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    Counts of surface cyclone centers in equal area boxes of 250 km × 250 km for (a) winter [December–February (DJF)] and (b) summer [June–August (JJA)] totaled over the period of 1958–2005 for the region north of 60°N. Results are based on NCEP reanalysis data. Fields have been smoothed by first averaging counts for each grid cell, along with counts at the four adjacent cells, and then applying a nine-cell center-weighted average. Bold contours indicate regions with at least 135 cyclone centers. Due to spurious cyclones associated with reduction of surface pressure to sea level, Greenland is masked.

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    The annual cycle of number of cyclone centers within the 135 count contour over the central Arctic Ocean depicted in Fig. 1b is shown. Values represent sums over the period of 1958–2005.

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    Latitude by height (pressure) cross section of the mean zonal wind for July, averaged over the period of 1958–2005 and for 60°–270°E longitude.

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    Gridcell counts over the period of 1958–2005 of events of cyclogenesis for systems that at any point in their life cycle were found within the central Arctic summer cyclone maximum. The colors of the circles correspond to the number of events. Greenland is masked.

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    Same as Fig. 4, but for events of maximum cyclone deepening.

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    Same as Fig. 4, but for events of cyclolysis.

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    Fields of mean (left) 500-hPa height (gpm) and (right) sea level pressure (hPa) over the period of 1958–2005 for (a), (b) January and (c), (d) July.

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    Time series of 1958–2005 of counts of cyclone centers for summer (total length of bars), and broken down for June (black bars), July (gray bars), and August (lightly shaded bars), for the region defining central Arctic Ocean cyclone maximum.

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    Composite means of (a) 500-hPa height (gpm) and (b) sea level pressure (hPa) for the 21 months when the summer cyclone pattern over the central Arctic Ocean was most strongly expressed.

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    Same as Fig. 9, but for when the summer cyclone pattern was most weakly expressed.

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    Composite differences in 500-hPa height (gpm) between the 21 months when the summer cyclone pattern over the central Arctic Ocean was most strongly and most weakly expressed (strong minus weak).

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    Same as in Fig. 11, but for the composite differences in 925-hPa temperature (°C).

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    Same as in Fig. 11, but for the composite differences in sea level pressure (hPa).

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    Composite differences in 500-hPa height (gpm) between the 21 months when the summer NAM was most strongly positive and most strongly negative (positive minus negative).

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    Summer cyclone center counts for the region north of 60°N for the 21 months when the NAM was (a) most strongly positive and (b) most strongly negative. Fields have been smoothed as in Fig. 1. Greenland is masked.

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    Fields for June 1989 of (a) 500-hPa height (gpm), (b) sea level pressure (hPa), and (c) 500-hPa Eady growth rates (day−1); regions where the growth rate exceeds 0.5 day−1 are shown in bold contours.

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    Same as Fig. 16, but for July.

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    Composite differences in aerologically calculated net precipitation (PE) from ERA-40 between months when the summer cyclone pattern over the central Arctic Ocean was strongly and weakly expressed (strong minus weak).

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The Summer Cyclone Maximum over the Central Arctic Ocean

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  • 1 Cooperative Institute for Research in Environmental Sciences, University of Colorado, Boulder, Colorado
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Abstract

A fascinating feature of the northern high-latitude circulation is a prominent summer maximum in cyclone activity over the Arctic Ocean, centered near the North Pole in the long-term mean. This pattern is associated with the influx of lows generated over the Eurasian continent and cyclogenesis over the Arctic Ocean itself. Its seasonal onset is linked to the following: an eastward shift in the Urals trough, migration of the 500-hPa vortex core to near the pole, and development of a separate region of high-latitude baroclinicity. The latter two features are consistent with differential atmospheric heating between the Arctic Ocean and snow-free land. Variability in the strength of the cyclone pattern can be broadly linked to the phase of the summer northern annular mode. When the cyclone pattern is well developed, the 500-hPa vortex is especially strong and symmetric about the pole, with negative sea level pressure (SLP) anomalies over the pole and positive anomalies over middle latitudes. Net precipitation tends to be anomalously positive over the Arctic Ocean. When poorly developed, the opposite holds.

Corresponding author address: Mark Serreze, CIRES, University of Colorado, Boulder, CO 80309. Email: serreze@kryos.colorado.edu

Abstract

A fascinating feature of the northern high-latitude circulation is a prominent summer maximum in cyclone activity over the Arctic Ocean, centered near the North Pole in the long-term mean. This pattern is associated with the influx of lows generated over the Eurasian continent and cyclogenesis over the Arctic Ocean itself. Its seasonal onset is linked to the following: an eastward shift in the Urals trough, migration of the 500-hPa vortex core to near the pole, and development of a separate region of high-latitude baroclinicity. The latter two features are consistent with differential atmospheric heating between the Arctic Ocean and snow-free land. Variability in the strength of the cyclone pattern can be broadly linked to the phase of the summer northern annular mode. When the cyclone pattern is well developed, the 500-hPa vortex is especially strong and symmetric about the pole, with negative sea level pressure (SLP) anomalies over the pole and positive anomalies over middle latitudes. Net precipitation tends to be anomalously positive over the Arctic Ocean. When poorly developed, the opposite holds.

