highly related to EESC is effective stratospheric chlorine (ESC). ESC, asdefined in Q20, is calculated with chemistry-climate models rather than inferred fromtropospheric ODS observations.
Long-term changes in EESC. In the latter half of the 20th century up until the 1990s,
EESC values steadily increased (see Figure Q16-1), causing global ozone depletion.
As a result of the Montreal Protocol regulations, the long-term increase in EESC
slowed, values reached a peak, and EESC began to decrease in the late 1990s.
By 2012, EESC at midlatitudes had declined by about 15% from its peak value in1997. The initial decrease came primarily from the substantial, rapid reductions inemissions of methyl chloroform, which has a lifetime of only 5 years. The decrease iscontinuing with declining emissions of CFCs and other long-lived ODSs. Decreasing
EESC means that the potential for stratospheric ozone depletion is now lessening
each year as a result of the Montreal Protocol. Decreases in EESC are projected
to continue throughout the 21st century if all nations continue to comply with theprovisions of the Protocol. The decrease will continue because as emissions are
reduced, natural processes continue to gradually remove halogen-containing gases
from the global atmosphere. Reduction of EESC to 1980 values or lower will require
several more decades because the most abundant ODS gases now in the atmospherehave lifetimes ranging from 10 to 100 years.

58

Section IV: Controlling ozone-depleting substances

17

Does depletion of the ozone layer increase
ground-level ultraviolet radiation?
Yes, ultraviolet radiation at Earth’s surface increases as theamount of overhead total ozone decreases, because ozone absorbs
ultraviolet radiation from the Sun. Measurements by ground-based
instruments and estimates made using satellite data provide
evidence that surface ultraviolet radiation has increased in large
geographic regions in response to ozone depletion.
The depletion of stratospheric ozone leads to an increase in solar ultraviolet radiation
at Earth’s surface. The increase occurs primarily in the ultraviolet-B (UV-B) componentof the Sun’s radiation. UV-B is defined 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 estimatedfrom total ozone changes.
UV-B radiation can harm humans, other life forms, and materials (see Q3). Most of theeffects of sunlight on the human body are caused by UV-B exposure. A principal effectis skin erythema, which leads to sunburning. Excess exposure may lead to skin cancers.
Erythemal radiation is regularly reported to the public in many countries in the form of
the “UV Index.” The long-term changes in surface UV-B radiation are important to study
because of its potential harmful effects and relationship to ozone depletion.
Surface UV-B radiation. The amount of UV-B radiation reaching Earth’s surface ata particular location depends in large part on total ozone at that location. Ozone
molecules in the stratosphere and in the troposphere absorb UV-B radiation, thereby
significantly reducing the amount that reaches Earth’s surface (see Q3). If conditions
occur that reduce the abundance of ozone molecules somewhere in the troposphereor stratosphere, total ozone is reduced and the amount of UV-B radiation reaching
Earth’s surface below is increased proportionately. This relationship between totalozone and surface UV-B radiation has been confirmed at a variety of locations withdirect measurements of both quantities.
Additional causes of UV changes. The actual amount of UV-B radiation reaching
Earth’s surface at a specific location and time depends on a number of factors in
addition to total ozone. The primary additional factor is the position of the Sun in
the sky, which changes with daily and seasonal cycles. Other factors include local
cloudiness, the altitude of the location, the amount of ice or snow cover, and theamounts of atmospheric particles (aerosols) in the atmosphere above the location.
Changes in clouds and aerosols are partially related to air pollution and greenhousegas emissions from human activities. Measurements indicate that both increases anddecreases in UV radiation at certain locations have resulted from changes in one ormore of these factors. Estimating the impact of changes in these factors is complex.
For example, an increase in cloud cover usually results in a reduction of UV radiation
below the clouds and could at the same time increase UV radiation at a location in
the mountains above the clouds.
Long-term surface UV changes. Long-term changes in UV-B radiation have been
estimated from measurements made with special UV monitoring instruments at several
surface locations since about 1990. For example, as a consequence of Antarctic ozone
depletion, the average UV-B measured at the South Pole during spring between 1991
and 2006 was 55–85% larger than estimated for the years 1963–1980. In addition,satellite observations of ozone changes have been used to estimate changes insurface UV-B radiation that have occurred over the past three decades. With satellite
observations, the UV-B radiation changes can be separately attributed to changes
in ozone and clouds. The results show that erythemal radiation has increased by anaverage of 3% between 1979 and 2008 over a wide range of latitudes outside theSection V: Implications of ozone depletion and the Montreal Protocol

59

17
ChangesininSurface
Surface Erythemal
Erythemal UV
ChangesUV Radiation
Radiation
Changes between 1979 and 2008

12

Averages

Average change (%)

10

Uncertainty ranges

86. . . from totalozone changes

4
2
0
-2

. . . from cloud andaerosol changes

Average change (%)-4

Average
Uncertainty range

86. . . from combinedozone, cloud, andaerosol changes

4
2
0
-2

-60
South

-40

-20

0

20

Latitude (degrees)

40

60
North

Figure Q17-1. Long-term changes in surface erythemalradiation.
Solar ultraviolet (UV) radiation at Earth’s surface has increased
over much of the globe since 1979. Erythemal radiation,
which can lead to sunburning, is a component of surface UVradiation that is harmful to humans and other life forms. Surface
erythemal radiation responds to changes in total ozone as well
as clouds and aerosols. Satellite observations have been used
to estimate the long-term changes of erythemal radiation due tochanges in ozone, clouds, and aerosols. These estimates showthat erythemal radiation at Earth’s surface has increased overmuch of the globe over the period 1979 to 2008, particularly at
midlatitudes in both hemispheres (bottom panel). The increases
in the Southern Hemisphere would have been larger without theoffsetting changes due to clouds and aerosols (upper panel).
The smallest estimated changes in erythemal radiation are in
the tropics because observed total ozone changes over thisperiod are smallest there (see Figure Q13-1). The shaded areas
surrounding the figure lines represent the uncertainties in theestimated changes.

tropics (see lower panel of Figure Q17-1). The largest percentage increases haveoccurred at high polar latitudes in both hemispheres, where the largest annual
decreases in total ozone are observed (see Figure Q13-1). Over this time period the UV
increases due to ozone depletion are partially offset by changes in cloudiness, primarily
in the high latitudes of the Southern Hemisphere (see top panel in Figure Q17-1).
Without changes in cloudiness, the increases in erythemal radiation at these latitudes
would have reached a maximum close to 9%. The smallest changes in erythemal UV
have been in the tropics, where long-term changes in total ozone are smallest (seeQ13). These findings have not changed substantially between 2008 and 2014.
UV Index changes. The UV Index is a measure of the erythemal radiation that occurs
at a particular surface location and time. The index is used internationally to increase
public awareness about the detrimental effects of UV on human health and to guidethe need for personal protective measures. The maximum daily UV Index varies withlocation and season, as shown for three sites in Figure Q17-2. The UV Index increases
when moving from high to low latitudes and is highest in summer when the midday Sun
is closest to overhead. For example, UV Index values in San Diego, California, at 32°N,
are generally higher than those in Barrow, Alaska, at 71°N. At all latitudes, UV Index
values increase in mountainous regions and over snow- or ice-covered regions. The UV
Index is zero during periods of continuous darkness in winter at high-latitude locations.
The UV Index over Antarctica has increased dramatically due to ozone depletion, as
illustrated in Figure Q17-2. Normal index values for Palmer, Antarctica, at 64°S inspring were estimated from satellite measurements made during the period 1978–

60

Section V: Implications of ozone depletion and the Montreal Protocol

1980, before the appearance of the ozone hole over Antarctica. In the period from1990 to 2006, the severe and persistent ozone depletion that occurred in late winterand early spring over Antarctica increased the average UV Index well above normal
values for several months. Now, the spring UV Index measured in Palmer, Antarctica,
routinely equals or exceeds spring and summer values measured in San Diego,
California, which is located at a much lower latitude (32°N).
UV changes and human health. Over the past several decades, depletion of thestratospheric ozone layer together with societal changes in lifestyle have increased
UV-B radiation exposure for many people. Increased exposure has adverse healtheffects, primarily associated with eye and skin disorders. UV radiation is a recognizedrisk factor for some types of eye cataracts. For the skin, the most common threat isskin cancer. Over the past decades, the incidence of several types of skin tumors
has risen significantly among people of all skin types. Skin cancer in humans occurs
long after exposure to sunburning UV. With current Montreal Protocol provisions,projections of additional skin cancer cases associated with ozone depletion arelargest in the early to middle decades of the 21st century and represent a significant
global health issue. Since it is projected that the recovery of the ozone layer to1980 levels will not occur until the mid-21st century, extensive and unprotected
UV-B exposure over the next decades will continue to have adverse health effects.
An important human health benefit of UV-B radiation exposure is the production ofvitamin D, which plays a significant role in bone metabolism and the immune system.
Human exposure to solar UV-B radiation requires a careful balance to maintainadequate vitamin D levels while minimizing long-term risks of skin and eye disorders
now and well into the future.

