The Sun and the Earth’s Climate
- 5.4k Downloads
Variations in solar activity, at least as observed in numbers of sunspots, have been apparent since ancient times but to what extent solar variability may affect global climate has been far more controversial. The subject had been in and out of fashion for at least two centuries but the current need to distinguish between natural and anthropogenic causes of climate change has brought it again to the forefront of meteorological research. The absolute radiometers carried by satellites since the late 1970s have produced indisputable evidence that total solar irradiance varies systematically over the 11-year sunspot cycle, relegating to history the term “solar constant”, but it is difficult to explain how the apparent response to the Sun, seen in many climate records, can be brought about by these rather small changes in radiation. This article reviews some of the evidence for a solar influence on the lower atmosphere and discusses some of the mechanisms whereby the Sun may produce more significant impacts than might be surmised from a consideration only of variations in total solar irradiance.
KeywordsOzone Solar Activity Solar Cycle Solar Irradiance Radiative Force
At periods of higher solar activity the Earth is subject to enhanced solar irradiance, a greater incidence of solar energetic particles and fewer galactic cosmic rays. This review presents an assessment of the extent to which solar variability affects the climate of the Earth’s lower and middle atmosphere and considers how the apparent response is brought about.
Summary of routes through which solar variability may influence the climate of the lower atmosphere.
Total solar irradiance (variations due to orbital variations or to variable solar emission).
Radiative forcing of climate. Direct impact on sea surface temperatures and hydrological cycle.
Solar UV irradiance.
Heating the upper and middle atmosphere, dynamical coupling down to troposphere. Middle and lower atmosphere chemistry and composition; impacts temperature structure and radiative forcing.
Solar energetic particles.
Ionisation of upper and middle atmosphere; impact on composition and temperatures. Magnetosphere — ionosphere — thermosphere coupling.
Galactic cosmic rays
Ionisation of lower atmosphere; impact on electric field. Impact on condensation nuclei.
First, Section 2 considers how information on past climates is derived and look at some of the (mainly circumstantial) evidence that variations in solar activity have affected climate on a wide range of timescales.
2 Climate Records
2.1 Measurements and reconstructions
Assessment of climate variability and climate change depends crucially on the existence and accuracy of records of meteorological parameters. Ideally records would consist of long time series of measurements made by well-calibrated instruments located with high density across the globe. In practice, of course, this ideal cannot be met. Measurements with global coverage have only been made since the start of the satellite era about 25 years ago while reliable instrumental records, covering the past few centuries, are only available from a few locations, mainly in Europe. For longer periods, and in remote regions, records have to be reconstructed from indirect indicators of climate known as proxy data. A readable review of the methods involved in the accumulation of climate data, and in its interpretation, is provided by Burroughs (1992).
The longest homogeneous instrument-based temperature series in the world is the Central England Temperature record dating back to 1659. It was first constructed in the 1970s from an accumulation of measurements made by amateur meteorologists in central lowland England. The construction of a homogeneous record requires knowledge of how, and at what time of day, the measurements were made and how local conditions may account for regional variations. Other similar temperature records dating back to the mid 18th century are available for Munich, Vienna, Berlin and Paris among other sites in Europe and at least one in the Eastern United States. These datasets have been extensively analysed for periodic variations in temperature. Generally they show clear indications of variations on timescales of about 2.2 to 2.4 years and 2.9 to 3.9 years but on longer timescales individual records show peaks at different frequencies with little statistical significance.
The most complete record of rainfall comes from eastern China where careful observations of floods and droughts date back to the fifteenth century. Again spectral analysis results in periodicities which vary from place to place.
Other climate records suggesting that the climate has been changing over the past century include the retreat of mountain glaciers, sea level rise, thinner Arctic ice sheets and an increased frequency of extreme precipitation events (IPCC, 2007).
Proxy data provide information about weather conditions at a particular location through records of a physical, biological or chemical response to these conditions. Some proxy datasets provide information dating back hundreds of thousands of years which make them particularly suitable for analysing long term climate variations and their correlation with solar activity.
Much longer records of temperature have been derived from analysis of oxygen isotopes in ice cores obtained from Greenland and Antarctica. The ratio of the concentration of 18O to that of 16O, or 2H to 1H, in the water molecules is determined by the rate of evaporation of water from tropical oceans and also the rate of precipitation of snow over the polar ice caps. Both these factors are dependent on temperature such that greater proportions of the heavy isotopes are deposited during periods of higher global temperatures. As each year’s accumulation of snow settles the layers below become compacted so that at depths corresponding to an age of more than 800 years it becomes difficult to precisely date the layers. Nevertheless, variations on timescales of more than a decade have been extracted dating back over hundreds of thousands of years.
Evidence of very long term temperature variations can also be obtained from ocean sediments. The skeletons of calciferous plankton make up a large proportion of the sediments at the bottom of the deep oceans and the 18O component is determined by the temperature of the upper ocean at the date when the living plankton absorbed carbon dioxide. The sediment accumulates slowly, at a rate of perhaps 1 m every 40,000 years, so that changes over periods of less than about 1,000 years are not detectable but ice age cycles every 100,000 years are clearly portrayed.
2.2 Solar signals in surface climate
Many different approaches have been adopted in the attempt to identify solar signals in climate records. Probably the simplest has been the type of spectral analysis mentioned above, in which cycles of 11 (or 22 or 90 etc.) years are assumed to be associated with the Sun. In another approach time series of observational data are correlated with time series of solar activity. An extension of the latter method uses simple linear regression to extract the response in the measured parameter to a chosen solar activity forcing. A further development allows a multiple regression, in which the responses to other factors are simultaneously extracted along with the solar influence. More sophisticated statistical techniques, involving e.g. pattern-matching, have also been employed. Each of these approaches gives more faith than the previous that the signal extracted is actually due to the Sun, and not to either some other factor or to random fluctuations in the climate system, and many interesting results emerge. It should be remembered, however, that such detection, while potentially robust in statistical terms, is not based on any understanding of how the presumed solar influence takes place. In the remainder of this section some results are presented from studies involving correlations and regressions of meteorological data with solar activity indices. Some of the mechanisms which have been proposed to explain how these changes take place are the subjects of Sections 3 to 7.
