1. Introduction
Improving our understanding of the dynamics of the atmosphere–ocean–sea ice system and the connecting mechanisms between the high and low latitudes has become increasingly important to climate science in the face of a rapidly warming world. The polar regions and the cryosphere in both hemispheres are active components in global climate. For example, changes within the polar regions dictate the strength of the thermal gradient between the tropics and the poles.
Climate changes have been observed in both of Earth’s polar regions over the past several decades and have been magnified in the last ~10 years in terms of recorded sea ice extent (SIE) and surface temperatures. For example, in September 2012 Arctic sea ice extent set a new record low in more than three decades of satellite observations, while Antarctic sea ice extent reached a record high for the same period, reported by the National Snow and Ice Data Center (http://nsidc.org/cryosphere/sotc/sea_ice.html). While September Arctic SIE has slightly recovered since 2013, Antarctic SIE continued to reach record highs until 2014, followed by a drop in 2016 and 2017.
Near-surface temperatures in the Arctic region show a warming of ~0.7 K decade−1 (Cavalieri and Parkinson 2012), which reflects primarily an ice-albedo feedback. Since Antarctic sea ice mostly melts back to the coast in summer when the sun returns to the Southern Hemisphere (SH) and albedo is highly correlated with sea ice concentration (SIC; Shao and Ke 2015), the ice-albedo feedback is not as important as it is in northern high latitudes. There is no consistent sea surface temperature (SST) warming. Instead, surface temperatures show a strong spatial contrast, with rapid warming (~0.35°C decade−1) around the Antarctic Peninsula (AP) and a net cooling in East Antarctica (Chapman and Walsh 2007; O’Donnell et al. 2011; Turner 2016). Studies had attributed the Antarctic cooling and sea ice expansion to ozone depletion (Thompson and Solomon 2002; Turner et al. 2009). However, recent work has shown that the response to ozone depletion should be less sea ice, while the opposite is observed (Bitz and Polvani 2012). Moreover, most climate models produce declined Antarctic sea ice in the recent decades (Shu et al. 2015). On the other hand, more recent studies suggest that decadal changes in the Antarctic Peninsular surface climate may be linked to the changes in tropical SSTs (Ding and Steig 2013; Li et al. 2014). These results refocus attention on the connections from the tropics to the poles.
It has been long known that polar climate is strongly influenced by the tropical SST variability related particularly to the El Niño–Southern Oscillation (ENSO) phenomenon, owning to its far-reaching impact on many aspects of the global climate system. The influences of ENSO can be observed in the atmosphere, ocean, sea ice, and glacial ice in polar regions of both hemispheres (Turner 2004). The atmospheric responses at southern high latitudes include anomalies in sea level pressure, troposphere height, streamfunction, blocking events, precipitation, surface winds, and temperature (van Loon and Shea 1985; Mo and White 1985; Krishnamurti, et al. 1986; Trenberth and Shea 1987; Karoly 1989; Smith and Stearns 1993; Cullather et al. 1996; Sinclair et al. 1997; Villalba et al. 1997; Kiladis and Mo 1998; Marshall et al. 1998; Noone et al. 1999; Bromwich et al. 2000; Renwick 1998; Liu et al. 2002). Similarly, the atmospheric responses can be traced in the Northern Hemisphere (NH) extratropics (Trenberth et al. 1998; Pozo-Vázquez et al. 2005; Cassou and Terray 2001; Jevrejeva et al. 2003). The mechanisms fostering the connection between ENSO and the high latitudes through the troposphere include 1) the Rossby waves generated by tropical convection (Karoly 1989; Mo and Higgins 1998; Kiladis and Mo 1998; Garreaud and Battisti 1999); 2) jet stream changes in response to tropical SST changes (Chen et al. 1996; Bals-Elsholz et al. 2001); 3) anomalous mean meridional and zonal circulations and associated heat fluxes (Carleton and Whalley 1988; Cullather et al. 1996; Kreutz et al. 2000; Liu et al. 2002; Seager et al. 2003; Liu et al. 2004; Yuan 2004); and 4) altered transient eddy activity (Carleton and Carpenter 1990; Carleton and Fitch 1993; Sinclair et al. 1997; Carleton and Song 2000). Through the atmospheric connection, ENSO events significantly influence sea ice variability in the Antarctic (Simmonds and Jacka 1995; Yuan and Martinson 2000, 2001; Harangozo 2000; Kwok and Comiso 2002; Martinson and Iannuzzi 2003) and in the Arctic (Gloersen 1995; Loewe and Koslowski 1998; Venegas and Mysak 2000; Jevrejeva et al. 2003).