Seasonal
Changesininthe
the UVUV Index
Index
Seasonal
Changes
Northern Hemisphere months

J
16

F

M

WinterA

M

J

SpringJASOSummer

N

D

Fall

14

Daily UV Index

12
10
8
6
4
2
0

J

A

S

O

N

D

J

F

M

A

Southern Hemisphere months
Long-term averages of the UV Index
Palmer, Antarctica (64°S)
1990 – 2006
1978 – 1980

(pre-ozone hole)

San Diego, California (32°N)
1992 – 2006
Barrow, Alaska (71°N)
1991– 2007

M

J

Figure Q17-2. Long-term changes in the UV Index.
The UV Index is a measure of the erythemal radiation thatreaches the surface at a particular location and time. Sunburningin humans is caused by erythemal radiation. The UV Index varies
with total ozone above a location, as well as the Sun’s midday
position in the local sky, which depends on latitude and season.
The highest values of the UV Index occur in the tropics, where the
midday Sun is highest throughout the year and where total ozone
values are low. The lowest average UV Index values occur at highlatitudes. The figure shows daily maximum UV Index values atseveral locations. Values are higher throughout the year in San
Diego, a low-latitude location, than in Barrow, a high-latitude
location. UV Index values are zero at high latitudes in winter whendarkness is continuous. Measurements at Palmer, Antarctica,
demonstrate the effect of Antarctic ozone depletion. Normal
values estimated for Palmer are shown for 1978–1980, a period
before the appearance of the “ozone hole” (see red dashed line).
In the period 1990–2006, Antarctic ozone depletion increased
the UV Index at Palmer throughout spring and into summer (seeyellow shaded region). Springtime values at Palmer now routinely
equal or exceed those measured in spring and summer in San
Diego, which is located at much lower latitude.

Section V: Implications of ozone depletion and the Montreal Protocol

61

18

Is depletion of the ozone layer the principal
cause of global climate change?
No, ozone depletion itself is not the principal cause of global
climate change. Changes in ozone and climate are directly linked
because ozone absorbs solar radiation and is also a greenhousegas. Stratospheric ozone depletion leads to surface cooling,
while the observed increases in tropospheric ozone and othergreenhouse gases lead to surface warming. The cooling fromozone depletion is small compared to the warming from thegreenhouse gases responsible for observed global climate change.
Ozone depletion and global climate change are indirectly linked
because both ozone-depleting substances and their substitutes
are greenhouse gases. The Antarctic ozone hole has contributed tochanges in Southern Hemisphere surface climate through indirect
effects on the atmospheric circulation.
While stratospheric ozone depletion is not the principal cause of climate change,
aspects of ozone depletion and climate change are closely linked. Both processes
involve gases released to the atmosphere by human activities. The links are best
understood by examining the contribution to climate change of the gases involved:
ozone; ozone-depleting substances (ODSs) (or halogen source gases) and theirsubstitutes; and other leading greenhouse gases.
Greenhouse gases and the radiative forcing of climate. The warming of Earth by
the Sun is enhanced by the presence of natural greenhouse gases, of which water
vapor is an important example. The natural abundances of these gases in Earth’s
atmosphere change the balance between incoming solar radiation and outgoing
infrared radiation, warming the atmosphere and surface. Without this natural
greenhouse effect, Earth’s surface would be much colder. Human activities in the
Industrial Era have led to long-term increases in the atmospheric abundancesof a number of long-lived and short-lived greenhouse gases. This group includes
carbon dioxide (CO2), methane (CH4), nitrous oxide (N2O), tropospheric ozone, andhalocarbons. ODSs and their substitutes make up a large fraction of the halocarbons
in today’s atmosphere. Increases in the abundances of these gases from humanactivities cause more outgoing radiation to be absorbed, which further warms theatmosphere and surface. This change in the Earth’s radiative balance caused byhuman activities is called a radiative forcing of climate or, more simply, a climate
forcing. The magnitude of this energy imbalance is usually evaluated at the top of
the troposphere (tropopause) and is expressed using units of watts per square meter(W/m2). The potential for climate change increases as this radiative forcing increases.
A summary of radiative forcings in 2011 resulting from the increases in the principal
long-lived and short-lived greenhouse gases during the Industrial Era is shownin Figure Q18-1. All forcings shown relate to human activities. Positive forcings
generally lead to warming and negative forcings lead to cooling of Earth’s surface.
Climate forcings also lead to other changes, such as in precipitation patterns andextreme weather events. International climate assessments conclude that much of
the observed surface warming and changes in other climate parameters over thelast decades is due to increases in the abundances of carbon dioxide and othergreenhouse gases caused by human activities.
Carbon dioxide, methane, and nitrous oxide. All three of these greenhousegases have both human and natural sources. The accumulation of carbon dioxideduring the Industrial Era represents the largest climate forcing related to human62

Section V: Implications of ozone depletion and the Montreal Protocol

Radiative Forcing
Forcing of of
ClimateRadiative
ClimateFrom changes in greenhouse gases caused byhuman activities between 1750 and 2011
Figure Q18-1. Radiative forcingof greenhouse gases.
Human activities since the start of theIndustrial Era (around 1750) have caused
increases in the abundances of several
short-lived and long-lived gases, known
as greenhouse gases, that all contributeto the radiative forcing of climate, also
known as climate forcing. Radiative
forcing is expressed in units of watts
per square meter (W/m2). As shown in
the figure, the largest forcings are thoseof carbon dioxide (CO2), followed bymethane (CH4), ozone increases due to
air pollution gases, halocarbon gases,
and nitrous oxide (N2O). These positive
forcings lead to a warming of Earth’s
surface. Halocarbons include all ODSs,
their substitutes, and a few other gases(see Figure Q18-2). In contrast, ozone
decreases due to ODSs represents a
negative forcing, which leads to coolingof Earth’s surface and tends to partially
offset the positive forcing from thehalocarbons. The black whiskers on eachbar show uncertainties in the values.

Carbon dioxide (CO2)Methane (CH4)

Halocarbons

Nitrous oxide (N2O)

Ozone decrease
due to ODSs

-1.0

-0.5

Ozone increase
due to other gases

0

0.5

1.0

1.5

2.0

2.5

2

Radiative forcing (W/m )Surface coolingSurface warming

activities. Carbon dioxide concentrations continue to increase in the atmosphereprimarily as the result of burning fossil fuels (coal, oil, and natural gas) for energy
and transportation, as well as from cement manufacturing. The global mean
atmospheric abundance of carbon dioxide is now more than 40% above what it wasin preindustrial times (before 1750), approaching values of 400 parts per million
by volume (ppmv). Carbon dioxide is considered a long-lived gas, since a significant
fraction remains in the atmosphere 100–1000 years after emission.
Methane is a short-lived climate gas (atmospheric lifetime of about 10 years). Humansources include livestock, fossil fuel extraction and use, rice agriculture, and landfills.
Natural sources include wetlands, termites, and oceans. Methane has increased byabout 150% since preindustrial times.
Nitrous oxide is a long-lived climate gas (atmospheric lifetime of about 110 years).
The largest human source is agricultural activities, especially related to fertilization.
Microbial processes in soils that are part of natural biogeochemical cycles represent
the largest natural source. In the stratosphere, nitrous oxide is the principal source ofreactive nitrogen species, which participate in ozone destruction cycles (see Q2 andQ7). Nitrous oxide has increased by about 20% since preindustrial times.
Halocarbons. Halocarbons in the atmosphere contribute to both ozone depletionand climate change. As used here, halocarbons represent those gases containing
chlorine, bromine, or fluorine atoms that are now controlled substances under theMontreal Protocol or the Kyoto Protocol. ODSs are the halocarbons controlled under
Section V: Implications of ozone depletion and the Montreal Protocol

