Europe, Climate of

Analysis of climatic data

Most climate data are assembled and kept in databanks held at regional climate centers or in research centers. The data are then analyzed using powerful statistical tools that enable the researcher to sarch for patterns and correlations. An example of one such data set is NACD — the North Atlantic Climatological Dataset containing temperature and precipitation series. Regional or subregional datasets are then compiled and adjusted to build a homogeneous dataset (Tuomenvirta, 2001). Other regional datasets are compiled for use in climate analysis computational programs covering a wide range of aspects from global warming through to hydrological management. This item has been compiled from a combination of descriptive sources and of analytical studies. In places there is some difficulty in linking together the climatic ensemble and this must be borne in mind by the reader. Many datasets are now available on websites.

Factors affecting the climate of Europe

The continent of Europe extends from the mid-Atlantic islands of the Azores (29°W) across the open North European Plain to the Ural Mountains of the Soviet Union (60°E) and from the islands of the Mediterranean Sea in the south (35°N) to the islands of the Barents Sea, which lie well within the Arctic Circle. Europe, along with its islands and peninsulas, is the most maritime of the continents. Large areas to the west of the 20°E longitude experience a temperate oceanic climate where extremes of temperature and rainfall are rarely experienced. The chief factors promoting this comparatively unique assemblage of climates are:

• 1. The predominant west-to-east movement of weather systems.

• 2. The extensive area of abnormally warm surface waters of the North Atlantic.

• 3. The virtual absence of north-south-aligned mountain ranges between latitudes 45°N and 60°N that might otherwise alter the nature of the westerly flow.

• 4. The presence of large inland seas such as the warm Mediterranean, the cooler Baltic, and the smaller, less warm, Caspian and Black seas.

Elements of the climate of Europe

Energy

In common with other midlatitude regions, Europe experiences a net deficit in radiation except at its southern margins. Values derived from satellite information for the Earth-atmosphere system indicate that annual deficits increase from −20 W m² over central Italy to around −80 W m² over southern Scandinavia and northern Norway (Raschke et al., 1973). The deficit cannot be attributed principally to the excessive loss of heat from the land surface, though this loss may be significant in cold snowy winters over central and eastern Europe. Compensation for this deficit in radiant energy is effected by the transfer of sensible and latent heat from lower latitudes. Radiative cooling is at its maximum at mid-tropospheric levels around latitude 60°N, and it is in this region that maximum cooling takes place in winter (Figure E6). The effects of this cooling on European climates are profound. In the long term such cooling produces latitudinal zones of temperature gradient in the atmosphere. These zones are associated with high winds and disturbed weather conditions so frequently experienced over Europe.

Figure E6

Change in the heat content of the troposphere from mid-November through mid-December in W/m². Average values for 1955–1959. Values are relatively small with maximum cooling representing 1.8% of the solar constant (after Crowe, 1971; Tucker, 1965).

Temperature

The influence of the warm waters of the Atlantic Ocean is most clearly seen in the winter pattern of surface temperatures (Junge and Stephenson, 2003). Average January values of 7°C over western Ireland decline steadily eastward to around minus 10°C in the vicinity of Moscow, producing a north-south thermal pattern between latitudes 45°N and 60°N. In summer the heating of the continental area leads to a more zonal thermal pattern with isotherms tilted WSE-ENE (see Figure E7). The 17°C isotherm for July (2005) now lies through southern England across to the southern Baltic Sea toward Leningrad. Summers in north European Russia, though warm, are usually short and are constantly being threatened by outbreaks of cold air of Arctic origin.

Figure E7

Mean temperature for January (solid lines) and July (dashed lines) in °C.

