Marine cold air outbreaks (MCAOs) are extreme weather events that occur when a cold air mass advects over open water, with strong winds supporting intense surface sensible and latent heat fluxes capable of reaching 1000 Wm-2 over open water near the Arctic sea ice . Mid- to high-latitude MCAOs often contain mixed-phase clouds , initially organized as convective rolls or “cloud streets” . As surface turbulent fluxes and wind shear weaken and the MCAO ages, the convective rolls can evolve into closed-cell stratocumulus, transitioning into open-cells with further fetch . These transitions impact the cloud cover, which in turn impacts the radiative budget e.g., that the high-latitude climate is sensitive to. MCAOs can also spawn polar lows: small, intense and unpredictable cyclones that are consequential for coastal communities and shipping. The MCAO lifecycle is thus impactful to both high-latitude weather and climate.

MCAO cloud transitions are to first order dictated by environmental variability in surface fluxes and free tropospheric (FT) conditions, summarized by the marine cold air outbreak index . For example, mid-latitude air mass trajectories located under lower FT temperatures and weaker subsidence will transition sooner to open-cells . Numerical weather prediction models and reanalyses can depict the synoptic patterns that generate MCAOs. The models and reanalyses nevertheless struggle to represent the dynamic, thermodynamic, and microphysical processes and aerosol-cloud interactions that inform the transition on the sub-kilometer scale .

In warm stratocumulus, precipitation is a necessary condition for cloud transitions and provides a useful metric that can be compared between models and observations . The precipitation helps drive density-driven gust fronts that help initiate new updrafts by developing surface cold pools through evaporation . In the trade-wind regions, a rain rate threshold of 2 mmh-1, averaged over the rain's catchment area, corresponds well to downdrafts of air with equivalent potential temperatures (θe) lower than their surroundings . These support moderate updraft velocities (<2 ms-1) that can initiate new convection . Whether such relationships also hold for Arctic MCAOs is, to our knowledge, not known, as cold pool behavior has not been extensively assessed in suppressed environments with strong surface turbulent fluxes. , in a modeling study intended for the lower-latitudes, find strong, interactive surface fluxes induce smaller but more numerous cold pools with weaker gust front velocities, compared to simulations with horizontally-homogenized surface fluxes. One recent Arctic-related study finds most of the variance in flight-level turbulent fluxes occurs at scales <5 km for a separate MCAO case . The modeling of sub-cloud distributions of vertical velocity, moisture and temperature in marine environments is still challenging, however, even in warmer, better-documented environments .

Precipitation is also altered by aerosols. Higher cloud condensation nuclei concentrations (NCCN) can reduce droplet size, delaying the transition to lower cloud cover . Smaller drop sizes delay precipitation formation, well-accepted by now for liquid-only transitions . Smaller drops evaporating closer to the cloud base help destabilize the sub-cloud layer, better coupling the cloud to the surface turbulent fluxes . In mixed-phase clouds, smaller drops also rime less efficiently , which also helps maintain the cloud. More of the ice particle growth, especially at ice-saturated temperatures near -15 °C, is likely to occur through the Wegener-Bergeron-Findeisen (WBF) process , wherein smaller drops evaporate readily, and the vapor is deposited on neighboring ice particles. Ice growth through vapor deposition can also occur in water-saturated updrafts, capable of producing new cloud droplets . Where larger drops support riming, the larger rimed ice particles fall faster than the pristine dendrites and can more easily deplete the cloud liquid . Variations in ice-nucleating particle concentrations (NINP) furthermore compete with NCCN: some modeling suggests higher NINP facilitate earlier cloud deck breakups , but observations over northern Alaska indicate that if higher NINP are accompanied by elevated NCCN, the ice mass production is reduced . Observational analysis within sub-Arctic MCAOs suggests NINP may increase as more open water becomes available, although the INP source remains unclear . The ice particle number concentrations (Ni) can also increase through secondary ice production (SIP) processes over time, most convincingly at cloud temperatures >-10 °C . , in a modeling study, found that moderate updrafts (2–3 ms-1) are best able to support SIP, as weaker updrafts primarily deplete liquid through precipitation, and stronger updrafts do not allow sufficient hydrometeor growth.

