Hi @健-伍, thanks for your question and feedback.
You are right that the lower part of the atmosphere was shown in a little more detail in the previous meteogram than it is now.

The values on the bottom line of the meteogram, in your case 1000, 1300 and 800 m, indicate the ceiling height, that is, the height of the lowest significant cloud layer above the model terrain. It is a model derived parameter, often referred to as Ceiling. For example, in the ECMWF model, significant cloudiness is defined as cloud cover exceeding 50% of the sky and condensate content exceeding 1.0 × 10⁻⁶.

The value 2,009 m shows the elevation of the location. This elevation forms the bottom of the vertical cross section of the meteogram.

The other levels shown in the vertical cross section are approximately 300 m / 1000 ft above the terrain, followed by the flight levels FL115, FL165 and FL300, which correspond to approximately 3.5 km, 5 km and 9 km. The values in kilometres are only an approximate conversion of the flight levels and are referenced to AMSL (Above Mean Sea Level).

The grey area in the vertical graph of the meteogram is calculated from relative humidity at individual pressure levels. It therefore shows where the atmosphere is very humid and where clouds can be expected to occur.

The darker the grey area, the higher the relative humidity and therefore the more likely clouds are to be present.

In your case (the circled area on the right), the grey area appears to extend from near the ground up to less than 3.5 km / FL115. It is therefore probably fog or clouds with a very low base, which may have a vertical extent of less than 1500 m (3,500 m − 2,009 m), with the cloud base perhaps at around 0 or 100 m above the terrain.

However, no value is shown on the bottom line because according to the model, this cloudiness does not cover more than 50% of the sky. So although the air near the ground is very humid and clouds or fog can therefore be expected, the condition required for calculating the Ceiling is not met.

I hope the new meteogram is clearer now.

I wish you many enjoyable trips to the mountains with beautiful views.

Jari

Hi @gsallman, thank you for your feedback. We hope you’ll be pleased to hear that we are working on adding a 0–6 km wind shear layer, as well as CIN and storm-relative helicity. Please bear with us a little longer.

Hi @texsun41, you can easily find a heat dome on Windy.com.

As you correctly mention, it is associated with an area of high pressure or a pronounced ridge. This feature is usually most clearly visible in the mid- and upper troposphere. At around 500 hPa, for example, it typically appears as a ridge or high-pressure system with unusually high geopotential heights.

You can display a heat dome as follows:

1. Select a parameter that can also be displayed at higher levels of the atmosphere, such as temperature or wind.
2. Select a pressure level, for example 500 hPa. On desktop, you can change it using the slider next to the aircraft symbol. On mobile, use the slider in the menu below the parameter selection.
3. Change the isolines to geopotential height.

The isolines will show you approximately how many metres above sea level the selected pressure level is located.

A heat dome can be identified by an area of unusually high values, often marked by closed contours or a pronounced northward bulge in the isolines. This pattern is clearly visible over the southern United States today.

At the surface, a heat dome may not be apparent because the surface pressure pattern is influenced by local factors, including the daily heating cycle, terrain, local circulations such as sea breezes and mountain–valley winds, boundary-layer mixing and soil moisture. A shallow thermal low may also form at the surface, as is the case today.

When television stations show a heat dome, they typically display geopotential heights in the mid-troposphere, usually at around 500 hPa.

@Kimjisung Hi, I agree with you. “Solar Power” can be misleading here, as it may suggest solar energy production rather than incoming solar radiation at the surface. “Solar irradiance” would be more appropriate. We’ll discuss this with the team. Thank you for pointing this out and for the suggestion.

@danielvfj Hi, please check that you are logged in on the mobile app, the alert is saved under Menu → My Alerts, and notifications and location permission are enabled for Windy.com in your device settings.

If everything is set correctly and you still do not receive alert notifications, please let us know what type of alerts you have set up, what device you have the app installed on, and which version of the app you are using.

@tornadouy Hi, thank you for your Lightning Jump suggestion and for the detailed description. I looked up different algorithms used to calculate it, and we will discuss it with our colleagues to see whether it could be reliably implemented with our lightning data.

Hi, the UKV model has been imported, so it is up to date again.

@Zack-Brown Hi, thank you very much for your suggestions. I like your idea. A filter showing lightning strikes within a selected radius, together with related notifications, could help users monitor nearby storm activity more precisely. We’ll keep your feedback in mind for future improvements.

For your information, lightning data is already used as part of our Live Storm Alerts.

@bobluck Hello, thank you for the report. Could you please confirm whether the problem appears in Safari only? If so, please send us your Safari version and macOS or iOS version. This will help us investigate the issue. Thank you.

@翊隆-鄭 Hello, we are sorry about the possible duplicate charge. Could you please send this information about your purchase, including the payment screenshot and the email address used for the purchase, to support@windy.com?

Our support team will check the transactions and advise you on the next steps. Thank you.

Hi @toddmpeterson, thank you very much for the suggestion. We appreciate the idea. At the moment, however, we are not planning to add a dedicated 500–1000 thickness layer.

Geopotential height and thickness are related, but they are not the same thing.

Geopotential height shows roughly the altitude of a specific pressure level, for example, how high the 500 hPa surface is above sea level.

Thickness, on the other hand, represents the vertical distance between two pressure levels, for example, between 1000 hPa and 500 hPa, and is therefore a proxy for the average temperature of that atmospheric layer. A larger distance between the levels generally indicates warmer air, while a smaller distance indicates colder air. Strong gradients in thickness usually correspond to strong temperature contrasts and often mark frontal zones.

You can derive a 500–1000 thickness by subtracting the geopotential height of the 1000 hPa level from that of the 500 hPa level.

If you let us know what you typically use a 500–1000 thickness for, we may be able to suggest another parameter that could serve a similar purpose.

Meteorological spring began in the Northern Hemisphere on Sunday, 1 March. Astronomical spring follows with the vernal equinox on 20 March. Across many parts of Europe, the first low concentrations of alder pollen are already being observed, and as temperatures rise and days grow longer, pollen levels are expected to increase rapidly.

Bee collecting pollen on a dandelion; photo by Nadine Eggenberger

After winter, flowering plants are a welcome sign of the new season for many of us, but for allergy sufferers, they also mark the start of symptoms. Pollen grains can trigger hay fever, asthma, or conjunctivitis. It is estimated that about one in four Europeans suffers from pollen allergy, making it one of the most common chronic conditions.

Scanning electron micrograph showing pollen grains of paper birch (Betula papyrifera) on the left and common ragweed (Ambrosia artemisiifolia) on the right; Cellimagelibrary.org

The timing of flowering and the amount of pollen in the air are closely linked to the weather. Meteorological conditions determine when plants start to bloom, how much pollen they release, how far the wind carries it, and when it gradually settles back to the ground. If we can describe when and how much pollen is released, numerical atmospheric models can be used to forecast pollen concentrations.

Pollen forecasts help allergy sufferers plan everyday activities and also allow them to start or adjust treatment in time, following medical advice, thereby reducing the risk of symptoms worsening during the season.

On Windy.com, you can view pollen concentration forecasts for selected pollen types at any location in Europe. These forecasts are provided by the Copernicus Atmosphere Monitoring Service (CAMS).

In this article, we will take a simplified look at how this pollen forecast is produced.

What is CAMS, and what does it provide?

The Copernicus Atmosphere Monitoring Service (CAMS) is part of the Copernicus programme and focuses on atmospheric composition monitoring. It is operated by the European Centre for Medium-Range Weather Forecasts (ECMWF). Its goal is to provide high-quality, well-documented, and openly accessible information on air pollution and other atmospheric components at global and European scales.

