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Kerry H. Cook, Gerald A. Meehl, and Julie M. Arblaster

near the Cameroon highlands (near 8°E and 5°N) and just off the west coast near 10°W and 5°N. The 1 mm day −1 precipitation contour lies near 13°N. The West African westerly jet ( Grodsky et al. 2003 ; Pu and Cook 2010 ) has not yet formed on the west coast at 10°N. Fig . 2. (a) MJ climatological precipitation (contours; mm day −1 ) from TRMM with 900-hPa wind vectors (m s −1 ) from the ERA-Interim reanalysis. (b) Precipitation contours (mm day −1 ) and 930-hPa wind vectors from the CCSM4

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Wilbert Weijer, Bernadette M. Sloyan, Mathew E. Maltrud, Nicole Jeffery, Matthew W. Hecht, Corinne A. Hartin, Erik van Sebille, Ilana Wainer, and Laura Landrum

century, despite the increasing trend in wind stress evident in Fig. 3b . This lack of response is partially explained by Gent and Danabasoglu (2011) , who show that the more sophisticated formulation of isopycnal tracer transport used in CCSM4 is consistent with the eddy saturation of circumpolar transport discussed by Hallberg and Gnanadesikan (2006) . d. Antarctic circumpolar current structure The ACC is highly filamented and comprises a number of meandering jets with meridional scales on the

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Semyon A. Grodsky, James A. Carton, Sumant Nigam, and Yuko M. Okumura

. The northward shift of the ITCZ also leads to a seasonal strengthening of the alongshore winds off southwest subtropical Africa. A low-level atmospheric jet along the Benguela coast is driven by the South Atlantic subtropical high pressure system, with topographic enhancement of winds west of the Namibian highland ( Nicholson 2010 ). This coastal wind jet drives local upwelling as well as the coastal branch of the equatorward Benguela Current, causing equatorward advection of cool Southern

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K. J. Evans, P. H. Lauritzen, S. K. Mishra, R. B. Neale, M. A. Taylor, and J. J. Tribbia

-SE show close agreement with CAM-FV compared to NCEP observations. We display the seasonally averaged zonal wind fields ( Figs. 12 and 13 ) because the polar winter jet stream governs many important aspects of global climate, including the meridional transport of heat, the frequency and location of storm tracks, and the global distribution of surface pressure and temperature. Figure 12 shows the zonally averaged zonal wind field averaged over December–February (DJF) for CAM-SE, CAM-FV, and NCEP

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Ernesto Muñoz, Wilbert Weijer, Semyon A. Grodsky, Susan C. Bates, and Ilana Wainer

Sea, the low-level (~925 hPa) zonal winds have a semiannual cycle with peaks in February and July ( Muñoz et al. 2008 ), given the seasonal changes in pressure and thermal wind. The Caribbean low-level jet affects precipitation in the Caribbean and Central America, and it also affects regional SST ( Small et al. 2007 ; Muñoz et al. 2008 ; Wang et al. 2008 ) through coastal upwelling. A significant aspect of the tropical Atlantic seasonal cycle is its warm pool. Warm pools have been defined as

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Christine A. Shields, David A. Bailey, Gokhan Danabasoglu, Markus Jochum, Jeffrey T. Kiehl, Samuel Levis, and Sungsu Park

) are −9, −7, −5, −4, −3, −2, −1, 0, 1, 2, 3, 4, 5, 7, and 9. Solid (dashed) lines are positive (negative) values. The T31X3_20C annually averaged zonal mean wind is shown in Fig. 11 . Note that in the Southern Hemisphere the jet is displaced too far equatorward, which is reflected in the bias in the zonal surface stress ( Fig. 3 ). The equatorward displacement of the jets is a signature of the low-resolution model. With increased horizontal resolution the Southern Hemisphere jet strengthens and

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Richard B. Neale, Jadwiga Richter, Sungsu Park, Peter H. Lauritzen, Stephen J. Vavrus, Philip J. Rasch, and Minghua Zhang

; ) and CAM chemistry (e.g., Lamarque et al. 2012 ). For CAM4 the default number of levels remains at 26 as a result of an undesirable nonconvergent response from boundary layer and shallow convection interactions when levels were significantly increased ( Williamson 2013 ). Because of the presence of grid-scale noise and excessive polar night jets, two new filtering/diffusion operators have been implemented in CAM4-FV ( Lauritzen et al. 2011 ). First

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Aneesh C. Subramanian, Markus Jochum, Arthur J. Miller, Raghu Murtugudde, Richard B. Neale, and Duane E. Waliser

and it reduces many of the model biases ( Neale et al. 2008 ). For example, the annual mean, the tropical easterly bias, subtropical westerly bias, and the excessive Southern Hemisphere midlatitude jet, seen in CCSM3, are improved in CCSM3.5, as shown by Zhou et al. (2012) . In combination, these modifications to the deep convection parameterization lead to significant improvements in the phase, amplitude, and spatial anomaly patterns of the modeled El Niño, also documented by Neale et al. (2008

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A. Gettelman, J. E. Kay, and J. T. Fasullo

LWP, along with smaller changes in NC860 and larger decreases in NC860 at high latitudes than simulations with low sensitivity. REL860 is not as coherent as cloud feedback or condensate correlations. The radiative effect of smaller changes in NC860 is to minimize any increase in SWCRE (reducing any cooling)—hence the contribution to increased sensitivity at these high latitudes poleward of the jet cores near 60°S/N. As noted however, correlation is not causation, and this might be a reaction to

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Gerald A. Meehl, Warren M. Washington, Julie M. Arblaster, Aixue Hu, Haiyan Teng, Jennifer E. Kay, Andrew Gettelman, David M. Lawrence, Benjamin M. Sanderson, and Warren G. Strand

in the geographical plots in Fig. 6 , this is more evident in zonal mean changes of temperature and winds discussed below with regards to Fig. 7 . With the greater tropical precipitation response in CESM1(CAM5), relatively higher pressure dominates in most subtropical regions corresponding to a greater expansion of the Hadley cell in CESM1(CAM5) (e.g., Meehl et al. 2007 ). This is seen in Figs. 7c and d as a poleward shift of the subtropical jets particularly in the Northern Hemisphere

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