Introduction
The Arctic is subject to an amplification of the global warming, as the
observed temperature increase has been almost twice the global average
(IPCC, 2013; Serreze and Barry, 2011; Shindell and
Faluvegi, 2009). This resulted in the first complete opening of the
Northwest Passage in 2007 (Serreze et al., 2007) together
with a greening of the coastal tundra (Bhatt et al., 2010) and altered wind
patterns (Overland and Wang, 2010). The Arctic amplification is the
result of complex global feedbacks (acting at different spatial and temporal
scales): the impact of sea ice changes on the heat fluxes between the ocean
and the atmosphere (Screen and Simmonds, 2010a, b), the effect of
changes in the cloud cover and water vapor on the longwave radiation fluxes
(Francis and Hunter, 2006), the changes in atmospheric and oceanic heat
transports (Yang et al., 2010), the black carbon (BC) deposition on the snow
(Hansen and Nazarenko, 2004) and the changes in the atmospheric BC and
aerosol concentrations themselves (Flanner, 2013;
Serreze and Barry, 2011; Shindell and Faluvegi, 2009).
Many of these processes depend on aerosol absorption and scattering of the
solar radiation (direct effect). Additionally, indirect effects play an
important role as the aerosols seed and modify the cloud properties. Lastly,
light absorption by BC can alter the atmospheric thermal structure within,
below or above clouds, consequently affecting cloud distributions (IPCC,
2013; Bond et al., 2013; Ramanathan and Feng, 2009; Koren et al., 2004, 2008; Kaufman et al., 2002). Shindell and
Faluvegi (2009) estimated that globally the decreasing concentrations of sulfate
aerosols and the increasing concentrations of BC contributed (during
1976–2007) 1.09 ± 0.81 ∘C to the Arctic surface
temperature increase of 1.48 ± 0.28 ∘C.
Aerosol particles are short-lived pollutants (∼ 1 to several weeks
of residence time) and act as short-lived climate forcers; thus, their
effect could be employed in short-term climate strategies
(Ødemark et al., 2012; Shindell et al., 2012;
Jacobson, 2010; Quinn et al., 2008). To
adopt the right mitigation strategies, key scientific issues in the study of
Arctic aerosols have to be solved. They include the identification of the
relative importance of long-range advection with respect to local emissions
(Flanner, 2013; Sand et al., 2013;
Shindell and Faluvegi, 2009). Most importantly, the seasonal characterization of the aerosol vertical
structure, a very poorly determined piece of information, is required.
Indeed, the aerosol properties (size distribution, chemical composition,
optical properties) in the Arctic exhibit a pronounced seasonal variation
due to an interplay of dominating sources (outside or inside the Arctic
region) with meteorological conditions that allow or inhibit the transport
from source regions (Quinn et al., 2008;
Eckhardt et al., 2003). The spring period is
characterized by the presence of the Arctic haze dominated by the
accumulation-mode aerosol (enriched in BC). During the Arctic haze, an
inflow of pollution (aerosol and gases) from northern midlatitudes (during
winter–spring) results in a reduction in visibility (Jacob et al., 2010;
Stohl et al., 2006; Radke et al., 1984; Barrie and Hoff, 1985; Brock et al.,
1989; Shaw, 1995). The Arctic haze occurs under meteorological conditions
with stable stratifications and the frequent and persistent occurrences of
surface-based inversions. According to Stohl et al. (2006), within these
conditions, the air pollution can be transported into the Arctic at
low level (followed by ascent in the Arctic or low level alone) or with an
uplift outside the Arctic, followed by descent in the Arctic itself. The
summer period is dominated by fresh Aitken particles, locally formed, with a
negligible BC content
(Tunved
et al., 2013; Spackman et al., 2010; Eleftheriadis et al., 2009; Ström et
al., 2003, 2009; Udisti et al., 2013; Viola et al., 2013).
In addition to the aforementioned seasonality, the same type of aerosol can
produce different climatic effects (warming or cooling) and local feedbacks
(snow/ice albedo, clouds) depending on its vertical location
(Flanner, 2013; Sand et al., 2013; Ban-Weiss et al., 2011;
Shindell and Faluvegi, 2009). For example, it is well known that BC aerosol
absorbs solar radiation and heat the surrounding air (Ferrero et al., 2011a,
2014; Samset et al., 2013, 2014; Ramana et al., 2007). The surface
temperature response varies considerably with the altitude of the induced
heating. BC may potentially warm the Arctic if it is located immediately
above snow and ice while it has a cooling effect, if it is located in the
free troposphere. In the latter case, BC may reduce the surface air
temperature and promote the increase in the sea-ice fraction (Flanner, 2013;
Brock et al., 2011; Seinfeld and Pandis, 2006; Hansen and Nazarenko, 2004).
The latter phenomenon results from a combination of the weakening of the
northward heat transport (due to a reduction in the meridional temperature
gradient) and the increasing of atmospheric stability (caused by the
contemporary dimming of the surface and heating aloft) which turns into a
reduction of the downward sensible heat flux (Flanner, 2013;
Sand et al., 2013; Shindell and Faluvegi, 2009).
In addition to the vertical distribution of BC, that of the total aerosol
particles is important: it can influence the indirect effect and the related
feedbacks. Changes in the cloud cover (especially low-level Arctic stratus)
increase the downward longwave flux to the surface as a function of the cloud
base temperature and cloud phase (liquid, mixed or ice)
(Serreze and Barry, 2011; Francis and Hunter, 2006). Low
clouds mainly warm the surface in the Arctic (with the exception of a brief
period in summer) (Vavrus et al., 2009; Intrieri et al., 2002) due to the
stable stratified conditions that often prevail in the Arctic (Manabe and
Wetherald, 1975). Because the highest number density of aerosol particles
observed in the Arctic is due to a locally formed aerosol (mainly in summer
as stated above) (Tunved et al., 2013; Engvall et al., 2008a;
Ström et al., 2003) it is important to assess the vertical behavior of
the aerosol concentration as a function of its size and the season. It is therefore
necessary to measure the vertical profiles in the Arctic. For this purpose, several
field campaigns have been performed in the Arctic in recent years with the aim to characterize aerosol properties along the
vertical direction.
The ARCTAS mission (Jacob et al., 2010, and references therein) showed highly layered air pollution
transport from North America and east Asia in the spring, characterized by
anthropogenic aerosol below 2 km and by biomass burning in the 2–4 km
layer. The ARCPAC campaign (Brock et al., 2011) grouped the aerosol affecting
the Arctic in spring in four categories: background troposphere (relatively
diffuse, sulfate-rich aerosol); depleted aerosol within the surface
inversion layer over sea ice; layers of organic-rich biomass burning aerosol
(above the top of the inversion layer) (see also Warneke et al., 2010) and
layers dominated by fossil fuel combustion. The ASTAR campaign (Engvall et
al., 2008), focused on the spring to summer transition period in Svalbard,
found Aitken and accumulation-mode particles more concentrated in the free
troposphere compared to the boundary layer. Kupiszewski et al. (2013)
reported new particle formation events in the near-surface layer (possibly
related to biological processes) during the summer ASCOS campaign.
Considering the BC, the springtime PAM-ARCMIP (Stone et al., 2010) and
HIPPO (Schwarz et al., 2010) campaigns showed high BC concentrations close
to the ground, below the thermal inversion, but also dense pollution and BC
at high altitudes over the Arctic (Wofsy et al., 2011). Interestingly, the
PAM-ARCMIP results show a decrease of BC compared to past measurements (i.e.,
AGASP; Hansen and Novakov, 1989). In addition, the HIPPO campaign revealed
that in the lower troposphere the BC vertical gradient can change seasonally
from positive to negative (Schwarz et al., 2013). In this respect, Spackman
et al. (2010) and Koch et al. (2009) reported BC located mainly in the
Arctic free troposphere with a positive gradient in the lower troposphere.
The aforementioned campaigns were conducted mainly using aircraft (or
helicopters) that are limited to intensive
observational periods (Kupiszewski et al.,
2013; Bates et al., 2013;
Spackman et al.,
2010; Schwarz et al.,
2010; Koch et al., 2009). Thus, aerosol vertical profiles in
the Arctic appear scarce if compared with the number of available data
collected at ground level
(Samset et al., 2013;
Koch et al., 2009). There is the need for regular vertical
aerosol profiling campaigns to improve the description of a seasonally
resolved aerosol and BC vertical behavior.
In addition to this, aerosol vertical distribution could be affected in the
future by changes in the aerosol emissions within the Arctic itself. The
increasing of shipping emission in the Arctic is a good example. Shipping
emissions inject the BC directly into the Arctic planetary boundary layer
(probably warming the surface and depositing on snow and ice). The
importance of the increasing shipping emission in the Arctic has been
recently underlined (Eckhardt et al., 2013;
Corbett et al., 2010; Granier et
al., 2006). Although the final impact is debated
(Browse et al., 2013), the effective vertical
distribution of these emissions has not yet been investigated. Thus, there is a
clear need to also improve the knowledge about aerosol
vertical profiles in the Arctic during week-long campaigns along years to
find common rules of behavior.
The Arctic site of Ny-Ålesund (Svalbard Islands) is particularly
suitable for such measurements, featuring long-term data series of ground-based aerosol properties, lidar profiles, radiometric and meteorological
data
(Maturilli and Kayser, 2016; Tunved et al., 2013; Di Liberto et al., 2012;
Vihma et al., 2011; Hoffmann et al., 2009; Eleftheriadis et al., 2009; Stock et al., 2012; Ström et al., 2003, 2009). Long-term
upper-air observations by daily radiosondes provided an overview of the
atmospheric vertical structure above Ny-Ålesund, including the planetary
boundary layer (PBL) altitude range (Maturilli and Kayser, 2016; Vihma et
al., 2011). In this climatological approach, stable atmospheric conditions,
radiative surface-based inversions were frequently found during polar night
conditions, indicating stable atmospheric conditions with suppressed
vertical exchange. Once the snowmelt leads to considerable sensible and
latent heat fluxes at the surface, atmospheric stratification becomes
neutral or instable, allowing convection and vertical mixing.
These observations point towards the need to understand how the aerosol is
vertically layered as a function of the meteorological changes along seasons.
Despite this, as stated above, aerosol and BC measurements along vertical
profiles are reported to be sparse; even recent UAV applications could
improve the available data sets
(Bates et al., 2013).
Thus, this paper reports new data of aerosol and BC vertical profiles
measured over Ny-Ålesund (Svalbard islands) in 2 successive years
(2011–2012) during an extensive field campaign (200 vertical profiles).
Vertical profile measurements were conducted in the framework of the
PRIN2009 ARCTICA project. The main part of the scientific activities at
Ny-Ålesund was aimed at studying the chemical and physical properties of
the aerosols and the long-range transport processes relevant for the
measurements of organic and inorganic species at the site and along vertical
profiles (Moroni et al., 2015; Udisti et al., 2013).
We describe first the sampling sites and the vertical profile measurements
(Sect. 2). Results and discussion follow in Sect. 3, with the
conclusions in the final Sect. 4.
