Sites

The Swedish station Wasa and the Finnish station Aboa in Dronning Maud Land (DML), Antarctica, are situated on the nunatak Basen (73˚03’ S, 13˚25’ W). Basen lies in the northeastern part of Vestfjella, which is a 135 km long range of nunataks. the mountain range is almost parallel to the coast and located about 120 km from the coast. Basen reaches 584ma.s.l.

The EC measurements inJanuary 2000 were made at two different sites, LillaVardet and Aerosol Hill, over periods of 4 and 19 days, respectively. LillaVardet is situated 1km east of the research stations, and the measurements were performed over a smooth snow surface that had a considerable fetch over snow in the main wind direction (i.e. east-northeast). the footprint of the EC system consisted entirely of a snow-covered area. Aerosol Hill is located a few hundred metres east-northeast of the research stations, and the ground on which the EC system was standing consists mainly of rocks in the 5–20 cm diameter range. the footprint of the EC system at Aerosol Hill was dominated by a rocky surface interrupted by snow patches.

The EC flux system

A Gill Solent 1012R research model ultrasonic anemometer was mounted at 5.25 m above the ground on a mast. A TSI 3010 Condensation Particle Counter (CPC) was located in an isolated aluminium box at the foot of the mast, with its sampling intake located just below the sensing heads of the ultrasonic anemometer. the box also contained a heating system together with a computer and an interface to the ultrasonic anemometer. the sampling flow rate was 1 L min^–1 through a 7.8 m heated stainless-steel sampling tube with an inner diameter of 4 mm.

The anemometer measured the vertical wind component (w) together with temperature (T) and the horizontal wind components (u and v) at 20.8 Hz. the CPC measured the number concentration (N) of aerosol particles with a dry diameter (D_p) larger than 10 nm. the maximum concentration that this device could measure was 10 000 particles cm^–3.

EC flux calculations

The fast measurements by the system allow us to separate N and w into mean values hNi and hwi (resulting from the mean flow defined by a 120 s running average) and turbulent fluctuations and

The aerosol flux as well as momentum and turbulent sensible-heat fluxes were calculated following Reference Buzorius, Rannik, Mäkelä, Vesala and KulmalaBuzorius and others (1998) and Reference Nilsson and RannikNilsson and Rannik (2001). A three-dimensional co-ordination of wind vector to the local streamlines was performed according to Reference Kaimal and Finnigan.Kaimal and Finnigan (1994). the turbulent vertical aerosol number flux (F) is given by

where the covariance was calculated over 30 min periods from the turbulent fluctuations and By convention, a positive value for F implies an upward flux. the aerosol deposition velocity (v_d) was calculated from the measured flux normalized by the mean aerosol number concentration at a certain reference height:

A positive deposition velocity is defined as being downwards. the response time of the CPC was of the order of 1Hz. the peak contribution to the aerosol flux will be at frequencies of 0.1 Hz or lower. the underestimation of particle-number flux due to the limited frequency response of the CPC was corrected for according to Reference MooreMoore (1986). the correction factor was typically <10%, since eddies at frequencies >1Hz are responsible for a relatively small part of the flux.

Themomentum flux and friction velocity (u_*) were calculated from u, v and w, while the turbulent sensible-heat flux (H) was calculated from w and T. Th. Monin–Obukhov length (L) was subsequently calculated from u_* and H. Th. surface roughness height (z₀) was solved numerically from:

where κ is von Kármán’s constant (κ =0.4), (U) is the averaged wind speed, z is the height of measurement, ϕ = (1 – 15z/L)^1/4 and ϕ₀ = (1 – 15z₀/L)^1/4.

Differential-mobility particle sizer

A differential-mobility particle sizer (DMPS) system at Aerosol Hill was continuously monitoring the aerosol particle-size distribution from 3 to 850 nm in diameter during the period 5–20 January 2000. the measured aerosol particle-number concentrations (for D_p >10 nm) from the DMPS system agrees well with the CPC data from the EC measurements. We have utilized the DMPS data to separate the EC measurement data from Aerosol Hill into groups dominated by one of three particle-size modes − ultrafine, Aitken, or Aitken plus accumulation mode−and have thereby been able to identify dry-deposition velocities for the different particle-size modes. the results from the DMPS measurements will be reported extensively elsewhere (paper in preparation by I.K.Koponen).