Corresponding author address: Mark Serreze, CIRES, University of Colorado, Boulder, CO 80309. Email: serreze@kryos.colorado.edu

1. Introduction

The past decade has seen an explosion of literature concerning the atmospheric circulation of the north polar region. To a considerable degree, this stems from the recognition that rapid changes observed in the Arctic, including rises in surface air temperature and declining sea ice extent, can be explained in part by attendant shifts in atmospheric patterns.

Most of this interest has focused on winter. Studies of links with the North Atlantic Oscillation (NAO) and its hemispheric-scale counterpart, the northern annular mode [NAM; also known as the Arctic Oscillation (AO)], have been especially prominent. From about 1970 through the mid-1990s, the NAO and NAM exhibited winter-season shifts from their negative to positive phases, associated with lower sea level pressure over the Arctic, especially in the vicinity of the Icelandic low. Numerous studies (e.g., Hurrell 1995; Thompson and Wallace 1998; Moritz et al. 2002) document the attendant changes in wind fields that help to explain strong winter warming over large parts of northern Eurasia during this period, as well as partly compensating cooling over eastern Canada and parts of Greenland. Through impacts on sea ice circulation, behavior of the winter NAM has also contributed, along with general warming and changing oceanic heat transport, to the observed downward trend in Arctic sea ice extent (Rigor et al. 2002; Comiso 2003; Rigor and Wallace 2004; Rothrock and Zhang 2005; Polyakov et al. 2005; Shimada et al. 2006). Recent warming in Alaska can be allied with a shift in the Pacific decadal oscillation from a generally negative phase from 1951 to 1976 to a primarily positive phase from 1977 to 2001. The deeper Aleutian low during the later positive phase helped to transport warm, moist air into the region (Hartmann and Wendler 2005).

Given that summer is the most synoptically active period over the central Arctic Ocean, it is surprising that this season has not garnered more attention. To illustrate this point, Fig. 1 shows the spatial distribution of counts of closed surface low pressure centers over the Arctic for winter and summer totaled over the period of 1958–2005, based on an algorithm (Serreze et al. 1997) applied to 6-hourly sea level pressure (SLP) fields from the National Centers for Environmental Prediction–National Center for Atmospheric Research (NCEP–NCAR; hereafter NCEP) reanalysis (Kalnay et al. 1996). Slow-moving cyclones could be counted twice or more in a given grid cell. The map patterns have been smoothed to highlight the major features. Details of the algorithm and screening of the database follow shortly.

Winter cyclone activity (Fig. 1a) is maximized in the vicinity of the mean Icelandic low, at the southern boundary of the map, reflecting both the migration of lows into the region from the south and frequent cyclogenesis (e.g., Whittaker and Horn 1984). High counts also extend northeastward into the Norwegian and Barents Seas, broadly defining the terminus of the primary North Atlantic storm track. Winter cyclones are, of course, very common in the vicinity of the Aleutian low, but relatively few of these migrate into the Arctic. The summer pattern (Fig. 1b) is very different. Activity in the Norwegian and Barents Seas is much less prominent. There is much stronger activity over land areas than in winter, especially over Eurasia.

Particularly striking, however, and representing the focus of the present paper, is the distinct summer pattern of high cyclone center counts over the central Arctic Ocean, focused at about 85°N along the date line (the summer cyclone maximum herein). Figure 2 gives the annual cycle of the cyclone center counts summed over the period of 1958–2005 for the region enclosing the 135 count contour in Fig. 1b. Onset of the pattern is seen as a sharp increase in activity between May and June. Activity peaks in July and August and then declines sharply between August and September. There is over a factor of 2 difference in counts of low pressure centers between January and July.

Obtaining a better understanding of the summer cyclone maximum and its variability is justified on at least two counts—its link with the Arctic heat budget and sea ice conditions, and its link with the freshwater budget of the Arctic Ocean. Under cyclonic wind stress, Ekman transport at the sea surface to the right of the wind forcing promotes ice divergence (Thorndike and Colony 1982). Prominent reductions in ice concentration over the central Arctic Ocean have been observed for many summers when the cyclone pattern was well developed, such as in 1980 and 2002 (Barry and Maslanik 1989; Serreze et al. 2003b). The open-water areas both reduce the regional albedo and foster enhanced heat fluxes to the atmosphere in autumn. A strongly developed cyclone pattern also spreads the existing ice over a larger area, contributing to the high variability in Arctic ice extent that has attended the overall downward trend (Ogi and Wallace 2007). The cyclone pattern impacts the freshwater budget of the Arctic Ocean through promoting a summer/early autumn peak in precipitation and net precipitation [precipitation minus evaporation (PE); Walsh et al. 1994; Cullather et al. 2000; Serreze et al. 2006].

In the former Soviet Union, recognition of the summer cyclone maximum can be traced to the pioneering work of Dzerdzeevskii (1945). Through the early 1950s, the prevalent theme in western climatology literature was of a rather quiescent circulation over the central Arctic Ocean in all seasons. While more modern western views can be traced to studies such as those by Wilson (1958), Hare and Orvig (1958), and Keegan (1958), there was limited recognition of the summer cyclone maximum until the seminal work of Reed and Kunkel (1960). Subsequent papers (e.g., Serreze and Barry 1988; Serreze et al. 1993, 2001) have addressed the feature as part of general studies of the Arctic circulation. LeDrew et al. (1992) examined the strongly cyclonic month of August 1980 as a case study.

