Will Sea Breeze Continue Inland for a Long Distance

Sea Breeze

The sea breeze is a surface manifestation of a thermally driven mesoscale circulation called the sea-breeze circulation, which often includes a weak return (land to sea) flow aloft.

From: Atmospheric Science (Second Edition) , 2006

Future Regional Climates

Jason Evans , ... Kendal McGuffie , in The Future of the World's Climate (Second Edition), 2012

9.2.2 Sea Breezes and Monsoons

A sea breeze is a coastal wind generated by the differential heating of the land and the water (Estoque, 1961; Rotunno, 1983). When the land is at a higher temperature than the neighbouring water, the air above it is heated and rises. The air at lower levels is replaced by cooler air advected from the adjacent ocean areas. Coastal temperatures are regularly impacted by sea breezes. If an offshore flow exists when a sea breeze forms, then there will be a region where the two opposing surface winds meet and are forced to ascend; this region is called a sea breeze front (Simpson et al., 1977). If there is enough moisture in the atmosphere, clouds and precipitation can form.

Sea breezes typically impact the land within some tens of kilometres of the coastline, a scale much too small to be resolved in today's global models. They are consistently occurring phenomena that can result in the coastal climate being several degrees cooler in summer than nearby inland areas, as well as having higher levels of precipitation. While sea breezes affect all coastlines to some degree, there are locations where the effect is larger, often because of the geometry of the coastline. Long, relatively thin peninsulas and islands can form sea breezes from opposing sides that will converge, greatly enhancing the potential for precipitation (Figure 9.4). The inter-relationships between sea breezes, the boundary layer, and larger-scale convection were studied during MCTEX in 1995 (Keenan et al., 2000; Schafer et al., 2001).

FIGURE 9.4. The local-scale sea breeze can have consequences that feed into larger-scale phenomena. In this sequence of radar images, two sea breeze fronts converge from the north and south coasts of Bathurst Island, Australia during the Maritime Continent Thunderstorm Experiment in 1995. The figure shows horizontal reflectivity at 2 km altitude at (a) 0216, (b) 0315, (c) 0416, (d) 0515, (e) 0614, and (f) 0715. The sea breeze fronts are evident in (a) along the north and south coasts of the island, and the convergence of the fronts prompts the development of more intense convection in the east, which develops, drifts in the environmental flow, and intensifies, forming an intense storm by 0515, which continues west and then dissipates. Carey and Rutledge, 2000. Refer to Carey and Rutledge, 2000, for discussion of other symbols on the figures and to Schafer et al., 2001, for further analysis of boundary layer development associated with these storms.

(Source: Carey, L. D., and S. A. Rutledge, 2000. © American Meteorological Society. Reprinted with permission.)

The spatial scale of sea breezes is generally too small to be captured by GCMs and hence high-resolution models have been used to study the role that the phenomenon plays in regional climate, at least as far back as Anthes (1978). One of the first studies to quantify the ability of a three-dimensional regional model to simulate a sea breeze was by Steyn and McKendry (1988). Since then, many studies have used regional-scale dynamical models to investigate aspects of the sea breeze and its interaction with other atmospheric phenomena (Crosman and Horel, 2010).

Monsoon is a name for seasonally reversing winds (Glickman, 2000). As for sea breezes, this phenomenon is caused, at least in part, by the differential heating of land and sea, though on a much larger scale. While sea breezes often occur as part of the diurnal cycle, monsoons are caused by the annual variation in temperature over large land areas relative to the neighbouring ocean surfaces. This temperature difference results in lower pressure over land during summer (driving an onshore wind) and over the water during winter (driving an offshore wind). Factors such as the topography of the land also have a considerable effect on monsoonal winds. Monsoon processes have been studied extensively (Webster et al., 1998; Wang et al., 2003), as has monsoon variability (Prell and Kutzbach, 1987; Wang et al., 2001). The onshore phase of the monsoon brings large amounts of precipitation and often occurs when the ITCZ is co-located. The major monsoon systems are the India, Asia–Australia, western Africa, and the North and South American monsoons.

