The Polar Marine Climate Revisited
by
Thomas J. Ballinger1, Thomas W. Schmidlin1, Daniel F. Steinhoff2
1
2
Department of Geography, Kent State University, Kent, Ohio
Research Applications Laboratory, National Center for Atmospheric Research, Boulder, CO
Planned Submission to Journal of Climate or Journal of Applied Meteorology and Climatology
May 2012
Abstract
As an additional classification to Köppen’s Climate Classification for polar (E) climates,
the Polar Marine (EM) climate was presented nearly five decades ago and is revisited in this
paper. The EM climate was traced to the North Atlantic, North Pacific and Southern Polar
Ocean and recognized as fairly wet, cloudy and windy especially in the respective winter
seasons. These areas are encompassed by coldest monthly mean temperatures of -6.7°C (20°F)
and warmest monthly mean temperatures of 10°C (50°F). Since the initial analysis was
performed, data availability has improved, and climate variability and change over polar regions
are better understood. Here we use three global reanalyses (ERA-Interim, CFSR and JRA-25) to
produce a modern depiction of EM climate. General agreement is found between original and
new EM boundaries. The poleward boundary is approximated by the climatological coldestmonth sea ice maximum and the equatorward boundary is approximated by warmest-month
SSTs. Additional variables are analyzed to gain a better understanding of regional mechanisms
that also play a role in formulating these boundaries. Interannual variability reveals Northern
Hemisphere (Southern Hemisphere) high/low EM area years during 1985/2003 (1986/1983) with
a general decline in EM area strongly influenced by northern hemisphere trends of summer SST
anomalies.
1. Introduction
The Köppen Climate Classification system defines polar climates as those regions where
the mean temperature of the warmest month is below 10oC. Köppen further divided the polar
climates into Polar Tundra (ET) where the warmest month was above 0oC but below 10oC and
Polar Ice Cap (EF) in which the warmest month was below 0oC. The upper limit of 10oC for the
warmest month corresponds roughly with the poleward limit of tree growth. The warmest month
limit of 0oC corresponds roughly with the equatorward limit of permanent snow and ice on land
(Rohli and Vega 2011).
Shear (1964) suggested that the Polar Tundra climate be further divided into the Polar
Marine (EM) climate, in which the mean temperature of the coldest month is above -6.7oC
(20oF), and the remaining Polar Tundra (ET) in which the coldest month is below -6.7oC. Shear
chose 20oF as the lower limit of coldest month in the EM climate to limit its occurrence to
marine environments. Shear expected that the poleward margins of the EM climate would
coincide with the maximum winter extent of pack ice as the pack ice boundary is the seasonal
projection of the pseudocontinent whose role in terms of energy exchange processes is more like
snow-covered land than open water. He used data from about 20 weather stations on islands or
continental coasts and various marine climate atlases from the early 20th century to define the
EM regions. He noted that an absence of data in many polar regions made the location of
boundaries difficult to determine.
Based on the data available in the early and mid-twentieth century, Shear (1964)
identified three primary regions with EM climate – (a) the southern Bering Sea and Aleutian
Islands in the North Pacific, (b) a southwest-to-northeast trending region in the North Atlantic
from south of Greenland through the Denmark Strait across much of Iceland and to the Barents
Sea north of Norway and the Kola Peninsula, and (c) a circumpolar zone over the Southern
Ocean and sub-Antarctic islands between roughly 49oS and 60oS (Fig. 1).
Polar marine climates may exist at high elevations outside of the polar regions, as noted
by Shear (1964). Mark et al. (2000) reported on a short temperature record above treeline in the
Southern Alps of New Zealand (44oS). Mean temperatures of the warmest and coldest months
were 8.9oC and -1.4oC, respectively, at 1600 m and 6.4oC and -4.8oC at 2000m elevation.
Noguchi et al. (1987) reported the highest summits on the tropical island of Hawaii (~19oN),
Mauna Loa (4169 m) and Mauna Kea (4205 m), have a mean temperature of the warmest month
between 0oC and 10oC with the coldest month warmer than -6.7oC and placed them in Shear’s
polar marine climate type.