63

18
Figure Q18-2. Halocarbons and radiative
forcing of climate.
Halocarbon gases in the atmosphere
represent an important contribution to theradiative forcing of climate between 2011
and the start of the Industrial Era (see Figure
Q18-1). Halocarbons are all gases containing
chlorine, bromine, or fluorine atoms thatare now controlled as ozone-depleting
substances (ODSs) by the Montreal Protocol
or as climate change gases by the Kyoto
Protocol (see color shading). Shown in the
figure are the separate contributions ofeach gas or group of gases, as estimatedusing atmospheric abundance histories
and Global Warming Potentials (GWPs)(see Figure Q18-3). The gases listed in theright hand labels begin with the largest
contribution in each group, except for
CFC-13 and CFC-115, which are shownas one total value. The individual forcing
terms add together to form the bottom bar
representing the total halocarbon forcing.
The forcings of CFC-11 and CFC-12, thelargest halocarbon contributions, are alreadydecreasing and will continue to decreaseas CFCs are gradually removed from theatmosphere (see Figure Q16-1). In contrast,
the contributions of the intermediate-term
ODS substitute gases, HCFCs, are projected
to grow for another two decades before
decreasing. The future contributions of thelong-term ODS substitute gases, HFCs, arealso expected to increase. In this case, thetotal contribution will depend strongly on
which HFCs are used because the GWPs
of individual HFCs vary widely (see Figures
Q18-3 and Q19-2).

Radiative
Forcing
Halocarbons
Radiative
ForcingofofClimate
Climate byby
Halocarbons
From increases in all controlled gases containing chlorine, bromine,and fluorine from human activities between 1750 and 2011
CFC-12,
CFC-11,
CFC-113

Major CFCs

Minor CFCs

(CFC-115 and CFC-13)

Other ODSs:
Carbon tetrachloride
Methyl chloroform

CCl4,
CH3CCl3
HCFC-22,
HCFC-142b,
HCFC-141b

Hydrochlorofluorocarbons(HCFCs)Hydrofluorocarbons (HFCs)
Perfluorocarbons (PFCs)
Sulfur hexafluoride (SF6)

HFCs,
PFCs,
SF6

Montreal Protocol Gases
Kyoto Protocol Gases

Total halocarbon
climate forcing
0

0.1

0.2

0.3

Radiative forcing (W/m 2 )0.4

Increasing surface warming

the Montreal Protocol. HFC substitute gases, perfluorocarbons (PFCs), and sulfur
hexafluoride (SF6) are controlled under the Kyoto Protocol. In 2011, the halocarbon
contribution to climate forcing was 0.36 W/m2, which is the fourth largest forcing
following carbon dioxide, methane, and tropospheric ozone (see Figure Q18-1).
The contributions of individual halocarbon gases are highlighted in Figure Q18-2.
Within the halocarbons, CFCs contribute the largest percentage (73%) to the 2011
climate forcing. HCFCs, the intermediate-term ODS substitutes, make the next largest
contribution (14%). The atmospheric abundance of HFCs, the longer-term ODS
substitutes, contributes 5% to the 2011 halocarbon climate forcing.
The large contribution of the CFCs has been gradually decreasing following thedecline in their atmospheric abundance and is expected to further decrease (seeFigure Q16-1). Based on their long lifetimes, CFCs will still make a significant
contribution, and most likely the largest ODS contribution, to halocarbon climate
forcing at the end of the 21st century. Halocarbons controlled under the Kyoto Protocol

64

Section V: Implications of ozone depletion and the Montreal Protocol

Evaluationof Selected
Ozone-DepletingSubstances
Substances and
Gases
Evaluationof Selected
Ozone-Depleting
andSubstituteSubstitute
Gases
Relative importance of equal mass emissions for ozone depletion and climate change
Ozone-depleting substances(1)

Chlorine-containing
gases

CFC-11
CFC-113
CFC-12
Carbon tetrachloride
HCFC-22(0.03)

Halon-1301

Bromine-containing
gases

Substitute gases

(0)
(0)

0

Halon-1211
Methyl bromide(2)

HFC-23

Hydrofluorocarbons(HFCs)

HFC-143a(0)

HFC-125(0)

HFC-134a(0)

HFC-152a

(150)

(0)

HFC-1234yf

(<1)

5
10
15
Ozone Depletion Potential (ODP)
Increasing ozone depletion

20

0

4000
8000
12000
16000
Global Warming Potential (GWP, 100-yr)
Increasing surface warming

Figure Q18-3. ODPs and GWPs.
ODSs and their substitutes can be compared via their Ozone Depletion Potentials (ODPs)
and Global Warming Potentials (GWPs) (see Table Q7-1). Larger ODPs or GWPs indicate
greater potential for ozone depletion or climate warming, respectively, when a gas isemitted to the atmosphere. The ODP and GWP values are derived assuming an equalmass of each gas is emitted. The GWPs shown here are evaluated for a 100-year timeinterval after emission. The ODP of CFC-11 and the GWP of CO2 are assigned reference
values of 1. The CFCs, halons, and HCFCs are ozone-depleting substances (see Q7) while
HFCs, used as ODS substitutes, do not destroy ozone (ODPs equal 0). The ODPs of thehalons far exceed those of the CFCs. All ODSs and their substitutes shown here have a
non-zero GWP, with values spanning the wide range of 2 to 13,000.

(HFCs, PFCs, and SF6) represent about 8% of halocarbon climate forcing in 2011.
With the projected growth of HFC production and consumption in developing nations,this percentage contribution is expected to continue to increase in the coming
decades (see Figure Q19-2).
Stratospheric and tropospheric ozone. Ozone in both the stratosphere and thetroposphere absorbs infrared radiation emitted from Earth’s surface, trapping heatin the atmosphere. Stratospheric ozone also significantly absorbs solar radiation.
As a result, increases or decreases in stratospheric or tropospheric ozone induce a
climate forcing and, therefore, represent direct links between ozone and climate. Air
pollution from human activities has led to increases in global tropospheric ozone,
causing a positive radiative forcing (warming) since preindustrial times (see Figure
Q18-1). The large uncertainty in tropospheric ozone forcing reflects the difficulty in
Section V: Implications of ozone depletion and the Montreal Protocol

