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Xiaofang Feng, Qinghua Ding, Liguang Wu, Charles Jones, Ian Baxter, Robert Tardif, Samantha Stevenson, Julien Emile-Geay, Jonathan Mitchell, Leila M. V. Carvalho, Huijun Wang, and Eric J. Steig

is commonly referred to as the warm Arctic–cold Eurasia (WACE) pattern ( Cohen et al. 2012 ; Mori et al. 2014 ). However, there are diverse opinions and a lack of consensus about the ultimate driver of the WACE and how the WACE interacts with Arctic warming, extratropical circulation, and tropical forcing ( Mori et al. 2014 ; Overland et al. 2015 ; Sun et al. 2016 ). Although many theories have been proposed to explain individual regional climate anomalies in the past decades, an integrated

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N. Fauchereau, B. Pohl, and A. Lorrey

lagged response of these KTs to successive MJO phases is thus about 5 days. This is consistent with the average propagation speed of the MJO signal itself around the equatorial waveguide, and thus the average life cycle of an MJO event (30–60 days). This finding suggests that MJO effects on NZ regional circulation are not only through synchronous modulations of atmospheric fluxes that interact with the topography and act to modulate associated rainfall patterns, but also through delayed responses of

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Ryan L. Fogt and Alex J. Wovrosh

1. Introduction Recent studies note marked regional changes in both Antarctic sea ice extent/concentration and near-surface temperatures across the Antarctic continent. In terms of sea ice, the Ross Sea sector has been displaying increases in sea ice extent ( Lefebvre et al. 2004 ; Comiso and Nishio 2008 ), while the neighboring Amundsen–Bellingshausen (AB) Seas sector has experienced decreasing sea ice extent as well as concentration ( Yuan and Martinson 2000 ; Zwally et al. 2002

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Xiaojun Yuan, Michael R. Kaplan, and Mark A. Cane

temperature anomalies over eastern Antarctic coastal areas. ENSO-induced circulation anomalies in high latitudes inevitably interact with regional climate modes, such as the SAM and the wavenumber-3 (wave-3) pattern, and interfere with the natural variability of these modes. These interactions actively modulate ENSO’s influence on sea ice and surface climate in Antarctica ( Stammerjohn et al. 2008 ; Yuan and Li 2008 ; Fogt et al. 2011 ; Ding et al. 2012 ; Clem and Fogt 2013 ; Yu et al. 2015 ; Wilson

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Kyle R. Clem, James A. Renwick, and James McGregor

and West Antarctic warming are both linked to regional circulation changes associated with teleconnections stemming from the tropical Pacific and Atlantic and related reductions in sea ice concentration in the Amundsen and Bellingshausen Seas. Autumn warming of the western peninsula has been linked to a deepening of the Amundsen Sea low (ASL; Raphael et al. 2016 ; Turner et al. 2013a ; Fogt et al. 2012a ) and associated reductions in sea ice concentration along the western peninsula coast that

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Kyle R. Clem and James A. Renwick

austral spring [September–November (SON)]. Statistically significant warming is also found in observations across the western Antarctic Peninsula during SON ( Turner et al. 2005 ), and after 1979 SON is found to be the only season with widespread significant warming across both West Antarctica and the Antarctic Peninsula ( Schneider et al. 2012 ; Bromwich et al. 2013 ; Nicolas and Bromwich 2014 ; Clem and Fogt 2015 ). Consistent with the SON warming are significant changes in regional atmospheric

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Bradley P. Goodwin, Ellen Mosley-Thompson, Aaron B. Wilson, Stacy E. Porter, and M. Roxana Sierra-Hernandez

1. Introduction The Antarctic Peninsula (AP) is a climatologically complex region that includes ice-free ocean, sea ice, land ice, and significant topographic relief within a relatively small area ( Fig. 1 ). Air temperatures have increased, particularly along the west coast since the 1950s (e.g., ~2.5°C; King 1994 ), reflecting one of the strongest positive regional trends recorded globally ( Marshall et al. 2002 ; Turner et al. 2005 ). Rapid warming observed over the AP has been associated

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Robert A. Tomas, Clara Deser, and Lantao Sun

; Sun et al. 2015 ) and an increase in warm extremes ( Screen et al. 2015b ) as a result of Arctic sea ice loss. In addition to local thermodynamic effects, diminished Arctic sea ice cover will weaken the tropospheric westerly winds along the poleward flank of the jet stream in association with a reduced north–south temperature gradient due to enhanced lower-tropospheric warming in the Arctic ( Deser et al. 2010 ; Peings and Magnusdottir 2014 ; Deser et al. 2015 , hereafter D15 ; Harvey et al

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Aaron B. Wilson, David H. Bromwich, and Keith M. Hines

-century trends toward a high-polarity SAM have been linked to decreases in stratospheric ozone over Antarctica with propagating effects into the troposphere during austral summer (e.g., Thompson and Solomon 2002 ; Gillett and Thompson 2003 ; Thompson et al. 2011 ), as well as increasing greenhouse gases ( Fyfe et al. 1999 ; Kushner et al. 2001 ; Marshall et al. 2004 ; Simpkins and Karpechko 2012 ; Zheng et al. 2013 ). Low-frequency forcing of the atmospheric circulation in the SH is also tied to the

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Ariaan Purich, Matthew H. England, Wenju Cai, Yoshimitsu Chikamoto, Axel Timmermann, John C. Fyfe, Leela Frankcombe, Gerald A. Meehl, and Julie M. Arblaster

. Based on this, the assimilation runs have the same climatology and externally forced component as the historical runs. Therefore, the difference between assimilation runs and historical runs is the first order representation of internal climate variability. Four CESM1 ensemble experiments with a partial assimilation approach in global or regional oceans similar to Chikamoto et al. (2015b) are conducted for the time period 1979–2013, with 10 ensemble members each: historical scenario (CESM1-HIST

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