Other characteristics of the EM climate, as described by Shear, distinguish the EM
climate from the continental ET climate. The EM climate is wetter than ET, has a winter
precipitation maximum rather than a summer maximum, a larger proportion of the annual
precipitation falls as rain, and there are more days of precipitation. The EM climate also has
greater storm frequency, more cloud cover, and stronger winds than the ET climate.
The EM climate classification has been incorporated into some text books and
descriptions of regional climates. Christopherson (2009) described the Köppen EM Polar Marine
climate as having all months above -7oC and the warmest month 0o to 10oC. Oliver and Hidore
(2002) described the EM climate in the Köppen system as a “Polar Wet” climate with mean
monthly summer temperature up to 10oC and winter means between -6.7oC and 0oC. Stern et al.
(2000) proposed several modifications to Koppen’s climate classification for application in
Australia, including a polar maritime subdivision reflecting the climate of the sub-Antarctic
islands. They do not cite Shear and imply a minimum temperature of the coldest month of -3oC
for their polar maritime climate.
These EM domains have experienced varying degrees of warming temperatures over
time. Warming trends have been especially robust in the Arctic where observations have shown
terrestrial low Arctic (64-70°N) annual temperature trends +0.38°C dec-1 from 1970-2008
(Chylek et al. 2009), while temperature trends over land north of 60°N have been estimated at
nearly +0.64°C from 1979-2008, indicative of an amplified warming signal (Bekryaev et al.
2010). The Arctic SSTs likewise rose from 1965-1995, however during roughly the last two
decades have notably increased, especially in the Western Arctic (Steele et al. 2008). These
ocean/atmosphere warming trends have coincided with recent, rapid deteriorating sea ice extent
and thickness (Maslanik et al. 2007; Kwok and Rothrock, 2009). This has led to anomalous
upward latent and sensible heat flux during late fall and winter and warmed the lower
troposphere (Serreze et al. 2009; Kumar et al 2010; Screen and Simmonds, 2010), not to mention
delayed refreeze and promoted earlier melt onset (Markus et al., 2009) and decreased March
maximum ice cover (Nghiem et al. 2007).
There is much less consensus on southern hemisphere warming as temperature trends are
spatially and temporally variable. Monaghan et al (2008) found statistically insignificant
positive temperature trends over most of Antarctica during most months from 1960-2005 only to
find weak, negative temperature trends 1970-2005. However, the widespread negative summer
and autumn temperature trends during the latter period show a positive signal over 1992-2005.
Walsh and Chapman (2007) found warming trends over 60-90°S during all seasons from 19582002, especially winter (0.172°C decade-1), most pronounced over the Antarctic Peninsula. Steig
et al. (2009) constructed a 50-year climatology (1957-2006) that revealed significant annual
warming (0.18°C decade-1) over West Antarctica, largely focused on winter and spring seasons.
Moreover, these authors found substantial increases in annual temperatures over this area 19792003 coincided with declines in sea ice fraction upwards of 20% in the adjacent western
Weddell, Amundsen and Bellingshausen Seas (their Fig. 4). Schneider et al. (2012) also found
robust warming over the region during spring, largely forced by atmospheric and SST factors,
that also coincided with noticeable Amundsen and Bellingshausen sea ice declines. Similar to
temperature trends, sea ice is highly variable in the Antarctic. Studies of Antarctic sea ice cover
have shown a slight positive annual trend during the past couple decades (Cavalieri et al. 2003),
estimated at nearly 1% increase per decade from 1979-2006 (Cavalieri and Parkinson 2008).
More recently, Parkinson and Cavalieri (2012) found that sea ice area has been increasing at a
rate of 17100±2300 km2 yr-1 from 1979-2010 with notable gains in the Ross Sea outweighing
substantial losses in the Bellingshausen and Amundsen Seas.