65

18
quantifying tropospheric ozone trends and in modeling the complex production andloss processes that control its abundance. On the other hand, rising ODS abundancesin the atmosphere since the middle of the 20th century have led to decreases instratospheric ozone and in the amounts of stratospheric ozone transported into
the troposphere, causing a negative radiative forcing (cooling) since preindustrial
times. The negative radiative forcing resulting from stratospheric ozone depletion willdiminish in the coming decades as ODSs are gradually removed from the atmosphere.
Stratospheric ozone depletion cannot be a principal cause of present-day global
climate change for two reasons: first, the climate forcing from ODS-induced ozone
depletion acts to cool Earth’s surface. Second, the total forcing from other long-lived
and short-lived gases in Figure Q18-1 is positive (leading to warming). The total
forcing from these other greenhouse gases (carbon dioxide, methane, halocarbons,and nitrous oxide) is the principal cause of observed climate change.
Ozone Depletion Potentials and Global Warming Potentials. An important way ofcomparing the influence of individual emissions of halocarbons on ozone depletionand climate change is to use Ozone Depletion Potentials (ODPs) and Global Warming
Potentials (GWPs). The ODP and GWP are the effectiveness of an emission of a gas
in causing ozone depletion and climate forcing, respectively, relative to a referencegas (see Table Q7-1). The principal halocarbon gases are contrasted with each otherin Figure Q18-3. The ODP of CFC-11 and the GWP of carbon dioxide are assigned
reference values of 1. The CFCs and carbon tetrachloride all have ODPs near 1,
indicating comparable effectiveness in causing ozone depletion per mass emitted.
The principal halons have ODPs greater than 7, making them the most effective
ozone-depleting substances per mass emitted. HFCs have ODPs of zero since theycause no ozone depletion (see Q7).
All halocarbons have non-zero GWPs and, therefore, contribute to climate forcing.
The GWP does not correspond strongly with the ODP of a gas because thesequantities depend on different chemical and physical properties. For example, while
HFC-134a does not destroy ozone (ODP equal 0), each gram emitted is 1,360 times
more effective than a gram of carbon dioxide in causing climate forcing. The future
selection of specific HFCs as ODS substitutes or for use in new global applications will
have important consequences for climate forcing. When these HFCs are eventually
released to the atmosphere, their contribution to climate forcing will depend on their
GWPs, which could vary over a wide range (2 to 13,000).
Montreal Protocol regulations have led to reductions in CFC emissions and increases
in HCFC emissions (see Q16). As a result of these actions, the total radiative forcing
from ODSs stopped increasing and is now slowly decreasing (see Q19). Overall
halocarbon radiative forcing, however, is slowly increasing because of growing
contributions from HFCs, PFCs, and SF6. It is important to note that, despite having a
GWP that is small in comparison to many other halocarbons and other greenhousegases, carbon dioxide is the most important greenhouse gas related to humanactivities because its emissions are large and its atmospheric abundance is far
greater than the abundances of other emitted gases.
The Antarctic ozone hole and Southern Hemisphere climate. While stratosphericozone depletion is not the principal cause of global climate change, the reoccurring
Antarctic ozone hole has indirectly contributed to observed changes in Southern
Hemisphere summertime surface climate through circulation changes. These newresearch findings are explained in more detail in the box below.

66

Section V: Implications of ozone depletion and the Montreal Protocol

The Antarctic Ozone Hole and Southern Hemisphere
Surface Climate
Links between stratospheric ozone depletion and changes in surface climate werefirst found in research studies in the early 2000s, based on both observations andmodels. While increasing greenhouse gases (such as carbon dioxide, methane,and nitrous oxide) are the primary drivers of global climate change, the Antarctic
ozone hole, which reoccurs every spring since the early 1980s, was shown tocontribute indirectly to observed changes in Southern Hemisphere surface climate
during summer due to its effects on atmospheric circulation.
The severe springtime ozone depletion over the Antarctic leads to a strong
cooling of the polar lower stratosphere persisting into early summer. This
increases the temperature contrast between the tropics and the polar region,
which strengthens stratospheric winds. For reasons that are not well understood
but are consistently reproduced in models, this leads to a poleward shift of
tropospheric circulation features including the tropical Hadley cell (which
determines the location of the subtropical dry zones) and the midlatitude jet(which is associated with weather systems). There is some evidence from bothmodels and observations that subtropical and midlatitude summer precipitation
patterns have been affected by these changes. The observed wind changes over
the Southern Ocean have also likely driven significant changes in ocean currents.
Model studies suggest that long-lived greenhouse gases that cause climatechange exacerbate this shift in summertime Southern Hemisphere troposphericcirculation, but that ozone depletion has been the dominant contributor to thechanges over the last few decades.
During the 21st century, ozone recovery is expected to lead to the reversal of
the above climate impacts driven by the ozone hole. This reversal will result insome cancellation of future Southern Hemisphere circulation changes driven
by greenhouse gas increases. The extent of such a cancellation depends on thegreenhouse gas emissions assumed in future climate projections. The Southern
Hemisphere surface climate response to ozone depletion in other seasons is
weaker. No such links between ozone depletion and regional climate change
have been observed for the Northern Hemisphere.

Section V: Implications of ozone depletion and the Montreal Protocol

67

19

Have Montreal Protocol controlsof ozone-depleting substances also
protected Earth’s climate?
Yes. Many ozone-depleting substances are also potent greenhousegases that contribute to climate forcing when they accumulate
in the atmosphere. Montreal Protocol controls have led toa substantial reduction in the emissions of ozone-depleting
substances (ODSs) over the last two decades. These reductions,while protecting the ozone layer, have the additional benefit ofreducing the human contribution to climate change. Without
Montreal Protocol controls, the contribution to climate forcing fromannual ODS emissions could now be 10-fold larger than its present
value, which would be a significant fraction of the climate forcing
from current carbon dioxide (CO2) emissions. Increases in ODS
substitute gases, which are also greenhouse gases, could offset
much of this climate benefit by substantially contributing to humaninduced climate forcing in the coming decades.
The success of the Montreal Protocol in controlling the production and consumptionof ozone-depleting substances (ODSs) has protected the ozone layer (see Q15). The
resulting reductions in atmospheric abundances of ODSs also reduced the human
influence on climate because all ODSs are greenhouse gases (see Q18). By protecting
both ozone and climate, the Montreal Protocol has provided a dual benefit to society
and Earth’s ecosystems. In the following, the dual benefit of the Montreal Protocol ishighlighted by considering long-term baseline and world-avoided scenarios of ODS
emissions that use Ozone Depletion Potentials (ODPs), Global Warming Potentials(GWPs), equivalent effective stratospheric chlorine (EESC), and the radiative forcingof climate. However, the climate benefit of the Montreal Protocol may be significantly
offset by increasing future use of ODS substitute gases such as HFCs.
Baseline ODS scenarios. The baseline scenarios of past and future ODS emissions
presented here include the emissions of principal halogen source gases assuming
compliance with the Montreal Protocol provisions. They are constructed from (1)historical annual production and consumption of individual ODSs reported to theMontreal Protocol, (2) projected annual production and consumption of ODSs for
future years based on provisions of the Protocol, (3) estimates of ODS banks, (4)atmospheric observations of ODSs and some naturally occurring halogen sourcegases, such as methyl chloride (CH3Cl), and (5) weighting factors related to ozone
depletion and climate change.
In forming two of the baseline scenarios shown in Figure Q19-1 (upper panels),
the emissions of each gas are added together after being weighted (multiplied) by
the Ozone Depletion Potential (ODP) or the Global Warming Potential (GWP) of therespective gas (see Q18 and Table Q7-1). In the ODP-weighted scenario, the sumof emissions is expressed as CFC-11-equivalent emissions because CFC-11 is thereference gas, with an assigned ODP value of 1. For example, in the sum, 1 kg of
halon-1211 emissions is added as 6.9 kg of CFC-11-equivalent emissions because
the ODP of halon-1211 is 6.9. Similarly, the GWP-weighted sum is expressed as
CO2-equivalent emissions because CO2 is the reference gas, with an assigned GWP of
1. For example, in the sum, 1 kg of carbon tetrachloride emissions is added as 1730
kg of CO2-equivalent emissions because the GWP of carbon tetrachloride is 1730.
World-avoided ODS scenarios. The baseline scenario of ODS emissions can becontrasted with a scenario of ODS emissions that the world has avoided by agreeing
to the Montreal Protocol (see Figure Q19-1). These world-avoided emissions are

68

Section V: Implications of ozone depletion and the Montreal Protocol

The Montreal
ProtocolOzone
ClimateThe Montreal
ProtocolProtection
Protection ofofOzone
andandClimateFrom global emissions of all ozone-depleting substances (ODSs) and CO2
Emissions weighted by
Global Warming Potentials (GWPs)

6(gigatonnes CO2-equivalent per year)

4

2

0
1960

8

Emissions

Emissions(megatonnes CFC-11-equivalent per year)

Emissions weighted by
Ozone Depletion Potentials (ODPs)

Montreal
Protocol1980

Ozone layer
protection

Year

2000

40

for 2008 – 2012

30

20

10

0
1960

2020

Kyoto Protocol

2.0 reduction target

Equivalent effective stratosphericchlorine (EESC)

Montreal
Protocol1980

Climateprotection

Year

2000

2020

Radiative forcing of climate

6

Ozone layer
protection
2

0
1960

Radiative forcing

4(watts per square meter)

EESC (relative amounts)