Given the dependence of monsoons on the interaction between coastlines and nearby land orography, GCMs struggle to reproduce monsoon systems successfully (Cook and Vizy, 2006). Regional-scale dynamical models have been used extensively to study monsoon systems, beginning with two-dimensional studies (Dudhia, 1989). Recent studies driven by re-analyses have shown that regional models can simulate the monsoons better than coarser global models (Kumar et al., 2010; Zou et al., 2010; Kim and Hong, 2010), as well as providing insights into their physical processes. An example of this is a study over the maritime continent by Qian et al. (2010).

Read full chapter

URL:

https://www.sciencedirect.com/science/article/pii/B9780123869173000099

MESOSCALE METEOROLOGY | Overview

D.J. Parker , in Encyclopedia of Atmospheric Sciences, 2003

Sea Breezes and Other Thermally Generated Winds

The sea breeze is an example of a flow generated by the relatively rapid generation of a baroclinic zone (density gradient). Warming of the land surface after sunrise is rapid, while the sea surface temperature remains almost constant, so the boundary layer inland becomes relatively warm over the space of a few hours. This leads to a baroclinic overturning ( Figure 3A) and the development of a flow resembling a density current, as a cold sea breeze front pushes inland. The front may be perceived in cloud formed due to the forced ascent, or in visibility changes. Continued heating of the land surface means that there is likely to be relatively strong convective turbulence inland, which tends to dissipate the sea breeze front and render it a more diffuse baroclinic zone. However, as the surface heating diminishes in the evening, the convective turbulence decays and the front may intensify and propagate further inland.

Figure 3. Baroclinic tendencies in horizontal vorticity lead to thermally driven flows: (A) the sea breeze and (B) a katabatic, downslope wind.

A density current in the atmosphere is in a state of balance between the pressure gradient force due to the density change across the front, and drag on the current due to turbulent stresses (principally Kelvin–Helmholtz instability at the head). However, the sea breeze is also influenced by the Coriolis acceleration, over a time scale 1/f, and can be expected to turn with the Coriolis acceleration as the day progresses. Sea breezes are quite sensitive to the larger-scale flow, and will not develop if the ambient winds are strong.

After sunset it is possible for a land breeze to develop, as the land surface cools more rapidly than the sea surface. However, the ensuing surface inversion, which suppresses turbulence, does not develop as deeply as the daytime convective boundary layer, so the land breeze tends to be less active than the sea breeze.

On sloping terrain, the diurnal cycle of surface heating leads to baroclinicity relative to the background air, and tends to cause upslope, or anabatic, flow in the daytime and downslope, katabatic, flows at night (Figure 3B). Again, these flow regimes only develop under conditions of light ambient wind, but when they do occur they may dominate the local meteorology. Accurate representation of these flows in numerical models requires close attention to surface and boundary layer conditions.

Read full chapter

URL:

https://www.sciencedirect.com/science/article/pii/B0122270908004784

MESOSCALE METEOROLOGY | Overview

D.J. Parker , in Encyclopedia of Atmospheric Sciences (Second Edition), 2015

Sea Breezes and Other Thermally Generated Winds

The sea breeze is the best example of a flow generated by the relatively rapid generation of a baroclinic zone (a horizontal temperature or density gradient) at low levels of the atmosphere. Warming of the land surface after sunrise is rapid, while the sea surface temperature remains almost constant, so the boundary-layer inland becomes relatively warm over the space of a few hours. This leads to a 'baroclinic overturning' ( Figure 5(a) ) in which the warm air tends to rise and the cool air to subside, with corresponding horizontal flows to balance mass. As this circulation develops, the development of a flow resembling a density current is observed, as a cold sea breeze front pushes inland. The front may be perceived in cloud formed due to the forced ascent, or in visibility changes. Continued heating of the land surface means that there is likely to be relatively strong convective turbulence in the boundary-layer inland, which tends to dissipate the sea breeze front and render it a more diffuse baroclinic zone. However, as the surface heating diminishes in the evening, the convective turbulence decays and the front may intensify and propagate further inland. Like a cold pool from a convective storm, the sea breeze front is also sensitive to the ambient winds. In strong background, winds a sea breeze will not be observed. In lighter winds, an intense sea breeze front tends to occur in conditions of light offshore flow.

