Atmospheric Science for Environmental Scientists E-Book
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- Herausgeber: John Wiley & Sons
- Kategorie: Wissenschaft und neue Technologien
- Sprache: Englisch
Enlightens readers on the realities of global atmospheric change, including global warming and poor air quality
Climate change and air pollution are two of the most pressing issues facing Mankind. This book gives undergraduate and graduate students, researchers and professionals working in the science and policy of pollution, climate change and air quality a broad and up-to-date account of the processes that occur in the atmosphere, how these are changing as Man’s relentless use of natural resources continues, and what effects these changes are having on the Earth’s climate and the quality of the air we breathe.
Written by an international team of experts, Atmospheric Science for Environmental Scientists, 2nd Edition provides an excellent overview of our current understanding of the state of the Earth’s atmosphere and how it is changing. The first half of the book covers: the climate of the Earth; chemical evolution of the atmosphere; atmospheric energy and the structure of the atmosphere; biogeochemical cycles; and tropospheric chemistry and air pollution. The second half looks at cloud formation and chemistry; particulate matter in the atmosphere; stratospheric chemistry and ozone depletion; boundary layer meteorology and atmospheric dispersion; urban air pollution; and global warming and climate change science.
- Provides succinct but detailed information on all the important aspects of atmospheric science for students
- Offers the most up-to-date treatment of key issues such as stratospheric chemistry, urban air pollution, and climate change
- Each chapter includes basic concepts, end-of-section questions, and more in-depth material
- Features contributions from the best experts and educators in the field of atmospheric science
Atmospheric Science for Environmental Scientists, 2nd Edition is an invaluable resource for students, teachers, and professionals involved in environmental science. It will also appeal to those interested in learning how the atmosphere works, how humankind is changing its composition, and what effects these changes are leading to.
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Table of Contents
Cover
List of Contributors
Preface
Abbreviations, Constants, and Nomenclature
Constants
Nomenclature
1 The Climate of the Earth
1.1 Basic Climatology
1.2 General Atmospheric Circulation
1.3 Palaeoclimates
1.4 Polar Climates
1.5 Temperate Latitude Climates
1.6 Tropical Climates
Questions
References
Further Reading
2 Chemical Evolution of the Atmosphere
2.1 Creation of the Planets and Their Earliest Atmospheres
2.2 Earth's Atmosphere before Life Began
2.3 Comparison of Venus, Earth, and Mars
2.4 Life and Earth's Atmosphere
2.5 Carbon Dioxide in Earth's Atmosphere
2.6 The Rise of Oxygen Concentrations
2.7 Protection of Life from Ultraviolet Radiation
2.8 The Great Oxidation Event and Related Issues
2.9 The Future
Questions
References
Further Reading
3 Atmospheric Energy and the Structure of the Atmosphere
3.1 The Vertical Structure of Earth's Atmosphere
3.2 Solar and Terrestrial Radiation
3.3 Solar Radiation, Ozone, and the Stratospheric Temperature Profile
3.4 Trapping of Longwave Radiation
3.5 A Simple Model of Radiation Transfer
3.6 Light Scattering
3.7 Conduction, Convection, and Sensible and Latent Heat
3.8 Energy Budget for Earth's Atmosphere
3.9 Aerosols, Clouds, and Climate
3.10 Solar Radiation and the Biosphere
3.11 Summary
Questions
References
Further Reading
4 Biogeochemical Cycles
4.1 Sources
4.2 Sinks
4.3 Carbon
4.4 Nitrogen
4.5 Sulphur
4.6 Halogens
4.7 Hydrogen
4.8 Summary
References
Further Reading
5 Tropospheric Chemistry and Air Pollution
5.1 Sources of Trace Gases in the Atmosphere
5.2 Key Processes in Tropospheric Chemistry
5.3 Initiation of Photochemistry by Light
5.4 Tropospheric Oxidation Chemistry
5.5 Night‐Time Oxidation Chemistry
5.6 Halogen Chemistry
5.7 Air Pollution and Urban Chemistry
5.8 Summary
Questions
References
Further Reading
6 Cloud Formation and Chemistry
6.1 Clouds
6.2 Cloud Formation
6.3 Particle Size and Water Content
6.4 Dissolved Solids in Cloud Water and Rainfall
6.5 Dissolution of Gases
6.6 Reactions and Photochemistry
6.7 Radical and Photochemical Reactions
6.8 Summary
References
Further Reading
Websites
7 Particulate Matter in the Atmosphere
7.1 Aerosol Properties
7.2 Aerosol Sources
7.3 The Role of Atmospheric Particles
7.4 Aerosol Measurements
7.5 Summary
Acknowledgement
Questions
References
8 Stratospheric Chemistry and Ozone Depletion
8.1 Ozone Column Amounts
8.2 Physical Structure of the Stratosphere
8.3 Gas‐Phase Chemistry of the Stratosphere
8.4 Aerosols and Clouds in the Stratosphere
8.5 Heterogeneous Chemistry of the Stratosphere
8.6 Future Perturbations to the Stratosphere
8.7 Summary
Questions
References
9 Boundary Layer Meteorology and Atmospheric Dispersion
9.1 The Atmospheric Boundary Layer
9.2 Flow over Vegetation
9.3 The Urban Boundary Layer
9.4 Dispersion of Pollutants
9.5 Summary
Questions
Hints and Answers
References
Further Reading
10 Urban Air Pollution
10.1 Introduction
10.2 Urban Air Pollution – A Brief History
10.3 Scale of Urban Air Pollution
10.4 Air Pollutants and Their Sources in the Urban Atmosphere
10.5 From Emissions to Airborne Concentrations
10.6 Urban‐Scale Impacts
10.7 Means of Mitigation
10.8 Summary
Acknowledgement
Questions
References
Further Reading
11 Global Warming and Climate Change Science
11.1 Historical Evidence of the Impact of Human Activities on Climate
11.2 Future Outlook of Climate Change
11.3 The Integrated Science Assessment Modelling (ISAM)
11.4 Potential Impacts of Climate Change
11.5 Summary
Acknowledgements
Questions
References
Appendix: Suggested Web Resources
Index
End User License Agreement
List of Tables
Chapter 1
Table 1.1 Mean temperature data (°C) for South Pole (1957–2010): latitude 90°...
Table 1.2 Tropical storm
a
development regions.
Chapter 3
Table 3.1 Typical values of key quantities measured during the
South African R
...
Table 3.2 Stability criteria for a moist air parcel. The ambient reduction in...
Chapter 4
Table 4.1 Examples of rate coefficients for the reaction of hydrocarbons with...
Table 4.2 Global budget for an anthropogenic perturbation of CO
2
.
Table 4.3 Sources and sinks for methane in the atmosphere expressed as Tg CH
4
Table 4.4 Global budget for carbon monoxide in the atmosphere.