2.0

Montreal
Protocol

1980

Year

2000

2020

1.5

1.0

0.5

0
1960

Climateprotection
Montreal
Protocol

1980

Year

2000

Baseline ODS scenarios

World-avoided scenarios

CO2 scenarios

Natural halogen sources

3% Annual growth from 1987
2% without the Montreal Protocol

Lower limit inCO2 scenarios

2020

Figure Q19-1. Montreal Protocol
protection of ozone and climate.
The provisions of the Montreal Protocol
have substantially reduced ozone-depleting
substances (ODSs) in the atmosphere.
This has protected the ozone layer and also
reduced the potential for climate change
because ODSs are greenhouse gases.
The scenarios and comparisons shown
here demonstrate this dual benefit of theMontreal Protocol. Baseline scenarios
for ODS emissions include all principal
gases weighted by their Ozone Depletion
Potentials (ODPs) or Global Warming
Potentials (GWPs) (top panels). With theseweightings, emissions are expressed
as CFC-11-equivalent or CO2-equivalentmass per year. The lower panels show
EESC and radiative forcing of climate asderived from the respective ODP- and GWPweighted scenarios. The world-avoided
emission scenarios assume ODS emission
growth of 2 or 3% per year beyond 1987
abundances. Shown for reference are theemissions and radiative forcing of CO2, andthe emissions reduction target of the firstcommitment period of the Kyoto Protocol.
The contributions of natural halogen sourcegases are shown in the ODP-weighted and
EESC scenarios (red dashed lines) and are
negligible in the GWP-weighted and radiative
forcing scenarios. The magnitude of thedual benefit has increased since about
1987 as shown by differences between
the world-avoided and baseline scenarios(blue shaded regions in each panel). For
completeness, these differences can be
adjusted by offsets due to additional ozone
depletion and HFC emissions (see text).
(A megatonne = 1 billion (109) kilograms.
A gigatonne = 1 trillion (1012) kilograms.)

estimated by assuming that emissions of ODSs in the baseline scenario increasebeyond 1987 values with a 2 or 3% annual growth rate. These growth rates are
consistent with the strong market for ODSs in the late 1980s that included a wide
variety of current and potential applications and that had potential for substantial
new growth in developing countries.
CO2 emission scenarios. Long-term CO2 emission scenarios are also shown forcomparison, as derived from past and projected CO2 emissions, because CO2 is theprincipal greenhouse gas related to human activities. The projected CO2 emissions
have high and low scenarios that are derived using different basic assumptions
about future economies, technical progress, and societal decisions.

Section V: Implications of ozone depletion and the Montreal Protocol

69

19
ODP-weighted emissions scenarios. The ODP-weighted emissions in the ODS
baseline scenario are a measure of the overall threat to stratospheric ozone fromODSs (see Figure Q19-1, upper left panel). When ODP-weighted emissions increase(decrease) in a given year, more (less) ozone will be destroyed in future years.
ODP-weighted emissions increased substantially in the baseline scenario between
1960 and 1987, the year the Montreal Protocol was signed (see Figure Q19-1 andQ0-1). After 1987, ODP-weighted emissions began a long and steady decline topresent-day values. The decline in emissions is expected to continue, causing theatmospheric abundances of individual ODSs to decrease (see Figure Q16-1). The
reduction in ODP-weighted emissions from the 1987 value is a conservative measureof the annual emissions avoided by the Montreal Protocol since 1987 and, hence, ofthe success of the Montreal Protocol in protecting the ozone layer.
Annual ODP-weighted emissions in the world-avoided scenario are about double the 1987
values by 2020. The annual differences between the world-avoided emissions and thebaseline scenario (blue shaded region in Figure Q19-1) provide reasonable upper limits to
the ODP-weighted emissions avoided by the Montreal Protocol each year since 1987.
GWP-weighted emissions scenarios. The GWP-weighted emissions in the ODS
baseline scenario are a measure of the overall threat to climate from ODSs (seeFigure Q19-1, upper right panel). As ODS emissions accumulate in the atmosphere,
their climate forcing contribution increases. The long-term changes in the
GWP-weighted scenario are very similar to those in the ODP-weighted scenario. Bothshow an increase before 1987 and decrease afterwards. The similarity follows fromthe predominant role that CFC-11 and CFC-12 emissions play in ozone depletion andclimate forcing from ODSs. The reduction in GWP-weighted emissions since 1987 is aconservative measure of the substantial success of the Montreal Protocol in reducing
the potential for climate change from human activities. The annual differences
since 1987 between the world-avoided emissions and the baseline scenario (blueshaded region in Figure Q19-1) provide reasonable upper limits to the GWP-weighted
emissions avoided by the Montreal Protocol each year since 1987.
The climate protection calculated using differences between world-avoided emissions
and the baseline scenario has two offsetting effects. The first is the additional ozone
depletion that would be caused by world-avoided ODS emissions. Ozone depletion
offsets ODS climate forcing because a greenhouse gas (ozone) is being removed
from the atmosphere in response to ODS emissions (see Q18). The second effect is
the increase in emissions of HFC substitute gases that occurred in response to ODS
reductions from Montreal Protocol controls. More HFCs in the atmosphere have thepotential to offset much of the gain in climate protection from ODS reductions (seeFigure 19-2) because HFCs are also greenhouse gases (see Q18).
The combined magnitude of these offsets in 2010, for example, is about 30% of
the difference between the baseline and world-avoided scenarios. The resulting
net GWP-weighted emission reduction in 2010 is about 9.7–12.5 gigatonnes
CO2-equivalent per year. In contrast, the annual emissions reduction target adopted
by the Kyoto Protocol during its first commitment period (2008–2012) is estimated as
2 gigatonnes CO2-equivalent per year (see Figure Q19-1). The reductions are expected
to result from controlling the Kyoto Protocol basket of gases that includes HFCs anddoes not include ODSs (see Q18). As a result, the upper limit for the net reduction in
annual GWP-weighted emissions achieved by the Montreal Protocol in 2010 is 5- to
6-fold larger than the Kyoto Protocol target.
Annual GWP-weighted emissions of ODSs were a large percentage (about 20–40%)
of CO2-baseline emissions between 1960 and 1987 (see Figure Q19-1). Thereafter,
this percentage has steadily decreased and is projected to reach 2–3% by 2020.

70

Section V: Implications of ozone depletion and the Montreal ProtocolThis projection stands in sharp contrast to the world-avoided scenario, in which thepercentage increases to 40–75% of CO2-baseline emissions by 2020.
EESC scenarios. The EESC scenario in Figure Q19-1 (lower left panel) provides
a measure of the year-to-year potential of the atmospheric abundances of ODSs
to destroy stratospheric ozone. Changes in historical and projected atmosphericemissions of ODSs cause changes in their atmospheric abundances. The derivationof EESC from ODS atmospheric abundances is discussed in Q16 and similar EESC
baseline scenarios are shown in Figures Q14-1, Q15-1, and Q16-1 for different timeintervals. An increase in ODP-weighted emissions always leads to some increase in
EESC in the years following the emissions. When ODS-weighted emissions decreased
after 1987, EESC did not proportionally decrease because of the long atmospheric
lifetimes of the principal ODSs. In Figure Q19-1, for example, EESC reached its peaknearly a decade after the peak in ODP-weighted emissions, and by 2010 the decrease
in EESC from its peak value was only about 10%, compared to the 70% decrease in
ODP-weighted emissions achieved by 2010.
Implications of HFC Usage for Climate

Emissions and radiative forcing of CFCs, HCFCs, and HFCs
Emissions weighted by
Global Warming Potentials (GWPs)

3.5

Radiative forcing

40
30
20
10
0
1960

Radiative forcing of climate

3.0

50(watts per square meter)

Emissions(gigatonnes CO2 - equivalent per year)

60

2.5
2.0
1.5
1.0
0.5

1980

2000

Year

2020

2040

World-avoided scenarios
3% Annual growth from 1987
2% without the Montreal Protocol

0.0
1960

1980

CO2 scenarios
Lower limit inCO2 scenarios

2000

Year

2020

2040

Baseline scenarios
HFCs
CFCs
HCFCs

Figure Q19-2. Implications of HFC usage for climate.
The provisions of the Montreal Protocol have reduced ozone-depleting substances (ODSs),
but as a result increased the demand for ODS substitute gases. Most of the hitherto
available substitute gases such as HCFCs and HFCs have large Global Warming Potentials(GWPs), and hence future emissions of these gases have the potential to offset the climate
benefit gained from the Montreal Protocol. A baseline scenario is calculated using current
HCFC consumption in leading applications, patterns of replacement of HCFCs by HFCs
in developed countries, and economic growth. The baseline scenario for HFC emissions
includes four major HFCs weighted by their GWP (expressed as CO2-equivalent mass peryear) (left panel). Note that the used baseline scenario represents average increases in
HFC emissions only. Radiative forcing as calculated from the GWP-weighted scenarios
for CFCs, HCFCs and HFCs is shown in the right panel. By 2050, the projected increases
in HFC abundances are sufficient to significantly offset the climate benefit achieved by
the Montreal Protocol and produce a radiative forcing that is approximately 15% of the
projected CO2 value.