EM climates are also strongly influenced by regional atmospheric and oceanic circulation
patterns of variability. North Atlantic climate is largely modulated by the North Atlantic
Oscillation (NAO; van Loon and Rogers 1978) and Arctic Oscillation (AO; Thompson and
Wallace 1998). The positive phase of the NAO teleconnection (manifested by a deepened
Icelandic Low) most frequently occurs during winter, produces increased surface winds that
propagate poleward fluxes of warm air and water through the Fram Strait and Barents Sea
(Karcher et al. 2003; Rogers et al. 2004) and can lead to late winter ice extent declines (Dickson
et al. 2000; their Fig. 13). Perhaps a larger scale contributor than the NAO, the positive AO has
also been linked to interannual Arctic temperature increases and zonal ice transport into the
eastern Arctic (Rigor et al. 2002; Serreze and Barry 2005).
The North Pacific underwent a massive climatic shift in the mid-1970s centering around
the Pacific Decadal Oscillation (PDO; Mantua et al. 1997) switch from negative to positive
phase yielded a rise of regional marine and terrestrial temperatures on seasonal and annual scales
as an amplified Aleutian Low prompted an increased southerly windfield, warm air advection
and increased storminess (Hartman and Wendler 2005). Warm water advection poleward and
positive SST anomalies in recent years have also paralleled this event and the decreasing western
Arctic ice cover (Woodgate et al. 2006, 2010).
Southern Ocean climate is largely modulated by ocean-atmosphere interactions between
the zonal, circumpolar pressure anomalies of the Southern Annular Mode (SAM; Thompson and
Wallace 2000) and tropical Pacific ocean-atmosphere feedbacks of El Niño-Southern Oscillation
(ENSO; Mo and Ghil 1987). Since the 1990s, El Niño and negative SAM and La Niña and
positive SAM have been common covarying modes, especially during spring and summer (Fogt
and Bromwich 2006). The latter combination during spring yields negative pressure anomalies
near 90-130°W, northerly winds, earlier sea ice retreat and later advances around the western
Antarctic Peninsula and southern Bellingshausen Sea and westerly winds culminating with later
spring retreat in the western Ross Sea (Stammerjohn et al. 2008).
The purpose of our study is to map and explore the modern EM climate using
contemporary reanalyses to expand Shear’s work over the period of 1979-2010. Global coupled
atmosphere/ocean reanalyses represent the most prudent way to assess the EM climate for
several reasons including the fact that these hindcast simulations combine a multitude of data
sets, including weather stations, buoys, aircrafts, rawinsondes, satellites and other sources, in
order to depict atmospheric and oceanic conditions over large, remote areas where direct
observations are lacking. Reanalysis output since 1979 are more reliable due to the fact that
model integration of satellite data has afforded high quality and near global observational
coverage (Serreze et al 2009; Serreze and Barrett 2011), especially over the Southern Ocean
where there are few insitu meteorological observations (Bromwich and Fogt 2004). Further
sections of the paper are organized as follows. In Section 2, data and methodology will be
addressed. In Section 3, we compare reanalyses, justify selection of one reanalysis for
comparison with Shear’s original outputs and forthcoming plots, determine the poleward and
equatorward EM boundaries, look at some variables that impact this climate regime, and
examine EM interannual variability over the period of study. Section 4 will assess differences
from Shear’s study and briefly address the possibility of future EM changes.
2. Data and Methodology
Global reanalyses use fixed numerical weather prediction models and data assimilation
schemes to produce gridded fields over time periods suitable for climate research. Reanalyses
are particularly useful over polar regions, providing a coherent representation of weather and
climate where relatively short temporal spans of data records and areas of sparse observations
exist. However, caution must be exercised when using reanalyses to study climate trends, as
output is sensitive to changes of the observing system and how observations are processed
(Bengtsson et al. 2004a,b; Sterl 2004; Thorne and Vose 2010; Screen and Simmonds 2011).
Such changes result in erroneous trends, particularly over Southern Hemisphere polar regions
(e.g., Hines et al. 2000; Marshall and Harangozo 2000; Marshall 2002), limiting the viability of
reanalysis products in these regions to the post-1978 modern satellite era (e.g., Bromwich and
Fogt 2004; Renwick 2004; Trenberth et al. 2005; Bromwich et al. 2007). There are also
substantial differences between reanalyses, based on different models and parameterizations,
observations, and data assimilation systems (e.g., Bromwich and Fogt 2004; Bromwich et al.