Table 4.5 Estimates of the global sources and sinks for N
2
O in Tg N per year....
Table 4.6 Atmospheric budget for NO
x
and NH
3
(TgN year
−1
).
Table 4.7 Global annual sources and sinks of OCS, CS
2
, H
2
S, DMS and SO
2
in Tg...
Table 4.8 Historical applications (prior to controls over use) of chlorofluor...
Table 4.9 Atmospheric budget for CH
3
Cl and CH
3
Br.
Chapter 5
Table 5.1 Natural and anthropogenic sources of a selection of trace gases.
Table 5.2 Calculated fractional contribution of various photolysis rates to r...
Table 5.3 Global turnover of tropospheric gases and fraction removed by react...
Table 5.4 Range of OH and HO
2
yields from the reaction of ozone with alkenes....
Table 5.5 Percentage of non‐methane hydrocarbon (NMHC) classes measured in th...
Table 5.6 Source strength of various types of aerosol particles.
Chapter 6
Table 6.1 Some typical size and water contents of liquid water in the atmosph...
Table 6.2 Henry's law constants (
K
H
[T
o
] at 298 K) and temperature dependencie...
Table 6.3 Henry's law constants and hydrolysis constants for gases that under...
Chapter 7
Table 7.1 Typical mass composition (μg m
−3
) of various chemical species...
Table 7.2 Representative values of observed aerosol optical properties in the...
Table 7.3 Global emission source strengths for atmospheric aerosols.
Table 7.4 Summary of sea‐salt aerosol size distribution above 100 nm and cons...
Chapter 8
Table 8.1 Example source gases which are decomposed in the stratosphere to re...
Table 8.2 Comparison of gravity and Rossby waves in the stratosphere.
Table 8.3 Some heterogeneous reactions of importance in the lower stratospher...
Table 8.4 Examples of
very‐short‐lived substances
(VSLS) containin...
Chapter 10
Table 10.1 Pollutant concentrations from various urban areas in 2017 (Unit: μ...
Table 10.2 EU air quality standards.
Chapter 11
Table 11.1 Projections of global mean temperature increases and global mean s...
Table 11.2 Lifetime estimates for various
greenhouse gases
(
GHGs
).
List of Illustrations
Chapter 1
Figure 1.1 The climate system (Houghton 2005).
Figure 1.2 Schematic representation of the meridional circulation and associ...
Figure 1.3 Global distribution of the mean height (1963–1973) of the 200 hPa...
Figure 1.4 Mean sea‐level pressure (hPa) (1963–1973) averaged for (a) Decemb...
Figure 1.5 The figure shows measurements deduced from ice‐cores drilled from...
Figure 1.6 Simulated (lines) and reconstructed (grey shading) Northern Hemis...
Figure 1.7 Global average surface temperature anomaly 1850–2017, relative to...
Figure 1.8 Natural factors cannot explain recent warming. Shaded area shows ...
Figure 1.9 The Hadley Centre climate model is now driven by changes in anthr...
Figure 1.10 Mean surface air temperature for (a) January and (b) July.
Figure 1.11 Global distribution of precipitation in decimetres per year.
Figure 1.12 Schematic representation of the Walker circulation along the Equ...
Chapter 2
Figure 2.1 Abundances of gases in the atmospheres of Venus, Earth, and Mars....
Figure 2.2 Gases released to the atmosphere for (a) homogeneous and (b) hete...
Figure 2.3 Hydrogen in the atmosphere of the prebiotic Earth. Release and es...
Figure 2.4 Life's influence on Earth's atmosphere. The diagram shows, on a l...
Figure 2.5 The effects on the early atmosphere of methanogenesis and anoxyge...
Figure 2.6 Inventory of oxidizing and reducing species in the Earth's crust,...
Figure 2.7 Evolution of atmospheric carbon dioxide concentrations, given rel...
Figure 2.8 Anomalous fractionation of sulphur isotopes shown as a function o...
Figure 2.9 Isotopic composition of volcanic O and S in Antarctic ice‐core sa...
Figure 2.10 The ‘three‐stage–three‐box’ model (Kasting 1993) used to represe...
Figure 2.11 The proposed timing for the three stages of Figure 2.10 (Kasting...
Figure 2.12 Percentage of oxygen in the atmosphere over the past 550 Myr inf...
Figure 2.13 Vertical distribution of ozone for different total atmospheric o...
Figure 2.14 Evolution of ozone (dashed line) on Earth in response to the cha...
Chapter 3
Figure 3.1 The average vertical temperature and molecular density structure ...
Figure 3.2 The spectral irradiance from the Sun compared with that of a blac...
Figure 3.3 (a) The blackbody curves for 5777 K and 280 K, representing the s...
Figure 3.4 The average vertical profile of ozone and temperature through the...
Figure 3.5 The extent that
ultraviolet
(
UV
) solar radiation penetrates throu...
Figure 3.6 The absorption cross‐section of molecular oxygen as a function of...
Figure 3.7 The absorption cross‐section of ozone at 273 K as a function of w...
Figure 3.8 Schematic of the variation of upwelling and downwelling radiation...
Figure 3.9 The temperature structure of the grey atmosphere. The temperature...
Figure 3.10 Schematic representation of an aerosol layer overlying a surface...
Figure 3.11 Vertical profiles of carbon monoxide, aerosol particle number co...
Figure 3.12 The lapse rates of the two environment curves A and B are respec...
Figure 3.13 The effect of moisture on the atmospheric lapse rate. A parcel o...
Figure 3.14 The average energy budget of the Earth–atmosphere system. The co...
Figure 3.15 (a) The calculated short and longwave energy budgets of the Eart...
Figure 3.16 Aerosols can have an impact on the radiation budget through thei...
Chapter 4
Figure 4.1 (a) Source and sink terms are equal; the concentration remains co...
Figure 4.2 Schematic of the rise in atmospheric CO
2
over the past 50 years r...
Figure 4.3 Distribution of isoprene emissions derived from a land‐surface mo...
Figure 4.4 A selection of the most common monoterpenes found in the atmosphe...
Figure 4.5 Hydrogen (H
2
) measured at the surface in the Northern Hemisphere ...
Chapter 5
Figure 5.1 Vertical structure of the atmosphere. The vertical profile of tem...
Figure 5.2 UK emission statistics by United Nations Economic Commission of E...
Figure 5.3 Emission maps (2002) for the UK on a 1 km × 1 km grid for (a) NO
2
Figure 5.4 A typical atmospheric ozone profile through the atmosphere (WMO 2...
Figure 5.5 A schematic representation of the atmosphere's role in the Earth ...
Figure 5.6 Solar flux outside the atmosphere and at sea level, respectively....
Figure 5.7 A simplified scheme of tropospheric chemistry. The figure illustr...
