|The Courier N° 133 May-june 1992 - Dossier : Environment and Development - Country Reports - Côte d'Ivoire - Papua New Guinea (EC Courier, 1992)|
|Dossier: Environment and development|
by Pascale DELECLUSE
Oceans are essential to shaping climate, be it over seasons, decades or millenia. The vast pool of energy constituted by the oceans shapes the way climate develops and modern research focuses on determining and understanding the complex movements which lead to the mass distribution of heat and salt observed in the ocean today. The machinery of exchange on the ocean-atmosphere interface upstream of these ocean movements gives the different masses of water their physical, chemical and dynamic properties.
It is at their surface that oceans exchange with the atmosphere the matter and energy which determine their movement. Energy supply from sources other than the atmosphere (thermal sources on the seabed, for example) is negligible. The atmosphere is very largely transparent to solar radiation-which is absorbed by the surface strata of the ocean, which radiates in turn, like a black body, towards the atmosphere. So the surface strata are warmer than the atmosphere, overall, and there is a substantial transfer of heat and latent heat from ocean to atmosphere. Latent heat is transferred when the surface waters evaporate, affecting not just the temperature of the ocean but salinity too. The other factors which affect salinity are rainfall, ice formation in the polar seas and, to a lesser extent, water from rivers. So salinity and temperature are determined at the surface by the multiple exchanges between ocean and atmosphere and the degrees they reach can only then be altered by advection or diffusion within the ocean. We shall therefore now look at ocean currents.
The ocean can store a great deal more heat than the atmosphere (a
2.5m column of water, for example, has the same thermal capacity as the whole of
the atmospheric column above it). The net
ocean heat budget is positive in quite specific regions (the tropics, particularly the east of the Pacific and of the Atlantic, accumulate a large amount of heat), while the warm western boundary currents (the Gulf Stream and Kuroshio) and high latitudes send back a large flow of heat towards the atmosphere. But the distribution of springs and wells is very spatially and temporally heterogeneous.
The ocean gets its mechanical energy on the surface too, from wind friction. Ocean circulation, which is very active, is measured in Sverdrups (million m³ per second-a current such as the Gulf Stream reaches 90). Circulation redistributes the heat world-wide and makes the ocean a powerful regulator of climate, tempering the zonal temperature gradients. It makes up for unbalanced thermal flows by transporting as much heat from the equatorial regions to the poles as the atmosphere-several Petawatts (or 10 to the power of 15 watts).
Elements of ocean physics
The ocean can be roughly divided into three layers-the well-mixed surface layer, the main thermocline (in which the temperature fades noticeably) and the deep waters-which shape the climate on different time scales and are distributed.
Main, seasonal thermocline
The seasonal development of the ocean-atmosphere pair is controlled by the well-mixed surface layer, which stores heat during the summer and returns it during the winter, tempering the atmospheric conditions. In seasons in which warming occurs (spring and summer in the northern hemisphere), the heat which the ocean receives from the atmosphere builds up in this layer-whose physical properties (temperature and salinity) are homogenised vertically by strong, turbulent vertical diffusion. The depth of this layer varies considerably with region and season (but 200m is a reasonable figure) and it is separated from the underlying layers by a zone of strong thermal gradient which isolates the sub-surface ocean waters. The ocean communicates with the atmosphere solely through this surface layer, which accumulates kinetic and thermal energy. Many digital ocean study models have been designed to include this ocean layer at the bottom of a general atmospheric circulation model. This is the first way of describing the thermal inertia of the ocean in the development of the climatic pair.
In autumn (in the northern hemisphere), this surface layer gets deeper under the influence of wind and cooling and dissipates part of the energy built up over the previous seasons to the atmosphere in the form of heat and latent heat. The thermal energy is distributed over greater depths. The maximum depth, reached during the winter, corresponds to the position of the principal thermocline, the permanent thermal gradient zone in the oceans. This is when the process of dispersal of the thermocline may occur. In a region where the vorticity of the wind friction on the surface is negative (i.e. no waves), the Ekman spiral takes surface waters into the main thermocline with all their properties-temperature, salinity and vorticity-intact, a process which generates horizontal circulation on an ocean basin scale.
Only part of the thermocline is distributed each year and the circulation into which these waters are drawn is complex. Balance takes several decades (the typical adjustment time for a basin in the temperature latitudes) and involves many interactions between wind-driven circulation and density-driven circulation. We shall now look at the structure of the polar and tropical regions before moving on to the global scale.
In the tropical oceans, seasonal thermocline and principal thermocline are one and the same, because very vigorous local dynamics make for efficient vertical diffusion to the principal thermocline, which is very shallow and near to the surface (50-200m) in these parts. Note, however, that the concept of the wellmixed surface layer may be based on the salt stratification in some regions (e.g. the western parts of the tropical Pacific). A shallow, clearly defined mixed layer in a zone where the Coriolis force is nil means that equatorial basins can adjust to atmospheric changes in only a few months, because the tropics develop wave-guiding properties. The speed of adjustment means that there can be an efficient pairing of atmospheric circulation and ocean circulation over a few months, with spectacular reversals of tropical circulation (which we shall be discussing when we get to the major natural variability of the climate-El Nino).