2007; Walsh et al. 2009; Screen and Simmonds 2011; Bromwich et al. 2011). In this study we
use three reanalyes, described below, for the 32-year period from 1979-2010.
The European Centre for Medium-Range Weather Forecasts (ECMWF) “Interim”
Reanalysis (ERA-Interim, Dee et al. 2011) supersedes the ERA-40 reanalysis (Uppala et al.
2005), and improves upon ERA-40 in several regards (see Dee et al. 2011). ERA-Interim uses a
12-hourly 4D-Var data assimilation system, and also uses an automated satellite radiance
variational bias correction scheme (Dee and Uppala 2009). The observational sources for polar
regions are listed in Andersson (2007) and Dee et al. (2011). ERA-Interim features spectral
T255 (~0.7°) horizontal resolution and 60 vertical levels. Output on a regular 512x256 0.7°
Gaussian grid from the National Center for Atmospheric Research Data Support Section (NCAR
DSS) is used in this study.
The National Centers for Environmental Prediction (NCEP) Climate Forecast System
Reanalysis (CFSR, Saha et al. 2010) is a coupled atmosphere-ocean-land surface-sea ice model
that supersedes the NCEP/Department of Energy Atmospheric Model Intercomparison Project 2
reanalysis (Kanamitsu et al. 2002). CFSR uses a 3D-Var gridpoint statistical interpolation (GSI)
data assimilation system (Kleist et al. 2009), and ingests a wide array of satellite observations in
radiance form. The CFSR atmospheric component features spectral T382 (~0.31°) horizontal
resolution with 64 vertical levels. Output on a 720x361 0.5° latitude/longitude grid from the
NCAR DSS is used here.
The Japan Meteorological Agency (JMA) 25-year reanalysis (JRA-25, Onogi et al. 2007)
uses the JMA numerical weather prediction and data assimilation systems. JRA-25 uses a 3DVar data assimilation scheme that ingests satellite radiances and features spectral T106 (~1.125°)
horizontal resolution with 40 vertical levels. Output on a regular 320x160 1.125° Gaussian grid
from the NCAR DSS is used in this study.
The previously described Polar Marine (EM) climate classification from Shear (1964)
was applied to each gridpoint of 2-m temperature from the three reanalyses. The union of the
regions of warmest month mean temperature greater than 32°F but less than 50°F and coldest
month mean temperature greater than 20°F represent EM climate. In addition to 2-m
temperature, precipitation, mean sea-level pressure, sea-surface temperature, sea ice fraction, 10m wind, and total cloud fraction are also analyzed to provide a more complete description of EM
climate. While all three reanalyses are used to characterize the spatial distribution of EM area, to
simplify the analysis, we solely use ERA-Interim for the detailed description of EM climate.
ERA-Interim is the only reanalysis to use 4D-Var data assimilation, and it along with its
predecessor ERA-40 compare favorably against other global reanalyses in both the Northern
Hemisphere (e.g., Bromwich and Wang 2005; Bromwich et al. 2007; Walsh et al. 2009; Screen
and Simmonds 2011) and Southern Hemisphere (Bromwich and Fogt 2004; Monaghan et al.
2006; Bromwich et al. 2011; Hodges et al. 2011) high latitudes. The primary findings of this
study are not critically dependent upon which reanalysis is used for detailed analysis.
3. Results
a. Intercomparison of the Reanalyses
Figure 2a shows EM area for all three reanalyses over the Northern Hemisphere. The
area east of Newfoundland extending south of Greenland, across Iceland, and over the
Norwegian and Barents Seas matches well between all three reanalyses. The second area over
the Bering Sea also shows general agreement, although the JRA-25 and CFSR areas are slightly
larger than ERA-Interim. There are also scattered small EM regions along the southern Alaska
and western Canadian coastlines. Additional small high-altitude mid-latitude regions primarily
show up in CFSR, likely due to the enhanced horizontal resolution compared to the other
reanalyses. The two large-scale EM areas in Figure 2a are along the primary high-latitude storm
tracks (e.g., Hoskins and Hodges 2002), where warm and moist air is advected into these areas
from the south. Notice the eastern offset of EM areas from the Canadian, Greenland, and
Siberian coasts, where continental effects prevent establishment of EM climate until a marine
influence dominates offshore.