Figure 5.8 The sources, interconversions, and sinks for HO
x
(and RO
x
) in the...
Figure 5.9 Simplified mechanism for the photochemical oxidation of CH
4
in th...
Figure 5.10 Schematic representation of the dependence of the net ozone (
N
(O
Figure 5.11 Isopleths giving net rate of ozone production (ppb hour
−1
Figure 5.12 A simplified reaction scheme for night‐time chemistry involving ...
Figure 5.13 Concentrations of N
2
O
5
and NO
3
from airborne measurements off th...
Figure 5.14 Formation of a primary ozonide followed by decomposition to a ca...
Figure 5.15 Main production and loss pathways for the hydroxyl radical in an...
Figure 5.16 A schematic representation of the so‐called Bromine explosion me...
Figure 5.17 Primary sources of reactive halogen species or their precursor r...
Figure 5.18 [IO], tidal height and solar radiation measured at Mace Head in ...
Figure 5.19 Monthly averages of the change of ozone to the mean (ΔO
3
) in the...
Figure 5.20 Three year average satellite instrument measured SO
2
over Easter...
Figure 5.21 Summary of emission, oxidation, and deposition of S(IV) and S(VI...
Figure 5.22 Schematic representation of dominant acidification proceses; VOC...
Figure 5.23 Size and relative number distribution of typical urban aerosols....
Figure 5.24 Air Quality and Climate Change Interactions.
Chapter 6
Figure 6.1 The equilibrium relative humidity and supersaturation as a functi...
Figure 6.2 The conceptual model of Ogren, as developed by Prupacher and Klet...
Figure 6.3 The structural formulae of α‐pinene (with its double bonds), the ...
Figure 6.4 The pH of a set of melted snow samples is collected from Norway, ...
Chapter 7
Figure 7.1 Schematic of the aerosol formation, processes and removal within ...
Figure 7.2 SEM/ESEM images of different aerosol types from the University of...
Figure 7.3 Summary of the dominant effects on particles of different sizes w...
Figure 7.4 Transmission of particles as a function of size.
Figure 7.5 Growth curves for two different salts (a) and two different sizes...
Figure 7.6 (a) Number, (b) surface, and (c) volume distributions for various...
Figure 7.7 Effect of a mixed aerosol layer on an absorbing (low
ω
0
) and...
Figure 7.8 The integrated aerosol backscatter (a measure of the total strato...
Figure 7.9 Schematic of proposed heterogeneous reactions of marine halogens ...
Figure 7.10 Figure SPM.5 from IPCC (2013). Radiative forcing estimates in 20...
Figure 7.11 Schematic of aerosol–cloud interactions for (a) marine air and (...
Figure 7.12 Schematic of a TSI Inc. Model 3790 condensation particle counter...
Figure 7.13 Schematic of an inertial cascade impactor.
Chapter 8
Figure 8.1 The mean column ozone as a function of latitude and season based ...
Figure 8.2 October mean column ozone over Halley Bay (76° S) and average ove...
Figure 8.3 The residual circulation in the stratosphere and associated trans...
Figure 8.4 A latitudinal cross‐section of CF
2
Cl
2
from the satellite instrume...
Figure 8.5 A high‐resolution simulation of the lower stratosphere showing mi...
Figure 8.6 A schematic horizontal–distance/pressure cross‐section of stratos...
Figure 8.7 A model of the chemical evolution in the winter‐time polar lower ...
Figure 8.8 Past and future column ozone. Multi‐model mean
total column ozone
Chapter 9
Figure 9.1 The surface energy balance measured at the University of Reading ...
Figure 9.2 Diurnal cycle of the boundary layer.
Figure 9.3 Schematic profiles of potential temperature, wind, and specific h...
Figure 9.4 Turbulence data measured at the Chilbolton Observatory (UK) on 21...
Figure 9.5 Carbon dioxide in the Amazon rainforest. (a) Profile of
leaf area
...
Figure 9.6 Wind speed profile above and within a canopy. Note the logarithmi...
Figure 9.7 Roughness parameters (a)
z
0
/
h
and (b)
d
/
h
as a function of plant ...
Figure 9.8 Sideview of street canyon flow for isolated roughness (top), wake...
Figure 9.9 Roughness and inertial sublayers. (a) Flow over buildings – verti...
Figure 9.10 Roughness change from rural to urban surface.
Internal boundary
...
Figure 9.11 Large eddy simulation of a thermally induced circulation over a ...
Figure 9.12 Daily temperature variation in urban and rural areas.
Figure 9.13 Dispersion processes near a buoyant release.
Figure 9.14 Schematic showing Gaussian plume from an elevated source.
Figure 9.15 Variation of ground‐level concentration with distance for (a) gr...
Chapter 10
Figure 10.1 Daily concentration of smoke and sulphur dioxide and daily morta...
Figure 10.2 The scale of air pollution in the urban atmosphere. In addition ...
Figure 10.3 Typical urban aerosol number, surface, and volume (proportional ...
Figure 10.4 Flow pattern of air pollutants in a street canyon.
Figure 10.5 Variations of NO, NO
2
, and O
3
at different times of day in a wee...
Figure 10.6 Mortality‐PM
10
dose–response curves for total (TOTAL) mortality ...
Figure 10.7 Deposition of particles in the human respiratory system. Note th...
Figure 10.8 Direct and indirect costs of air pollution.
Figure 10.9 Comparison of NO
x
emission from different euro class vehicles.....
Figure 10.10 Hypothetical air pollution Kuznets curve. Delhi, Beijing, Londo...
Figure 10.11 Annual average smoke and sulphur dioxide concentrations in Lond...
Figure 10.12 Annual concentrations of PM
10
at urban background and roadside ...
Figure 10.13 Annual concentrations of O
3
in the UK, 1987–2017. The O
3
index ...
Figure 10.14 Annual mean concentrations of NO
2
at roadside, urban background...
Figure 10.15 Weighted annual mean concentration of PM
2.5
in selected cities ...
Figure 10.16 Annual arithmetic mean concentration of hourly concentrations o...
Figure 10.17 Annual average PM
10
and SO
2
mass concentration in Beijing.
Chapter 11
Figure 11.1 Climate and atmospheric composition records from the Vostok ice‐...
Figure 11.2 Global CO
2
emissions from fossil‐fuel burning, cement manufactur...
Figure 11.3 Observed monthly average CO
2
concentration (ppmv) from Mauna Loa...
Figure 11.4 Observed global annual average (a) CH
4
, (b) N
2
O, and (c) CFC‐11 ...
Figure 11.5 Globally averaged (a) combined land and ocean surface temperatur...
Figure 11.6 Main components of a full‐scale
integrated assessment model
(
IAM
Figure 11.7 Anthropogenic emissions in the four scenarios of the
Representat
...
Figure 11.8
Integrated science assessment model
(
ISAM
) estimated atmospheric...
