Chapter 2
Basics of Global Climate
2.1 Components and phenomena in the climate system
2.2 Basics of radiative forcing
2.3 Globally averaged energy budget
2.4 Gradients of rad. Forcing and energy transport by atm.
2.5 Atmospheric circulation
2.6 Ocean circulation
2.7 Land surface processes
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
2.1 Components and phenomena in the climate system
 "Components" of the climate system:
• atmosphere
• ocean
• land surface
• cryosphere (land ice (including ice shelves and glaciers), snow
and sea ice)
• biosphere
• lithosphere (solid earth)
• chemical composition (biogeochemistry; biology  chemistry
of climate system).
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
 Climate processes
• solar radiation tends to get through the atmosphere;
absorbed at land surface and upper 10 m of ocean.
• ocean heated from above  stable to vertical motions.
• warm surface layer, colder deep waters
• mixing near ocean surface by turbulent motions created,
e.g., by wind  upper mixed layer ~ 50 m depth.
• mixing carries surface warming down as far as
thermocline, layer of rapid transition of temperature to
the colder abyssal waters below
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Figure 2.1
Schematic of components of the climate system
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Table 2.1
Atmosphere
Overall response time to heating
Typical spin-down time of wind if nothing is forcing it
Frontal system lifetime (1000s of km)
Convective cloud lifetime (100m to km horizontal;
Time scale for typical upper level wind (20 m s-1) to
Ocean
Response time of upper ocean (above thermocline) to
heating
Response time of deep ocean to atmospheric changes
Ocean eddy lifetime (10s to 100 km)
Ocean mixing in the surface layer
Time for typical ocean current (cm s-1) to cross ocean
(1000s of km)
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Typical time scales
months
days
days
hours
days
months to years
decades to millennia
month
hours to days
decade
Cryosphere
Snow cover
Sea ice (extent and thickness variations)
Glaciers
Ice caps
Land surface
Response time to heating
Response time of vegetation to oppose excess
evaporation
Soil moisture response time
Biosphere
Ocean plankton response to nutrient changes
Recovery time from deforestation
Lithosphere
Isostatic rebound of continents (after being
depressed by weight of glacier)
Weathering, mountain building
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Typical time scales
months
months to years
decades to centuries
centuries to millennia
hours
hours
days to months
weeks
years to decades
10,000s of years
1,000,000s of years
Figure 2.2
Satellite image based
on visible light
Courtesy of NASA
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Figure 2.2
Satellite image based
on visible light with
5°x5° grid overlay
Courtesy of NASA
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
The parameterization problem
 For each grid box in a climate model, only the average
across the grid box of wind, temperature, etc. is
represented.
 In the observations, many fine variations occur inside,
• e.g., squall lines, cumulonimbus clouds, etc.
 The average of these small scale effects has important
impacts on large-scale climate.
• e.g., clouds primarily occur at small scales, yet the
average amount of sunlight reflected by clouds affects
the average solar heating of a whole grid box.
 Average effects of the small scales on the grid scale must
be included in the climate model.
 These averages change with the parameters of large-scale
fields that affect the clouds, such as moisture and temp.
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Method of representing average effects of clouds (or other
small scale effects) over a grid box interactively with the
other variables known as parameterization.
Successes and difficulties of parameterization important to
accuracy of climate models.
grid used in Figure 2.2 to illustrate a climate model is not
as fine as current models (half in lat. and lon.).
finer grid implies greater computational costs (or shorter
simulation)
 As computers become faster  finer grids.
But there are always smaller scales.
Scale interaction is one of the main effects that makes
climate modeling challenging.
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Rainfall from the
TRMM-based
merged data
(3B42RT)
Weekly accumulation
Rain rate from a 3hourly period within
the week shown above
Related animations
Visible image
Western Pacific
1652Z
Rainfall animation from
the TRMM-based merged
data (3B42RT)
Figure 2.3
A cumulonimbus cloud in the tropics
 area, thickness and
height of different
parts of the cloud
affect how it will
reflect solar
radiation and absorb
or emit infrared
radiation (hard to
parameterize).
 (less hard)
transports heat from
surface through deep
layer in atm.
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Image courtesy of T. A. Toney
Figure 2.4
Satellite image of sea
surface temperature off
the east coast of the
United States
Gulf stream, ocean eddies.
