Posted on May 16th, 2013 in Isaac Held's Blog
Schematic of the response of tropical rainfall to high latitude warming in one hemisphere and cooling in the other or, equivalently, to a cross-equatorial heat flux in the ocean. From Kang et al 2009.
When discussing the response of the distribution of precipitation around the world to increasing CO2 or other forcing agents, I think you can make the case for the following three basic ingredients:
- the tendency for regions in which there is moisture convergence to get wetter and regions in which there is moisture divergence to get drier (“wet get wetter and dry get drier”) in response to warming (due to increases in water vapor in the lower troposphere — post #13);
- the tendency for the subtropical dry zones and the mid-latitude storm tracks to move polewards with warming;
- the tendency for the tropical rainbelts to move towards the hemisphere that warms more.
There are other important elements we could add to this set, especially if one focuses on particular regions — for example, changes in ENSO variability would affect rainfall in the tropics and over North America in important ways . But I think a subset of these three basic ingredients, in some combination, are important nearly everywhere. I want to focus here on 3) the effect on tropical rain belts of changing interhemispheric gradients.
My exposure to this issue started some time ago listening to Suki Manabe discussing early coupled atmosphere-ocean model simulations in which the latitude of the Pacific intertropical convergence zone (ITCZ) was sensitive to the cloud cover in midlatitudes of the Southern Hemisphere — these were the days in which cloud cover was prescribed in the models so it was easy to manipulate. By increasing the cloud cover in the South, well away from the tropical rainbelts themselves, one could move the ITCZ from south of the equator to north of the equator (where it is in reality).
In the late 80’s a flurry of work on Sahel rainfall and particularly the severe drought in the preceding decade, starting with Folland et al 1986, argued that much of the decadal variability in the Sahel is tied to the differential warming of the hemispheres. Relatively cool Northern Hemisphere, as in the 70’s, results in less Sahel rainfall, thinking of the Sahel as marking the northernmost extension of the ITCZ, or monsoonal rainfall, over Africa, which retreats due to the pull of the differential cooling of the Northern with respect to the Southern Hemisphere. While there are other things going on in the Sahel, most recent research supports this picture of variations in interhemispheric temperature gradients, whether produced by variability in Atlantic overturning or aerosol forcing, as being a big part of the Sahel drought picture. I talk about this here, although that page is a bit dated.
John Chiang and collaborators have emphasized the importance of this mechanism for paleoclimate as well as higher frequency climate variations in a series of papers, see this recent review (Chiang and Friedman, 2012). Another paper that affected my own work on this issue was that of Broccoli et al 2006, which helped shift the picture of the underlying dynamics from one focused on surface energy balances and changes in tropical ocean temperatures to one focused on the requirements of atmospheric energy balance. Two former students of mine, Sarah Kang and Dargan Frierson, have picked up on this energy balance perspective in more recent work, starting with Kang et al 2008.
Sarah has focused on a setup in which one takes an atmospheric model of the type used for climate simulations and couples it to a “slab ocean” of uniform depth with no ocean currents, which just provides some heat capacity and a saturated surface. Starting with the case in which there is no heat flux through the bottom of the slab ocean, the resulting climate is independent of longitude and symmetric about the equator, with most tropical precipitation confined to a sharp ITCZ located over the equator (the solid line in the lower panel). The model determines its own surface temperature, and the energy flowing into the slab will be zero everywhere if you average long enough. (One of the nice things about this setup is that, unlike models in which surface temperatures are prescribed, you never have to worry about generating a double ITCZ.) Heat is then added poleward of 40N in one hemisphere and the same amount is removed from the other hemisphere (as pictured in the upper left panel.) . This is equivalent to prescribing a cross-equatorial heat flux in the ocean underneath the slab (upper right). No heat is being input or extracted equatorward of 40 degrees. After the model equilibrates, the ITCZ has moved into the warmed hemisphere. The larger the heating, the larger the displacement of the ITCZ. The lower panel shows the precipitation from simulations in which the peak in the imposed subpolar heating/cooling is 10, 20 and 40 W/m2.
One way of thinking about this is to focus on how the surface temperatures in the tropics are affected by the extratropical heat sources/sinks, assuming that the ITCZ will follow the warmest surface temperatures. But I prefer a perspective based on the atmospheric energy budget, as in the papers by Broccoli et al and Kang et al linked to above.