Figure 11.9
Integrated science assessment model
(
ISAM
) estimated radiative f...
Figure 11.10
Integrated science assessment model
(
ISAM
) estimated global mea...
Figure 11.11
Integrated science assessment model
(
ISAM
) estimated sea‐level ...
Guide
Cover
Table of Contents
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Atmospheric Science for Environmental Scientists
Edited by
C.N. Hewitt
Lancaster UniversityLancaster, UK
and
Andrea V. Jackson
University of LeedsLeeds, UK
Second Edition
This edition first published 2020© 2020 John Wiley & Sons Ltd
Edition HistoryBlackwell Publishing Ltd (1e, 2009)
All rights reserved. No part of this publication may be reproduced, stored in a retrieval system, or transmitted, in any form or by any means, electronic, mechanical, photocopying, recording or otherwise, except as permitted by law. Advice on how to obtain permission to reuse material from this title is available at http://www.wiley.com/go/permissions.
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List of Contributors
Janet Barlow
Department of MeteorologyUniversity of ReadingReading, UK
Peter Brimblecombe
School of Energy and EnvironmentCity University of Hong KongHong Kong
Martyn P. Chipperfield
School of Earth and EnvironmentUniversity of LeedsLeeds, UK
Hugh Coe
School of Earth, Atmospheric, and Environmental SciencesThe University of ManchesterManchester, UK
Nick Hewitt
Lancaster Environment CentreLancaster UniversityLancaster, UK
Atul Jain
Department of Atmospheric SciencesUniversity of IllinoisUrbana, IL, USA
Anwar Khan
School of ChemistryUniversity of BristolBristol, UK
John Lockwood
Formerly University of LeedsLeeds, UK
A. Rob MacKenzie
School of GeographyEarth and Environmental SciencesUniversity of BirminghamBirmingham, UK
Paul Monks
Department of ChemistryUniversity of LeicesterLeicester, UK
Dudley Shallcross
School of ChemistryUniversity of BristolBristol, UK
Zongbo Shi
School of Geography, Earth and Environmental ScienceThe University of BirminghamBirmingham, UK
Natalie Theeuwes
Department of MeteorologyUniversity of ReadingReading, UK
Joshua Vande Hey
Department of ChemistryUniversity of LeicesterLeicester, UK
Richard Wayne
Physical and Theoretical Chemistry LaboratoryDepartment of ChemistryUniversity of OxfordOxford, UK
Paul I. Williams
School of Earth and Environmental Sciences & National Centre for Atmospheric ScienceThe University of ManchesterManchester, UK
Xiaoming Xu
Department of Atmospheric SciencesUniversity of IllinoisUrbana, IL, USA
Preface
When we wrote the Preface to the first edition of ‘Atmospheric Science for Environmental Scientists’ in 2008, we noted that never before had the teaching, learning, and researching of atmospheric science been so important. We said that society must face up to the realities of global atmospheric change, including global warming and poor air quality, and that the education of students and provision of accessible information to policy makers and the public were priorities.
More than a decade later, we can only reiterate these sentiments. In 2018, the Intergovernmental Panel on Climate Change warned that the planet will reach the crucial threshold of 1.5 °C above pre‐industrial levels by as early as 2030, precipitating the risks of extreme drought, wildfires, floods, and food shortages for hundreds of millions of people. And in 2018, the World Health Organization reported that 90% of the world's population lived in places where air quality exceeded WHO guideline limits, and that more than 4 million people a year died prematurely from outdoor air pollution and a further 3 million a year from indoor air pollution.
What further warnings are needed? To help society cope with the unprecedented changes that humankind is causing to our fragile atmosphere, education must be key and policy makers must act. We hope this book helps both causes.
In putting this book together, we have drawn on some of the best experts and educators in the field of atmospheric science. We hope their knowledge and enthusiasm shines through in these chapters. Our aim is to provide succinct but detailed information on all the important aspects of atmospheric science for students of environmental science and to others who are interested in learning how the atmosphere works, how humankind is changing its composition, and what effects these changes might lead to.
We are grateful to all the experts who have contributed to this book, for all reviewers' comments, and to all our students over the years who have demonstrated the need for this volume.
October 2019
Nick HewittAndrea V. Jackson
Abbreviations, Constants, and Nomenclature
ADMS
Atmospheric Dispersion Modelling System
CEE
Central and eastern Europe
CCN
cloud condensation nuclei
CFC
chlorofluorocarbons
CO
2
carbon dioxide
DMS
dimethyl sulphide
DNA
deoxyribonucleic acid
EC
elemental carbon
EM
electromagnetic
ENSO
El Niño–Southern Oscillation
EPA
Environmental Protection Agency
EU
European Union
GDP
global domestic product
GEMS/AIR
Global Environmental Monitoring System/Air
GHG
greenhouse gas
HAP
hazardous air pollutant
HCFC
hydrochlorofluorocarbons
HFC
hydrofluorocarbons
IAM
integrated assessment models
IBL
internal boundary layer
IPCC
Intergovernmental Panel on Climate Change
IR
infrared
ISAM
integrated science assessment model
ITCZ
intertropical convergence zone
LAI
leaf‐area index
LW
longwave
NMHC
non‐methane hydrocarbons
MAP
major air pollutant
MTBE
methyl‐
tert
‐butyl ether
NDVI
normalized difference vegetation index
OCS
carbonyl sulphide
OECD
Organization for Economic Cooperation and Development
PAH
polycyclic aromatic hydrocarbons
PAN
peroxyacetyl nitrate
PAR
photosynthetically active radiation
PCB
polychlorinated biphenyls
PFC
perfluorogenated carbon
PM
particulate matter
PM
10
particles with aerodynamic diameter less than 10 μm
ppm
parts per million
ppmv
part per million by volume (1 × 10
−6
)
ppbv
part per billion by volume (1 × 10
−9
)
pptv
part per trillion by volume (1 × 10
−12
)
PSS
photostationary state
SAFARI
South African Regional Science Initiative
SW
shortwave
TSP
total suspended particulates
UNEP
United Nations Environmental Programme
UV
ultraviolet
VSLS
very short‐lived substances
VOC
volatile organic compounds
WHO
World Health Organization
WMO
World Meteorological Organization
Constants
c
speed of light in vacuum 2.998 × 10
8
m s
−1
g
acceleration due to gravity 9.8 m s
−2
h
Planck’s constant 6.626 × 10
34
J s
k
Boltzmann constant 1.381 × 10
34
J K
−1
R
gas constant 8.314 J K
−1
mol
−1
(1.3 × 10
5
l atm mol
−1
K
−1
)
Γ
d
dry adiabatic lapse rate 9.81 K km
−1
π
3.14159
σ
Stefan–Boltzmann constant 5.67 × 10
−8
W m
−2
K
−4
Nomenclature
a
radius of a particle
A
albedo
A
s
surface albedo
B
radiative intensity of a blackbody
c
p
specific heat capacity of dry air at constant pressure (1004 J kg
−1
K
−1
)
c
v
specific heat at constant volume