Eddy have spatial scale
~50 km  computationally
costly.
Eddies tend to twist into
very small scale features,
become mixed  transport
heat poleward ("eddy
transports")
Courtesy of Ocean Remote Sensing Group, Johns Hopkins U., Applied Phys. Lab.
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
2.2 Basics of radiative forcing
 Solar radiation comes in, mostly reaching the surface
 Infrared radiation (IR) is the only way this heat input can
be balanced by heat loss to space
 Since IR emissions depend on the Earth's temperature, the
planet tends to adjust to a temperature where IR energy
loss balances solar input:
• i.e., outgoing flux of long wavelength infrared radiation*,
integrated over the Earth, balances flux of incoming short
wavelength solar radiation. *"outgoing longwave radiation" (OLR).
• intensity of radiation as a flux in units of Watts per square
meter (W/m2).
 Blackbody radiation: approximation for how radiation
depends on temperature (that does not depend on the substance
doing the emitting; black = perfect absorber/emitter).
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Figure 2.5
Blackbody radiation
curves & absorption
of radiation at each
wavelength
Fraction of radiation
absorbed at each
wavelength as it
passes through whole
depth of atmosphere
Fraction absorbed
above 11 km
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
After Goody and Yung, 1989, Atmospheric Radiation, Theoretical Basis
by permission of Oxford University Press (www.oup.com).
 Total energy flux integrated across all wavelengths of light
(for a blackbody emitter) depends on absolute temperature
T by the Stefan-Boltzmann law
R = sT4
s = 5.67x10-8 Wm-2K-4, T in Kelvin (celsius + 273.16)
 Actual surfaces or gases do not absorb or emit as a perfect
blackbody  define an emissivity e for each substance
R = esT4
 absorptivity = emissivity.
 later bulk emissivity ea for an atmospheric layer.
 Full climate models do detailed computation as a function
of wavelength, for every level in the atmosphere, …
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Figure 2.6
Schematic of Sun's rays arriving at disk and spherical Earth
 Solar energy flux (integrated across all wavelengths) at Earths orbit
So  1366 Wm-2
 Insolation averaged over one day, and over all latitudes = global
average solar flux
So/4  341.5 Wm-2 (round to 342 Wm-2 below)
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Supplemental Figure
Daily total solar irradiance averaged over Earth’s surface (So/4)
Variations from mean.
Multi-satellite estimates
following Frohlich and
Lean, 1998; Frohlich 2003.
(NB: corrections between
satellites applied)
Note: 11 yr solar cycle (little
climate impact due to ocean
heat storage & large
natural variability)
Figure 2.7
Seasonal and latitude dependence of insolation
Orbital parameters vary on paleoclimate time scales, e.g., ice ages.
(Axial tilt precesses around orbit 23 kyr; tilt changes 41 kyr; eccentricity 100, 400 kyr;
kyr=1000 yr)
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
2.3 Globally averaged energy budget
Pathways of energy transfer in a global average
Figure 2.8
After Kiehl and Trenberth, 1997, Bull. Amer. Meteor. Soc.
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
 Albedo: fraction of incident solar
radiation that is reflected.
 Global average "planetary
albedo" 0.31 (=107/342).
 Deep clouds: albedo roughly 0.9,
ocean albedo 0.08.
 Some absorption of solar
radiation, e.g., in ozone layer
(UV).
 Aerosols = suspended particles.
 Most incoming sunlight that is
not reflected passes through the
atmosphere, absorbed at surface.
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
 Heat transfer from the surface upward: sensible heat, latent
heat and infrared emission.
 Sensible heat: contact between molecules, subsequent upward
transfer by parcels of hotter air (e.g. hot plumes known as
thermals; dry convection).
 Evaporation more effective means of cooling the surface;
stores energy as latent heat.
 Latent heat subsequently released when
water vapor condenses into clouds.
 Wherever there is precipitation, latent heat
remains in atmosphere.
 Cloud formation is most often associated
with overturning motions, known as moist
convection  transfers heat through a deep layer.
 Overall effect convective heating.
 Over land, vegetation plays such an important role in
evaporation that process is called evapotranspiration.
• plants access ground water, actively regulate water loss.
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
 Atmosphere emits IR downward  absorbed at surface.
