Most of the ion production in the atmosphere is attributed to ionising radiation. In the lower atmosphere, ionising radiation consists mainly of the decay emissions of radon and its progeny, gamma radiation of the terrestrial origin as well as photons and elementary particles of cosmic radiation. These types of radiation produce ion pairs via the ionisation of nitrogen and oxygen as well as trace species in the atmosphere, the rate of which is defined as the ionising capacity. Larger air ions are produced out of the initial charge carriers by processes such as clustering or attachment to pre-existing aerosol particles. This study aimed (1) to identify the key factors responsible for the variability in ionising radiation and in the observed air ion concentrations, (2) to reveal the linkage between them and (3) to provide an in-depth analysis into the effects of ionising radiation on air ion formation, based on measurement data collected during 2003–2006 from a boreal forest site in southern Finland. In general, gamma radiation dominated the ion production in the lower atmosphere. Variations in the ionising capacity came from mixing layer dynamics, soil type and moisture content, meteorological conditions, long-distance transportation, snow cover attenuation and precipitation. Slightly similar diurnal patterns to variations in the ionising capacity were observed in air ion concentrations of the cluster size (0.8–1.7 nm in mobility diameters). However, features observed in the 0.8–1 nm ion concentration were in good connection to variations of the ionising capacity. Further, by carefully constraining perturbing variables, a strong dependency of the cluster ion concentration on the ionising capacity was identified, proving the functionality of ionising radiation in air ion production in the lower atmosphere. This relationship, however, was only clearly observed on new particle formation (NPF) days, possibly indicating that charges after being born underwent different processes on NPF days and non-event days and also that the transformation of newly formed charges to cluster ions occurred in a shorter timescale on NPF days than on non-event days.
Ambient radioactivity in the lower atmosphere supplies ionising energy for
the production of electric charges in the air. It consists of natural and
anthropogenic radioactivity. The anthropogenic fraction comes mainly from
routine and accidental emissions from nuclear power plants and related
facilities as well as nuclear detonations. Minor emissions of natural
radioactivity occur also in connection with various mining activities.
Natural radioactivity is composed of the decay emissions of naturally
occurring radionuclides and cosmic radiation. Alpha and beta particles as
well as the associated gamma and X-ray photons constitute the ionising
radiation from natural radionuclides.
Due to the atmospheric abundance of nitrogen (N
A schematic demonstration of the relationship between primary ions, molecular ions and cluster ions, as well as processes governing their formation and loss.
Air ions were historically concerned in the discipline of atmospheric electricity (Israël, 1970), because their flow in the electric field of the Earth–atmosphere system serves as the measurable conduction current in the atmosphere (Harrison and Carslaw, 2003; Tinsley, 2008; Wilson, 1921). The interest in atmospheric electricity could be traced back to the early 18th century, when thunderstorms were suggested to be electrical phenomena (Herbert, 1997). However, only after Benjamin Franklin proposed the idea to draw electricity down from lightning in 1752 was this theory confirmed, and the study of atmospheric electricity became popular (Herbert, 1997; Tinsley, 2008). Early efforts in this field were substantially invested in understanding lightning and electrification of clouds (e.g. Canton, 1753a, b; Franklin, 1751), even though there were reports on observations of atmospheric electricity under fair weather conditions (Bennett and Harrison, 2007; Canton, 1753b; Read, 1792). Why the air was conductive could not, however, be explained. Meanwhile, Charles-Augustin de Coulomb observed gradual discharge of a well-insulated electroscope around 1785 and he attributed his observation to the contact of suspending particles present in the air (De Angelis, 2014; Walter, 2012). This phenomenon was reproduced by Michael Faraday half a century later in 1835 (De Angelis, 2014). Thanks to the further improvement of the electroscope by William Thomson and Lord Kelvin (De Angelis, 2014; Flagan, 1998), Crookes (1878) found that the discharge rate of an electroscope decreases with a decreasing air pressure, suggesting that it is the air inside the instrument that manipulates the discharge. However, the reasoning remained undisclosed until the discovery of radioactivity by Wilhelm Röntgen, Henri Becquerel and Marie and Pierre Curie enabled Julius Elster and Hans Geitel from Germany and Charles Thomson Rees Wilson from Scotland to relate the spontaneous discharge of the electroscope to ionisation of the air by radioactive sources (Carlson and De Angelis, 2011; De Angelis, 2014; Wilson, 1895, 1899). Therefrom, the importance