Vertical profiles of trace gas and aerosol properties over the Eastern North Atlantic: Variations with season and synoptic condition

Because of their extensive coverage, marine low clouds greatly impact the global climate. Presently, the response of marine low clouds to the changes in atmospheric aerosols remains a major source of uncertainty in climate simulations. One key contribution to this large uncertainty derives from the poor understanding of the properties and processes of marine aerosols under natural conditions, and the perturbation by anthropogenic emissions. The Eastern North Atlantic (ENA) is a region of 25 persistent but diverse subtropical marine boundary layer (MBL) clouds, where cloud albedo and precipitation are highly susceptible to perturbations in aerosol properties. Here we examine the key processes that drive representing an indirect major at The impact of synoptic condition on the aerosol properties is examined. Under pre-front and post-front 40 conditions, shallow convective activity often leads to a deep and decoupled boundary layer. Coalescence scavenging and evaporation of below clouds leads to much reduced N CCN and larger accumulation-mode particle sizes in the cloud-containing decoupled indicating that surface measurements overestimate the N CCN relevant to the formation of clouds under decoupled

because the average FT NCCN is the same or even lower than that in the MBL. The particle number flux due to FT entrainment is dominated by pre-CCN (particles that are too small to form cloud droplets under typical conditions, i.e., particles with sizes below the Hoppel minimum) due to the elevated Npre-CCN in the lower FT. Once these pre-CCN are entrained into the MBL, they can grow and reach CCN size range through condensational growth, representing an indirect and major source of MBL CCN at ENA. The impact of synoptic condition on the aerosol properties is examined. Under pre-front and post-front 40 conditions, shallow convective activity often leads to a deep and decoupled boundary layer. Coalescence scavenging and evaporation of drizzle below clouds leads to much reduced NCCN and larger accumulation-mode particle sizes in the upper, cloud-containing decoupled layer, indicating that surface measurements overestimate the NCCN relevant to the formation of MBL clouds under decoupled conditions.

Introduction 45
Remote marine low-cloud systems have a large spatial coverage, and are particularly susceptible to aerosol perturbations because of their relatively low optical thickness and low background cloud condensation nuclei (CCN) concentrations. The response of low cloud systems to changes in aerosols is a major source of uncertainty in the simulations of climate change (Bony and Dufresne, 2005;Wyant et al., 2006;Turner et al., 2007;Carslaw et al., 2013). One major contribution to this large uncertainty derives from the poor understanding of the properties and processes of marine aerosols under natural conditions, 50 and the perturbation by anthropogenic emissions. The processes that control CCN population in the marine boundary layer (MBL) have been examined by a number of studies. These processes include entrainment of free troposphere aerosols (Raes, 1995;Clarke et al., 2013), new particle formation (NPF) (Bates et al., 1998;O'Dowd et al., 2010), production of sea spray aerosols (O'Dowd et al., 2004;Russell et al., 2010;Prather et al., 2013;Quinn et al., 2017), condensational growth of Aitkenmode particles (Sanchez et al., 2018;Zheng et al., 2018;Zheng et al., 2020a), interstitial particle scavenging by cloud droplets 55 (Pierce et al., 2015), and the removal of CCN by coalescence scavenging Wood et al., 2017). In addition, synoptic conditions also strongly influence entrainment and coalescence scavenging, and therefore, the population of MBL aerosols Wood et al., 2015;Wood et al., 2017).
The Eastern North Atlantic (ENA) is a region of persistent but diverse subtropical MBL clouds Rémillard and Tselioudis, 2015). The origins of the aerosols arriving at the ENA are diverse, varying from clean marine air masses to 60 those that are strongly influenced by continental emissions (O'Dowd and Smith, 1993;Wood et al., 2015;China et al., 2017;Zawadowicz et al., 2020). Zheng et al. (2018) examined MBL aerosol in the ENA using three years of measurements (2015 -2017) at the US Department of Energy Atmospheric Radiation Measurement (ARM) Facility site on Graciosa Island in the Azores, Portugal. In the ENA, MBL aerosol concentrations in different size ranges exhibit strong seasonal variations. For example, larger accumulation-mode aerosols (Dp > 300 nm) are dominated by sea spray aerosols with a higher concentration 65 in winter than in summer, largely due to the seasonal variation in wind speed. The growth of nucleation-and Aitken-mode aerosols in the MBL represents a substantial source of accumulation-mode aerosols with the highest contribution potentially reaching 60% during summer (Zheng et al., 2018). Using the long-term measurement data from the size-resolved CCN measurement system, Zheng et al. (2020a) further found that over the ENA, organics represent an important or even dominant condensing species for the growth of pre-CCN particles. Potential precursors of the secondary organics are volatile organic 70 compounds from the ocean biological activities and those produced by the air-sea interfacial oxidation. The properties of aerosols from the ENA were also studied using long-term observations at the Mace Head Atmospheric Research Station located on the west coast of Ireland (O'Dowd et al., 2004;Ovadnevaite et al., 2014;Yoon et al., 2007). These studies show that a major driver of the seasonal variation of midlatitude marine aerosol is the biological activity, because the majority of the aerosol mass, both inorganic sea salt and organic matter, was linked to bubble-mediated aerosol production. The organic fraction was 75 also linked to organic matter concentrated at the ocean surface resulting from plankton bloom (Behrenfeld et al., 2019;Ovadnevaite et al., 2014).
While previous studies have greatly advanced our understanding of MBL aerosol properties and processes, they are mostly based on measurements at ground/sea level, whereas the vertical profiles of aerosol properties are needed to understand some of the key processes that drive CCN populations in the MBL, including long-range transport of continental emissions, 80 entrainment of free tropospheric (FT) air, and the interactions between aerosols and clouds. Airborne measurements were carried out during several field campaigns in the 1990s, including the North Atlantic Regional Experiment (NARE) (Parrish et al., 1998), the Atlantic Stratocumulus Transition Experiment (ASTEX) (Albrecht et al., 1995), the Marine Aerosol and Gas Exchange (MAGE) campaign (Huebert et al., 1996) that organized the chemical experiment within ASTEX, and the Aerosol Characterization Experiment (ACE-2) (Raes et al., 2000). The emphasis of NARE was mostly on ozone chemistry and ASTEX 85 was focused on the transition of marine stratocumulus clouds. During ACE-2, the variation of aerosol size distribution and chemical composition was examined during three Lagrangian experiments over periods of ~30 hours as air masses advected over the North Atlantic . These experiments show that the production of sea-salt particles at elevated wind speeds leads to an increase in accumulation-mode particle concentration (Hoell et al., 2000). The reduction of the Aitkenmode particle concentration is attributed to collision with cloud droplets and accumulation-mode particles (Johnson et al., 90 2000). Dilution by entrainment, in particular when the depth of the MBL increases as the air mass moves over a warmer ocean, is the main reason for a general reduction in the aerosol concentration.
In this study, we present comprehensive airborne measurements of aerosols and trace gases in both summer and winter seasons during the Aerosol and Cloud Experiments in the Eastern North Atlantic (ACE-ENA) campaign (Wang et al., 2021). The large number of the flights provides statistically robust characterization of the vertical profiles of aerosol properties and allows for 95 understanding of the aerosol properties under natural conditions (i.e., aerosols mostly produced by natural processes) and those strongly influenced by anthropogenic emissions. Key processes that control the population of CCN in the MBL are investigated by examining the vertical profiles of aerosol properties and their variations between the seasons. The impact of synoptic conditions on MBL structure and the vertical profiles of the aerosol populations are examined, and its implication on studying aerosol cloud interactions using ground-based aerosol measurements is discussed. 100