Example of a CAMS product: aerosol forecast; Windy.com

CAMS produces near-real-time analyses and forecasts of atmospheric composition, as well as long-term reanalyses that enable assessment of past developments and trends. Key products include forecasts of major air-quality pollutants such as particulate matter (PM₁₀ and PM₂.₅), nitrogen dioxide (NO₂), and ozone (O₃), as well as aerosol forecasts (including Saharan dust and wildfire smoke), estimates of fire emissions (GFAS), and information on the ozone layer and UV radiation. On the European scale, CAMS also provides pollen forecasts.

Why are CAMS pollen forecasts available only for Europe?

CAMS pollen forecasts are part of the CAMS Regional Production System. This system is designed specifically for the European domain, with a horizontal resolution of 0.1° × 0.1° (about 10 km). It uses detailed European vegetation datasets and a dense observation network, and is therefore not part of the global CAMS system.

European hourly air quality forecast of alder pollen concentrations; CAMS

The CAMS Regional Production System is based on 11 European chemical transport models for air quality: CHIMERE (INERIS, France), EMEP (MET Norway, Norway), EURAD-IM (Jülich IEK, Germany), LOTOS-EUROS (KNMI and TNO, the Netherlands), MATCH (SMHI, Sweden), MOCAGE (Météo-France, France), SILAM (FMI, Finland), DEHM (Aarhus University, Denmark), GEM-AQ (IEP-NRI, Poland), MONARCH (BSC, Spain), and MINNI (ENEA, Italy).

Ground-level alder pollen concentrations; CAMS

Coordination is provided by the Central Regional Production Unit (CRPU), led by Météo-France. Each model produces its own pollen forecast. Forecasts are run at 00 UTC and are available hourly up to four days ahead (T+96). The CRPU then combines the individual outputs into an ensemble product, typically as the median at each grid point.

On the Copernicus website, both the individual model outputs and the combined ensemble product are available.

Which pollen types are included in the forecast?

Six pollen types are included in the forecast: alder, birch, olive, grasses, mugwort, and ragweed.

These are among the most important pollen allergens in Europe, both because of their wide geographic distribution and because they can trigger allergic reactions in a substantial part of the European population.

Distribution of black alder (Alnus glutinosa) in Europe and the frequency of records in national forest inventories. Native range: EUFORGEN; European Commission Tree Atlas

In early spring, alder and birch are among the most significant pollen sources, especially across northern and central Europe. In many regions, alder can start the pollen season in the winter months, while birch dominates spring and is among the most allergenic tree pollens. Pollen from these trees most commonly causes allergic rhinitis and conjunctivitis, and in more sensitive individuals, it can also worsen asthma symptoms. Birch pollen is also well known for cross-reactivity with certain foods, which can trigger symptoms after eating some types of fruit or nuts. In Europe, around 8–16% of the population is sensitised to birch pollen.

Distribution of silver birch (Betula pendula) in Europe and the frequency of records in national forest inventories. Native range: EUFORGEN; European Commission Tree Atlas

From late spring into early summer, grasses dominate and are widespread across almost all of Europe. They are the most widespread source of airborne pollen and one of the most common triggers of pollen allergy. Typical symptoms include hay fever, allergic conjunctivitis, and, often, asthma-like symptoms. Grass pollen is recognised as one of the principal causes of pollen allergy in Europe, with national sensitisation rates of up to ~26% reported in some countries, although values vary by region and study population. In the Mediterranean region, olive tree pollen is another important spring and early-summer allergen, and in some areas it can be the main driver of symptoms, with high sensitisation rates.

Distribution of olive tree (Olea europaea) in Europe and the frequency of records in national forest inventories. Native range: EUFORGEN; European Commission Tree Atlas

Late summer and autumn are typically associated with mugwort and ragweed pollen. Mugwort is native to Europe and widely distributed, while ragweed is an invasive species that is spreading, especially in central and south-eastern Europe. Ragweed pollen is among the most potent aeroallergens and is often linked to severe allergic symptoms and asthma exacerbations. Sensitisation rates to ragweed pollen in Europe vary markedly by region, with some national estimates in the ~10–20% range (e.g., ~16% in Germany). These late-season allergens can extend the pollen season well into autumn.

Birch and alder pollen exposure in Europe. Maps were created using data from the Medical University of Vienna, Vienna, Austria, and represent medium to very high pollen levels; T. Biedermann, 2018

How is a pollen forecast produced?

In simplified terms, producing a pollen forecast can be broken down into a few key steps that together determine the resulting pollen concentrations in the air. First, emissions are estimated, that is, the spatio-temporal pattern of pollen release. Next comes atmospheric transport, dispersion, and deposition to the surface. Finally, the forecast is evaluated by comparing it with observations.

The individual operational models in the CAMS system may differ in how they numerically solve transport (also known as advection) and deposition, but the emission schemes are, to some extent, harmonised and coordinated across the system.

Common ragweed (Ambrosia artemisiifolia) in the early flowering stage, a highly allergenic invasive species in Europe; ECMWF

How does the model estimate pollen emissions?

In the model, the pollen season is not determined by real-world plant development, but by the accumulation of temperature above a threshold, expressed as temperature sums (degree days). The CAMS system uses two degree-day thresholds. Once the accumulated temperature sum exceeds the first threshold, the season starts. After the second threshold is reached, the season is considered to have ended.

The threshold values are derived from real-world measurements. However, these measurements are not ingested into the model operationally every day; instead, they are used retrospectively for validation and, if needed, for tuning parameters (such as the degree-day thresholds) during calibration. The operational forecast, therefore, relies on the meteorological forecast and predefined, calibrated parameters.

Map showing the temperature-sum thresholds for the onset and termination of the birch pollen emission season; M. Sofiev, 2013

Because in reality the season does not start abruptly, the model applies a gradual, probabilistic ramp-up and ramp-down of emissions. As the temperature sum increases, an increasingly larger fraction of the seasonal pollen reservoir is released.

Next, an emission flux is calculated for each grid cell, that is, the amount of pollen released into the air at a given moment (the number of pollen grains released per unit area per unit time). This requires input data on the distribution of individual pollen types (alder, birch, grasses, etc.) in the landscape and on their production potential.

How much pollen is actually released on a given day depends on the current weather. Emissions are therefore adjusted based on meteorological conditions. Rain and high humidity suppress release (pollen becomes moist and remains on the plant), whereas wind and turbulence enhance it by mechanically liberating pollen into the air. The wind effect is capped so that under very strong winds, emissions do not increase without limit but instead gradually level off.

At the same time, information on production potential is used to determine how much pollen is available for a given pollen type and season in a given area (the seasonal reservoir). Once this reservoir is depleted, pollen is no longer released, even if, according to the temperature thresholds, the season has not yet formally ended.

The output of the emission module is the pollen emission flux at each grid point and time step.

Photo by Alex Jones

What controls pollen transport in the atmosphere?

After emission, transport determines how pollen is distributed in the atmosphere. In the model, pollen is treated as a chemically inert particle, meaning it does not undergo chemical changes. It is carried horizontally by the wind (advection) and mixed within the atmospheric boundary layer by turbulence, which drives both vertical and horizontal dispersion. Turbulent mixing is often enhanced by convection, especially during daytime surface heating, when updrafts and downdrafts distribute pollen more efficiently across different heights. Transport is therefore mainly governed by meteorological conditions, such as the wind field, atmospheric stability, and the depth of the boundary layer.

In parallel, pollen is removed from the air by deposition. Dry deposition includes gravitational settling as well as turbulent impaction and interception on surfaces. For example, birch pollen typically has a settling velocity of a few centimetres per second.

Wet deposition is also very important, namely, precipitation scavenging within and below clouds. Precipitation has a major impact on pollen concentrations, and even relatively small rainfall amounts can substantially reduce airborne pollen levels.