The basic view is that the summer cyclone maximum pattern arises primarily from systems generated over Eurasia and along the weakened North Atlantic track that migrate into the Arctic Ocean and subsequently occlude (e.g., Reed and Kunkel 1960; Serreze 1995). Building on Reed and Kunkel (1960), Yoshimura (1967), Kurashima (1968), and others, Serreze et al. (2001) argued for a link with the summer “Arctic frontal zone” that appears to develop in response to differential atmospheric heating between the Arctic Ocean and snow-free land. This frontal zone, which extends from roughly 60° to 270°E, is expressed aloft as a weak but distinct high-latitude maximum in the zonal wind at about 300 hPa, separate from the middle-latitude jet. Figure 3, a latitude-by-height cross section of the July zonal wind that is averaged over this longitude range and is based on NCEP reanalysis data over the period of 1958–2005, shows this feature clearly. Cross sections for June and August (not shown) are similar, but there is less separation between the middle- and high-latitude wind maxima. One neither sees separation in other months, nor does one see a separate high-latitude wind maximum over the Atlantic sector in any month.

The coastal baroclinicity is sharpest over northeastern Siberia and north of Alaska, where topography appears to help “trap” the cold Arctic Ocean air. These are also areas of summer cyclogenesis, with the eastern Eurasian feature having an expression in Fig. 1 in the regional maximum in cyclone center counts extending from about 65° to 70°N, and from 140° to 170°E. At least some of the systems generated over eastern Eurasia migrate into the central Arctic Ocean and contribute to the summer cyclone maximum, while those formed over northern Alaska instead tend to track eastward or southeastward into the Canadian Arctic Archipelago (Serreze et al. 2001).

More recent work suggests a more general framework, namely, the summer NAM. Thompson and Wallace (2000) defined the NAM using a single empirical orthogonal function (EOF) analysis of the zonal mean geopotential height fields for all calendar months. In this framework, the NAM is most strongly expressed in winter, with its center of action near the mean Icelandic low. Ogi et al. (2004), by contrast, defined the NAM separately for each calendar month through EOF analysis of zonally averaged geopotential height fields from 1000 to 200 hPa, poleward of 40°N, which is an approach that better captures seasonality in the structure of the mode. The most notable feature is that the Arctic center of action in summer is shifted north to lie almost over the pole, near the region of peak cyclone activity shown in Fig. 1. As discussed by Ogi et al. (2004), the winter NAM is maintained primarily by stationary waves, contrasting with summer when both stationary and transient waves contribute. When the summer NAM is positive, the Arctic jet over Eurasia and North America is prominent, and clearly separate from the jet at lower latitudes, that is, the Arctic frontal zone is stronger.

Despite these insights, our understanding of the summer cyclone maximum is far from complete. The objective of the present paper is to reexamine some of the above ideas through a more complete analysis of the source regions of systems comprising the cyclone maximum, the development and decay characteristics of these systems, and links with both regional- and large-scale aspects of the circulation. We show that while the summer pattern is, in part, associated with the influx of lows generated along northeastern Eurasia, where the Arctic frontal zone is especially well expressed, the broader picture involves an eastward shift in the Urals trough and migration of the 500-hPa vortex core to near the pole, associated with the influx of systems generated along a wide swath of the Eurasian continent, augmented by cyclogenesis within the Arctic Ocean itself. Variability in the cyclone pattern from month to month and year to year is indeed shown to be broadly related to the phase of the summer NAM.

2. Primary data and analysis tools

Extensive use is made of output from the cyclone detection/tracking algorithm used to compile Figs. 1 and 2. The 6-hourly outputs include the location of each cyclone center, cyclone central pressure, and 6-h deepening rate. It represents the most recent incarnation of the algorithm described by Serreze et al. (1997). This newer version has been used to examine Northern Hemisphere cyclone trends (McCabe et al. 2001), links between cyclone activity and precipitation variability over northern Eurasia (Serreze and Etringer 2003), cyclone activity associated with the summer Arctic frontal zone (Serreze et al. 2001), and the tracks and deepening characteristics of cyclones in the northern North Atlantic (Tsukernik et al. 2007).

The algorithm is applied to Northern Hemisphere SLP fields from the NCEP reanalysis. Cyclone detection is based on a series of search patterns that test whether a gridpoint SLP value is surrounded by gridpoint values at least 1 hPa higher than the central point tested. Cyclone tracking employs a nearest-neighbor approach that compares system positions for a given 6-h chart with those for the next 6-h chart. Cyclogenesis (cyclolysis) represents the first (last) appearance of a closed 1-hPa isobar.

The original algorithm was applied to 12-hourly SLP fields on the old National Meteorological Center (NMC) 47 × 51 octagonal grid. For application to the NCEP reanalysis data, SLP fields on the 2.5° × 2.5° latitude/longitude NCEP grid are reprojected to a 250 × 250 km version of the National Snow and Ice Data Center north polar Equal-Area Scalable Earth (EASE) grid (Armstrong and Brodzik 1995). Cyclone detection and tracking is performed on the reprojected fields. The procedure avoids the problem of strong convergence of the meridians at high latitudes, which would have been difficult to deal with given the architecture of the algorithm.

Cyclones that lasted less than a day (less than four 6-h charts) and those that remained stationary during their life cycle (always appeared at the same grid cell) are eliminated from consideration. This discards most spurious systems that appear from the reduction of surface pressure to sea level over high or complex topography. The first check also tends to eliminate smaller mesoscale features. To further limit the analysis to robust systems, a cyclone had to deepen sometime during its life cycle (a total of at least 2 hPa from the computed 6-h SLP changes over the course of its life cycle). Figures 1 and 2 are based on this screened database. Despite the screening, some spurious systems still remain, especially over Greenland. As such, the Greenland ice sheet is masked in Fig. 1 and in all subsequent depictions of the location of cyclone centers.