Figure 5. Baroclinic tendencies in horizontal vorticity lead to thermally driven flows; (a) the sea breeze and (b) a katabatic, downslope wind.

A density current in the atmosphere is in a state of balance between the pressure gradient force due to the density change across the front, and drag on the current due to turbulent stresses (principally Kelvin–Helmholtz instability at the head). However, the sea breeze is also influenced by the Coriolis acceleration, over a timescale 1/f, and can be expected to turn with the Coriolis acceleration as the day progresses. Sea breezes are quite sensitive to the large-scale flow, and will not develop if the ambient winds are strong.

After sunset it is possible for a land breeze to develop, as the land surface cools more rapidly than the sea surface. However, the ensuing surface inversion, which suppresses turbulence, does not develop as deeply as the daytime convective boundary layer, so the land breeze tends to be less active and less intense than the sea breeze.

Under conditions of light ambient winds, comparable circulations can be observed at boundaries between land surface types. These circulations are generally weaker, less coherent and harder to observe than the sea breeze, because the contrasts in surface heating between different land surface types is weaker than the land–sea contrast. However, coherent sea breeze-like circulations occur over cities (related to the 'urban heat island'), at forest-crop boundaries, and in semiarid regions at contrasts in soil moisture. There is increasing evidence, notably from the Sahel zone of West Africa, that these circulations forced by land surface contrasts in vegetation or soil moisture influence the generation of cumulonimbus rainfall. This is an important way in which the land surface patterns exert some control on rainfall, and therefore lead to feedbacks between the slow evolution of the land surface and the rapid mesoscale processes causing rainfall. Similarly, other results from the Amazon basin have indicated that mesoscale patches of deforestation may result in increased rainfall over the deforested areas, with a consequent promotion of forest regeneration in those areas.

On sloping terrain, the diurnal cycle of surface heating leads to baroclinicity relative to the background air, and tends to cause upslope, or anabatic, flow in the daytime and downslope, katabatic, flows at night (Figure 5(b)). Again, these flow regimes only develop under conditions of light ambient wind, but when they do occur they may dominate the local meteorology.

Thermally generated winds can be very important to the local meteorology and climate of a specific location, influencing the daily temperatures, cloud cover, winds, and rainfall on scales of a few to hundreds of kilometers. Therefore, these mesoscale features need to be accounted for in making assessments of the weather and climate of locations where surface contrasts are strong, such as coastal regions or large cities. Accurate representation of all of these thermally generated flows in numerical models requires close attention to the surface and boundary-layer conditions, as well as suitable model resolution to represent the mesoscale dynamics.

Read full chapter

URL:

https://www.sciencedirect.com/science/article/pii/B9780123822253004783

The Physics of the Atmosphere

Daniel Vallero , in Fundamentals of Air Pollution (Fifth Edition), 2014

2.4.1 Sea and Land Breezes

The sea breeze is a result of the differential heating of land and water surfaces by incoming solar radiation. Since solar radiation penetrates several meters of a body of water, it warms very slowly. In contrast, only the upper few centimeters of land are heated, and warming occurs rapidly in response to solar heating. Therefore, especially on clear summer days, the land surface heats rapidly, warming the air near the surface and decreasing its density. This causes the air to rise over the land, decreasing the atmospheric pressure near the surface relative to the pressure at the same altitude over the water surface. The rising air increases the pressure over the land relative to that above the water at altitudes of approximately 100–200  m. The air that rises over the land surface is replaced by cooler air from over the water surface. This air, in turn, is replaced by subsiding air from somewhat higher layers of the atmosphere over the water. Air from the higher-pressure zone several hundred meters above the surface then flows from over the land surface out over the water, completing a circular or cellular flow (see Figure 2.19). Any general flow due to large-scale pressure systems will be superimposed on the sea breeze and may either reinforce or inhibit it. Ignoring the larger-scale influences, the strength of the sea breeze will generally be a function of the temperature excess of the air above the land surface over that above the water surface.

