The transfer of particles

The ocean surface is subject to a continual barrage of particles from the atmosphere. Some of these owe their origin to previous expulsion from the ocean surface, as discussed in §2.9.2. Others derive from chemical and physical processes within the atmosphere causing coagulation of smaller particles or gas molecules. Yet more have been swept, or thrown, up from a land surface and carried perhaps thousands of kilometres in the atmospheric circulation.

The cascade of these particles into the sea affects the chemistry of the ocean in many ways. For some of the trace constituents this will be their major source (§4.1.4). However, in this chapter we are not studying the general chemistry of the ocean, but only those aspects relevant to climate. In this context our interest will be restricted to those particles likely to affect the nutrient supply of phytoplankton or the cycling of greenhouse gases.

Particles are also being continually ejected from the ocean by evaporation, bubble bursting and spray. We have already considered the physical processes involved in this production of marine aerosol (§2.9.2); here we will consider the physical and chemical influences these particles exert on climate.

3.5.1 Aerosols, plankton, and climate

The size distributions for aerosols in a marine, and heavily polluted terrestrial, atmosphere are shown in Fig. 3.12. There are millions of very small particles of less than 0.1 ^m radius in every cubic metre, but only a few thousand particles larger than 1 ^m. The former are called Aitken nuclei, after their 1880 discoverer, while the latter are known as Giant nuclei. Those whose size lies between the two extreme sets are known as Large nuclei. The figure shows that the terrestrial atmosphere has far more Aitken and Large nuclei, but a similar number of Giants, to the marine atmosphere. The Giant nuclei are largely

Fig. 3.12. Aerosol size-distribution spectra characteristic of continental and maritime air. [Using data from Ludlam, 1980.]

10-4

10-4

, Aitken , nuclei

_ Large t nuclei

nuclei

1

/

\

Nucleu

s-rich continen

al air

/ /

\

\

/

\

Maritim componer

J \ t \

) \

\

10"10 10 s 10-0 10"7 10-® Radius in m sea salt particles, whose significant impact on climate is discussed in the next section.

The Large and Aitken nuclei consist of a mixture of many minerals and chemical compounds. Our interest here lies in those with carbon, nitrogen, sulphur, phosphorus and iron content, as these contribute to biological processes if they are deposited in the sea. Iron is mostly to be found in aerosols derived from weathering of clays or other rocks. There is a small, but biologically significant, mid-oceanic aerial component, but most of this material is relatively large in size and likely to be lost through deposition in coastal zones. Atmospheric phosphate concentrations are also very low and, unlike for iron, not thought to make a significant contribution to oceanic levels.

Nitrogeneous compounds are particularly common in combustion and bacterial decay. The resulting aerosols can be small, and while marine concentrations (0.1-0.4 ^gkg-1) are smaller than over land by an order of magnitude, they are not negligible. Remote marine areas tend to have concentrations at the lower end of this range. Some of this nitrate will be from re-cycling of material of marine origin. However, seas close to shore, or exposed to air that has recently passed over continental areas, have enhanced concentrations which will contribute towards the supply of nitrate in the surface waters of the oceans. This is particularly true in the Northern Hemisphere, where man's activities are a dominating source of atmospheric nitrate and land masses occupy a much higher proportion of the surface area. Nitrate concentrations in waters subject to active algal growth will be of the same order as atmospheric levels, suggesting that aerial input could be of importance in supplying nitrate to fuel planktonic growth. It is believed, however, that this source is considerably less important than re-cycling of nutrients within the mixed layer. In §4.2.2 we

Fig. 3.13. Number size distributions of sulphate (broken line) and organic aerosols (solid line) at an elevation of 1000 m in an exposed location on Puerto Rico in the Caribbean. [Adapted from Fig. 1b of Novakov and Penner (1993). Adapted with permission from Nature, 365, p. 824. Copyright (1993) Macmillan Magazines Limited.]

will see that the gas nitrous oxide has a marine source as well as a terrestrial source.

About half of the sub-micron-sized aerosols are composed of sulphates. Sulphate is an aqueous phase reaction product of the gas sulphur dioxide, SO2, through the two reactions

Sulphur dioxide has many terrestrial sources through combustion, volcanic emissions and biological activity. However, about a quarter of the atmospheric budget is the end result of biological activity in the ocean. This source, and the effect of sulphate on precipitation chemistry, will be discussed in §4.4.