Methodology
Tethered balloon soundings were carried out during spring 2011 and two
summers (2011 and 2012) over Ny-Ålesund. The site is located at the
Kongsfjorden, a fjord that develops in the northwest–southeast (NW–SE) direction.
Northwards, Ny-Ålesund faces the sea, while a small chain of 400–500 m
high mountains is located to the south (Fig. 1a).
(a) Ny-Ålesund, the Kongsfjord and the surrounding orography;
(b) Gruvebadet sampling site; (c) the tethered balloon in Ny-Ålesund.
Dates, UTC, number of profiles, maximum altitude and sky
conditions reached during the 2011–2012 spring–summer campaign in
Ny-Ålesund.
Date (dd-mm-yyyy)
Time in UTC
No. profiles
Max altitude (m)
Cloud base (m)
Spring 2011
30-03-2011
12:40–15:18
6
741
No clouds
01-04-2011
06:30–17:25
10
788
No clouds
04-04-2011
18:17–20:18
6
748
1152
06-04-2011
16:42–19:22
8
716
No clouds
07-04-2011
12:51–19:23
10
712
No clouds
08-04-2011
08:25–19:44
14
740
No clouds/15341
10-04-2011
12:45–14:38
6
300
No clouds/9842
14-04-2011
15:02–15:58
5
738
No clouds
22-04-2011
19:17–20:09
4
846
No clouds
23-04-2011
12:10–13:34
6
1008
2414
26-04-2011
16:07–22:00
8
1152
No clouds
30-04-2011
09:46–10:48
6
855
4018
Summer 2011
06-07-2011
07:40–17:55
10
1143
No clouds/28133
08-07-2011
16:43–20:53
2
1208
1787
12-07-2011
08:19–10:01
6
724
506
Summer 2012
21-06-2012
15:12–16:11
2
980
No clouds
23-06-2012
05:55–11:07
12
1024
622
24-06-2012
11:02–15:16
8
1076
No clouds
26-06-2012
07:43–13:00
10
948
No clouds
29-06-2012
07:58–13:19
10
1144
No clouds
30-06-2012
09:20–20:30
8
1100
No clouds/8214
01-07-2012
08:35–21:40
10
1212
No clouds
04-07-2012
13:30–18:05
8
1192
654
10-07-2012
08:44–21:07
8
1268
No clouds
11-07-2012
08:37–23:20
14
1196
No clouds/7225
1 From 16:00 UTC (last six profiles). 2 Variable for
half of the time. 3 Variable for half of the time. 4 Clouds until
11:30 UTC. 5 From 22:19 UTC (last two profiles).
Vertical profiles were measured from two sampling sites: during spring, the
vertical profiles were taken at the Italian CNR Gruvebadet sampling site
(78∘55′03′′ N, 11∘53′40′′ E; Fig. 1b) to assure a large
distance to the Ny-Ålesund village. During summer, the tethered balloon
measurements were operated at the German–French AWIPEV research base
(78∘55′24′′ N, 11∘55′15′′ E) to lie in the proximity of
the Ny-Ålesund harbor (600 m) allowing the measurement of ship plume
diffusion (Sect. 3.3). Table 1 lists the dates of the campaign (25 measurement days), the number of flights (197 measured profiles), the
maximum altitudes (∼ 700–1300 m) and the cloud base height
(clouds present for 48 % of campaign). The aerosol and meteorological
measurements were carried out both at ground and along the profiles as
described in the following sections.
Ground-based measurements
Ground-based measurements were carried out at the Gruvebadet laboratory
(Fig. 1b) where the distance (1.2 km southern Ny-Ålesund) and the
limitations established for snowmobile traffic and other potentially
contaminant activities limits the impact from local emissions.
The Gruvebadet laboratory is equipped with a series of instruments aimed at
measuring aerosol physical and optical properties, and collecting samples
for chemical analysis (Sect. 2.1.1). The aerosol size distribution was
measured using a scanning mobility particle sizer (TSI-SMPS 3034, 54 size
classes, 10–487 nm) coupled with an aerodynamic particle sizer (TSI-APS
3321, 52 classes, 0.5–20 µm). The two coupled systems measure one size
spectrum every 10 min (Giardi et al., 2016). PM samples were collected by
high-volume and low-volume samplers. For the purpose of the present paper,
PM10 samples collected using two TECORA SkyPost low-volume samplers (EN
12341; PM10 sampling head, flow 2.3 m3 h-1; PTFE and
quartz microfiber filters, Ø = 47 mm) were considered. Sampling was
carried out in ambient conditions: pressure and temperature were
continuously monitored in order to maintain the constant flow rate of 2.3 m3 h-1. The first sampler collected PM10 for 24 h on Teflon
filters (Pall R2PJ047) to determine the ionic fraction, while the second one
collected PM10 for 96 h on pre-fired quartz microfiber filters (CHM QF1
grade) to determine organic and elemental carbon (Sect. 2.1.1). The Teflon
filters were conditioned for 48 h (25∘C and 50 % relative
humidity) before and after the sampling, then weighted by a five-digit
microbalance (Sartorius ME235P). The reproducibility error for filter
weighing was lower than 5 % (experimentally evaluated). After sampling,
filters were individually sealed in pre-washed (with Milli-Q water, 18.3 MΩ cm)
polystyrene filter containers and stored at -20 ∘C until analysis.
Aerosol chemistry measurements at ground level
PM10 samples collected at ground level at Gruvebadet were analyzed to
determine first the water-soluble ionic fraction. Half of each PM10
Teflon filter was extracted in 10 mL of ultrapure water (Milli-Q, 18.3 MΩ cm resistivity) by ultrasonic bath for 20 min. Filters
manipulation was carried out under a class-100 laminar-flow hood, in order
to minimize contamination risks. Inorganic cations and anions together with
organic anions, were simultaneously measured by a triple Dionex
ion-chromatography system, equipped with electrochemical-suppressed
conductivity detectors. Cations (Na+, NH4+, K+,
Mg2+ and Ca2+) have been determined by a Dionex CS12A-4 mm
analytical column with 20 mM H2SO4 eluent. Anions (Cl-,
NO3-, SO42- and C2O42-) were measured by
a Dionex AS4A 4 mm analytical column with a 1.8 mM Na2CO3 / 1.7 mM
NaHCO3 eluent, while F- and some organic anions (acetate,
glycolate, formate and methanesulfonate) were determined by a Dionex AS11
separation column by a gradient elution (0.075 to 2.5 mM
Na2B4O7 eluent) (Udisti et al., 2004; Becagli et al., 2011).
The detection limit (ng m-3) of each analyzed chemical component is
reported in the Supplement (Table S1) together with the measured
ambient ion concentrations. All the analyzed chemical components were
largely above the detection limit.
The contribution of sea salt and crustal components in Ny-Ålesund is
not-negligible (Udisti et al., 2016; Giardi et al., 2016; Moroni et al.,
2015). Thus, Na+, Ca2+ and SO42- (which originate from
both these sources) were apportioned between sea-salt (ss) and non-sea-salt
(nss) fractions on the basis of known w/w (weight / weight) ratios in sea
water and Earth crust (Udisti et al., 2012, 2016; Giardi et al., 2016; Becagli et
al., 2012):
tot-Na+=ss-Na++nss-Na+tot-Ca2+=ss-Ca2++nss-Ca2+ss-Na+=tot-Na+-0.562nss-Ca2+nss-Ca2+=tot-Ca2+-0.038ss-Na+,
where 0.562 represents the w/w Na+ / Ca2+ ratio in the crust (Bowen,
1979) and 0.038 is the Ca2+ / Na+ w/w ratio in seawater (Nozaki,
1997). Similarly, the ss-SO42- fraction was calculated from the
ss-Na+ using the 0.253 SO42- / Na+ w/w ratio in seawater
(Bowen, 1979). The crustal fraction of sulfate (cr-SO42-) was
determined from the nss-Ca2+ using the 0.59 SO42- / Ca2+
w/w ratio in the uppermost Earth crust (Wagenbach et al., 1996). Finally,
the nss-nc-SO42- fraction, which can be due to anthropogenic or
secondary formed aerosol, was calculated by subtracting the ss-SO42-and cr-SO42- contributions from the total
SO42- concentrations.
The organic carbon (OC) and elemental carbon (EC) fractions were determined
in PM10 samples using a thermo-optical transmission (TOT) method
following the NIOSH protocol. The organic matter (OM) was calculated by
multiplying the OC fraction by 2.1 (Turpin and Lim, 2001) typical for remote
sites with a large fraction of secondary aerosols. Table S1 also reports the detection limit for EC and OC.
Meteorological context
Meteorological parameters are currently measured at different sites in
Ny-Ålesund. The German–French AWIPEV research base operates surface
meteorology measurements with 1 min time resolution, including temperature
and relative humidity at 2 m height, wind speed and direction at 10 m
height and pressure at station level close to the summer campaign balloon
launch site (Maturilli et al., 2013). The cloud base height above the station
is retrieved using a Vaisala LD-40 ceilometer. Daily radiosoundings
(11:00 UTC) by the AWIPEV observatory provide auxiliary data for the aerosol
profile analysis.
Since 2009, the Italian National Research Council (CNR) has operated the
Amundsen–Nobile climate change tower (CCT), providing meteorological,
micrometeorological, radiation and snow measurements continuously all
year long (Mazzola et al., 2016a). Conventional and micrometeorological
parameters are measured at different heights (4 and 3 levels, respectively)
in order to investigate their vertical variations in different conditions.
Dai et al. (2011) and Mazzola et al. (2016b) found that both in the
Adventfjorden and in Kongsfjorden, where Ny-Ålesund is located, the
atmosphere is stable for about 50 % of the time during the year, by
analyzing micrometeorological data. The term stability refers to the
propensity of air masses to move vertically: stable air resists any vertical
motion, while unstable air masses are prone to vertical movements. A parcel
of air is stable/unstable if the temperature lapse rate is
lower/higher than the adiabatic one, i.e., if the potential temperature is
increasing/decreasing with height, respectively. In stable stratification,
turbulence and vertical mixing are suppressed, leading to trapping of
pollutants near ground level. On the above grounds, the spring 2011, summer 2011 and summer 2012 campaign
periods can be put into a climatological context.
Vertical profile measurements
Vertical profile measurements have been carried out by means of a
kytoon-shaped, helium-filled tethered balloon (length 8 m, Ø = 3 m, volume
55.0 m3, payload 25 kg; Fig. 1c). The tethered balloon was designed
to fly in severe wind conditions. However, the presence of the payload
limits the balloon flights from low to moderate wind conditions (< 10 m s-1).
The tethered balloon was equipped with an instrumental package consisting
of
an optical particle counter (OPC GRIMM 1.107; 31 size classes between 0.25 to
32 µm, 6 s sampling time) for the particle number size distribution
determination;
a miniaturized electrical particle detector (miniDiSC, Matter Aerosol) to
measure the total particle number concentration (1 s sampling time);
two micro-aethalometers: the microAeth® AE51
and a prototype (1–60 s sampling time);
a meteorological station (LSI-Lastem: pressure, temperature and relative
humidity, 6 s sampling time).
During the period 11–30 April 2011, the Vaisala tethersonde TTS111
(pressure, temperature, relative humidity, wind speed and wind direction; 1 s sampling time) was also used.