The screened cyclone database is used in conjunction with NCEP fields of 500-hPa height, SLP, and temperature. Two case studies employ fields of the Eady growth rate (σ), an index of baroclinic instability widely used in studies of synoptic development (e.g., Hoskins and Valdes 1990; Paciorek et al. 2002). The index relates baroclinic instability to the vertical wind shear (the thermal wind) and static stability. Larger values imply stronger instability. It is defined as
i1520-0442-21-5-1048-eq1
where ƒ is the Coriolis parameter, v is the horizontal wind vector, z is vertical distance, and N is the buoyancy frequency. Following Hall et al. (1994) and Paciorek et al. (2002), growth rates are calculated at the 500-hPa level, with the required gradients based on data at 400 and 600 hPa.

Six-hourly NCEP fields are available back to 1948. While fields should be most reliable for the period from 1979 onward, when the model can draw from modern satellite data streams, the rawinsonde database in northern high latitudes is fairly dense back to 1958. Through 1991, the Russian North Pole program provided rawinsonde coverage over the central Arctic Ocean. The good spatial and temporal coverage of soundings justifies using reanalysis data prior to the satellite era. Apart from the two case studies, all evaluations that follow based on NCEP data are for the period of 1958–2005.

A variety of other cyclone detection and tracking algorithms are in existence (e.g., Murray and Simmonds 1991; Grigoriev et al. 2000; Hoskins and Hodges 2002 and references therein; Zhang et al. 2004), each with different architectures. The widely used Hodges algorithm (for a recent application see Hodges et al. 2003) merits note. Instead of using SLP, it identifies and tracks systems based on tropospheric vorticity maxima. An issue relevant to the results that follow is that tropospheric vorticity maxima can often be identified before or after surface cyclogenesis and cyclolysis, based on the first and last appearance of a closed isobar in the SLP field. As such, track lengths of individual systems from the Hodges approach are often longer compared to those from the Serreze algorithm (M. Tsukernik 2007, personal communication). Drawing from the study of Hodges et al. (2003), one should also be aware that cyclone characteristics as assessed here using NCEP fields could differ somewhat from those based on the companion 40-yr European Centre for Medium-Range Weather Forecasts (ECMWF) Re-Analysis (ERA-40; Uppala et al. 2005).

3. General characteristics

a. Cyclogenesis, deepening, and cyclolysis

The screened cyclone database is first used to assess the origin of systems comprising the summer cyclone maximum. Figure 4 shows counts by EASE grid cell over the period of 1958–2005 of the location of cyclogenesis events for systems that at any time during their life cycle were present within the summer cyclone maximum (i.e., they “contributed to” the cyclone maximum). As with Fig. 2, the cyclone maximum is defined as the region enclosing the 135 count contour in Fig. 1b. The cyclogenesis event had to occur between June and August. A total of 584 systems met these criteria. The cyclone maximum region comprises an area of about 1.63 × 106 km2 (twenty-six 250 km × 250 km grid cells).

According to our algorithm, the majority of systems contributing to the summer cyclone maximum originated outside of the region (79% of the total). Most of these externally formed systems were generated over a broad swath of the Eurasian continent, sometimes from distant source regions, or over the Arctic Ocean in close proximity to the cyclone maximum region. Very few contributing systems were generated over either North America or the North Pacific. A few entered from the northern North Atlantic, and Norwegian and Greenland Seas. The remaining 21% of the contributing systems formed within the cyclone maximum region itself.

The location of maximum cyclone deepening (the location with the largest 6-h deepening rate) is a good indicator of where the strongest development occurs. Summed counts by grid cells of the location of the maximum deepening for any system contributing to the summer cyclone maximum appear in Fig. 5. Some cyclones that formed over Eurasia show their maximum deepening before entering the Arctic Ocean. However, there is a peak in the occurrence of maximum deepening over the Arctic Ocean. This is partly associated with systems entering from the outside that deepened in the area. It also reflects the observation that systems generated over the Arctic Ocean tend to spend their entire life cycle in the Arctic Ocean.

Figure 6 shows the number of cyclolysis events of systems contributing to the summer cyclone maximum (as with Figs. 4 and 5, “contributing to” means any system that at any time in its life cycle was present in the cyclone maximum region). Cyclolysis events are strongly clustered within the cyclone maximum region and the surrounding area. Put differently, once systems either enter the region from the outside or form within the region, they are largely destined to die in the region.

To summarize, based on the cyclone detection and tracking algorithm used here, the summer cyclone maximum of the central Arctic Ocean appears largely as a “collection zone” for systems, especially those generated over Eurasia. There is also some generation within the Arctic Ocean and the cyclone maximum region itself. Lows migrating into, or formed within, the cyclone maximum region almost invariably decay within, or in close proximity to, the region. Collectively, these processes result in a high count of cyclone centers.