High-quality in situ observations of Arctic mixed-phase cloud transitions are slowly becoming more available and analyzed thanks to recent field campaigns. During an aircraft sampling of a MCAO over the Norwegian Sea, observed a sharp decrease in total aerosol number concentration (Na) from approximately 100 cm-3 to 25 cm-3 across a closed-to-open cell transition, due to precipitation scavenging. concluded that the depletion in NCCN leading up to the transition acted to increase drop sizes, supporting SIP at temperatures between -8 and -3 °C (attributed to Hallett-Mossop rime splintering). The increased observed Ni correlated with more precipitation and was followed by the transition . Follow-up modeling suggests that ice-on-ice collisional breakup and freezing-induced-drop fragmentation were additionally necessary to reproduce the observed Ni . This is also consistent with and . Several other field campaigns have examined aspects of Northern Hemisphere/Arctic MCAO clouds , but those collecting in situ microphysical aircraft data have only occurred in cloud temperatures >-10 °C to date, to our knowledge . This is the temperature regime at which Hallett-Mossop rime-splintering is active, often the only SIP process represented in models. Observations in colder temperatures that preclude this SIP process remain rare.

During the Cold-Air outbreak Experiment in the Sub-Arctic Region CAESAR; campaign, an unusual observation was made of a closed-to-open cell cloud transition occurring over the Nordic Seas, far away from land influences, with cloud temperatures <-15 °C, under high aerosol loading conditions. The measured aerosol concentrations surpassed those from other recent campaigns . Given a Siberian source region, high winter and spring aerosol loadings may be more common than is often presumed . The polar front's more southerly location during winter allows anthropogenic emissions to advect into the central Arctic, creating what is known as Arctic haze . Aerosols in the Arctic can have long lifetimes and remain close to the surface due to the low moisture and high static stability . As the northern high-latitudes become more trafficked by ships within an increasingly ice-free Arctic , elevated aerosol concentrations from shipping emissions, and the resulting aerosol-cloud interactions e.g.,, should also become more common.

We use the instrumentation aboard the NSF NCAR C-130 during CAESAR to provide insights into the characteristics of this MCAO transition occurring on 29 February 2024 during Research Flight 2 (RF02). New observations include that of the sub-cloud thermodynamic vertical structure using a Raman lidar, and of liquid water paths (LWPs) made at a dynamically-relevant 500 m spatial scale. Aerosol measurements made upwind of the flight at Villum Research Station in northeast Greenland confirm the high aerosol loading, and further help constrain the aerosol lifecycle. Ice microphysical observations provide insight into the processes influencing the cloud morphological changes.

The RF02 flight track consisted of three surface legs (SL) flown at 300 m above sea level with spiral ascents/descents reaching 1750–2200 m in altitude bounding each SL (Fig. ; the final ascent away from SL3 was in-line). Table S1 in the Supplement shows the sampling time as a function of altitude for each vertical profile. Each SL is approximately aligned with the southwest cold air advection, with SL1 and SL3 following approximately the same flow line. This allows SL1 and SL3 to be treated as a single interrupted SL, which we call SL13. SL2 lies beneath a different free tropospheric airmass than SL13, and is not the main focus of this analysis. Fetch distance along SL13 is defined as the distance downstream from the beginning of SL1. Distances are calculated following the flight path along SL1, across the gap between SL1 and SL3, and continuing along SL3. Table S2 provides a lookup table showing the time at various fetch distances. The instrumentation is summarized in Table A1, with a brief description provided below.

On 26 February 2024, a sharp shortwave trough is situated southeast of Greenland, with a large area of diffluent flow over Iceland and the surrounding waters (not shown). A remnant surface low situated just west of Iceland rapidly strengthens underneath the large-scale divergence aloft, further enhanced by upward motion in the left exit region of the jet streak (not shown). Both the trough and surface low propagate towards the east-northeast. By 12:00 UTC on 27 February 2024, the trough becomes cutoff and vertically stacked with the surface low around 200 km northeast of Iceland, limiting upper level support. Research flight RF01 on 28 February and RF02 on 29 February sampled different portions of this surface low's MCAO sector.