Section V: Implications of ozone depletion and the Montreal Protocol

71

19
Radiative forcing of climate. The radiative forcing derived for the ODS baseline
scenario in Figure Q19-1 (lower right panel) provides a measure of the year-to-year
contribution to climate forcing from atmospheric ODS abundances. The radiative
forcing of an ODS is proportional to its radiative efficiency (that is, the radiative
forcing per abundance change) and the net increase in its atmospheric abundance
during the Industrial Era. Increases in abundance up to the present are derived
from atmospheric observations. Future abundances rely on projected emissions andatmospheric lifetimes of each gas. In Figure Q19-1, radiative forcing due to ODSs
increases smoothly from 1960 onward, peaks in 2003, and decreases very graduallyin subsequent years. Radiative forcing responds to ODS emission reductions ina manner similar to EESC, with the current slow decline attributable to the two
principal contributing gases, CFC-11 and CFC-12, and their long atmospheric lifetimes(50–100 years).
The differences in ODS climate forcing between the world-avoided and baseline
scenarios are offset by additional ozone depletion and HFC emissions in a manner
similar to that noted above for differences in GWP-weighted emissions. Afteraccounting for these two offsets, the climate forcing due to ODSs in the world-avoided
scenario is approximately 70% higher than that in the baseline scenario in 2010 andapproximately 30% of that due to CO2.
The considerable contributions that ODSs could have made to climate forcing, if notcontrolled by the Montreal Protocol, attests to their potency as greenhouse gases.
ODSs had negligible atmospheric abundances 50–60 years ago and, as a group,
represent chlorine amounts that currently are about 100,000 times less abundant in
the atmosphere than CO2.
Implications of HFC usage for climate. HFCs are currently principal substitutegases for ODS uses. Many of these HFCs are also potent greenhouse gases. In
the coming decades, the projected growth in HFC emissions has the potential tooffset much of the climate benefit gained from the Montreal Protocol controls,
depending on emissions scenarios. An average future projection scenario basedon current production and consumption patterns and projected economic growthsuggests that increasing HFC demands could result in GWP-weighted emissionsof around 5 gigatonnes CO2-equivalent per year by 2050 (see Figure Q19-2),
primarily in developing nations. The 2050 value is comparable to half of the peak in
GWP-weighted ODS emissions in 1987 (see Figure Q19-1). The resulting radiative
forcing from these projected emissions would reach levels of the early 2000s andcontribute almost 10% of that due to CO2 to the total anthropogenic climate forcing
by 2050.
This projection assumes that HFC application demand in developing countries would
be met using the same suite of HFCs currently used in developed countries. If future
HFC demand is met instead with lower-GWP substances, the 2050 projection would
be substantially reduced.
Increasing the benefits of the Montreal Protocol into the future. Fewer control
options are available to increase the dual benefit of the Montreal Protocol beyond
2020 because the most effective and abundant ODSs have already been phased
out under Montreal Protocol provisions (see Q16). The most recent Montreal Protocol
action increased ozone and climate protection by accelerating the phase-out of
HCFCs (Montreal in 2007) (see Q15). This provision is expected to reduce total
GWP-weighted emissions of HCFCs by about 50% between 2010 and 2050,
corresponding to about 18 gigatonnes of CO2-equivalent emissions.

72

Section V: Implications of ozone depletion and the Montreal Protocol

Ozone and climate protection could be further enhanced with Montreal Protocol
provisions that increase the effectiveness of capturing and destroying ODSs
contained in halon, CFC, and HCFC banks, namely, those ODSs currently being usedin refrigeration, air conditioning, and fire protection equipment, or stockpiled forservicing long-term applications. Current banks are projected to contribute more toozone depletion in the coming decades than the limited future production of ODSs
allowed by the Montreal Protocol provisions.
The available options, such as cutting the remaining production of HCFCs, halting
emissions of CCl4 and quarantine or pre-shipment CH3Br, as well as destroying halon,
CFC, and HCFC banks, would together advance the return of the ozone layer to 1980
levels by 11 years.

Section V: Implications of ozone depletion and the Montreal Protocol

73

20

How is ozone expected to change
in the coming decades?
Substantial recovery of the ozone layer from the effects of ozonedepleting substances (ODSs) is expected near the middle ofthe 21st century, assuming global compliance with the Montreal
Protocol. Recovery will occur as ODSs and reactive halogen gasesin the stratosphere decrease in the coming decades. In addition
to responding to ODSs, future ozone amounts will increasingly be
influenced by expected changes in climate. The impacts of futureclimate change on the ozone layer will vary between the tropics,
midlatitudes, and polar regions, and strongly depend on future
emissions of carbon dioxide (CO2), methane (CH4), and nitrous oxide(N2O). During the long recovery period, large volcanic eruptions could
temporarily reduce global ozone amounts for several years.
Substantial recovery from the depletion of global and polar ozone caused by ODSs is
expected in the later decades of this century. The recovery follows from the successof the Montreal Protocol in reducing the global production and consumption of ODSs.
Now, the atmospheric abundances of most major ODSs and the associated valuesof equivalent effective stratospheric chlorine (EESC) are in decline (see Q16). In
contrast to the diminishing role of ODSs, changes in climate are expected to have an
increasing influence on future ozone. Climate change is driven by the projected growthin greenhouse gas abundances, primarily carbon dioxide (CO2), methane (CH4), andnitrous oxide (N2O), which will lead to changes in temperatures, chemistry, and thecirculation 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 indifferent geographical regions during the recovery period. Global events, such as major
volcanic eruptions and geoengineering actions, may also influence future ozone levels.
Using chemistry-climate models. Projections of total ozone presented here arebased on the results from a group of chemistry-climate models that take into
account the influences of changes in ODSs and climate. These models show howchanges in ozone are expected to vary across global regions by evaluating thecomplex interactions of the processes that control ozone and climate involving
radiation, chemistry, and transport. Required model inputs include, for example,
historical and projected concentrations of ODSs, CO2, CH4, and N2O, and changes
in sea surface temperatures. The results from chemistry-climate model simulations
are used to identify aspects of the models that are particularly important for futureozone abundances. For example, model projections for the coming decades show
a strengthening in the atmospheric circulation that brings air to the stratosphere in
the tropics, moves air poleward into both hemispheres, and then returns it to thetroposphere at middle to high latitudes. These circulation changes will significantly
alter the global distribution of ozone and the atmospheric lifetimes of ODSs andother long-lived gases. Also, while Earth’s surface is expected to continue to warm
in response to positive radiative forcing from greenhouse gases, the stratosphere is
expected to continue to cool. A colder upper stratosphere leads to increases in ozone
because lower temperatures decrease the effectiveness of gas-phase ozone-loss
reactions. Finally, CH4 and N2O are both involved in the chemistry that determines
stratospheric ozone levels.
Simulating recent ozone changes. Comparisons of model results with observations
help confirm the causes of ozone depletion and increase confidence in model
projections of future ozone amounts. Two important measures of ozone, global totalozone and minimum total ozone values in the Antarctic (available since 1980), are74