of air ions in the atmosphere emerged. Contemporaneously, the interest of Joseph John Thomson, director of the Cavendish Laboratory, in the charge carriers produced by ionising radiation motivated the development of instrumentations for measuring electrical charges in gases, leading to various valuable outcomes, e.g. the cloud chamber designed by C. T. R. Wilson, as well as techniques for measuring ion mobility by Ernest Rutherford and John Zeleny and for studying gaseous ion diffusion by John Sealy Townsend (Flagan, 1998; Robotti, 2006). These works laid the theoretical and instrumental foundation for later aerosol studies. The experimental results from C. T. R. Wilson's cloud chamber measurements in 1895 and 1899 on the influence of ionising radiation on the formation of cloud droplets brought interest in air ions to the atmospheric aerosol community. Inspired by these early works, advancements in atmospheric aerosol studies progressed both instrumentally and theoretically over the century (e.g. Aplin and Harrison, 2000; Hewitt, 1957; Hinds, 1999; Hogg, 1939; Mason and McDaniel, 1988; Millikan, 1923; Nolan, 1924; Reischl, 1991; Rosell-Llompart and Fernández de la Mora, 1993; Tammet, 1970, 1995, 2006).
Devices employed for ion studies comprise different types of aspiration condensers, ion mobility spectrometers (IMSs) and mass spectrometers (Cumeras et al., 2015; Hirsikko et al., 2011; Laskin et al., 2012; Tammet, 1970). Notably, modern key instruments for field observations of air ions are mainly aspiration contender-based devices and mass spectrometers, such as the Gerdien counter – an integral aspiration condenser (Aplin and Harrison, 2000; Gerdien, 1905; Vojtek et al., 2006) – ion spectrometers designed by Airel Ltd. – single or multiple channel aspiration condensers (Kulmala et al., 2016; Manninen et al., 2009; Mirme et al., 2007; Tammet, 2006, 2011) – and the atmospheric pressure interface time-of-flight mass spectrometer (APi-ToF) (Junninen et al., 2010). While aspiration condensers provide information on the concentration and mobility of charge carriers, mass spectrometers reveal mainly the chemical properties of them. The IMS, however, has a limited application in studying ambient ions due to difficulties in spectrum interpretation (Hirsikko et al., 2011). The purpose of these instrumentations is not only air ion or air conductivity observations but also nano-material synthesis (Kruis et al., 1998) and the improvement of the fundamental understanding on the relationship between mobility, mass and size (Ku and Fernández de la Mora, 2009).
At present, the number size distribution of air ions can be measured down to
about 0.8 nm in Millikan mobility size by using ion spectrometers
(Manninen et al., 2009; Tammet, 2006, 2011). Most air ions concentrate at
the lowest size band with a diameter of 0.8–1.7 nm (Manninen et
al., 2009), which is generally known as the cluster size range
(Tammet, 1995, 2012). In principle, this size range contains
large molecular ions and clusters of molecular ions. Ions smaller than
0.8 nm comprise mostly relatively simple molecular ions, which are either
primary ions or originate from the survived fraction of primary ions from
recombination. The critical cluster size was found to be 1.5
Although it is known that ionising radiation creates ion pairs via ionisation in the atmosphere (Flagan, 1998; Harrison and Carslaw, 2003; Israël, 1970), except for a few attempts (Hirsikko et al., 2007; Laakso et al., 2004), only minor efforts have been invested in understanding the connection between ionising radiation and observed air ions in the lower atmosphere. Moreover, there is a lack of quantification on the underlying processes. Such deficiencies prompt the motivation of this work to examine how variations in ionising radiation are reflected in observed air ions based on ambient measurement. The aims of this study are (1) to identify the key factors responsible for the variability in ionising radiation and in observed air ion concentrations, (2) to reveal the linkage of observed air ions to the variations in ionising radiation and (3) to provide an in-depth analysis on the effects of ionising radiation on air ion formation. We will first introduce factors that cause the seasonal and diurnal variability in ionising radiation and air ions and then exposit the connection of observed air ions to variations in ionising radiation and the influence of different atmospheric conditions on this relationship. To assist our analysis, we will determine theoretically the potential maximum production rate of ion pairs by ionising radiation, based on our ionising radiation measurements and an assumed average energy expenditure of 34 eV for creating an ion pair, which is termed as the ionising capacity. The ionising capacity can be viewed as a measure of the theoretical maximum ionisation rate, which, however, may not well capture the true ionisation rate due to uncertainties in ionising radiation measurements, possible energy dissipation of ionising radiation in excitation and the invalidity concern associated with the use of 34 eV per production of an ion pair at near ground level in our calculation.
