the Energy (Im)Balance of Our Current Atmosphere.

Natural Climate Change Radiative Forcing Global Warming Potential Summary  

Driving Questions:

  • What has controlled past changes in Earth's temperature?
  • What processes lead to temperature change?
  • What gases comprise the atmosphere and how do they influence the temperature?
  • What are the global warming potentials for greenhouse gases?

1. Natural Climate Change

We believe that the temperature of the earth has varied wildly over the evolution of the earth. Figure 1 shows an estimate of temperature changes as complied by Scotese. So how can it be that the climate has changed so over the ages and what processes could lead to these changes?

Figure 1. Estimated changes in global temperature.

The processes for changing climate naturally over long timescales include:

  1. Plate tectonics
  2. Long-term carbon cycle
  3. Solar Variations, and
  4. Orbital Variations

Plate Tectonics

The movement of the continents has obviously influenced the climate at specific locations (Figure 2), but could also influence the global temperature by redistributing the collection of solar radiation and/or providing land masses on which continental glaciers could form.

Figure 2. Location of continents during the Devonian period from Scotese.

Long-term carbon cycle

On long timescales of millions of years, the CO2 composition of Earth's atmosphere has been controlled by the exchange of carbon between the atmosphere, life, and rocks. Volanic eruptions (both on land and in the ocean) and metamorphism are sources of CO2 to the atmosphere. Silicate weathering and organic carbon burial are processes that have consumed atmospheric CO2. The balance between these three processes have maintained Earth's climate within a habitable zone.

Cartoon of the long-term carbon cycle

Figure 3. The long-term (inorganic) carbon cycle (Bice, 2001). Missing from this figure is the portrayal of organic carbon burial.

Solar Variability

On geological timescales of millions of years, the Sun's luminosity has increased at about 1% per 100 million years. It is surprising then that Earth's climate has remained relatively stable. It is thought that the increase in solar radiation has been balanced by a long-term decrease in atmosphere CO2 through an increase in silicate weathering.

Orbital Variability

The Earth's orbit changes over time in ways that could influence the amount of energy received at the surface. Thes e include changes in eccentricity, precession of the equinox, and changes in the Earth's tilt (obliquity).

Eccentricity
The eccentricity of the Earth's oprbit changes with a period of 100,000 years. At the moment the Earth's orbit is fairly circular but in 50,000 years it will be more eccentric with the difference between aphelion (farthest) and perihelion (nearest) points in the orbit will become larger.
Precession of the Equinox

Also over time the time of year when the Earth reaches perihelion changes. Now the Earth is closest the sun in January and faerthest in July. The combination of changing eccentricity and precession of the equinox leads to changing available solar radiation.

Obliquity

Finally, the Earth wobbles on its axis of rotation changing the tilt of the Earth (and hence its seasonality) over a 41,000 year period. The tilt is now 23.5 but changes between 22.5 and 24.5.


2. Radiative Forcing

The temperature of the Earth's surface and atmosphere are dictated by a balance between incoming energy and outgoing energy. When more energy is received than lost, temperatures rise. The Earth's surface, for example, absorbs radiation from the Sun. This energy is then redistributed by the atmospheric and oceanic circulations and radiated back to space at longer (infrared) wavelengths. For the annual mean and for the Earth as a whole, the incoming solar radiation energy is balanced approximately by the outgoing terrestrial radiation. Any factor that alters the radiation received from the Sun or lost to space, or that alters the redistribution of energy within the atmosphere and between the atmosphere, land, and ocean, can affect climate. A change in the net radiative energy available to the global Earth-atmosphere system is termed a radiative forcing. Positive radiative forcings tend to warm the Earth’s surface and lower atmosphere. Negative radiative forcings tend to cool them.

Increases in the concentrations of greenhouse gases will reduce the efficiency with which the Earth’s surface radiates to space. More of the outgoing terrestrial radiation from the surface is absorbed by the atmosphere and re-emitted at higher altitudes and lower temperatures. This results in a positive radiative forcing that tends to warm the lower atmosphere and surface. Because less heat escapes to space, this is the enhanced greenhouse effect – an enhancement of an effect that has operated in the Earth’s atmosphere for billions of years due to the presence of naturally occurring greenhouse gases: water vapor, carbon dioxide, ozone, methane and nitrous oxide. The amount of radiative forcing depends on the size of the increase in concentration of each greenhouse gas, the radiative
properties of the gases involved (indicated by their global warming potential), and the concentrations of other greenhouse gases already present in the atmosphere. Further, many greenhouse gases reside in the atmosphere for centuries after being emitted, thereby introducing a long-term commitment to positive radiative forcing.