By 03:00 UTC on 28 February, the low has moved northward (Fig. a). Despite lacking upper-level support, the low is still situated along a strong baroclinic zone, maintaining a maximum strength of approximately 960 hPa. A warm frontal zone stretches across Svalbard and cloud cover from a warm air intrusion spans from northern Scandinavia towards Svalbard. MCAO conditions are limited to the area between the island of Jan Mayen (70°59^′ N, 8°32^′ W) and Iceland behind a cold front with cold air advection off the sea ice supporting cloud streets.

By 16:00 UTC on 28 February, the low has moved to the north while maintaining its minimum central pressure (Fig. b). Strong low-level cold air advection off of the Greenland marginal ice zone creates a broad region of MCAO conditions. The cold front bounding this air mass contains a line of convection just off the Norwegian coast stretching towards Svalbard. RF01 sampled the post-frontal MCAO clouds, but this air mass moved north of the region sampled 1 d later during RF02.

By 03:00 UTC on 29 February, the center of the low is encircled by the MCAO, completing the occlusion process and causing the low to rapidly weaken (Fig. c). MCAO conditions stretch nearly all the way from the sea ice edge off eastern Greenland to the Norwegian coast, with a closed-to-open cell cloud transition occurring at around 5° E. The strong pressure gradient between the surface low and Greenland constrains the air flowing south past Villum to remain over the marginal ice zone. A new synoptic low (centered near 62° N, 10° W) further south encroaches, with warm air advection bounding the southern edge of the MCAO region. Twelve hours later, by 15:00 UTC, the occluded surface low driving the MCAO has weakened by 16 hPa (Fig. d). Warm air advection from the southern synoptic low reduces the southern extent of the MCAO and supports a cloud shield, also sampled by the remote sensors during the ferry legs of RF02 at 6 km altitude . The lower-level RF02 sampling took place within the MCAO persisting to the north of the encroaching warm air intrusion (orange track in Fig. d).

Despite the close proximity of SL13 to SL2, AROME-Arctic depicts different capping free tropospheric conditions (Fig. a, b). The 700 hPa air above most of SL13 at 15:00 UTC is moist, with relative humidities (RH) >90 %, compared to RH <10 % over SL2 (Fig. a). In situ FT measurements made during the cross legs from SL1 to SL2 and SL2 to SL3 confirm the sharp horizontal RH gradients in AROME-Arctic (Fig. a). During the cross leg from SL2 to SL3 flown at 2050 m, just 27 km south of the spiral descent to SL3, WVMR increases from 0.05 to 1.25 gkg-1 over 9 km. The 700 hPa temperature at 09:00 UTC was also warmer above SL13 (>-18 °C versus <-20 °C above SL2; not shown), prior to becoming modified by warm air advection (Fig. b). This is because more of the FT air above SL13 originates from the mid-latitudes (red back-trajectories in Fig. a, b) before wrapping around the occluded low. In contrast, SL2 resides below a modified cold, dry FT air mass originating from Greenland (blue back-trajectories in Fig. a, b). The air mass boundary is also apparent in in situ FT gas concentrations measured during both cross legs, with O3 decreasing from 134 to 51 ppbv and CO increasing from 81 to 115 ppbv between SL2 to SL3 (not shown). The two airmasses do become linked closer to Svalbard, where the cold airmass above SL2 forms a lobe of cold air located southwest of Svalbard (Fig. b). This creates a horizontal FT temperature gradient at the end of SL13 that ultimately weakens the cloud-top inversion (Fig. c).

A vertical AROME-Arctic cross section along SL13 (Fig. c) depicts the gradual deepening of the boundary layer (BL), driven by modeled buoyancy fluxes averaging 260 Wm-2 during SL1 (sensible/latent heat fluxes of 260/140 Wm-2), decreasing to 200 Wm-2 during SL3 (sensible/latent heat fluxes of 190/140 Wm-2) (not shown). The 700 hPa temperatures decrease by approximately 3 K along SL13, evident in the downward slope of isotherms. The cooling reduces the inversion strength, evident in the increased spacing of isentropes (thin solid lines) with fetch. The FT RH increases with fetch (dotted regions in Fig. c), consistent with observed mid-level clouds (Fig. ). In situ profiles (discussed in Sect. 5) confirm the gradient in the AROME-Arctic FT thermodynamic values, although the inversion is both ∼400 m lower and less sharp in AROME-Arctic compared to our observations (Fig. c).