 Greenhouse effect:
•The upward IR from the surface is
mostly trapped in the atmosphere, rather
than escaping directly to space, so it tends
to heat the atmosphere.
•The atmosphere warms to a temperature
where it emits sufficient radiation to
balance the heat budget, but it emits both
upward and downward, so part of the
energy is returned back down to the
surface where it is absorbed.
•This results in additional warming of the
surface, compared to a case with no
atmospheric absorption of IR.
•Both gases and clouds contribute to
absorption of IR and thus to the
greenhouse effect.
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
 At the top of the atmosphere, in the global average and for a
steady climate:
• IR emitted balances incoming solar.
 Global warming involves a slight imbalance:
• a change in the greenhouse effect slightly less IR emitted
from the top (chap. 6).
• small imbalance slow warming.
 Three roles for clouds and convection:
1. heating of the atmosphere (through a deep layer)
2. reflection of solar radiation (contributing to albedo)
3. trapping of infrared radiation (contributing to the
greenhouse effect)
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
2.4 Gradients of rad. forcing and energy transport by atm.
 Differences in input of solar energy between latitudes 
temperature gradients.
 These gradients would be huge if it were not for heat transport
in ocean & atmosphere and heat storage in ocean.
• e.g., North Pole in winter time would be extremely cold.
 Net solar energy input (Dec. climatology; Fig 2.9)
• net = incoming minus part reflected to space.
• essentially zero north of the Artic circle to about 385 Wm-2 near 30°S .
 Outgoing longwave radiation (Fig 2.9) varies much less as a
function of latitude because:
• atmospheric and oceanic transports are very effective at reducing
temperature gradients.
• the ocean stores some heat from the previous summer.
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Figure 2.9
Net solar energy input and output of longwave radiation
for December climatology as a function of latitude
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Figure 2.10
Annual average net solar energy input and output
of infrared radiation to space from Earth
 In the annual average climatology: rate of heat storage is small.
• differences between solar input and compensating IR is due to northsouth heat transport.
 Net loss at poles approximately balances net energy gain in the tropics.
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
2.5 Atmospheric circulation
Figure 2.11
Temperature as a
function of height
or pressure
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Figure 2.12
Latitude structure of the circulation
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Latitude structure of the circulation (cont.)
 Hadley cell: thermally driven,
overturning circulation, rising in
the tropics and sinking at
slightly higher latitudes (the
subtropics).
• transports heat poleward (to
roughly 30°N).
• rising branch assoc. with
convective heating and heavy
rainfall.
• subtropical descent regions: warm
at upper levels  hard to convect,
little rain.
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Latitude structure of the circulation (cont.)
 In midlatitudes, the average effect of the transient weather
disturbances transports heat poleward.
 Trade winds in the tropics blow westward (i.e., from east, so
known as easterlies).
• they converge into the Intertropical Convergence Zone (ITCZ),
i.e., the tropical convective zone.
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Latitude structure of the circulation (cont.)
 At midlatitudes surface winds are westerly (from the west).
 Momentum transport in Hadley cell and midlatitude transients
important to wind patterns.
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Figure 2.13
January
precipitation
climatology
July
precipitation
climatology
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Intertropical convergence
zones (ITZCs) or tropical
convection zones: heavy
precipitation features deep in
the tropics, (convergence
refers to the low level winds
that converge into these
regions).
Monsoons: tropical convection
zones move northward in
northern summer, southward
in southern summer,
especially over continents.
•e.g., Asian-Austral monsoon,
Pan-American monsoon,
African monsoon.
•traditionally monsoon was
defined by local reversals of
wind; now generalized.
Figure 2.13
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Not perfectly symmetric about
the Equator.
Variations in longitude
(departures from "zonal
symmetry"). e.g., eastern
Pacific has little rain, western
Pacific has intense rainfall.
More convection and
associated rising motion in the
western Pacific 
overturning circulations along
the equatorial band known as
the Walker circulation (see
Figure 2.14).
Storm tracks around 30-45°N.
Figure 2.13
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Figure 2.14
Equatorial Walker circulation
Adapted from Madden and Julian, 1972, J. Atmos. Sci. and Webster, 1983, Large-Scale Dynamical Processes in the Atmosphere
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Figure 2.15
DJF climatology
for upper (200 mb)
level winds
•Note subtropical jets;
strong near storm
tracks.