C
concentration (ppm or kg m
−3
)
d
preface to variable indicating incremental quantity
d
displacement height (m)
d
Q
incremental change in heat
d
u
incremental change in internal energy
d
v
incremental change in volume
d
w
incremental change in work
e
turbulent kinetic energy per unit mass (J kg
−1
)
F
B
the total flux from a black body radiator
F
s
incoming solar radiation absorbed at the surface
F
net flux
F
↑
upwelling radiative flux
F
↓
downwelling radiative flux
net flux leaving an element or layer
F
r
total upward reflected shortwave flux
F
0
incident solar flux
G
ground heat flux (W m
−2
)
h
mean height of roughness elements (m)
h
b
depth of internal boundary layer (m)
H
sensible heat flux (W m
−2
) or mean building height (m)
H
s
scale height
H
/
W
aspect ratio (‐)
I
intensity of light
I
0
initial intensity of light
k
von Kármán’s constant ≈ 0.4 (‐)
k
a
absorption coefficient
k
e
extinction coefficient
k
s
scattering coefficient
l
distance through a gas interacting with light
L
latent heat of vaporization
L
Obukhov length (m)
L
w
Liquid water content
L
x
integral (or decorrelation) lengthscale (m)
m
refractive index
M
molar mass (of air unless otherwise specified)
n
number concentration, typically of absorbers or scatterers
p
(u)
probability of windspeed u (‐)
P
plant area index (‐)
P
,
p
pressure
q
specific humidity (kg kg
−1
, sometimes g kg
−1
) or source emission rate (kg s
−1
)
q
*
scaling parameter for specific humidity profile (kg kg
−1
, sometimes g kg
−1
)
r
radial distance
r
v
mass mixing ratio of water vapour
r
w
saturation mixing ratio
r
r
scattered fraction reflected upwards
R
radius of the Earth
R
(
λ
)
action spectrum
R
n
net radiation (W m
−2
)
S
Solar constant
Sk
u
skewness statistic for downstream component of the wind (‐)
Sk
w
skewness statistic for vertical component of the wind (‐)
t
time (s)
t
r
scattered fraction transmitted downwards
t
t
total fraction of radiation transmitted downwards
T
temperature
T
a
transmittance of the atmosphere
T
e
effective blackbody temperature of the Earth
T
L
integral timescale (s)
T
*
scaling parameter for temperature profile (°C)
u
downstream velocity component (m s
−1
)
mean component (m s
−1
)
u
′
fluctuation around mean component (m s
−1
)
u
*
friction velocity (m s
−1
)
U
velocity vector (m s
−1
)
(z)
vertical mean wind profile (m s
−1
)
mean windspeed at the top of a canopy (m s
−1
)
v
lateral velocity component (m s
−1
)
w
vertical velocity component (m s
−1
)
W
along wind building spacing (m)
x
horizontal distance in downstream direction (m)
y
horizontal distance in lateral direction (m)
z
distance, usually altitude (m)
z
h
roughness length for heat transfer (m)
z
m
height of maximum plant area index (m)
z
0
roughness length for momentum transfer (m)
z
0r
roughness length for rural surface (m)
z
0u
roughness length for urban surface (m)
z
q
roughness length for moisture transfer (m)
z
*
roughness sublayer depth (m)
z
/
L
Monin‐Obukhov stability parameter (‐)
σ
standard deviation (depends on quantity)
σ
x
downstream plume spread (m)
σ
y
lateral plume spread (m)
σ
z
vertical plume spread (m)
σ
(
λ
)
absorption cross section as a function of wavelength
β
upscatter or backscatter fraction
γ
s
saturated adiabatic lapse rate
∆S
storage term (W m
−2
)
∆T
temperature difference between rural and urban areas (or heat island intensity) (°C)
∂
preface to variable indicating incremental quantity
θ
scattering angle, or potential temperature (°C or K)
θ
0
potential temperature at the surface (°C or K)
ε
emissivity
λ
wavelength of light
λE
latent heat flux (W m
−2
)
λ
F
frontal area index (‐)
ρ
air density (1.2 kg m
−3
)
ρ
v
density of water vapour
ρ
a
density of dry air
ρ
NIR
ratio of emitted to incident near infra‐red radiation (‐)
ρ
VIS
ratio of emitted to incident visible radiation (‐)
χ
optical depth
τ
momentum flux, or shear stress (kg m
−1
s
−2
)
υ
e
extinction coefficient (‐)
ω
0
single scattering albedo
ψ
m
stability function for momentum (‐)
1The Climate of the Earth
John Lockwood
Formerly University of Leeds, Leeds, United Kingdom
The causes, history, and distributions of the Earth's climates are introduced in this chapter. The combination of the distribution of incoming solar radiation across the Earth's surface and the Earth's rotation both drive and shape the observed atmosphere–ocean circulation. Important factors determining changes in climate include palaeogeography, greenhouse gas concentrations, changing orbital parameters, and varying ocean heat transport. One of the major controls of climatic changes is the greenhouse gas concentration of the atmosphere, in particular that of carbon dioxide. Before the Eocene–Oligocene boundary (≈34 Myr ago) the atmosphere–ocean circulation supported a warm atmosphere and ocean, with both poles free of permanent ice. At the Eocene–Oligocene boundary, the atmosphere–ocean circulation changed to a form similar to the present, and the first evidence of an Antarctic ice sheet is found. Falling atmospheric carbon dioxide levels probably caused this change. The waxing and waning of massive temperate latitude continental ice sheets characterize the climate of the past million years. This chapter discusses recent climate changes and evidence that they are largely driven by anthropogenic generated atmospheric carbon dioxide. In particular, recent climate changes are causing the expansion of the tropical zone and a retreat of the polar zones.
The major climate zones of the world are described, with particular attention to interannual variability, and the causes of droughts and heavy rainfalls. This includes discussions of the climatic effects of the North Atlantic Oscillation (NAO) and El Niño‐Southern Oscillation (ENSO).
For more specific information on global warming and climate change science, the reader is referred to Chapter 11 in this book and to the latest reports of the Intergovernmental Panel on Climate Change, available at www.ipcc.ch.
1.1 Basic Climatology
The climate of a particular place is the average state of the atmosphere observed as the weather over a finite period (e.g. a season) for a number of different years. The so‐called climate system, which determines the weather, is a composite system consisting of five major interactive adjoint components: the atmosphere, the hydrosphere, including the oceans, the cryosphere, the lithosphere, and the biosphere (Figure 1.1). All the subsystems are open and not isolated, as the atmosphere, hydrosphere, cryosphere, and biosphere act as cascading systems linked by complex feedback processes. The climate system is subject to two primary external forcings that condition its behaviour: solar radiation and the Earth's rotation. Solar radiation must be regarded as the primary forcing mechanism, as it provides almost all the energy that drives the climate system.