FIGURE 2.19. Sea breeze due to surface heating over land, resulting in thermals, and subsidence over water.

Just as heating in the daytime occurs more quickly over land than over water, the radiational cooling at night occurs more quickly over land. The pressure pattern tends to be the reverse of that in the daytime. The warmer air tends to rise over the water, which is replaced by the land breeze from land to water, with the reverse flow (water to land) completing the circular flow at altitudes somewhat aloft. Frequently at night, the temperature differences between land and water are smaller than those during the daytime, and therefore the land breeze has a lower speed.

Read full chapter

URL:

https://www.sciencedirect.com/science/article/pii/B9780124017337000025

Mesoscale Meteorological Modeling

Roger A. PielkeSr., in International Geophysics, 2013

9.5.3 Sea Breeze System

The 2-D sea-breeze system adopted for comparison between Lagrangian simulations and analytical solutions is the one described by Defant's equations (Section 5.2.3.1 ). The non-linear advection terms are neglected in order to derive Defant's analytical solutions. Therefore, the analytical solutions are expected to more closely approximate the sea breeze system subject to smaller forcing, represented by the amplitude of the maximum mesoscale perturbation surface potential temperature $M$ (a measure of the land-sea temperature contrast). $M=1°\text{C}$ is adopted in the results shown here.

Read full chapter

URL:

https://www.sciencedirect.com/science/article/pii/B9780123852373000098

BOUNDARY LAYER (ATMOSPHERIC) AND AIR POLLUTION | Air Pollution Meteorology

X.-M. Hu , in Encyclopedia of Atmospheric Sciences (Second Edition), 2015

Evaluation of meteorological predictions

Meteorological processes (e.g., sea breeze, low-level jets) play a vital role in O ₃ events in Houston. Without properly capturing these meteorological processes, it is unlikely for the model to accurately predict the O₃ events in terms of time of occurrence, location, and peak values. The meteorological predictions are therefore first evaluated before the presentation of the chemical predictions.

Simulated temperature at 2 m (T2) is evaluated at 32 observational sites. The observed mean temperature during the simulation period is as high as 31.2 °C. Maximum surface temperature exceeded 40 °C on several days at certain sites during this episode. High temperatures accelerate chemical reactions and favor rapid production of secondary pollutants such as O₃. WRF/Chem-MADRID captures the diurnal variation of temperature quite well for most sites (with a high correlation of 0.92 with the observation) and only overpredicts T2 by 0.15 °C.

On the average, wind speed is overpredicted by 8.7%. The mean observed wind direction is south-southwestly while the simulated mean wind is biased by 25° to be more westly. WRF/Chem-MADRID captures the diurnal variations of the wind fairly well at most sites. In addition to the overall statistics, the statistics for wind speed are calculated for nighttime and daytime separately. The performance of wind speed at night (with a correlation coefficient of 0.374, and a Normalized Mean Bias (NMB) of 39.8%) is worse than that during daytime (with a correlation coefficient of 0.492, and an NMB of −10.9%). The worse performance during nighttime may be due to the well-known model deficiency in accurately simulating nocturnal turbulent mixing near the surface, which is a common problem for all numerical weather prediction models.

Planetary boundary layer (PBL) height determines the vertical extent of dilution of pollutants and significant uncertainties are associated with the estimation of PBL height in current AQMs. Evaluation of PBL height at five radar wind profiler sites shows that overall the PBL height is overpredicted by 72.1%.

The high O₃ mixing ratios in the Houston–Galveston area are at times associated with the occurrence of sea breezes. Sea breeze circulations were clearly observed on 29–31 August 2000 over this area. WRF/Chem-MADRID reproduced the observed sea breeze development sequence fairly well even though there is underestimation of the strength of the sea breeze (Figure 5). Large-scale offshore flow (westerly wind) near the surface persists for most of the morning in the Houston–Galveston area. A sea breeze develops around noontime. The front of the sea breeze reaches around Houston at 12:00 LT and a confluence line forms there, when the wind field is nearly stagnant around Houston. Such stalled sea breeze favors the buildup of high pollutant concentrations in this region. The sea breeze continues in the afternoon and the sea breeze front reaches more inland until 18:00 LT.