Figure 2b shows EM area for the Southern Hemisphere from all three reanalyses. The
Southern Hemisphere contains 90% of global EM climate area. EM area exists over much of the
Southern Ocean, and farther equatorward in the Southern Hemisphere compared to the Northern
Hemisphere. Differences between reanalyses are again small, with CFSR extending slightly
farther south, ERA-Interim extending slightly north, and some discrepancies over southern Chile.
CFSR also identifies EM over southwestern New Zealand. The Southern Hemisphere EM area
generally follows the Southern Hemisphere storm track, which dips poleward from the south
Atlantic eastward to regions south of New Zealand and into the south Pacific. However, the EM
area occurs on the northern edge of the primary circumpolar storm track (e.g., Simmonds et al.
2003; Hoskins and Hodges 2005), where equatorward incursions of Antarctic airmasses allow
for establishment of EM climate in otherwise marine environments.
b. Assessment of the EM Boundaries
The ERA-Interim EM boundaries are mapped in Figure 3a and 3b along with the winter
sea ice maximum, represented by the 25% ice concentration extent, as the poleward boundary
and the warmest-month SST 10ºC isotherm (month?) as the equatorward boundary. Figure 3a
shows the SST threshold values match the southern EM boundary well in the North Atlantic
from approximately 45ºW northeastward until the northwestern Norwegian coastline at about
15ºE before ending just north of Scandinavia in the Norwegian Sea. The maximum sea ice
extent fit as the poleward EM boundary improves east of Greenland from the Denmark Strait and
Greenland Sea and represents an acceptable border southeast of Svalbard and southwest of
Novaya Zemlya in the Barents Sea. The SST boundaries of the North Pacific (Fig. 3a) almost
encircle the EM and trace the part of the Aleutian Islands with better fit on the eastern periphery
near northern Kuskokwim Bay off the southwestern Alaskan coastline. The sea ice maximum
roughly matches the northern extent of this region’s EM at 60ºN. These regions contrast each
other’s EM latitudinal position as the southernmost North Atlantic EM (~55ºN) almost matches
the northernmost EM in the North Pacific.
The Antarctic EM boundaries are also well represented by SST and maximum sea ice as
shown in Figure 3b. The SST boundary (represented by month?) is slightly farther south in the
South Atlantic between 30ºW and 60ºW, but otherwise stays pretty consistent between 40º-60ºS
around the Southern Ocean as it mirrors the equatorward (northern) EM area. The maximum sea
ice poleward (southern) boundary (month?) is generally within 1-2° of the EM climatic extent
roughly between 58°S and 65ºS around the Antarctic continent.
c. EM Climatic Characteristics
Shear (1964) argued that the Polar Marine (EM) climate had distinctive characteristics
that distinguished it from the continental Polar Tundra (ET) climate. The mild winter and
smaller annual temperature range were the chief distinguishing features upon which Shear
defined the EM climate. In addition, he noted the EM climate has more precipitation and more
days of precipitation than the ET, the EM climate has a tendency toward a winter maximum of
precipitation instead of a summer maximum, and EM precipitation is dominantly rain. Shear
noted that the EM climate has greater storm frequency, more cloud cover, and stronger winds
than the continental ET. To examine these distinguishing characteristics of the EM climate in
the modern reanalysis data, we present patterns of annual mean sea level pressure, wind speed,
precipitation, and cloud over the Northern Hemisphere (Figs 4a-d) and Southern Hemisphere
(Figs. 5a-d).
The Polar Marine climates of the North Atlantic and North Pacific oceans are clearly
regions of low mean sea level pressure, indicating prevailing tracks of cyclonic storms (Fig. 4a).