DJF climatology
for lower (925 mb)
level winds
•Note strong tropical
Pacific trade winds.
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
2.6 Ocean circulation
Sea surface temperature
climatology - January
Sea surface temperature
climatology - July
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Figure 2.16
 SST is warmest in tropics;
approaches freezing in
higher latitudes ( -2 C);
strongest SST gradient
occurs at midlatitudes.
 SST is not perfectly
symmetric about the
Equator.
 Variations in longitude. e.g.,
eastern Pacific is relatively
cold.
 Equatorial cold tongue:
along the equator in the
Pacific.
• maintained by upwelling
of cold water from below.
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Figure 2.16
 Rainfall over oceans has a
close, though not perfect,
relationship to SST pattern
(compare to Figure 2.13).
• in the tropics it tends to
rain over SST that is
warmer than SST in
neighboring regions.
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Figure 2.16
Figure 2.17
Ocean surface currents
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Ocean surface currents (cont.)
 Along the Equator, currents are in direction of the wind
(easterly winds drive westward currents [note terminology!]
 Off the Equator, currents need not be in the direction of the
wind. Currents set by change of the zonal wind with latitude
and Coriolis force (chap. 3). (“zonal" = east-west direction)
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
 Just slightly off the Equator, small component of the current
moves poleward; important because it diverges  produces
upwelling (chap. 3).
 Circulation systems known as gyres. In the subtropical gyres,
currents flow slowly equatorward in most of basin.
Compensating return flow toward poles occurs in narrow, fast
western boundary currents (Gulf stream, the Kuroshio, Brazil
currents).
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Ocean vertical structure
 Ocean surface is
warmed from above 
lighter water over
denser water (“stable
stratification”).
• incoming solar
radiation warms upper
10 m. Turbulence near
the surface mixes some
of this warming
downward.
• mixing driven by windgenerated turbulence
and instabilities of
surface currents.
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Figure 2.18
Ocean vertical structure (cont.)
 At thermocline, any
mixing of the denser
fluid below into lighter
fluid above requires
work  limits the
mixing.
 Deep waters tend to
remain cold
• on long time scales,
import of cold waters
from a few sinking
regions near the poles
maintains cold
temperatures.
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Figure 2.18
Ocean vertical structure (cont.)
 Ocean surface is
directly warmed by
solar radiation  loses
heat to atmosphere.
• air temperature a few
meters above the
surface tends to be
slightly colder than the
surface temperature.
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Figure 2.18
Figure 2.19
The thermohaline circulation
 Salinity (concentration of salt) affects ocean density in addition
to temperature.
 Waters dense enough to sink: cold and salty
 Thermohaline circulation: deep overturning circulation is
termed the (thermal for the temperature, haline from the greek
word for salt, hals).
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Figure 2.19
The thermohaline circulation (cont.)
 Deep water formation in a few small regions that produce
densest water
• e.g., off Greenland, Labrador Sea, regions around Antarctica.
 Small regions control temperature of deep ocean potential
sensitivity.
• likely player in past climate variations.
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
2.7 Land surface processes
Land surface impacts on physical climate; main effects:
 Land does not transport or store heat significantly.
 produces land-ocean contrast.
 Albedo depends on vegetation, snow, ice, & sometimes on
surface characteristics (e.g., Sahara Desert has a bright sandy
surface). Annual average albedo (next slide) is highest where there
is year-round snow/ice cover; intermediate values typically from
winter snow/summer vegetation; rainforests low albedo.
 Evapotranspiration and surface hydrology.
 "soil moisture" in subsurface layers of soil
 catchment basins, lakes, rivers...
 evapotranspiration: vegetation actively regulates
moisture loss.
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Figure 2.20
Annual average Albedo
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Traditional climate models specify characteristics of
vegetation from observations.
 Recently, interactive vegetation models.