The data presented in this work were collected from a boreal forest site,
which is known as the SMEAR II station located at Hyytiälä, Southern
Finland (61
The ionising radiation measurement system consists of a radon monitor and a
gamma spectrometer. Both devices are maintained by the Finnish
Meteorological Institute (FMI). The air ion data were obtained by a Balanced
Scanning Mobility Analyser (BSMA) (Tammet, 2006), which is part
of the aerosol monitoring system at the station. A differential mobility particle sizer (DMPS) has been responsible for observing the ambient aerosol
number size distribution on this site since 1996. For the study period of
2003–2006, the data availability for air ions, particles, radon and gamma
was 91, 99, 71 and 76 %, respectively, allowing the coverage
of every single day-of-year by each parameter. Data are presented in local
winter time (UTC
The radon monitor is a dual fixed filter-based instrument and measures the
aerosol beta activity. Its inlet is fixed at 6 m above the ground and this
device has been described in detail by Paatero et al. (1994).
Briefly, it is made up of two cylindrical Geiger–Müller counters covered
with glass-fibre filters in lead shielding. A pump controller directs the
airflow to each counter alternatively for a 4 h period, allowing the beta
activity on the other counter to decay. The counting efficiency for beta
particles is determined by the geometric configuration of the counting
system together with the intrinsic detection efficiency of the GM tubes,
which is 0.96 and 4.3 % for
The gamma spectrometer is a scintillation-type detector using a piece of 76 mm
The most state-of-the-art measurement techniques nowadays, usually employing
a mass spectrometer as the detector, such as the APi-ToF (Junninen et al.,
2010, 2016), are able to track air ions down to molecular
sizes and characterise their chemical composition, but they lack the
capability for providing information on the number concentration. Ion
spectrometers are the most deployed type of devices for the study of air ion
concentrations in segregated mobility channels
(Hirsikko et al., 2011). The BSMA (Tammet, 2006) and the
Neutral Cluster and Air Ion Spectrometer (NAIS) (Kulmala et al., 2007;
Manninen et al., 2016; Mirme and Mirme, 2013) are examples in this category.
The highest measurable mobility with these ion spectrometers is
3.2 cm
Ambient air ion data used in this work were measured by a BSMA. The inlet of
this instrument was at a height of 1.5 m above the ground. The BSMA is an
integral-type counter (Tammet, 1970), which offers air ion mobility
spectra by a continuous voltage-scanning system. It is composed of two plain
aspiration condensers, one for each polarity. An electrofilter is installed
at the inlet of each condenser. The BSMA has a total flow rate of 2640 L min
A DMPS gives the information on the aerosol number size distribution (Wiedensohler et al., 2012). At the SMEAR II station, a twin-DMPS system is deployed, with one responsible for the mobility size range of 3–50 nm and the other for larger sizes (Aalto et al., 2001; Kulmala et al., 2012). Each DMPS consists primarily of a differential mobility analyser (DMA) and a condensation particle counter (CPC). The sample air passes through a common bipolar diffusion charger, where aerosol particles are brought to a thermal charge equilibrium. Subsequently, the sample stream is divided into two to be directed to individual DMPSs, where aerosol particles are size segregated in the DMA by changing the voltage step-wisely and then counted in the CPC. The DMPS covered 3–500 nm until December 2004, after which the size range was expanded to 3–1000 nm. The determination of the CS was conducted following the method presented by Kulmala et al. (2012). CS accounts for the loss rate of vapours due to condensational uptake by aerosol particles in the atmosphere (Kulmala et al., 2001).