A measure of the increasing warmth from north to south through European Russia has been visualized using summer daily air temperatures that exceed 10°C. Accumulated summer temperature values of 750°C occur at around 70°N, increasing to 2000°C at 55°N and exceeding 3000°C around the Black Sea at 45°N (Budyko, 1974). In the cloudier climate of the British Isles these values are usually lower. They range from below 500 in northern England (55°N) and approach 1000 in the Channel Islands (49°N). Using this method the equivalent value for Moscow would be 722. Analysis of temperature data by Gerasimov (1964) showed that average daily temperatures rose above 10°C by 21 April in the Crimea, but not until July along the coastlands of the Barents Sea. The average date when temperatures have fallen below 10°C in the fall have been 15 of October for the Crimea and 11 August for North Russia. Using a slightly different definition of the growing season, Mitchell and Hulme (2002) show that for Central England there has been an increase in length of the growing season over the twentieth century of 28 days, much of this in the period 1980–2005. Recent research by Zhou et al. (2001) and Menzel and Fabian (1999) has confirmed that this is also true for most of Europe. A more detailed daily temperature dataset has been compiled for the ECA (European Climate Assessment) at De Bilt in The Netherlands by Klein Tank et al. (2002), and shows a good correlation with existing gridded data sets for Europe.

Frost-free period

Closely associated with the growing season is the average length of the frost-free period. Over the open plains of European Russia this increases from at least 75 days in the north to around 200 along the northern shore of the Black Sea. Further west, across the European Plain, the length of the frost-free period becomes less easy to determine because of the alternating airflows from cold and warm sources (Lednicky, 1985). Over lowland France mid-twentieth-century values have ranged from 160 days in Alsace in eastern France to 240 days along the coastlands of the Bay of Biscay (Garnier, 1954). In contrast, the determination of the frost-free period in mountainous countries such as Austria is largely a function of elevation. In the Austrian Alps the frost-free season at 800 m above mean sea level lasts for about 155 days per year. At a height of 1600 m this is reduced to 104 days, whereas above 2500 m spells of below freezing temperature may even be expected in cold summers. In the Iberian Peninsula the frost-free period is governed both by altitude and proximity to the Mediterranean Sea and Atlantic Ocean. Much of interior Spain may experience less than 250 days frost-free. Only the extreme southeast and southwest of the peninsula are likely to be entirely free of frost during most years.

Extreme values

Extreme values of temperature tend to reflect the local controls of topography as well as distance from the North Atlantic Ocean. Very low temperatures show a much greater variation in time and space than high temperatures. Very dry, cold air with little wind is required for temperatures to fall below −15°C. Such an occurrence is far more common in winter over countries such as Finland, where temperatures below −30°C are recorded nearly every winter. The lowest temperature indicated in the standard climatic tables is −51°C at Syktyvkar (61°40′N, 50°51′E) in northeast European Russia. Moscow, along with many stations in northern Russia and Finland, has experienced temperatures of −40°C at screen height. The penetration of cold air into parts of eastern Europe also means that very low temperatures have occurred at relatively low latitudes. This may be seen from cities such as Iasi (Rumania) at 47°N, recording −30°C and Sofia (Bulgaria) 42°30′N, −27.5°C. There were also reports of −26°C from several locations in Thrace (Greece) in December 1972. Only the southernmost part of Spain and a part of the Mediterranean island of Crete have absolute minimums at or above 0°C. Comparatively, maritime countries such as Spain experience frost in severe winters. Madrid recorded −10.1°C during the period 1901–1930, and Rome (Italy) reported −7.4°C. Over the British Isles, minimums of −10°C have been recorded along the western seaboard. Inland values of −25°C are very rare, but were registered in the Dee Valley of the Grampians in Scotland and in the West Midlands of England during the winter 1981–1982. An extensive climatology of extreme temperatures has been assembled by Yan et al. (2002).

In contrast, there is little variation in absolute maximums. Over much of northern Europe, temperatures have occasionally climbed into the low 30s (Jenkinson, 1985) whilst over the remainder of Europe temperatures between 35°C and 39°C have sometimes occurred, as in the summer of 2003 (Eden, 2003). Apart from isolated events, temperatures in excess of 40°C have only been recorded south of latitude 42°N. Absolute maximums of 45°C (Palermo, Sicily) and 47°C (Malaga, Spain) testify to occasional incursion of very warm desert air from the Sahara. Trends in minimum temperatures (Weber et al., 1997) over the period 1901–1990 for Central Europe, show significant increases across the plains, whereas little change was found across the mountains.