 Biological activity on land impacts trace gases (e.g.,
carbon storage)
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Values based on Denman et al., 2007, IPCC; format follows
Sarmiento and Gruber, 2002, Physics Today
Figure 2.21a
The carbon cycle
Carbon transfers among reservoirs without human perturbation
 1 Petagram (Pg)=1 trillion kg=1 billion metric tons=1gigaton
(Gt); keeping track of the mass of carbon atoms in both organic
and inorganic (e.g., carbonate, bicarbonate, CO2) compounds
 Ocean 38,000 PgC; 3 PgC as marine biota; upper ocean 900 PgC
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Values based on Denman et al., 2007, IPCC; format follows
Sarmiento and Gruber, 2002, Physics Today
Figure 2.21a, The carbon cycle, cont’d
 land biomass reservoir (vegetation, soils and vegetation detritus
such as leaf litter) ~2300 PgC; preindustrial atm.<600 PgC
 coal, oil and natural gas est. > 4,000 PgC, much larger if
nonconventional sources like tar sands are included
 Dark arrows = Gross fluxes; Photosynthesis versus respiration;
CO2 dissolved in some ocean surface regions, released in others
 light arrows = net fluxes
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Human perturbation
Values based on Denman et al., 2007, IPCC; format follows
Sarmiento and Gruber, 2002, Physics Today
Figure 2.21b
The carbon cycle
 Fossil fuels 6.4 PgC/yr (1990s) (incl. 0.1 PgC/yr cement production)
 ~40% from coal (& growing), 40% from oil and oil derivatives
such as gasoline, 20% from natural gas
 1.6 PgC/yr land-use change; e.g., deforestation, with replacement
by agricultural land with smaller carbon storage
Values based on Denman et al., 2007, IPCC; format follows
Sarmiento and Gruber, 2002, Physics Today
Figure 2.21b, The carbon cycle, cont’d
 Based on 1990s data; anthropogenic emissions ~8 PgC/yr (6.4
fossil fuels + cement, 1.6 land use change)
 Fortunately, less than half remains in the atmosphere
 increased atmospheric concentrations yield 2.2 PgC/yr increased
flux into ocean
 + 2.5 PgC/yr taken up by land vegetation (e.g., forest regrowth?
Details uncertain, estimated as residual)
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Values based on Denman et al., 2007, IPCC; format follows
Sarmiento and Gruber, 2002, Physics Today
Figure 2.21b, The carbon cycle, cont’d
 Atmospheric concentrations of CO2 rise by about 1 ppm for each
2.1 PgC that remains in the atmosphere (from # molecules per
petagram of carbon compared to # of molecules in the atmosphere)
 So 8-2.2-2.5 = 3.3 PgC/yr = 1.6 ppm/year CO2 increase in atm.
(but this is variable in time)
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Figure 2.22
Values are from Denman et al., 2007, IPCC; format follows
Sarmiento and Gruber , 2002, Physics Today
Fossil fuel emissions and increases in atmospheric CO2 concentrations
 Fossil fuel emissions converted to CO2 concentration change if all
remained in the atmosphere (1 ppm for each 2.1 PgC); rising as
energy consumption increases
 Actual rate of accumulation (change in concentration each year;
all positive = rising concentration, but variable rate of increase)
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
Values are from Denman et al., 2007, IPCC; format follows
Sarmiento and Gruber , 2002, Physics Today
Figure 2.22 (cont.)
 Variations primarily due to land biosphere, e.g., droughts
associated with El Niño/southern oscillation
 Accumulation rate ~ 55% of fossil fuel emissions on average, but
not guaranteed to continue
 Anthropogenic land use change not separated due to estimation
challenges year by year ( e.g., deforestation versus regrowth)
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP
 For later reference: emissions sometimes given in gigatons of
CO2/yr. In those units, emissions differ by a factor of 3.66 x
PgC/yr because it also counts the weight of the two oxygen atoms.
 (molecular weight of CO2 44.01)/(atomic weight of carbon 12.01)= 3.66
 i.e. 12 gigatons of carbon in the fuel burned produces 44 gigatons of CO2.
 For perspective:
 1 US gallon of gasoline yields about 9 kg of carbon dioxide
 E.g., 1000 gallons/yr not unusual for personal vehicle (20,000
miles/20 mpg) gives 9 metric tons of CO2 from one person, just
from driving
 per capita CO2 emission each year:
 US 20 tons per person
 United Kingdom, Germany, Japan roughly 10 tons per person
 World average 4 tons per person (2002 values, and climbing)
Neelin, 2011. Climate Change and Climate Modeling, Cambridge UP