Section 2.1 mentioned how temperature records may be extracted from ice cores and ocean sediments. These media may also preserve information on cosmic ray flux, and thus solar activity, in isotopes such as 10Be and 14C. Thus simultaneous records of climate and solar activity may be retrieved. An example is given in Figure 4 which shows fluctuations on the 1,000 year timescale well correlated between the two records, suggesting a long-term solar influence on climate.
Many studies have purported to show variations in meteorological parameters in phase with the “11-year” solar activity cycle. Some of these are statistically questionable and some show signals that appear over a certain period only to disappear, or even reverse, over another period. The evidence of a solar influence on climate on solar cycle timescales is, however, becoming increasingly robust.
Multiple regression analysis of the surface Northern Annular Mode (NAM) (in winter) and Southern Annular Mode (SAM). The annular modes are the first empirical orthogonal functions of 90-day low-pass filtered anomalies, poleward of 20 in each hemisphere, of 1000 hPa geopotential height in normalised units. Positive values indicate a stronger equator-to-pole temperature gradient and more zonal flow. Columns show regression coefficients for linear trend (N.H. only), stratospheric chlorine (S.H. only), El Nino-Southern oscillation (ENSO), volcanic (stratospheric) aerosol loading, solar variability (10.7 cm index) and the Quasi-Biennial Oscillation (QBO). Sol*QBO indicates that an index composed of a product of the solar and QBO indices was used in place of those two factors individually. The data cover the period 1958–2001. Colours indicate the statistical significance levels of the values, derived using a Student’s t-test: 99%, 95%, 90%, 80%, < 80%. From Haigh and Roscoe (2006).
2.3 Solar signals throughout the atmosphere
Another important aspect of the Sun’s influence on the stratosphere is in the modulation of ozone concentrations. This is described in Section 5.4.
3 Earth Radiation Budget and Radiative Forcing of Climate Change
This section introduces the concept of radiative forcing as a tool for estimating the impact on surface temperature of perturbations to the Earth’s radiation budget, such as might be due to changes in solar irradiance.
3.1 Earth radiation budget
The temperature and emissivity of the surface are such that 390 Wm−2 of infrared energy are emitted into the atmosphere. Only 40 Wm−2 of this, however, escapes to space with the remainder being absorbed by atmospheric gases and cloud. The atmosphere returns 324 Wm−2 to the surface. The energy balance at the surface is achieved by non-radiative processes such as evaporation and convection. The radiation balance at the top of the atmosphere is achieved by the 195 Wm−2 of heat energy emitted to space by the atmosphere and clouds.
3.2 Radiative forcing of climate change
In equilibrium the net global and annual mean radiative flux at the top of the atmosphere (TOA) is zero. Consider, however, a situation in which some new factor (such as changes in solar irradiance, planetary albedo or the concentrations of radiatively active gases, aerosols or cloud) acts to perturb the absorbed solar or emitted infrared radiation. Then, before surface and atmospheric temperatures have had time to adjust and to establish a new equilibrium, the net flux at TOA is not zero. The simplest definition of radiative forcing (RF) is just the (hypothetical) instantaneous change in the value of the net downward flux. If the RF is positive then there is an increase in energy entering the system (or equivalently a decrease in energy leaving the system) and it will tend to warm until the outgoing energy matches the incoming and the net flux is again zero.
The large, factor 3, range in the value of λ given above represents the spread of values given by different GCMs. This gives an indication of the uncertainties in climate prediction. It should be noted, however, that for each particular GCM the range of λ found using different sources of radiative forcing is much narrower. This suggests that, while absolute predictions are subject to large uncertainty, the forecast of the relative effects of different factors is more reliable.
It has been found that the value of λ, and thus the usefulness of the radiative forcing concept, is more robust if, instead of using the instantaneous change in net flux at the top of the atmosphere, RF is defined at the tropopause with the stratosphere first allowed to adjust to the imposed changes. Thus a formal definition of radiative forcing, as used by the Intergovernmental Panel on Climate Change (IPCC) is the change in net flux at the tropopause after allowing stratospheric temperatures to adjust to radiative equilibrium (but with surface and tropospheric temperatures held fixed). The effects of the stratospheric adjustment are complex as can be illustrated by the case of changes in stratospheric ozone: an increase in ozone masks the lower atmosphere from solar ultraviolet, i.e. reducing the net flux at the tropopause and thus RF. However, the presence of ozone in the lower stratosphere increases the downward infrared emission (and thus RF) both directly through the 9.6 band and also indirectly through the increase in stratospheric temperatures which it produces. Whether the net effect is positive or negative depends on whether the shortwave or longwave effect dominates and this is determined by the vertical distribution of the ozone change.
The direct effect of an increase in total solar irradiance is to increase the radiative forcing. The heating of the stratosphere by the additional irradiance will enhance this by increasing the downward emission of thermal radiation. However, the sign of the radiative forcing due to any solar-induced increases in ozone is not clear — published estimates show both positive and negative values — because of the uncertainties in the distribution of the ozone change (see Section 5.5).