Dry-deposition velocity

The dry-deposition velocity can also be expressed in terms of resistances (e.g. Reference Seinfeld and Pandis.Seinfeld and Pandis, 1998):

where r_a is the aerodynamic resistance, r_b is the resistance in the quasi-laminar sub-layer and v_s is the settling velocity due to gravitation (e.g. Reference HindsHinds, 1982).

The aerodynamic resistance is a function of u_*, z₀, z and stability (z/L). It can be expressed as (e.g. Reference Seinfeld and Pandis.Seinfeld and Pandis, 1998):

The sub-layer resistance r_b is a function of D_p, z₀, u*, the diffusion coefficient (D) and kinematic viscosity (ν). This resistance increases with increasing surface roughness and decreasing diffusivity of the aerosol. for fine particles, the influence of Brownian diffusion increases, and r_b decreases with decreasing particle size. Reference Schack, Pratsinis and Friedlander.Schack and others (1985) give an expression for r_b for neutral conditions and a completely rough surface:

A and B are empirical constants of proportionality and depend on the shape of the elements composing the collecting surface. the right side of the equation is for the diffusion and interception ranges, respectively. the dry-deposition velocity is indirectly dependent on (U), z/L and z₀ through the u_* dependency of Equation (4).

Snow-covered area at Lilla Vardet

The EC measurements at Lilla Vardet were performed during 4 days at the beginning ofJanuary, andthemeasured and calculated variables are shown in Table 1. Both downward and upward particle fluxes were registered, but the downward flux was one order of magnitude larger on average. Ninety per cent of the aerosol particle-number concentrations were found in the range 307–444 particles cm^–3. the stratification of the surface boundary layer was near neutral most of the time, as indicated by a median for z/L of –0.005. Ninety per cent of the dry-deposition velocities ranged from 0.08 to 1.89 cm s^–1, with a median of 0.33 cm s^–1, for all particle sizes (D_p >10 nm).

Table 1.

Measured and calculated variables at Lilla Vardet, given as arithmetic mean plus or minus 1σ, the median, the range where 90% of the data are found and the number of data points (0.5 hour averages)

Snow-free area at Aerosol Hill

The EC measurements at Aerosol Hill were performed over a 19 day period later inJanuary, andthemeasured and calculated variables are shown in Table 2. At Aerosol Hill the downward and upward particle fluxes were of the same order of magnitude, with a small net flux downwards. A frequency analysis showed that two-thirds of the measured particle fluxes were negative, which also indicates that deposition dominated. the positive fluxes, i.e. upward fluxes, may be explained as resuspension of particles from the ground. However, it should be noted that upward fluxes were also measured over the snow surface at LillaVardet. Particle-number concentrations at Aerosol Hill were found in the range 270–750 particles cm^–3.The low albedo of the nunatak on which the EC system stood favours unstable conditions in the surface boundary layer. This is indicated by the negative values for z/L in 90% of cases. An explanation for this is the relatively large net radiation, which at noon peaked at approximately 90Wm^–2. the average wind speed for the different measuring periods was higher at Aerosol Hill than at Lilla Vardet, and reached >10ms^–1 during 8% of the measuring period. Ninety per cent of the dry-deposition velocities ranged from 0.2 to 2.4 cm s^–1, with a median of 0.80 cm s^–1, for all particle sizes (D_p >10 nm).

Table 2.