The storm frequency in the EM climates is greater than in the northern continental ET climates.
This greater storm frequency should be associated with greater precipitation, wind speed, and
cloud cover. Wind speeds are greater in the EM climates than in the ET climates of the
continents, however even greater wind speeds occur south of the EM climates in the Atlantic and
Pacific (Fig. 4b). The EM climate region northeast of Iceland displays a regional minimum of
wind speed. Annual precipitation is greater in the EM climates than in the continental ET
climates (Fig. 4c). Precipitation is greater in the Atlantic EM climate (600-900 mm east of
Iceland, 900-1200 m west of Iceland) than in the Pacific EM climate (600-900 mm) and in both
ocean basins the precipitation increases southward from the EM climates. Cloud cover is 8090% in the EM climates, somewhat higher than over the high latitude continents (Fig. 4d).
In the southern hemisphere, the lowest mean sea level pressure is on the poleward
margins of the EM climate region (Fig. 5a). There is scant area of ET climate in the southern
hemisphere for comparison. The southern storm track appears to be displaced southward rather
than situated in the core of the EM climate as occurs in the northern EM climates. Mean wind
speeds reach a maximum through the core of the southern EM climate (Fig. 5b) and are stronger
than in the northern EM climates. Wind speeds are strongest in the eastern hemisphere and reach
a peak between 50oE and 100oE. Precipitation is 600-1200 mm in the southern EM climate and
decreases toward the pole (Fig. 5c). The regions of greatest cloud cover (>90%), like the mean
pressure, are on the poleward margins of the southern EM climate or even south of the EM
climate (Fig. 5d). The southern EM climate has 80-90% cloud cover, similar to the northern EM
climate.
As expected and as predicted by Shear, the EM climates are stormy, windy, wet, and
cloudy. Where comparisons can be made in the northern hemisphere, these EM climates are
distinctive from their continental ET counterparts in all of these parameters.
d. Interannual Variability
Figure 6a shows the general downward trend of the EM area (-0.004241 x 1013 m2 yr-1,
significant at 97.1%) from 1979-2010, however the northern and southern hemisphere EM areas
display much different trends. The Northern Hemisphere EM area (Fig. 6b) displays an annual
trend of -0.045746 x 1012 m2 yr-1 (significant at 99%) that solely represents the EM decline, as
the Southern Hemisphere area (Fig. 6c), even being an order of magnitude higher, displays an
insignificant positive trend (Figure 6c; 0.000333 x 1013 m2 yr-1).
The interannual variability of the respective hemispheres is also much different
temporally and spatially. Figures 7a and 7b show the high (1985) and low (2003) EM years for
the northern hemisphere. During 1985, compared to the 32-year climatology, the EM extended a
few degrees farther south into the Labrador Sea and North Atlantic and also slightly farther north
into the Davis Strait. Slight increases are also found just east of Iceland and southeast of
Svalbard. The largest Northern Hemisphere EM increases are found in the North Pacific where
the EM expands to cover the Aleutians as well as the southern tip of the Kamchatka peninsula,
encroaching into the Sea of Okhotsk. In contrast, 2003 shows EM declines (blue) are noticeable
around both its equatorward and poleward boundaries in the North Atlantic. The largest declines
of this low area year undoubtedly occur in the North Pacific where almost the entire EM area is
lacking with the exception of a sliver at 60°N near the Bering Strait. Southern hemisphere EM
variability during high/low area years is much less distinct (Figs. 7c-d). During the high EM
(1986) there are slight poleward increases near 60°S just west of the Antarctic Peninsula and
Ross Sea and equatorward increases in the South Pacific between 90°W and 150°W. The low
EM year (1983) shows a decline in a similar region of the South Pacific that expressed growth in
1986. All of the maps, regardless of high or low year, show areas of increase/decline for the
respective years.
Behavior of the EM boundaries is one factor controlling these areas. Figure 8a and 8b
show the winter ice cover of the high/low years relative to 1979-2010. The negative anomalies
during 1985 in EM areas are much more pronounced versus the smaller negative anomalies of
2003, especially in the North Pacific, which would indicate that the ice concentration was
anomalously low during that winter allowing the possibility for northward expansion of EM area.