Recall from earlier discussion that the summer Arctic frontal zone is expressed along a wide swath of the Eurasian coast, but is especially prominent over northeastern Eurasia (roughly the 65°–70°N, 140°–170°E region). However, the results in Fig. 4 show no clear dominance of contributions to the summer cyclone maximum from systems generated along the coast. With respect to eastern Eurasia, further inspection shows that systems formed in this region are more likely to track eastward, away from the summer cyclone maximum. This said, part of what we are seeing also seems to reflect shortcomings in the cyclone detection and tracking algorithm, namely, defining cyclogenesis as the first appearance of a closed isobar in the SLP field. When individual 6-hourly charts are examined, it is found that many cyclogenesis events defined as having occurred over the Arctic Ocean, but yet to enter the cyclone maximum region, can be associated with preexisting tropospheric waves originating along the coast. By the same token, the conclusion that systems entering the cyclone maximum region tend to decay in or very near the region must be tempered by the recognition that tropospheric vorticity maxima may exist after cyclolysis at the surface.

b. Relationships with large-scale circulation

Mean fields of 500-hPa height and SLP for January and July (Fig. 7) help to place both the strong seasonal cycle in cyclone activity and the results just presented into context. During January, the 500-hPa circulation (Fig. 7a) is dominated by the major eastern North American and East Asian troughs, the less prominent Urals trough over western Asia, and the longwave ridges separating these features. The core of the circumpolar vortex at 500 hPa is centered over the Canadian Arctic Archipelago, and height gradients (hence, geostrophic winds) over the central Arctic Ocean are slack. The SLP field (Fig. 7b) is dominated by the two subpolar lows and the Siberian high. The central Arctic Ocean lies under a ridge of high pressure.

With the exception of the sector east of about 80°E, the 500-hPa flow over Eurasia is strong. However, this strong flow has a southerly component, so that lows generated over the continent will tend to be steered away from the central Arctic Ocean. This is confirmed from an analysis like that used to compile Fig. 4 (but not shown here), in which cyclogenesis locations are plotted for systems that any point in their life cycle were within the central Arctic Ocean during the winter season. While there are far fewer systems over the central Arctic Ocean in winter than in summer, almost none formed over the Eurasian continent. Compared to summer, more of the systems originated in the Atlantic subpolar seas and the northern North Pacific; however, even in winter, cyclogenesis may occur over the central Arctic Ocean.

The 500-hPa and SLP fields for July are very different. Compared to January, the 500-hPa flow is weaker over much of Eurasia but the Urals trough has shifted eastward, is more pronounced in terms of curvature of the height field, and the flow ahead of it is more zonal (Fig. 7c). Cyclones generated over Eurasia are, hence, more likely to migrate toward the Arctic Ocean. Particularly pronounced changes are seen in higher latitudes. The core of the 500-hPa circumpolar vortex has shifted from the Canadian Arctic Archipelago to lie almost over the pole. While the flow over northern Eurasia east of about 80°E is weak in January, it is fairly strong in July. Cyclone development will be favored along this broad region. This is part of a larger change: from about 60° to 270°E there is separation between relatively tight height gradients in high latitudes and those in lower latitudes. The line of separation is close to the Arctic coastline in this sector, and corresponds to the Arctic frontal zone. As just discussed, many systems contributing to the summer cyclone maximum originate as waves along the Eurasian sector of this zone.

Building on Serreze et al. (2001), the migration of the 500-hPa vortex to over the pole and the development of the frontal zone both point to a role of differential atmospheric heating between the Arctic Ocean and snow-free land. While there is 24-h daylight over the Arctic Ocean in summer, the high albedo of the sea ice cover and the low-level stratus that dominate the region mean that much of the downwelling solar radiation is returned back to space. Furthermore, as shown in the recent Arctic energy budget analysis by Serreze et al. (2007), based on data from ERA-40, there is a strong downward net surface heat flux in July of about 100 W m−2, and of 75 and 45 W m−2 in June and August, respectively (a flux from the atmosphere into the underlying ocean), for the Arctic Ocean as a whole. This flux is associated with melt of the sea ice cover and (primarily over coastal open water areas) seasonal replenishment of oceanic sensible heat storage. For July, the net surface flux is actually larger that the horizontal convergence of atmospheric energy into the Arctic Ocean region. It follows that the 500-hPa vortex will want to shift to over the cold Arctic Ocean with broadly symmetric flow around it.

Cyclones migrating into the cyclone maximum region, along with those generated over the region itself, will tend to track around the vortex and eventually occlude. This is consistent with the distribution of cyclolysis events (Fig. 6). On the North American side, the high-latitude 500-hPa flow has a strong southward component, which tends to steer systems away from the Arctic Ocean. This aligns with the observation that relatively few systems found in the summer cyclone maximum originate from the North American side.

In the mean, the July SLP field is rather flat, with a weak low centered almost directly under the core of the 500-hPa vortex (Fig. 7d). For summer averages as a whole, there is no obvious surface feature. As discussed below, however, strong mean lows are found for individual months when the cyclone pattern is well developed.

4. Variability in the cyclone pattern

a. Time series

Time series of the counts of cyclone centers within the cyclone maximum for June, July, August appear in Fig. 8. The most obvious feature is strong variability. For example, June counts range from a low of 0 in 1965 to a high of 81 in 1989. For the summer as a whole, there is almost a factor of 3 range from a low of 39 in 1965 to high of 151 in 1994. Some years show fairly high counts for all three summer months—2002 serves as an example. Others, such as 1965, show below-average counts for all months. More typically, however, there is a mix of high and low counts between the summer months. There are no obvious trends, only a suggestion of more high total summer counts since the early 1980s.

b. Composite fields and NAM links

Composite analyses reveal distinct differences in the regional- and large-scale circulation associated with months of strong versus weak cyclone activity. The period of 1958–2005 yields a total of 144 summer months. Figures 9 and 10 are composites of 500-hPa height and SLP for the 21 months with the highest and lowest counts of cyclone centers, respectively, in the summer cyclone maximum region. This equates to the top and bottom 15% of the distribution.