The maximum height reached during each flight depended on atmospheric
conditions and, for the majority of the profiles, was between 0.7 and 1.3 km. An electric winch controlled the ascent/descent rates that were set at
40.0 ± 0.1 m min-1.
A deeper description of each instrument is reported below.
Size distribution data
In this study, the total aerosol concentration and the number size
distribution along height were measured using a coupled miniDiSC–OPC
(λ= 655 nm) system.
The miniDiSC is a miniature diffusion size classifier, a small and portable
instrument (4 × 9 × 18 cm, 670 g, 8 h of battery supply)
(Fierz et al., 2011). The aerosol is first charged in a
standard positive unipolar diffusion charger (the average charge is
approximately proportional to the particle diameter). The charged particles
flow through a diffusion stage (an electrically insulated stack of stainless
steel screens connected to a sensitive electrometer that collect the finest
particles) and into a second stage (equipped with a HEPA filter) where the
current of larger particles is measured with an electrometer. The miniDiSC
has a d50 cutoff at 14 nm. Thus, the instrument underestimates particle
number concentrations for particles smaller than 20 nm (nucleation mode). As
a result, the miniDiSC counts only partially the nucleation mode, while it
allows a whole determination of Aitken and accumulation-mode particles. As
demonstrated by Fierz et al. (2011), a bimodal lognormal aerosol size
distribution with a fixed accumulation mode at 100 nm and a varying
nucleation mode at 20 nm introduces an underestimation of about -2 to 10 % of the total aerosol concentration in the miniDiSC response. The
particle number determination is robust, and the error never exceeds 20 %.
The OPC used in the campaign was the model Grimm 1.107 that counts and
classifies the aerosol in 31 size classes between 250 nm and 32 µm.
As reported in the literature (Ferrero et al., 2014; Howell et al., 2006; Heyder
and Gebhart, 1979), OPC size classification of the aerosol particles is a
function of their ability to scatter the laser light under the assumption of
spherical particles. The aerosol particles are classified in terms of their
optical equivalent diameter, which is defined as “the diameter of a sphere
of known refractive index (that of polystyrene latex spheres used for of
calibration) that scatters light as efficiently as the real particle in
question”. This effect usually results in an undersizing of the size
classification, due to the higher refractive index of the polystyrene latex
spheres (PSL spheres, m=1.58 at 655 nm; Ma et al., 2003)
used in the OPC calibration compared to ambient aerosol
(Guyon et al., 2003; Liu and Daum, 2008;
Schumann, 1990). In order to derive a proper size classification of the
aerosol over Svalbard, the undersizing issue was solved by correcting
the OPC size channels to account for the ambient aerosol refractive index
m. The OPC response function (S: the partial light-scattering cross section of
the particle related to the specific optical design of the OPC) was computed
at 655 nm as follows (Baron and Willeke, 2005; Heyder and
Gebhart, 1979):
Sθ0,ΔΩ,x,m=λ24π2∫∫ΔΩiθ,Φ,x,msinθdθdΦ,
where θ0 represents the mean scattering angle of the optical
arrangement, ΔΩ the receiver aperture, x the dimensionless
size parameter, m the refractive index and i(θ, φ, x, m) the Mie
scattering function composed by the perpendicular and parallel components
i1(θ, x, m) and i2(θ, x, m), respectively. The optical
arrangement of the OPC 1.107 consists of (1) a wide angle parabolic mirror
(121∘, from 29.5 to 150.5∘, θ0= 90∘) that focuses scattered light on the photodetector
located on the opposite side; (2) 18∘ of direct collected scattered
light on the photodetector (from 81 to 99∘, θ0= 90∘) (Heim et al., 2008).
The response function was calculated both for PSL spheres (SPSL) and for
ambient aerosol (SAMB). The refractive indexes of ambient aerosol used
in SAMB calculations were obtained from the closest AERONET site
(Horsund; 77∘00′04′′ N, 15∘33′37′′ E) for
spring 2011 and summer 2011–2012: 1.544 + 0.013i and 1.535 + 0.015i, respectively.
These refractive indexes were determined at 674 nm (the closest AERONET
wavelength to the OPC laser wavelength of 655 nm) and were close to those
determined at 530 nm at Gruvebadet site (range 1.4–1.8 during 2010 and 2011;
Lanconelli et al., 2013). Table S2 shows the new size-corrected channels in
comparison with the PSL spheres' equivalent ones. The new channels were used
to define three broad size ranges (detailed below) to evaluate the
vertical behavior of aerosol.
The coupled miniDiSC–OPC (λ= 655 nm) system measurement range
covers the relevant region of the aerosol number size distribution. In order
to study the behavior of different size classes along height, three aerosol
number concentration size ranges were selected:
the number concentration of aerosol between 14 nm (d50 of miniDiSC)
and 260 nm (cf. Table S2) obtained as the difference between the total
number concentration measured by the miniDiSC and that measured by the OPC,
hereinafter indicated as N14-260;
the number concentration of aerosol between 260 nm (lower limit of OPC)
and 1200 nm hereinafter indicated as N260-1200;
the number concentration of aerosol above 1200 nm, hereinafter indicated
as N>1200.
The mode N14-260 includes a small fraction of the nucleation mode (from
14 to 20 nm), the totality of the Aitken mode (20–100 nm) and a fraction of
the accumulation mode (from 100 to 260 nm). The mode N260-1200 includes
most of the accumulation-mode particles. Finally, mode N>1200 covers
the totality of giant nuclei mode.
The accuracy of both miniDiSC and OPC measurements was investigated
comparing the lowermost portion of their measurements along vertical profile
with SMPS + APS data collected at ground level at Gruvebadet. This
comparison was performed during spring 2011 to avoid any contamination from
ship plumes arriving from Ny-Ålesund harbor towards Gruvebadet in summer
(balloon soundings were conducted from the Koldewey station instead of from
Gruvebadet; Sect. 2.1). The comparison of N14-260 (miniDiSC vs. SMPS)
and of N>260 (OPC vs. SMPS + APS) was characterized by an excellent
correlation (R2 > 0.9; linear best fit close to the ideal
one) with an average error of 7 and 16 % for both N14-260 and
N>260, respectively (Fig. S1a–b in the Supplement). These
results highlight the reliability of measurements carried out along the
vertical profiles, an important feature considering the low aerosol
concentration values and their variation, which are present within the
Arctic (Sect. 1).
Number concentration data were also used in Sect. 3.2.4 to estimate the
contribution of locally formed aerosol. The method is based on the N / BC
ratio, developed by Rodríguez and Cuevas (2007) and successfully
applied in Europe by Reche et al. (2011).
The basic concept of this method is that highest values of N / BC ratios
(i.e.,
the lowest BC fraction values) occur during secondary aerosol formation in
the atmosphere (Reche et al., 2011; Dall'Osto et
al., 2011, 2013). The methodology is as follows:
N2=N-N1N1=S1×BC,
where N2 represents the secondary aerosol concentration locally formed
in the atmosphere, N is the measured aerosol number concentration and
N1 is the aerosol number concentration already present in the
background air. S1 represents a reference value for the N / BC ratio
(expressed as particles cm-3 ng-1 m-3 of BC) in the background air.
The parameter N1 is calculated from the parameter S1 multiplied by
the measured BC concentration (see Sect. 2.2.2). S1 can vary
from ∼ 2 to ∼ 9, while the N / BC ratio during
secondary aerosol formation reaches values higher than ∼ 15–20
and up to ∼ 100–200 (Reche et al., 2011; Dall'Osto et al.,
2011, 2013). The differences in S1 values determined in different
sites can be caused by (1) the use of different particle counters (with
different d50 cutoff), as lowest S1 values are usually observed
when devices with largest d50 cutoff are used; (2) the influence of
the ambient air conditions on the secondary aerosol formation. Thus, S1
is site instrument specific and has to be determined onsite depending on
the used particle counter. If the Rodríguez and Cuevas (2007) method is
applied to ground-based temporal data series, S1 can be obtained as the
minimum N vs. BC slope observed during the day (Reche et al., 2011).
However, in the case of measured vertical profiles, the values of S1
were taken as that of background aerosol above the aerosol stratifications
described in Sect. 3.2.4.
Black carbon
BC have been determined using two micro-aethalometers: the
microAeth® AE51 and a prototype (Magee
Scientific; 250 g, 117 × 66 × 38 mm3). Adopting the nomenclature
recommended by Petzold et al. (2013), by Andreae and Gelencser (2006) and by
other authors (Gilardoni et al., 2011; Stohl et al.,
2013; Eckhardt et al., 2013), we refer to the measured parameter as
equivalent black carbon (eBC) due to the absence of an overall agreed
reference material, linking light absorption to the empirically defined BC
mass concentration. In agreement with the above cited literature, we also
report the absorption coefficient values.
AE51 and the prototype were identical with the exception that the prototype
measured at 2λ (370 and 880 nm) while AE51 only at 880 nm. At the
time of campaign the prototype had just been developed and was used instead
the AE51 on the balloon platform during the spring 2011 campaign only when
necessary (i.e., AE51 in charge) to ensure the continuity of measurements
during the campaign.
In both the aethalometers, the aerosol containing BC was continuously sampled
onto a PTFE-coated borosilicate glass fiber filter
(Fiberfilm™
filters, Pall Corporation) where the light attenuation (ATN) was measured at
880 nm relative to a clean part of the filter. ATN was calculated as
ATN= 100⋅ln(I0/I),
where I0 and I are the light intensities transmitted throughout a
reference blank spot and the aerosol-laden 3 mm diameter sample spot of the
filter, respectively.
The attenuation coefficient of the particles collected on the filters,
bATN, was derived from ATN as follows (Weingartner
et al., 2003):
bATN=A100QΔATNΔt,
where ΔATN indicates the ATN variation during the time period Δt, A is the sample spot area (7.1 × 10-6 m2) and Q is the
volumetric flow rate (2.5 × 10-6 m3 s-1 for the AE51 and
4.42 × 10-6 m3 s-1 for the prototype).
Finally, to determine the eBC ambient concentration the apparent mass
attenuation cross section (σATN=12.5 m2 g-1) is
needed; it is defined for the eBC collected on the PTFE-coated borosilicate
glass fiber filter. The σATN value (12.5 m2 g-1) was
obtained by comparing the eBC values measured with the
microAeth® model AE51, with an AE31 aethalometer (880 nm
wavelength) operating in a test chamber with different eBC concentrations at
low attenuation values. The comparison was then repeated using ambient air
(Ferrero et al., 2011a). This value is not far from the σATN
values of 15.2 and 15.9 m2 g-1 reported in
Eleftheriadis et al. (2009) which recorded 10 years of eBC measurements in
Ny-Ålesund at the Zeppelin station with the aethalometers AE9 and AE31.
The difference between these values results from the use of different filter
materials to collect the sample in the different aethalometers, which was
quantified in Ferrero et al. (2011a) and Drinovec et al. (2015).
The eBC concentrations were determined as follows:
eBC=bATNσATN.
The accuracy of eBC measurement was investigated. The AE51 and the prototype
measurements carried out simultaneously agreed very well (R2=0.852;
slope = 0.976; Fig. S1c). This result was important
as it was obtained with two different flow rates (2.5 × 10-6 m3 s-1 for the AE51 and 4.42 × 10-6 m3 s-1 for the
prototype).