Before the system is disturbed, the northward heat flux F in the atmosphere is zero at the equator and has some slope in latitude as pictured below. The Hadley cells, symmetric about the equator, have poleward flow in the upper troposphere and equatorward flow near the surface, with rising motion mostly confined to the ITCZ. These cells transport energy in the direction of their upper tropospheric flow. In response to the high latitude heating and cooling, the atmosphere tries to resist the resulting interhemispheric asymmetry by transporting energy across the equator from the heated to the cooled hemisphere. In this setup, you can equivalently talk about how much of the prescribed oceanic flux is compensated by an atmospheric flux in the opposite direction. In the schematic at the top of the post, the fraction of the flux that is compensated is denoted by C. Putting aside how C is determined, we can estimate the new latitude of the “energy flux equator” where the atmospheric flux vanishes. (See sketch below.) If simple Hadley cells continue to dominate the horizontal energy fluxes in the tropics, with most of the rising motion in a sharp ITCZ, then the ITCZ will need to be close to this energy flux equator so that energy flows away from this latitude in both directions.
But how do you estimate C? Start by ignoring any responses in clouds. Part of the input of energy into the warmed hemisphere is balanced more or less locally by an increase in the energy radiated away to space and the rest is transported to low latitudes. I picture the transport as a diffusive process (post #37), with a diffusivity that weakens as one approaches the tropics, where the mean meridional circulation (the Hadley cell) takes over a lot of the energy transport. The subtropical heating/cooling by midlatitude storms creates a problem because the tropical atmosphere can’t sustain large horizontal temperature gradients. If the change in the net radiation at the top of the atmosphere is primarily a function of tropospheric temperature (ie if clouds don’t change), then the changes in this net radiation have to be very uniform with latitude across the tropics. So the key from this perspective is the extent to which the eddy diffusive-like fluxes in midlatitudes manage to extract or input energy into the subtropics of each hemisphere, which the circulation must then redistribute.
As discussed in the Kang et a papers linked to above, if we either fix clouds in the GCM or use an idealized moist GCM with no clouds, this degree of compensation at the equator, C, seems to be of the order of 25-40%, a value you can get from a simple diffusive model with the diffusivity tuned to the atmospheric fluxes in the control climate. If you use the standard AM2 model that was used in our contribution to the CMIP3/AR4 database, you get something like 80%, but this number can be changed by manipulating the closure scheme for moist convection. It’s not the convection per se that matters, but the effect of the convection scheme on the cloud feedbacks — the response of clouds to this extratropical heating/cooling perturbation and the effect of these changes on the top-of-atmosphere (TOA) balance. It is still the TOA that matters here, because the net surface fluxes are prescribed — one can only change the net atmospheric poleward fluxes if the TOA fluxes change. (It is this emphasis on the TOA fluxes that distinguishes this perspective from those focusing on surface temperatures.)
There are two distinct kinds of cloud feedbacks that come into play. First, there can be changes in clouds in the high latitude regions which are directly being heated or cooled. These modify the heating/cooling that the atmosphere feels, so they effectively renormalize the forcing. But in addition, once the tropical circulation is modified clouds in the tropics will react to these changes in circulation to alter the energy transports needed to homogenize the tropical temperatures. For example, the strength of the subsidence increases in the tropics and subtropics of the cooled hemisphere, which might result in an increase in low level cloudiness (due to the suppression of vertical mixing of vapor into the upper troposphere)– a positive feedback on the initial cooling. (The movement of the ITCZ would also directly generate changes in long and shortwave fluxes at the TOA, but these tend to cancel — the effects of shallow clouds are often dominant.) So this is a hard problem to get right quantitatively, as are all cloud-related issues it seems. But the qualitative effect is clear.
This mechanism is important when thinking about the tropics during past glacial periods, given the large cooling associated with Northern ice sheets. It is important for the response to aerosol forcing that preferentially cools the Northern Hemisphere. It is important for the response to variations in the Atlantic meridional overturning, which directly modifies the cross-equatorial ocean flux. And it can be important for understanding the mean climatology, as indicated by the reference above to Suki Manabe’s early experience with coupled atmosphere-ocean models (see also Hwang and Frierson, 2013 and Marshall et al 2013).
[The views expressed on this blog are in no sense official positions of the Geophysical Fluid Dynamics Laboratory, the National Oceanic and Atmospheric Administration, or the Department of Commerce.]