Another significant contributor to cloud condensation nuclei comes from a range of organic carbon-based aerosols. These contribute relatively little to the total aerosol mass, but their size distribution clusters below 0.05 ^m and so they may have a greater impact on cloud droplet formation than their mass would suggest (Fig. 3.13). The particles are derived from direct emission from terrestrial combustion, or from gas-phase reactions of hydrocarbons within the atmosphere. Their contribution to the oceanic carbon budget is likely to be small, however. Their possible impact on cloud processes will be considered in the next section.

3.5.2 Sea spray, clouds, and climate

Sea salt is the largest contributor by mass to particulate material in the marine atmosphere. Between 109 and 1010 tonnes are cycled through the atmosphere each year. Residence times for particles vary from seconds to days, depending on the height to which the particles penetrate into the troposphere and the distance over which they are carried. Giant sea salt particles can travel a thousand kilometres; concentrations of such particles, chemically unaltered, can be almost unchanged hundreds of kilometres from the nearest ocean, provided no precipitation has occurred from the air mass in transit. Aerosols are returned to the Earth's surface by direct particulate, or dry, deposition, and through precipitation.

Fig. 3.13. Number size distributions of sulphate (broken line) and organic aerosols (solid line) at an elevation of 1000 m in an exposed location on Puerto Rico in the Caribbean. [Adapted from Fig. 1b of Novakov and Penner (1993). Adapted with permission from Nature, 365, p. 824. Copyright (1993) Macmillan Magazines Limited.]

Geometric diameter d (urn)

Geometric diameter d (urn)

Fig. 3.14. Sea salt concentrations as a function of wind speed. The upper (solid) line shows the best fit to some of Woodcock's observations near the surface of the sea off Oahu, Hawaii; the lower (broken) line shows earlier results of Woodcock from observations near the top of the planetary boundary layer.

Fig. 3.14. Sea salt concentrations as a function of wind speed. The upper (solid) line shows the best fit to some of Woodcock's observations near the surface of the sea off Oahu, Hawaii; the lower (broken) line shows earlier results of Woodcock from observations near the top of the planetary boundary layer.

The latter process is known as wet deposition. The importance of sea salt for weather and climate was not appreciated until the 1940s when Woodcock began a series of pioneering projects, culminating in his presentation of the theory of sea-salt-driven formation of raindrops in 1952. This theory is highly relevant to climate change and will be discussed in §3.5.3, taking into account more modern views on the importance of sub-micron, non-sea-salt particles for complete explanation of the observed droplet spectrum.

Over the oceans, in the planetary boundary layer, it is found that the concentration of salt particles in the atmosphere is strongly dependent on wind speed: the stronger the wind speed the more sea salt there is in the atmosphere. This is shown in Fig. 3.14, a summary of observations near Hawaii. The size distribution of particles is also dependent on wind speed, although the tendency for many very light, but few heavy, particles is not. Fig. 3.15 illustrates this size-weight-wind speed relationship in the same experiment.

Within the boundary layer over the ocean there is often a sea salt inversion. At moderate wind speeds a rise in sea salt concentration occurs between 300 and 600 m. This peculiar phenomenon occurs both in the presence and absence of clouds. There are at least two possible mechanisms that might explain the inversion. Beneath clouds any fine drizzle often evaporates. This could leave the salt condensation nuclei at the evaporation level. Vertical wind shear could then allow the clouds to move faster than the sub-cloud air, leaving the salt-enriched air behind. Another mechanism involves the transport of sea salt from the surface into a more humid part of the boundary layer. Sea salt is hygroscopic, so the particles will grow in such an environment. Eventually some particles will become sufficiently large so that their fall speed, due to their mass, will be similar to the updraught raising them from the surface. Salt particles will then accumulate at this level of stability.

The hygroscopic nature of sea salt aerosol, its generally large size, and its abundance, are the principal factors giving this aerosol its climatic influence. Sea salt particles are large: greater than 0.1 ^m, and very often 1 ^m, in diameter.