Figure 1.1 The climate system (Houghton 2005).
The distribution of climates across the Earth's surface is determined by its spherical shape, its rotation, the tilt of the Earth's axis of rotation in relation to a perpendicular line through the plane of the Earth's orbit around the Sun, the eccentricity of the Earth's orbit, the greenhouse gas content of the atmosphere, and the nature of the underlying surface. The spherical shape creates sharp north–south temperature differences, whilst the tilt is responsible for month‐by‐month changes in the amount of solar radiation reaching each part of the planet, and hence the variations in the length of daylight throughout the year at different latitudes and the resulting seasonal weather cycle.
The present orbit of the Earth is slightly elliptical with the Sun at one focus of the ellipse, and as a consequence the strength of the solar beam reaching the Earth varies about its mean value. At present, the Earth is nearest to the Sun in January and farthest from the Sun in July. This makes the solar beam near the Earth about 3.5% stronger than the average mean value in January, and 3.5% weaker than average in July. The gravity of the Sun, the Moon, and the other planets causes the Earth to vary its orbit around the Sun (over many thousands of years). Three different cycles are present, and when combined, produce the rather complex changes observed. These cycles affect only the seasonal and geographical distribution of solar radiation on the Earth's surface, yearly global totals remaining constant. Surplus in one season is compensated by a deficit during the opposite one; surplus in one geographical area is compensated by simultaneous deficit in some other zone. Nevertheless, these Earth orbital variations can have a significant effect on climate and are responsible for some major long‐term variations.
Firstly, there are variations in the orbital eccentricity. The Earth's orbit varies from almost a complete circle to a marked ellipse, when it will be nearer to the Sun at one particular season. A complete cycle from near circular through a marked ellipse back to near circular takes between 90 000 and 100 000 years. When the orbit is at its most elliptical, the intensity of the solar beam reaching the Earth must undergo a seasonal range of about 30%. Second, there is a wobble in the Earth's axis of rotation causing a phenomenon known as the precession of the equinoxes. That is to say, within the elliptical orbits just described, the distance between Earth and Sun varies so that the season of the closest approach to the Sun also varies. The complete cycle takes about 21 000 years. Lastly, the tilt of the Earth's axis of rotation relative to the plane of its orbit varies at least between 21.8° and 24.4° over a regular period of about 40 000 years. At present, it is almost 23.44° and is decreasing. The greater the tilt of the Earth's axis, the more pronounced the difference between winter and summer. Technically, these three mechanisms are known as the Milankovitch mechanism.
If the Earth did not rotate relative to the Sun – that is, it always kept the same side towards the Sun – the most likely atmospheric circulation would consist of rising air over an extremely hot, daylight face and sinking air over an extremely cold, night face. The diurnal cycle of heating and cooling obviously would not exist, since it depends on the Earth's rotation. Surface winds everywhere would blow from the cold night face towards the hot daylight face, whilst upper flow patterns would be the reverse of those at lower levels. Whatever the exact nature of the atmospheric flow patterns, the climatic zones on a nonrotating Earth would be totally different from anything observed today. Theoretical studies suggest that if this stationary Earth started to rotate, then as the rate of rotation increased, the atmospheric circulation patterns would be progressively modified until they resembled those observed today. In very general terms, these circulation patterns take the form of a number of meridional overturning cells in the atmosphere, with separate zones of rising air motion at low and middle latitudes, and corresponding sinking motions in subtropical and polar latitudes.
1.2 General Atmospheric Circulation
A schematic representation of the mean meridional circulation between Equator and pole is shown in Figure 1.2. A simple direct circulation cell, known as the Hadley cell, is clearly seen equatorward from 30° latitude (Lockwood 2003). Eastward angular momentum is transported from the equatorial latitudes to the middle latitudes by nearly horizontal eddies, 1000 km or more across, moving in the upper troposphere and lower stratosphere. This transport, together with the dynamics of the middle latitude atmosphere, leads to an accumulation of eastward momentum between 30° and 40° latitude, where a strong meandering current of air, generally known as the subtropical westerly jet stream, develops (Figure 1.3). The cores of the subtropical westerly jet streams in both hemispheres and throughout the year occur at an altitude of about 12 km. The air subsiding from the jet streams forms the belts of subtropical anticyclones at about 30° to 40° N and S (Figure 1.4). The widespread subsidence in the descending limb of the Hadley cell should be contrasted with the rising limb, where ascent is restricted to a few local areas of intense convection. More momentum than is necessary to maintain the subtropical jet streams against dissipation through internal friction is transported to these zones of upper strong winds. The excess is transported downwards and polewards to maintain the eastward‐flowing surface winds (temperate latitude westerlies) against ground friction. The middle latitude westerly winds are part of an indirect circulation cell known as the Ferrel cell. The supply of eastward momentum to the Earth's surface in middle latitudes tends to speed up the Earth's rotation. Counteracting such potential speeding up of the Earth's rotation, air flows from the subtropical anticyclones towards the equatorial regions, forming the so‐called trade winds. The trade winds, with a strong flow component directed towards the west (easterly winds), tend to retard the Earth's rotation, and in turn gain eastward momentum.
Figure 1.2 Schematic representation of the meridional circulation and associated jet‐stream cores in winter. The tropical Hadley cell and middle latitude Ferrel cell are clearly visible.
Source: from Palmen 1951.
Figure 1.3 Global distribution of the mean height (1963–1973) of the 200 hPa pressure field represented as mean height minus 11 784 gpm for (a) December, January, February; (b) June, July, August. Wind speed and direction shown by arrows. Each barb on the tail of an arrow represents a wind speed of 5 m s−1
Source: from Peixoto and Oort (1992).
Figure 1.4 Mean sea‐level pressure (hPa) (1963–1973) averaged for (a) December, January, February; and (b) June, July, August.
Source: from Henderson‐Sellers and Robinson (1986) and Oort (1983).
The greatest atmospheric variability occurs in middle latitudes, from approximately 40° to 70° N and S, where large areas of the Earth's surface are affected by a succession of eastward‐moving cyclones (frontal depressions) and anticyclones or ridges. This is a region of strong north–south thermal gradients with vigorous westerlies in the upper air at about 10 km, culminating in the polar‐front jet streams along the polar edges of the Ferrel cells (Figure 1.2). The zone of westerlies is permanently unstable and gives rise to a continuous stream of large‐scale eddies near the surface, the cyclonic eddies moving eastward and poleward and the anticyclonic ones moving eastward and equatorward. In contrast, at about 10 km, in the upper westerlies, smooth wave‐shaped patterns are the general rule. Normally, there are four or five major waves around the hemisphere, and superimposed on these are smaller waves that travel through the slowly moving train of larger waves. The major waves are often called Rossby waves, after Rossby who first investigated their principal properties. Compared with the Hadley cells, the middle latitude atmosphere is highly disturbed and the suggested meridional circulation shown in Figure 1.2 is largely schematic.