Figure 5. Predicted wind fields on 29 August 2000 by WRF/Chem-MADRID.

Read full chapter

URL:

https://www.sciencedirect.com/science/article/pii/B9780123822253004990

Mesoscale Meteorological Modeling

In International Geophysics, 2002

9.1 Introduction

In many mesoscale systems such as the sea breeze and squall line, phase changes of water occur as mesoscale and/or subgrid-scale circulations lift air above its condensation level and as water falls back out of or detrains from clouds and begins to evaporate. The presence of water as solid, liquid, and gas necessitates that the complete form of the conservation equations for water substance [e.g., Eq. (4-25)] be included in a mesoscale model. In addition, the proper representation of the source–sink term for diabatic heating [i.e., S_θ in Eq. (4-24)] is required. This chapter discusses procedures to allow the effects of the phase change of water to be included in grid-volume–averaged conservation equations, such as given by Eqs. (4-24) and Eqs. (4-25).

To parameterize the effects of phase changes in a mesoscale model, it is helpful to catalog the grid-volume–averaged atmosphere in a vertical column as convectively stable if $\partial {\overline{\theta }}_E/\partial z>0$ everywhere above the condensation-sublimation level of z _cl, or convectively unstable if for at least one level above, $z_{c1},\partial {\overline{\theta }}_E/\partial z\le 0.$ When a layer is convectively stable, forced lifting of the layer must continue to sustain the conversion of water vapor to liquid or solid once the specific humidity equals the saturation-specific humidity. If the layer is convectively unstable, however, clouds continue to grow without further forced lifting of the layer once saturation is reached. Convective instability is also called potential instability, or layer instability.

The variable ${\overline{\theta }}_E,$ the grid-volume–averaged equivalent potential temperature, is used to determine grid-volume–averaged convective instability. This temperature is derived as follows from the conservation-of-heat relation expressed by Eq. (2-23). Let the contribution resulting from the first three terms in Eq. (2-24) for the source–sink term S_θ be written as

(9-1) $\frac{C_{\text{P}}}{\theta }{S}_{\theta }^{*}=-\left({\delta }_{\text{f}}{L}_{\text{f}}+{\delta }_{\text{c}}{L}_{\text{c}}\right)T_{\text{v}}^{-1}\frac{d{q}_{\text{s}}}{dt}=\frac{C_{\text{p}}}{\theta }\frac{d\theta }{dt},$

where q _s is the saturation-specific humidity ¹ and L _c and L _f are the latent heats of condensation and freezing, respectively (L _c = 2.5 × 10⁶ J kg⁻¹ and L _f = 0.33 × 10⁶ J kg⁻¹ at 0°C). ² The parameters δ_f = 1 if freezing or melting occurs, δ_c = 1 if condensation or evaporation occurs, δ_c = δ_f = 1 if deposition or sublimation occurs, and 0 otherwise. Using Eq. (2-22), Eq. (9-1) can also be written as

(9-2) $C_{\text{P}}\frac{d\theta }{\theta }=-L{T}_{\text{v}}^{-1}d{q}_{\text{s}}\simeq -Ld\left(q_{\text{s}}/T_{\text{v}}\right),$

where the approximation $T_{\text{v}}^{-1}\left|d{q}_s\right|\gg q_s{T}_{\text{v}}^{-2}\left|d{T}_{\text{v}}\right|$ has been used ³ and L is equal to either L _c or L _f + L _c.