Continentality

Various indices have been formulated to express the degree of continentality over Europe. In a general sense the indices make use of the range of the mean temperature between the coldest and warmest months of the year, and this is recorded as the annual range of mean temperature. The range increases from the Atlantic coastlands eastward across Europe into Russia (Figure E8). A relatively simple index has been used by Tsenker (in Borisov, 1965), who proposed that K=A/φ where K is the index, A is the annual range and φ the latitude. Similar values are obtained from Conrad’s index (Conrad, 1946). More complex formulae appear in Russian literature, an example being that of Ivanov (1959). Using Conrad’s equation a value of 4 is obtained for western Ireland, 12.5 for London, 24 for Berlin, and 40 for Moscow. Indices of over 30 are also found in parts of the Mediterranean countries, especially where inversions of temperature occur in winter such as in North Italy (Milan: 32). A different index may be calculated for an individual year to describe the degree of Atlantic influence, such as by adding the number of days of westerly flow either using Lamb’s weather types or the North Atlantic Oscillation Index (Jones and Hulme, 1997).

Figure E8

Annual range of mean monthly air temperature in °C.

Precipitation

Annual precipitation

Values of annual lowland rainfall show little spatial variation in the zone stretching from the English Midlands through to the Urals of the USSR (Figure E9). The passage of fronts usually ensures moderate precipitation across the European Plain, excessive falls being rare and usually associated with slow-moving disturbances. A slow decline from 640 mm to 500 mm can be observed eastward, and is evidence of the ease with which moist Atlantic air can penetrate across Europe. A few areas such as Bydgoszcz (northwest of Warsaw) and coastal regions of Lithuania record annual values less than 500 mm. As a general rule, annual rainfall over lowlands declines both northward toward the Arctic and southward toward the Mediterranean. Everywhere the existence of hills, mountains, enclosed basins, and peninsulas complicates this simple pattern. Highest annual values are found in mountainous areas where valleys are open to the southwest or westerly airflow. A few localities have recorded in excess of 5000 mm mean annual precipitation for the period 1961–1990, notably south of Nordfjord (Norway) and Sprinkling Tarn (English Lake District) and possibly inland from Boka Kotorska near the Yugoslav-Albanian border. A value of 4000 mm may have occurred at Monchsgrat in the Bernese Alps, although heavy winter snow makes measurement difficult. It is probable that one or two localities in the Bavarian Alps (West Germany) and in the Sierra de Gredos (west of Madrid, Spain) receive 3000 mm, whereas 2000 mm would seem feasible for the highest parts of the Central Massif and Jura mountains in France.

Figure E9

Mean annual precipitation in cm.

Seasonal pattern

Invasions of moist air across Europe are most frequent when the circulation is vigorous and depressions are moving quickly east across Scotland and Scania into the Baltic region. Such a situation is most likely to occur in early winter and is least likely in spring or early fall when anticyclones are usually present over some parts of Europe. In fact, over much of Europe the percentage of annual precipitation falling in spring is remarkably constant, lying between 16% and 21% (Table E3). A winter maximum in rainfall might well be expected to occur fairly widely. That this is not so indicates the many other factors that influence the seasonal distribution. In particular, higher summer temperatures over much of central and eastern Europe can lead to enhanced convection. This shower activity may be random if pressure gradients are weak, or may be organized along cold fronts as incursions of cool maritime polar air from the Atlantic travel southeast over much of the region. As may be seen from five of the stations listed in Table E3, a summer maximum occurs over eastern England and central, eastern, and northern Europe. A zone of fall maximum occurs from Portugal through

Table E3 Seasonal precipitation as a percentage of the annual precipitation (1931–1960)

western France, central Britain, the coastlands of the Low Countries and Norway (see Bergen, Table E3). In a climatic context this maximum is the result of relatively high sea surface temperatures leading to upward transfer of water vapor into the boundary layer. The moisture may then be ingested into the circulation of midlatitude depressions via a conveyor belt mechanism (Barry and Chorley, 2003) These depressions become more vigorous and active from September through November across central and northern Europe.