In the strong composite (based on months with the highest cyclone center counts) high-latitude winds from about 60° to 270°E, as inferred from the 500-hPa height field, are fairly strong about a symmetric polar vortex (Fig. 9a). Over Eurasia, there is a pronounced Urals trough. At sea level (Fig. 9b) there is a distinct mean low of about 1005 hPa underlying the center of the 500-hPa vortex, which is an equivalent barotropic structure. This is in sharp contrast to the weak composite. High-latitude winds at 500 hPa are relatively slack, and the polar vortex in high latitudes is asymmetric (Fig. 10a). The relevant feature at sea level is a pronounced anticyclone with a central pressure of about 1017 hPa centered over the Beaufort Sea (Fig. 10b). The center of the high is just east of a weak 500-hPa ridge.

Fields of composite differences (strong minus weak) are also revealing. Figure 11 is the composite difference field of 500-hPa height. There are pronounced height differences centered over the Arctic Ocean, just south of the pole; the maximum composite difference is about 135 gpm. This is paired with differences of opposing signs over lower latitudes, most prominent over the North Pacific (60 gpm) near Hudson Bay (45 gpm) and north-central Siberia (45 gpm). This difference pattern has an equivalent barotropic structure. The composite difference field of 500-hPa temperature (not shown) has a very similar structure, and the same basic pattern extends to higher-tropospheric levels and the surface. Figure 12 is the composite difference field of 925-hPa temperatures, which is a useful level that captures lower-tropospheric conditions and (at least in the Arctic) relates fairly strongly to surface temperature anomalies. Figure 13 provides the composite field of SLP. Putting these results together, it is evident that variations in the strength of the summer cyclone pattern are part of the large-scale mass oscillation between the Arctic Ocean and lower latitudes.

The patterns shown in Figs. 11, 12 and 13 resemble the summer signature of the NAM discussed earlier with reference to the study of Ogi et al. (2004). Figure 14 shows the composite difference field of 500-hPa height based on the 21 most positive and 21 most negative normalized NAM index values for the summer months over the period of 1958–2005. It can be usefully compared to the composite difference field based on the same number of cases from the cyclone center count time series. The similarity is obvious. In turn, the distribution of cyclone center counts based on months comprising the positive NAM composite has a clear maximum over the central Arctic Ocean near its climatolological location as shown in Fig. 1, while the negative composite does not (Fig. 15).

The NAM link with cyclone activity is nevertheless not entirely robust. There are differences between Fig. 14 and Fig. 11, notably that the negative composite differences in 500-hPa height over high latitudes based on the cyclone time series are more strongly focused over the Arctic Ocean. In the positive composites, this is manifested as a somewhat more symmetric high-latitude vortex. Of the 21 months with the highest cyclone counts in the central Arctic maximum, all but 2 are paired with a positive NAM index. Of the 21 months with the lowest cyclone counts, all but 3 are paired with a negative NAM index. However, the linear correlation between time series of the summer NAM and cyclone counts in the central Arctic maximum region (144 months) is a rather modest 0.55. In this analysis, the cyclone center counts were first converted to z scores based on the mean and standard deviation for each month, from June through August. This modest correlation is not surprising. Cyclones center counts are a noisy secondary component of the circulation. Furthermore, because the cyclone time series is based on counts over a fairly small region, it can be sensitive to even small month-to-month variations in the location of maximum cyclone activity. Regarding the latter, the correlation of the NAM index against a cyclone time series based on counts for the larger region north of 75°N rises to 0.73.

c. Case studies

To round out the discussion of variability, case studies are presented of two sharply contrasting months. The first is for June 1989, when the cyclone pattern was very strongly expressed. The NAM was also in a strongly positive phase, with an index value of +2.31. The second case study is for August of 1999; when the cyclone pattern was extremely weak and the NAM was modestly negative (index value of −0.84).

Fields of monthly mean 500-hPa height, SLP, and 500-hPa Eady growth rates (calculated from daily mean data, see section 2) for June 1989 are provided in Fig. 16. Consistent with earlier results, the 500-hPa height field (Fig. 16a) shows a strong vortex centered near the pole and a strong Urals trough shifted eastward from its mean position. The SLP field (Fig. 16b) features a pronounced mean low with a central pressure of 1000 hPa. Pointing to the equivalent barotropic structure in the monthly mean, the center of the surface low almost directly underlies the 500-hPa low. Higher latitudes are characterized by a band of fairly strong Eady growth rates extending from Greenland and across northern Eurasia to eastern Siberia that have obvious correspondence to the 500-hPa height gradients. Over the central Arctic Ocean itself, near the center of the 500-hPa vortex, growth rates are small (Fig. 16c).

Daily fields for June 1989 show a series of lows moving into the central Arctic Ocean generated over Eurasia and along its coast where the Eady growth rates are strong (as a reflection of the Arctic frontal zone), which often deepen in their passage. At 500 hPa, a closed low persisted over the Arctic Ocean for the entire month, meandering about the region. On a daily basis, it was typically highly asymmetric, and the surface lows developed and deepened ahead of the troughs. In this particular case study, surface temperatures over the central Arctic Ocean averaged for the month are near normal, contrasting with other months with a strongly developed cyclonic system, such as August 1980, when surface temperatures were strongly below norms.