The solar contribution is assessed to be in the range 0.06–0.30 Wm−2. Note that when calculating solar radiative forcing it is necessary to scale the total solar irradiance at the Earth by a factor taking into account geometric considerations as well as the planetary albedo. Thus the RF due to a change in TSI of 1 Wm−2 is about 1/6 Wm−2, or a change in TSI of 0.7 Wm−2 since 1750 is equivalent to RF = 0.12 Wm−2. The actual variations in TSI over the past few centuries is very uncertain (see Section 4.2) and the change in TSI depends crucially on the starting date (chosen as 1750 by the IPCC to represent the pre-industrial atmosphere): choice of earlier or later in the 18th century would have given an increased solar RF. Thus the value of solar radiative forcing in the IPCC figure is largely indicative. Taking a value of the climate sensitivity parameter λ of 0.6 K (Wm−2)−1 suggests that a global λ of 0.6 K (Wm−2)−1 suggests that a global average surface warming of less than 0.1 K since 1750 could be ascribed to the Sun. However, the IPCC gives the assignation “very low” to the LOSU associated with solar radiative forcing, thereby acknowledging that there may be factors as yet unknown, or not fully understood, which may act to amplify (or even diminish) its effects.
4 Total Solar Irradiance
4.1 The Earth’s orbit around the Sun
Averaged over the globe and over a year the solar energy flux at the Earth depends only on e but seasonal and geographical variations of the irradiance depend on θt and e sin p. However, it is not just the temperatures of individual seasons that are at stake: the intensity of radiation received at high northern latitudes in summer determines whether the winter growth of the ice cap will recede or whether the climate will be precipitated into an ice age. Thus changes in seasonal irradiance can lead to much longer term shifts in climatic regime.
Figure 3 showed temperature variations associated with transitions between ice ages and interglacials but also presented the concentrations of methane and carbon dioxide preserved in the ice core showing a strong correlation between these and temperature. (NB neither concentration reaches anything like the present day values of about 1700 ppbv and 380 ppmv respectively). One theory (Petit et al., 1999) suggests that the warming of southern high latitudes caused by the orbital variations is amplified by the release of CO2 from the southern oceans and this warming is then further amplified through a reduction in albedo resulting from the melting of Northern Hemisphere ice sheets. Such positive feedback mechanisms might explain the sharp increases in temperature seen in the record.
Cyclical variations in climate records with periods of around 19, 23, 41, 100 and 413 kyr are generally referred to as Milankovitch cycles after the geophysicist who made the first detailed investigation of solar-climate links related to orbital variations.
4.2 Variations in total solar irradiance
Direct measurements of TSI made outside the Earth’s atmosphere began with the launch of satellite instruments in 1978. Previous surface-based measurements did not provide sufficient accuracy, as they were subject to uncertainties and fluctuations in atmospheric absorption that may have swamped the small solar variability signal.
There is a related uncertainty, however, in the existence of any underlying trend in TSI over the past 2 cycles. Figure 21b presents one attempt to composite the measurements into a best estimate. It shows essentially no difference in TSI values between the cycle minima occurring in 1986 and 1996. The results of Willson and Mordinov (2003), however, show an increase in irradiance of 0.045% between these dates. The discrepancy hinges on assumptions made concerning the degradations of the Nimbus7/ERB and ERBS/ERBE instruments, data from which fills the interval, from July 1989 to October 1991, between observations made by the ACRIM I and II instruments. If such a trend were maintained, it would imply an increase in radiative forcing of about 0.1 Wm−2 per decade. Compared in terms of climate forcing, this is appreciable, being about one-third that due to the increase in concentrations of greenhouse gases averaged over the past 50 years. These discrepancies are important because all the available TSI reconstructions discussed below either use directly, or are scaled to fit, the recent satellite measurements, using one or other of the TSI composites.
It is clear that the estimates diverge as they go back in time so that the value of solar radiative forcing over this period is highly uncertain. Furthermore, this plot does not include some estimates which suggest a much larger difference between the present level and that prevailing during the Maunder Minimum in sunspots which occurred during the latter part of the 17th century (e.g. Hoyt and Schatten, 1998).
5 Solar spectral irradiance
The interaction of solar radiation with the atmosphere is fundamental in determining its temperature structure and in controlling many of the chemical processes which take place there. This section outlines the role of solar UV radiation in determining the budget of stratospheric ozone and discusses how the composition and thermodynamic structure of the stratosphere are modulated by solar variability. It goes on to consider how perturbations to the stratosphere may impact solar radiative forcing of tropospheric climate.
5.1 Absorption of solar spectral radiation by the atmosphere
The spectral composition of solar irradiance is thus important in determining at what altitudes it is absorbed and produces local heating.
5.2 Variability of solar spectral irradiance and heating rates
It is apparent that much larger fractional changes take place at shorter wavelengths than the ∼ 0.1% variation in TSI, shown in Figure 22, which represents the changes in the visible portion of the spectrum. Thus the direct impact of solar irradiance variability is larger in the middle and upper atmosphere than it is at lower altitudes.
Also shown in Figure 27 (lower panels) are the differences in fluxes and heating rates between 11-year solar cycle minimum and maximum conditions, based on the spectral differences shown in Figure 26. At the top of the atmosphere the increases in incoming radiation in the visible and UV regions are of similar magnitude but the stronger absorption of UV produces much greater heating. The spectral changes prescribed for these calculations were such that near-infrared radiation was actually weaker at solar maximum so decreases in heating rate are shown. This is contentious but the changes are anyway very small being about one part in ten thousand.
5.3 Stratospheric photochemistry
If the solar spectral irradiance varies then, without a change composition, the spectral heating rate varies in proportion to this change. If, however — as is actually the case, the atmospheric composition also responds to solar variability then this will affect solar fluxes and heating rates in a non-linear fashion. Particularly important in this context is stratospheric ozone.