Studies of the tropical–polar teleconnection have advanced rapidly in the recent decade since earlier reviews on the subject (Trenberth et al. 1998; Turner 2004), given the accumulation of polar cryosphere observations, enhanced atmospheric data assimilation, and improved climate models. Emerging studies suggest that tropical climate variability at other (“non ENSO”) time scales also reaches the polar regions. For example, tropical variability on an intraseasonal time scale, namely, the Madden–Julian oscillation (MJO), impacts extratropical regions as far as the high-latitude Arctic (Yoo et al. 2011). At multidecadal (or long-term trend) time scales, tropical Pacific SST variability and associated Rossby wave trains influence the southern annular mode (SAM) and surface temperature around the AP and West Antarctica (Ding et al. 2011, 2012; Ding and Steig 2013; Schneider et al. 2012a,b; Clem and Fogt 2013, 2015; Clem and Renwick 2015; Yu et al. 2015). The tropical forcing that influences the high latitudes has been also found in the equatorial Atlantic (Li et al. 2014; Simpkins et al. 2014) and Indian Ocean (Nuncio and Yuan 2015). In addition to the connective mechanisms through the troposphere, teleconnections can also take stratospheric pathways (Butler and Polvani 2011). Here, we summarize some of the recent studies, including those mentioned above, which are organized by time scale from intraseasonal, interannual, decadal to multidecadal, and centennial and longer, including paleoclimate studies. In summary, this review tries to follow an overarching theme: that the problems and research findings of recent and past climates can often be appreciated better within the context of tropical–polar linkages. Also, studies of tropical–polar connections based on instrumental records can greatly help paleoclimate researchers better understand inferred processes in the past.
2. Connections at intraseasonal time scales
Tropical–polar teleconnections at intraseasonal time scales were discovered in the last decade. The most profound tropical intraseasonal variability is the MJO, which can modulate high-latitude atmospheric circulation (Matthews and Meredith 2004; Zhou and Miller 2005; Cassou 2008; L’Heureux and Higgins 2008; Flatau and Kim 2013). Based on reanalysis data, satellite observations, polar bottom pressure records, and tide gauge station data, Matthews and Meredith (2004) found that the SAM peaks approximately seven days after the MJO. Three days after this atmospheric response to the MJO, an increase in the ocean transport through Drake Passage was found. This associated oceanic intraseasonal variability is well captured by the Ocean Circulation and Climate Advanced Modelling (OCCAM) global ocean model (Webb and de Cuevas 2002), which indicates that up to 15% of oceanographic intraseasonal variance can be linearly attributed to the MJO (Matthews and Meredith 2004). On the other hand, Pohl et al. (2010) claimed that the MJO does not have a significant impact on SAM. Contrary to the conclusion of Pohl et al. (2010), Flatau and Kim (2013) found that the MJO is capable of forcing the leading modes of mid–high-latitude atmospheric circulation, namely, the northern annular mode (NAM) and SAM in both hemispheric winter and summer. They showed that Indian (Pacific) Ocean convection related to the MJO precedes the increase (decrease) in the NAM and SAM indices during the respective hemispheric cold seasons. During the hemispheric warm seasons, NAM has a similar response to that in the cold season, but SAM has the opposite response (Fig. 1), although the observational uncertainties are larger for the warm season.
The average change (tendency) of the index over 1 day for the (a) NAM index and (b) SAM index (gpm day−1). The cool colors denote the changes for the winter cold season, and the warm colors denote the changes for the warm season. The minimum MJO amplitudes considered in the calculation of each average tendency are indicated (Flatau and Kim 2013). MJO phases 1–3 occur in the Indian Ocean and phases 4–8 occur in the Pacific.
Citation: Journal of Climate 31, 15; 10.1175/JCLI-D-16-0637.1
The impact of the MJO on polar atmospheric circulation extends to the surface climate, such as surface temperature. S. Lee et al. (2011) investigated the connection of Arctic temperature variability to MJO using ERA-40 reanalysis data. They used the coupled self-organizing map (SOM) between the NH 250-hPa streamfunction and tropical convective precipitation to detect dynamic links between the low and high latitudes. They found that the tropical convection associated with MJO causes extratropical circulation changes through atmospheric Rossby wave propagation, and the circulation anomalies can be established at high latitudes in 3–6 days as the Pacific–North America (PNA) pattern. These MJO-induced circulation changes alter poleward heat transport. Together with adiabatic warming and downward infrared radiative fluxes, the anomalous poleward heat transport is capable of influencing the variability of winter surface temperature in the Arctic. This result is consistent with the results of Graversen (2006). Yoo et al. (2011) further found in observations that MJO heating in the tropical Pacific is followed 1–2 weeks later by Arctic warming through atmospheric circulation changes [similar to the finding by S. Lee et al. (2011)], whereas MJO heating in the equatorial Indian Ocean is followed by Arctic cooling (Fig. 2).