Measurement overview
During the ACE-ENA campaign, the ARM Aerial Facility Gulfstream-1 (G-1) research aircraft was deployed in the Azores, Portugal as part of two intense operation periods (IOPs) during summer 2017 (June to July, summer IOP) and winter 2018 (January to February, winter IOP). The G-1 aircraft was stationed at the Lajes airport on Terceira Island, and a total of 39 flights (20 in summer and 19 in winter) were conducted. The dates and durations of the flights are summarized in Fig. 1. The 105 deployments during both seasons allow for the examination of key aerosol and cloud processes under a variety of representative meteorological and cloud conditions. Each flight consisted of four to six vertical profiles (excluding those leaving and arriving at the Lajes airport), providing the aerosol and trace gas properties as a function of altitude. The flights also included horizontal legs near the surface of the ocean (~ 100 m AGL), just below cloud base, within the cloud, at cloud top, and above clouds in the lower FT. To maximize the synergy between the G-1 and ground measurements at the ARM ENA observatory on Graciosa Measurements onboard the G-1 included meteorological parameters, trace gas species, aerosol, and cloud properties (Wang et 115 al., 2021). Measurements used in this study are summarized in Table 1 and described briefly below. The height of the MBL is derived from the measured vertical profile of potential temperature, from which the boundary between the MBL and FT is often clearly defined by an abrupt increase of the potential temperature with altitude (i.e., temperature inversion). When the inversion is not obvious, liquid water content (LWC) and water vapor mixing ratio (w) profiles are used to help identify the cloud top and define the MBL height (rapid increase of LWC or decrease in w). Water vapor mixing ratio is calculated from 120 the ambient temperature and dew point measured onboard the G-1. The mixing ratios of CO and O3 were measured by trace gas monitors (Los Gatos Research, Inc., N2O/COR-23r and Thermo Scientific model 49i, respectively). The aerosols particles were sampled using an isokinetic inlet that has a 50% upper cut-off size of 5 μm (Schmid et al., 2014) and were subsequently dried with a Nafion tube. Aerosol size distributions from 10 nm to ~600 nm were measured by a fast-integrated mobility spectrometer (FIMS) . Total number concentration of particles larger than 10 nm in diameter (N>10) was 125 measured by a condensation particle counter (CPC 3772, TSI Inc.). Another CPC (CPC 3010, TSI Inc., modified to achieve an lower cut-off size of 10 nm) was operated downstream of a thermal denuder operated at 300 o C (Fierz et al., 2007), which allows volatility-based separation by exploiting the higher volatility of organics and sulfates versus sea salt and refractory black carbon (BC) (Clarke et al., 2013). The volatile particle number fraction (fvol) is derived as 1 -N>10,TD/N>10, where N>10,TD represents the number concentration of thermally denuded aerosols measured by the modified CPC 3010. 130 BC mass concentration was characterized by a single-particle soot photometer (SP2, DMT Inc., Longmont, CO). A highresolution time-of-flight aerosol mass spectrometer (HR-ToF-AMS) (Jayne et al., 2000;DeCarlo et al., 2006;Zawadowicz et al., 2020) and a particle-into-liquid sampler (PILS) coupled with off-line ion chromatography (Orsini et al., 2003;Sullivan et al., 2019) were deployed to characterize sub-micrometer non-refractory aerosol composition. Both the HR-ToF-AMS and the PILS were deployed downstream of an impactor with a cut-off size of 1 μm. Cloud LWC is calculated by integrating the 135 droplet size distribution measured by a fast cloud droplet probe (FCDP, SPEC Inc., Boulder CO) and validated by a multielement water content system (WCM-2000, SEI Inc., Tolland, CT). To minimize artifacts due to droplet shattering on the aerosol sampling inlet, we exclude aerosol measurements inside clouds (i.e., LWC > 10 -3 g m -3 ) from our analysis. In addition, only measurements collected at least 5 km away from Graciosa Island and Terceira Island are included in the analysis to minimize the potential impact of local island emissions. 140

Air mass origins
The vertical profiles of potential temperature and LWC show that the MBL heights during the summer and winter IOPs are 1220 ± 450 and 1640 ± 480 m (mean ± standard deviation), respectively, indicating a strong seasonal variation. The shallower MBL during summer is due to the presence of a stronger Azores high pressure system that favors MBL low cloud occurrence 145 through enhanced synoptic-scale subsidence, lower tropospheric stability, and increased low-level relative humidity. To understand the air mass origin and its impact on the vertical profiles of trace gases and aerosols, we calculated the 10-day back-trajectories of air masses arriving at three altitudes (500, 1500, and 3000 m) above the ENA site during the G-1 flight days using the Hybrid Single-Particle Lagrangian Integrated Trajectory (HYSPLIT) version 4 model (Stein et al., 2015). The back trajectories were simulated with a time step of 6 h using the National Center for Environmental Prediction (NCEP) Global 150 Data Assimilation System (GDAS) meteorological data as input. A cluster analysis of the trajectories was then performed (Abdalmogith and Harrison, 2005) for each IOP and arriving altitude (i.e., 500, 1500, or 3000 m) (Fig. 3). Based on these solutions, the air masses are classified into three main clusters of different originating locations: (1) North America, (2) the recirculating flow around the Azores high, and (3) the Arctic. This classification is consistent with previous studies of the air mass origins over the ENA (O'Dowd and Smith, 1993;Zheng et al., 2018). Among the three clusters, the Arctic cluster has the 155 lowest frequency of occurrence. A large fraction of the air masses arriving at 1500 m over the ENA site (i.e., lower FT) originated from the North America, suggesting a strong influence of continental emissions. Nearly half of the air masses arriving at the ENA MBL during the summer had circulated around the Azores high over the ocean for more than ten days.
Air mass trajectories during the summertime show relatively weaker vertical motion, whereas most of the wintertime air masses arriving at all three altitudes over the ENA site originated from above 3000 m ten days before, suggesting that the wintertime 160 air masses arriving at the ENA are less influenced by anthropogenic emissions near the surface.