However, a large scatter is present at low eBC concentrations (i.e.,
10–20 ng m-3; see Fig. S1c). Thus, the absolute error (in percentage)
of each eBC value (considering the average of the two aethalometers) was
calculated for intervals of 5 ng m-3 of concentrations. At low
concentrations, the error can reach up to 90 % and more (Fig. S2a). This
error decreases with increasing concentration, dropping below
20 ng m-3 for eBC concentrations above 5 ng m-3. The relative
error lies below 20 % for both the average and the 90th percentile at eBC
concentrations above 20 ng m-3. Thus, it is possible to consider this
value as the limit above which a single eBC measurement point is not affected
by instrumental noise. Nevertheless, this limit is close to the BC
concentrations that have been previously measured in the Arctic
(Eleftheriadis et al., 2009). In this respect, we
note that the BC profiles presented in the paper are an average of many
measurements, hence the effect of the noise on the reported eBC
concentrations is further reduced. The aim of this paper is to determine the
seasonal phenomenology of the aerosol behavior along vertical profiles,
classifying the collected experimental data according to their shape and
averaging them for each season. This is very important as even though the error in
percentage of each data point can reach high values (especially at low
concentrations), the average of the data stabilizes the instrumental
fluctuations. This effect is demonstrated by Fig. S2b which reports the correlation between the BC concentrations (AE51
and prototype) averaged on the same intervals of 5 ng m-3 used in
Fig. S2a (R2= 0.986; slope = 1.017).
The above-reported analysis underlines a critical situation for summer
because, as reported in Eleftheriadis et al. (2009), the eBC concentration
range expected in summer is ∼ 0–10 ng m-3. Therefore,
summer eBC data were used here only to highlight the impact of shipping
emissions on the Arctic background concentrations along the atmospheric
column. Due to high ship impact (Sect. 3.4), the performance of the
micro-aethalometers was suitable and reliable for the purpose of this
application.
In addition to eBC, the micro-aethalometers also allow the determination of
the aerosol absorption coefficient, babs, that was calculated as
follows:
babs=bATNC⋅R(ATN),
where C and R(ATN) are the multiple scattering optical enhancement factor and the
aerosol loading factor, respectively. Briefly, the constant optical
enhancement factor C compensates for the enhanced optical path through the
filter caused by multiple scattering induced by the filter fibers themselves
(Schmid et al., 2006; Arnott
et al., 2005; Weingartner et al., 2003). The
parameter R(ATN) compensates for the nonlinearity – the loading effect due to
reduction of the measurement sensitivity due to the saturation caused by the
collected sample on the filter. The compensation with the parameter
R(ATN) is needed only when ATN becomes higher than 20
(Schmid et al., 2006; Arnott
et al., 2005; Weingartner et al., 2003). In this
study, the experimental design allowed us to neglect the use of R(ATN): all eBC
vertical profiles were conducted in the clean Arctic environment and the
filter tickets were changed regularly to always keep ATN lower than 20 as
recommended by Weingartner et al. (2003). For the
AE51 and the prototype, the only parameter C available in the literature is
2.05 ± 0.03 (at λ=880 nm)
(Ferrero et
al., 2011a), even though recently Ran et al. (2016) proposed a C value of
2.52 for ground-based measurements in China.
The C value of 2.05 ± 0.03 was determined over Milan in Ferrero et al. (2011a) and thus a brief description is necessary to determine its
applicability in the Arctic area. The parameter C was determined using data
collected both below the mixing layer and above it, in a cleaner atmosphere,
along the vertical profiles (Ferrero et al., 2011a). During the C
determination, a new filter ticket was used for each profile. As a result,
ATN never reached values higher than 20 (average ATN was 5 ± 1) and the
total amount of aerosol collected on each filter during the C determination
was negligible. Therefore, the determined C was exclusively dependent on the
filter material and the AE51 instrumental geometry. This ensures the
negligible influence of the particles in the filter matrix on the C value.
The reliability of the obtained C (2.05 ± 0.03) was demonstrated in
Ferrero et al. (2014) below the mixing layer and in free troposphere.
Finally, it should be noted that absorbing non-BC particles may contribute to
the signal in aethalometers (i.e., brown carbon, dust). However, BrC is
characterized by negligible absorption in the infrared (Andreae and
Gelencsér, 2006), which is the wavelength range of the
eBC measurements (microAeth AE51 uses 880 nm). In this respect, Massabò
et al. (2015) showed the potential contribution of BrC to the determination
of eBC to be below 10 %.
To estimate the possible influence of BrC on eBC measurements carried out
during the spring 2011 campaign, the data collected with the microAeth
prototype at 370 and 880 nm were considered. They highlighted a BrC positive
artifact on eBC measurements less than 10 % during the campaign. Details
are reported in the Supplement.
Vertical profiles of N14-260 (green line), N260-1200
(blue line), potential temperature (red line) and relative humidity (light
blue line) measured over Ny-Ålesund on (a) 7 April 2011 (12:51–13:10 UTC); (b) 1 April 2011 (06:30–06:57 UTC); (c) 23 April 2011
(13:21–13:34 UTC); (d) 6 April 2011 (18:03–18:22 UTC).
Meteorological data and aerosol stratifications
Meteorological data along height allowed the determination of the absolute
height of the balloon using the hypsometric equation; due to a change in
April 2011 in the measuring system (Sect. 2.2), a comparison of the
altitude obtained by the LSI-Lastem and Vaisala tethersonde was conducted
during several target flights. The result (R2=0.997; slope = 0.999;
Fig. S1d) demonstrated the accuracy of the height
determination. The measurement of altitude is fundamental in the study of
the vertical aerosol properties in relationship to meteorological
parameters. In fact, vertical aerosol profiles allow the determination of
the height of aerosol stratifications by means of a gradient method applied
to aerosol concentration profiles, as suggested by Seibert
et al. (2000).
The gradient method is based on the determination of the minimum value of
the vertical derivative of the aerosol concentration. The use of gradient
method to determine the aerosol mixing height has been demonstrated at lower
latitudes in previous works (Ferrero et al., 2007, 2011a, b, and 2012;
Sangiorgi et al., 2011; Di Liberto et al., 2012). However, in remote areas
such as the Arctic, several processes other than dispersion can shape the
aerosol profiles. The two most important ones are (1) differential advection
(Tunved et al., 2013) and (2) a lack of emission of aerosol from ground.
These processes should generate a vertical structure not directly related to
the PBL height.
Therefore, in the present work the gradient method has been limited to
individuate aerosol stratification heights (ASh), even if these related
to the behavior of meteorological variables governing the behavior of the
PBL as will be addressed in Sect. 3.1.
The ASh will be used in the next sections to calculate averaged
aerosol and eBC profiles (Sect. 3.2 and 3.3). In fact, in order to
investigate the variation of aerosol properties with height, vertical
profiles were statistically averaged. As reported in previous works
(Ferrero
et al., 2011a, 2012, 2014), a way to average vertical profile data by
taking their main gradients (ASh) into account is to consider the
relative position of each measured data point in respect to the ASh.
Thus, vertical profiles were first normalized, introducing a standardized
height (Hs) calculated as follows:
Hs=z-AShASh,
where z is the height above ground. Hs assumes a value of 0 at the
ASh and values of -1 and 1 at ground level and at twice the ASh,
respectively.
Examples of ASh, accompanied with the corresponding potential
temperature (θ) and RH profiles, are presented in Fig. 2a–d. The
presented data accurately describe the vertical distribution of the aerosol
and its properties in the first kilometer above Ny-Ålesund. Moreover,
they allowed obtaining different pieces of information.
The absence or the presence of marked aerosol stratifications (AS) is
notable. When present, the altitude at which they occur (ASh) was
determined by the gradient method, described above. This is the first
obtainable information.
A second piece of information was the size-dependent vertical behavior of
aerosol concentrations. Figure 2b and c highlight a similar behavior for
both N14-260 and N260-1200, while Fig. 2d shows different
behavior for N260-1200 due to a concentration change located at a
different altitude. Finally, we obtained the magnitude of the observed
concentration change at each ASh for each size range (N14-260,
N260-1200and N>1200) and season.
The analysis and combination of these three types of information allowed to
classify, as a function of seasonality, the altitude, magnitude and
frequency of aerosol stratifications. Furthermore, it has been possible to
shed some light on the dynamics underlying the seasonal phenomenology found
during the field campaigns. A detailed discussion of these points is
reported in the following section.
Results and discussion
Vertical profiles of aerosol number size distribution and eBC concentrations
were measured to assess changes in aerosol properties within the vertical
column in the Arctic region. The results obtained along vertical profiles
are discussed in order to first highlight the vertical behavior of the
ASh in relation to the main atmospheric meteorological parameters
(Sect. 3.1). Then, vertical aerosol properties are discussed in details
for springtime (Sect. 3.2) and summertime (Sect. 3.3). All averaged data
are reported hereinafter as mean ± mean standard deviation.
Aerosol stratifications: seasonal vertical frequency distribution and
relationship with meteorology
As reported in Table 1, about 200 profiles were measured during three campaigns
in spring 2011, summer 2011 and summer 2012. Here, the ambient
conditions under which the vertical profiles were measured are briefly
described.
First of all, the observational periods (spring 2011, summer 2011 and summer 2012) were addressed in a climatological context. In this respect, the
temperature measured in spring 2011 was within the standard deviation range
of the long-term observations, while a 10–day period at the end of April 2011 was slightly warmer than the climatological mean (Fig. S3a). The
temperatures during the summer seasons 2011 and 2012 were mostly within the
range of the long-term observations (Fig. S3b). Neither of the campaign
periods was conducted under exceptional meteorological conditions, so the
vertical profile measurements can be considered to have been obtained under
typical meteorological conditions representative for the Ny-Ålesund
environment.
We note, however, that the tethered balloon measurements have limitations
with respect to the balloon's launch conditions (Sect. 2.2). Particularly, balloon
profiles were measured in low-wind conditions, as it is very difficult to
launch the balloons during high winds. This introduces a bias in respect to
average meteorological conditions above the launch site. The maximum wind
speed measured at the Amundsen-Nobile climate change tower (Sect. 2.1.2)
during balloon flights was lower than that during the whole period of the
campaign (April 2011, June and July 2011 and 2012): 4.9 and 10.7 m s-1 (springtime and summertime balloon profiles) compared to 27.9
and 16.3 m s-1 (full spring 2011 and summers 2011 and 2012). Table 1 resumes the conditions for all the measured profiles. The majority of
vertical profile measurements was conducted under clear sky conditions (no
clouds) or with clouds with base height above the balloon payload.
Thus, it is possible to assert that the measured profiles and the seasonal
phenomenology described hereinafter are representative of typical Arctic
springtime and summertime periods mainly for low wind and clear sky
conditions.
Figure 2a–d highlight different atmospheric dispersal conditions upon
Ny-Ålesund. Although these four case studies are not illustrative and
comprehensive of the whole data set, their discussion helps to illustrate
the seasonal and size-dependent behavior of the frequency distribution of
the ASh with altitude.