The first of these reactions represents the photodissociation of oxygen at wavelengths less than 242 nm. This process is the key step in ozone formation because the oxygen atoms produced react with oxygen molecules to produce ozone molecules, as depicted in the second reaction (the M represents any other air molecule whose presence is necessary to simultaneously conserve momentum and kinetic energy in the combination reaction). Because the short wavelength ultraviolet radiation gets used up as it passes through the atmosphere, concentrations of atomic oxygen increase with height. This would tend to produce a similar profile for ozone but the effect is counterbalanced by the need for a 3-body collision (reaction 2) which is more likely at higher pressures (lower altitudes). Thus a peak in ozone production occurs at around 50 km. The third reaction is the photodissociation of ozone, mainly by radiation in the Hartley band (λ < 310 nm), into one atom and one molecule of oxygen. This does not represent the fundamental destruction of the ozone because the oxygen atom produced can quickly recombine with an oxygen molecule. The fourth reaction represents the destruction of ozone by combination with an oxygen atom. The fifth and sixth reactions represent the destruction by any catalyst X, which may include OH, NO and Cl. The various destruction paths are important at different altitudes but the combined effect is an ozone concentration profile which peaks near 25 km in equatorial regions.
The minimum in ozone column which occurs near the south pole in spring (October) has deepened considerably during the past thirty years into what has become known as the “ozone hole”. Observations and theoretical studies have shown that the depletion occurs mainly in the lower stratosphere and is due to catalytic destruction of ozone by chlorine radicals activated on the surface of polar stratospheric cloud particles during the polar night. This chlorine is a product of the breakdown of anthropogenic chlorofluorocarbons (CFCs) which are now banned under international agreement. Because of the long atmospheric lifetime of CFCs, and exacerbation of the situation by colder stratospheric temperatures associated with greater CO2 concentrations, it may be several decades before the winter polar stratosphere recovers to its unperturbed state.
5.4 Solar cycle signal in stratospheric ozone
5.5 Impact of the stratosphere on solar radiative forcing of climate
A summary of published estimates of solar radiative forcing. 1st column: reference; 2nd: nominal solar variability; 3rd and 4th: solar UV radiative forcing at the top of atmosphere and at the tropopause; 5th solar-induced ozone change; 6th, 7th and 8th: impact of ozone change on shortwave and longwave components of radiative forcing and the net effect; 9th: percentage amplification of solar forcing due to change in ozone.
ΔS RF (toa)
ΔS RF (tpse)
O3 SW effect
O3 LW effect
Net O3 effect
+ve peak near 40 km
Hansen et al. (1997)
+ve 10–150 hPa
Myhre et al. (1998)
Wuebbles et al. (1998)
0.49 to 0.70
0.42 to 0.60
+ve peak near 40 km
−30 to −21
Larkin et al. (2000)
+ve (as Haigh, 1994)
Shindell et al. (2001)
0.30 to 0.39
0.26 to 0.33
−ve (upper strat)
+6 to +8
To calculate radiative forcing a knowledge of changes to the temperature of the stratosphere are also required, adding a further complication. The next section considers how well the effects of solar ultraviolet variability on the thermal structure of the middle atmosphere are understood.
6 Models and Mechanisms
This section outlines potential mechanisms involved in producing some of the solar signals outlined in Sections 2.2 and 2.3. It discusses how the atmosphere may respond to changes in both total and spectrally-resolved solar irradiance through radiative and dynamical processes. It also considers how climate models can be used to test potential mechanisms and to what extent climate models are able to simulate the observed signals.
6.1 Climate change in response to variations in total solar irradiance
Figure 34b, however, shows the derived magnitudes of the forcings. Here the value 1 indicated that the derived magnitude equals that the model gives using standard radiative forcing estimates. The model appears to be underestimating the solar influence by a factor of 2 or 3 implying that some amplification factors of the solar influence are not incorporated into the model’s representation.
This result, however, is sensitive to the choice of TSI reconstruction. The Stott et al. (2003) work used the Hoyt and Schatten (1998) series, which has a large secular variation, although more recent work (see the discussion in Section 4.2) is suggesting much smaller long-term variability in TSI. If the latter were used then an even larger amplification of the model’s response would be required to match observations. This is an intriguing suggestion but this work needs to be reassessed using different GCMs to check that it is not an artefact of the specific one used in that study.
Nevertheless, most GCM runs which only include variations in TSI are unable to reproduce the distribution of temperature response shown in Sections 2.2 and 2.3, confirming that something is lacking in their ability to simulate the response of climate to solar activity. One recent study by Meehl et al. (2003), however, presented results from a model experiment which produced changes in SST similar to those shown in Figure 9. The authors explained these in terms of a response of the mean overturning circulations to sea surface temperature gradients enhanced between cloudy and clear regions. This intriguing possibility remains to be validated.
To explain the model underestimate it is necessary to find some factors which amplifies the effect from that derived simply by consideration of total solar irradiance as the primary driving mechanism behind the impact of solar variability on climate. Potentially one such amplification mechanism is through the effects of variations in solar UV radiation on the stratosphere.
6.2 Model studies of the influence of varying UV in the middle atmosphere
In Section 2.3 the thermal response of the stratosphere/troposphere to solar variability was shown to be largest near the tropical stratopause with lobes of warming in the sub-tropics in the lower stratosphere and bands of warming through mid-latitudes in the troposphere. One route to understanding this structure is to see if it can be reproduced in atmospheric models.
A recent intercomparison of coupled chemistry climate models (Austin, 2007) suggests that this may be true for short model runs but finds that the mid-stratospheric minimum can be reproduced in simulations which include time-varying solar irradiance and prescribed sea surface temperatures. These differ from the 3D models in Figure 37 which used fixed solar max/min scenarios with climatological SSTs. Austin (2007) argue that the transience in the simulations allows better reproduction of the mean meridional circulation of the stratosphere and thus the transport of lower stratospheric ozone. Nevertheless, another state-of-the-art coupled chemistry climate model, with very high vertical resolution (Schmidt and Brasseur, 2006) does reproduce the vertical structure with time-slice (i.e. not transient) runs. The response of stratospheric ozone to solar variability remains an active topic of research.