If at low temperatures, q _s /T _v → 0 (i.e., the saturation-specific humidity goes to 0 faster than temperature approaches absolute 0 [see, e.g., Eq. (9-8)], then Eq. (9-2) can be integrated to yield

$C_{\text{p}}{\displaystyle {\int }_{\theta }^{{\theta }_{\text{ES}}}d\text{In}\stackrel{\cap }{\theta }=-L{\displaystyle {\int }_{q_{\text{s}}/T_{\text{v}}}^{0}d\left({\stackrel{\cap }{q}}_{\text{s}}/{\stackrel{\cap }{T}}_{\text{v}}\right)}},$

where C _p and L are treated as constants. The upper limit of integration θ_ES is called the saturation equivalent potential temperature, since specific humidity is given by its saturated value q _s. The integrated form of this relation can then be written as

(9-3) ${\theta }_{\text{ES}}=\theta \text{exp}\left(\frac{L}{C_{\text{p}}}\frac{q_{\text{s}}}{T_{\text{v}}}\right)=\theta \text{exp}\left(\frac{L{q}_{\text{s}}}{\pi \theta }\right),$

where the definition of π given after Eq. (4-36) is used. With L = L _c, θ_ES represents the saturation equivalent potential temperature with respect to liquid water, and with L = L _c + L _f, the temperature is defined with respect to ice. This formulation for θ_ES is a measure of the change in potential temperature if all of the moisture is condensed (L = L _c), or deposited, or condensed and frozen (L = L _c + L _f), with the heat released used to warm a parcel of air. Because of the approximations made [e.g., Eq. (9-2)], the expression is not exact (see Simpson 1978 and Bolton 1980 for a precise derivation of θ_ES); however, it is in a suitable form for use in most mesoscale model calculations. ⁴

The grid-volume–averaged form of Eq. (9-3) is defined by replacing the instantaneous values of the dependent variables in Eq. (9-3) with their grid-volume–averaged counterparts. Expressed formally,

(9-4) ${\overline{\theta }}_{\text{ES}}=\overline{\theta }\text{exp}\left(L{\overline{q}}_{\text{s}}/C_{\text{p}}{\overline{T}}_{\text{v}}\right)=\overline{\theta }\text{exp}\left(L{\overline{q}}_{\text{s}}/\overline{\pi }\overline{\theta }\right).$

When an air panel is not saturated, ${\overline{q}}_s$ in Eq. (9-4) is replaced by the specific humidity of the parcel $\overline{q},$ yielding

(9-5) ${\overline{\theta }}_{\text{E}}=\overline{\theta }\text{exp}\left(L\overline{q}/\overline{\pi }\overline{\theta }\right),$

where ${\overline{\theta }}_E$ is the equivalent potential temperature. A layer with $\partial {\overline{\theta }}_E/\partial z\le 0$ will become less stable as it is lifted (as shown graphically by, e.g., Byers 1959:191), whereas a layer with $\partial {\overline{\theta }}_E/\partial z>0$ will become more stable. It is the vertical distribution of ${\overline{\theta }}_E$ that is used to assess convective stability.

Betts (1974) has demonstrated that vertical profiles of the difference ${\overline{\theta }}_{ES}-{\overline{\theta }}_E$ is a measure of convective regimes. Over Venezuela, he found that in the lowest levels (i.e., the 10-mb layer nearest the ground) ${\overline{\theta }}_{ES}-{\overline{\theta }}_E={40}^{o}C$ or so on dry days, and this difference was reduced by about half on disturbed days with extensive cumulus connection.

Read full chapter

URL:

