a. An isobaric geopotential height in σ-coordinates

The vertical coordinate is defined by σ = p/p∗, where p∗(x, y, t) is the surface pressure. The vertical distribution of variables is shown in Fig 2a. In the vertical, 16 σ-levels at σ = 0.025, 0.075, 0.125, 0.170, 0.240, 0.325, 0.430, 0.550, 0.610, 0.670, 0.775, 0.855, 0.915, 0.955, 0.980, and 0.995 are used. Let U, V, Ω, and D be

Uum,Vνm,ζm²Dδm²(1)

where m is the map scale factor, δ the horizontal divergence, and ζ the vertical component of the relative vorticity. A variable X ↓ is a column vector denoted by

XX₁X_kX_N^T(2)

where (. . .)^T is for transpose. Using the continuity equation and vertical finite differencing, the pressure vertical velocity in σ coordinates can be expressed by

View Expanded

where C is a lower-triangular matrix (Chen et al. 1997).

Equation (3) can be used to compute ω directly from the observed wind, and it can be referred to as a kinematic method. However, the horizontal divergence derived from the kinematic method is sensitive to small errors in the observed wind. The divergence derived from the following vorticity, divergence, and thermodynamic equations is determined from terms of similar magnitude and is less sensitive to observational errors than the kinematic method; it is referred to as an ω-equation method. Because the ω-equation method involves computations of higher-order derivatives than the kinematic method, this diminishes the advantages if the computation of derivatives is not accurate enough. Thus, in order to reduce computational errors, a harmonic-Fourier spectral method (Chen and Kuo 1992a,b) is used.

Based on the harmonic-sine series expansion (Chen and Kuo 1992a; Chen et al. 1996), the streamfunction and velocity potential for a limited area can be divided into the harmonic and inner parts:

ψψ_hψ_iχχ_hχ_i(4)

where the inner parts satisfy the Poisson equations

∇²ψ_i∇²χ_iD,(5)

with zero Dirichlet boundary values. The harmonic parts, ψ_h and χ_h, are the harmonic functions satisfying the Laplace equations

∇²ψ_h∇²χ_h(6)

with the coupled boundary condition

View Expanded

whereu_E | _Σ and v_E | _Σ are the external wind at the boundary.

In order to extend the ω-equation to σ-coordinates, it is important to know how the horizontal pressure gradient force is expressed in σ-coordinates. The transformation of the horizontal pressure gradient from p- into σ-coordinates is expressed by

View Expanded

where ϕ(x, y, p, t) is the geopotential height in p-coordinates, and ϕ(x, y, σ, t) is the geopotential height in σ-coordinates. It is seen from (8) that the vector on the left-hand side is irrotational, but the total vector on the right-hand side of (8) is not irrotational because T is a function of x and y. Partitioning a wind field into the irrotational and nondivergent parts can easily be done by Chen and Kuo’s (1992a,b) method. If the total vector on the right-hand side of (8) is partitioned by this method, its irrotational part is expressed by ∇ϕ^′_i(x, y, σ, t), where ϕ^′_i is referred to as the inner part of an isobaric geopotential height in σ-coordinates. Based on this method, the inner part ϕ^′_i can be derived from the solution of the following Poission equation:

View Expanded

with zero Dirichlet boundary value. The isobaric geopotential in σ-coordinates can also be utilized to improve synoptic analysis and model predictions for high mountain regions. Over the Tibetan Plateau, analyses on isobaric surfaces below the 500-hPa level hardly identify weather systems because data are extrapolated from higher levels. However, these systems do exist and they can affect the weather in the downstream region. It is found (Chen and Bromwich 1996) that the contours of the isobaric geopotential at the σ = 0.995 level over the Tibetan Plateau are very similar to those of the streamfunction for synoptic and mesoscale systems, which is the same as the geostrophic wind approximation used on isobaric surfaces over plains.

b. A generalized ω-equation in σ-coordinates

The vorticity and divergence equations can be transformed into the equations of the inner parts of the streamfunction and velocity potential, respectively, and they are expressed by

View Expanded

View Expanded

where E_i ↓ = [m²(U² + V ²)_i/2] ↓, and the terms ψ_adv,i ↓ and χ_adv,i ↓ are the variation rates of the inner parts of the streamfunction and velocity potential caused by advection, respectively. The equation of the ageostrophic isobaric geopotential can be written as

View Expanded

where

D_hDD_i(15)

and Φ^′_had,ia ↓ is referred to as the variation rate of the ageostrophic isobaric geopotential caused by advection and diabatic heating, and A is a matrix (Chen et al. 1997).