The Mediterranean region with its mountains, peninsulas, islands, and inland sea presents a more complex pattern of seasonal precipitation (Figure E10). It is more convenient to state the driest period of the year, which is summer, when extensions of the subtropical high pressure lie over the region. During the remainder of the year, although frontal depressions can sometimes be identified, many disturbances are ill-defined and rather transitory. Moreover, their activity in terms of producing rain varies on both a seasonal and annual basis. One or two very wet Octobers can lead to an imprint on the climatic record that can be misleading if arithmetic averages are employed. Bearing this in mind, it is generally true that a fall maximum is more likely in the western and northern Mediterranean and that a winter maximum is more evident in the southern and eastern Mediterranean (see Athens, Greece, Table E3). A bimodal maximum (spring, fall) is a feature of much of Spain, at least when the westerly flow is relatively strong. Comprehensive analysis of seasonal rainfall frequency in the Mediterranean was carried out by Reichel and Huttary (in Trewartha, 1981). More recent investigations have employed indices and statistical analyses in attempts to describe present rainfall patterns and to search for regular periodicities (see Fukui, 1966; Tabony, 1981). It should also be noted that there is a tendency for the general characteristics of a rainfall regime to change over periods of 30 years. Thus fall may be considered “wet” in one decade, spring in the next (Boucher, 1994). More detailed gridded analysis of precipitation has recently been carried out for a 25 km grid over the entire region of the European Alps and then extended over the time period 1901–1990. This reveals an increase in winter precipitation of 20–30% per 100 years over the western part of the Alps and a decrease of 20–40% fall precipitation in the southern part of the Alps (Schmidli, 2002).

Figure E10

Seasonal distribution of maximum rainfall. The figure shows months of maximum rainfall over southern Europe based on World Meteorological Organization Data 1961–1990. Months are numbered from January (1) through December (12). Complex regimes with no distinct seasonal maximum are indicated by C.

Types of rainfall and duration of heavy falls

Attempts have been made to classify rainfall into three categories of cyclonic (predominantly frontal uplift), convective (thunderstorm type), and orographic (uplift over mountains). It is not easy to distinguish clearly between these categories, though convective thunderstorms contribute significant amounts to summer rainfall over much of Europe. Precipitation over mountainous areas is enhanced by uplift of airmasses on windward slopes. Extreme maximum precipitation over periods of less than 24 hours is highest at lowland stations where convection is dominant. Continuous heavy rain depends on steady inflow of moisture into a region usually associated with slow-moving low-pressure systems. Recent analysis of heavy rainfall events in Central Europe (Ulbrich et al., 2001, Ulbrich et al., 2003a, Ulbrich et al., 2003b) has shown that extreme rainfall periods have return periods of between 300 and 1000 years with rainfall totals over 5–10 days exceeding the monthly rainfall by 300%. Data from Ulbrich and others are shown in Table E4. Such widespread rainfall usually combines three components: uplift of moist air over mountains, vigorous wave activity along a quasi-stationary frontal zone, and convective instability. Two examples are given in Table E5, in which over 600 million cubic meters of rain fell over a limited area.

Thunderstorms and related phenomena

Over much of continental Europe, thunderstorms are typically a summer phenomenon. Within European Russia about 50% of the thunderstorms are associated with cold fronts, 22% with warm fronts, and 28% with convergence within airmasses. Thunderstorm frequency increases with the rise in surface temperatures in early summer and reaches a maximum in June in parts of southern European Russia; in July over much of central and northern Europe, including northern Italy; in August in parts of Sweden; in September and October in the western

Table E4 Maximum precipitation sample extreme values-Europe 1960–2005

Table E5 Data relating to two notable rainfall events with return periods in excess of 100 years

Mediterranean and in winter over southern Italy as well as parts of northwestern Britain.

The occurrence of hail is most closely associated with airmass thunderstorms, especially where these develop in hilly or mountainous regions. They may be accompanied by squally winds that are best attributed to storm downdraughts. Small-scale tornadoes occasionally accompany storms, but nowhere have they been reported to attain the magnitude and frequency of those in the United States. Over western Europe they may be closely associated with cold fronts at any time of the year. Their tracks rarely exceed 1 km. There have been fairly frequent sightings of water spouts over the western and central Mediterranean during thundery weather. Small-scale whirlwinds and tornadoes occur fairly regularly during summer months.