6.3 Dynamical mechanisms in the middle atmosphere
Gray (2003) pointed out that such a relationship between equatorial winds in the upper stratosphere and the timing of sudden stratospheric warmings could help to explain the interaction between the solar cycle and QBO influences on polar temperatures (as identified by Labitzke and van Loon, 1992, see Figure 15). Modelling evidence to support this idea has been provided by Matthes et al. (2004) who found they could reproduce the observed solar cycle/QBO polar temperature relationship if typical QBO wind profiles were imposed through the depth of the tropical stratosphere.
6.4 Model studies of the influence of varying solar UV on the troposphere
The modelled signals for zonal wind (Figure 40) and mean meridional circulation (Figure 41) are broadly similar to those deduced from observational (NCEP reanalysis) data (Figure 16 and Figure 17), although somewhat weaker in magnitude. The models used in these studies had fixed sea surface temperatures which essentially restricted them from responses involving feedbacks between SSTs, clouds and circulation and the results suggest that at least part of the solar signal in the troposphere comes from a response to changes in the atmosphere above.
Recently an atmosphere-ocean GCM with fully coupled stratospheric chemistry has been run (despite huge computational demands) to simulate the effects of changes in solar irradiance between the Maunder Minimum and the present (Shindell et al., 2006). As in the previous studies the results show a weakened Hadley circulation when the Sun is more active, and they also suggest an impact on the hydrological cycle with greater tropical precipitation. Furthermore, they provide additional evidence that coupling with stratospheric chemistry enhances the solar signal near the surface.
6.5 Coupling between the stratosphere and troposphere
The Kodera and Gray studies discussed in Section 6.2 provide evidence, and mechanisms, whereby solar perturbations to the upper stratosphere may affect the lower stratosphere, in both cases enhancing any direct solar heating in this region. However, they do not explain the apparent subsequent propagation of the solar signal downwards into the troposphere.
There is some observational evidence that variations in the strength of the polar vortex in the upper stratosphere may subsequently influence surface climate. A study of polar temperature trends by Thompson et al. (2005) suggests a downward influence, and modelling experiments by Gillett and Thompson (2003) demonstrate that depletion of stratospheric ozone over the south pole can affect the troposphere after about one month. Neither of these studies is specifically concerned with a solar influence but the accumulating evidence suggests that any factor influencing the strength of the polar stratospheric jet may be able to influence surface climate, at least at high latitudes.
Similarly the apparent success of the tropospheric GCM studies (see Section 6.4) in simulating the observed response to solar variability provides intriguing evidence that changes to the stratosphere, specifically induced by variations in solar UV radiation and resulting changes in ozone, can influence the troposphere. But they do not provide a detailed understanding of the mechanisms whereby these effects take place.
Recently some effort to advance understanding of the mechanisms of stratosphere-troposphere coupling has been made through the use of simplified general circulation models (Polvani and Kushner, 2002; Kushner and Polvani, 2004; Haigh et al., 2005; Haigh and Blackburn, 2006). These models include a full representation of atmospheric dynamics but only highly-parameterised representations of radiative and cloud processes so that multiple runs can be carried out. These experiments are not intended to simulate solar (or any other specific) forcing factors but to identify and investigate possible mechanisms for stratosphere-troposphere coupling.
In order to investigate the chain of causality involved in converting the stratospheric thermal forcing to a tropospheric climate signal another experiment used an ensemble of model spin-ups to analyse the time development of the response to an applied stratospheric perturbation (Haigh and Blackburn, 2006). It was found that the initial effect of the change in static stability at the tropopause is to reduce the eddy momentum flux convergence in this region. This is followed by a vertical transfer of the momentum forcing anomaly by an anomalous mean circulation to the surface, where it is partly balanced by surface stress anomalies. The unbalanced part drives the evolution of the vertically integrated zonal flow. It was concluded that solar heating of the stratosphere may produce changes in the circulation of the troposphere even without any direct forcing below the tropopause and that the impact of the stratospheric changes on wave propagation is key to the mechanisms involved.
This work is beginning to provide us with an understanding of how, through the spectral composition of solar irradiance, apparently small changes in solar irradiance may significantly impact the circulation of the lower atmosphere.
Clouds play a major part in establishing the heat and radiation budgets of the atmosphere. They transport latent heat from the oceans to the atmosphere, they reflect solar radiation back to space, reducing the net incoming radiative flux, and they trap infrared radiation, acting in a similar way to greenhouse gases. Any factor influencing cloud cover thus has the potential to seriously affect climate.
In this section the role of cloud in the radiation budget, and how this impacts radiative forcing, is reviewed. There follows a brief description of how clouds are formed and how their radiative properties relate to their microphysical structure. Finally the potential for cloud cover to be modified by variations in atmospheric ionisation by cosmic rays is briefly discussed.
7.1 Clouds and the Earth radiative budget
Figure 18 showed that on a global average clouds increase the planetary albedo by reflecting incoming solar irradiance back to space but decrease outgoing longwave radiation by acting in a similar way to greenhouse gases. The magnitude of the reflectance depends on the optical thickness of the cloud, the water phase (liquid or ice) and the cloud particle sizes and shapes. The degree of longwave trapping depends on the transmissivity of the cloud and also its temperature: high (cold) cloud is more effective because it emits less radiation to space while trapping the (warm) radiation from below. The net effect of cloud on the radiation budget depends on whether the shortwave or longwave effect is larger and thus on the location, height and microphysical properties of the cloud.