From the solution χ_i of (16), the divergence can be computed from (5), and then ω is calculated from (3). Using the continuity equation, the sigma vertical velocity is computed by

View Expanded

c. A simplified method for computation ofprecipitation

The thermodynamic equation is written as

View Expanded

and the continuity equation for specific humidity is

View Expanded

where C and E are the rates of condensation and evaporation per unit mass and L is the condensation heat per unit mass. The vertical advection of a variable X is approximated by

View Expanded

The advection and adiabatic variations of the temperature and specific humidity are computed by

View Expanded

where the horizontal advection is computed by a spectral method (Chen et al. 1997) and the vertical advection is computed by (20). The vertical velocity σ̇ is obtained from (17) and ω is derived from (16) and (3). Time differencing for (21) and (22) is computed by the Euler-backward or Matsuno scheme.

Large-scale condensation occurs when q_i,j,k(t + Δt) is greater than r_cq^*_i,j,k, where q_i,j,k(t + Δt) is the specific humidity predicted by (22); q^*_i,j,k is the saturation specific humidity at temperature T_i,j,k(t + Δt) and pressure p_i,j,k. Here, T_i,j,k(t + Δt) is computed from (21) and r_c is a critical relative humidity. Excess water in a layer can be obtained by removing supersaturation from the specific humidity field and modifying the temperature through latent heat release. The excess water removed from an atmospheric layer precipitates into the layer immediately below. The falling precipitation either evaporates completely in that layer or brings the layer to saturation and then the excess water passes to the next layer below again. The process is repeated. When the lowermost layer is saturated, the condensed water precipitates onto the ground as snow or rain. The computation procedure is similar to the large-scale condensation and precipitation discussed by Arakawa and Lamb (1977). In this method, both the precipitation and latent heat release are obtained.

At the beginning of the computation for the pressure and sigma vertical velocity, the diabatic heating (L/C_p)(C − E) ↓ vanishes. After the latent heat release is derived, the diabatic heating, (L/C_p)(C − E) ↓, is set to be equal to the computed latent heat release, and then the vertical velocity and precipitation process is calculated again. The whole process is repeated once. Because the effect of latent heat release on the vertical motion is important (Smith et al. 1984; Pauley and Nieman 1992), the above method is used to get a corrected precipitation amount that accounts for this effect.

a. The distribution of the annual precipitation for 1987 and 1988

Data from the EuropeanCentre for Medium-Range Weather Forecasts (ECMWF) at 2.5° × 2.5° resolution (TOGA Archive II from NCAR) for 1987 and 1988 are used. The precipitation is calculated twice per day based on the analyzed data at 0000 and 1200 UTC. The topography of Greenland is presented in Fig. 2b, which shows that the steepest slopes are near the southeast coast. The mesh size of the limited region is 111 × 81 and the grid spacing is 50 km. The time step Δt used here is 30 min. The precipitation rate is only computed for one time step and then it is applied to a 12-h period. According to tests of the precipitation over Greenland, it is better to assume that the critical relative humidity is 0.85.

The retrieved precipitation for 1987 and 1988 is shown in Figs. 3a and 3b, respectively. Figure 3c is the mean annual precipitation for these two years. In order to check if the computed annual precipitation distribution is reasonable, we compare it with the observed mean annual accumulation over the Greenland Ice Sheet shown in Fig. 1b. Although two years are too short for a climatic depiction, the comparison gives a preliminary evaluation of the method.

From Figs. 3a–c, it can be seen that the maximum values of more than 100 cm yr⁻¹ are located along the southeastern coast and southwestern edge of Greenland. A secondary band of relatively high precipitation is present along the western coast of Greenland from Jakobshavn to Thule, which is underestimated in the simulation by Bromwich et al. (1993). In comparison with the observed distribution in Fig. 1b and the results of Bromwich et al. (1993) shown in Fig.1a, it is found that the major shortcoming in the simulation by Bromwich et al. (1993) is corrected by the new method. A large area of very low precipitation about 10 cm yr⁻¹, which dominates the central interior region of Greenland to the north of 70°N, is also shown in Figs. 3a–c.