The snow cover depth was measured manually on a weekly basis on seven
different locations at the SMEAR II station. Measurements on soil
temperature and soil volumetric water content were described by
Pumpanen et al. (2003) and Ilvesniemi et al. (2010). Only the
organic horizon data (5 cm depth, above the mineral layer; Pumpanen et
al., 2003) were used in this work. The organic horizon is in direct contact
with the atmosphere, the condition of which exerts the primary influence on
radon exhalation. The ambient relative humidity and air temperature data
were taken from the mast measurement at 16 and 4.2 m, respectively. More
detailed description of the mast instrumentation can be found from the home
page of the measurement site
(
The ionising capacity (
Decay modes and energy of
The conversion from the activity concentration (Bq m
For conciseness and clarity, hereafter the ionising capacity results from
the alpha and beta decay of
The natural ionising capacity has generally the same dynamical variations as
ionising radiation, from which the ionising capacity was derived. A decline
in the gamma ionising capacity was seen in the seasonal profile prior to the
lowest value (4.5 cm
Seasonal patterns of radon and gamma ionising capacities as a
function of day-of-year over the years 2003–2006. The radon ionising
capacity was determined from the alpha and beta radioactivity associated
with
The diurnal cycle in the ionising capacity originated mainly from the radon
component and followed variations in the radon activity concentration
presented by Chen et al. (2016), which was
attributed to the mixing layer development. However, the contribution by
gamma radiation shifted the relative levels of the ionising capacity from
the seasonal pattern of the radon activity concentration shown by
Chen et al. (2016). A clear diurnal cycle was
observed in both spring and summer, with high ionising capacities found in
the morning and low ones in the afternoon (Fig. 3). The ionising capacity
was generally high in summer and autumn and low in spring and winter, with
the median values being 8.9, 11.2,
11.3 and 8.7 cm
Diurnal cycles of the total ionising capacity presented as medians in different seasons in the years 2003–2006. Spring: March–May; summer: June–August; autumn: September–November; winter: December–February.
The share of radon ionising capacity in the total ionising capacity was in
the range of 10–20 % (Fig. 4), with the lowest fraction obtained in spring
and a progressive increase through the year reaching the highest share in
winter. Interestingly, when separating the data according to the
classification of NPF events defined by
Dal Maso et al. (2005), low radon ionising capacities were
found in association with NPF events. The statistical contribution of radon
in ion pair production was below 10 % on NPF days in all seasons (Fig. 4)
and the median radon ionising capacities on NPF event days were typically
one-third to half of those on non-event days (Supplement Fig. S1 and Table 2). This
observation is likely related to the fact that marine air masses from Arctic
and North Atlantic oceans, which favour NPF (Nilsson et al., 2001),
typically have a low radon content (Chen et al.,
2016). Radon comes from the radioactive decay of radium. Since marine
surface water has a significantly lower radium content than the continental
surface layer (Wilkening and Clements, 1975), only minor amount
of radon can be collected by air masses traversing over the ocean.
Of the airborne
The relative importance of alpha and beta activities from
Median radon ionising capacity in cm
The seasonal and diurnal variations in the ionising capacity originate from features in ionising radiation, i.e. the atmospheric radon concentration and environmental gamma radiation. Presumably, the seasonality of the ionising capacity comes from both of the gamma and radon components, whereas the diurnal feature is primarily related to the dynamical response of radon as a result of the mixing layer evolution.
Energy in eV m
The atmospheric radon concentration, and consequently the derived radon ionising capacity, is affected by mixing layer dynamics, soil type, soil and meteorological conditions, long-distance transportation, etc. The atmospheric radon concentration is generally related to the mixing layer depth, which is also influenced by varying atmospheric conditions, in terms of the air temperature, wind speed, intensity of solar radiation, etc. These connections are further complicated by the arrival of continental air masses at the measurement site, which brings extra radon in addition to the local sources exhaled from the ground. Such aspects have been discussed in our earlier work (Chen et al., 2016).
Radon exhalation from the ground typically depends on the availability of
The median radon ionising capacity as a function of the soil
temperature in 1
In order to examine the impact of SWC on the radon ionising capacity, a soil
temperature window (
The effect of soil conditions on the radon ionising capacity.