An example of homogeneous dispersion of aerosol (independent from its size)
in the lower troposphere is shown in Fig. 2a. Decreasing potential
temperature with height indicates atmospheric instability, allowing the
vertical mixing of air masses by convection. Homogeneous aerosol profiles
with convective conditions were found in 15 % of the profiles in spring
and 37 % of the profiles in summer, respectively. Convective conditions
generally are observed more frequently during summer in Ny-Ålesund related to
the different level of radiation energy at disposal and surface properties.
In summer, homogenous profiles were observed often (37 %) than in spring
(15 %), due to a synergy of the higher solar power density at disposal
(186.4 ± 71.2 W m-2 in summer and 109.2 ± 35.9 W m-2 in
spring) together with a lower albedo (0.15 ± 0.01 in summer and
0.87 ± 0.04 in spring) induced by the summer snowmelt in Svalbard
(Mazzola et al., 2016b). The resulting change in surface energy balance
affects the atmospheric stability, from the more stable conditions and
inversion situations with snow cover to the unstable conditions favoring
mixing within the boundary layer once the snow cover has disappeared.
Vertical frequency distribution of the ASh for
N14-260, N260-1200, N>1200, θ and RH during spring in
panels from (a) to (d) and in summer in panels from (e) to (h).
In spring, the presence of different layers of aerosol, separated by abrupt
changes in the aerosol properties (i.e., concentration) along height were
observed more frequently than the homogeneous mixing conditions (Fig. 2b–d). The presence of aerosol stratification was complementary to the
homogenous profiles and thus occurred for 85 and 63 % of cases during
spring and summer, respectively.
Within stratified conditions it was possible to determine the ASh for
each aerosol size and season together with the altitude of sharp changes
observed also for θ and RH. For example, Fig. 2c shows ASh at 630 m in agreement with the vertical gradients of θ and RH. On
the other hand, Fig. 2d shows a first ASh at 134 m only for
N14-260, while N260-1200 appeared homogenously distributed until
a second ASh was located at 674 m. The first ASh, detected for the
smallest particles, was related to a ground-based θ inversion, while
the second ASh, detected for the accumulation-mode particles, was
related to an elevated θ inversion.
The aforementioned case studies helped us to introduce the description of
the ASh frequency distribution with altitude and season. The ASh
were used independently from the sign of the aerosol concentration change
(either positive or negative; Fig. 2b–c) and were computed separately for
each broad size range (i.e., Fig. 2d). The resulting frequency distribution
with altitude of the first ASh is reported in Fig. 3a–d and
e–h for spring and summer, respectively. It has to be underlined
that sometimes it was possible to detect up to two ASh for each
profile. This situation was observed for ∼ 30 % of profiles
(characterized by the presence of aerosol gradients) in spring and summer,
due to the limited maximum altitude reached by the balloon during the flight
(usually between 0.7 and 1.3 km). The behavior of the second ASh is
shown in the Supplement (Fig. S4) to support the
description of the behavior of the first aerosol stratification with
altitude.
Focusing on each season and considering first the springtime ASh
vertical frequency distributions (Fig. 3a–d) a common behavior can be
first observed for both the three size ranges (N14-260, N260-1200
and N>1200) and meteorological parameters (θ and RH): all of
them showed a bimodal distribution characterized by a minimum within the
∼ 400–500 m height range. Maturilli and Kayser (2016)
identified a frequent occurrence of a temperature inversion layer in the
shear zone above the mountain ridges; this phenomenon is typically present
throughout the year, leading to a decoupling of the lowermost kilometer of
the atmosphere from the free troposphere above. In between the mountains,
the atmosphere is characterized by wind channeling along the fjord axis,
disturbed by, e.g., glacier outflow or land–sea breeze. Thus, the observed
separation of ASh most likely relates to the separation of the
atmospheric flow.
Springtime statistical mean profiles of N14-260 (green line),
BC (black line), N260-1200 (blue line), N>1200 (magenta line),
θ (red line) and RH (light blue line) along height over
Ny-Ålesund for the four typologies of vertical profiles:
(a)–(c)
homogeneous profiles (HO); (d)–(f) positive gradient profiles (PG);
(g)–(i) negative gradient profiles (NG); (l)–(n) decoupled negative gradient profiles
(DNG).
Moving forward, some differences were then found in the behavior of each
aerosol size range below and above the minimum at 400–500 m. While
N260-1200 and N>1200 (Fig. 3b–c) appeared equally distributed
below and above the minimum at 400–500 m, N14-260 showed highest
frequencies of ASh below 400 m (Fig. 3a). Particularly, the 82 % of
ASh for N14-260 were located below 400 m and showed a clear
maximum peak close to the ground (0–100 m) with a frequency of 38 %.
Moreover, the average values of ASh for N14-260 below and above
the minimum at 400–500 m were 143 ± 13 m and 669 ± 32 m,
respectively. Conversely, ASh for N260-1200 and N>1200
occurred for 52 and 51 % below 400 m peaking in the 100–200 m range
(22 and 21 %, respectively) with average values of 208 ± 19 m and
177 ± 14 m, respectively. Above 400 m the ASh for N260-1200
and N>1200 peaked in the 600–700 m range (22 and 19 %,
respectively) with average values of 672 ± 16 and 652 ± 20 m.
The observed lack of symmetry between N14-260 and N260-1200–N>1200 is explained in Fig. 2d, where a decoupled trend for
N14-260 and N260-1200 is shown. As stated above, the behavior of
smallest particles was in that case related to a ground-based θ
inversion. Considering the whole data set, the behavior of the vertical
frequency distribution of the first gradient of both θ and RH was in
agreement with that of ASh for N14-260. In this respect, a maximum
peak for both θ and RH was found close to the ground (0–100 m), with
a frequency of 49 and 38 %, respectively (average values below the
minimum at 400–500 m for gradients for θ and RH of 117 ± 12
and 669 ± 32 m). Interestingly, the frequency of ground-based θ
inversions (49 %) was higher than that of ASh for N14-260
(38 %). This feature is due to the vertical profile along which, even in the
presence of a ground-based θ inversion, N14-260 did not show
any variation of concentration; an example is reported in Fig. 2b. Thus,
the presence of a ground-based θ inversion appears as a necessary
but not sufficient condition to observe the aforementioned behavior (resumed
by Figs. 2d, 3a and d). The phenomenology and the aerosol dynamic
responsible of this behavior (together with that of N260-1200–N>1200) will be addressed and discussed in the following Sect. 3.3.
We describe below the summer ASh behavior. Figure 3e–h (and Fig. S4)
show that even in summer, the multilayered structure persisted and was also
characterized by a bimodal distribution, as in spring, but with a higher
minimum (than in spring) that ranged approximately between 500 and 600 m.
The summer ASh for all size ranges and the gradient for θ and
RH peaked between 100 and 300 m: 78 % (N14-260; average value
276 ± 19 m), 71 % (N260-1200; average value 269 ± 18 m),
76 % (N>1200; average value 272 ± 18 m), 83 % (θ;
average value 262 ± 18 m) and 79 % (RH; average value 268 ± 18 m). Figure 3e–h and the aforementioned data highlight that the vertical
frequency distribution for ASh of all sizes and gradients for θ
and RH behave similarly. This phenomenon, different from that observed in
spring, will be addressed and discussed in Sect. 3.3. As a final
conclusion, a multilayered structure was found over Ny-Ålesund both in
spring and summer (see also Fig. S4), and the most important atmospheric
thermodynamic parameters (θ and RH) indicated the role of
meteorology in shaping the aerosol vertical profiles. This result is of
great importance as the majority of the aerosol measurements conducted in
the Arctic area are ground based and thus, it is necessary to understand
their validity with altitude.
Meteorological parameters (temperature, relative humidity, wind
speed, pressure) measured at the CCT at different levels (33, 10, 5 and 2 m)
and averaged (timely coincident) for each profile class.
Season
Profile type
T (∘C)
RH (%)
WS (m s-1)
P (hPa)
33 m
10 m
5 m
2 m
33 m
10 m
5 m
2 m
33 m
10 m
5 m
2 m
Ground
Spring
HO
mean
-8.5
-8.9
-9.7
-9.7
61.0
61.2
61.8
62.7
0.6
0.7
0.6
0.6
997.4
σm
0.5
0.5
0.5
0.5
1.0
1.0
0.9
0.9
0.1
0.1
0.1
0.1
0.4
PG
mean
-4.7
-5.2
-5.8
-5.9
64.8
65.0
65.9
67.3
2.3
2.1
1.8
1.7
996.2
σm
0.2
0.3
0.3
0.3
0.4
0.4
0.4
0.4
0.1
0.1
0.1
0.1
0.5
NG
mean
-8.9
-9.3
-9.9
-9.9
57.2
58.5
59.1
60.1
1.3
1.1
1.0
0.9
992.8
σm
0.2
0.2
0.2
0.2
0.4
0.3
0.3
0.3
0.1
0.1
0.1
0.1
0.2
DNG
mean
-15.8
-16.5
-17.2
-17.2
72.6
73.2
72.9
73.6
0.7
0.9
1.0
0.9
994.8
σm
0.3
0.3
0.3
0.3
0.5
0.3
0.3
0.3
0.1
0.1
0.1
0.1
0.4
Summer
HO
mean
5.7
5.8
5.5
6.0
76.5
75.3
74.7
74.9
4.1
3.8
3.6
3.4
1004.1
σm
0.1
0.1
0.1
0.1
0.2
0.2
0.2
0.2
0.1
0.1
0.1
0.1
0.2
SP
mean
5.8
5.9
5.6
6.1
78.0
76.4
75.6
75.8
2.6
2.6
2.5
2.4
1001.9
σm
0.1
0.1
0.1
0.1
0.2
0.2
0.2
0.2
0.1
0.1
0.1
0.1
0.1
Average concentrations of N14-260, N260-1200,
N>1200, eBC and babs along height over Ny-Ålesund for the
springtime and summertime typologies of vertical profiles.