https://www.sciencedirect.com/science/article/pii/S007461420280136X

MESOSCALE METEOROLOGY | Density Currents

P.G. Baines , in Encyclopedia of Atmospheric Sciences (Second Edition), 2015

Examples

Atmospheric examples of such flows are common. They include sea breezes , in which heating of the air over the land by convection from the solar-heated surface below causes a lateral density difference between that and the relatively cooler air over the ocean. The onshore-flowing sea breeze is the result. Similarly, the weaker offshore nocturnal land breeze occurs due to radiative nocturnal cooling over the land. Radiative cooling of the air causes flow down sloping topography, forming drainage flows, and valley winds, particularly at night. Another source is thunderstorms, which contain downdrafts of cold air due to the drag of falling raindrops, and cooling because of their evaporation. On reaching the ground, these downdrafts spread and form density currents. Cold fronts usually contain one or more squall lines, which consist (by definition) of a line of thunderstorms, so that the associated downdraft-produced density currents are conspicuous features of cold fronts. Such flows are sometimes made visible by suspended dust, which mark out the features of the cold density current. Noted examples of these are seen in the Sudan (where they are known as haboobs), and in India, Australia, and Arizona during dry summers. In these flows the dust usually makes a negligible contribution to the density difference, but this is not the case in another example of density currents – powder snow avalanches. Here turbulence causes the snow to be suspended in the air, producing the density difference that causes the downflow, which in turn produces the turbulence.

Read full chapter

URL:

https://www.sciencedirect.com/science/article/pii/B9780123822253001341

Megacities and the Coast: Global Context and Scope for Transformation

Sophie Blackburn , ... César Marques , in Coasts and Estuaries, 2019

Sea/land breezes : The differential heating of water and land generates daytime sea breezes (onshore flows of air) and land breezes (offshore flows of air). In summer, sea breezes are important in coastal cities in mitigating heat stress, with implications for thermal stress and air quality. Tokyo, for instance, is considering removing buildings to allow the sea breeze to penetrate and aerate the city. Other cities are also considering the orientation of the buildings in order to affect the inflow of sea air (e.g., Hong Kong and Singapore). This needs to take into account both wind flow and solar gain (e.g., Ng et al., 2011). Sea/land breezes also serve to concentrate and recirculate pollutants across coastal cities with important implications, particularly at night when the urban boundary layer (and thus atmospheric mixing) is diminished.

Read full chapter

URL:

https://www.sciencedirect.com/science/article/pii/B9780128140031000381

Coastal Meteorology

S.A. Hsu , in Encyclopedia of Physical Science and Technology (Third Edition), 2003

II.A Land and Sea Breezes

The best example of local winds in the coastal zone is perhaps the land–sea breeze system. This coastal air-circulation system brings fresh air from the sea in the afternoon to cool coastal residents, whereas farther inland hot and still air is the general rule. On coasts and shores of relatively large lakes, because of the large diurnal temperature variations over land as compared to that over water, a diurnal reversal of onshore (sea breeze) and offshore (land breeze) wind occurs.

A sea breeze develops a few hours after sunrise, continues during the daylight hours, and dies down after sunset. Later, a seaward-blowing land breeze appears and continues until after sunrise. The sea breeze may extend up to 50 or 100 km inland, but the seaward range of the land breeze is much smaller. In the vertical, the sea breeze reaches altitudes of 1300–1400   m in tropical coastal areas, with a maximum speed at a few hundred meters above the ground. In contrast, the nocturnal land breeze is usually rather shallow, being only a few hundred meters deep. Typical horizontal speeds of the sea breeze are of the order of meters per second, while the vertical components are only a few centimeters per second. At specific locations, large and abrupt temperature and relative humidity changes can occur with the passage of the sea-breeze front. An example of the land- and sea-breeze system is shown in Fig. 7. The onshore and offshore wind components are shown at 3-hr time intervals during the day. The lower portion of the onshore flow is the sea breeze and that of the offshore flow is the land breeze. The maximum wind speed and its approximate height in each current are depicted by arrows. The elliptical shapes in the figure illustrate the horizontal and vertical extent of the land–sea breeze circulation. The dashed horizontal line represents the 900-mb pressure surface (approximately the convective condensation level). At 0900 LST (local standard time), the air temperature over land is still cooler than over the sea and the land breeze is still blowing. By 1200 LST, the land has become warmer than the water, and the circulation has reversed. At this time, a line of small cumulus may mark the sea-breeze front. At 1500 LST, the sea breeze is fully developed, and rain showers may be observed at the convergence zone, 30 to 40 km inland. Because of a low-level velocity divergence, there is a pronounced subsidence and thus a clear sky near the coastal area at this time. At 1800 and 2100 LST, the sea breeze is still clearly present but is gradually weakening in intensity. By midnight or 0000 LST, the sea breeze is barely evident aloft, and the surface wind is nearly calm over land. At this time, a temperature inversion and occasionally fog appear over land. After land again becomes cooler than the water, a land breeze becomes well developed by 0300 LST and reaches its maximum intensity near 0600 LST. A weak land-breeze convergence line and associated line of cumulus clouds develop offshore near sunrise. The land breeze continues until midmorning, when the sea-breeze cycle starts over again. It is interesting to note that in this model the maximum strength of the land breeze in the near-surface layer is comparable to that of the sea breeze. Because of day–night differences in stability and frictional effects, however, the observed strength of the daytime sea breeze at the surface is considerably greater than the nighttime land breeze.