Generation of the individual forecasts

By combining emissions, transport, and deposition, each of the 11 European chemical transport models produces its own forecast of pollen concentrations across the grid and over time, expressed in pollen grains per cubic metre.

These individual outputs are then combined into a single representative ensemble product, usually computed as the median value at each grid point.

Daily mean concentrations of birch pollen grains from the CAMS regional ensemble forecast; Copernicus

This ensemble approach improves forecast stability and reliability by reducing the influence of extreme errors from individual models and by blending their different representations of emissions, transport, and deposition into a more robust result.

Are pollen concentration forecasts compared with real-world conditions?

The final and crucial step in producing a pollen forecast is validation. Modelled pollen concentrations are retrospectively compared with real measurements from aerobiological monitoring stations of the European Aeroallergen Network (EAN). The evaluation considers, for example, overall forecast error, systematic bias, and whether the model correctly captures the onset of the season, its development, and peak concentrations.

Comparison of observed birch pollen concentration (MeteoSwiss/EUMETNET/Autopollen) with the CAMS forecast for the same period in Zurich; ECMWF

The EAN is a pan-European database of pollen and fungal spore measurements. Operating since 1988, it now includes data from more than 600 stations in 38 countries. It serves as a central platform for collecting, standardising, and sharing both long-term and near-real-time pollen data from national networks and laboratories.

CAMS Evaluation Report 2024: RMSE of birch, alder, and mugwort pollen forecasts at individual EAN stations (model error, pollen grains m⁻³); Copernicus

These data allow continuous assessment of model performance and the identification of weaknesses. Validation results support further model development and are used to adjust parameters and schemes in subsequent versions, including phenological thresholds, temperature sensitivity, the influence of humidity and precipitation, and source maps. In this way, forecasts can be progressively refined in future seasons, both at the level of individual models and of the overall ensemble product.

Pollen forecast on Windy.com

On Windy.com, you can view CAMS pollen forecasts in the detailed forecast for a specific location by switching to Pollen & Air Quality. If higher pollen levels are expected over the next four days, the chart will show the forecast concentrations for the relevant pollen types. The displayed value represents the CAMS ensemble median.

Pollen is shown alongside air quality to better assess the overall allergy load. Air pollution can worsen symptoms and, in some cases, interact with pollen grains to increase their allergenic potential. Pollutants such as particulate matter (PM₁₀ and PM₂.₅), ozone (O₃), nitrogen dioxide (NO₂), and sulfur dioxide (SO₂) can lead to more intense reactions than pollen alone. As a result, days with high pollen combined with poor air quality can feel more challenging than high-pollen days with cleaner air.

In January, winter climatically peaks in the Northern Hemisphere. Around 22–24 January, the Northern Hemisphere mean daily air temperature (1991–2020 climatological normal) falls to about 8.6 °C, the lowest daily average of the year.

Daily surface air temperature in the Northern Hemisphere; Climate Reanalyzer

Snow is an inseparable part of winter weather. In the following sections, we take a closer look at how a snowflake forms and its “journey” through the atmosphere. To do so, we briefly explore the basics of cloud and precipitation microphysics.

Modelled snow depth, 16 January 2026, 08 UTC (ECMWF); WIndy.com

Cloud Particle Formation

Our story begins in a cloud. For a cloud to form, that is, for tiny cloud droplets or ice crystals to begin forming, the air must first cool, most commonly through ascent.

Schematic diagram of a rising and sinking air parcel; Climate Water Project

For simplicity, let us imagine an air parcel, a small volume of air that we follow as it rises. As it ascends into lower-pressure levels, it expands and therefore cools. The amount of water vapor changes little initially, so the water vapor mixing ratio remains nearly constant, whereas relative humidity increases. This is because the saturation vapor pressure decreases rapidly with temperature (as given by the Clausius–Clapeyron relation), so colder air reaches saturation with less water vapor. At a certain altitude, the air cools sufficiently to reach saturation (100%), and water vapor can begin to condense or deposit to form cloud particles. In the real atmosphere, this process is strongly facilitated by tiny aerosol particles.

Warmer air reaches saturation with more water vapor, so it can hold more water vapor than colder air.

Homogeneous nucleation, the formation of cloud droplets and ice particles directly from water vapor without aerosol particles, is extremely rare in the real atmosphere. It is an energetically demanding and statistically unlikely process that would require very high supersaturation (on the order of hundreds of percent).

The relative size of water molecules to condensation nuclei; NOAA

In practice, therefore, condensation and, in colder clouds, ice formation occur mainly on aerosol particles, which are abundant in the atmosphere and provide surfaces where cloud particles can readily form. For droplet formation, we refer to cloud condensation nuclei (CCN); for ice crystal formation, we use the term ice-nucleating particles (INP). These processes are known as heterogeneous nucleation.

A single aerosol particle can act as both a condensation nucleus and an ice nucleus. The most effective ice nuclei are typically mineral dust particles (e.g., clays and desert dust), soil particles, volcanic ash, and certain biological particles (e.g., pollen, spores, and specific bacteria). By contrast, other aerosols, such as sea salt and many organic particles, often serve as effective condensation nuclei for droplets but are less effective at initiating ice formation.

Schematic illustration of ice-nucleating particle formation from mineral dust aerosols; EGU

Formation of Ice Crystals in Clouds

Because low temperatures alone do not automatically induce the freezing of water droplets, and suitable ice nuclei are often absent, cloud droplets at temperatures below 0 °C (32 °F) are frequently supercooled and remain liquid. As long as ice nuclei or other freezing triggers are not present, such as direct contact with ice, they can persist in the liquid phase well below the freezing point. Particularly small and very clean droplets can remain liquid even at temperatures tens of degrees below freezing. Only at very low temperatures, typically below about −38 °C (−36 °F), do supercooled droplets freeze spontaneously even without ice nuclei (homogeneous ice nucleation).

Occurrence of supercooled liquid water droplets and ice crystals in clouds as a function of air temperature; COMET via NOAA

Current meteorological understanding suggests that, in real clouds, ice particles form primarily by freezing supercooled droplets. The most common pathways are immersion freezing and condensation freezing. In immersion freezing, an ice-nucleating particle is immersed inside a supercooled droplet; once the temperature is low enough, ice forms on the particle and the droplet freezes. In condensation freezing, an aerosol particle first acts as a cloud condensation nucleus on which a droplet forms, and then serves as an ice nucleus that triggers freezing.

Schematic illustration of how MPs (microplastics) can promote atmospheric ice nucleation (a).Experimental images of a gradually cooled MP particle, showing the immersion-freezing process (b). The scale bar in (b) is 20 µm; Philip Brahana 2024

An ice crystal can sometimes grow directly from water vapour without an intermediate liquid-water stage; this is known as deposition nucleation. However, it is relatively uncommon because it typically requires supersaturation with respect to ice and highly effective ice-nucleating particles.

Another possibility is contact freezing, in which a supercooled droplet freezes upon contact with a suitable particle; its importance depends on the frequency of such contacts within the cloud.

Finally, there is homogeneous freezing, which we described above. This process typically occurs in very cold clouds.

Ice nucleation occurs through four mechanisms responsible for forming primary ice crystals in the atmosphere; Wikipedia

From Cloud Particles to Precipitation

For cloud particles to become precipitation particles, i.e., droplets or ice crystals that fall out of a cloud, they must first grow sufficiently. Only when their fall speed, determined by the balance between gravity and air resistance, exceeds the speed of the updrafts do they begin to precipitate. Precipitation particles are typically considered to have radii greater than about 0.1 mm.