Snowfall

The relatively high sea surface temperatures in winter mean that snowfall in winter is variable over low ground in western Europe. Snow cover may persist for more than 7 days only in the coldest winters. As a general rule, monthly average temperatures must lie below 1°C for snow cover to persist. The number of individual days with complete snow cover increases northeast across the European Plain from 5 days in the Midlands of England to 40 over central Poland, 135 days around Moscow and up to 200 along the Arctic fringe. The mean maximum depth of snow cover increases from 20 cm in southwest Russia to about 75 cm on the lowlands to the west of the Urals. Most mountain areas are liable to receive heavy snowfall, though aspect and character of individual winters along with recent rise in temperatures have reduced snow accumulations. Prolonged snowfall exceeding 25 cm over parts of western Europe is only likely to the north of an active frontal zone. Such a situation leads to very cold continental air being drawn west over northern Germany toward central Britain, while humid air is forced to ascend over the cold easterlies. Snow enhancement occurs if small disturbances run east, typically along the English Channel into central Germany. Polar lows occasionally bring blizzards to Scotland, and eastern parts of the North Sea, but their activity depends on unstable cyclonic disturbances developing in cold northerly airflows crossing warm ocean waters.

Drought

Lack of precipitation is a normal part of the seasonal climatic cycle within the Mediterranean in summer. An appropriate index is used to describe the rainless period. Aridity indices, such as that of De Martonne, have traditionally been employed by Italian and French climatologists to describe summer drought (Pinna, 1957). The existence of a rainless period is closely associated with midtropospheric subsidence. Therefore the existence and maintenance of summer drought depend on the strength and position of the Azores anticyclone and its eastward extension. When this anticyclone is well developed, drought is complete and may last for up to 4 months in southern parts of Spain, Italy, and Greece. Occasionally during summers such as those of 1959, 1976, 1992 and 2003 the belt of high pressure may be located farther north over northern France and the North Sea, bringing drought to areas usually well supplied with precipitation and delivering summer rainfall to the Mediterranean (Perry, 1976). Such annual variations are of deep concern, not only to countries such as Spain and Portugal (Estrela et al., 2000), but also to the grain-growing areas of Poland and Russia where the failure of adequate snowmelt and spring-summer rain may seriously reduce yields, as in 1972 and 2003. Hydrological droughts (low streamflow) have continued to be cause for anxiety (EEA, 1998) but show both spatial and temporal variations (Lloyd- Hughes and Saunders, 2002). Increasing drought deficit volumes have occurred in the eastern sector of East Europe and large parts of the UK, but decreasing deficits have characterized many parts of central Europe (Hisdal et al., 2001).

Severe droughts are associated with rainfall deficiencies of 50% over periods of 6 months or more. The impact of such droughts is greatest if they occur during the soil moisture recharge period (Hounam et al., 1975). As elsewhere, the intensity of the drought is increased by high winds, low humidity, and high temperatures. Areas likely to suffer such conditions lie along the southern border of Europe and across the steppes of the Ukraine and Kazakhstan (north Caspian Sea). In the latter area the steppe climate is characterized by extreme annual range of temperature approaching 80°C, frequent droughts, and desiccating, dust-laden winds known as sukhovei. During the period 1880–1950, serious drought occurred on average every 3.5 years (Poltarus, in Borisov, 1965). The effect of one year’s drought is magnified by the tendency of one drought year to follow another. Such was the case over much of European Russia from 1889 through 1892, and in southern Spain from 1917 to 1919.

Within the midlatitude belt of Europe the occurrence of drought is closely associated with the position, height and persistence of ridges in the middle part of the troposphere. Persistence of high pressure tendency over a number of years may produce general deficits in precipitation. The droughts of the early 1970s in European Russia prompted a number of studies (Buchinskiy, 1975; Chistyakova, 1975). In an earlier investigation Davitaya (1958) found that droughts over southern Russia reflected shifts in the midtropospheric ridge-trough pattern. A thorough investigation by Rauner (1980) into the simultaneous recurrence of drought in different regions of Europe revealed that drought affected the whole of the grain zone of the USSR in 5 years during the period 1891–1975. On a regional basis droughts were found to have affected the Ukraine in 43 out of 106 years (1870–1976) and the Volga region, a 1000 km further east, in 41 out of 96 years (1880–1976) of which 19 years were common to both regions. The occurrence of drought simultaneously in European Russia and Western Europe evidently has taken place in 20 years during the period 1700–1976, three of which occurred consecutively from 1747 to 1749.