Measured and calculated variables at Aerosol Hill, given as arithmetic mean plus or minus 1σ, the median, the range where 90% of the data are found and the number of data points (0.5 hour averages)

Diurnal cycles

Figure 1 shows 0.5 hour averages in meteorological variables from the period of measurement at Aerosol Hill. the turbulent sensible-heat flux (Fig. 1a) had a diurnal cycle that peaked at noon when the solar zenith angle was at minimum, with an average H maximum of 260 Wm^–2.The temperature cycle (Fig. 1b) lagged the H cycle by about 2 hours, and the average diurnal temperature variation ranged from –6.6˚ to –3.7˚C. During the morning hours, the surface boundary layer was close to neutrally stratified, but an unstable boundary layer was formed during the day due to the radiative warming at the surface (Fig. 1c). When the wind speed increased during the day (Fig. 1d) and the stability diminished, the friction velocity increased (Fig. 1e). the particle concentration increased rapidly during the morning hours and culminated at noon (Fig.1f). A similar but more scattered trend was seen in the downward aerosol particle flux (Fig.1g).

Fig. 1

Averaged diurnal cycles at Aerosol Hill of (a) turbulent sensible-heat flux (H), (b) temperature (T), (c) surface boundary-layer stability (z/L), (d) wind speed (U), (e) friction velocity (u_*), (f) particle-number concentration (N), (g) particle flux (F) and (h) dry-deposition velocity (v_d). Full lines connect 0.5 hour averages, and dotted lines are polynomials fitted to the diurnal averages.

Hence, we can see a clear diurnal cycle in the surface boundary-layer stability, temperature and fluxes. This is not due to the contrast between night and day, since we had sunlight 24 hours d^–1. Instead, the driving force must be the diurnal cycle in the solar zenith angle, which is still rather large at 73˚ S. This agrees with the measurements by Reference Nilsson and RannikNilsson and Rannik (2001) over the pack ice of the Arctic Ocean, although the diurnal amplitudes were smaller in their case due to the higher latitude (>80˚N). the dry-deposition velocity, or normalized flux, had no significant trend (Fig. 1h). Since most of the large downward fluxes occurred simultaneously with high values of aerosol particle concentrations, the value of the normalized flux remains practically the same when data points are compared.

Because the measurements over the snowfield at Lilla Vardet lasted only a few days, the average diurnal cycles are much too influenced by the weather conditions on individual days to produce a figure corresponding to Figure 1.However, the data from Lilla Vardet indicate a diurnal behaviour similar to that for Aerosol Hill, with some differences. At Lilla Vardet, H was negative at ``night’’ and positive at noon, with a smaller amplitude of the diurnal cycle. This is in agreement with a more stably stratified surface boundary layer and obviously caused by the higher albedo of the snowfield and the smaller absorption of solar radiation as compared to Aerosol Hill. the downward aerosol fluxes and the dry-deposition velocity over the snowfield were smaller than at Aerosol Hill, except in the afternoons.

Dry-deposition velocities

Previously reported dry-deposition velocities show a large variability depending on the atmospheric conditions and characteristics of the deposited species. Examples of dry-deposition velocities for different species to snow surfaces on Greenland are shown in Table 3. None of these studies were made with the EC method. All these previous studies emphasized the dry-deposition velocity for the size range that is most important for mass deposition, rather than number deposition, hence they refer to a larger size range. the dry-deposition velocities in this study are generally higher than the values in Table 3, especially the ones measured over the snow-free surface at Aerosol Hill. However, an order-of-magnitude difference is found in the reported values for K, Mg and Na from Dye 3, hence the absolute value of the dry-deposition velocity can vary strongly between different situations.

Table 3.

Examples of dry-deposition velocities (+1σ) to snow surfaces

The dry-deposition velocity is dependent on the turbulence intensity, atmospheric stability and particle size (see Equations (5–7)). Unlike Reference Nilsson and RannikNilsson and Rannik (2001), we were unable to identify an increased dry-deposition velocity with increased friction velocity in this dataset. the reason is probably a large scatter in v_d due to changes in the aerosol size distribution, but also the variability in z₀ and z/L.