The annual precipitation over the Greenland region simulated by the NCEP/NCAR reanalysis for 1987 and 1988 is shown in Figs. 3d and 3e, respectively. It can be seen that the annual precipitation pattern over Greenland for an individual year shown in Figs. 3d or 3e is very similar to that of the mean annual precipitation for 1982–94 in Fig. 1c. Comparing Figs. 3d and 3e with Figs. 1b, 3a, and 3b, it can be seen that the two important shortcomings in the mean precipitation over Greenland shown by Fig. 1c are also present in Figs. 3d and 3e. The precipitation amounts from the reanalysis are about twice those observed in Fig. 1b and those retrieved in Figs. 3a–c. There are high precipitation areas in the central region of Greenland with maxima of 111 and 101 cm yr⁻¹ in Figs. 3d and 3e, respectively. They are about 10 times larger than the observed amount and the retrieved values in Figs. 3a–c. A wave train pattern extending southeastward from the southeast coast of Greenland is also found in the annual precipitation in Figs. 3d and 3e for 1987 and 1988.

It is seen fromthe above that the precipitation distributions over Greenland retrieved for 1987 and 1988 by the simple dynamic method presented in this paper are superior to those obtained by NCEP/NCAR reanalysis and are accurate enough to be used for study of the basic features of precipitation over Greenland.

b. The annual mean circulation at the 700-hPa level and vertical motion for 1987 and 1988

The height of the Greenland Ice Sheet in the central interior region is approximately at the 700-hPa level. The annual mean wind, geopotential height, and temperature at the 700-hPa level are shown in Fig. 3f. There is a pronounced minimum in the temperature field over the central interior region. Figure 3g is the annual mean vertical velocity at the 700-hPa level. The vertical velocity is calculated twice per day at 0000 and 1200 UTC, and then these computed values are averaged over two years of 1987 and 1988 to obtain the annual mean vertical motion. Figure 3g shows that a relatively large area with mean descending motion occurs in the central interior region.

The large area of very low precipitation that dominates the central interior region of Greenland to the north of 70°N (Figs. 3a–c) may be closely related to the mean descending motion in the central region shown in Fig. 3g.

Snow and ice surfaces typically have an albedo larger than 80%, implying that more than three-quarters of the incident shortwave radiation from the sun is reflected. Little of this reflected solar radiation is absorbed by the overlying atmosphere. In addition, the ice surface acts nearly like a blackbody for longwave radiation. Thus, the ice surface efficiently radiates thermal energy to outer space. The cold source at the ice and snow surface will reduce the air temperature in the boundary layer by downward sensible heat transfer. On the other hand, an adiabatic warming must be produced by the descending motion in the central region shown in Fig. 3g, and it may primarily be compensated by the temperature decrease due to the diabatic cooling. Furthermore, the descending motion would minimize cloud cover and promote radiative cooling. Thus, the cold center shown in Fig. 3f is probably also related to the intense cooling at the ice surface, which markedly contrasts with the relatively warm air at some horizontal distance away in the free atmosphere.

Scorer (1988) published a paper entitled “Sunny Greenland.” In comparison to other places along the same parallel of latitude, the precipitation over the central interior region of Greenland is very small, thus the weather in the summer half year in this region may be very sunny. However, it is also seen from Figs. 3a–c that most of the southern part of Greenland, south of about 68°N, has heavy precipitation of more than 100 cm yr⁻¹. Thus, “Sunny Greenland” does not apply to the entire island. Scorer (1988) also mentioned that two major parties of explorers were trapped by deep snowfalls for a few months in the 1920s. These falls had not been expected because of the belief that the southern part had the same sunny weather.

c. The annual precipitation over the entire Greenland Ice Sheet

As a check on the magnitudes of the simulated annual precipitation, the annual precipitation over the entire Greenland Ice Sheet was computed. This was done by finding all grid points that lie over Greenland, multiplying their annual precipitation values by the area of the grid square centered on that point, and then summing all these products. Our computed resultsand those obtained by other investigators including those based on glaciological observations are compared in Table 1. Examination of Table 1 shows that the model’s values of annual accumulation are close to those obtained previously.