Climate change

Rising surface temperatures in the European sector have been shown by Parker and Alexander (2002) whilst their annual analysis of rainfall reveals high variability on time scales of 10 years with no easily discernible trends.

Synoptic climatology

Weather types

Regional assessment of climate has often been carried out on the basis of a classification of air movement and associated weather conditions. British research (Lamb, 1972) has favored a less rigorous system than a number of German and Russian climatologists (see Hess and Brezowsky, 1969; Chubukov, 1977). Despite its limitations, an example of such an annual regime is included. Other circulation indices include that of Namias (1950) and some variants used by Dole and Gordon (1983). Research has also focused on indices that affect part of, or all of, the northern hemisphere, such as the North Atlantic Oscillation (NAO). This fluctuation in pressure gradient between the Icelandic low (65°N) and the Azores high (40°N) is believed to influence weather patterns over much of Europe in winter (van Loon and Rogers, 1978). Since 1989 the NAO has been mostly positive, leading to warmer winters over many parts of Europe (Barry and Chorley, 2003).

Spells of weather

The term “singularities” has been applied to periods of distinctive weather lasting between 7 and 14 days that have a tendency to recur in most years on or about the same dates. Trewartha (1981) associates this calendar of weather episodes with the Grosswetterlagen of German climatologists. For a certain region of Europe a calendar sequence of probable types of weather may be assembled. In essence, singularities represent adjustments of the general circulation to various distributions of energy within the boundary layer and in the troposphere above. A singularity such as an anticyclonic spell may be regarded as a statistical probability of a circulation pattern occurring for a given period of the climatic record. An example of such a calendar is presented in Table E6, which refers chiefly to the British Isles. Most singularities shown also occur over Germany; however, the region of European Russia lies too far downstream in the westerlies for statistical correlations to be significant.

Due to the ridge-trough nature of the air flow at 5000 m periods of anticyclonic activity over western Europe are likely to be associated with cyclonic activity over eastern Europe. It might appear from Table E6 that periods of cyclonic activity alternate

Table E6 Abbreviated calendar of singularities

with anticyclones, but such a deduction can be misleading. The singularity table is not a forecasting tool. For reasons not well understood, certain singularities become statistically significant for limited periods of the record before declining, even disappearing, for a number of years only to re-emerge at a later period. It is interesting to postulate that such changes form part of a climatic signal in the atmospheric system, indicating fundamental adjustments of the general circulation (see Lamb, 1982).

The general circulation

Pressure patterns at the surface

The mean centers of both anticyclonic and cyclonic activity lie over the Azores islands, west of Portugal, and southwest of Iceland respectively — outside continental Europe. In periods when the westerly flow is well established, ridges and troughs move eastward away from these centers across central Europe. These may be associated with closed isobaric highs and lows at the surface. On the average monthly mean pressure charts for 1951–1966 (Meteorological Office, 1975), the shape of the North Atlantic low changes, being at its most intense in December (997 mb); thereafter it declines until July when there is no easily identifiable center of low pressure close to Europe. The Azores anticyclone appears weakest in March (1020 mb) and lies equatorward of 30°N. It reasserts its dominance and is at its most extensive in July (1025 mb) with ridges extending east and northeast across Europe (Figure E11). It has been noted that July is also the wettest month over much of Europe, indicating that breakdowns in the high pressure are frequent in some summers. The greatest monthly variation in pressure during the period 1951–1966 occurred west of Ireland in February and November, with standard deviation values of 11 mb and 8mb, respectively. This indicates marked changes in flow patterns in these months from one year to another, indicating the variable nature of West European climate (Table E7 and E8).

Figure E11

Normal sea level pressure in mb for (A) January and (B) July (after Meteorological Office, 1975, pp. 110–111).

The jet flow across europe

Mean airflow over Europe in the troposphere is predominantly from the west between high pressure to the south and low pressure to the north. Imposed on this pattern are waves of large amplitude having a wavelength of several thousand kilometers. Their position, movement, and intensity produce an ever-changing pattern of airflow. Embedded in this flow lie narrow zones of fast-moving air known as jet streams that extend downstream for several thousand kilometers. They may be about 400 km wide, and are associated with airmass boundaries of the Polar Front. They appear and disappear with the changes in the circulation pattern. Within the core of the Polar Front jet, winds may exceed values of 60 m/s over Europe.