Anthropogenic aerosols (microscopic airborne particles or droplets) in the troposphere, such as those derived from fossil fuel and biomass burning, can reflect solar radiation, which leads to a cooling tendency in the climate system. Because it can absorb solar radiation, black carbon (soot) aerosol tends to warm the climate system. In addition, changes in aerosol concentrations can alter cloud amount and cloud reflectivity through their effect on cloud
properties and lifetimes. In most cases, tropospheric aerosols tend to produce a negative radiative forcing and a cooler climate. They have a much shorter lifetime (days to weeks) than most greenhouse gases (decades to centuries), and, as a result, their concentrations respond much more quickly to changes in emissions. Volcanic activity can inject large amounts of sulphur-containing gases (primarily sulfur dioxide) into the stratosphere, which are transformed into sulfate aerosols. Individual eruptions can produce a large, but transitory, negative radiative forcing, tending to cool the Earth’s surface and lower atmosphere over periods of a few years.

When radiative forcing changes, the climate system responds on various time-scales. The longest of these are due to the large heat capacity of the deep ocean and dynamic adjustment of the ice sheets. This means that the transient response to a change (either positive or negative) may last for thousands of years. Any changes in the radiative balance of the Earth, including those due to an increase in greenhouse gases or in aerosols, will alter the global hydrological cycle and atmospheric and oceanic circulation, thereby affecting weather patterns and regional temperatures and precipitation.

Solar radiation budget
Figure 5. Global average flow of shortwave (solar) radiation through the Earth's atmosphere.

Shortwave Radiation Budget

Solar radiation entering the Earth's atmosphere (called "shortwave" radiation) can be reflected off clouds, the surface, and air molecules and dust. On a global average this accounts for about 30% of incoming radiation (see Figure 5). This percentage is quantified as the albedo of the system.

Albedo = percentage of incoming radiation that is reflected back into space = 30% for Earth

Another 19% on average is absorbed by the atmosphere, mainly by ozone in the Earth's stratosphere. The remaining 51% is absorbed by the Earth's surface.

Over a long term average, the Earth and its atmosphere must radiate as much energy out to space as it receives from the sun, but over the course of a year or a day or as one moves geographically it is likely that such a balance will not be present. At night and in the winter, for example, there is less solar radiation producing an energy deficit and leading to lower temperatures at those times in general.

In order to understand the whole energy balance we must also consider the other means for exchanging energy between the Earth's surface, atmosphere and space.


Distribution of earth radiation
Figure 6: Energy exchange between the Earth's surface and its atmosphere.

Gains

Losses


51 Visible from Sun 7 Conduction, Convection
96 IR from atmosphere 23 Evaporation


117 IR radiation

147 net 147 net

Longwave Radiation Budget

As was learned earlier all objects emit radiation in an amount and at a wavelength dictated by the object's temperature. The 51% of shortwave radiation absorbed by the Earth's surface (Figure 5) heats the surface. But as the surface heats it emits radiation in the infrared back into the atmosphere.

Figure 6 shows the annual global average exchange of energy between the Earth's surface and the atmosphere. Note the 51% of original solar radiation is absorbed, but 117% of the original solar input is emitted to the atmosphere, how can this be?

The answer makes sense when we consider that the surface of a planet receives a great deal of energy from its own atmosphere. Thus the effect of the atmosphere is to warm the surface over the temperature above that resulting from the Sun's energy.

The atmosphere warms the Earth by "trapping" radiation, allowing the surface to warm to 300°K. At that temperature, the black body surface radiation is large enough to ensure that an equilibrium condition pertains. The atmosphere traps radiation through the action of certain gases, called Greenhouse Gases. These gases (e.g., CO2, H2O, NO, CFCs, CO) are very good at absorbing and re-emitting infrared radiation. They intercept the IR radiation from the ground and reflect some of the energy back to the ground, warming it up more than would occur otherwise.

 

 

3. Global Warming Potential

The Global Warming Potential (GWP) of a greenhouse gas is the ratio of global warming, or radiative forcing – both direct and indirect – from one unit mass of a greenhouse gas to that of one unit mass of carbon dioxide over a period of time. Hence this is a measure of the potential for global warming per unit mass relative to carbon dioxide.