However, the overall lack of EM poleward growth during 1985 would indicate that perhaps
winter maximum ice extent is not the best controller of this climatic regime during this particular
year. SSTs during summer of 1985 were also slightly lower than average around the
equatorward EM limit (60°N) across most of the North Atlantic and North Pacific, which may
have allowed the EM to expand southward. In contrast, during summer 2003 the same areas
exhibited SSTs up to 3 K warmer, while that winter’s positive sea ice concentration anomalies in
the Barents Sea likely co-contributed to the decreased EM area. As previously mentioned, warm
water advection into both regions has been well documented in recent years, but the role of
higher ice concentrations indicative of heavy ice conditions, as what transpired in the Barents
Sea during winter 2003, is likely also a major contributor to past low EM years.
Southern hemisphere ice cover during high/low EM years (Figs. 9a-b) shows strong
negative anomalies in sea ice concentration in the Ross Sea during 1986 prompted a slight
increase in EM area whereas positive anomalies just west of the Antarctic Peninsula coincided
with the EM declines during 1986 (Fig. 7c). The circum-Antarctic sea ice concentration
anomalies are much less pronounced in 1983 and the equatorward boundary declines cannot be
conclusively tied to sea ice behavior. On the other hand, the largest EM declines in both Figures
7c and 7d can be tied to positive SST anomalies typically ≥1 K over those areas. During 1986, a
warm SST pocket between 40-60°S and 150°E and the International Dateline is the likely culprit
for a declining EM. Moreover, during the low year of 1983, large EM declines between 90°W
and 150°W (Fig. 7d) match well with general SST warming along the equatorward Southern
Ocean equatorward boundary between 40-60°S.
4. Discussion/Conclusion
The modern EM climate regions shown in Figures 3a and 3b are similar in general
locations and shapes to the regions presented by Shear in 1964 (Fig. 1) but differ in some aspects
as would be expected given the different time periods and methods used to develop the maps.
In the North Atlantic, the southern margin of the EM climate depicted by Shear (Fig. 1)
extends from off the coast of Newfoundland at about 53oN, 51oW, northeastward to southern
Iceland, and then northeastward to the North Cape of Norway and ending in the Barents Sea at
72oN, 41oE. The modern EM region (Fig. 3a) has a similar position off the coast of
Newfoundland but extends farther westward into the Labrador Sea than shown by Shear. The
southern boundary is similar to Shear’s across Iceland to North Cape but the modern EM climate
extends farther eastward along the Kola Peninsula and northward into the Barents Sea. The
northern boundary in the modern EM climate (Fig. 3a) is similar to Shear’s around southern
Greenland but extends hundreds of kilometers north of Shear’s northern boundary over the
Greenland Sea to 79oN near Svalbard.
In the North Pacific (Fig. 3a), Shear showed the EM climate in the Bering Sea extending
southward to 50oN across the Aleutian Islands from Umnak Island (167oW) westward beyond
Attu Island (173oE) to the Komandorskyie Islands and ending along the 50th parallel at about
153oE. The EM climate depicted in Fig 3a does not extend as far south or west. It does not
include the Aleutians, except near 180o, and does not extend west of about 172oE. The northern
extent of the EM climate in the Bering Sea depicted by Shear was just south of 60oN from
Kuskokwim Bay westward. Fig 3a shows a similar position in Kuskokwim Bay but the EM
climate extends westward while trending north of 60oN to near 175oE, well north of Shear’s
region along the Russian coast. The modern EM region in the Bering Sea is somewhat smaller
and displaced northward from Shear’s region.
The poleward (southern) boundary of the EM climate in the southern hemisphere (Fig.
3b) is near 60oS from 30oE to 100oE, where Shear depicts the boundary 2o to 3o north of 60oN.
From 30oE to 40oW, the modern EM boundary is north of 60oS, similar to Shear’s. From the
Antarctic Peninsula (~50oW) westward to 120oE, the modern EM boundary is well south of 60oS
and south of Shear’s southern boundary. The equatorward (northern) boundary of the EM region
is near 50oS in the western hemisphere, but near 45oS elsewhere and this is in general agreement
with Shear’s northern boundary.