Monthly fields of 500-hPa height, SLP, and Eady growth rates for a case study of weak cyclone activity appear in Fig. 17. Compared to June 1989, high-latitude winds are slack, and there are separate closed 500-hPa lows (Fig. 17a). The surface circulation (Fig. 17b) is characterized by an anticyclone, centered just off the pole at about 120°E, with a maximum pressure of 1018 hPa. Surface temperature anomalies over the Arctic Ocean are mostly positive (not shown). Eady growth rates are small over most of the Northern Hemisphere extratropics (Fig. 17c).

5. Discussion and conclusions

Based on output from a cyclone detection and tracking algorithm applied to NCEP reanalysis data, most of the individual lows contributing to the summer cyclone maximum of the central Arctic Ocean are imported from outside of the region. While many are generated over the Eurasian continent, others form outside of the region over the Arctic Ocean. Pointing to the impact of the Arctic frontal zone, many of the latter can be identified as preexisting tropospheric waves along the Eurasian coast. The North Pacific, the North Atlantic, and North America are insignificant as source regions. Cyclogenesis may also occur within the cyclone maximum region itself.

Seasonal onset of the frontal zone is allied with a tendency for the summer 500-hPa circumpolar vortex at high latitudes to become broadly symmetric about the pole. Building on earlier studies, this argues for a role for differential atmospheric heating between the land and Arctic Ocean in the development of the summer cyclone maximum. In turn, there is an eastward shift of the Urals trough, and the flow ahead of it becomes more zonal than in winter. Systems entering the central Arctic Ocean from the outside, or forming within the Arctic Ocean, migrate around the 500-hPa vortex, and tend to decay within the cyclone maximum region or in fairly close proximity. Not examined in our study are changes within the summer season that may affect the pattern. While the summer stratospheric circulation is anticyclonic, LeDrew et al. (1992) note that the transition back to a cyclonic stratospheric circulation tends to occur around mid-August, increasing the potential vorticity that is available for surface development.

The strength of the summer pattern is highly variable from year to year. When well developed, the 500-hPa vortex is especially strong and symmetric. Negative SLP anomalies centered near the pole are attended by positive anomalies over middle latitudes. When poorly developed, the 500-hPa vortex is asymmetric, and positive SLP anomalies over the central Arctic Ocean are paired with negative anomalies in middle latitudes. This pattern resembles the summer northern annular mode (Ogi et al. 2004).

As noted in the introduction, links have been observed between the cyclone pattern and variability in sea ice conditions, as well as net precipitation (PE) over the Arctic Ocean. While the reader is referred to papers cited in the introduction regarding the former, it is useful to briefly summarize the PE link here.

To this end, Fig. 18 shows the composite difference field of aerologically calculated PE for summer from ERA-40, based on months of strong and weak cyclone activity over the years of 1958–2001. The PE is computed from the aerological budget, that is, through adjusting the vertically integrated vapor flux convergence by the time change in column water vapor. Fields of PE from ERA-40 are preferred over those from the NCEP reanalysis, because those in the latter contain spurious noise associated with topography (Serreze et al. 2003a). Given the shorter record, the composites are based on the same cases used earlier, except for eliminating August 2005 (in the strong composite) and June 2005 (in the weak composite). Hence, each composite is based on 20, rather than 21, cases. The results are quite clear—when the cyclone pattern is strongly (weakly) developed, PE tends to be more positive (less positive) over most of the Arctic Ocean, with the strongest signal in the Beaufort Sea. Simply put, enhanced cyclone activity promotes water vapor convergence and uplift. Other regions of the Arctic show a mix of positive and negative differences. Not surprisingly, results based on the summer NAM index (not shown) are similar.

While there have been no trends in the strength or persistence of the summer cyclone pattern over the period of 1958–2005, it is natural to speculate on its future behavior. Climate models are in near-universal agreement that Arctic warming in response to greenhouse gas loading will be especially strong. Results from the present study suggest that, at least in part, the summer cyclone pattern owes its existence to differential atmospheric heating between the Arctic Ocean and snow-free land. If patterns of differential heating change substantially, such as through earlier springtime loss of snow cover over land, or through changes in the presently strong summer net surface heat flux over the Arctic Ocean as the sea ice cover disappears, this may invoke changes in the summer circulation.

One must also consider links with changes in the winter circulation. Ogi et al. (2003, 2004) offer evidence that a positive winter phase of the NAM tends to be followed by a summer circulation of the positive phase, and vise versa. Surface boundary conditions provide a plausible link. For example, a positive winter phase of the NAM fosters less snow cover over the Arctic coast of Eurasia and North America, which can then enhance the summer thermal contrast between the land and Arctic Ocean, favoring a positive summer NAM by triggering enhanced eddy activity along the Arctic frontal zone (Ogi et al. 2003). There is also evidence, albeit still controversial, that despite recent regression of the winter NAM to neutral conditions, higher concentrations of atmospheric greenhouse gases may favor the positive mode (e.g., Gillett et al. 2003; Kuzmina et al. 2005). If so, expressions in the summer circulation can be expected. Further clarifying the role of differential atmospheric heating on the Arctic’s summer circulation and linkages with the NAM seems well suited to sensitivity studies via coupled global models.