However, the available supply of water vapor is not sufficient for all cloud particles to grow to precipitation size, because a liter of cloudy air can contain hundreds of thousands to millions of droplets and ice crystals. Only a fraction of them grow large enough to become precipitating particles, often at the expense of the others.

The primary mechanism by which small cloud particles grow is the diffusion of water vapor. In supersaturated air, water vapor deposits on their surfaces, causing the particles to grow. By contrast, in an unsaturated environment, water evaporates from their surfaces (or sublimates from ice), and the particles shrink and may even disappear entirely.

An illustrative example of the Bergeron–Findeisen process is a fallstreak hole (cavum). It forms when ice crystals develop locally within a thin cloud layer composed of supercooled water droplets, for example, after an aircraft passes through the cloud; NOAA

In clouds where ice crystals and supercooled droplets coexist, the Bergeron–Findeisen process takes place (a diffusion-driven growth process). Because the saturation vapor pressure over ice is lower than over liquid water at the same temperature, the air can be supersaturated with respect to ice while remaining unsaturated with respect to liquid water. As a result, ice crystals grow, while the supercooled droplets gradually evaporate.

Once ice crystals grow large enough for their fall speed to exceed the speed of the updrafts, they begin to fall out of the cloud. During their descent, they can clump together (aggregation) and collide with supercooled droplets, which freeze onto them almost instantly. The crystals then become coated with a layer of frozen droplets, a process known as riming, which can obscure their original shape. Some ice particles can also fragment, increasing their number within the cloud.

Sounding examples: rain, a sufficiently deep warm layer near the surface (left), and snowfall, with temperatures below freezing throughout the profile (right); Windy.com

If the air remains below the freezing point throughout its fall, the precipitation reaches the ground as snow; if it passes through a warmer layer above 0 °C (32 °F), it begins to melt and may fall as rain.

The Remarkable Variety of Snowflakes

A snowflake can consist of a single ice crystal or, more often, an aggregate of ice crystals that forms within a cloud and falls to the ground, where it accumulates as snow. Although snowflakes form through broadly similar processes, which we have largely described above, each one is unique—no two snowflakes are exactly alike.

Because the crystal lattice of ice is hexagonal, ice crystals naturally grow with sixfold symmetry, typically as plates or columns. In plates, growth tends to be strongest at the six corners, while in columns it is strongest along the six edges.

The morphology diagram shows how snow-crystal growth depends on temperature and water-vapour supersaturation; the water saturation line indicates liquid-water supersaturation relative to ice (a). Classic thin, flat stellar-dendrite snowflakes are typically photographed within a narrow range around −15 °C, while slender columns and needles form mainly near −5 °C (b); Kenneth G. Libbrecht 2017

What an ice crystal looks like depends mainly on the temperature and humidity in the cloud: sometimes delicate, highly branched dendrites dominate, while at other times plates, columns, needles, or grains are more common.


As a crystal grows, water vapor deposits onto it, and the crystal retains its hexagonal structure. During its fall, other ice crystals often attach to it, or supercooled droplets freeze onto it. Once a crystal or aggregate becomes visible to the naked eye, it is called a snowflake and can continue to grow. The average diameter of snowflakes is around 5 mm (≈0.2 in), and the largest documented size is reported to be 38 cm (≈15 in).

Close-up of a real snowflake; Photo: Jaroslav Fous

Cloud Types That Most Often Produce Snow

In temperate and subpolar latitudes (and, more rarely, even in the subtropics), snowfall most often comes from mixed-phase clouds containing both supercooled droplets and ice crystals, where snow formation is highly efficient due to the Bergeron–Findeisen process and subsequent aggregation. This is particularly true in the dendritic growth zone, roughly between −12 and −18 °C (about 10 to 0 °F), where crystals grow rapidly and readily aggregate into large flakes.

Cloud type: Significant snowfall can be produced by Ns, Cu con, Cb, and occasionally Sc; Valentin de Bruyn via Wikimedia Commons

Snow most often falls from the following cloud types:

• Nimbostratus (Ns) most commonly produces long-lasting, widespread snowfall. It occurs mainly along warm and occluded fronts and within the broad precipitation area of mid-latitude lows, including the wrap-around region (the “comma head”).

• Stratus (St) and stratocumulus (Sc) usually bring only light snow or brief snow showers, often referred to as low-cloud snowfall.

• Altostratus (As) is often a sign of an approaching warm front. Altostratus itself may produce light snow, but it often gradually transitions into nimbostratus, with precipitation intensifying and typically lasting longer.

• Cumulus congestus produces snow showers. It occurs mainly along cold fronts or in the cold air behind them. During lake-effect snow, it can form long bands (“cloud streets”), often together with stratocumulus (Sc).

• Cumulonimbus (Cb) is associated with very intense snow showers and, in rare cases, thundersnow, which can occur along strong cold fronts.

A striking lunar halo formed by diamond-dust ice crystals; Photo: Jaroslav Fous

Snow can also fall from ice clouds, but less often and usually less intensely than from mixed-phase clouds. These are mainly high-level clouds composed of ice crystals, especially cirrus (Ci), cirrostratus (Cs), and cirrocumulus (Cc). From these clouds, precipitation is typically limited to very light snow or isolated ice crystals, often as virga or diamond dust during dry, very cold weather. Heavy snowfall is rare because ice clouds typically have low ice-water content, the air beneath them is often dry, and they lack the efficient growth and aggregation processes typical of mixed-phase clouds.

Note: In one of our upcoming articles, we will focus on snowfall-favourable situations, covering both synoptic setups and locally driven mechanisms (e.g., orographic snowfall, lake-effect snow, and industrial/anthropogenic snow).

Snowfall and Snow Cover on Windy.com

On Windy.com, you can get a quick overview of current conditions and the expected evolution of snowfall in just a few clicks, or explore a more detailed analysis.

A meteorological radar is available, clearly showing not only precipitation coverage and intensity, but also, more recently, an estimated precipitation type.

Lake-effect snow over the Great Lakes as seen on Windy.com radar (10 November 2025); Windy.com

In addition, you can browse forecasts from numerical weather models. For selected models, you can display layers such as Snow depth and New snow, which represent accumulation over a time period of your choice.

From the meteogram for a specific location, you can track how forecast snowfall accumulation evolves over time, either at an hourly resolution for premium users or in three-hour intervals.

If you want to dive deeper into snowfall forecasting, Windy also offers tools for assessing the atmospheric temperature profile. You can use both measured soundings and model-based (forecast) soundings.

Freezing altitude layer; WIndy.com

A practical feature is the display of the freezing level (0 °C isotherm), either as a standalone layer or as isolines. Because the freezing level often determines whether precipitation falls as snow or rain, it provides a quick way to estimate the snowline. If multiple freezing levels appear in the profile, Windy displays the topmost one.

In mid-January, the far south of Russia’s Kamchatka Peninsula was affected by a series of deep cyclones. The associated winter storms produced exceptional snowfall. In Petropavlovsk-Kamchatsky, 131.6 mm (≈ 5.2 in) of precipitation was recorded from 13–16 January 2026 (120% of the monthly normal), with snow depth at the meteorological station reaching 170 cm (≈ 67 in). According to available records, this was the city’s highest snow depth in 55 years.

A snow-covered Kamchatka Peninsula after mid-January winter storms, 17 January 2026; Windy.com

In Canada’s largest city, Toronto, Toronto Pearson International Airport recorded a daily total of 46 cm (≈ 18.1 in) of fresh snow on Sunday, 25 January, setting a new one-day snowfall record. Observations at the airport date back to 1938.

Observations from Toronto Pearson International Airport, 26 January 2026 (rule of thumb: 1 mm liquid precipitation ≈ 1 cm snow); Windy.com

Even parts of northwestern and western Japan have experienced record-breaking snowfall this winter. Above-average snow depths are the result of frequent January snowfall, which affected Hokkaido in particular and the western part of Honshu.