In comparison with the radon ionising capacity accounting for alpha and beta
emissions of radon decay, the gamma ionising capacity exhibits a simpler
pattern. Little diurnal variations exist in the total gamma radiation and
therefore in the derived gamma ionising capacity. However, occasionally high
gamma radiation is perceivable during rain events on a temporary basis,
generally of about 2 h. Such observations are typically ascribed to
gamma emissions of the washed-out short-lived progeny of radon (Brunetti
et al., 2000; Dwyer et al., 2012; Paatero and Hatakka, 1999). As for the
seasonal cycle, the low gamma ionising capacity in winter results from the
attenuation effect of snow cover on the terrestrial fraction of gamma
radiation. An exponential reduction was typically seen in the gamma ionising
capacity along with snow cover thickening (Fig. 7). A similar feature has
been reported on the relationship between snow water equivalents and
gamma dose rates for two other Finnish measurement sites in Sodankylä
(67
The attenuation effect of total gamma radiation by snow cover for
the years 2003–2006. Pit 70 and 100 are two measurement points of snow cover
depth. Exponential fittings were made with the goodness of fit denoted as
Cluster ions are produced from primary and more complex molecular ions via their attachment to pre-existing small neutral clusters and the growth by vapour uptake. Molecular ions include both primary ions and those originating from the fraction of primary ions that have survived from the recombination or other sinks, and they are therefore in a close linkage with the ionising capacity. However, due to technical limitations, no reliable measurement can be carried out to acquire the concentration of molecular ions. Ions in the cluster size range (0.8–1.7 nm) are the smallest detectable air ion group based on the current counting technology. Since the formation of these ions are directly related to the dynamics of molecular ions and, to certain extent, to the ionising capacity, the focus in this section is the analysis of variations in the air ion concentration of the cluster size range in association with the ionising capacity.
The cluster ion concentration exhibited some degree of association with the
ionising capacity (Fig. 8). The median cluster ion concentration showed
little diurnal variations during the cold months: from January to March, it
remained at a low level (Fig. 8a), with a median of 513 cm
Median variations in
The enrichment in the median cluster ion concentration in the evening between May and August occurred typically a few hours ahead of the recovery of the median ionising capacity (Fig. 8). Such increases in cluster ion concentrations could result from either an augmentation in the production or a recession in the consumption or sink of these ions. In the former case, a balance between electric charge production and acquisition is re-established towards a higher production of cluster ions. Typically, this can be achieved via either a promotion in the production of electric charges or an enhancement in the charge acquisition. However, since no remarkable increase in the ionising capacity was observed, when cluster ion concentrations started to increase (Fig. 8), the production of cluster ions could be attributed to an enhancement in the charge acquisition. For the formation of cluster ions, the charge acquisition may occur at three stages: (a) prior to clustering via the formation of molecular ions from primary ions, (b) during the actual clustering process from both primary ions and molecular ions and (c) after clustering via the charge uptake from both primary ions and molecular ions. Recombination consumes part of these acquired charges, leaving the rest retained eventually in the form of cluster ions. In the latter case, certain removal processes of ions from the cluster size range are inhibited, which can be either the growth of cluster ions to sizes bigger than 1.7 nm or the loss of cluster ions by the attachment to bigger particles. All of these mechanisms, however, are manipulated primarily by atmospheric conditions. Atmospheric conditions, such as the temperature and relative humidity, can directly influence the rate of clustering and growth. Nonetheless, they also modify atmospheric compositions via the production of functional vapours involved in clustering or growth and via altering the availability of precursor gases of these vapours. As a consequence, these phenomena, seen in Fig. 8, are likely brought by a synergy of complicated atmospheric dynamic processes.
Although high ionising capacities were found in the morning, on average the morning cluster ion concentration was not so high as the evening level during the relatively warm months between April and October (Fig. 8a). Especially in autumn months, the enhancement in air ion production from the high morning ionising capacity was not reflected in the cluster ion concentration. Such observations may result from the cluster formation process becoming inferior to the preferred particle growth due to photochemical processes, while facing the dilution led by the expansion of the mixing volume. Autumn was the second peak period for the occurrence of NPF events at the SMEAR II station after spring (Nieminen et al., 2014). The dissimilar autumn and spring patterns in the cluster ion concentration originate likely from differences in atmospheric conditions and vapour sources. For example, spring UVA radiation intensities are higher than the autumn ones, while the RH shows an opposite behaviour (Lyubovtseva et al., 2005). Biogenic VOC emissions have a strong seasonality (Hakola et al., 2012; Tarvainen et al., 2005), which is reflected in their atmospheric concentrations. Hakola et al. (2012) demonstrated that monoterpenes tend to dominate VOCs in late summer and autumn, while aromatic hydrocarbons dominate in spring and early summer. Sesquiterpene emissions and concentrations were found to be high in late summer and autumn (Hakola et al., 2012; Tarvainen et al., 2005, 2007).