Fig. 3.15. Variation of sea salt particle cumulative distributions with wind speed, from Woodcock's observations at the top of the planetary boundary layer.

Fig. 3.15. Variation of sea salt particle cumulative distributions with wind speed, from Woodcock's observations at the top of the planetary boundary layer.

Aerosols visible through a microscope are predominantly composed of sea salt, which is largely sodium chloride, NaCl. At relative humidities greater than 75% NaCl particles become hygroscopic and attract water vapour to form small droplets. Other salts can become hygroscopic at much lower relative humidities - potassium carbonate, K2CO3, at 44%, for instance - but the abundance of NaCl makes this a significant hygroscopic aerosol.

Other, significantly smaller aerosols are also important in the formation of cloud droplets. Sulphate particles, derived from biological or combustion sources (Chapter 4), form the major aerosol in the size range just below 1 ^m radius (see Fig. 3.13).2 Even smaller sizes (below 0.1 ^m) have been found to be predominantly organic aerosols, some with an iodine component (§4.2.4). Both of these source materials will actively contribute to cloud droplet formation, and are likely to dominate it, in terms of total number. Nevertheless, the larger droplets of sea salt origin have a crucial role in the precipitation process, as will be seen in §3.5.3.

Hygroscopic particles will make the atmosphere hazy, but it is only when the relative humidity reaches 100%, and the air becomes saturated with respect to water vapour, that clouds or fog appear. This is because hygroscopic particles, in absorbing water vapour, deplete the vapour concentration of the surrounding air. In any mix of particles this will mean that only a portion will be actively

2 Sulphate aerosols have several sources. Some, particularly of larger size, are derived from the sulphate salts of seawater spray (see Table 1.4). The remainder are termed non-sea-salt sulphate (or nss sulphate) aerosols. In general in this book, when sulphate aerosols are referred to they will be of nss origin.

growing; the rest will have local microclimates with humidities below their threshold growth level. When the air is at saturation, or, as is common in clouds, slightly super-saturated, sufficient moisture is present for an optically dense mix of droplets to be generated. There will, however, still be a significant proportion of non-growing particles in such a mix.

Hygroscopic particles are essential for cloud formation in the atmosphere. It would require super-saturation humidities of 220-340% for spontaneous, or homogeneous, nucleation of water droplets from vapour alone. In providing a hygroscopic surface the aerosols allow heteorogeneous nucleation to occur in natural conditions. For small particles (less than 1 ^m in diameter) such as nss sulphate or organic aerosols, and very marginal super-saturations, the radius of curvature of the forming droplet is important in determining growth. For larger, sea salt, droplets the rate of growth in the droplet radius, r, is a simple function of r and the level of super-saturation, SS:

dt r

In (3.12) G is an almost linearly increasing function of temperature. At 0°C G = 6 x 10-9 m2s-1, and 20°C G = 1.23 x 10-8 m2s-1. This simple formula is applicable to most cloud formation processes, where moisture is supplied by motion external to the cloud. For fogs, and possibly clouds forming through internal (radiative) cooling, smaller droplets are more common and the radius of curvature of the droplets must be considered in studying the evolution of such clouds.

Condensation of water vapour around the larger aerosols is a necessary part of cloud formation leading to rainfall, as we shall see in §3.5.3. Because salt particles form the majority of the Giant condensation nuclei over the 70% of the globe that is covered by ocean (Table 2.2), such aerosols play a key role in the production of marine rainfall. Note, however, that the majority of the Large condensation nuclei over the ocean are sulphate particles (see Chapter 4). Over land there are many other sources of nuclei, from dust raised by the wind to particulates emitted by factories or formed by reactions within the terrestrial atmosphere. Cloud formation over the oceans requires strongly hygroscopic condensation nuclei, as the particle concentrations are rather lower than over continental interiors. Within marine clouds there are perhaps 20-200 droplets per cubic centimetre, while terrestrial clouds contain 200-2000.