Vertical profiles of trace gas mixing ratios
Mixing ratios of water vapor, CO and O3 measured onboard the G-1 aircraft are grouped into 200-m altitude bins, and their statistics are shown as a function of altitude in Fig. 4 for both the summer (red) and winter (blue) IOPs. The vertical profiles indicate a strong seasonal variation. Due to the lower ambient temperature and thus the lower saturation vapor pressure, the 165 average water vapor concentration during winter is significantly lower than that during summer (by 2 -5 g m -3 ). The CO mixing ratio during winter (>100 ppbv) is higher than that during the summertime (<85 ppbv) at all altitudes sampled, a result of the seasonal variation of hydroxyl radical (OH) concentration. The major sink of atmospheric CO is oxidation by OH, which accounts for about 90% of the loss, the remainder being mainly due to dry deposition (Thompson, 1992). Because there is no substantial emission of CO over the open ocean in the ENA, the vertical profile of CO mixing ratio is largely controlled by the 170 long-range transport of CO (i.e., source) and its oxidation by OH (i.e., sink). Due to the high actinic flux, the concentration of OH is higher during the summer months, leading to a lower CO mixing ratio (Novelli et al., 1998). The vertical trends of CO also differ between summer and winter. The average CO mixing ratio during wintertime decreases with increasing altitude, from 110 ppbv near the ocean surface to 100 ppbv at an altitude of 3000 m. The temperature of the land surface is substantially colder than that of the ocean during winter, therefore continental outflow tends to stabilize the lower atmosphere, limiting 175 vertical mixing to the lowest portion of the atmosphere. Because the major CO sources over continents are near the surface, the ENA CO mixing ratio during winter decreases with increasing altitude as a result of the weaker impact of continental emissions aloft (Holloway et al., 2000). In addition, higher OH concentration and thus a higher CO sink in the FT also contributes to the vertical trend (Spivakovsky et al., 1990). The CO mixing ratio exhibits an opposite vertical trend during summer, increasing from 76 ppbv near the ocean surface to 85 ppbv at an altitude of 3000 m. This reverse of the vertical trend 180 is due to stronger influences of biomass burning and pollution from North America (Honrath et al., 2004). In North America, biomass burning is more frequent and its emissions are stronger during summertime. A large fraction of the biomass burning emissions are lofted into the FT (Zheng et al., 2020b). In addition, due to the warmer land surface temperature compared to that of the ocean, the lower atmosphere is destabilized and continental outflow is often lifted above the MBL as it is transported over the Atlantic Ocean. 185 Both biomass burning and anthropogenic pollution generate the precursors of O3, including nitrogen oxides (NOx), CO, and volatile organic compounds (VOCs) (Thompson et al., 2001;Jaffe et al., 2004). As a result, during the summer IOP, the G-1 detected increased O3 concentrations as a function of altitude, from 23 ppbv at the ocean surface to 37 ppbv at 3000 m above the ocean. The wider range of O3 concentration during summer (23 to 37 ppbv) compared to that during winter (35 to 37 ppbv) 190 suggests episodic influence of long-range transported plumes in summer months. There is also a relatively strong correlation between O3 and CO in the FT during summer (R 2 = 0.63), due to the long-range transport of continental emissions (Fig. S1).
The correlation coefficient is substantially lower during winter (R 2 = 0.23), revealing that the transport of CO and O3 may be decoupled during winter due to a weaker impact of pollution events. This observation generally agrees with the data collected during the NARE campaign over the North Atlantic region, where anthropogenic pollution leads to a positive correlation 195 between O3 and CO (Parrish et al., 1998). The presence of O3 of stratospheric origin in the lower FT is unlikely, as the transport of O3 from the stratosphere is generally associated with a rapid and significant increase of O3 under special meteorological conditions (Jaeglé et al., 2017), which were not encountered during ACE-ENA. As the signature of such stratospheric air is elevated O3 with low CO levels, the transport of stratosphere air leads to a negative O3 and CO correlation (Parrish et al., 1998), which was not observed in the lower FT over the ENA either (Fig. S1b). 200