It is not possible to represent a mean position of the Polar Front jet, but some idea of its strength and location at 700 mb is given in Figure E12, which shows the axes of maximum wind at around 3000 m. The direction and movement of midlatitude depressions in any one month are closely associated with the mean position of the Polar Front jet during that month. The Subtropical jet is far more persistent. Its mean winter position lies over the central Sahara. Figure E13 shows cross-sections of the average zonal wind components in January and July: (1) 40°E; (2) 15°E; and (3) 10°W. The position of the Polar

Table E7 Meridional mean pressure difference 40°N to 60°N in mb at the surface

Table E8 Meridional mean difference in the height of the 500mb level in geopotential meters between 40°N and 60°N

Front jet is purely schematic. In winter it may be found anywhere between 30°N and 70°N, but it is too ephemeral to imprint its presence on average maps of zonal flow. In summer only one jet flow is recognizable at 300 mb over Europe even if there is a suggestion of two axes at 700mb.

Figure E12

Envelopes showing the position and strength of the monthly mean maximum geostrophic wind at 700 mb for winter (December to February) and summer (June to August) along the Greenwich Meridian. Plots for individual months are also shown for 1969–1975 (partial record).

Figure E13

Meridional cross-sections at 10°W, 15°E, and 40°E showing the westerly (zonal) component of the wind in meters per second. Positive speeds represent westerly wind, negative values easterly wind. Schematic position of the tropopause (Tr) is also shown (after Meteorological Office, 1975, p. 121). PFJ indicates the likely position of the Polar Front Jet.

Blocking patterns

One of the most frequent interruptions to the westerly flow over Europe results from “blocking” high-pressure patterns developing either upstream in the mid-North Atlantic or over Europe. The presence of a large anticyclone, persisting on average for 16 days, causes disruption of and breakdown in the mid-level flow. This blocking may be dominant in some years (such as 2003) or only weakly developed in most other years. The effect of blocking on the distribution of precipitation is potentially important. Abnormally low temperatures in winter may also result, as in February 1986. Much attention has been given to these features in the literature. Rex (1950) has provided a comprehensive review of the climatic data relating to blocking over Scandinavia.

Obstruction to westerly flow below 3000 m also results from the development and persistence of a cold anticyclone or thermal high over snow-covered Siberia. The westward extension of this high into European Russia causes continental temperate air to flow around the southern margin of the high in winter. On average, air associated with the Siberian high is found on 24 days in January over the lower Volga Plain, decreasing to 12 days in the vicinity of Leningrad (Lydolph, 1977).

Surface anticyclones and depressions

Anticyclones

Throughout the year, “cells” may break away from the semipermanent Azores anticyclone over the east central North Atlantic and travel slowly northeast, close to the English Channel and across the European Plain or central Europe into southern European Russia. Two to four new anticyclones a year tend to form along this axis. Another favorite center for anticyclogenesis lies over Scandinavia, while high values over the Mediterranean represent the relocation and development of the Azores high. Annual mean speeds of eastward (progressive) moving highs are of the order of 6° of longitude per day at 55°N. Somewhat surprisingly, anticyclones moving westward (retrogressing) away from Europe traveled at 7° per day on average (Meteorological Office, 1975) and accounted for 40% of mobile anticyclones. There was little significant seasonal pattern evident. Figure E14 shows the total number of days in which the center of a high was located within a 5° × 5° grid. Counts were made daily. Climatological interest in this empirical aspect of climatology has since declined.

Figure E14

Total number of days with anticyclonic centers at 1230 GMT, 1899–1938 (after Meteorological Office, 1975, p. 63).

Midlatitude depressions

In European literature the term “cyclones” usually is applied to tropical storms. Midlatitude cyclonic disturbances are referred to as depressions or lows. Frontal depressions affect much of Europe north of 45°N throughout the year, and parts of the Mediterranean during the winter half. A wide variety of nonfrontal depressions also occur and include heat lows (thunder lows), polar lows (Businger, 1985), and lee depressions. Most of these are associated with infrequent synoptic weather patterns. Major depressions that develop central pressures below 980 mb originate over the western North Atlantic and attain maximum intensity before reaching land. This is evident from Figure E15, which shows the diminishing influence of low pressure southeastward into central Europe. A feature of great importance over the sea areas bordering northwestern Europe is the ability of some lows to suddenly deepen, creating high winds and storm surges as in 1987.