Clouds can form for a variety of reasons; each involves bringing air to saturation. The most familiar perhaps is the ascent of moist air because of local convection, typically seen on a warm summer's day. Air warmed by a patch of ground hotter than its surroundings - perhaps a bare field, a city, or an island - becomes less dense and is forced to rise to seek its equilibrium density level. Rising air cools because it expands as it moves into regions of lower pressure. This parcel of air, sometimes known as a thermal, rises fast enough that there is little mixing with its surroundings. It can be shown that air rising without the gain or loss of heat, or adiabatically, loses temperature at a rate of 9.8°C/km. If the ascent is allowed to proceed unhindered the air will eventually reach saturation (see Fig. 2.5) and cloud formation will begin. Clouds formed in this way often display signs of the vigorous upward motion by

producing distinct puffs of cloud: these are characteristic of cumulus clouds (Fig. 3.16).

Ascent of moist air can also be created by large-scale dynamical influences associated with fronts. This can occur over short distances, giving rise to large cumulus clouds, or over hundreds of kilometres, when layered clouds result. Air forced to rise over mountain ranges can also give rise to cloud, if the mountains are high enough and there is sufficient moisture content in the air.

Clouds can also be formed by cooling of the air in situ. This tends to produce low level, layer clouds, but in unstable air low level cooling can contribute to cumulus development when the cooled air mass passes over a warmer surface. This happens if air flowing over a cold sea surface moves over warmer land. The air has both cooled and gathered water vapour so that when it is made more buoyant by heating from the warm land it can rise to produce cumulus cloud. Such cloud formation can occur readily around the North Sea or the west coast of North America in the spring, when the land is warming but the sea surface is still cold.

Another mechanism for producing cloud, encountered in §2.11, is the input of moisture from a warm sea surface into cold air. This can produce low cloud or fog.

In our discussion so far we have assumed that clouds are composed of water droplets alone. Ice clouds, or mixtures of water and ice, have been neglected.

Fig. 3.17. Variation of saturation vapour pressure

Fig. 3.17. Variation of saturation vapour pressure with temperature, over liquid water (solid line) and over ice (dotted line). [Using data from Ludlam, 1980.]

to to with temperature, over liquid water (solid line) and over ice (dotted line). [Using data from Ludlam, 1980.]

to to

High clouds, or clouds at high latitudes, are, however, likely to be composed of ice crystals or snowflakes rather than water droplets. Aerosols that allow ice to sublime directly to the particle tend to be of terrestrial, rather than marine, origin; clays and heavy salts are most active as ice nuclei. Nevertheless, marine aerosols play a part in producing ice through the freezing of water droplets. Depending on the humidity and the size of the drop, water can exist in its liquid phase at sub-zero temperatures. In laboratory conditions tiny supercooled droplets can exist at temperatures of -40°C, but even in the atmosphere supercooling to below - 10°C is not uncommon. Once a droplet has been frozen by local cooling or through transport to a colder environment, the new ice crystal will begin to grow through sublimation. This is because the saturation vapour pressure for ice is lower than for water; air saturated with respect to an ice surface can be under-saturated with respect to water. This is shown in Fig. 3.17. In the upper troposphere the concentration of sea salt particles is drastically reduced from its boundary layer concentration. However, the necessity for saturation with respect to water, rather than ice, before cirrus clouds form suggests that instantaneous freezing of water droplets forming about salt, nss sulphate or organic carbon nuclei is an important mechanism for high cloud formation.

Stratospheric clouds, known to be important as sites for chlorine activation in the chemical process resulting in depletion of high level ozone, have sulphate particles as their main source of condensation nuclei. Most of these aerosols are thought to be of terrestrial origin, but as Chapter 4 will discuss, there are important marine sources of sulphate which may contribute to the stratospheric loading via entrainment of tropospheric air into the stratosphere in frontal regions or the ITCZ.

Clouds contribute to the climate in two ways. They play an important role in the radiational balance of the Earth. They are also the source for precipitation. We have encountered the radiational aspects of cloud physics several times in earlier chapters. Clouds reflect, scatter, and absorb solar radiation. They also reflect, absorb, and re-radiate the Earth's radiation. Different types and thicknesses of clouds will affect the radiation balance in different ways. High clouds will have less water content, and thus will be optically 'thinner', allowing more transmission of solar radiation. They are also cold, thus emitting less infrared radiation to space than a lower, warmer cloud. The cloud water content can have a double-edged contribution to climate: high water content makes a cloud a better solar reflector, but it will also act as a stronger greenhouse absorber. High water content also warms the cloud by releasing greater quantities of latent heat. Such warming is an example of diabatic heating. In Chapter 7 we will see how the reaction of cloud physics and chemistry to greenhouse gas emissions is a major area of uncertainty for the prediction of climatic change.