Vertical profiles of aerosol properties
The average dry aerosol size distributions within the MBL are bimodal with a clear Hoppel minimum for both seasons (Fig.   5a). The Hoppel minimum represents an average particle size at which particles become CCN. Although the Hoppel minimum shows some variation from day to day, its value is relatively constant during the same flight day. To facilitate the discussion 205 of aerosol processes that influence the MBL CCN population, we define pre-CCN as particles with diameters smaller than the Hoppel minimum (i.e., particles that are too small to form cloud droplets under typical conditions in the MBL). CCN are defined as the particles with diameters larger than the Hoppel minimum. Therefore, both nucleation-and Aitken-mode particles belong to the pre-CCN. During ACE-ENA, since the Aitken-mode particles often dominated the pre-CCN population, the concentrations of pre-CCN (Npre-CCN) and CCN (NCCN) are close to the concentrations of Aitken and accumulation modes, 210 respectively, and they are used interchangeably in this study. In this study, Npre-CCN and NCCN are derived by integrating the aerosol size distributions below and above the Hoppel minimum that was determined for each flight. Compared to the CCN concentrations measured by a CCN counter at fixed supersaturations, the derived NCCN based on the Hoppel minimum takes into consideration the variation of the supersaturation relevant to MBL cloud formation, and therefore, is a more realistic representation of the CCN concentration. 215 Figure 6 shows the vertical profiles of particle number concentrations normalized to standard temperature and pressure (STP, 273.15 K and 101.325 kPa, Fig. 6a-6c), mean particle diameter of Aitken and accumulation modes ( Fig. 6e-6f), and the number fraction of volatile particles (Fig. 6d). During both the summer and winter IOPs, the total aerosol number concentration (N>10) in the FT is higher than that in the MBL (Fig. 6a), largely due to the elevated Npre-CCN (Fig. 6b) in the FT. In contrast, NCCN in 220 the FT (e.g., at altitude of 2000-3000 m) is about the same as or even slightly lower than that in the MBL during both summer and winter seasons (Fig. 6c). It is worth noting that there are some vertical variations of NCCN within the MBL, because the MBL in the ENA is often decoupled (see Sect. 3.5.2). Particle number concentrations exhibit strong seasonal variations, and on average, NCCN, Npre-CCN, and N>10 are higher in summer than winter at all altitudes sampled. The average number fraction of volatile particles (fvol) increases from ~ 50% near the ocean surface to ~60% and ~75% near the top of the MBL (i.e., ~ 1500 225 m) during summer and winter, respectively. The value of fvol remains relatively constant in the lower FT (Fig. 6d). Free tropospheric fvol is higher in the winter, consistent with less influence of long-range range transported aerosols discussed below.
The mean Aitken-mode particle size (Dp,Ait) during summer is approximately 20 nm larger than that during winter in the MBL (Fig. 6e), The vertical profiles of accumulation mode diameters (Dp,Acc) are similar for the two seasons, expect at altitudes between 500 and 1500 m (i.e., the upper MBL), where mean Dp,Acc is larger during winter (200 nm) than summer (170 nm). 230 Figure 7 shows the vertical profiles of the mass concentrations of non-refractory species measured by HR-ToF-AMS and BC mass concentration measured by SP2 for both seasons. Sulfate, organics, and ammonium constitute almost 99% of the nonrefractory sub-micrometer aerosol mass, whereas nitrate concentration is negligible. The sulfate concentration maximizes near the ocean surface, reaching approximately 0.5 μg m -3 during the summer. The decrease of sulfate with increasing altitude 235 indicates a surface source. The organics, ammonium and nitrate mass concentrations show elevated values near the surface (i.e., below ~ 1000 m) and in the lower FT (i.e., between 1500 and 2500 m). BC mass concentration increases with altitude and peaks around 2200 meter in the lower FT, indicating that the major source of BC is long-range transport in the FT. All species show higher concentrations during the summer season.

Differences between MBL and FT aerosols 240
The vertical gradients of Npre-CCN and NCCN shown in Fig. 6 suggest that the entrainment of FT air may impact the aerosol properties in the MBL. During summer, at an altitude of around 2000 m, the mean NCCN and mBC are substantially higher than the corresponding median values ( Fig. 6c and Fig. 7e), indicating occasional abnormally high NCCN and mBC in the lower FT.
Given the much lower mBC in the MBL, the abnormally high NCCN and mBC are attributed to long-range transport of continental emissions. In addition, nitrate, organics, and ammonium ( Fig. 7) also show a local maximum in the similar altitude range (i.e., 245 1500 to 2500 m), consistent with the influence of continental emissions. Both urban pollution and biomass burning in North America likely contribute to the long-range transported aerosol layers, which will be discussed further in Sect. 3.4. The influence of the plumes is also evident from the comparison of the vertical profiles with those under background conditions.
In this study, we define the background conditions (i.e., with minimum influence from continental emissions) as those when mBC is less than 5 ng m -3 . The vertical profiles of the aerosol properties under the background conditions are shown in Fig. S2  250 and S3. Once the influence from continental plumes is excluded, the average NCCN in the FT during the summer show a significant decrease. As a result, NCCN under the background condition (NCCN,bg) in the FT becomes substantially lower than that in the MBL. For the aerosol chemical composition, both sulfate and organic mass concentrations at altitudes between 1500 and 2500 m reduce substantially when the measurements are limited to those with mBC less than 5 ng m -3 ( Fig. S3a and S3b).
This indicates that sulfate, organics, and BC coexist in the long-range transported aerosol layers. For the winter IOP, NCCN 255 only exhibits a minor peak at the altitude of around 2200 m and there is little difference between NCCN and NCCN,bg in the lower FT, suggesting relatively weak influence from the long-range transported plumes over the ENA. This is consistent with the 10-day back-trajectories showing that, during winter, air masses arriving at the lower FT in the ENA mostly descended from higher altitudes (Fig. 3). As MBL aerosol in the ENA is continually being modified by air entrained from the FT, these vertical profiles indicate that except for occasional periods during summer when long-range transported plumes are present, the 260 entrainment of FT air does not serve as a direct source of CCN in the MBL. Instead, the entrainment of FT air dilutes and acts to reduce MBL CCN concentrations. We note that aerosols under the background conditions during ACE-ENA are likely influenced by diluted and aged continental plumes. Therefore, NCCN under natural conditions (i.e., during the pre-industrial era) is expected to be even lower, possibly leading to a more pronounced difference in NCCN between the lower FT and the

MBL. 265
In contrast, during both seasons, Npre-CCN in the FT is substantially higher than that in the MBL, leading to an increasing N>10 with altitude from the MBL to the lower FT. Therefore, entrainment of FT air increases MBL pre-CCN concentrations and total particle number concentrations in the MBL. Compared to NCCN, the vertical profiles of Npre-CCN,bg are very similar to those of Npre-CCN, implying that, on average, continental emissions have relatively weaker impact on Npre-CCN. This is consistent with 270 the picture that aged aerosols in long-range transported continental plumes are dominated by accumulation-mode particles (Zheng et al., 2020b) and a substantial fraction of Aitken-mode particles in the FT is produced by NPF in the outflow of convective and frontal clouds followed by the coagulation and condensational growth (Clarke et al., 1998;Andreae et al., 2018;Williamson et al., 2019;McCoy et al., 2020) .