Season
Profile type
N14-260 (cm-3)
N260-1200 (cm-3)
N>1200 (cm-3)
eBC (ng m-3)
babs (Mm-1)
Spring
HO (column)
mean
80.2
17.5
0.20
35
0.22
σm
16.4
2
0.10
21
0.13
PG (Hs < 0)
mean
205.4
19.9
0.26
24
0.14
σm
12.5
0.2
0.02
3
0.02
PG (Hs > 0)
mean
557.6
22.3
0.16
26
0.16
σm
45.9
1.4
0.02
4
0.02
NG (Hs textless 0)
mean
252.3
23.1
0.53
71
0.43
σm
17.5
0.4
0.05
4
0.02
NG (Hs > 0)
mean
118.9
9.7
0.18
39
0.24
σm
9.3
0.3
0.02
2
0.01
DNG (Hs < 0, ground Aitken plume)
mean
601.3
32.4
0.17
36
0.22
σm
19.9
0.8
0.01
11
0.06
DNG (Hs < 0, above-ground plume)
mean
260.6
32.4
0.17
121
0.74
σm
13.1
0.8
0.01
5
0.03
DNG (Hs > 0)
mean
187.8
22.3
0.08
102
0.62
σm
17.8
0.5
0.01
11
0.07
Summer
HO (column)
mean
435.9
2.1
0.04
–
–
σm
5.8
0.1
0.004
–
–
SP (Hs < 0)
mean
9000
7.0
0.06
319
1.95
σm
250
0.7
0.004
14
0.09
SP (Hs > 0)
mean
648.1
1.4
0.01
24
0.15
σm
27.3
0.1
0.001
1
0.01
Springtime phenomenology
The previous section introduced the vertical behavior of sized aerosol in
terms of frequency distribution of ASh. However, it is also
necessary to describe the intensity of the aerosol concentration changes at
the ASh and the possible dynamics underlying these changes. Thus,
in this section, the springtime vertical aerosol phenomenology will be
investigated. All the profiles measured in spring 2011 were classified based
on their vertical behavior (i.e., shape) and averaged considering the
relative position of each measured data point with respect to the
ASh. The obtained averaged vertical profiles were referred to a
standardized height (Hs; Eq. 12) as described in Sect. 2.2.3. As
the size classes can behave differently with height, Hs=0 was
referred to that observed for the intermediate N260-1200 size class. The
result of the classification and averaging procedure is reported in
Fig. 4a–n (all profile data are reported in Fig. S5). Four main typologies
of vertical profile were found. According to their shape they were named as
follows:
Type 1, homogeneous profiles (hereinafter referred to as HO), Fig. 4a–c;
Type 2, profiles characterized by a positive gradient at Hs=0
(hereinafter referred to as PG), Fig. 4d–f;
Type 3, profiles characterized by a negative gradient at Hs=0
(hereinafter referred to as NG), Fig. 4g–i;
Type 4, profiles characterized by negative gradients located at different
altitude as a function of size (hereinafter referred to as decoupled negative
gradient, DNG), Fig. 4l–n.
Average concentrations of aerosol (N14-260, N260-1200,
N>1200) and eBC below and above Hs=0 for each profile class are
summarized in Table 3.
We first report here the columnar averages of both total aerosol number and
eBC concentrations obtained by averaging all the aforementioned profile
classes: 236.1 ± 23.9 cm-3 (N14-260), 21.1 ± 1.3 cm-3 (N260-1200), 0.2 ± 4 × 10-2 cm-3 (N>1200)
and 52 ± 8 ng m-3 (eBC). They perfectly agree with long-term data
series collected over Ny-Ålesund at the Zeppelin observatory
(Eleftheriadis et al., 2009; Tunved et al., 2013)
during spring (∼ 100–250 cm-3 and 50–70 ng m-3 of
eBC during April). This agreement indicates that all the profile classes
discussed below can be considered characteristic (with their occurring
frequencies and altitudes) for the background Arctic aerosols measured by
Arctic observatories within GAW, AMAP and EMEP observation programs.
Moreover, the eBC data also agreed with results from PAM-ARCMIP (Stone et
al., 2010) which showed a 40–90 ng m-3 range of eBC within the surface
inversion layer and 30–50 ng m-3 above.
Wind rose obtained from the measured wind speed and direction at
the CCT (33 m). Springtime wind rose timely coincident with (a) homogeneous
profiles (HO); (b) positive gradient profiles (PG); (c) negative gradient
profiles (NG); (d) decoupled negative gradient profiles (DNG). Summertime
wind rose timely coincident with (e) homogeneous profiles (HO); (f) profiles
impacted by shipping emissions (SP).
All the CCT wind data were used to compute wind rose graphs timely
coincident with each profile typology (Fig. 5a–d) and will be used in the
following sections. The fjord direction into which the wind is often
channeled is NW–SE. Here, we underline that Fig. 5a–d show the absence of
wind from the north during the profile measurements; thus, any influence from the
Ny-Ålesund village is negligible. In addition, Fig. 6 shows the
ground-based number size distribution measured at Gruvebadet (Sect. 2.1)
for HO, PG, NG and DNG profiles, respectively.
Finally, a brief discussion of the air masses' origin for the four categories
is summarized in the Supplement (see also Fig. S6).
Springtime aerosol number size distribution measured at ground and
timely coincident with homogeneous profiles (HO), positive gradient
profiles (PG), negative gradient profiles (NG) and decoupled negative
gradient profiles (DNG).
Homogeneous (HO) profiles
HO profiles (Type 1; Fig. 4a–c) were observed in 15 % of cases (for
7 days) and were characterized by a homogenous vertical distribution of
aerosol and eBC upon Ny-Ålesund. HO profiles are reported with an
absolute height AGL, because they did not show any ASh to calculate
Hs. They appear in some way analogous to the relatively diffuse
background aerosol reported in the springtime ARCPAC campaign (Brock et al., 2011). During HO profiles, local wind (Fig. 5a) was blowing mainly from
the SW direction, from the glaciers behind Ny-Ålesund (Fig. 1a) and
not along the predominant NW–SE direction. Moreover, HO profiles featured
the lowest wind speed (range 0–2 m s-1, average of 0.6 ± 0.1 m s-1 at 33 m; Table 2).
At the same time, as shown in Fig. 4c, slightly positive θ
profiles characterized, on average, the HO profiles. Profiles of θ with
both positive and negative vertical gradients (as shown in Fig. 2a) contributed to this average.
Negative θ gradients allowed vertical
mixing. Positive θ gradients instead favored stable conditions.
However, they do not contrast the absence of an aerosol
stratification. In fact, just the presence of an important aerosol source
(either local or transported) allow the formation of a distinct aerosol
layer. This process is detailed in Sect. 3.2.3.
The number concentrations in HO profiles were 80.2 ± 16.4 cm-3
(81.9 ± 16.8 % of the total aerosol concentration), 17.5 ± 2.0 cm-3 (17.9 ± 2.0 % of the total aerosol concentration) and
0.2 ± 0.1 cm-3 (0.2 ± 0.1 % of the total aerosol
concentration) for N14-260, N260-1200 and N>1200, respectively.
The aerosol number concentration was thus dominated by the N14-260 size
fraction along the whole profile. eBC and the related babs (Sect. 2.2.2; Eq. 4) reached values of 35 ± 21 ng m-3
and 0.22 ± 0.13 Mm-1.
Aerosol number concentration in HO profiles lies close to the lower values
registered at Zeppelin in April (Tunved et al., 2013) and were characterized
by a ground-based size distribution dominated by the accumulation-mode
particles (Fig. 6). eBC was close to the 25–50th
percentiles reported in Eleftheriadis et al. (2009) for April and to the
lower troposphere value of refractory BC (rBC) measured in spring during the
HIPPO campaign (Schwarz et al., 2013).
Within HO profiles, the aerosol pollution, previously transported from
midlatitudes, affected the first kilometer of the atmosphere. The observed
homogenous mixing conditions allowed us to assume that the aerosol
properties measured at ground level in Ny-Ålesund were representative
of the lower troposphere.
Positive gradient (PG) profiles
PG profiles (Type 2; Fig. 4d–f) occurred in 17 % of cases (for 6 days) and were characterized by an increase of aerosol number concentrations
above Hs=0 and moderate eBC concentrations (24 ± 3 ng m-3;
babs was 0.15 ± 0.02 Mm-1). The average value of ASh
(corresponding to Hs=0) was 417 ± 266 m. During PG profiles, local
wind (Fig. 5b) was blowing mainly from the SE along the predominant NW–SE
direction. This situation is common in Kongsfjorden (Vihma et al., 2011).
PG profiles featured the highest wind speed (range 0–5 m s-1, average
of 2.3 ± 0.1 m s-1 at 33 m; Table 2). At the same time, as shown
in Fig. 4f, a positive θ profile was present with a +1.5 ± 0.4 K increase from Hs=0. A stable atmosphere was present and the
aerosol was brought to the site by long-range transport in this stable
situation. The increment of aerosol number concentrations with altitude was
particularly evident for N14-260 that increased by +171.5 ± 25.4 % (going from 205.4 ± 12.5 cm-3 below the ASh to
557.6 ± 45.9 cm-3 above it) while N260-1200 experienced a
more modest increase of 11.8 ± 7.0 % (going from 19.9 ± 0.2 cm-3 below the ASh to 22.3 ± 1.4 cm-3 above it).
Conversely, the coarse fraction (N>1200) decreased with altitude of
-38.1 ± 10.0 % (going from 0.26 ± 0.02 cm-3 below the
ASh to 0.16 ± 0.02 cm-3 above it). The observed increase of
Aitken and accumulation-mode fractions (N14-260 plus N260-1200),
and the corresponding decrease of the coarse fraction (N>1200) appear
to be in agreement with the observation that during transport events wet
removal processes and dry deposition decrease the coarse particle
concentration by scavenging and, at the same time, establish conditions that
favor secondary aerosol formation due to the lowering of the condensational
sink (Tunved et al., 2013).
The aerosol number concentration values above Hs=0 were close to those
reported in Engvall et al. (2008b) for the Arctic free troposphere during
ASTAR. The ground-based size distribution, not influenced by the pollution
layer at high altitude, was dominated by accumulation-mode particles as in
HO profiles (Fig. 6).
Case study of 23 April 2011 (12:00–13:30 UTC): (a) air mass
back-trajectories; (b) time evolution of N14-260 obtained through the
interpolation of six vertical profiles, each of which lasted ∼ 15 min and was distanced ∼ 1 min from the following profile.
The PG profile data suggested that high-altitude transport events could be
the origin of these types of profiles during springtime. An example of this
process is the interesting case study of 23 April 2011 (12:00–13:30 UTC), when an intense plume of aerosol was transported over Ny-Ålesund.
Figure 7a–b show this event with the associated air mass back trajectories,
and the time evolution of the event obtained through the interpolation of six vertical profiles, each of which lasted ∼ 15 min and was
removed about 1 min from the following profile.
Altitude layers of pollution were documented in Brock et al. (2011) and in
Kupiszewski et al. (2013) and Wofsy et al. (2011). Jacob et al. (2010)
reported that transport from North America and east Asia takes place mainly
at higher altitudes, as also documented in Stohl et al. (2006). In this
respect, the back-trajectories reported in Fig. 7a described high altitude
air from North America and Asia that descended in the Arctic. High aerosol
concentrations at high altitude are important because aerosols can act as CCN
and thus impact on climate via the indirect aerosol effects.
Negative gradient (NG) profiles
NG profiles (Fig. 4g–i) were observed in 48 % of all cases (for 9 days) making the NG the dominant typology of profiles. The average value of
ASh (corresponding to Hs=0) was 506 ± 212 m. The predominant
wind direction was the same as in PG profiles (SE–E direction) with a
component also from SW (Fig. 5c). Figure 4i shows that a strong, positive
θ profile (+4.1 ± 0.3 K increase from Hs=0)
characterized NG profiles.
Within this condition, the case study reported in Fig. 7b showed the
origin of NG profiles. Figure 7b shows first a transport event that
generated PG profiles (Sect. 3.2.2). Afterwards, the transported aerosol
was mixed downward within the PBL until ground. Most important, at the end
of the process (13:30 UTC) a negative concentration gradient with altitude
was established generating a NG profile.