3.5.3 Mechanisms for precipitation formation

Precipitation occurs from clouds composed of water droplets or ice crystals. The previous section showed how, in a cloud of water droplets, differential growth rates will operate for different drops, depending on the local moisture content of the air and the radius of the drop. Very small droplets near larger droplets will have difficulty growing because their large radius of curvature increases the saturation vapour pressure compared to that for larger drops. This is illustrated in Fig. 3.18. Once a few larger drops form, therefore, they will scavenge vapour from nearby small droplets by forcing the latter's local humidity below saturation. Even if there is competition for vapour between two relatively large drops, the larger will capture more vapour. This is because the rate of change of volume, V, for large drop growth is a linear function of radius:

Note that the typical distance between cloud droplets may be several hundred droplet radii, that is 0.1-1 mm.

The larger drops will eventually become heavy enough for their gravitationally-induced fall speed to become greater than the ascending motion within the cloud. For example, a drop of radius 10 ^m has a fall speed of 1 cms-1. As these drops fall through the cloud they will grow by sweeping up smaller drops in their path. If they gain enough mass the drops will become unstable and splinter into a number of smaller, but still relatively large, drops. These will continue to descend, capturing more and more water, perhaps undergoing repeated shattering, until they fall out of the cloud. This process of coalescence is shown schematically in Fig. 3.19. The resulting raindrops will then reach the ground as precipitation, unless evaporation below the cloud base converts the drops back to vapour before impact. In the latter case an observer will see trails of rain seemingly suspended beneath the cloud in what are known as virga. The time required for rain to form depends on the initial size distribution of the droplets and the water content of the cloud. The fastest rates of 10-30 minutes occur in very moist air with few condensation nuclei - and hence large droplets. Large numbers of nuclei with relatively small moisture levels lead to clouds which will never reach the stage of precipitating. The warm marine atmosphere of the tropics is ideal for potential precipitation.

Clouds in the winter mid-latitudes and polar regions will be a mix of a few ice particles and a large number of tiny cloud droplets. We have seen that this is an unstable situation because of the saturation vapour pressure difference over

Fig. 3.18. Variation of equilibrium relative humidity with droplet radius, r, for a fixed mass of sodium chloride in water, for three different masses. Note that the scale has been expanded above 100% relative humidity to clarify the structure of within-cloud processes. [After Fig. 6.13 of Mcllveen (1992). Reproduced with permission of Chapman and Hall from Fundamentals of Weather and Climate by R. Mcllveen (1992).]

Fig. 3.18. Variation of equilibrium relative humidity with droplet radius, r, for a fixed mass of sodium chloride in water, for three different masses. Note that the scale has been expanded above 100% relative humidity to clarify the structure of within-cloud processes. [After Fig. 6.13 of Mcllveen (1992). Reproduced with permission of Chapman and Hall from Fundamentals of Weather and Climate by R. Mcllveen (1992).]

Fig. 3.19. Schematic illustration of the coalescence process.

ice and water, causing the ice particles to grow preferentially. Eventually the large ice crystals will become large enough to precipitate as snow, or rain if the temperature of the lower atmosphere is somewhat above freezing. This is known as the Bergeron-Findeisen process of rain formation, after the Norwegian scientists who proposed the mechanism in the 1930s. At one time it was thought to be the major mechanism for rain production, but it is now realized that the coalescence process is more widespread. Even within cold clouds, where the Bergeron-Findeisen process is operating, collisions between ice crystals and water droplets add mass to the crystals much faster than preferential vapour attraction. The Bergeron process thus requires a moist atmosphere with adequate supplies of ice and water nuclei. Continental margins would seem to be the most likely areas for these conditions to be found.