In Sapporo, Hokkaido, snow depth reached 110 cm (≈ 43.3 in) on 26 January. The last time it exceeded the one-metre mark was in January 2005, 21 years ago.

Aomori captured by a webcam (3 February 2026). In early February, snow depth here reached 183 cm (≈ 72 in), the highest value in the past 40 years. Windy.com; webcam source: rab.co.jp

In Aomori City, in northern Honshu, snow depth reached 183 cm on Sunday, 1 February 2026. According to the Japan Meteorological Agency, this is the highest value in the past 40 years and the first time since 1986 that snow depth has exceeded 180 cm (≈ 70.9 in). The all-time record in Aomori is 209 cm, observed on 21 February 1945; records at this site date back to 1893.

Key facts about snow; WMO

Extreme snowfall in these regions was linked to winter storms during frigid Arctic air outbreaks. In addition to synoptic-scale mechanisms such as low-pressure systems, local and regional processes also played an important role in further enhancing snowfall totals.

In this article, we examine the processes that determine how much snow from a cloud actually reaches the ground and highlight selected historical episodes of extreme snow accumulation. Loosely building on the article Winter Weather: From Water Vapour to Unique Snowflake, we shift from snowflake formation to the conditions that allow snow to accumulate at the surface.

Typical Synoptic Setups for Snowfall

Snowfall most often occurs during synoptic setups that typically produce precipitation. If the atmosphere is sufficiently cold and temperatures remain below 0 °C (32 °F) throughout the entire vertical profile from the cloud layer down to the surface, precipitation will fall predominantly as snow. This includes various parts of low-pressure systems, from frontal boundaries (warm fronts, cold fronts, and occlusions) to the broad precipitation shields of cyclones, such as wrap-around/deformation zones (e.g., the comma head).

Schematic of a developing low-pressure system with the precipitation area highlighted; NOAA

In addition, some situations can trigger snowfall or enhance it locally, such as industrial snowfall, orographic snowfall, and lake-/sea-effect snow.

Schematic vertical cross-section of a cold front and a warm front with typical cloud types and precipitation; Met Office via Future Learn

Industrial Snowfall (Anthropogenic Snow)

Industrial snowfall most often occurs at low temperatures (often around −5 °C, 23 °F), under calm or light winds, with low clouds or fog, and with a high concentration of both condensation nuclei and ice nuclei.

Schematic of industrial snow formation; Velle Toll 2024

The low cloud layer typically has a base near the ground (below ~150 m, ≈ 500 ft) and a thickness of about 200–300 m (≈ 650–1,000 ft); higher-level clouds are often absent above it. Nocturnal radiative cooling of the upper part of the cloud/fog layer promotes mixing and the formation of supercooled droplets. Additional aerosols, moisture, and waste heat from industrial sources can further enhance mixing, accelerate ice particle growth, and trigger light snowfall, most often just before sunrise.

Usually, only a thin layer of snow falls, often less than 2 cm (≈ 0.8 in), but documented cases exist, for example, in West Virginia, USA (near the John E. Amos power plant, December 1975) or in the Texas Panhandle, USA (February 2014), where around 10 cm (≈ 4 in) of snow accumulated overnight.

Localized industrial snow over the Netherlands, observed by Landsat 8 (19 January 2017); NASA Earth Observatory via Edward Graham 2024

Orographic Snowfall

Orographic snowfall occurs in mountainous terrain when moist air is forced to rise along the windward side of a mountain range. As the air ascends, it cools, condenses, and forms clouds, from which snow can fall when the atmospheric column is sufficiently cold. The key factor here is not atmospheric instability, but the mechanical forcing of upward motion as airflow passes over the terrain.

Schematic of upslope snowfall and snow depth over a mountain range (areas B and C not included); Rebecca Mott 2018

A pronounced spatial variability is typical: tens of centimeters of snow (several inches to over a foot) can fall in a short time on windward slopes, while on the leeward side, where air descends and dries out, a so-called snow shadow forms with much lower totals, sometimes with no precipitation at all.

Another characteristic is persistence: under steady airflow and with sufficient moisture, snowfall can continue for hours or even days, often leading to very high accumulations in mountainous areas.

Mean daily precipitation: statistically, the highest totals are found in the tropics and mountainous regions; Copernicus Interactive Climate Atlas

Orographic snowfall is therefore responsible for the highest long-term snowfall totals in many mountainous regions worldwide, including the Alps, the Rocky Mountains, the Cascade Range, the Canadian Coast Mountains, the Japanese Alps, and New Zealand’s mountain ranges. Rapid snowpack buildup during persistent orographic snowfall also directly affects avalanche danger and hydrological conditions in mountain watersheds.

Orographic snowfall can also enhance precipitation that is already present. It can significantly intensify snowfall along fronts or amplify lake-effect snow in areas where snow bands encounter mountainous terrain.

Schematic of Lake-Effect Snow; NOAA

Lake-Effect Snow in the Great Lakes Region

In the North American Great Lakes region, lake-effect snowfall can develop as heat and moisture from the lake surface are transferred into the atmosphere. It occurs when very cold Arctic air flows over the relatively warm water, warming and moistening as it passes over the lakes.

Great Lakes surface-water temperatures and air temperature isotherms. A strong air–water temperature contrast supported the season’s first significant lake-effect snow event (10 November 2025); Windy.com

A commonly used rule of thumb is that for lake-effect snow to develop, the temperature difference between the lake surface and the air at about 1.5 km altitude (≈ 1 mi, ≈ 850 hPa) should reach at least 13 °C (23 °F). This promotes instability and the formation of cumulus (Cu) and stratocumulus (Sc) clouds, and in more extreme cases, even cumulus congestus (Cu con) and locally cumulonimbus (Cb).

Wind direction is also important because it determines the length of the airflow path over open water (fetch). The longer the fetch, the more heat and moisture the air can pick up from the lake. To produce long snow bands and “cloud streets,” a fetch of at least about 100 km (≈ 60 mi) is usually needed, although the exact value depends on the temperature contrast between the water and the air, atmospheric instability, and wind shear.

Lake-effect snow on radar: precipitation bands over Lake Michigan affecting the Chicago–South Bend area (10 Nov 2025).; Windy.com

A typical feature of these bands is pronounced spatial variability. They are often very narrow (only a few kilometers wide), so while more than 75 cm (≈ 30 inches) of snow may fall in one area, only a few centimeters may accumulate just a few kilometers away.

One of the most remarkable multi-day episodes occurred from 3 to 12 February 2007 across the Tug Hill Plateau near Redfield, New York, when up to 141 inches (≈ 358 cm) of snow fell over 10 days.

The largest reported 24-hour lake-effect total is often cited as 77 inches (≈ 196 cm) in Montague, New York, during 11–12 January 1997; however, because standard measurement procedures were not followed, this value was not accepted by an NWS review committee as an official 24-hour record.

Lake-effect storm (Feb 3–12, 2007): Max snowfall – Lake Ontario 141 in (≈358 cm; Redfield); Lake Erie 42 in (≈107 cm; East Aurora); NOAA

Lake/Sea/Ocean-Effect Snowfall

Lake-effect snow is not limited to the Great Lakes region, but can occur wherever very cold air flows over a sufficiently large, relatively warm body of water and then moves onshore. When the water body is a sea or an ocean, the same mechanism is referred to as sea-effect or ocean-effect snow. A classic example is Japan’s Sea of Japan coast, where the winter northwesterly monsoon produces heavy snowfall, often further enhanced by orographic lifting. Similar conditions occur, for instance, around the Black Sea (especially in northern Turkey, including Istanbul) and around the Caspian Sea, particularly along the southwestern coast of northern Iran.