A factor which induces a change in cloud cover, drop size or altitude will introduce a radiative forcing. If, however, such a change is brought in response to another forcing factor then it should be viewed as a feedback on the initial forcing. For example, an increase in greenhouse gases might cause a surface warming, enhanced convection and an increase in cloud cover. The thick convective cloud produced would have a negative radiative forcing and thus reduce that due to the greenhouse gases alone. Such feedback effects, however, are implicitly included in the value of the climate forcing parameter λ. Thus the cloud produced by a dynamical response to other forcings can not be viewed as an additional forcing component. Only if changes to cloud properties are induced in situ by chemical or microphysical processes can they produce a radiative forcing, in the climate change sense, in their own right.
7.2 Cloud formation
Clouds form when the water vapour in the air condenses. Generally this occurs in air-masses which are rising, expanding and therefore cooling. The cooler air has a lower saturated vapour pressure so that the air-mass, with a given humidity, becomes saturated. Air masses rise either due to convection or through being forced to pass over some sort of barrier.
Cooling the air to saturation point is not, however, a sufficient condition for cloud to form: it is possible for the relative humidity to reach 500% without any spontaneous condensation taking place. Due to the energy associated with the surface tension of a droplet it is energetically unfavourable for a small droplet to grow and the water vapour requires a suitable surface, called a condensation nucleus, on which to condense. If the condensation nucleus is not a water surface then heterogeneous nucleation is said to take place; if it is then homogeneous nucleation occurs. In the free atmosphere, however, heterogeneous nucleation is the only important process because homogeneous nucleation requires prohibitively high relative humidities. (For a comprehensive discussion of cloud formation see the classic text by Ludlam, 1980).
Particles which act as condensation nuclei include sea salt, sulphates, mineral dust and aerosols produced from biomass burning. The concentration and composition of atmospheric aerosol vary geographically with, for example, sulphate aerosol being more abundant in the northern hemisphere as it is generated in industrial regions. A region with a higher concentration of condensation nuclei will produce a larger number of smaller cloud droplets than a remote area with clean air which will produce fewer larger droplets for the same total water content.
7.3 Atmospheric ionisation and cloud
It has been proposed, originally by Dickinson (1975) who acknowledged that his idea was entirely speculative at the time, that variations in cosmic rays could provide a mechanism whereby solar activity would produce a direct impact on cloud cover by modulating atmospheric ionization, resulting in the electrification of aerosol and increasing its effectiveness to act as condensation nuclei. The processes involved are complex but if they do take place then there is scope for considerable amendment to the value for solar radiative forcing of climate based on incident irradiance alone. Other processes, whereby changes in the Earth’s electric field might modify cloud cover have also been proposed (see e.g. Tinsley, 2000) but these are not discussed here.
Ionisation of molecules in the lower atmosphere is brought about by cosmic rays and by naturally-occurring radioactivity. The latter consists of airborne alpha-particle emitters (such as radon gas) and direct gamma radiation from the soil. Cosmic radiation consists of extremely high energy (> GeV) particles, mostly protons and helium nuclei. Both sources cause an electron to separate from a molecule of nitrogen or oxygen; the electron then being captured by a neutral molecule on a very short timescale. Thus equal numbers of positive and negative ions are produced. Other processes can introduce a net charge into the atmosphere; these include combustion, rainfall and breaking ocean waves.
Solar activity modulates the heliospheric magnetic field which acts as a shield to cosmic rays. Thus, during periods of higher solar activity fewer cosmic rays reach the Earth, although the modulation primarily affects lower energy cosmic radiation. At the Earth’s surface cosmic rays are monitored by neutron monitors which detect the disintegrated particles (e.g. pions, muons) produced when cosmic radiation impacts atmospheric particles. Figure 46b presents a time series of the neutron count rate at two surface stations. The lower latitude station clearly shows lower counts and a weaker solar cycle modulation.
The air ions produced by the cosmic rays may act as sites for the nucleation of new ultrafine aerosol (or condensation nuclei, CN). The mechanism then hinges on the extent to which these CN may grow into particles large enough (> 80 nm) to become cloud condensation nuclei (CCN) and whether this process is enhanced by the particles being charged. Growth may occur through condensation of water vapour or other soluble gases or through coagulation among neutral and charged particles. Some observational evidence (Yu and Turco, 2000) suggests that charged molecular clusters grow faster than neutral clusters and chemical box models (Yu and Turco, 2001) have been able to simulate this effect. However, to reach CCN size would take several days and whether the growth can be maintained depends on the supply of vapour and competitive sources of new aerosol and CCN all of which vary with location, altitude and time of year. Yu (2002) suggests that conditions in the lower troposphere may be more favourable than at higher altitudes but Arnold (2006) concludes that such processes are most likely to occur in the upper troposphere, providing there is a sufficient supply of SO2.
Even if by such a mechanism it proves feasible to produce a measurable effect on cloud cover or properties, the magnitude, and even the sign, of the impact on radiative forcing remains uncertain as it will depend on the cloud location, altitude and physical properties.
Radiation from the Sun ultimately provides the only energy source for the Earth’s atmosphere and changes in solar activity clearly have the potential to affect climate. There is statistical evidence for solar influence on various meteorological parameters on all timescales, although extracting the signal from the noise in a naturally highly variable system remains a key problem. Changes in total solar irradiance undoubtedly impact the Earth’s energy balance but uncertainties in the historical record of TSI mean that the magnitude of even this direct influence is not well known. Variations in solar UV radiation impact the thermal structure and composition of the middle atmosphere but details of the responses in both temperature and ozone concentrations are not well established. Various theories are now being developed for coupling mechanisms whereby direct solar impacts on the middle atmosphere might influence the troposphere but the influences are complex and non-linear and many questions remain concerning the detailed mechanisms which determine to what extent, where and when the solar influence is felt. Variations in cosmic radiation, modulated by solar activity, are manifest in changes in atmospheric ionisation but it is not yet clear whether these have the potential to significantly affect the atmosphere in a way that will impact climate.