Section VI: Stratospheric ozone in the futurecompared to a group of chemistry-climate model simulations in Figure Q20-1. Bothozone measures show substantial depletion since 1980. The average model value
follows the observed decline in both ozone measures, indicating that the mainprocesses involved in ozone depletion are reasonably well represented by the models.
Some differences between the simulations and observations can be explained byunusual meteorological conditions, volcanic eruptions, changes in solar activity, orother natural influences, which are not all fully accounted for in the various models
used here. In Figure Q20-1, global total ozone has increased in the last 20 years,
while Antarctic minimum ozone values have been relatively constant. The global totalozone increases cannot be attributed solely to reductions in ODSs that began in the
1990s, because total ozone has also recovered from additional depletion caused by
the 1991 Mt. Pinatubo eruption (see Q14).
Equivalent stratospheric chlorine (ESC) projections. Equivalent stratosphericchlorine (ESC) values are also projected in chemistry-climate models to represent
how the potential for reactive halogen gases (see Q8) to destroy ozone varies overtime. ESC is similar to equivalent effective stratospheric chlorine (EESC) (see Q16)
in that both represent a weighted sum of reactive chlorine and bromine gases inthe stratosphere. The weighting in each accounts for the greater effectiveness ofbromine atoms in destroying ozone. ESC can be calculated more precisely andcomprehensively than EESC because the chemistry-climate models can derive ESC
as a function of altitude, latitude, longitude, and time. The ESC calculations are
based on the history and projections of ODS surface abundances and the chemical
and transport processes that control (1) the conversion of ODSs to reactive halogen
gases, (2) the distribution of reactive halogen gases in the global stratosphere, and(3) their ultimate removal from the stratosphere.

Simulationsof Stratospheric
OzoneOzone
Depletion
Simulationsof Stratospheric
Depletion
Results
fromfromchemistry-climate
modelsResultschemistry-climate
models5

Global total ozone change from 1980

0
-5
-10

Annually averaged
observations-15

Average and rangeof atmosphericmodel predictions

-20
-25

Minimum total ozone (DU)

Figure Q20-1. Simulations of ozone depletion.
Chemistry-climate models of the atmosphere that account forchanges in ozone-depleting substances (ODSs) and climateparameters are widely used to assess past ozone changes andproject future ozone. Agreement in comparisons of model resultswith observations increases confidence in the model projections
and our understanding of the processes leading to ozone depletion.
Observed values of global total ozone (top panel) and minimum totalozone values over Antarctica (bottom panel) decreased beginning in
the early 1980s (red points). Ozone is no longer decreasing in either
region. The observations lie mostly within the range of projections
derived from a group of chemistry-climate models (blue shading).
Differences between models and observations can occur due tounusual meteorological conditions, volcanic eruptions, changes
in solar activity, or other influences, which are not fully accounted
for in the different models. As ODS abundances decrease in the
21st century, chemistry-climate models project global total ozone toincrease steadily and exceed 1980 values, and Antarctic minimum
ozone to return to 1980 values.

Total ozone change (DU)

10

250

Antarctic minimum total ozone

200
150
100

Averages fromobservations

50

Average and rangeof atmosphericmodel predictions

0
1980

2000

2020

2040

Year

Section VI: Stratospheric ozone in the future75

20
The long-term changes in ESC in the global and regional analyses shown in Figure
Q20-2 are highly similar. In all regions, changes in ESC from 1960 values increasesmoothly with time, reach a peak near the end of the 20th century, and decrease
gradually until the end of the 21st century. Values at the end of the 21st century
are approaching those in 1960, indicating that ODSs are largely removed from thestratosphere by that time. ESC returns to 1980 values several decades sooner than
to 1960 values because of the slow rate of decline.
Peak ESC values around the year 2000 are highest in polar regions and lowest in
the tropics. In the tropics, stratospheric air has only recently been transported fromthe troposphere, with the result that only a small fraction of ODSs has undergone
conversion to reactive halogen gases (see Q8). In polar regions, the fraction is muchlarger because stratospheric air requires several years on average to journey fromits entry point in the tropics to the polar lower stratosphere. During this time, a muchlarger fraction of ODSs undergoes conversion to reactive halogen gases.
Long-term total ozone projections. Total ozone changes derived from chemistryclimate models, as referenced to 1960 values, are shown in Figure Q20-2. The rangeof values from the group of models is included in the figure as one measure of theuncertainty in the model projections. Total ozone changes in the different regions aredescribed as follows:• Antarctic. Total ozone changes are largest in the Antarctic region in springtime(October). Chemistry-climate models show that ODSs are the predominant factor
in Antarctic ozone depletion in the past and in the coming decades. Changes inclimate parameters have a smaller role. As a result, total ozone changes mostly
mirror changes in ESC: as ESC increases, ozone proportionately decreases; as
ESC decreases, ozone proportionately increases. Antarctic total ozone is projected
to return to 1980 levels after midcentury and later than in any other region.
Meteorological variability in the Antarctic in late winter/early spring when ozone
depletion occurs causes a large range in the model projections.
• Arctic. Total ozone changes in the Arctic region in springtime (March) areconsiderably smaller than in the Antarctic. In contrast to the Antarctic, ozone
changes do not closely mirror changes in ESC. After midcentury, Arctic total ozone
increases to values above those that would be expected from ESC reductions alone
because of the strengthening of the atmospheric circulation and the enhanced
stratospheric cooling associated with increases in CO2. By 2100 Arctic total ozone
is projected to be well above both 1960 and 1980 values. The large range inprojections compared to other nonpolar regions is due to greater meteorologicalvariability, as noted for the Antarctic. Arctic total ozone is projected to return to1980 levels between 2020 and 2035, which is two to three decades before ESC
returns to 1980 levels, and to continue to increase until the end of the century.
• Northern and southern midlatitudes. The annual averages of total ozone
changes in midlatitudes are much smaller than the springtime losses in polar
regions. Both midlatitude regions resemble the Arctic, with total ozone returning
to 1960 and 1980 values much sooner than ESC. In the northern midlatitudes,
the models predict a return of total ozone to 1960 values in 2030, whereas ESC
requires the full century to return close to 1960 values. Total ozone in southern
midlatitudes changes in a manner similar to that in the northern midlatitudes.
Differences are that in southern midlatitudes, the 1960 return date for total ozone
is somewhat later (2055) and the maximum ozone depletion observed near 2000
is greater. Both aspects reflect the influence of Antarctic polar ozone depletion,with stratospheric air low in ozone being transported to southern midlatitudes inspring after the breakup of the polar vortex (see Q11). The more rapid return of76

Section VI: Stratospheric ozone in the futureTotalOzone
andEquivalentStratospheric
Stratospheric Chlorine
Total
Ozone
and
Equivalent
Chlorine

Results from atmospheric chemistry-climate models for 1960 to 2100
GlobalArctic (March)

80

4

ESC

8
4
0

-40

-80

ESC

4

Ozone
0

0

-40
0

Northern midlatitudes

0 40

0

ESC

40

40

1980

2020

2060

Year

Ozone

2100
-40
1980

Ozone

2020

2060

2100

Tropics
1980

2020

2060

2100

40

Antarctic (October)
ESC

40
0

4

ESC
0

0

4

Ozone
-40

0

1980

-40

200

300

-80
-120

Ozone

-160
1980

2020

Year

2060

2100

(2090 – 2099 average)

40

Annual (or monthly) meanchanges from 1960 values
Total ozone (DU)
Equivalent stratosphericchlorine (ESC) (ppb) near
21- km altitude1980 values

0

Ranges of model values

2060

2100

Southern midlatitudes

400 450

Total ozone (DU)

2020

ESC

4
0

Ozone

-40
1980

2020

Year

2060

2100

Figure Q20-2. Long-term changes in ozone and equivalent stratospheric chlorine (ESC).
Chemistry-climate models are used to make projections of total ozone amounts that
account for the effects of ozone-depleting substances (ODSs) and climate change.
Regional and global projections are shown for total ozone and ESC for the period
1960–2100, referenced to 1960 values. The globe in the center shows average
total ozone projections for the last decade of the 21st century. Total ozone depletion
increased after 1960 as ESC values steadily increased throughout the stratosphere.
ESC values have peaked and are now in a slow decline. All the projections show
maximum total ozone depletion around 2000, coincident with the highest abundancesof ESC. Thereafter, total ozone increases, except in the tropics, as ESC slowly declines.
In all the projections except the Antarctic and the tropics, total ozone returns to 1960
values by midcentury, which is earlier than expected from the decrease in ESC alone.
The earlier returns are attributable to climate change, which influences total ozone
through changes in stratospheric transport and temperatures. In the tropics, in contrast,
climate change causes total ozone to remain below 1960 values throughout thecentury. In the Antarctic, the effect of climate change is smaller than in other regions.
As a result, Antarctic total ozone in springtime mirrors the changes in ESC, with bothclosely approaching 1960 values at century’s end. The dots on each curve mark theoccurrences of 1980 values of total ozone and ESC. Note that the equal vertical scales
in each panel allow direct comparisons of ozone and ESC changes between regions.