The interannual variations of precipitation over Greenland were studied by Bromwich et al. (1993) and also estimated from the atmospheric moisture budget by Robasky and Bromwich (1994). These two computationally independent estimates matched closely for the period 1980–89. The 2-yr mean values for 1987 and 1988 from Bromwich et al. (1993) and Robasky and Bromwich (1994) are 651 and 664 km³, respectively, while in our computation, this 2-yr mean value is 638 km³. The values for 1987 from Bromwich et al. (1993) and Robasky and Bromwich (1994) are 19 and 66 km³ larger than those for 1988, respectively, while in our computation, the value for 1987 is also 26 km³ larger than that for 1988. Thus, the mean values for these two years computed by the above three methods are very close, and the tendency of the variations between these two years has the same sign.

d. Precipitation over Greenland assimilated by the NCEP/NCAR reanalysis and T126 model

To determine if the precipitation over Greenland predicted by the global model of NCEP is affected by the horizontal resolution of the model, the precipitation rates derived from the operational global T126 (105 km) model for NCEP (Kanamitsu et al. 1991) are also examined. The precipitation rates with a 1° × 1° resolution predicted by the T126 model are obtained from the NCEP GRIB (Gridded Binary) dataset. In this paper we do not want to study the difference of the precipitation over Greenland predicted by the T62 and T126 models, and we only have the NCEP GRIB data based on the T126 model since 1994. Thus, only some examples of the precipitation retrieved by the simple dynamic method are compared with the computed precipitation rates derived from the operational global T126 model.

The NCEP GRIB dataset only has daily precipitation rates. The precipitation rates at 1200 UTC 24 and 25 July 1994 retrieved by the dynamic method are shown in Figs. 4a and 4b, while the corresponding precipitation rates from the NCEP GRIB data based on the T126 model are shown respectively in Figs. 4c and 4d. Comparing Figs. 4a and 4b with 4c and 4d, respectively, it can be seen that the distribution of the major precipitation centers obtained from these two methods is similar. These dates for comparison were chosen at random. The monthly precipitation for July 1994 retrieved by the dynamic method is shown in Fig. 4e. The monthly precipitation computed from the precipitation rates of the NCEP GRIB dataset is based on a method that is the same as that used in the dynamic method, and its value for July 1994 is shown in Fig. 4f. The monthly precipitation for July 1994 from the NCEP/NCAR reanalysis based on the T62 model is shown in Fig. 4g. Comparing Figs. 4e and 4f, it is seen that the distribution of monthly precipitation amounts obtained from the dynamic method and from the NCEP GRIB data based on the T126 model are similar, but the precipitation over the southeast coast of Greenland retrieved by the dynamic method is a little larger. As shown in Fig. 4g, the two important shortcomings and thefalse wave train pattern shown in Fig. 1c are also present in the monthly precipitation for July 1994 based on the NCEP/NCAR reanalysis. However, these shortcomings do not occur in the predicted total precipitation for July 1994 based on the T126 model shown by Fig. 4f. Therefore, in order to obtain a better assimilation of precipitation over Greenland in the NCEP/NCAR reanalysis, it seems necessary to use a higher resolution model, such as T126, instead of T62.

Recently, Genthon and Braun (1995) used the predicted precipitation of the ECMWF T106 model (about 1.125°) from 6 yr of the ECMWF archive (May 1985 to April 1991) to study the annual mean accumulation (precipitation minus evaporation or sublimation) over Greenland. Comparing Fig. 3c with the annual mean accumulation shown in Fig. 6 of Genthon and Braun (1995), the distributions of these two figures are quite similar. The amounts in the high precipitation region shown in Fig. 3c are about 10 cm yr⁻¹ larger than those shown in Fig. 6 of Genthon and Braun (1995) because the precipitation in Fig. 3c does not take the evaporation/sublimation loss into account.

a. The major system producing precipitation over Greenland

Based on evaluation of the daily isobaric analyses and retrieved precipitation for 1987 and 1988, it is found that the major weather system producing precipitation over Greenland is the frontal cyclone. Cyclonic activity, in general, includes the distribution of cyclone frequency, the locations of cyclogenesis and the tracks followed by cyclones. Cyclonic activity is very important for the climatological features of precipitation over Greenland.

b. Three major regions of Greenland

According to the synoptic survey of precipitation and possible effects of orography on moving cyclones, the Greenland Ice Sheet may be divided into three major regions separated by the solid lines shown in Fig. 5. The first is referred to as the sourthern region and mostly lies to the south of about 68°N. When a cyclone passes close to the southwest coast of Greenland or Cape Farewell on the southern coast, heavy precipitation often occurs over the southern region. This situation sometimes causes lee cyclogenesis near the east coast.

The second region is located to the north of the southern region, and it is referred to as the central region. Its width is about 1200 km, and the highest elevation is more than 3000 m. Based on the survey of 1987 and 1988, it is found that the central region of Greenland has an important blocking effect on moving cyclones. No cyclone moving from west to east across this region was found during these two years. According to the precipitation characteristics, the central region can also be divided into three subregions, which are the central west coastal region, central interior region, and central east coastal region, as shown in Fig. 5. It has been pointed out in the previous section that the central interior region has annual average descending motion and very low precipitation.