Further advances in this field require work on a number of fronts. One important issue is to establish the magnitude of any secular trends in total solar irradiance (TSI). This may be achieved by careful analysis and understanding of the satellite instruments involved in collecting data over the past two-and-a-half solar cycles, and must be continued through analysis of data from current and new satellites. For longer periods it requires a more fundamental understanding of how solar magnetic activity relates to TSI. This would not only facilitate more reliable centennial-scale reconstructions of TSI, from e.g. sunspot records, but also advance understanding of how cosmogenic isotope records may be interpreted as historical TSI.
With regard to the climate, further data-mining and analysis are required to firmly establish the magnitude, geographical distribution and seasonality of its response to various forms of solar activity. Understanding the mechanisms involved in the response then becomes the overriding objective. Current ideas suggest three main avenues where further research is needed. Firstly, the means whereby solar radiative heating of the upper and middle atmosphere may influence the lower atmosphere through dynamical coupling needs to be better understood. Secondly, it needs to be established whether or not variations in direct solar heating of the tropical oceans can be of sufficient magnitude to produce apparently observed effects. Thirdly, more work is needed on the microphysical processes involved in ion-induced nucleation, and, probably more importantly, the growth rates of the condensation nuclei produced.
Perhaps when these questions are answered we will be confident that we really understand how changes in the Sun affect the climate on Earth.
- Andrews, D.G., 2000, An Introduction to Atmospheric Physics, Cambridge University Press, Cambridge, U.K.; New York, U.S.A.Google Scholar
- Austin, J. et al., 2007, “Coupled chemistry climate simulations of the solar cycle in stratospheric ozone and temperature”, J. Geophys. Res., submittedGoogle Scholar
- Boberg, F., Lundstedt, H., 2002, “Solar Wind Variations Related to Fluctuations of the North Atlantic Oscillation”, Geophys. Res. Lett., 29, 1718. Related online version (cited on 06 September 2007): http://sunspot.lund.irf.se/. ADS: http://adsabs.harvard.edu/abs/2002GeoRL..29o..13BADSGoogle Scholar
- Bond, G., Kromer, B., Beer, J., Muscheler, R., Evans, M.N., Showers, W., Hoffmann, S., Lotti-Bond, R., Hajdas, I., Bonani, G., 2001, “Persistent Solar Influence on North Atlantic Climate During the Holocene”, Science, 294, 2130–2136. ADS: http://adsabs.harvard.edu/abs/2001Sci...294.2130BADSGoogle Scholar
- Burroughs, W.J., 1992, Weather Cycles: Real or Imaginary, Cambridge University Press, Cambridge, U.K.; New York, U.S.A.Google Scholar
- Egorova, T., Rozanov, E., Manzini, E., Haberreiter, M., Schmutz, W., Zubov, V., Peter, T., 2004, “Chemical and dynamical response to the 11-year variability of the solar irradiance simulated with a chemistry climate model”, Geophys. Res. Lett., 31, L06 119. ADS: http://adsabs.harvard.edu/abs/2004GeoRL..3106119EGoogle Scholar
- Foster, S., 2004, Reconstruction of solar irradiance variations, for use in studies of global climate change: application of recent SoHO observations with historic data from the Greenwich Observatory, Ph.D. Thesis, University of Southampton, Southampton, U.K. ADS: http://adsabs.harvard.edu/abs/2004PhDT....... 5FGoogle Scholar
- Gray, L.J., 2003, “The influence of the equatorial upper stratosphere on sudden stratospheric warmings”, Geophys. Res. Lett., 30. ADS: http://adsabs.harvard.edu/abs/2003GeoRL..30d..15G
- Gray, L.J., Phipps, S.J., Dunkerton, T.J., Baldwin, M.P., Drysdale, E.F., Allen, M.R., 2001, “A data study of the influence of the equatorial upper stratosphere on northern-hemisphere stratospheric sudden warmings”, Quart. J. R. Meteorol. Soc., 127, 1985–2003. ADS: http://adsabs.harvard.edu/abs/2001QJRMS.127.1985GADSGoogle Scholar
- Gray, L.J., Haigh, J.D., Harrison, R.G., 2005, The Influence of Solar Changes on the Earth’s Climate, Hadley Centre technical note 62, Met Office, Exeter, U.K. URL (cited on 03 September 2007): http://www.metoffice.gov.uk/research/hadleycentre/pubs/HCTN/index.htmlGoogle Scholar
- Haigh, J.D., Roscoe, H.K., 2006, “Solar influences on polar modes of variability”, Meteorol. Z., 15, 371–378Google Scholar
- Haigh, J.D., Austin, J., Butchart, N., Chanin, M.-L., Crooks, S., Gray, L.J., Halenka, T., Hampson, J., Hood, L.L., Isaksen, I.S.A., Keckhut, P., Labitzke, K., Langematz, U., Matthes, K., Palmer, M., Rognerud, B., Tourpali, K., Zerefos, C., 2004, “Solar variability and climate: selected results from the SOLICE project”, SPARC Newsletter, 23, 19–29. URL (cited on 03 September 2007): http://www.atmosp.physics.utoronto.ca/SPARC/News23/23_Haigh.htmlGoogle Scholar
- Houghton, J.T., 1977, The Physics of Atmospheres, Cambridge University Press, Cambridge, U.K.; New York, U.S.A.Google Scholar