Ion concentrations in different sub-size ranges (0.8–1, 1–1.2 and
1.2–1.7 nm) of cluster sizes showed distinct patterns (Fig. 9). While the
positive polarity dominated the overall cluster ion concentration, more
negative ions were seen in the first two sub-size ranges (0.8–1 and 1–1.2 nm).
The former results from the electrode effect of the negatively charged
Earth surface, which repels negative ions in its vicinity, and it is a
well-known phenomenon in the atmospheric electricity community (Harrison
and Carslaw, 2003; Israël, 1970; Tinsley, 2008; Wilson, 1921). The
latter agrees with observations that generally negative ions possess higher
mean mobility than positive ions (Dhanorkar and Kamra, 1992; Hõrrak,
2001; Israël, 1970); i.e. on average, negative ions are of smaller sizes
than positive ions. Ion concentrations of 0.8–1 nm were found to be the lowest
(around 100 cm
Median seasonality of the cluster ion concentration in 0.8–1, 1–1.2 and 1.2–1.7 nm ranges over the years 2003–2006. Negative ion concentrations are depicted by the solid line and positive ion concentrations by the dash line.
Slight diurnal patterns were found in 0.8–1 and 1–1.2 nm ions. The 0.8–1 nm ion concentration showed features similar to those in the ionising capacity (Figs. 9 and 3), being high in the morning and low in the afternoon. This observation possibly indicates that the dominant population in the size range of 0.8–1 nm are molecular ions, which have not been heavily influenced by dynamical processes of cluster formation and therefore retained the characteristics of primary ions. The seasonal pattern of 1–1.2 nm ions was inconsistent with that of the ionising capacity. Two bumps existed in the concentration of 1–1.2 nm ions in spring and summer; they were found more separated in summer than in spring. However, no identifiable diurnal variation in the 1–1.2 nm ion concentration was seen in autumn; in winter, the 1–1.2 nm ion concentration showed only a small valley in the late afternoon. For 1.2–1.7 nm, which usually is the size range of critical clusters (Kulmala et al., 2012, 2013; Lange et al., 1996; Sipilä et al., 2010), clearer variations could be identified. In all the seasons, a peak close to the sunset, evolving from noon in winter to late evening (21:00) in summer, was observed.
The ionising capacity contains radon and gamma fractions, both of which were
observed to have the capability to promote the production of 0.8–1 nm ions
(Fig. 10). The radon ionising capacity showed a slightly better correlation
with the negative 0.8–1 nm ion concentration than with the positive
polarity. Along with the increase in the radon ionising capacity, more of
the 0.8–1 nm ions were detected, but the degree of dispersion intensified in
the correlation plots (Figs. 10a and S2). This dispersion could come from a
joint effect of the temperature (
Relationship between the hourly ionising capacity and the 0.8–1 nm
ion concentration when the condensation sink (CS) is below 0.001 s
At high radon ionising capacities, the 0.8–1 nm ion concentration,
especially in the positive polarity, dropped to a medium level in winter
months under moist conditions (Fig. 10a–d). This observation might be
related to the proton affinity of water molecules (H
The relationship between the gamma ionising capacity and 0.8–1 nm ion
production was temperature dependent (Fig. 10e and f). When the
The connections between the overall ionising capacity and the 0.8–1 nm ion concentration are reflected in their diurnal behaviour, with minor dissimilarities associated with varying atmospheric conditions and dynamical processes. The median diurnal variation of the 0.8–1 nm ion concentration was very similar to that of the ionising capacity in spring, summer and autumn: both the lowest median 0.8–1 nm ion concentration and the ionising capacity occurred typically at 14:00 when the mixing layer was fully developed (Fig. 11). In comparison with the ionising capacity, however, the 0.8–1 nm ion concentration seemed to follow more instantly the changes in the MLH. The reason behind this observation may be related to the fact that the diurnal variation in the ionising capacity is primarily contributed by radon decay emissions; for example, Chen et al. (2016) showed that the atmospheric radon concentration at near ground level does not respond immediately to mixing layer expansion or shrinkage. In contrast, 0.8–1 nm ion concentrations are related, in addition to the ionising capacity, to photochemical processes and availability of nucleating vapours influenced by solar intensity and atmospheric conditions.