The large surface elements on Aerosol Hill gave a median surface roughness height of 0.11 cm, while the median surface roughness height at Lilla Vardet was 0.03 cm. the median dry-deposition velocity was >100% higher over the rougher surface than over the smooth surface at Lilla Vardet. A study by Reference Schack, Pratsinis and Friedlander.Schack and others (1985) consists of a number of wind-tunnel experiments over various surfaces, weighting the diffusion and interception effects with the empirical constants A and B(see Equation (7)). They report a surface that consisted of gravel in the 3.8–5.1cm size range, which resulted in surface roughness heights of 0.3–0.6 cm and deposition velocities ranging from 0.05 to 33 cm s^–1. the constants A and B for gravel were according to them 4.0±0.4 and 121±8, respectively.To calculate the quasi-laminar sub-layer resistance r_b at Aerosol Hill, the empirical parameters had to be approximately 10 times larger (A = 75, B =1000) to parameterize themedian dry-deposition velocity.

Analysis of the DMPS data showed that there were three distinct particle-size modes in the sub-micron aerosol (Fig. 2). These number modes had a geometric mean diameter of around 14, 42 and 134nm, referred to as ultrafine, Aitken and accumulation-mode, respectively. This agrees approximately with the range found by Reference Covert, Wiedensohler, Aalto, Heintzenberg, McMurry and LeckCovert and others (1996) in the Arctic marine boundary layer. the data from the EC measurements at Aerosol Hill were divided into three different categories based on the aerosol number size distributions:

Fig. 2

Three examples of particle-number size distributions, dN/d log(D_p), showing the three categories into which the EC dataset was divided: ultrafine mode (solid line), Aitken mode (solid line with stars) and Aitken plus accumulation mode (solid line with circles).

(a) ultrafine (D_p =10–20 nm) whenever there was an ultrafine mode present

(b)Aitken (D_p =20–90nm) when there was an Aitken mode but no ultrafine or accumulation mode present; and

(c) Aitken plus accumulation (20 nm< Dp <200 nm) when both modes were present at the same time.

There were no cases with only an accumulation mode present. Reference Nilsson and RannikNilsson and Rannik (2001) used a similar approach for the v_d over the Arctic Ocean pack ice and also could not separate a specific dry-deposition velocity for only accumulation-mode particles. Since Aitken mode was present most of the time, the third category could not consist of a pure accumulation mode but had to be a combination of both. This was also the case for the ultrafine mode, where an Aitken and an accumulation mode were also present. Since the measured F includes a wide range of particle sizes, it is difficult to relate v_d to D_p. the number of Aitken-mode particles may have dominated the category with the accumulation mode, while the number of ultrafine-mode particles probably dominated the category of ultrafine mode.

Median dry-deposition velocities to the snow-free surface at Aerosol Hill were 1.02, 0.75 and 0.40 cms^–1 for the ultrafine, Aitken and Aitken plus accumulation-mode category, respectively. These are larger than the values given by Reference Nilsson and RannikNilsson and Rannik (2001) over sea ice (Table 4). the measured median dry-deposition velocity of each size category agrees with that calculated from Equations (5–7) with A and B set to 75 and 1000, respectively, through empirical evaluation. Since we have tuned A and B in Equation (7) to fit the observations, the agreement between the magnitude of the measured and calculated v_d is not a solid result. However, the dv_d/dD_p slope is much less influenced by the choice of Aand Bthan is v_d. the dv_d/dD_p slope between the dry-deposition velocities of the ultrafine and Aitken modes as measured (–0.7×10⁵ cm s^–1m^–1) and given by Equations (5–7) (–1.1×10⁵ cm s^–1m^–1) agrees well. This is a confirmation from field data that indeed the small ultrafine-mode particles have a higher deposition velocity than the larger Aitken-mode particles, as predicted by theory, owing to a stronger Brownian diffusion.

Table 4.

Dry-deposition velocity (cm s^–1) at Aerosol Hill for the three size categories ultrafine, Aitken, and Aitken plus accumulation, given as arithmetic mean plus or minus 1s, the median, the range where 90% of the data are found and the number of data points (0.5 hour averages)