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Tristan S. L’Ecuyer and Greg McGarragh

System (CERES) clouds and radiative swath (CRS) product ( Wielicki et al. 1996 ) offers estimates of Q R that are constrained to match top of the atmosphere (TOA) flux measurements but with reduced temporal sampling, whereas Cloudsat’s level-2B radiative fluxes and heating rates algorithm (2B-FLXHR; L’Ecuyer et al. 2008 ) offers improved cloud boundary information and spatial resolution but at greatly reduced spatial and temporal sampling. All of these algorithms are built on the same basic

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Wei-Kuo Tao, Stephen Lang, Xiping Zeng, Shoichi Shige, and Yukari Takayabu

dominated by phase changes between water vapor and small liquid or frozen cloud-sized particles. It consists of the condensation of cloud droplets, evaporation of cloud droplets and raindrops, freezing of cloud droplets and raindrops, melting of snow and graupel/hail, and the deposition and sublimation of ice particles. In addition, eddy heat flux convergence from cloud motions can also redistribute the heating or cooling vertically and horizontally. LH cannot be measured directly with current

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T. N. Krishnamurti, Arindam Chakraborty, and A. K. Mishra

. The two diagrams show, respectively, the predicted maxima of heating distributions for hours 12–36 and hours 36–60. The units of heating are here expressed in K day −1 . The first 12 hours are not included in the analysis presented here; for reasons of the initial spinup within the three models that utilize different convection schemes, those are different from the Tiedtke (1989) mass flux scheme, which was implicit in the data assimilation that defines our initial states and came from the

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Yasu-Masa Kodama, Masaki Katsumata, Shuichi Mori, Sinsuke Satoh, Yuki Hirose, and Hiroaki Ueda


The large-scale distribution of precipitation and latent heating (LH) profiles in the tropics, subtropics, and part of the midlatitudes was studied using a 9-yr dataset derived from Tropical Rainfall Measuring Mission precipitation radar observations, with emphasis on the contribution of warm rain. The distribution of warm rain showed features unique from those of rain in other categories and those of outgoing longwave radiation. Warm rain was weak over land but widely distributed over oceans, especially along the intertropical convergence zone (ITCZ) and the western part of the subtropical oceans. The observed amount of warm rain depended on the rainfall intensity rather than on the frequency of warm rain events. The amount of warm rain over ocean was positively correlated with sea surface temperature (SST); this dependency was found in the tropics, subtropics, and part of the midlatitudes, whereas dependency of SST on total rain was confined to the tropics. Both total rain and warm rain were concentrated in the ITCZ, which elongated along the local SST maximum. Small amounts of warm rain were found along subtropical convergence zones (the baiu frontal zone and subtropical portions of the South Pacific convergence zone and the South Atlantic convergence zone) with ample total rainfall. However, larger amounts of warm rain were observed at the lower-latitude sides of these zones in the upstream portions of low-level moisture flow toward the zones. Warm rain may cultivate the subtropical convergence zones by deepening the moist boundary layer and increasing moisture flux toward the zones. The statistical relationship between warm rain and low-level cloudiness showed that the warm rain amount was large when low-level cloudiness was 20%–30% and small when low-level cloudiness was greater than 40%. This indicates that intense warm rain is provided by convective clouds, not by stratiform clouds, in conditions of substantial cloudiness. Despite the small contribution to total rain, warm rain maintained positive LH values over most of the tropical and subtropical oceans. The LH by warm rain masked low-level cooling observed in stratiform rain and maintained positive LH in the lower atmosphere below the melting layer. Because warm rain was confined to oceans, a strong LH contrast was maintained along the coast; this contrast reached values of 1–2 K day−1 in certain places and may affect local and monsoonal circulation across continental coasts.

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Shaocheng Xie, Timothy Hume, Christian Jakob, Stephen A. Klein, Renata B. McCoy, and Minghua Zhang

top-of-the-atmosphere (TOA) observations as constraints to adjust atmospheric state variables from soundings by the smallest possible amount to conserve column-integrated mass, moisture, and static energy so that the final analysis dataset is dynamically and thermodynamically consistent. The required observation constraints include the surface and TOA radiative fluxes, surface latent and sensible heat fluxes, and surface precipitation. The variational analysis has been successfully used to process

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Mircea Grecu, William S. Olson, Chung-Lin Shie, Tristan S. L’Ecuyer, and Wei-Kuo Tao

from these different perspectives, several methods for estimating latent heating from satellite observations have been developed. Tao et al. (1990) and Smith et al. (1994) used satellite estimates of precipitation vertical structure to infer latent heating rates in discrete atmospheric layers, assuming that the net flux of precipitation into or out of a given layer is balanced by microphysical processes under steady-state conditions. Tao et al. (1993) later simplified their approach by

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Shoichi Shige, Yukari N. Takayabu, Satoshi Kida, Wei-Kuo Tao, Xiping Zeng, Chie Yokoyama, and Tristan L’Ecuyer

crystals, snow, and graupel; deposition of ice crystals; and sublimation of all ice hydrometeors, respectively. The first term on the right-hand side of Eq. (2) is the vertical eddy heat flux convergence from upward and downward cloud-scale motions, while the second term is the horizontal eddy heat flux convergence. The precipitation falling at a given time is not related to the heating/cooling that is occurring at that instant but rather to the accumulated heating/cooling that led up to the

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Richard H. Johnson, Paul E. Ciesielski, Tristan S. L’Ecuyer, and Andrew J. Newman

effects of mesoscale convective systems ( Chen and Houze 1997 ). Over land, daytime heating exerts the primary control on the diurnal cycle of precipitation; however, factors influencing the development and organization of convection—surface fluxes, surface heterogeneity, low-level jets, orography, convective downdrafts, etc.—are varied and complex, complicating its treatment in models ( Betts and Jakob 2002 ; Bechtold et al. 2004 : Khairoutdinov and Randall 2006 ). Observations prior to NAME have

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Xianan Jiang, Duane E. Waliser, William S. Olson, Wei-Kuo Tao, Tristan S. L’Ecuyer, Jui-Lin Li, Baijun Tian, Yuk L. Yung, Adrian M. Tompkins, Stephen E. Lang, and Mircea Grecu

sublimation, respectively. Term I in Eq. (1) represents the latent heat due to phase changes, and term II is the vertical and horizontal eddy sensible heat flux convergence. It is noted that both TRMM/TRAIN Q 1 − Q R and TRMM/CSH Q 1 are not estimated under conditions of zero surface rainfall. As a result, the radiative cooling effect during nonrainy days could be largely underestimated in both TRMM heating estimates. Therefore, in order to facilitate a more direct comparison between these two

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Samson Hagos, Chidong Zhang, Wei-Kuo Tao, Steve Lang, Yukari N. Takayabu, Shoichi Shige, Masaki Katsumata, Bill Olson, and Tristan L’Ecuyer

heat fluxes from the surface to the atmosphere are influenced by the ambient surface winds as well as temperature and moisture distributions. Hence, the three-dimensional structure of the diabatic heating is closely related to the atmospheric circulation because it not only drives the circulation, but also receives feedback from it. This is particularly true for the diabatic heating associated with tropical precipitation, which on the one hand is a result of instability due to the accumulation of

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