The extension into very high latitudes and the northward narrowing of the northern North Atlantic have consequences on the Atlantic Ocean circulation, which in turn has a series of unique effects on the climate system. This is in complete contrast to the much more benign North Pacific Ocean. Warm, saline surface water flows into the northern North Atlantic, after travelling from the Caribbean Sea, via the Gulf Stream and the North Atlantic Drift. This inflowing water, which is more saline than anywhere else in the high‐latitude oceans, is finally advected to sites in the Greenland and Norwegian Seas, where extreme cooling to the atmosphere occurs and surface water sinks to the ocean depths. When cooled, water with the salinity normal in the world's oceans becomes denser, but unlike fresh water does not reach its maximum density until near its freezing point, at about −2 °C. Thus, the saltwater of the deep oceans, when cooled at the surface, goes into convective patterns, the coldest and densest portions gradually sinking from the surface to the ocean depths. Low‐density surface layers in the oceans can arise either because of surface heating, or the addition of relatively fresh continental runoff or precipitation onto the ocean surface. In the cold oceans, sea‐ice will form only when a layer of the ocean close to the surface has a relatively low salinity. The existence of this layer allows the temperature of the surface water to fall to freezing point, and ice to form, despite the lower levels of the ocean having a higher temperature.
1.3 Palaeoclimates
The major controls of very long‐term climatic change include palaeogeography, greenhouse gas concentrations, changing orbital parameters, and varying ocean heat transport. One of the major controls on long‐term climatic changes is the greenhouse gas concentration of the atmosphere and in particular that of carbon dioxide. Atmospheric carbon dioxide concentration has varied markedly during the Earth's history. Atmospheric CO2 concentrations are controlled by the carbon cycle, and the net effect of slight imbalances in the carbon cycle over tens to hundreds of millions of years has been to reduce atmospheric CO2. Atmospheric CO2 concentrations remained relatively high up to about 60 Myr ago when there was a very marked fall. Atmospheric concentrations continued to fall after about 60 Myr ago, and there is geochemical evidence that concentrations were less than 300 ppm by about 20 Myr ago.
Available evidence is that during the Mesozoic Era, temperatures ranged from 10° to 20 °C at the poles to 25–30 °C at the Equator – that is, the poles were free of permanent ice fields, and the atmosphere–ocean circulation was different in some important aspects from that observed today. Slight cooling took place at the start of the Jurassic Period and marked high‐latitude warming during the first half of the Cretaceous Period. Global cooling again took place towards the end of Cretaceous time, and a long‐term cooling trend commenced at the start of the Eocene Epoch, some 55 Myr ago.
The sudden, widespread glaciations of Antarctica and the associated shift towards colder temperatures at the Eocene–Oligocene boundary (approximately 34 Myr ago) is one of the most fundamental reorganizations of global climate and ocean circulation known in the geological record. Prior to the Eocene–Oligocene boundary, there is little evidence of the deep cold water in the world ocean that is so common today. Indeed, before the boundary, atmospheric, and particularly oceanic circulation conditions were probably different from those observed today. After the boundary they are probably rather similar to present‐day conditions. Oceanic bottom water is formed in small regions by convective buoyancy plumes that transfer relatively dense ocean water from near the surface to the ocean depths. Deep‐ocean temperatures are therefore closely related to ocean surface temperatures in key regions. The surface density and salinity is usually increased by evaporation and heat transfer to the atmosphere; therefore, virtually all deep‐water formation seems to be over continental shelves in low latitudes or at high latitudes. During the Cretaceous Period and up to the end of the Eocene Epoch, the ocean bottom‐waters were warm, saline, and formed in shallow subtropical marginal seas. At the Eocene–Oligocene boundary, ocean bottom‐water temperatures decreased rapidly to approximately present‐day levels. Deep‐sea cores suggest that this change occurred within 100 000 years, which is remarkably abrupt for preglacial Tertiary times, and is considered to represent the time when large‐scale freezing conditions developed at sea‐level around Antarctica, forming the first significant sea‐ice. At this time, cold‐water plumes forming off Antarctica started to dominate ocean bottom‐water formation, and together with Arctic Ocean plumes they have dominated until the present day. Thus, from early Oligocene times onwards it may be considered that world climates were in the present cold or semi‐glacial state.
The initial growth of the East Antarctic Ice Sheet near the Eocene–Oligocene boundary is often attributed to the opening by continental drift of ocean gateways between Antarctica and Australia (Tasmanian Passage) and Antarctica and South America (Drake Passage), leading to the organization of the Antarctic Circumpolar current and the ‘thermal isolation’ of Antarctica. This notion has been challenged because although most tectonic reconstructions place the opening of the Tasmanian Passage close to the Eocene–Oligocene boundary, the Drake Passage may not have provided a significant deep‐water passage until several million years later. Recent model simulations (DeConto and Pollard 2003) of the glacial inception and early growth of the East Antarctic Ice Sheet suggest that declining Cenozoic carbon dioxide first leads to the formation of small, highly dynamic ice caps on high Antarctic plateaux. At a later time, a carbon dioxide threshold is crossed, initiating various feedbacks that cause the ice caps to expand rapidly with large orbital variations, eventually coalescing into a continental‐scale East Antarctic Ice sheet. According to this simulation, the opening of the two Southern Ocean gateways plays a secondary role in this transition, relative to changing carbon dioxide concentration.
1.3.1 Quaternary Glaciations
Continental ice‐sheets probably appeared in the Northern Hemisphere about 3 Myr ago, occupying lands adjacent to the North Atlantic Ocean. The time of the formation of the Greenland ice‐sheet is not well known from terrestrial evidence, but the presence of glacial marine sediments in North Atlantic marine cores first appeared around 3 Myr ago. The oldest glacial moraines in Iceland are dated to approximately 2.6 Myr ago. For at least the past million years, the Earth's climate has been characterized by an alternation of glacial and interglacial episodes, marked in the Northern Hemisphere by the waxing and waning of continental ice‐sheets and in both hemispheres by rising and falling temperatures (Figure 1.5). The present dominant cycle is one of about 100 kyr and is seen in the growth and decay of the continental ice‐sheets. The last major glacial episode started about 110 kyr ago and finished only about 10 kyr ago. These fluctuations or cycles are found in a large number of proxy data records, analysis of which suggests that Antarctic air temperature, atmospheric CO2 content, and deep‐ocean temperatures are dominated by variance with a 100 kyr period and vary in phase with orbital eccentricity. In contrast, global ice volume lags changes in orbital eccentricity (Shackleton 2000). Hence, the 100‐kyr ice‐sheet cycle does not arise from ice‐sheet dynamics; instead, it is probably the response of the global carbon cycle to changes in orbital eccentricity that generates the eccentricity signal in the climate record by causing changes in atmospheric carbon dioxide concentrations.