CCN concentrations during SL1 were the highest observed during CAESAR, occasionally surpassing 800 cm-3. HYSPLIT back-trajectories and a FLEXPART footprint analysis of BC, initialized within the SL13 boundary layer (Fig. ), indicate the airmass likely originated from near Norilsk, Russia. Norilsk hosts the world's largest producer of Nickel and also smelts other metals, making Norilsk one of the world's largest SO2 emission hotspots and Russia's most polluted city . GEOS-CF simulations (Fig. S1 in the Supplement) further corroborate the back-trajectories, by indicating a buildup of SO2 originating from Russia over the central Arctic 2 d before RF02. In the day prior to RF02, a filament advects along the east coast of Greenland before deflecting around the occluded low and into RF02's flight domain.

Back-trajectories initialized along SL1 and SL3 pass within 60 km of Villum Research Station on 01:00–02:00 UTC, 29 February, approximately 13 h prior, and 01:00–22:00 UTC, 28 February, approximately 17–31 h prior, respectively (Fig. ). The longer travel times for the SL3 back-trajectories reflect the weaker winds near the center of the low, compared to the winds encountered by the SL1 air mass.

Na, CH4, and CO first increase at Villum on 27 February, and continue to increase to 326 cm-3, 2054 ppbv, and 142 ppbv respectively until 00:00 UTC, 29 February, before decreasing again (Fig. ). The correlation coefficients between Na and CH4/CO are 0.84/0.87, respectively, consistent with an anthropogenic origin. A decrease in NCCN at a 0.3 % supersaturation from 150 to 41 cm-3 from 21:00 UTC, 28 February, to 00:00 UTC, 29 February, coincides with an increase in Na, suggesting the CO and CH4-enriched air mass also contains less hygroscopic aerosol (Fig. a).

The mean aerosol number size distribution at Villum corresponding to the SL1 back-trajectories has a dominant accumulation mode with a peak diameter near 110 nm (Fig. ), similar to central springtime Arctic observations . A small enhancement in aerosol concentrations from 10–30 nm (Fig. ) is potentially indicative of new particle formation. The mean aerosol number size distribution at Villum corresponding to the SL3 back-trajectories is unimodal with an accumulation mode peak near 110 nm (Fig. ). The time-mean Na corresponding to the SL1 back-trajectories at Villum of 266 cm-3 exceeds that of 194 cm-3 corresponding to the SL3 back-trajectories.

Overall Na values are less at Villum than during RF02 (Fig. ). By the beginning of SL1, 4 h after the air mass has left the sea ice edge and 13 h after flowing past Villum, the aerosol number size distribution is bimodal, with a larger primary accumulation mode centered at 100 nm, and an additional smaller secondary Aitken mode centered at 20–40 nm, separated by a minimum at 40–50 nm (Fig. ). The Aitken and accumulation mode aerosol number concentrations (Nait and Nacc) exceed those measured on Svalbard in MCAO conditions . Observed (collocated) Na/NCCN decrease from 650/680 cm-3 during SL1, to 380/350 cm-3 at the beginning of SL3, to 125/90 cm-3 by the end of SL3 (Fig. ). NCCN is sometimes higher than Na, attributed to instrument uncertainty. BL CO concentrations of 140–150 ppbv measured during vertical profiles are consistent with values measured at Villum between 21:00 UTC, 28 February, and 01:00 UTC, 29 February (Figs. b, a–c). This suggests new aerosol sources, such as from sea spray production during transit over open water can help explain the Na increase. Sampling differences between the two locations (surface observations in a vertically stratified environment at Villum vs. airborne observations in a mixed BL during RF02) may also contribute.

By the time the air is sampled during RF02, CO and O3 are both well-mixed within the SL1 BL profiles (Fig. a–c), underneath a FT with more O3 and less CO. The FT gas concentrations remain relatively constant during SL1. Some rBC is also present in the boundary layer, consistent with an anthropogenic origin. By the last vertical profile, both the BL rBC and NCCN have declined substantially. The last ascent, at the end of SL3, is into a distinct FT airmass with higher O3 concentrations and lower CO concentrations as the aircraft turned towards the south (Figs. , d).