Diurnal patterns in median 0.8–1 nm negative and positive ion concentrations, ionising capacities, global and UVB radiation intensities as well as modelled mixing layer heights (MLH) in different seasons over 2003–2006.
In summer and spring, the median 0.8–1 nm ion concentration built up with the increase in the ionising capacity in the early morning before the solar irradiance started to intensify and the mixing layer to grow. The turbulence introduced by mixing layer development may assist the production of vapours for nucleation and growth from photochemical reactions, possibly initiating the growth of these 0.8–1 nm ions. As can be seen from Figs. 9 and S3, peak concentrations in the 1–1.2 nm size range occurred typically later than those in the 0.8–1 nm size range. Also, tiny bumps in 1.2–1.7 nm ion concentrations could be discerned with some time lag at around 08:00–12:00 in spring and 05:00–07:00 in summer.
After the mixing layer had fully developed in spring, summer and autumn, there could be observed a transient boost in the 0.8–1 nm ion production (especially of the positive polarity) along with the shrinkage of the mixing layer, even prior to a clear recovery of the ionising capacity (Fig. 11). This observation may be attributed to the production of certain vapours that compete with the recombination process and other sink mechanisms for electric charges either via clustering or simple charge binding. Ionising radiation can potentially free a large number of electric charges, which would “sacrifice” themselves mostly in recombination, if not otherwise become detectable air ions. The survived electric charges take part in the formation of 0.8–1 nm ions mainly in the form of primary ions and molecular ions. Certain vapours can cluster among themselves around primary ions to form 0.8–1 nm ions. Charge-binding vapours, however, are able to take over charges from primary ions to form molecular ions. This charge transfer process may occur either via charge exchange ionisation or via chemical reactions between the vapour molecules and primary ions. Some of the molecular ions are possibly born with a size falling in the 0.8–1 nm size range. Ions in the 0.8–1 nm size range may additionally originate from (a) further growth of molecular ions via chemical reactions, (b) clustering of primary ions or molecular ions with nucleating vapours or among themselves or (c) charge uptake by small neutral clusters. Accordingly, the enhanced production of 0.8–1 nm ions after the complete development of the mixing layer, seen in Fig. 11, may be related to changes in the availability of nucleating or charge-binding vapours, altered likely by atmospheric conditions and mixing volume reduction.
Although the ionising capacity continued to increase ever since the recovery in the late afternoon, the enrichment of the 0.8–1 nm ion population ceased typically when little solar radiation was left (Fig. 11). At the same time, bursts in 1–1.2 nm ion concentrations and subsequently in 1.2–1.7 nm ion concentrations were seen (Figs. 9 and S3), which were probably linked to the nocturnal cluster formation events (Ehn et al., 2010; Lehtipalo et al., 2011; Mazon et al., 2016). The emergence of this phenomenon lies at the basis that the production rate of 0.8–1 nm ions by clustering or charge binding is overtaken by the consumption rate of them via either coagulation or condensational growth. The 1.2–1.7 nm ion concentration typically peaked nearly right after the die-out of photochemical reactions when dark hours came (Fig. S3). Concurrently, the 0.8–1 nm ion concentration started to increase again, as the ionising capacity intensified (Fig. 11).
There existed a weak relation between the total ionising capacity and the ion concentration of the whole cluster size (0.8–1.7 nm) range (Fig. 12). On NPF event days, the cluster ion concentration showed a relatively clear increase with the intensification of the ionising capacity (Fig. 12a and b). However, since the cluster ions are very small, they can preserve some properties of gaseous molecules and therefore may sink onto bigger particles. Consequently, corresponding to an ionising capacity value, the cluster ion concentration spanned over a wide range, with high cluster ion concentrations occurring at low CSs. On non-event days, however, the connection between the ionising capacity and the cluster ion concentration became even less identifiable, probably due to the fact that CSs on non-event days are typically higher than those on event days by a factor of 3.6 on average (Dal Maso et al., 2005). In addition to this, meteorological and atmospheric conditions are also more versatile on non-event days than on event days, because NPF events are usually localised in spring, but non-events are spread all over the year (Dal Maso et al., 2005; Nieminen et al., 2014).