Global Warming Potentials are presented in Table 1 for an expanded set of gases. GWPs are a measure of the relative radiative effect of a given substance compared to CO2, integrated over a chosen time horizon. New categories of gases in Table 1 include fluorinated organic molecules, many of which are ethers that are proposed as halocarbon substitutes. Some of the GWPs have larger uncertainties than that of others, particularly for those gases where detailed laboratory data on lifetimes are not yet available. The direct GWPs have been calculated relative to CO2 using an improved calculation of the CO2 radiative forcing, the SAR response function for a CO2 pulse, and new values for the radiative forcing and lifetimes for a number of halocarbons. Indirect GWPs, resulting from indirect radiative forcing effects, are also estimated for some new gases, including carbon monoxide. The direct GWPs for those species whose lifetimes are well characterized are estimated to be accurate within ±35%, but the indirect GWPs are less certain.

Table 1. Direct Global Warming Potentials (GWPs) relative to carbon dioxide (for gases for which the lifetimes have been adequately characterized). GWPs are an index for estimating relative global warming contribution due to atmospheric emission of a kg of a particular greenhouse gas compared to emission of a kg of carbon dioxide. GWPs calculated for different time horizons show the effects of atmospheric lifetimes of the different gases.
    Lifetime Global Warming Potential
    (years) (Time Horizon in Years)
 GAS     20 yrs 100 yrs 500 yrs
Carbon Dioxide CO2
 
1
1
1
Methane CH4
12.0
62
23
7
Nitrous Oxide N2O
114
275
296
156
Chlorofluorocarbons
 
 
 
 
CFC-11  
55
4500
3400
1400
CFC-12  
116
7100
7100
4100
CFC-115  
550
5500
7000
8500
Hydrofluorocarbons
 
 
 
 
HFC-23 CHF3
260
9400
12000
10000
HFC-32 CH2F2
5
1800
550
170
HFC-41 CH3F
2.6
330
97
30
HFC-125 CHF2CF3
29
5900
3400
1100
HFC-134 CHF2CHF2
9.6
3200
1100
330
HFC-134a CH2FCF3
13.8
3300
1300
400
HFC-143 CHF2CH2F
3.4
1100
330
100
HFC-143a CF3CH3
52
5500
4300
1600
HFC-152 CH2FCH2F
0.5
140
43
13
HFC-152a CH3CHF2
1.4
410
120
37
HFC-161 CH3CH2F
0.3
40
12
4
HFC-227ea CF3CHFCF3
33
5600
3500
1100
HFC-236cb CH2FCF2CF3
13.2
3300
1300
390
HFC-236ea CHF2CHFCF3
10
3600
1200
390
HFC-236fa CF3CH2CF3
220
7500
9400
7100
HFC-245ca CH2FCF2CHF2
5.9
2100
640
200
HFC-245fa CHF2CH2CF3
7.2
3000
950
300
HFC-365mfc CF3CH2CF2CH3
9.9
2600
890
280
HFC-43-10mee CF3CHFCHFCF2CF3
15
3700
1500
470
Fully fluorinated species
 
 
 
 
SF6  
3200
15100
22200
32400
CF4  
50000
3900
5700
8900
C2F6  
10000
8000
11900
18000
C3F8  
2600
5900
8600
12400
C4F10  
2600
5900
8600
12400
c-C4F8  
3200
6800
10000
14500
C5F12  
4100
6000
8900
13200
C6F14  
3200
6100
9000
13200
Ethers and Halogenated Ethers
 
 
 
 
CH3OCH3  
0.015
1
1
<<1
HFE-125 CF3OCHF2
150
12900
14900
9200
HFE-134 CHF2OCHF2
26.2
10500
6100
2000
HFE-143a CH3OCF3
4.4
2500
750
230
HCFE-235da2 CF3CHClOCHF2
2.6
1100
340
110
HFE-245fa2 CF3CH2OCHF2
4.4
1900
570
180
HFE-254cb2 CHF2CF2OCH3
0.22
99
30
9
HFE-7100 C4F9OCH3
5
1300
390
120
HFE-7200 C4F9OC2H5
0.77
190
55
17
H-Galden 1040x CHF2OCF2OC2F4OCHF2
6.3
5900
1800
560
HG-10 CHF2OCF2OCHF2
12.1
7500
2700
850
HG-01 CHF2OCF2CF2OCHF2
6.2
4700
1500
450

Summary

  • Greenhouse gases selective absorb infrared radiation, thus trapping energy in the atmosphere.
  • The atmosphere radiates energy to the surface at an average rate greater than the rate of incoming solar radiation.
  • Each greenhouse gas is characterized by its atmospheric lifetime and global warming potential.

Greenhouse Gases and the Greenhouse Effect Self Test

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