Growth of pre-CCN into CCN size range 275
Once pre-CCN are entrained into the MBL, they can grow and reach CCN-active sizes through condensation (Yoon et al., 2007;Sanchez et al., 2018;Zheng et al., 2018;Zheng et al., 2020a). Therefore, the entrainment of FT Aitken-mode aerosol represents an indirect source of MBL CCN in the ENA. It has long been recognized that sulfates produced from dimethyl sulfide (DMS) oxidation are major species for the condensational growth of pre-CCN in remote marine environments.
Another mechanism for the formation of CCN within the MBL is the activation of Aitken-mode particles in a stronger than average updraft, which causes a higher peak supersaturation (Kaufman and Tanré, 1994). These Aitken-mode particles would otherwise remain in the interstitial air of clouds. Once activated, sulfate and organics can be produced through aqueous 290 chemistry inside droplets. Unless these droplets are removed by precipitation, they become CCN upon evaporation outside of the clouds, and readily participate in subsequent cloud formation. The effect of this mechanism on the MBL CCN budget is difficult to evaluate with measurements only and will be a subject of future studies. The vertical profile of sulfate mass concentration indicates a surface source, consistent with the picture that over the open ocean, most submicron sulfate is derived from DMS through both gas phase and in-cloud oxidation (Hegg and Hobbs, 1981;Gurciullo et al., 1999;Ovadnevaite et al., 295 2014;McCoy et al., 2015). The higher MBL sulfate mass concentration during the summer season is a result of stronger DMS emission (Zawadowicz et al., 2020) and higher oxidant (e.g., OH) concentrations. During summer, nearly half of the air masses arriving in the ENA MBL had been circulating around the Azores high over open ocean for more than ten days, indicating that Aitken-mode aerosols have extended time to grow by condensation and in-cloud processes.

Contribution from marine primary aerosols 300
For both seasons, fvol is largely altitude-independent in the lower FT, and it deceases nearly linearly from the top of the MBL to the ocean surface (Fig. 6d). Due to the lower existing condensation sink and higher radiation intensity, NPF often occurs in the FT (Clarke et al., 1998). Sulfuric acid is recognized to be the major component of these freshly formed particles, while ammonia, amines, and biogenic VOC may also participate in the particle formation process (Dunne et al., 2016). The newly formed particles can subsequently grow to Aitken-mode size through coagulation and condensation. These particles can 305 represent a large fraction of the FT particle number and are volatile at 300 o C, leading to a relatively higher fvol in the FT. The volatile fraction in the FT is lower during summer (Fig. 6d), due to the stronger influence of long-range transported continental plumes that consist of refractory aerosol components (i.e., BC). This is also supported by the comparable FT fvol values for both seasons under background conditions (Fig. S2). The decrease of fvol from the top of the MBL towards the ocean surface is attributed to sea spray aerosol emitted from the ocean, which contributes to the MBL aerosol population (Pirjola et al., 2000) 310 and is mostly non-volatile (Rasmussen et al., 2017;Bates et al., 2012). One can notice the elevated Npre-CCN and NCCN values near the ocean surface ( Fig. 6b and 6c), which is mainly due to the cloud scavenging that reduced Npre-CCN and NCCN values in the upper MBL. The ocean is unlikely a source of Aitken-mode aerosols because NPF near the ocean surface is rare due to the large condensation and coagulation sinks (Pirjola et al., 2000).

315
Enhanced organic mass concentration was observed in the MBL during summer (Fig. 7b). Comparison of the vertical profiles of mBC and organic mass concentration indicates that long-range transported continental emissions have a minor contribution to the organics in the MBL during the summer, suggesting a dominant surface source of the organics. The enhanced organic mass concentration is attributed to both primary marine aerosol and secondary organic aerosol formed from oceanic VOCs.
Previous long-term HR-ToF-AMS measurements at Mace Head station in the North Atlantic show that the aerosol organics 320 are similar to those observed during primary marine organic "plume" events (Ovadnevaite et al., 2011), and the mass fingerprints and H:C and O:C features were consistent with organics originating from primary marine sources. Ocean emitted VOC can also lead to formation of secondary organic aerosol as discussed in the previous section. The seasonal trend of ammonium is consistent with the contribution from marine sources shown by previous isotopic analysis (Jickells et al., 2003).

Seasonal variation 325
NCCN, Npre-CCN, and N>10 are higher in the summer than the winter at all altitudes ( Fig. 6a-6c). The higher NCCN in the FT during summer is to a large degree due to more frequent occurrence of long-range transported continental plumes from North America.
It is important to note that FT NCCN,bg is consistently higher during the summer (Fig. S2). This higher background is likely due to the greater influence of diluted continental emissions, as evidenced by the slightly higher mBC,bg in the FT during summer.
The seasonal variation of the FT Npre-CCN is likely due to stronger NPF during summer as a result of the higher DMS emissions 330 over the open ocean (Clarke et al., 1998;Williamson et al., 2019). The higher Npre-CCN and NCCN in the FT contribute to the elevated values in the MBL during the summer through entrainment. The higher MBL NCCN during summer is also partially due to the increased growth rate of nucleation and Aitken-mode particles as a result of stronger oceanic VOC emission (Zawadowicz et al., 2020;Zheng et al., 2020b). Stronger precipitation and thus coalescence scavenging of CCN can also contribute to the seasonal variation of NCCN in the MBL. 335 The spectral shape of submicron aerosol size distributions shows a strong variability between seasons and between the MBL and FT (Fig. 5)