Acknowledgments

This study was supported by NSF Grants ARC-0531302, ARC-0531040, OPP-0240948, ARC-0229651, and OPP-0138018, and NASA Contacts NNG04GH04G, NNG04GJ39G, and NNG06GB26G. M. Ogi is thanked for the summer NAM index time series.

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Fig. 1.
Fig. 1.

Counts of surface cyclone centers in equal area boxes of 250 km × 250 km for (a) winter [December–February (DJF)] and (b) summer [June–August (JJA)] totaled over the period of 1958–2005 for the region north of 60°N. Results are based on NCEP reanalysis data. Fields have been smoothed by first averaging counts for each grid cell, along with counts at the four adjacent cells, and then applying a nine-cell center-weighted average. Bold contours indicate regions with at least 135 cyclone centers. Due to spurious cyclones associated with reduction of surface pressure to sea level, Greenland is masked.

Citation: Journal of Climate 21, 5; 10.1175/2007JCLI1810.1

Fig. 2.
Fig. 2.

The annual cycle of number of cyclone centers within the 135 count contour over the central Arctic Ocean depicted in Fig. 1b is shown. Values represent sums over the period of 1958–2005.

Citation: Journal of Climate 21, 5; 10.1175/2007JCLI1810.1

Fig. 3.
Fig. 3.

Latitude by height (pressure) cross section of the mean zonal wind for July, averaged over the period of 1958–2005 and for 60°–270°E longitude.

Citation: Journal of Climate 21, 5; 10.1175/2007JCLI1810.1

Fig. 4.
Fig. 4.

Gridcell counts over the period of 1958–2005 of events of cyclogenesis for systems that at any point in their life cycle were found within the central Arctic summer cyclone maximum. The colors of the circles correspond to the number of events. Greenland is masked.

Citation: Journal of Climate 21, 5; 10.1175/2007JCLI1810.1

Fig. 5.
Fig. 5.

Same as Fig. 4, but for events of maximum cyclone deepening.

Citation: Journal of Climate 21, 5; 10.1175/2007JCLI1810.1

Fig. 6.
Fig. 6.

Same as Fig. 4, but for events of cyclolysis.

Citation: Journal of Climate 21, 5; 10.1175/2007JCLI1810.1

Fig. 7.
Fig. 7.

Fields of mean (left) 500-hPa height (gpm) and (right) sea level pressure (hPa) over the period of 1958–2005 for (a), (b) January and (c), (d) July.

Citation: Journal of Climate 21, 5; 10.1175/2007JCLI1810.1

Fig. 8.
Fig. 8.

Time series of 1958–2005 of counts of cyclone centers for summer (total length of bars), and broken down for June (black bars), July (gray bars), and August (lightly shaded bars), for the region defining central Arctic Ocean cyclone maximum.

Citation: Journal of Climate 21, 5; 10.1175/2007JCLI1810.1

Fig. 9.
Fig. 9.

Composite means of (a) 500-hPa height (gpm) and (b) sea level pressure (hPa) for the 21 months when the summer cyclone pattern over the central Arctic Ocean was most strongly expressed.

Citation: Journal of Climate 21, 5; 10.1175/2007JCLI1810.1

Fig. 10.
Fig. 10.

Same as Fig. 9, but for when the summer cyclone pattern was most weakly expressed.

Citation: Journal of Climate 21, 5; 10.1175/2007JCLI1810.1

Fig. 11.
Fig. 11.

Composite differences in 500-hPa height (gpm) between the 21 months when the summer cyclone pattern over the central Arctic Ocean was most strongly and most weakly expressed (strong minus weak).

Citation: Journal of Climate 21, 5; 10.1175/2007JCLI1810.1

Fig. 12.
Fig. 12.

Same as in Fig. 11, but for the composite differences in 925-hPa temperature (°C).

Citation: Journal of Climate 21, 5; 10.1175/2007JCLI1810.1

Fig. 13.
Fig. 13.

Same as in Fig. 11, but for the composite differences in sea level pressure (hPa).

Citation: Journal of Climate 21, 5; 10.1175/2007JCLI1810.1

Fig. 14.
Fig. 14.

Composite differences in 500-hPa height (gpm) between the 21 months when the summer NAM was most strongly positive and most strongly negative (positive minus negative).

Citation: Journal of Climate 21, 5; 10.1175/2007JCLI1810.1

Fig. 15.
Fig. 15.

Summer cyclone center counts for the region north of 60°N for the 21 months when the NAM was (a) most strongly positive and (b) most strongly negative. Fields have been smoothed as in Fig. 1. Greenland is masked.

Citation: Journal of Climate 21, 5; 10.1175/2007JCLI1810.1

Fig. 16.
Fig. 16.

Fields for June 1989 of (a) 500-hPa height (gpm), (b) sea level pressure (hPa), and (c) 500-hPa Eady growth rates (day−1); regions where the growth rate exceeds 0.5 day−1 are shown in bold contours.

Citation: Journal of Climate 21, 5; 10.1175/2007JCLI1810.1

Fig. 17.
Fig. 17.

Same as Fig. 16, but for July.

Citation: Journal of Climate 21, 5; 10.1175/2007JCLI1810.1

Fig. 18.
Fig. 18.

Composite differences in aerologically calculated net precipitation (PE) from ERA-40 between months when the summer cyclone pattern over the central Arctic Ocean was strongly and weakly expressed (strong minus weak).

Citation: Journal of Climate 21, 5; 10.1175/2007JCLI1810.1

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