The third region is the north coastal region located to the north of 80°N and is a narrow area that slopes steeply down to sea level. A few cyclones influence this region during summer.

c. Seasonal variations of precipitation over the north coastal and central interiorregions, and composite charts for high precipitation months in summer

In order to show some intraannual variations of precipitation over Greenland, the monthly mean precipitation for 1987 and 1988 averaged for the five regions and for all of Greenland are given in Figs. 6a–f. This mean precipitation is computed by finding all N grid points that lie in the region, summing their monthly precipitation values, and then dividing by N.

Figures 6a and 6b show the mean monthly precipitation averaged over the north coastal and central interior regions. It is seen from these two figures that mean monthly precipitation has an obvious seasonal variation, especially in the north coastal region. The precipitation in the north coastal region is much larger in summer (June, July, and August) than in winter (December, January, and February). In the central interior region, the amplitude of the variation is very small, but the values in summer are also larger.

Figures 7a and 7b are the composite monthly mean sea level pressure (SLP) and monthly mean wind, temperature, and geopotential height at the 700-hPa level, respectively, for the high precipitation months in the north coastal and central interior regions. Because the sea level pressure over Greenland can only be extrapolated and computed from the temperature at upper levels, it is too high over the Greenland Ice Sheet.

The monthly total precipitation is produced by the sum of the behavior of individual synoptic systems within the month and cannot be completely deduced from the monthly mean synoptic maps. However, a relationship between the monthly total precipitation and monthly mean circulation approximately exists, similar to what has been found for the coastal regions of Antarctica (Bromwich et al. 1995). It is seen from Figs. 7a and 7b that there is a mean cyclone located in the Arctic Ocean close to the northeast coast of Greenland and Spitzbergen during the high precipitation months in summer. In this case, the north coast is affected by northwesterly winds associated with the mean cyclone. This flow is forced to ascend over the slopes of the north coastal region and causes precipitation.

In order to show possible cyclone track information for high precipitation months over the north coastal region, a Lanczos filtering scheme (Duchon 1979) for 2.5–6-day variations of SLP was applied to the twice daily ECMWF data for the summer months (June–August) of 1987–88. The average synoptic SLP variability for summer is shown in Fig. 7c. The axis of maximum variability corresponds to the cyclone track. It is seen from Fig. 7c that a relatively large center is located over Baffin Bay and trends northeastward through Nares Strait. These features reveal that the cyclones often develop over Baffin Bay in summer and then move northeastward through Nares Strait. There are other centers of variability in Fig. 7c located to the southeast of Iceland and over the Arctic Ocean and Scandinavia. These maximum value regions are related to the tracks of the Icelandic and other cyclones, whose intensity is stronger than that of the cyclones over Baffin Bay. It should be noted that axes of variance maxima do not always coincide with traditional storm tracks because alterations in the position and intensiy of anticyclones can also contribute significantly to variance maxima.

d. Composite charts for high and low precipitation months over the southern region: Labrador Sea and Icelandic cyclones

Except for the monthly precipitation over the north coastal and central interior regions discussed above, it is very difficult from Figs. 6c–f to find obvious seasonal precipitation variations. However, if composite charts for high and low monthly precipitation over a region are constructed without making any allowance for seasonal variation, some interesting results are found.

Figure 6c is the mean monthly precipitation averaged over the southern region of Greenland. The majority of the precipitation over all of Greenland occurs in the southern region; thus, the monthly mean precipitation over this region can be used to represent the monthly mean precipitation over the entire ice sheet. Figure 6f is the mean monthly precipitation averaged over all of Greenland. Comparing Figs. 6c and 6f, it can be seen that the values are smaller in Fig. 6f due to the averaging over a larger area, but the months with maximum and minimum values are nearly the same.