. A clear separation between the Aitken and accumulation modes by a Hoppel minimum is evident in the MBL.
What stands out in the wintertime aerosol size distribution is the larger proportion of particles below 20 nm, potentially resulting from more NPF events due to low existing surface area concentrations. A recent study showed that over the ENA, 340 NPF takes place in the upper part of the decoupled MBL following the passage of cold fronts, when open-cell convection and scattered cumulus clouds frequently occur (Zheng et al., 2021). The NPF is due to the combination of low existing aerosol surface area, cold air temperature, availability of reactive gases, and high actinic fluxes in the clear regions between scattered cumulus clouds. The larger fraction of particles below 20 nm in the MBL during the winter is attributed to, at least partially, the more frequent passage of cold fronts over the ENA and NPF in the upper MBL (Kolstad et al., 2009). These newly formed 345 particles can continuously grow into the Aitken mode, and contribute to the CCN in the MBL (Zheng et al., 2018;Zheng et al., 2020a). The mean Aitken-mode particle size ( !,#$% ) during summer is approximately 20 nm larger than that during winter in the MBL (Fig. 5, and Fig. 6e), likely due to a combination of the following two reasons. First, a faster Aitken-mode particle growth is expected given stronger summertime emissions of ocean biogenic precursors. Second, both stronger wintertime convective activities and low CCN concentration lead to higher supersaturation, which allows the activation of smaller 350 particles, leading to a smaller Hoppel minimum size and thus smaller !,#$% . The vertical profiles of !,#&& of the two seasons are similar except at altitudes between 500 and 1500 m, where !,#&& is substantially larger during the winter (200 nm) compared to the summer (170 nm). The larger !,#&& during winter is attributed to the formation of large accumulation-mode particles by evaporating drizzle (see Sect. 3.5.2 for further discussion). The average size distributions in the MBL and FT during both IOPs are fitted with lognormal size distributions, and the Aitken-and accumulation-mode diameters and 355 concentrations, along with the geometric standard deviations are listed in Table 2.
The contribution of FT entrainment to MBL particle concentrations are estimated from the entrainment velocity and the difference in the average particle concentrations between the lower FT and MBL. The entrainment flux ' is calculated as: where ( is the entrainment velocity, )* and +,-are the average of property X in the lower FT and MBL, respectively. The 360 value of ( can be estimated from the dynamics of the MBL height . (Caldwell et al., 2005;Russell et al., 1998). The time variation of . is described by: where / is the large-scale subsidence rate, is the horizontal wind vector, and ⋅ . represents the variation of boundary layer height due to the horizontal advection. Assuming steady state conditions for boundary layer height (i.e., . / = 0), we have: 365 In the 40º (North and South) latitude range (i.e., the latitude of the ENA site), the average lower tropospheric subsidence rate (700 hPa) during low cloud conditions is around 8 mm s -1 (McCoy et al., 2017). Assuming lower tropospheric divergence is constant with height (consistent with previous analyses (Wood et al., 2009)), the mean subsidence rates at the top of the MBL (typically ~ 1.5 km) will be around 4 mm s -1 . The advection term ⋅ . is more difficult to estimate in general. For subtropical Eastern ocean regions (e.g., SE Pacific, NE Pacific), the advection term is roughly about 1/3 of / (Wood and 370 Bretherton, 2004). Assuming this relationship also applies in the ENA, we estimate the entrainment velocity at the top of the MBL using Eq. (3) as ~ 5 mm s -1 . The change of concentration X due to the entrainment 1' 12 ent is given by: (4) Taking into account the seasonal variation of the average MBL height (i.e., ~1200 m and ~1600 m for summer and winter, respectively), we estimate the particle concentrations above the MBL as the averages between altitudes of 1600 and 2200 m during summer, and of 2000 to 2600 m during winter. Particle concentrations inside the MBL are averaged from 400 to 1000 375 m for both seasons (Table 3).
The rate of N>10 change due to FT entrainment mixing (i.e., 13 !"# 12ent ) is estimated as 40 cm -3 day -1 and 53 cm -3 day -1 for the summer and winter, respectively. The total source of particle number is balanced by total loss under steady state conditions.
The major particle number sinks in the MBL are intermodal coagulation (i.e., coagulation of Aitken and accumulation mode), 380 in-cloud coagulation of interstitial aerosol, and in-cloud coalescence scavenging of CCN (Zheng et al., 2018). Using the long-term data collected at the ENA site, Zheng et al. (2018) estimated the sink for summer and winter seasons are 59 and 53 cm -3 day -1 , respectively. The estimated 13 !"# 12 ent suggests that the FT entrainment represents the dominant source of particle number in the MBL for both seasons. The rate of N>10 change due to FT entrainment mixing is dominated by pre-CCN (i.e.,

12
ent ) and the contribution from CCN (i.e.,

12
ent ) is essentially negligible (Table 3). It is worth noting that whereas 385 pre-CCN concentration in the MBL is lower during the winter season, 13 $%&'(() 12 ent is higher than that during the summer. The combination of higher 13 $%&'(()

12
ent and lower concentration suggests more efficient removal of MBL pre-CCN in winter, which is likely due to the following two reasons. First, the diameters of pre-CCN particles during winter are smaller than during summer, leading to a higher coagulation coefficient with larger particles (i.e., CCN and droplets). Higher concentrations of sea salt particles, as a result of stronger winds may also contribute to a higher coagulation rates during winter (Zheng et al., 390 2018). Second, abnormally high updraft velocities may be more frequent due to more convective activity during the winter, therefore growing more pre-CCN into CCN size range through aqueous-phase reactions.

Long-range transport of continental aerosols
As shown earlier, aerosol layers with elevated NCCN and mBC were observed above the MBL in the ENA during the summer IOP. These aerosol layers may strongly influence the aerosol properties in the region. The back trajectories of the air masses 395 arriving at the altitudes of the observed aerosol layers are examined to understand their origins. One example is on June 29, 2017 (Fig. 8), when an aerosol layer was observed at altitudes between 1300 and 2000 m. On this day, the vertical profile of potential temperature shows a MBL height of 400 m. The aerosol layer exhibited elevated concentrations of accumulationmode particles (modal diameter of ~120 nm). Increased BC mass concentration and CO mixing ratio were also observed in the layer. A frequency analysis of hourly 10-day back trajectories shows that a large fraction of air masses arriving on June 29 400 had travelled over the northern United States, likely bringing anthropogenic pollution and biomass burning aerosols to the ENA. Taking into account the transport time of air mass from the North America to the ENA site (generally 3 to 6 days), we generated the fire emission map from June 23 to 26 (Fig. S4), which suggest relatively strong biomass burning emissions along the air mass trajectories over the North America. A similar case of the long-range transport on July 18, 2017 is also illustrated in Fig. S5. 405 Previous ground observations at the Pico mountaintop station showed frequent elevations of summertime CO and O3 concentrations (Honrath et al., 2004;China et al., 2017). Based on air mass back trajectories, these events were attributed to the long-range transport of biomass burning emissions and anthropogenic pollution from North America. Air masses arriving from the continental U.S., western Atlantic, or northern North America exhibited elevated CO and O3 concentrations. When 410 the long-range transported plume was dominated by North American pollution with minimum biomass burning influence, the slopes of the linear fit between O3 and CO (d[O3]/d[CO]) ranged from 0.81 to 1.28, and the correlation coefficients (R 2 ) ranged from 0.53 to 0.83. During biomass burning dominated events, the slopes were between 0.4 and 0.9, and the value of R 2 reduced to 0.4 to 0.5 (Honrath et al., 2004). The value of d[O3]/d [CO] observed in the lower FT during the summer IOP of ACE-ENA is 0.79 (Fig. S1a), which is close to or within the ranges previously observed during anthropogenic pollution and biomass 415 burning events. The correlation coefficient for this study is 0.63, within the range of R 2 when air masses observed at the Pico mountaintop station were dominated by North America pollution.
The composition of FT aerosol layers provides additional insight into the source of the long-range transported aerosols. There are relatively strong correlations among the mass concentrations of BC, sulfate, and organics in the FT aerosol layers compared 420 to those in the MBL (Fig. S6), consistent with previous measurements of summertime aerosols at the Pico mountaintop station, which show that the FT aerosols are generally internal mixtures, including soot and sulfate coated by organic matter (China et al., 2017). During ACE-ENA, on average, the sulfate mass fraction between 1600-2600 m (i.e., the altitude range of FT aerosol layers during the summer IOP) outside of the background conditions (mBC > 5 ng m -3 ) is about 35%. For the two aerosol layers observed on June 29th and July 18th, the aerosol sulfate mass fractions are 38% and 30%, respectively. Typically, fresh 425 biomass burning particles are >90% carbonaceous material and the sulfate mass fraction is very low (Reid et al., 2005). As biomass burning particles age in the atmosphere, the sulfate mass fraction could increase as a result of condensation of secondary sulfate and/or coagulation with particles with a higher sulfate fraction. However, the air mass trajectories (e.g., in the cases of June 29 and July 18, 2017) show that the aerosol layers were above the MBL over the Atlantic Ocean before arriving at the ENA, and hence, it is unlikely that the sulfate in the aerosol layers derived from oxidation of DMS emitted from 430 the ocean. In addition, the sulfate in the aerosol layers was fully neutralized, in contrast to the typical acidic sulfate observed in the MBL aerosols (Zawadowicz et al., 2020;Zheng et al., 2020a). These pieces of evidence indicate that the sulfate in the aerosol layers originated from continental instead of oceanic emissions.
The sulfate mass fraction (~ 35%) in these FT aerosol layers is also substantially higher than that of aged biomass burning 435 plumes elsewhere. For example, previous FT aerosol sampled at the Pico mountaintop station under the impact of the longrange transport of the North America biomass burning plumes had an average sulfate fraction of around 16.3% (Dzepina et al., 2015). Holanda et al. (2020) sampled long-range transported African biomass burning in the FT offshore near the Brazilian coast of the Atlantic Ocean and found that sulfate represented 16% of the submicron aerosol mass. Furthermore, the mass concentration of potassium measured by PILS, a tracer of biomass burning aerosol (Andreae, 1983;Sciare et al., 2008), shows 440 no correlation with organic mass loading for the measurements within the altitude range of the aerosol layers during the summer (i.e., from 1600m to 2600 m) (Fig. S7). Although the back trajectories suggest a contribution of biomass burning, the above evidence indicate that the biomass burning aerosols are likely mixed with anthropogenic pollutions as they travel through the North America continent. On average, continental pollutions represent a major, and potentially the dominant, source of the aerosol mass in the layers observed in the lower FT over the ENA. 445