EM climates are largely modulated by the boundary parameters, summer SSTs and
winter maximum sea ice extent, laid forth in this study. However, temperature fluctuations and
the response of these environmental variables over time will prompt EM areas to change. Should
summer SSTs warm and winter sea ice extent decline, we would expect the EM area to shift
poleward. However, as we have seen examining high/low EM years, these parameters do not
necessarily behave in tandem (increased SSTs, decreased ice cover). Therefore, an EM climate
shift, versus a change in area, in either hemisphere may not be realistic on interannual or
extended temporal scales moving forward. Continued monitoring of an array of
atmospheric/oceanic variables including warm water flux into the polar regions and the recovery
of sea ice during the winter, especially in the northern hemisphere, is essential to determining the
future of EM climates.
Acknowledgements: The authors would like to thank Andrew Monaghan for constructive
comments. NCAR is sponsored by the National Science Foundation and other agencies.
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Figures
FIG 1. The EM climate as depicted by Shear (1964).
FIG 2. (a)Overplots of EM area for the Northern Hemisphere (1979-2010) from ERA-Interim,
CFSR and JRA-25 reanalyses, (b) same as (a) but for the Southern Hemisphere.
FIG 3. (a) ERA-Interim EM area (green) from 1979-2010 with warmest month SST (red) and
maximum winter 25% sea ice concentration (blue) for the Northern Hemisphere, (b) same as (a)
but for the Southern Hemisphere
FIG 4. ERA-Interim annual mean (1979-2010) (a) MSLP (hPa), (b) 10-m windspeed (m s-1), (c)
precipitation (mm yr-1), and (d) cloud cover (%) for the Northern Hemisphere. ERA-Interim EM
area outlined by solid black contours.
FIG 5. As in Fig. 4, but for the Southern Hemisphere. MSLP not plotted at elevations over 1000
m over Antarctica.
FIG 6. Time series of ERA-Interim EM (a) total area (Northern and Southern Hemisphere; 1013
m2 yr-1), (b) Northern Hemisphere area (1012 m2 yr-1), and (c) Southern Hemisphere area (1013
m2 yr-1) from 1979-2010. Linear trends and statistical significance shown in bottom left of each
plot.
FIG 7. (a) ERA-Interim representation of highest EM area year (1985) in the Northern
Hemisphere, (b) lowest EM area year (2003) in the Northern Hemisphere, (c) same as (a), but for
Southern Hemisphere (1986), (d) same as (b), but for Southern Hemisphere (1983). Green
represents the EM area (1979-2010), red represents additional EM area for that specific year, and
blue represents missing EM area for that year, relative to 1979-2010 monthly average
temperatures.
FIG 8. (a) ERA-Interim mean winter (DJF) sea ice concentration anomalies for 1985 minus the
1979-2010 climatology in the Northern Hemisphere. ERA-Interim EM area for 1979-2010
average outlined by solid black contours, EM area for 1985 outlined by dashed black contours.
(b) same (a), but for 2003. (c) Mean summer (JJA) SST (K) anomalies for 1985 minus the 19792010 climatology in the Northern Hemisphere. ERA-Interim EM area for 1979-2010 average
outlined by solid black contours, EM area for 1985 outlined by dashed black contours. (d) same
as (c), but for 2003.
FIG 9. (a) ERA-Interim mean winter (JJA) sea ice concentration anomalies for 1986 minus the
1979-2010 climatology in the Southern Hemisphere. ERA-Interim EM area for 1979-2010
average outlined by solid black contours, EM area for 1986 outlined by dashed black contours.
(b) same (a), but for 1983. (c) Mean summer (DJF) SST anomalies for 1986 minus the 19792010 climatology in the Southern Hemisphere. ERA-Interim EM area for 1979-2010 average
outlined by solid black contours, EM area for 1985 outlined by dashed black contours. (d) same
as (c), but for 1983.