The four vertical profiles bounding SL1 and SL3 straddle 450 km (Fig. ), with the first three profiles done as spirals to minimize spatial variability. Over this distance, BL heights increase from 1270 to 2180 m, sub-cloud BL potential temperatures (θ) increase from 265 to 270 K, and flight-level temperature/moisture increases and moistens from -13 °C/1.2 gkg-1 to -8 °C/2 gkg-1 (Fig. ). Sub-cloud θ profiles indicate the BL is well-mixed (as do CO and O3 in Fig. ), while southwesterly BL winds of 12–17 ms-1 gradually weaken with fetch (Fig. ). The boundary layer is capped by a strong ∼ 10 K temperature inversion (Fig. a–c), with more water vapor in the FT than within the BL in all three spiral profiles (Fig. a–c). During the rapid in-line ascent from SL3 through open-cell convection, the aircraft traveled southward (Fig. ) and ascended into a much drier FT airmass, consistent with crossing the sharp gradient in FT moisture shown in AROME-Arctic (Figs. a, d).

The profiles bounding SL1 sampled liquid-dominated clouds with approximately adiabatic LWC profiles, corresponding to LWPs of 118 and 131 gm-2 (Fig. a, b). LWCs surpass 0.3 gm-3 at temperatures below -20 °C with cloud tops, reaching 1310–1320 m, overshooting the inversion by 10–40 m (Fig. a, b). Nd exceeds 600 cm-3 in both profiles bounding SL1 (Figs. a, b, ). An Na–Nd closure (Appendix C) using an adiabatic cloud parcel model establishes confidence in the unusually high Nd observations. Nearly all drops are <15 µm, and 80 % of all drops are <10 µm at both the beginning and end of SL1 (Fig. ). No drizzle (>75 µm) is observed by HOLODEC-II in either profile, although at least one drizzle drop was captured by the larger sampling volume of the F2DS during both profiles (not shown). The effective diameters within 100 m of cloud top at the beginning and end of SL1 are 11.6 and 8.5 µm respectively (not shown). HVPS-3 microphysical imagery indicates dendrites, aggregates, and small rimed ice throughout the cloud (Fig. a, b). The growth of dendrites and aggregates through vapor deposition is favored given cloud temperatures of -20 to -15 °C in the presence of small supercooled droplets. Ni remains below 3 L-1, with the exception of a spike to 6.5 L-1 at 750 m during the descent to SL1 (Fig. a).

By the spiral descent to SL3, the average in-cloud Nd has decreased to 302 cm-3 (Figs. c, ), with a sharply adiabatically-distributed supercooled LWC reaching a maximum of 0.58 gm-3 at the inversion height, corresponding to a LWP of 254 gm-2. The cloud-top height overshoots the inversion height by 20 m. The drop size distribution has shifted to larger sizes, with 60 % of drop diameters >10 µm (Fig. ), and an effective diameter within 100 m of cloud top of 16.1 µm (not shown). Drizzle (>75 µm) is observed by the HOLODEC-II between 880 m (-15.5 °C) and 1290 m (-18.6 °C) (Fig. c). Riming becomes more obvious, with rimed particles larger and more numerous (Fig. c). Ni remains <3 L-1 (Fig. c), however.

The rapid in-line ascent from SL3 through open-cell convection passes through clear sky prior to clipping the top of a convective cell (indicated in WCR reflectivity; not shown) with two discrete peaks of LWC: 0.23 gm-3 at 1790 m and 0.58 gm-3 at 2010 m (Fig. d). Average in-cloud Nd is now reduced to 32 cm-3, consistent with larger drop sizes: 80 % of drops have diameters between 20 and 30 µm (Fig. ), and drizzle (>75 µm) is observed by HOLODEC-II at temperatures as low as -22 °C (Fig. d). Large graupel are present, with diameters reaching 5 mm (Fig. d).

Over the course of SL13, the highest Na, NCCN and Nd reported during CAESAR (Na of 650 cm-3, NCCN of 680 cm-3 and Nd>600 cm-3), all decrease to less than 125 cm-3. These Na, NCCN and Nd are consistent within measurement uncertainties (Appendix C). Cloud droplet formation is primarily dependent on the activation of accumulation mode aerosols, which are depleted through entrainment with cleaner FT air and likely precipitation scavenging (Fig. ; Appendix D). Sea spray aerosol production is estimated to be an order of magnitude less than the aerosol depletion rate (Appendix D). The transfer of Aitken mode aerosols through activation into larger accumulation mode particles is also unlikely to keep up with the removal of accumulation mode particles, even when buffered through entrainment of FT Aitken mode aerosol (Appendix D).