3.6 Photochemical reactions in sea water

There are innumerable atmospheric photochemical reactions in which molecules are fragmented, or given different reactivity by changing their energetic state. In Chapter 1 we encountered the climatically important photolytic formation of ozone in the stratosphere (equation 1.4). In Chapter 4 we will see how a photolytic reaction removes hydrogen sulphide from the atmosphere rapidly, and also how sulphur dioxide is photolytically oxidized in forming sulphate aerosols. Reactions driven by the energy of a photon of appropriate energy are also important in the production of smogs over industrialized cities.

In the ocean the possibility of photolytic reaction is often neglected, yet Chapter 4 describes the consequences of just one: photosynthesis. In the right conditions appreciable amounts of solar radiation can penetrate tens of metres beneath the surface of the sea (Fig. 2.2). Sea water preferentially absorbs the longer, less energetic, red wavelengths of the visible part of the solar spectrum. Thus the shorter, high energy, visible wavelengths provide much of this penetrative power. Ultra-violet light, however, is also strongly absorbed and scattered. In the atmosphere ultra-violet light is a potent driver of many photochemical reactions; most of the solar photons in this energy band have been absorbed in the atmosphere before the solar radiation reaches the ocean surface; what remains is rapidly removed in the upper few metres.

Chemical reactions driven by photons can take a number of different forms. One of the most important is photo-dissociation: a fragmentation of the chemical species upon interaction with a photon. The photon must have sufficient energy to break the chemical bond; photons with greater energy additionally impart kinetic energy to the resulting fragments. One of these fragments is typically in an excited state, that is, its electrons will occupy orbitals above the atom's (lowest energy) standard electron configuration (see Appendix B). A special case of this type of reaction is known as photo-ionization. This occurs if the fragments from the reaction are not discrete atoms or molecules but merely an electron and the reacting atom's or molecule's positively charged ion.

Absorption of a photon can lead to the specie radiating energy in order to return to its standard energy state. This is called luminescence, but particular types of this reaction are known as fluorescence or phosphorescence, depending on the type of energy transition. Such a reaction concentrates radiation from a wide spectrum of incident wavelengths to one, making this one emission wavelength particularly intense. Both of these reactions are seen in sea water; fluorescence can be used as a chlorophyll tracer, chlorophyll being one of the products of photosynthesis (see §4.1.3).

Photons absorbed by a specie do not always result in fragmentation or luminescence. The raised energy state of the specie can be relatively stable, giving it a different reactivity and chemistry. The energy can also be given to other molecules via collision, converting the energy into kinetic energy, molecular vibrations, or transferring an excited state to the colliding molecule. Excited

Table 3.4. Climatically important photochemical reactions in sea water

NO2- + H2O + photon ^ NO + OH + OH-NO3- + photon ^ NO2- + O CH3I + C1- + photon ^ CH3CI + ICO + photon ^ CO* (an excited, more reactive state of CO) H2S + OH- + photon ^ H2O + HS (CH3)2S + 5O2 + photon ^ 2CO2 + 2H2O + H2SO4

molecules or atoms often have greatly enhanced reactivity, so that they react with other molecules very rapidly.

There are a number of photochemical reactions in the upper layers of the ocean that influence the concentrations of climatically sensitive gases, in addition to those mentioned at the beginning of this section. Nitrate, NO3-, and nitrite, NO2-, two nutrients important for phytoplankton growth, can be photolysed to produce different oxidation states. A number of trace gases produced in the ocean, such as methyl iodide, CH3I, dimethyl sulphide, (CH3)2S, carbon monoxide, CO, and hydrogen sulphide, H2S, will all undergo photochemical transformation in sea water as well as the atmosphere. A summary of presently known reactions is given in Table 3.4. This is a field of growing interest and new, climatically important, reactions are likely to be added in the future, for instance the photochemistry of organic iron colloids (see §4.1.1).

These photochemical reactions complicate the estimation of fluxes of gases to the atmosphere, as they introduce a diurnal cycle that may enhance or counter any similar biologically driven cycle. Observational study is as yet insufficient to determine the importance of the diurnal cycle in many air-sea gas exchanges.

This brief discussion of photochemistry provides useful background when considering biological processes in Chapter 4 and the atmospheric photolytic reactions important in the stratospheric ozone problem in Chapter 7.

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