Sea-effect snowfall over Japan (25 January 2026); Windy.com

Blizzards Are Formed by a Combination of Strong Winds and Snow

Snow plays a major role in a blizzard, but it does not necessarily have to be snowing during the event. The U.S. National Weather Service (NWS) defines a blizzard as a situation in which sustained winds or frequent gusts reach at least ~35 mph (≈ 56 km/h), and significant falling and/or blowing snow occurs, often reducing visibility to 1/4 mile (≈ 400 m) or less for at least 3 hours.

Weather warning layer: Blizzard Warning for parts of Minnesota and North Dakota (17 January 2026); Windy.com

A blizzard is typically associated with a deep low-pressure system (an intense winter storm), with the worst conditions often occurring in its rear sector and in the wrap-around/deformation zone (the comma head), where strong winds combine with snowfall. Another common scenario occurs behind a strong cold front, with gusty winds and snow showers. Hazardous conditions can also develop during a ground blizzard, when dry, powdery snow is lofted and blown around between a departing low and a strengthening high-pressure system.

During a blizzard, the main hazard is a rapid drop in visibility, sometimes to whiteout conditions, when falling and blowing snow makes it difficult to maintain orientation even over short distances. Strong winds quickly redistribute snow and create drifts, so even with relatively small new snowfall totals, roads can become impassable, vehicles can get stranded, and some areas may become isolated. The wind also greatly increases heat loss from the body (wind chill), raising the risk of hypothermia and frostbite.

Iran was hit by the world’s worst snowstorm in 1972: rescue operations (left); life in Tehran on 11 February 1972 (right). Ettela'at newspaper via Wikipedia

The early February 1972 blizzard in Iran (3–9 February) is often cited as one of the worst and, in some accounts, among the deadliest in modern history. Over the course of about a week, the country was hit by a series of exceptionally intense snowstorms accompanied by severe cold. In some areas, snowfall totals reached several metres, and snowdrifts in the south were estimated at up to 7.9 m (≈26 ft). The town of Ardakan and nearby villages were among the hardest hit, with no survivors reported in Kakkan and Kumar. During the cold spell, temperatures in some areas dropped to −25 °C (−13 °F). Estimates suggest the disaster claimed around 4,000 lives.

The Great Blizzard of 1978: Snow removal in Indiana (photo by Argil Shock / The Fort Wayne News-Sentinel), a car stranded in large snowdrifts in the Ohio Valley, and the synoptic setup over the eastern United States on 26 January 1978. NOAA; NOAA

In the United States, one of the most destructive blizzards was the Great Blizzard of January 1978, which occurred from 24 to 28 January. An exceptionally deep cyclone hit mainly the Ohio Valley and the Great Lakes region: its central pressure dropped to 955.5 hPa, and wind gusts in Cleveland reached 82 mph (132 km/h). The storm brought intense snowfall, massive snowdrifts, and widespread whiteout conditions. It claimed around 90 lives and caused an estimated $100 million in damage (in 1978 dollars).

A North Dakota DOT employee next to the top of a utility pole during a blizzard (March 9, 1966). Collection of Dr. Herbert Kroehl, NGDC via NOAA

One last interesting fact: Yuki-no-Otani

Yuki-no-Otani is a snow corridor along the Tateyama Kurobe Alpine Route, a mountain route across the Northern Japanese Alps that is world-famous for its giant snow walls. In the Tateyama massif, especially near the highest point of the route at Murodo (about 2,450 m (≈ 8,000 ft)), an exceptional amount of snow accumulates during winter. This is due to a combination of the winter northwesterly (monsoonal) flow, which brings moist air from the Sea of Japan, and strong orographic enhancement on the windward side of the Japanese Alps.

The Tateyama Kurobe Alpine Route runs through the Japanese Alps. Kyodonews/ZUMA Press via CNN

In winter, the route is closed due to the heavy snowfall and reopens in early spring, when heavy machinery must literally cut through the snow to create the corridor. In April 2025, the snow walls reached roughly 16 m (≈ 52 ft), while the historical maximum of around 20 m (≈ 66 ft) was recorded in 2000.

Snow walls on the Tateyama Kurobe Alpine Route; JNTO

Snowfall and Snow Cover on Windy.com

Windy.com offers a wide range of features to help you monitor snowfall, estimate how precipitation will evolve over the next few hours, and assess how much snow may accumulate on the ground.

View from a webcam in Livigno: Belvedere (1 February 2026); Windy.com; webcam source: livigno.panomax.com

For an up-to-date picture of current conditions, use radar and satellite observations, weather-station measurements, radiosonde observations, and webcams, which give a direct view of what is happening on the ground.

Sounding examples: rain, a sufficiently deep warm layer near the surface (left), and snowfall, with temperatures below freezing throughout the profile (right).

For the forecast, Windy.com includes multiple model layers. The key layers for snow conditions are New snow (forecast snowfall accumulation over the selected time period) and Snow depth (forecast snow cover on the ground). In general, the most useful level of detail is obtained by using the highest-resolution model available for a given area.

A vertical cross-section of the atmosphere along a selected route, using the Distance & Planning feature (VFR mode), showing flight levels, precipitation (blue for rain, cyan for snow), visibility, and freezing altitude (blue line); Windy.com

To estimate at what altitude snow may turn into rain, you can use the Freezing altitude information. It shows the approximate altitude of the highest freezing level (0 °C/32 °F) and can be viewed as a map layer, as isolines, or as a vertical cross-section along a selected route in VFR mode.

Hi @josepbc, I don’t have any new information yet on humidity-based alerts or alerts based on a combination of parameters. All the best in the New Year. Jari

As the strength of the Sun’s rays decreases and the polar night beyond the Arctic Circle expands day by day, winter is approaching in the Northern Hemisphere. Astronomically, it does not begin until the moment of the winter solstice, which this year falls on Sunday, 21 December 2025 at 15:03 UTC. Meteorological winter arrives earlier, on 1 December.

Even though autumn still officially reigns in the calendar, the first wintry days with freezing weather and, in places, a generous blanket of snow have already arrived in the Northern Hemisphere during the 2025/2026 season.

Solar Power: forecast solar radiation at Earth’s surface (ECMWF), 23–25 November 2025; Windy.com

The Northern Hemisphere has already experienced two notably cold episodes. The first came in the first half of November 2025, when Arctic air spread into the eastern two-thirds of the United States. The second affected western and central Europe and parts of southern Europe in the second half of November, bringing an early taste of winter weather with below-average temperatures and snow not only in the mountains but also in some places at lower elevations.

Forecast maximum temperatures and temperature anomalies for 11 November 2025; NOAA National Weather Service

According to current forecasts, another outbreak of cold Arctic air is on the horizon around the turn of November and December 2025, this time for much of Canada and probably part of the United States.

Forecast 2 m air temperature anomaly (°C) over successive 24-hour periods from 22 November to 1 December 2025; ECMWF

With winter on the way, terms like Arctic air outbreaks, the polar vortex, or sudden stratospheric warming will appear more and more often in weather forecasts. In this article, we will explain them in simplified terms, show how they are connected, and why they can influence the weather from North America across Europe to the Arctic.

Polar vortex

At its most basic, the polar vortex is a ring of strong westerly winds that, in winter, circulates the cold polar air near the pole. It forms as part of the general circulation of the atmosphere and, in terms of its size, is incomparable with any other vortical phenomenon on our planet.

Schematic of the stratospheric and tropospheric polar vortex over the Northern Hemisphere; NOAA Climate.gov

In the media, the term polar vortex began to appear frequently during the 2013/2014 winter season, in connection with the freezing winter in North America. However, it has been used as a technical meteorological term for a long time; it appeared as early as 1853 in the magazine Littell's Living Age.