Classification of synoptic conditions
The synoptic condition strongly influences the structure of the MBL and thus aerosol properties. The Azores consistently lie in an area of substantial variability in synoptic configuration, thermodynamic environment, and cloud properties. The ENA site is under a strong influence from the North Atlantic high-pressure system (Azores high) and is periodically subject to frontal 450 passages (Rémillard et al., 2012). The synoptic conditions for the 39 flight days during ACE-ENA are classified as Azores high, pre-front, front, post-front, or unclassified conditions following the method described in Mechem et al. (2018). The classification of the synoptic conditions is based on the reanalysis fields of geopotential height at 500-hPa pressure levels (Gelaro et al., 2017). The 6-hourly reanalysis products are examined to judge the category of the synoptic conditions. This classification process is further combined with the archived surface weather maps obtained from the National Meteorological 455 Service of Germany (Deutscher Wetterdienst (DWD), http://www2.wetter3.de/index_en.html). The fractions of different synoptic conditions during all 39 flight days are 38.5% for the Azores high, 23.1% for pre-front, 25.6% for post-front, and 12.8% for unclassified conditions, respectively. The fractions of the synoptic conditions in each IOP are further shown in Table   4. No flight day was classified as front because no research flight was conducted on a day of a frontal passage due to logistical challenges. Under the Azores high, a strong inversion is often present at the top of the boundary layer, and the MBL is often 460 shallower and more likely to be well mixed. While under the pre-front, front, and post-front conditions, strong convective activities often lead to deeper boundary layers. As the MBL deepens, the turbulence produced from surface-heating and cloudtop radiative cooling becomes insufficient to maintain a well-mixed layer. Consequently, the MBL begins to "decouple" into a surface mixed layer and an upper decoupled layer (Wood and Bretherton, 2004;Bretherton et al., 2010). Earlier studies show that the boundary layer over the ENA tends to be decoupled much of the time (Rémillard et al., 2012). Another study found 465 that only 14% of the soundings over Graciosa were well-mixed (Ghate et al., 2015). Figure 9 shows the vertical profiles of meteorological parameters, CO mixing ratio, and aerosol properties measured on July 8 th , an example of Azores high conditions. On this day, the potential temperature and LWC indicate a well-mixed MBL with 470 shallow clouds below a strong temperature inversion at around 1000 m. Inside the MBL, the bimodal aerosol size distribution shows a clear Hoppel minimum, a result of cloud processing (Hoppel et al., 1994). The aerosol size distribution was largely uniform at different altitudes within the MBL. The aerosol in the lower FT showed a layered structure, with properties clearly different from that in the MBL, demonstrating the heterogeneity within the FT under a stable atmospheric structure controlled by the Azores high. An elevated concentration of Aitken-mode particles with mode diameter of ~35 nm was observed at 475 altitudes ranging from 1600 m to 2400 m. The elevated Aitken-mode concentration (i.e., Npre-CCN, ~900 cm -3 ) also coincides with an increased volatile fraction (i.e., up to ~80%) and low NCCN and mBC, suggesting the Aitken-mode particles derived from NPF in the FT over the open ocean when the existing accumulation-mode particle concentration is low (Clarke et al., 1998;Williamson et al., 2019). An aerosol layer with elevated mBC (~25 ng m -3 ) was observed above 2300 m. This layer exhibits a higher CO mixing ratio and mBC and is attributed to long-range transported continental emissions. The gradual 480 change of CO mixing ratio and mBC at altitudes above 1500 m also suggests the mixing between the layers of high Aitkenmode concentration due to NPF and long-range transported continental emissions. Figure 10 shows an example of vertical profiles of meteorological parameters, CO mixing ratio, and aerosol properties when 485 the MBL is decoupled. The measurements were carried out on February 8 th , 2018, when the front and associated cloud band are located at north of the Azores (i.e., pre-front condition). The vertical profile of potential temperature indicates a deep decoupled MBL that consisted of the surface mixed layer below 500 m and the upper decoupled layer from 500 to 1900 m. A thin layer of stratus was observed near the top of the surface mixed layer, and cumulus clouds were observed in the upper decoupled layer (Fig. 10a). Inside the surface mixed layer, the aerosol size distribution is bimodal and independent of altitude. 490