Here we discuss how reductions in NCCN and Nd, and the resulting increasing droplet size, increase the riming efficiency with fetch. This in turn encourages more precipitation, as rimed particles fall faster than vapor-grown ice crystals, and thereby cold pools. In general, larger supercooled droplets and larger hydrometeors possess higher riming efficiencies. More specific values for the ice crystal collision/riming efficiencies as a function of ice crystal shape and size with the drop sizes measured during the three spiral profiles are provided in Table  using data from and .

The vertical profiles bounding SL1 (Fig. a, b) sampled aggregates and dendrites. Although some of the small ice at flight level appear rimed, dendrites are more common, indicating most ice particle growth is through vapor deposition. Despite the large LWPs, the small droplet sizes reduce the riming efficiencies . No cold pools are observed during SL1, consistent with a light liquid-equivalent precipitation rate and slow fall speeds of dendrites. Ice particles with lower fall speeds are more likely to sublimate closer to the cloud base, moistening the sub-cloud layer, which helps maintain coupling between the ocean and the cloud layer, similar to liquid drops .

By the beginning of SL3, the Nd has decreased to 300 cm-3, the cloud droplet sizes have increased, and some supercooled drizzle is observed (Figs. c, ). Although dendrites are present in the upper portion of the cloud, these become rimed at lower altitudes (Fig. c). Collision efficiencies increase by a factor of ∼2 between SL1 and the beginning of SL3 as drop sizes increase (Table ). Cold pools nevertheless do not appear until after NCCN decreases below 250 cm-3, after rapid BL growth (Fig. ). By km 282, Ni exceeds 5 L-1, and the ice habit suggests graupel (Fig. b, i). The LWPs have not increased since the beginning of SL3, indicating the dilution of NCCN, by entrainment, may be impactful for increasing the ice production (Fig. g). By the time open-cell convection sets in (Fig. b), and during the ascent away from SL3 (Fig. d), numerous large graupel particles are present with diameters reaching 5 mm, and cold pools are plentiful (Fig. d). The increase in ice particle size and Ni after the transition to open-cells is also apparent in the composite F2DS and HVPS-3 ice particle size distribution (Fig. ). The reduction in NCCN accompanying the rapid BL growth prior to km 282 – likely brought about by entrainment facilitated by a weakening temperature inversion, subsequently accelerated by aerosol deposition – appears critical for setting up the internal conditions facilitating riming, more precipitation, and thereby the transition in cloud morphology.

Turbulence also increases the collision likelihood and thereby riming efficiency between ice particles and cloud droplets, particularly when both particles have similar fall speeds . In situ flight-level turbulent kinetic energy across a 100 s (10 s) moving window decreases from 1.9 m2s-2 (1.0 m2s-2) during SL1, however, to 1.6 m2s-2 (0.74 m2s-2) during the open-cell segment of SL3, as horizontal wind speeds decrease. This suggests turbulence may increase overall riming efficiency, but cannot explain increased riming with fetch.

NINP during SL1 is comparable to other CAESAR flights, with concentrations around 2 L-1 at an instrument aerosol processing temperature between -28 and -26 °C . NINP decreases below the detection limit of the CFDC by the end of SL3 (Fig. i, Table ). Ni is 50–500× higher than NINP20 during closed-cell convection during SL1, consistent with SIP. Cloud temperatures are too low to support Hallett-Mossop splintering . Supercooled drizzle may be present during SL1, but only in small number concentrations that minimize the potential impact of freezing-induced-drop fragmentation. Another potential SIP mechanism is through ice-on-ice collisional breakup. One supporting observation is that, at an altitude of 750 m during the descent to SL1, Ni reaches a concentration of 6.5 L-1 (500× higher than NINP20), collocated with a spike in IWC to 0.25 gm-3 (and a depletion in LWC consistent with the enhanced removal of supercooled liquid through WBF processes and/or riming in the presence of higher Ni) at a temperature of -16.3 °C (Fig. a). This is near the temperature at which ice-on-ice collisional breakup is most effective . During the open-cell convection portion of SL3, Ni reaches concentrations of 12.5 L-1, more than 5000× higher than NINP20 (Fig. i, Table ). Supercooled drizzle (>75 µm) at temperatures below -18 °C within both profiles bounding SL3 (Fig. c, d) potentially support both freezing-induced-drop fragmentation and ice-on-ice collisional breakup. Ice is also present in updrafts with in situ vertical velocities >1 ms-1, crudely on par or exceeding the ice particle fall velocity (Fig. S2). Ni is greater than 0.1 L-1 (1 L-1) during 43 % (21 %) of the flight-level samples below open-cells with in situ vertical velocities >1 ms-1. Ni values exceeding 100× ambient NINP20 within updrafts indicates that their recirculation into open-cell convection could increase further Ni production .