Temperature anomalies in winter 2013/2014 show an extreme east–west contrast across the United States. The central and eastern regions were much colder than average due to repeated Arctic air intrusions, while the West, especially California, was significantly warmer than average; NOAA Climate.gov

In meteorological practice, we often distinguish between the stratospheric and tropospheric polar vortex. They are two layers of the same system, but their evolution over time need not be synchronized.

Stratospheric polar vortex

If only the term polar vortex is used, it usually refers specifically to the stratospheric polar vortex, that is, the circulation in the stratosphere, roughly between 100 and 1 hPa (about 15–50 km). You can think of it as a strong, relatively symmetric band of westerly winds encircling the pole, which is clearly visible, for example, on geopotential height maps at 10 hPa.

Animated map of Northern Polar region showing the polar vortex; NOAA Climate.gov

Tropospheric polar vortex

The tropospheric polar vortex, on the other hand, refers to the circulation of cold air and westerly flow in the troposphere, typically in the 500–200 hPa layer (the jet stream region), and also in the near-surface pressure field as a deep polar trough or a broad area of low pressure.

Schematic of the position of the polar night jet and the jet stream within the polar vortex; Met Office

The boundary between cold polar air and the warmer air of the mid-latitudes in the upper troposphere is formed by the polar jet stream. The polar jet stream (hereafter simply the jet stream) is a narrow, elongated band of powerful westerly winds at altitudes of roughly 8–12 km. It extends approximately between 50 ° and 60° latitude. This strong temperature and pressure contrast between cold and warmer air is the main reason why the jet is so fast: wind speeds in its core can exceed 200 km/h.

On a large scale, the jet stream is one of the main factors controlling weather in the mid-latitudes. It strongly influences the movement of low- and high-pressure systems, thereby determining the overall character of the weather. When it is relatively straight and strong, smooth westerly flow predominates; when it becomes strongly meandering, its troughs and ridges can bring deep intrusions of cold polar air towards the equator and, conversely, carry warm air far into the cold polar regions.

Schematic of a meandering jet stream and the positions of low- and high-pressure systems; NOAA Climate.gov

How the Polar Vortex Forms

In winter, the polar regions are under polar night conditions. The Sun does not rise at all for weeks to months. As a result, the air over polar regions cools strongly through radiative cooling. At lower latitudes, the stratosphere at middle and upper levels, roughly 30–10 hPa (about 25–30 km, or 16–19 miles), is significantly warmer. This creates a sharp temperature contrast between the warmer low-latitude air and the colder air over the polar regions.

The atmosphere tries to balance this difference. Warmer air tends to flow from the tropics toward the pole, but Earth’s rotation deflects it to the right in the Northern Hemisphere and to the left in the Southern Hemisphere. The closer the flow gets to the pole, the stronger this effect becomes. The result is a band of strong westerly winds (called the polar night jet) blowing from west to east around the cold air over the polar regions. These westerly winds form the stratospheric polar vortex.

As long as the pole remains much colder than the lower latitudes in the stratosphere, the ring of westerly winds is maintained.

Observed (bold blue) and forecast (red) polar vortex wind speeds at 60°N compared with the 1991–2020 average and the range of natural variability. Around 9 March 2025, a significant disruption of the polar vortex occurred; NOAA Climate.gov

Meteorologists most often monitor the zonal-mean westerly wind component at about 10 hPa and 60°N to determine when the stratospheric polar vortex has formed. Once a persistent westerly wind sets in during autumn and its speed stabilizes at values of at least 10–20 m/s (about 36–72 km/h, 22–45 mph) for several days, the stratospheric polar vortex is considered to have formed. Other parameters are also monitored.

Strength and stability of the polar vortex

When the westerly circulation in the polar regions is strong, the stratospheric polar vortex is strong and relatively stable. The jet stream in the troposphere tends to flow closer to the polar areas and has fewer large meanders. As a result, cold air remains confined over the polar regions (the Arctic during the Northern Hemisphere winter and Antarctica during the Southern Hemisphere winter), and winters in the mid-latitudes are usually milder, without pronounced cold outbreaks.

When the stratospheric polar vortex is strongly weakened or breaks down, the wintertime westerly circulation in the stratosphere is disrupted. The westerly winds weaken, or may even temporarily reverse to easterly. The polar vortex can be displaced away from the pole (a displacement event) or split into two or more parts (a split event). As a consequence, the jet stream in the troposphere often becomes highly meandering, facilitating deep intrusions of cold polar air into the mid-latitudes.

For example, when the polar jet stream dips far south over the United States, cold polar air can reach as far south as Texas and the Gulf Coast. Conversely, where the jet stream flows northward into higher latitudes, warm air is carried into those regions.

When the Arctic polar vortex is strong and stable, it keeps the polar jet stream farther north, allowing the coldest air to remain in the Arctic. When the vortex weakens, shifts, or splits, the jet stream becomes highly wavy, allowing warm air to enter the Arctic and polar air to spill into the mid-latitudes; NOAA Climate.gov

Causes of Polar Vortex Disruption

The westerly circulation around the cold polar air in the stratosphere can be disrupted for several reasons.

This regularly occurs due to the seasonal change in radiative conditions. With the onset of the polar day, ozone in the stratosphere above the pole once again begins to absorb UV radiation from the Sun. The polar stratosphere gradually warms, and the temperature difference between the pole and lower latitudes decreases until, at certain levels, the temperature gradient reverses and the westerly flow gradually changes to easterly (the so-called final warming).

During the polar summer, temperatures in the middle and upper stratosphere (roughly 30–10 hPa, that is, from about 25–30 km/16–19 miles upwards) above the pole gradually increase to the point where they eventually become warmer there than over the tropical latitudes. It is precisely this reversal of the temperature gradient in the middle and upper stratosphere that explains why easterly winds prevail there in summer instead of the westerlies typical of winter.

Rossby waves are large bends in the jet stream that turn an otherwise straight west-to-east flow into a wavy pattern; Wikimedia Commons

During the winter season, strong planetary Rossby waves are among the most common causes of disturbances to the stratospheric polar vortex. On upper-level height maps at 500 hPa or 300 hPa, they appear as large meanders of the jet stream in the form of deep troughs and pronounced ridges. They arise primarily as a consequence of Earth’s rotation and the fact that the effect of the Coriolis force changes with latitude. Their shape and strength are further influenced by mountain ranges, land–sea contrasts, and significant differences in heating between different regions, which distort the flow and can lead to the development of pronounced meanders.

When they are strong enough, they can propagate vertically upward across the tropopause into the stratosphere, where they gradually dissipate and transfer their energy and momentum to the surrounding flow. In doing so, they distort the wind field, slow down the zonal (westerly) winds at high latitudes, and thereby enhance the Brewer–Dobson circulation.

Average temperature profile in the lower atmosphere (left) and schematic of the Brewer–Dobson circulation (right); NOAA Jet Stream, EULIAA

The Brewer–Dobson circulation is the main large-scale circulation of air in the stratosphere. It provides a slow but crucial transport of air and trace gases (such as ozone). In this circulation, air rises from the troposphere into the stratosphere above the tropics and then, at higher altitudes, flows predominantly toward the poles (i.e., meridionally), where it gradually descends. This circulation exists year-round but is most pronounced during winter in the respective hemisphere. The concept of the Brewer–Dobson circulation was derived by Gordon Dobson and Alan Brewer from the distribution of ozone and water vapour in the stratosphere in the mid-20th century; it was only later confirmed by direct measurements.