Section VI: Stratospheric ozone in the future77

20
total ozone in both regions compared with ESC again is the result of the influenceof climate-induced changes in transport and upper stratospheric temperaturesin the model projections. The model range is significantly less in midlatitude
annual averages than in the polar springtime averages. After returning to either
1960 or 1980 values, total ozone continues to increase in both hemispheres andsignificantly exceeds those values by 2100.
• Tropics. Total ozone changes in the tropics are smaller than in any other region.
Ozone is less sensitive to ESC in the tropical stratosphere because of the dominantroles of production and transport in controlling ozone and the low amounts of ESC
available in this region. In contrast to other regions, chemistry-climate models
project total ozone to remain below 1960 values throughout the 21st century.
Total ozone gradually returns to 1980 values in 2040, peaks around 2060, anddecreases again until the end of the century. The different total ozone behavior in
the tropics is due to the changing balance between ozone increases in the upper
stratosphere and decreases in the lower stratosphere. The ozone increases in
the upper stratosphere occur from declining ODS abundances and decreases intemperatures caused by increasing greenhouse gases, primarily CO2. In the lower
stratosphere, ozone is reduced because the strengthening of the stratosphericcirculation reduces the time for ozone production before air leaves the tropics.
These circulation changes also influence the Arctic and midlatitude regions as
noted above.
• The globe. The annual averages of global total ozone are projected to return to1960 levels around the middle of the century (2040–2080) while ESC returns
to 1960 values near the century’s end. The return dates for 1980 values aresubstantially earlier. Chemistry-climate model analysis suggests that this early
return of total ozone is primarily a result of upper stratospheric cooling. The larger
total ozone changes shown for polar regions in springtime do not significantly
influence global total ozone values because these values are annual averages andpolar regions are a small geographical fraction of the globe.
Future ultraviolet radiation. Projections of long-term changes in total ozone can beused to estimate long-term changes in solar ultraviolet radiation reaching Earth’s
surface (see Q17). The UV-B component of ultraviolet radiation increases as totalozone decreases. Based on the global total ozone projections, clear-sky UV-B
radiation is expected to decrease below 1960 values by the end of the century
because ozone is above 1960 values in the later decades. The latitude regions
expected to experience increased ultraviolet radiation levels are the Antarctic andthe tropics, where total ozone remains lower than 1960 values until the end of the
century. Large increases or decreases in surface erythemal radiation away from 1960
values are expected to lead to adverse effects on human and ecosystem health.
Greenhouse gas emission scenarios and impact on future ozone. The climate
impacts on future ozone discussed above have been derived using a particulargreenhouse gas emission scenario. Emission scenarios specify the amounts ofgreenhouse gases emitted into the atmosphere in the future and are constructedbased on socio-economic assumptions on future population growth, economicaldevelopments, technology innovation, and political decisions. Future emissions arehighly uncertain and therefore various emission scenarios are designed to cover a wide
range of possible future outcomes. Chemistry-climate model simulations are run using
the different emission scenarios to determine the range of impacts climate change
may have on the ozone layer. Such simulations show that the evolution of ozone andthe resulting surface UV levels in different regions of the world are strongly dependent
on the mixture and relative abundances of greenhouse gases that will be used.

78

Section VI: Stratospheric ozone in the future

Global
Total
Changes in
GlobalTotalOzone
Ozone Changes
Response
totoIncreasing
Greenhouse
Gases
in Response
Increasing Greenhouse
Gases
4

10

CO2 only00

All gases
N2O only

-10

-2
-4

ODSs only

-6

-20
1960

2

2000

2040

Year

2080

Ozone change (%)

Ozone change (DU)

CH4 only

Figure Q20-3. Global total ozone changes in response toincreasing greenhouse gases.
Model simulations that include changes in each of thegreenhouse gases CO2, N2O, and CH4 separately illustrate
how a particular gas affects global total ozone. Increasing CO2
leads to increases in ozone, because CO2 lowers stratospherictemperatures and slows down ozone destruction reactions.
Increasing CH4 leads also to increases in ozone, while N2Oleads to decreases in ozone similarly to the ODSs. The results
are contrasted with a model simulation that includes changes
in all greenhouse gases together. The combined impact of thesegases on total global ozone is not simply the sum of the effectsof each gas, but depends strongly on the relative abundancesof the gases.

Simulations with a simplified model using increases in CO2, N2O, and CH4 separately
illustrate the reasons for these dependencies (see Figure Q20-3). The simulation
using increases in CO2 only shows that the main effect of CO2 is to increase global
ozone in the future. The ozone increases are explained mainly by decreases instratospheric temperatures. As detailed above, lower temperatures in the upper
stratosphere slow down the chemical gas-phase reactions that lead to ozone
depletion. Note that circulation changes only play a minor role for global mean ozone
since decreases in total ozone in the tropics and increases at higher latitudes largely
cancel. Simulations using increases in CH4 and N2O separately show that the main
effect of CH4 is to increase ozone, while that of N2O is to decrease ozone in the future.
The reason for this behavior is that N2O and CH4 are source gases of reactive gases(NOx and HOx) that strongly affect stratospheric ozone chemistry. The reactive gases
produced from N2O lead to stratospheric ozone depletion, while those from CH4 leadto the production of stratospheric ozone.
Volcanoes and geoengineering. Other factors not included in chemistry-climate
models can potentially affect future total ozone amounts. Explosive volcanic eruptions
have temporarily reduced global total ozone in the past (see Q14). Similar eruptions,especially in the early decades of this century when ESC values are highest, are also
expected to reduce total ozone for a few years. Volcanic eruptions are an additionalsource of uncertainty not included in the ozone projections in Figures Q20-2 and Q20-3.
Several geoengineering methods have been proposed to reduce climate forcing
from human activities. A widely discussed proposal is the enhancement of sulfate
aerosols in the stratosphere from direct injections of sulfur or sulfuric acid. The
expected response is a cooling of the climate system from increased aerosol
scattering of sunlight, similar to that observed after some explosive volcanic
eruptions. The required injections, if sustained over many years, are likely to haveunintended consequences such as reductions in total ozone amounts and changes
in stratospheric temperatures and circulation.

Section VI: Stratospheric ozone in the future79World Meteorological Organization
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Published January 2015
ISBN: 978-9966-076-02-1
The Scientific Ozone Assessment report including the Twenty Questions and Answers
is available electronically via the following links:
http://www.wmo.int/pages/prog/arep/gaw/ozone_2014/ozone_asst_report.html
http://ozone.unep.org/en/assessment_panels_bodies.php?committee_id=7
http://esrl.noaa.gov/csd/assessments/ozone/
Note: Figures from this report are in the public domain and may be used with proper
attribution to source.

Citation information
Michaela I. Hegglin (Lead Author), David. W. Fahey, Mack McFarland, Stephen A.
Montzka, and Eric R. Nash, Twenty Questions and Answers About the Ozone Layer:
2014 Update, Scientific Assessment of Ozone Depletion: 2014, 84 pp., World
Meteorological Organization, Geneva, Switzerland, 2015.

Sponsors
We are grateful to the UNEP Ozone Secretariat, the National Oceanic and Atmospheric
Administration, the European StratoClim Project, the University of Cambridge throughfunding from the Department for the Environment, Food, and Rural Affairs (Defra),and the University of Reading for providing funding for the layout, print, and designof the Twenty Questions and Answers About the Ozone Layer: 2014.

B08785 03.15