As now shown, NG profiles might have originated from the entrance of Arctic
haze into the PBL after a transport process. In the Arctic, in absence of an
important local aerosol source (i.e., nucleation which acts mainly in summer;
Tunved et al., 2013), only transported aerosol trapped within a thermal
inversion made possible the presence of this typology of profiles. The
presence of an intense θ inversion stabilizes the situation
maintaining a NG typology of profile since vertical mixing is prevented (see
also Fig. 2c). It has to be noticed that an intense θ inversion is
just a necessary condition to promote the formation of NG profiles, not a
sufficient one. This result explains the presence of HO profiles even in a
stable atmosphere (Sect. 3.2.1) and is in agreement with the observation
(reported in Sect. 3.1.1) that the frequency of ground-based θ
inversions (49 %; Fig. 3d) reached values higher than those for any
ASh for any aerosol size (Fig. 3a–c).
NG profiles were characterized by high pollution levels below Hs=0
where an intense decrease of both aerosol and eBC was observed. Crossing the
ASh, aerosol concentrations decreased by -52.9 ± 8.7 % (from
252.3 ± 17.5 to 118.9 ± 9.3 cm-3) for N14-260,
by -57.9 ± 2.6 % (from 23.1 ± 0.4 to 9.7 ± 0.3 cm-3) for N260-1200 and by -66.5 ± 11.5 %
(from 0.53 ± 0.05 to 0.18 ± 0.02 cm-3) for N>1200. eBC behaved
similarly, decreasing by -50.4 ± 6.8 % (from 71 ± 4
to 35 ± 2 ng m-3), with the same phenomenology for babs (from 0.43 ± 0.02 to 0.21 ± 0.01 Mm-1). The last
finding is very important because the altitude of eBC occurrence in the
atmosphere modulates its influence on the climate in the Arctic.
NG profiles exhibited characteristics in agreement with literature data.
Focusing on eBC, the vertical behavior and the observed concentrations
agreed with those found in the PAM-ARCMIP campaign. Stone et al. (2010)
reported rBC concentrations of 40–90 ng m-3 within the surface-based
temperature inversion layer, decreasing to 30–50 ng m-3 above it.
Results reported in Schwarz et al. (2013) for the HIPPO campaign in January
agree with our measurements. Aerosol number concentration and eBC are also
close to the higher values registered at the Zeppelin station in April
(Tunved et al., 2013; Eleftheriadis et al., 2009). In fact, NG profiles
represent the most polluted situation affecting the whole boundary layer,
quite the opposite to HO profiles.
Decoupled negative gradient (DNG) profiles
A particular kind of profiles characterized by a decrease in concentration
with altitude is the DNG typology. Within this class, observed in 20 % of
cases (for 5 days), a lack of symmetry between N14-260 and
N260-1200–N>1200 was observed (Fig. 4l–n). The average value
of ASh (corresponding to Hs=0) was 585 ± 90 m. The main wind
direction was the same as in PG profiles (SE–E direction; Fig. 5d) but
with a lower wind speed (close to that of HO profiles; 0–2 m s-1;
average of 0.7 ± 0.1 m s-1 at 33 m).
Figure 4n shows two strong, positive θ inversions. The first one,
ground based, resulted in +1.3 ± 0.4 K increase from ground; the
second one characterized by +1.1 ± 0.2 K increase from Hs=0.
Within this condition, N14-260 showed a concentration peak close to the
ground of 601.3 ± 19.9 cm-3 that was not present for
N260-1200 and N>1200. N14-260 quickly decreased above the
ground-based peak (-56.7 ± 4.4 %) to a concentration value of
260.6 ± 13.1 cm-3 analogous to that observed in standard NG
profiles (252.3 ± 17.5 cm-3) below the ASh (before reaching
Hs=0). N260-1200 and N>1200 instead remained quite constant
(32.4 ± 0.8 and 0.17 ± 0.01 cm-3, respectively)
from ground until Hs=0 where they decreased by -31.1 ± 2.9 % (to
22.3 ± 0.8 cm-3) and by -54.2 ± 4.7 % (to 0.08 ± 0.01 cm-3), respectively.
Interestingly, eBC concentrations behave contrarily to the N14-260
aerosol fraction. Lowest eBC concentrations were found close to the ground
(36 ± 11 ng m-3; babs was 0.22 ± 0.06 Mm-1) in
correspondence of the N14-260 concentration peak. Above this peak, eBC
concentrations were higher (121 ± 5 ng m-3; babs was
0.74 ± 0.03 Mm-1).
All the aforementioned observations suggest that a particular process could
have influenced ground-level concentrations for this size class only. In
order to shed light on this process, several parameters will be considered below,
namely meteorological parameters, the aerosol chemical
composition and the aerosol number size distribution. Starting with
meteorological parameters, and recalling first Fig. 2d, we see that the
behavior of smallest particles in the proximity of the ground can be
observed concomitantly to the presence of ground-based θ inversions,
a necessary condition (or a concurrent cause) to promote the presence of
ground-based concentration peaks for N14-260. The crucial point to
unravel this phenomenon is to understand the possible origin of these
particles.
Thus, ground-based aerosol and meteorological measurements collected at
Gruvebadet laboratory and at the CCT (Sect. 2.1), and temporally coincident
with the observation of DNG profile, were considered.
Figure 8 show the ground-level PM10 chemical composition determined for
the four categories (HO, PG, NG, DNG) of profiles.
Springtime aerosol chemical composition determined at ground
during (a) homogeneous profiles (HO); (b) positive gradient profiles (PG);
(c) negative gradient profiles (NG); (d) decoupled negative gradient profiles
(DNG). Data shown are the ambient concentrations of each individual aerosol
species.
The nss-nc-SO42- in DNG profiles (1349.9 ± 354.7 ng m-3)
was 3.0 ± 0.7 times higher than that observed in the other profile
classes (389.9 ± 113.2, 410.5 ± 104.3 and
622.1 ± 210.0 ng m-3 for HO, PG and NG). At the same time, the
ss-SO42- was 0.6 ± 0.2 times lower compared to the other
profile classes while the cr-SO42- remained quite constant (ratio
1.1 ± 0.4). The same pattern can be observed considering the
aforementioned sulfate fractions in the PM10 samples (Fig. S7).
These observations, coupled with the lowering of eBC fraction in proximity of
the ground (Fig. 4l), point towards the hypothesis that the ground-based
N14-260 concentration peak was secondary in origin. The
nss-nc-SO42- fraction during DNG profiles appeared in acidic form, as
it was just poorly neutralized by the ammonium. Particularly, the w/w
(weight/weight) nss-nc-SO42- / NH4+ ratio was
1.6 ± 0.4 times higher for DNG profiles (10.3 ± 1.5) than that
observed in the other profile classes: 6.1 ± 1.9, 5.8 ± 2.1 and
7.0 ± 2.2 for HO, PG and NG, respectively. As reported in the
literature (Udisti et al., 2016; Becagli et al., 2012; Udisti et al., 2012)
these values for DNG profiles feature the presence of sulfate in acidic form
(H2SO4). This is in agreement with the finding that “springtime
submicron aerosol in the Arctic surface sites is composed predominantly of
partially neutralized sulfate and sea salt, with lesser contributions from
nitrate, BC, soil and trace elements” as reported in Quinn et al. (2002).
This information is very important when coupled with meteorological data
measured at the CCT (Table 2). Focusing first on the air temperature, it can
be observed that during DNG profiles, the temperature close to the ground
(-17.2 ± 0.3 ∘C) was lower than that observed in the other
profile classes: -9.7 ± 0.5, -5.9 ± 0.3 and -9.9 ± 0.2 for HO, PG and NG, respectively. In addition, the RH
was 18.0 ± 2.0 % higher (73.6 ± 0.3 %) for DNG profiles,
compared to that observed in the other profile classes. Finally, the
wind speed during DNG profiles was half of that during the other profile
classes and was not affected by north direction, avoiding the influence of
Ny-Ålesund (Fig. 5d). All the aforementioned conditions, featured
during DNG profiles, can be resumed in higher acidic sulfate fraction,
lower eBC fraction, lower temperature, higher relative humidity and lower
wind speed during DNG profiles.
As reported in the literature (Kirkby et al., 2011;
Reddington
et al., 2011; Lovejoy, 2004) these conditions decrease
the height of the barrier for new particle formation just considering the
very simplified binary H2SO4–H2O system. Under these
conditions, the secondary aerosol formation can proceed at ambient acid
concentrations in the cooler midtroposphere and at lower altitude in polar
regions. However, it has to be underlined that, as recently reported
(Riccobono
et al., 2014), organics plays a fundamental role for secondary aerosol
formation. They were found in Ny-Ålesund even in spring
(Zangrando et al., 2013). These
figures, coupled with the aforementioned data indicated the N14-260
concentration peaks at ground as locally formed secondary aerosol.
The ground-based aerosol number size distribution (Fig. 6) shows a huge
Aitken mode for DNG profiles, while it is negligible for the other profile
classes. This mode was characterized by a geometric mean diameter Dg of
0.032 ± 0.001 µm and by a geometric standard deviation σg of 1.790 ± 0.006 that were in agreement with 10-year average
values reported in Tunved et al. (2013) at the Zeppelin observatory during
the month of April. The presence of a clearly visible Aitken mode in DNG
profiles supports the aforementioned hypothesis of the presence of a
ground-based plume of locally and newly formed aerosol particles.
In order to estimate (meaning the order of magnitude) the contribution of
locally formed aerosol, the method based on the N / eBC ratio (Sect. 2.2.1),
developed by Rodríguez and Cuevas (2007) was used. The value of S1
(2.4 ± 0.2) was taken as that of background aerosol above the
ground N14-260 plume in DNG profiles. This value was very similar to
that measured during homogeneous profiles (2.5 ± 0.1) when a pure
background aerosol was measured. These S1 values, obtained over
Ny-Ålesund, are close to the lowest values reported in the literature (Reche
et al., 2011; Dall'Osto et al., 2013), a fact resulting from the d50
cutoff size of the miniDiSC (14 nm; Sect. 2.2.1), which is higher than
d50 cutoff sizes (∼ 2–7 nm) usually present in the widely
used condensation particle counters.
Using this reference S1 value, N1 (background number
concentration) and N2 (locally formed secondary aerosol) were computed
as reported in Sect. 2.2.1. Their vertical behavior is reported in Fig. 9. The total amount of secondary aerosol close to the ground is clearly
visible and accounted on average for 63.7 ± 5.6 % (up to 95 % at
ground) of the total N14-260 plume. In fact, within these plumes, the
N14-260 / eBC ratio reached an average value of 22.5 ± 5.4 (maximum
value of 54.8 at ground) clearly indicating the presence of a secondary
formed aerosol (Reche et al., 2011; Dall'Osto et al., 2013). In addition to
this, Fig. 10 shows the temporal behavior of SMPS plus APS data collected at
Gruvebadet during April 2011 together with the percentiles (25,
50, 75 and 90th percentiles of the measured number size
distribution. The presence of nanoparticles (below 100 nm) is clearly evident, even in springtime in the Arctic.
Springtime statistical mean profiles of N14-260 (green line)
apportioned along height for the contribution of background N14-260
aerosol (blue line) and of secondary locally formed N14-260 aerosol
(red line) for the decoupled negative gradient profiles (DNG) category.
