Chapter 2 Basics of Global Climate

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

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).

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

Figure 2.1 Schematic of components of the climate system

Table 2.1 Atmosphere Typical time scales Overall response time to heating months Typical spin-down time of wind if nothing is forcing it days Frontal system lifetime (1000s of km) days Convective cloud lifetime (100m to km horizontal; hours Time scale for typical upper level wind (20 m s -1 ) to days Ocean Response time of upper ocean (above thermocline) to months to years heating Response time of deep ocean to atmospheric changes decades to millennia Ocean eddy lifetime (10s to 100 km) month Ocean mixing in the surface layer hours to days Time for typical ocean current (cm s-1) to cross ocean decade (1000s of km)

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 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

Satellite image based on visible light Figure 2.2 Courtesy of NASA

Satellite image based on visible light with 5 x5 grid overlay Figure 2.2 Courtesy of NASA

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.

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.

Rainfall from the TRMM-based merged data (3B42RT) Weekly accumulation Rain rate from a 3- hourly 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 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. A cumulonimbus cloud in the tropics 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 spatial scale ~50 km computationally costly. Eddies tend to twist into very small scale features, become mixed transport heat poleward ("eddy transports") Optional: climate model simulation (Geophysical Fluid Dynamics Lab model; 13 mon with 10-25km ocean grid size) Courtesy of Ocean Remote Sensing Group, Johns Hopkins U., Applied Phys. Lab.

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/m 2 ). Blackbody radiation: approximation for how radiation depends on temperature (that does not depend on the substance doing the emitting; black = perfect absorber/emitter).

Blackbody radiation curves & absorption of radiation at each wavelength Figure 2.5 Fraction of radiation absorbed at each wavelength as it passes through whole depth of atmosphere Fraction absorbed above 11 km 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 = σt 4 σ = 5.67x10-8 Wm -2 K -4, T in Kelvin (celsius + 273.16) Actual surfaces or gases do not absorb or emit as a perfect blackbody define an emissivity ε for each substance R = εσt 4 absorptivity = emissivity. later bulk emissivity ε a for an atmospheric layer. Full climate models do detailed computation as a function of wavelength, for every level in the atmosphere,

Figure 2.6 Schematic of Sun's rays arriving at disk and spherical Earth Solar energy flux (integrated across all wavelengths) at Earths orbit S o 1366 Wm -2 Insolation averaged over one day, and over all latitudes = global average solar flux S o /4 341.5 Wm -2 (round to 342 Wm -2 below)

Supplemental Figure Daily total solar irradiance averaged over Earth s surface (S o /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)

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.

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.

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.

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.

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)

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.

Figure 2.9 Net solar energy input and output of longwave radiation (in Wm -2 ) for December climatology as a function of latitude

Figure 2.10 Annual average net solar energy input and output (in Wm -2 ) 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.

Temperature as a function of height or pressure 2.5 Atmospheric circulation Figure 2.11

Figure 2.12 Latitude structure of the circulation

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.

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.

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.

Figure 2.13 January precipitation climatology July precipitation climatology

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

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

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

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.

2.6 Ocean circulation Figure 2.16 Sea surface temperature (SST) climatology - January Sea surface temperature climatology - July

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. 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. Figure 2.16

Figure 2.17 Ocean surface currents

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)

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).

Ocean vertical structure Figure 2.18 Ocean surface warmed by solar radiation lighter water over denser water ( stable stratification ). solar radiation warms upper 10m. Turbulence near surface mixes warming downward. mixing driven by windgenerated turbulence and instabilities of surface currents. air temperature just above surface slightly colder but close to SST

Ocean vertical structure (cont.) Figure 2.18 At thermocline, any mixing of 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.

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 (thermal for the temperature, haline from the greek word for salt, hals). Also termed the Meridional Overturning Circulation.

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.

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.

Figure 2.20 Annual average Albedo

Traditional climate models specify characteristics of vegetation from observations. Recently, interactive vegetation models. Biological activity on land impacts trace gases (e.g., carbon storage)

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, CO 2 ) compounds Ocean 38,000 PgC; 3 PgC as marine biota; upper ocean 900 PgC 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 resources, est. > 4,000 PgC, much larger if nonconventional sources like tar sands are included Dark arrows = Gross fluxes; Photosynthesis versus respiration; CO 2 dissolved in some ocean surface regions, released in others light arrows = net fluxes 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% coal, 40% oil+derivatives (e.g., gasoline), 20% natural gas 1.6 PgC/yr land-use change; e.g., deforestation, with replacement by agricultural land with smaller carbon storage 2002-11 avg: fossil fuels 8.3± 0.7 PgC/yr; land use 0.9 fossil fuel increase: 2002 7.0,, 2011 9.5, 2012* 9.7 PgC/yr *prelim est. 2015 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) (8.3+0.9=9.2 in 2002-11) Fortunately, only ~ half remains in the atmosphere (time dependent) increased atmospheric concentrations yield 2.2 PgC/yr increased flux into ocean (2.4 in 2002-2011) + 2.5 PgC/yr taken up by land vegetation (e.g., forest regrowth? Details uncertain, estimated as residual; similar in 2002-2011) 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 CO 2 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 CO 2 increase in atm. but this is variable in time. For 2002-11: (9.2-2.4-2.5)/2.1=2.0ppm/yr. Est. overall change from ~1750 to 1994: atm. +165 PgC, ocean +118 PgC, land changes complex, e.g. forest clearing/regrowth Values based on Denman et al., 2007, IPCC; format follows Sarmiento and Gruber, 2002, Physics Today

Figure 2.22 Fossil fuel emissions and increases in atmospheric CO 2 concentrations Fossil fuel emissions converted to CO 2 concentration change if all remained in the atmosphere (1 ppm for each 2.1 PgC); rising as energy consumption increases (2.0ppm/yr incr. in 2002-11) Actual rate of accumulation (change in concentration each year; all positive = rising concentration, but variable rate of increase) 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 Anthropogenic land use change not separated due to estimation challenges year by year ( e.g., deforestation versus regrowth) Accumulation rate in atm. ~ 55% of fossil fuel emissions in recent decades avg. Land+ocean not guaranteed to continue at 45% slow deep ocean uptake leaves ~15-40% longer than 1000 years Values are from Denman et al., 2007, IPCC; format follows Sarmiento and Gruber, 2002, Physics Today

For later reference: emissions sometimes given in gigatons of CO 2 /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 CO 2 44.01)/(atomic weight of carbon 12.01)= 3.66 i.e. 12 gigatons of carbon in the fuel burned produces 44 gigatons of CO 2. 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 CO 2 from one person, just from driving per capita CO 2 emission each year*: US 17.6 tons/person (2011; down from 19.2 in 2009); Canada 16 Germany, Japan ~9-10 tons; United Kingdom 7.8 (9.1 in 2008) China, 6.1 tons per person (2011, was 4.7 in 2008); India 1.5 tons Philippines 0.8, Nigeria 0.5, Rwanda 0.07, Chad 0.02, Mali 0.05 World average ~4.6 tons per person (2011, 4.4 in 2008, climbing) * US Energy Information Administration data, 2015