Pre-front and post-front conditions: Decoupled boundary layer
In contrast, the aerosol size distribution varied with altitude inside the upper decoupled layer. NCCN decreased with increasing altitude and exhibits a significantly lower value at the cloud level, a result of coalescence scavenging by the cumulus clouds.
Similar features are also evident from the averages of the vertical profiles during the IOPs (Fig. 6c). The lower NCCN in the upper decoupled layer where the cumulus clouds form indicates that surface measurements overestimate the concentration of CCN that are relevant for cumulus cloud formation when the MBL is decoupled. Npre-CCN also exhibits a decreasing trend with 495 altitude due to interstitial coagulation with cloud droplets. The accumulation-mode diameter inside the upper decoupled layer, below the cumulus cloud level (i.e., 600-1200 m), is substantially larger than that in the surface mixed layer (Fig. 10d). This larger accumulation mode diameter is likely due to the formation of drizzle drops through autoconversion and accretion of cloud droplets. The evaporation of the drizzle drops below clouds leads to fewer but larger accumulation-mode particles. Such a process also explains the vertical profile of accumulation mode diameter shown in Fig. 6f, given the deeper MBL and higher 500 drizzle rate during the winter season. Similar phenomenon can be observed under the post-front condition (Fig. S8), showing that the presence of the decoupled layer during pre-front and post-front conditions can lead to vertical heterogeneity of aerosol properties in the MBL.

Conclusions
In this study, we present aerosol properties, trace gas mixing ratios, and meteorological parameters characterized onboard the 505 G-1 aircraft during both the summer and winter IOPs of the ACE-ENA campaign. The key processes that drive the CCN population in the MBL are investigated by examining the variation of aerosol properties with altitude, season, and synoptic condition. On average, all particle concentrations (i.e., NCCN, Npre-CCN, and N>10) are higher in summer than winter at all altitudes. The elevated FT NCCN during summer is due to the periodic presence of long-range transported aerosol layers in the lower FT. The sources of aerosol in the layers include both biomass burning and anthropogenic pollution from North America, 510 with the contribution of anthropogenic pollution likely the dominant one for aerosol mass. In comparison, the influence of long-range transported continental emissions on FT Npre-CCN is weaker. Most of the observed seasonal variation in FT Npre-CCN is likely due to stronger NPF during summer as a result of a higher DMS emission rates over the open ocean. On average, FT NCCN is slightly lower than that in the MBL, indicating entrainment of FT air does not serve as a direct source of CCN in the MBL. However, entrainment of FT air is a major source of particle number (N>10) in the MBL for both seasons. The particle 515 number flux due to FT entrainment is dominated by the pre-CCN due to elevated Npre-CCN in the FT. Once the pre-CCN (i.e., nucleation and Aitken-mode particles) are entrained into the MBL, they can grow and reach CCN size ranges through condensational growth. The higher MBL NCCN during summer is also partially due to the increased growth rate of pre-CCN as a result of stronger oceanic VOC emissions. Stronger precipitation and thus coalescence scavenging of CCN may also contribute to the seasonal variation of NCCN in the MBL. 520 The chemical composition analysis shows that sulfate, organics, and ammonium dominate the non-refractory aerosol mass concentration (around 99% in both the summer and winter IOPs). The vertical profile of sulfate mass concentration indicates a surface source, consistent with the picture that over the open ocean, sulfate in submicron aerosol is mostly derived from DMS through both gas-phase and in-cloud oxidation. Stronger DMS emission and higher oxidant (e.g., OH) concentrations 525 lead to a higher MBL sulfate mass concentration during summer. Enhanced organic mass concentration was also observed in the MBL during summer, and is attributed to surface sources including stronger emission of primary organic marine aerosol and production of secondary organic aerosol from oceanic VOC.
The impact of synoptic conditions on the MBL structure and aerosol properties is examined. Under the pre-front and post-530 front conditions, stronger convective activities often lead to a deeper and decoupled boundary layer consisting of two sublayers, a surface mixed layer and an upper decoupled layer. In comparison, a well-mixed boundary layer is more prevalent under Azores high conditions. Aerosol in the decoupled boundary layers exhibits strong vertical variations. Coagulation scavenging and evaporation of drizzle below clouds lead to much reduced NCCN and larger accumulation mode size in the upper decoupled layer. Therefore, surface measurements (within the surface mixed layer) overestimate NCCN that is relevant for the formation 535 of MBL clouds under decoupled conditions.

Data availability
All ARM datasets used in this paper are publicly available on the ARM website (arm.gov/data). The data specifically related to the ACE-ENA campaign can be found at arm.gov/research/campaigns/aaf2017ace-ena 540 Author contributions YW, GZ, DAK, AL, AAM, FM, RM, AJS, JES, SS, AS, JT, R. Weber, R. Wood, MAZ, and JW collected and analyzed the aerosol and trace gas data aboard the G-1. YW, GZ, MPJ, DM, DV, R. Wood, and JW analyzed the cloud data and synoptic conditions. YW prepared the manuscript with contributions from all co-authors.          Table 1. Instrumentation deployed during the ACE-ENA that are used for data analysis in this study. All measurements were made at a frequency of 1 Hz except for the HR-ToF-AMS, PILS, and thermal denuder.

Instrument Measurement
Condensation particle counter (CPC) Total aerosol concentration > 10 nm Fast integrated mobility spectrometer (FIMS) Aerosol size distribution, 0.01 to 0.5 μm Fast-cloud droplet probe (FCDP) Cloud particles size distribution 2 to 50 µm High-resolution time-of-flight aerosol mass spectrometer (HR-ToF-AMS) Non-refractory aerosol composition Multi-element water content system (WCM) Liquid water content Particle in liquid sampler (PILS) Water soluble aerosol composition Single-particle soot photometer (SP2) Soot spectrometry Trace gas monitors Concentrations of CO and O3 Thermal denuder Sampling of non-volatile component of aerosol particles Ultrafine condensation particle counter (CPC) Total aerosol concentration > 3 nm Table 2. The fitting parameters of average aerosol size distributions for the Aitken mode and accumulation mode during the summer IOP and winter IOP in the MBL and FT. The particle concentrations are normalized to standard temperature and pressure (273.15 K and 101.325 kPa; STP).