- IPCC, 2001, Climate Change 2001: The Scientific Basis. Contribution of Working Group I to the Third Assessment Report of the Intergovernmental Panel on Climate Change, Cambridge University Press, Cambridge, U.K.; New York, U.S.A. Related online version (cited on 03 September 2007): http://www.grida.no/climate/ipcc_tar/wg1/index.htm. (Eds.) Houghton, J.T. and Ding, Y. and Griggs, D.J. and Noguer, M. and van der Linden, P.J. and Dai, X. and Maskell, K. and Johnson, C.A.Google Scholar
- IPCC, 2007, Climate Change 2007: The Physical Science Basis. Contribution of Working Group I to the Fourth Assessment Report of the Intergovernmental Panel on Climate Change, Cambridge University Press, Cambridge, U.K.; New York, U.S.A. Related online version (cited on 03 September 2007): http://ipcc-wg1.ucar.edu/wg1/wg1-report.html. (Eds.) Solomon, S. and Qin, D. and Manning, M. and Chen, Y. and Marquis, M. and Averyt, K.B. and Tignor, M. and Miller, H.L.Google Scholar
- Kiehl, J.T., Trenberth, K.E., 1997, “Earth’s Annual Global Mean Energy Budget”, Bull. Am. Meteorol. Soc., 78, 197–208. Related online version (cited on 06 September 2007): http://www.cgd.ucar.edu/cas/papers/KiehlTrenbBAMS97.pdfADSGoogle Scholar
- Larkin, A., 2000, Investigation into the effects of solar variability on climate using atmospheric models of the troposphere and stratosphere, Ph.D. Thesis, University of London, London, U.K.Google Scholar
- Ludlam, F.H., 1980, Clouds and Storms: The Behavior and Effect of Water in the Atmosphere, Pennylvania State University Press, University Park, U.S.A.Google Scholar
- Matthes, K., Kodera, K., Haigh, J.D., Shindell, D.T., Shibata, K., Langematz, U., Rozanov, E., Kuroda, Y., 2003, “GRIPS Solar Experiments Intercomparison Project: Initial Results”, Pap. Met. Geophys., 54, 71–90Google Scholar
- Matthes, K., Langematz, U., Gray, L.J., Kodera, K., Labitzke, K., 2004, “Realistic solar signal in the FUB-CMAM”, J. Geophys. Res., 109Google Scholar
- Myhre, G., Stordal, F., Rognerud, B., Isaksen, I.S.A., 1998, “Radiative forcing due to stratospheric ozone”, in Atmospheric Ozone, (Eds.) Bojkov, R.D., Visconti, G., Proceedings of the XVIII Quadrennial Ozone Symposium, L’Aquila, Italy, September 12–21, 1996, pp. 813–816, Edigrafital S.P.A. / Parco Scientifico e Tecnologico d’Abruzzo, Sant’Atto, ItalyGoogle Scholar
- Petit, J.R., Jouzel, J., Raynaud, D., Barkov, N.I., Barnola, J.M., Basile, I., Bender, M., Chappellaz, J., Davis, M., Delaygue, G., Delmotte, M., Kotlyakov, V.M., Legrand, M., Lipenkov, V.Y., Lorius, C., Pepin, L., Ritz, C., Saltzman, E., Stievenard, M., 1999, “Climate and atmospheric history of the past 420,000 years from the Vostok ice core, Antarctica”, Nature, 399, 429–436ADSGoogle Scholar
- Randel, W.J., Wu, F., 2007, “A stratospheric ozone profile data set for 1979–2005: Variability, trends, and comparisons with column ozone data”, J. Geophys. Res., 112. ADS: http://adsabs.harvard.edu/abs/2007JGRD..11206313R
- Shindell, D.T., Faluvegi, G., Miller, R.L., Schmidt, G.A., Hansen, J.E., Sun, S., 2006, “Solar and anthropogenic forcing of tropical hydrology”, Geophys. Res. Lett., 33. ADS: http://adsabs.harvard.edu/abs/2006GeoRL..3324706S
- Tourpali, K., Schuurmans, C.J.E., van Dorland, R., Steil, B., Brühl, C., 2003, “Stratospheric and tropospheric response to enhanced solar UV radiation: A model study”, Geophys. Res. Lett., 30, 1231. ADS: http://adsabs.harvard.edu/abs/2003GeoRL..30e..35T 37ADSGoogle Scholar
- van Loon, H., Meehl, G.A., Shea, D.J., 2007, “Coupled air-sea response to solar forcing in the Pacific region during northern winter”, J. Geophys. Res., 112, D02108 9Google Scholar
- Vlachogiannis, D., Haigh, J.D., 1998, “The impact of solar proton events on lower stratospheric ozone”, in Atmospheric Ozone, (Eds.) Bojkov, R.D., Visconti, G., Proceedings of the XVIII Quadrennial Ozone Symposium, L’Aquila, Italy, September 12–21, 1996, pp. 275–278, Edigrafital S.P.A. / Parco Scientifico e Tecnologico d’Abruzzo, Sant’Atto, ItalyGoogle Scholar
- WMO, 2007, Scientific Assessment of Ozone Depletion: 2006, Global Ozone Monitoring Project — Report No. 50, World Meteorological Organization, Geneva, Switzerland. URL (cited on 03 September 2007): http://www.mo.int/pages/prog/arep/gaw/ozone_2006/ozone_asst_report.htmlGoogle Scholar
- Yu, F., 2002, “Altitude variations of cosmic ray induced production of aerosols: Implications for global cloudiness and climate”, J. Geophys. Res., 107. ADS: http://adsabs.harvard.edu/abs/2002JGRA.107g.SIA8Y
Open AccessThis article is distributed under the terms of the Creative Commons Attribution 4.0 International License (https://creativecommons.org/licenses/by/4.0), which permits use, duplication, adaptation, distribution, and reproduction in any medium or format, as long as you give appropriate credit to the original author(s) and the source, provide a link to the Creative Commons license, and indicate if changes were made.