Figure E15

Total number of days with low pressure centers at 1230 GMT, 1899–1938, adjusted to unit area size (after Meteorological Office, 1975, p. 46).

Figure E16 shows the areas most likely to experience cyclogenesis. It should be noted that the map does not convey any dynamic information about the development, movement, and intensity of the lows (Campins et al., 2000). For example, the center over northern Italy reflects the formation of many shallow but active lee depressions (see Zenone and Lecce, in Wallén, 1977). Other “centers” are less geographically fixed and do not appear on all map interpretations of the data (see Trewartha, 1981; Borisov, 1970).

Figure E16

Twenty-year frequency distribution of cyclogenesis, 1909–1914 and 1924–1937 (after Meteorological Office, 1975, p. 57 and Trewartha, 1981).

The movement of depressions over the northern hemisphere has been investigated by Klein (1957) and others, and is shown in Figure E17. The continuous lines denote main tracks and the dashed lines indicate less frequent routes. The main Icelandic low is renewed by depressions originating further west. Tracks across northwestern Europe are usually followed by subsidiary lows circulating around the parent depression, although some do break away beneath the westerly jet and travel across the Barents Sea. Most European lows either follow the northern or southern boundary of the continent in winter, whereas in summer a central route across the North and Baltic seas has been established. Over European Russia, tracks are difficult to identify in winter when high pressure predominates, but low pressure is more in evidence in summer. This is particularly so in central European Russia — a region that stretches from the western borders with Poland eastward toward the southern Urals. In this region there are twice as many cyclonic centers in summer as in winter. Some of these lows will be associated with fronts and outbreaks of cool polar air. Others are more easily identified with zones of marked baroclinicity (Ulbrich et al., 2001) sometimes associated with unusual pressure features (Pearce et al., 2001; Barry and Chorley, 2003).

Figure E17

Main tracks of traveling depressions (after Klein, 1957, p. 15).

Depressions in the mediterranean

It is unusual for the whole of the Mediterranean, which extends 3700 km eastward from Gibraltar, to be free from depressions at any one time, even in summer. When depressions form within or travel through the Mediterranean, there is often a quasistationary anticyclone over central or western Europe. Other factors contributing to cyclogenesis in the Mediterranean are the presence of a baroclinic or frontal zone, lee effects in strong airflows, and instability in airmasses (Meteorological Office, 1975). Analysis of upper-air charts shows that depressions travel in the direction of the flow at 200 mb, whereas thunderstorms are more likely to follow the flow pattern at 700mb. On the other hand, Perry (1981) has drawn attention to the role of sea surface temperature anomalies in promoting cyclonic activity within the Mediterranean.

Cold pools

During periods of low index and meridional flow masses of cold air may be transported southward over Europe and may become entirely surrounded within the troposphere by relatively warm air. The severing of the tip of this tongue of cold air produces a tropospheric cold pool. When this occurs over the Mediterranean it becomes the source of instability, thunderstorms, and abnormal precipitation (Boucher, 1982). One notable cold pool and accompanying storm developed in late January 1986 with central pressure falling below 980 mb.

Conclusion

The ever-changing pattern of the climate of western Europe provided the stimulus for the development of the science of meteorology in the early part of the twentieth century — notably the establishment of the Bergen School founded in 1918 by the Norwegian physicist Vilhelm Bjerknes (Jewell, 1981). This day-to-day variability has posed a challenge for climatologists who seek to establish order out of seeming chaos (see Volkert, 1985 and Lamb, 1985). Others have turned their attention to archived records across Europe in an attempt to place the climate records across Europe into a meaningful sequence (New et al., 2001). Elsewhere the emphasis on classification and indices has enabled greater understanding of climatic processes to be combined with issues such as that of global warming. Even so, the quantitative expression of the regional characteristics of climate remains a daunting task for climatologists and statisticians alike (Trewartha, 1981).