The advanced and diverse instrumentation on-board the NSF NCAR C-130 during the CAESAR campaign characterized a closed-to-open cell transition within an MCAO made unusual by its high initial aerosol loading under cloud temperatures too cold to sustain Hallett-Mossop rime splintering, in the presence of surface buoyancy fluxes exceeding 250 Wm-2. In situ thermodynamic measurements confirm the synoptic pattern depicted by AROME-Arctic, supporting the inference that a cooling FT weakens the cloud-capping inversion, aiding the transition, as documented for warm-cloud transitions . The observations are summarized by the schematic in Fig. . Initially high CCN concentrations reduce the droplet diameters to <15 µm despite LWPs of 120 gm-2. The riming efficiency is low, and SIP through freezing-induced-drop fragmentation is unlikely. Ni remains mostly <1 L-1, and ice precipitation rates are too low to generate surface-based cold pools. A sharp cloud-top temperature inversion (∼10 K) limits BL growth by entrainment, despite strong surface fluxes. These factors support the maintenance of closed-cell convection.

By the beginning of SL3, as precipitation removes aerosol from the BL, and FT entrainment dilutes the BL aerosol number concentration, cloud droplet sizes increase. This supports increased riming. After a period of rapid BL growth likely facilitated by cooling FT temperatures that weaken the cloud-top inversion, SIP becomes more pronounced, with Ni>5 L-1. Given the cloud temperatures, the SIP mechanisms are most likely freezing-induced-drop fragmentation and ice-on-ice collisional breakup , aided by recirculation. A deeper BL paired with increased riming efficiency increases precipitation rates, falling into an increasingly subsaturated sub-cloud layer, supporting the formation of cold pools.

A strength of this study is the ability of the MARLi-derived measurements of sub-cloud thermodynamics and microwave-derived cloud liquid water paths to align with aircraft-measured vertical velocities, thereby depicting the cloud interactions with its environment. The cold pools appear to dynamically organize the convection through surface wind convergence, facilitating the transition to open-cell convection. The cold pools encourage convection at their upwind side at the expense of the downwind side, in contrast to trade-wind cumulus cold pools . However, the cold pools are weak and do not stabilize the sub-cloud air in the presence of the strong surface fluxes.

Our observations across the closed-to-open cell transition match some findings from , with both cases observing liquid-dominated closed-cells, ice-dominated open-cells with localized LWP peaks associated with updrafts, and a decrease in aerosol number concentrations across the transition. This is despite the case including warmer cloud tops of −14.5 °C and weaker surface fluxes. In our case study, SIP occurs with cloud base temperatures <-15 °C, outside the Hallett-Mossop rime-splintering temperature regime. Ice-on-ice collisional breakup and freezing-induced-drop fragmentation are hypothesized to dominate the SIP processes active during RF02. These SIP processes also likely contribute to the total Ni of the warmer case . A further distinction is that the BL during RF02 is much more polluted than that documented in , with aerosol concentrations of 650 cm⁻³ before the transition. The reduction in the BL aerosol number concentration through cloud-top entrainment, facilitated by a weakening cloud-top temperature inversion, appears to be crucial for establishing the timing of the closed-to-open cell transition.

While we are not aware of significant caveats to this observational study, the representativeness of this case has not yet been established. Further modeling work will establish the relative impact of aerosols, thermodynamics, and SIP on the transition timing as well as distinguish the importance of the environmental controls from internal mechanisms.