Proxy data records can be grouped into two climatic regimes with the transitional zone about 430 kyr ago (Figure 1.5). The earlier shows higher‐frequency cycles (dominance of 40‐kyr cycles), with less coherence amongst the various proxy climatic records than the later one (dominance of 100‐kyr cycles). This may be due to a decrease in average atmospheric CO2 levels over the past two million years (Brook 2005). There are differences in the amplitudes of deuterium and CO2 oscillations before and after 430 kyr ago. The atmospheric concentration of CO2 did not exceed 300 ppmv for the 650 000 years before the beginning (around 1750) of the industrial era. Since the Industrial Revolution, atmospheric carbon dioxide concentrations have increased by 33%. Before 430 kyr ago concentrations of CO2 did not exceed 260 ppmv.
Figure 1.5 The figure shows measurements deduced from ice‐cores drilled from the Antarctic ice‐sheet, and analysed by the British Antarctic Survey and others as part of the European programme EPICA. The actual measurement is of the concentration of deuterium in air bubbles, and this can be related to local temperatures. The figure shows that temperature rise between the depth of the last ice age 20 000 years ago and the current interglacial is about 9 °C.
Source: Hadley Centre for Climate Prediction and Research. From EPICA Community Members (2004).
The transition from glacial to interglacial conditions about 430 kyr ago resembles the transition into the present interglacial period at about 10 kyr ago in terms of the magnitude of changes in temperature and greenhouse gases. As commented above, the transition 430 kyr ago also delimits the frontier between two different patterns of climate, and has been identified by recent investigations as a unique and exceptionally long interglacial. Some workers (see Brook 2005) suggest that because the orbital parameters (low eccentricity and consequently weak precessional forcing) are similar to those of the present and next tens of thousands of years, the interglacial 430 kyr ago may be the best analogue available for present and future climate without human intervention. Long interglacials with stable conditions are not, therefore, exceptional, short interglacials such as the past three are not the rule and hence cannot serve as analogues of the present Holocene interglacial.
Sudden and short‐lived climate oscillations giving rise to warm events occurred many times during the generally colder conditions that prevailed during the last glacial period between 110 000 and 10 000 years ago (Lockwood 2001). They are often known as interstadials to distinguish them from the cold phases or stadials. Between 115 000 and 14 000 years ago there were 24 of these oscillations, as recognized in the Greenland ice‐core records where they are called Dansgaard–Oeschger oscillations. These can be viewed as oscillations of the climate system about an extremely ill‐defined mean state. Each oscillation contains a warm interstadial, which is linked to and followed by a cold stadial. Ice‐core and ocean data suggest that the oscillations began and ended suddenly, although in general the ‘jump’ in climate at the start of an oscillation was followed by a more gradual decline that returned conditions to the colder ‘glacial’ state. Warming into each oscillation occurred over a few decades or less, and the overall duration of some of these warm phases may have been just a few decades, whereas others vary in length from a few centuries to nearly 2000 years.
Of totally different nature to Dansgaard–Oeschger oscillations are extreme and short‐lived cold events, known as Heinrich events. These events occurred against the general background of the glacial climate and represent the climatic effects of massive surges of fresh water and icebergs from melting ice sheets into the North Atlantic, causing substantial changes in the thermohaline circulation (Lockwood 2001). Several massive ice‐rafting events show up in the Greenland ice‐cores as a further 3–6 °C drop in temperature from already cold glacial conditions. Many of these events have also been picked up as particularly cold and arid intervals in European and North American pollen records. The most recent Heinrich event is known as the Younger Dryas and appears as a time of glacial re‐advance in Europe after the end of the main ice‐age.
1.3.2 The Recent Climate Record
Extensive instrumental temperature records exist only for the period after about 1860, but recently multiproxy data networks (e.g. Mann et al. 1999; Intergovernmental Panel on Climate Change Fifth Assessment Report, 2014) have been used to reconstruct Northern Hemisphere temperatures back to AD 1000 or earlier (Figure 1.6). These reconstructions and simulations show a long‐term cooling trend in the Northern Hemisphere prior to industrialization of −0.02 °C per century, possibly related to orbital forcing, which is thought to have driven long‐term temperatures downward since the mid‐Holocene at a rate within the range from −0.01 to −0.04 °C per century. The temperature reconstruction also shows that the late eleventh, twelfth and fourteenth centuries rival mean twentieth century temperature levels, whilst cooling following the fourteenth century can be viewed as the initial onset of the cold period known as the Little Ice Age. There is general agreement that the Little Ice Age came to an abrupt end around 1850, whilst studies in Switzerland indicate that overall the coldest conditions of the past 500 years were in the late seventeenth and early nineteenth centuries. The early nineteenth century was especially cold and can be considered as the ‘climatic pessimism’ of the past 1000 years.
Figure 1.6 Simulated (lines) and reconstructed (grey shading) Northern Hemisphere average surface temperature changes since the Middle Ages. All data are expressed as anomalies from their 1500–1850 mean and smoothed with a 30‐year filter.
Source: data from, and more details available at, http://www.ipcc.ch/report/ar5/syr.
Global mean surface temperature has increased dramatically during the last one hundred years or so, but not in a uniform manner (Figure 1.7). The global increase in temperature since about 1880 occurred during two sustained periods, one beginning around 1910 and another beginning in the 1970s and continuing to the present day. Best available estimates (Jones and Moberg 2003) give global temperature trends from 1910 to 1945 of 0.11 °C per decade, −0.01 °C per decade from 1946 to 1975 and 0.22 °C per decade from 1976 to 2000. In the period 2001 to 2010 warming was again ~0.11 °C per decade. Nine of the 10 warmest years observed globally since reliable observations were begun over a century and a half ago have occurred since the year 2000, and all 10 warmest years have occurred since 1998. The 20 warmest years on record have all occurred since 1995, with the five warmest years occurring since 2010. The warmest year of all (prior to 2018) was 2016.
The largest recent warming is in the winter extratropical Northern Hemisphere, with a faster rate of warming over land compared with the ocean. Using satellite‐borne microwave sounding units, Qiang Fu et al. (2006) have examined atmospheric temperature trends for 1979–2005. They found that relative to the global‐mean trends of the respective layers, both hemispheres have experienced enhanced tropospheric warming and stratospheric cooling in the 15–45° latitude belt. This suggests a widening of the tropical circulation zone and a poleward shift of the subtropical jet streams and their associated subtropical dry zones.
Figure 1.7