If Rossby wave activity becomes very strong, the downward motion of air over the polar region intensifies. As the air descends, it warms rapidly adiabatically, and a sudden stratospheric warming can occur.

Sudden stratospheric warming

Sudden stratospheric warming (SSW) is a phenomenon in which the winter stratosphere over the polar region warms by several tens of degrees within just a few days. Most often, SSW is discussed in the context of the Arctic.

At levels around 10–30 hPa (about 20–30 km, roughly 12–19 miles above the surface), the temperature can jump by about 20–40 °C (35–70 °F). At the same time, the westerly zonal winds in the polar vortex weaken. During a so-called major SSW, the ring of westerly flow around the pole at 60°N and 10 hPa even temporarily reverses to easterly. The vortex then often shifts off the pole or breaks up into several parts, typically into two lobes.

Evolution of the stratospheric polar vortex: on 4 March 2025 (left) it is compact and more tightly centred over the pole, while by 10 March (right) it has been displaced off the pole, warm air is entering the Arctic, and easterly winds appear around 60°N, a hallmark of a sudden stratospheric warming; NOAA Climate.gov

At the surface, the effects of an SSW usually appear with a time lag of about 1–3 weeks. As the altered circulation in the stratosphere gradually propagates downward, the jet stream in the troposphere often weakens and becomes more wavy, and blocking high-pressure systems become more frequent in the polar regions. This increases the likelihood of significant outbreaks of cold Arctic air into the mid-latitudes (Europe, North America, Asia) and, conversely, intrusions of warm air into the polar regions. It is not the case, however, that every SSW automatically brings severe cold to a specific area in the mid-latitudes; rather, it increases the overall likelihood of cold polar air outbreaks in the weeks that follow.

Suppose an SSW and the subsequent weakening of the polar winds occur during winter, when solar radiation does not yet reach the polar regions directly (for example, in January or February). In that case, the polar stratosphere will cool radiatively again, become colder than the stratosphere at lower latitudes, and the westerly winds and thus the polar vortex will be re-established.

Average daily temperatures in the polar stratosphere (10-millibar pressure level) of the Northern Hemisphere from late 2022 into early 2023 (dark purple line); NOAA Climate.gov

Suppose the disruption occurs after the polar night ends and sunlight returns to the polar regions (for example, in March). In that case, the polar stratosphere is already beginning to warm, and the temperature difference between lower latitudes and the pole is weakening. In this situation, easterly flow tends to persist, and the polar vortex decays. It does not re-establish itself until the following autumn. Such an SSW is referred to as a final warming.

How to identify the polar vortex on meteorological maps and graphs

To monitor and describe the stratospheric polar vortex, several types of meteorological products are used, which together provide a picture of its strength, shape, and stability.

A basic indicator is the time series of the zonal-mean westerly wind at 60°N at 10 hPa. This graph shows how the average strength of the westerly flow in the mid-stratosphere changes over the course of winter: a strong, cold, and compact vortex corresponds to large positive values, while a weakened or disturbed vortex appears as a pronounced drop in wind speed or even a reversal of the wind direction.

Mean zonal wind at 10 hPa - Sub-seasonal range forecast; ECMWF

Equally important are maps of geopotential height and stratospheric flow, most often at the 10 hPa level. They clearly show whether the polar vortex is nearly circular and centred over the pole, or instead deformed, displaced toward lower latitudes, or split into several smaller vortices. Temperature and temperature anomaly maps then show how cold or warm the air inside the vortex is compared to climatology; a cold, low, “compressed” stratosphere indicates a strong vortex, whereas warming and positive anomalies over the pole signal that it is weakening.

For a comprehensive description of the state of the polar vortex, various indices (such as the Northern Annular Mode, NAM) and also “polar-cap” averages are often used. These are averages of quantities (for example, temperature anomalies or geopotential height) over the entire 60–90°N latitude band. These indices and polar-cap averages condense the complex spatial fields of pressure and temperature into a simple numerical time series, from which it is immediately apparent whether the polar vortex is powerful, near normal, or, conversely, significantly weakened.

How temperature affects the height of pressure; NOAA Jet Stream

Vertical cross-sections of the atmosphere are also very useful, often showing anomalies of geopotential height, either as latitude–height plots at a given time or as time–height plots over the polar cap (e.g., 60–90°N). This makes it possible to see how deeply into the troposphere the influence of a strong or weakened polar vortex extends, how its vertical structure changes throughout the winter, and whether a sudden stratospheric warming occurs.

Geopotential height anomalies over the polar cap during the 2024–25 polar vortex season. Positive atmospheric thickness anomalies over the Arctic are gradually propagating downward from the stratosphere into the troposphere; NOAA Climate.gov

When an SSW occurs over the polar region (typically shown on a vertical cross-section as an average over the polar cap 60–90°N), the air column in the middle stratosphere warms significantly, “puffs up,” and the 10–30 hPa pressure levels are shifted to higher altitudes. On vertical cross-sections of geopotential height anomalies, this appears as a column of positive anomalies over the pole, initially strongest in the upper stratosphere. Over time, this signal propagates downward into the lower stratosphere and the troposphere, eventually reaching the surface, where blocking anticyclones often form, allowing frigid air to spill into the mid-latitudes around them.

Daily geopotential height anomalies (17 pressure levels, last 120 days), normalized relative to the 1979–2000 daily climatology and averaged over the polar cap north of 65°N; NOAA Climate Prediction Center

Polar Vortex on Windy.com

You can easily view outbreaks of cold polar air in Windy in the Temperature or Wind (surface) layers. These views show where the cold air is coming from, how it spreads into the mid-latitudes, and allow you to switch on additional layers such as Clouds, Rain, thunder, or New snow. This gives you a much more complete picture of how the cold-air outbreak will affect the weather in your area.

If you are interested in the polar vortex itself, switch to higher levels of the atmosphere and display, in the same layers, for example, 150 or 10 hPa. In the standard atmosphere, the 10 hPa level corresponds to an altitude of roughly 30 km (19 mi) above sea level, i.e., the boundary between the middle and upper stratosphere. For more precise information about the altitude at each pressure level, select Geopotential height in the isoline settings.

Temperature (shading) and geopotential height (contours) at 10 hPa on 24 November 2025 and 8 December 2025. During this period, the displaced and weakened polar vortex undergoes a slight strengthening; Windy.com

Vertical cross-sections of the atmosphere (especially in the troposphere) are also a handy tool for visualizing temperature stratification and airflow. You can find them in the Airgram mode. In Windy, you can display them in two ways. The first option is to view the evolution of temperature and wind above a single point in the detailed forecast for a given location. The second option is a cross-section along a chosen route, which you can find in the Distance & planning tool. Draw a track, then select Airgram mode again.

On Windy.com, you can monitor the polar vortex and how it is forecast to change over time, from the Earth’s surface all the way up to the stratosphere – all in one place and with just a few clicks.

Vertical cross-section of the troposphere along the Distance & Planning route (Airgram mode); Windy.com

@Mhartain Martine, dobrý den,
článek bohužel existuje jen v anglické verzi. Na překlad ale můžete samozřejmě využít některý z dostupných online překladačů.
Moc děkuji za Váš zájem a milou zprávu. Jari

@mamankeericjoel Hi Eric, You are welcome to publish the article in SAT Magazine. Please note that it was initially published on Windy.com and that the author is Jaroslava Sochorova. No other formalities are needed. Thanks!

Hi @Fierrinho, thank you for your interest! The article is available only in the app, as we create and publish it there directly.
Have a nice day, Jari

@Wetterfranzi @jdebe @Wheats @idefix37 Hi, thanks for the feedback! If you have a tip for a weather or climate topic for one of our upcoming articles, please let us know. Kind regards, Jari