(a) Aerosol number size distribution measured at ground during
April 2011; (b) 25, 50, 75 and 90th percentiles of the measured number size distribution.
Interestingly, all the aforementioned results are analogous to data reported
in ARCPAC by Brock et al. (2011). Within the surface inversion layer over
sea ice, they found a region of depleted BC and organic mass concentrations
(lower than in the background case), while sulfate concentrations were
similar or higher. However, under this condition, Brock et al. (2011), did
not feature an increase in particle number concentration. Surprisingly, DNG
profiles are more similar to vertical aerosol profiles discussed by
Kupiszewski et al. (2013) during the summertime ASCOS campaign. In fact,
they found a plume of nanoparticles within the near-surface layer (not related
to the behavior of accumulation-mode particles) during new particle
formation events. They hypothesize that the origin of ultrafine particles
was related to biological processes. This observation becomes important when
considering again the number size distribution reported in Fig. 10. The
75 and 90th percentiles exhibited a summer-like behavior
when compared with Zeppelin data reported in Engvall et al. (2008a) and in
Tunved et al. (2013). These findings point towards the importance of
measuring the frequencies of these episodes, present in the surface layer of
Ny-Ålesund, together with their vertical development (i.e., vertical
mixing), to understand their importance of CCN influencing the Arctic
climate.
Summertime statistical mean profiles of N14-260 (green
line), BC (black line), N260-1200 (blue line), N>1200 (magenta
line), θ (red line) and RH (light blue line) along height over
Ny-Ålesund for the two typologies of vertical profiles: (a)–(c) homogeneous
profiles (HO); (d)–(f) profiles impacted by shipping emissions (SP).
Summer phenomenology
Vertical profiles measured during summers 2011–2012 were also classified
according to their vertical behavior (i.e., shape). They were averaged
considering the relative position of each measured data point with respect
to the ASh. The obtained averaged vertical profiles were referred to
the standardized height Hs. The result of the classification and
averaging procedure is reported in Fig. 11a–f (all profile data are
reported in Fig. S8). In summer, two main
categories were observed:
Type 1, homogeneous profiles (HO), Fig. 11a–c and
Type 2, profiles characterized by the presence of shipping emissions
(hereinafter referred to as SP), Fig. 11d–f.
Average concentrations of aerosol (N14-260, N260-1200, N>1200)
and eBC below and above Hs=0 for each profile class are
summarized in Table 3.
HO profiles
HO profiles (Type 1; Fig. 11a–c) were observed in 37 % of cases due to
the summer higher solar power density at disposal together with the low
albedo as discussed in Sect. 3.2. As already reported for springtime
results, these are the only averaged profiles referred to an absolute height
AGL because they did not show any ASh to calculate Hs. During HO
profiles, local wind (Fig. 5e) was interestingly blowing from the same
direction as in the case of HO springtime profiles: SW direction, from the
glaciers behind Ny-Ålesund (Fig. 1a) and not along the predominant
NW–SE direction of the Kongsfjord. However, summertime HO profiles featured
higher wind speed than in spring (range 0–11 m s-1, average of
4.1 ± 0.1 m s-1 at 33 m; Table 2). At the same time, as shown in Fig. 11c, slightly positive θ profiles characterized, on
average, the HO profiles.
The aerosol number concentrations were found to be 435.9 ± 5.8 cm-3 for N14-260, 2.1 ± 0.1 cm-3 for N260-1200 and
4 × 10-2 ± 4 × 10-3 cm-3 for N>1200, respectively.
Thus, N14-260 accounted for 99.5 ± 1.9 % of the total aerosol
number concentration, which is considerably higher than 81.9 ± 16.8 %
observed in springtime HO profiles. This is in agreement with the
observations reported in the literature that the sunlit summer period is
dominated by small, locally formed Aitken particles (Giardi et al., 2016;
Tunved et al., 2013; Ström et al., 2003, 2009; Udisti et al., 2013;
Viola et al., 2013). eBC and the related babs were negligible, as also
reported by Eleftheriadis et al. (2009). HO profiles were observed in the
absence of ships anchoring in the Ny-Ålesund harbor.
Ship impact along vertical profiles
Summer vertical profiles showed a considerable impact of ship emissions. The
number of ships and the number of passengers (a useful proxy of the ship
dimension) was registered by the Kings-Bay Kull Company and it is reported
in Fig. S9 for summer 2011 and 2012, respectively. Particularly, 57 days
with a total of 103 ship arrivals were registered during JJA of 2011 (62 %
of days; Fig. S9a) while 78 days (85 % of days) with 138 ships (Fig. S9b) were registered during JJA of 2012.
Case study of 6 July 2011 when four ships anchored (not
simultaneously) in the harbor of Ny-Ålesund. Vertical profiles of
N14-260 (green line) and BC (black line) (07:40, 09:01, 09:32
and 13:40 UTC) are reported in panels from (a) to (d) together with ground SMPS
data collected at Gruvebadet (e). Note the change in BC scale to
progressively increasing BC values during the peak of the ship activity at
midday.
Figure 12a–e report the case study of 6 July 2011, when four ships
anchored (not simultaneously) in the harbor of Ny-Ålesund from 07:00
to 19:00 UTC. The largest ship arrived in the morning with approx. 1000 passengers. Figure 12a–d
show four profiles (07:40, 09:01, 09:32
and 13:40 UTC) together with ground SMPS data collected at Gruvebadet (Fig. 12e). The ship impact in the Kongsfjord,
distant by about 1200 m, is clearly evident. N14-260 concentrations reached values up to
2–3 × 104 cm-3 and the eBC concentrations reached the maximum value
of 2000 ng m-3 (at 09:32 UTC).
To highlight the impact of ship emission along the 2 years, average
vertical profiles were calculated. The result is shown in Fig. 11d–f.
During SP, wind blew mainly from the N–NE direction (where the harbor is
located; Fig. 5f). The average wind speed was 2.6 ± 0.1 m s-1.
Once a ship plume arrived, SP profiles were characterized by high pollution
levels below Hs=0. Figure 11f also shows a positive gradient of
θ (+1.4 ± 0.2 K increase from Hs=0). This gradient
constrains the ship plume above ground to an altitude variable with time as
shown in Fig. 12a–d (from 103 to 592 m). Particularly, N14-260
showed a concentration peak close to the ground of 9.0 × 103 ± 2.5 × 102 cm-3, N260-1200 of 7.0 ± 0.7 cm-3 and
N>1200 of 6 × 10-2 ± 4 × 10-3 cm-3. eBC concentrations
behave similarly, reaching concentrations of 319 ± 14 ng m-3. These
concentration values were higher by a factor of 13.9 ± 0.7
(N14-260), 5.1 ± 0.5 (N260-1200), 4.8 ± 0.4
(N>1200) and 13.4 ± 1.0 (eBC) than those observed above
Hs=0 where an intense decrease of both aerosol and eBC is observed.
Crossing the ASh, aerosol concentrations decreased to 648.1 ± 27.3 cm-3 for N14-260, to 1.4 ± 0.1 cm-3 for N260-1200
and to 1 × 10-2 ± 1 × 10-3 cm-3 for N>1200. These
values were close to the background values observed during summer HO
profiles. eBC behave similarly, decreasing to 24 ± 1 ng m-3. As the
ship plume of eBC is located close to the ground, it may exert positive
forcing (Flanner, 2013;
Brock
et al., 2011; Seinfeld and Pandis, 2006; Hansen and Nazarenko, 2004). In
fact, the babs reached values of 1.95 ± 0.09 Mm-1 below
Hs=0.
It is important to note that SP profiles were observed in summer. In summer
(Browse et al., 2012; Quinn et al., 2008;
Stohl et al., 2006) and the locally formed aerosol becomes dominant (Giardi
et al., 2015; Tunved et al., 2013; Ström et al., 2003, 2009). Within
this context, the SP profiles show that the rising shipping emissions in the
Arctic (Corbett et al., 2010; Granier et al., 2006) could affect the
concentrations and the vertical distribution of aerosol, resulting in a
positive forcing, induced by a positive feedback through the local
anthropogenic impact on climate.
Conclusions
Vertical profiles of in situ aerosol number size distribution and black
carbon measurements were conducted by a tethered balloon in the atmosphere
over Ny-Ålesund. The balloon payload was equipped with an optical
particle counter (31 size classes, 0.25 to 32 µm), an electrical
particle detector (d50=14 nm), two micro-aethalometers and
meteorological sensors. Moreover, chemical analysis of filter samples,
aerosol size distribution and a full set of meteorological parameters at
ground were available. A systematic study of vertical profiles of aerosol
number size distribution (14 nm–32 µm) and equivalent black carbon
concentrations was conducted. A total of 200 vertical profiles were measured during
spring and summer along 2 years (2011–2012). Vertical aerosol profiles were
classified for each season according to their shape allowing to obtain a
description of the seasonal phenomenology of vertical aerosol properties in
the Arctic.
Focusing on spring, four main types of profiles were found.
The first one was the homogeneous profiles class (HO), characterized by
constant aerosol and eBC concentration with altitude and representative of
Arctic background conditions.
The second class was that of positive gradient profiles (PG), characterized
by an increase of aerosol concentration with altitude. The importance of
this class is related to the fact that aerosols can act as CCN influencing
the cloud cover and thus the longwave fluxes.
The third class was characterized by negative gradient profiles (NG) with a
decrease of aerosol concentration with altitude and thus high pollution
level close to the ground. This finding is very important because a eBC
layer located immediately above snow and ice may induce a positive forcing.
The fourth class of profiles was characterized by negative gradients located
at different altitude as a function of size (DNG). These profiles were
observed during ground-based events of locally formed secondary aerosol. It
is important, as locally formed aerosol can act as CCN. As low clouds play a
particular role in the sensitive Arctic climate system, the aerosol–cloud
interactions will be one focus of future research activities within the
Ny-Ålesund research community, manifested in the Ny-Ålesund
atmospheric flagship program.
The four categories described above are important when considering the large
amount of ground-based data available for comparisons with modeling
results. Particularly, for HO profiles, ground measurements were fully
representative of the vertical column (up until ∼ 1 km,
vertical limit of experimental activity). During NG and PG profiles, the
ground-based measurements were representative of the air column up to the
planetary boundary layer. Finally, DNG profiles showed that ground-based
measurements differ from those conducted aloft. However, the last case was
influenced by secondary aerosol formation that can be easily detected by an
SMPS (or similar experimental devices). Thus, ground-based measurements
(coupled with a proper PBL determination) are fundamental and very useful
for model comparison. In addition to these, vertical profiles shed light on
the phenomenology and dynamics of the vertical distribution of aerosols in
the Arctic.
During summertime, two main types of profiles were observed. The first class was
characterized by homogeneous background condition profiles, while the second
class reflected the impact of shipping emissions. The ship impact resulted
in a plume of aerosol and eBC pollution constrained close to the ground. In
summer, atmospheric transport from midlatitudes is minor. Increasing
shipping emissions in the Arctic could significantly increase anthropogenic
aerosol and eBC concentrations in the Arctic summer, enhancing the climate
change that this region is already experiencing.