A month for which the monthly mean precipitation averaged over the southern region of Greenland is larger than 80 cm yr⁻¹ or smaller than 50 cm yr⁻¹ is referred to as a high or low precipitation month, respectively. Figures 7d and 7e are the composite monthly mean SLP and wind, geopotential height, and temperature fields at the 700-hPa level, respectively, for the high precipitation months (4 months in the sample) over the southern region. It is seen from Fig. 7d that there is a mean cyclone located in the Atlantic Ocean and Labrador Sea close to the southwest coast of Greenland. At the 700-hPa level shown in Fig. 7e, there is a trough in the Labrador Sea close to the east coast of Labrador. The southern region of Greenland has warm air advection and is located in advance of this trough, where the cyclonic vorticity advection is often larger at upper levels than lower in the troposphere. Such conditions are favorable for dynamically supported ascending motion on individual synoptic maps (Chen 1981, 1987; Sanders and Hoskins 1990) and may approximately be applied to monthly mean conditions. In addition, both the southwest and southeast coasts of this region are directly impacted by the onshore flow from the Atlantic Ocean. This flow would have relatively high humidity and produces precipitation over the coastal slopes by orographic lifting.

Figures 7f and g are the composite monthly mean SLP and 700-hPa level charts, respectively, for the low precipitation months (5 months in the sample) over the southern region. Figure 7f shows that a monthly mean cyclone is located in the Atlantic Ocean close to Iceland. Comparing Figs. 7e and 7g, the northern part of the trough at the 700-hPa level shown in Fig. 7e has moved eastward and is located over Davis Strait, but its southern part is not clear and seems to have moved to the east of Greenland. With this circulation, the southern part of Greenland is located to the rear of the southern part of the 700-hPa trough and to the northwest of the surface cyclone, and the temperature advection is not clear and very small. Both the baroclinic and orographic effects favorable for precipitation in Figs. 7d and 7e are not present in Figs. 7f and 7g.

The average synoptic SLP variability for high and low precipitation months over the southern region are shown in Figs. 7h and 7i, respectively. The variability for high precipitation months over the Labrador Sea is larger than 6 hPa, and the axis of maximumvalues (cyclone track) has two routes. One goes northward through Davis Strait along the west coast of Greenland. The other turns eastward and is associated with lee cyclogenesis. This lee cyclogenesis is similar to the cyclones that form in the lee of the Rocky Mountains, known as “Colorado” and “Alberta” cyclones (Palmen and Newton 1969). Lee cyclogenesis is usually associated with a “parent” cyclone upstream along the west coast of the southern region. When the parent cyclone moves eastward across the southern region, a discontinuous movement may occur in the sea level pressure field. The parent cyclone often decays along the west coast or sometimes moves northward through Davis Strait, while a lee cyclone develops along the east coast. This lee cyclogenesis process is revealed by Fig. 7h, and an example will be shown in the appendix.

Figure 7i shows that the axis of maximum variability is over the Atlantic Ocean to the southeast of Greenland. This is the cyclone track for low precipitation over the southern part of Greenland, and it passes through the region near the southeast coast of Iceland from southwest to northeast.

The mean cyclone located to the southwest of Greenland over the Labrador Sea as shown in Fig. 7d is referred to as a Labrador Sea cyclone. For the same reason, the mean cyclone located to the southeast of Greenland over the Atlantic Ocean near the south coast of Iceland as shown in Fig. 7f is referred to as an Icelandic cyclone. The Icelandic cyclone in Fig. 7f is similar to the Icelandic low present on the long-term averaged mean SLP charts.

e. Composite charts for high precipitation months over the central west and central east coastal regions

Figures 6d and 6e are the mean monthly precipitations averaged over the central west coastal and central east coastal regions of Greenland, respectively. A month for which the monthly mean precipitation averaged over one of these coastal regions is larger than 40 cm yr⁻¹ is referred to as a high precipitation month for that coastal region. Figure 7j is the composite monthly mean SLP for the high precipitation months (6 months in the sample) over the central west coastal region. Figure 7j shows that there is a mean trough along Davis Strait and Baffin Bay and a small mean cyclone center located in Baffin Bay. This cyclone produces the precipitation over the central west coastal region.

Figure 7k is the composite monthly mean SLP for the high precipitation months (3 months in the sample) over the central east coastal region. It shows that there is a mean cyclone located near the southeast coast, but a mean trough extends northward along Denmark Strait. The easterly flow associated with this trough onto the east coast causes precipitation over this region.

f. A relationship between precipitation over Greenland and cyclonic activity

An SLP monthly mean cyclone located in a region means that, during this month, cyclones often pass through or they persist or develop with high frequency. Because most of the precipitation over Greenland falls in the southern region, from Figs. 7d and 7f, it can be seen if the Labrador Sea cyclone occurs, enhanced precipitation will fall over Greenland during that month. In contrast, if the Icelandic cyclone occurs, precipitation over Greenland will be reduced. This is an important relationship between precipitation over Greenland and cyclonicactivity.