The 1 h cluster (0.8–1.7 nm) ion concentration as a function of the total ionising capacity on new particle formation event days (upper panel) and non-event days (lower panel), with the condensation sink (CS) indicated on the colour scale.
By carefully setting constraints to focus on data obtained under relatively
uniform conditions, it was possible to observe a clear relationship between
the total ionising capacity and the whole cluster population (Fig. 13). A
time window between 00:00 and 03:00 was selected to minimise the effect of
diurnal variations. Since the value of CS is typically largest during
night-time (Kulmala et al., 2013), the constraint on the CS was set to be
below 0.002 s
The cluster (0.8–1.7 nm) ion concentration as a function of the
ionising capacity (radon ionising capacity
On both NPF event (Fig. 13a and b) and non-event (Fig. 13c and d) days,
the cluster ion concentration grew with an increase in the ionising
capacity. The dependency of the cluster ion concentration on the ionising
capacity was weaker on non-event days than that on NPF event days. Higher
ionising capacities were seen on non-event days, and correspondingly more
cluster ions were detected on such days than on event days. The cluster ion
concentration tended to level out on event days at high ionising capacities
for both polarities (
In this work, diurnal and seasonal cycles in ionising radiation were
presented and key influencing factors responsible for these features were
overviewed in order to investigate how observed air ions respond to these
variations and to improve our understanding on air ion formation. To assist
the analysis, a term, ionising capacity, was introduced to capture patterns
in ionising radiation. The ionising capacity was determined theoretically as
the potential maximum production rate of ion pairs in the atmosphere by
ionising radiation, based on the assumption that an ion pair is produced
upon every 34 eV energy dissipation of ionising radiation. The data used in
this study were collected from ambient measurements from a
boreal forest site in southern Finland during 2003–2006. In our analysis, the accounted
ionising radiation is composed of energy from alpha and beta decays of
Although ionising radiation is known to be responsible for air ion production, patterns in the measured air ion concentration in the cluster size range (0.8–1.7 nm) did not exhibit a highly comparability to those in the ionising capacity due to modifications of air ion properties exerted by different dynamical processes and chemical reactions during the evolution of charges in the atmosphere. Nevertheless, the connection of air ions to ionising radiation was seen for air ions detected in the lowest detected size band (0.8–1 nm) of the cluster size range (0.8–1.7 nm). The evolvement of these 0.8–1 nm ions with time to larger sizes in the cluster size band was also identified, affirming the primary role of ionising radiation in the production of air ions in the lower atmosphere. However, atmospheric conditions, such as temperature, humidity and pre-existing aerosol particles, brought complications into this relationship. By carefully constraining data to conditions of a similar meteorology, seasonality, diurnality and amount of background aerosol particles, a strong dependency of total cluster ion concentrations on the ionising capacity was identified on NPF days. However, the linkage was not visible on non-event days. These observations may suggest that charges, after being born, underwent different processes on NPF days and non-event days and possibly indicate also that the transformation of newly formed charges to cluster ions occurred faster on NPF days than on non-event days. These results could help to advance our understanding on the role of ions in atmospheric new particle formation.
However, to obtain further insights into the fate of charges created by ionising radiation in the atmosphere, i.e. ion balance, and into the role of air ions in the atmospheric new particle formation process, it is crucial to understand the transformation process of electric charges into detectable air ions. For this purpose, knowledge on the number size distribution of air ions smaller than 0.8 nm is of necessity. Additionally, theoretical understanding on the formation mechanisms of cluster ions from molecular ions needs to be deepened. Conjointly, also advancing instrumental development for the detection of sub-0.8 nm ions could be worth being brought onto the agenda.
The aerosol particle and
meteorological data used in this work are publicly accessible on the
SmartSmear website
(
This work received funding support from the Academy of Finland Centre of Excellence (project no. 272041 and 1118615), European Union's Horizon 2020 research and innovation programme under grant agreement no. 654109 (ACTRIS-2) as well as the European Union Seventh Framework Programme (FP7/2007-2013 ACTRIS) under grant agreement no. 262254. Also the CRyosphere-Atmosphere Interactions in a Changing arctic Climate (CRAICC) project of the Nordic Centre of Excellence is acknowledged. The authors appreciate the valuable communication with Jaana Bäck, Pasi Kolari, Anne Hirsikko and Juha Hatakka. Edited by: D. Spracklen Reviewed by: two anonymous referees