FIGURE 7. A simplified synthesized observed life cycle of the land–sea breeze system along the Texas Gulf Coast. Arrow lengths are proportional to wind speed. See text for explanation.

The importance of the effect of latitude on the sea-breeze circulation has been investigated numerically by scientists at the U.S. National Center for Atmospheric Research. They show that at the equator the absence of the Coriolis force results in a sea breeze at all times. At the other latitudes, the Coriolis force is responsible for producing the large-scale land breeze. At 20   °N, the slower rotation of the horizontal wind after sunset produces a large-scale land breeze that persists until several hours after sunrise. At 30   °N, the inertial effects produce a maximum land breeze at about sunrise, and the land breeze is strongest at this latitude. At 45   °, the rotational rate of the horizontal wind after sunset is faster, so that the maximum land breeze occurs several hours before sunrise. These results indicate that the Coriolis force may be more important than the reversal of the horizontal temperature gradient from day to night in producing large-scale land breezes away from the equator.

Onshore penetration of the sea breeze varies with latitude also. In midlatitude regions (generally above 40   °N), even under favorable synoptic conditions, the sea breeze may extend to 100 km. At latitudes equatorward of about 35   °, much greater penetrations have been reported, for example, 250 km inland of the Pakistan coastline. In tropical regions of Australia, the sea breeze can penetrate to 500 km. In such cases, the sea breeze traveled throughout the night before dissipation occurred shortly after sunrise on the second day.

The effect of the sea breeze on the long-range transport of air pollutants to cause inland nighttime high oxidants has recently been investigated by Japanese scientists. On clear nights with weak gradient winds, a surface-based inversion layer often forms between sunset and sunrise. A strong inversion forms mainly in basin bottoms in the inland mountainous region between Tokyo on the Pacific Ocean and Suzaka near the Japan Sea. An air mass that passes over the large emission sources along the coastline can be transported inland by the sea breeze in the form of a gravity current. In the case studied, a high-concentration layer of oxidants was created in the upper part of the gravity current. It descended at the rear edge of a gravity-current head because of the internal circulation within the head, thus yielding the highest concentration of oxidants near the ground.

An example of the sea-breeze system along the coasts of Texas and Louisiana is presented in Figs. 8 and 9.

FIGURE 8. The sea-breeze system along the Gulf coast of Mexico, Texas, and Louisiana. This visible imagery from the GOES satellite shows the sinking air (or subsiding and thus clearing) on both sides of the shoreline. Notice the existence of the sea-breeze front or the convergence line displaced inland from the shore.

FIGURE 9. Infrared imagery from the GOES satellite for Fig. 8. Two lines across the sea-breeze front are delineated, one in south Texas and the other in west Louisiana. They provide the horizontal temperature distribution of the cloud top at B (south Texas) and F (west Louisiana), sea surface (at D and H), and ground (at A, C, E, and G). Note that both B and F are located on the sea-breeze front. Also, there is 4   °C difference between A and C as well as between E and G, indicating the advancing of cooler air onshore associated with the sea-breeze system.

Read full chapter

URL:

https://www.sciencedirect.com/science/article/pii/B0122274105001149

thornhillactem1984.blogspot.com

Source: https://www.sciencedirect.com/topics/earth-and-planetary-sciences/sea-breeze

Share this post