The mixed layer also has a special part to play at high latitudes. During the winter in the hemisphere under scrutiny, very harsh conditions on the surface lead to the formation of only very slightly stratified masses of water, sometimes going right down to the ocean floor. The very cold air masses at these latitudes lead to cold, dry winds blowing over the ocean. The temperature difference between sea and air may then be more than 10 degrees, with intense heat transfers from ocean to atmosphere. Between - I and-2 degrees, sea water freezes at the surface, throwing back its salt content into the ocean. Variations in the volume of sea water are then controlled by salinity and not temperature. Excess salinity leads to the formation of dense water, which is heavier than the surrounding waters and sinks to the bottom - producing deep convection. The masses of water thus formed drop to the level of water of equal density. They fill the bottom of all the ocean basins-75% of the world's waters (more than 10 to the power of 9 cubic kiLométres!), although the areas in which they are formed only account for a few thousandths of the ocean surface. Deep convection processes, which are intermittent and localised, occur in difficult atmospheric conditions and therefore tend not to be under direct observation at the present time. The deep waters of the ocean have a fabulous storage potential which takes a very long time to respond-centuries, a thousand years even-so deep circulation is unlikely ever to be stationary.
Average ocean circulation
A major difficulty in studying ocean circulation is that of describing quantity and quality. In some, well-observed basins (the North Atlantic, for example), it is perfectly possible to map monthly temperature and salinity variation, but there is no exhaustive description of, circulation. Direct measurements are rare and widely separated in space and time.
Furthermore, average circulation is strongly affected by turbulence and is difficult to determine. At temperate latitudes, eddies cover an area equal to the internal radius of turbulence (50 km) and last for about a month. Their addition to the normal heat transport in the ocean may, in some parts, double the average figure. They are a potential major contributor to climatic balance, although estimates are still very hypothetical. Turbulence is very unevenly distributed over the surface of the oceans. Although the rings of the Gulf Stream were the first eddies to be detected, this type of activity occurs in all parts of the globe, albeit with varying morphology and energy levels. Satellite measurement is a very useful way of monitoring the global behaviour of the kinetic energy of turbulence. How exactly does it contribute to the general movement of the oceans ? That is still the question.
How oceans affect climate
In 1983, serious climatic anomalies occurred over the whole of the Pacific. There were torrential rains in Peru, forest fires in Australia and devastating hurricanes in Polynesia - not isolated events, but ones which hit a peak of particular intensity that year, with dramatic economic consequences. The ENSO (El Nino Southern Oscillation) anomaly occurs at intervals of 2-10 years, a three-year gap being the most common, I and, since it is the clearest manifestation of the natural variability of climate on an interannual scale, is of great interest to many oceanographers and meteorologists.
Southern Oscillation and El Nino are the manifestations in atmosphere and ocean of the same climatic anomaly- which, in its fully developed phase, takes the form of an extension of warm water throughout the tropical parts of the Pacific Ocean and a reversal of the sea slope. The ocean thermocline goes down I tens of metres in the eastern part of the basin and the currents are reversed, with strong eastward currents appearing on the surface and the Equatorial Undercurrent diminishing, possibly to nothing. In the atmosphere, the trade winds drop, often to the point where they reverse in the western part of the basin. The convection zone, usually over the islands, of Indonesia, moves eastwards and so rainfall declines in the western Pacific and , increases in the central and eastern parts.
The two zones in which the surface winds of the Pacific come together merge into a single convergence zone over the centre.
Meteorologists say that the trade winds weaken because of a change in the distribution of warm water in the tropical parts of the Pacific, with the east-west temperature gradient declining and the Walker cell no longer maintained. As this - warm water moves to the mid-Pacific, the zonal winds converge towards the warm anomaly-which is why the winds in the western parts of the Ocean reverse.
Oceanographers say that the appear- ' ance of warm waters in the central and eastern reaches of the Pacific has something to do with the waning of the trade winds which, under normal circumstances, would be maintaining an upwards slope from east to west. As soon as the trade winds stop, there is nothing to maintain the slope and so the warm waters flow over the whole of the equatorial region. The Equatorial Undercurrent wanes and the west wind creates surface ocean currents moving eastwards.
Neither atmosphere nor ocean can develop the anomaly independently of the other. Both elements are closely linked by easily destabilised interactions. An anomaly in the surface temperature of the ocean will immediately trigger an anomaly in atmospheric circulation- which will then maintain and develop the initial ocean temperature anomaly and vice versa. But the response is not confined to the zone of the original anomaly, because the equatorial waves soon spread along the equator, bringing the whole tropical ocean reservoir into play for more than a year.
North Atlantic Deep Water circuit
The uneven distribution of temperature and salinity for a given level of, pressure generates thermohaline circulation-important in view of the masses which it shifts and the time it takes to adjust. The circuit which Gordon put forward (1986) for inter-ocean exchange illustrates the space- and time-scales I involved here. North Atlantic Deep Water (NADW) is formed in very restricted convection zones-the Labrador Sea and on the edge of the Greenland Sea and the Norwegian Sea. The water mass flows southwards by the deep western border current to mix with water from the Mediterranean. It influences the Atlantic basin to a depth of more than 100m, flowing into it gradually, mixing with the neighbouring waters and upwelling slowly. In the southern parts of the Atlantic, it is taken eastwards by the Antarctic Circumpolar Current, with powerful upwelling around the Antarctic, and ends up in the intermediate Antarctic waters which flow into the thermocline. The return journey to the Atlantic is mainly via the 'hot route' through the tropics.
The waters of the thermocline in the Pacific flow between Indonesia and Australia into the Indian Ocean. This region, a maritime continent, has heavy precipitation and vertical mixes on the passage between the islands. The water we are monitoring gets warmer and loses a lot of its salinity (the minimum level, clearly, is on the Indian Ocean crossing between isotherms 10 and 20 deg.C). It follows the South Equatorial Current and gradually joins the much more saline, thermocline waters from the southern part of the Indian Ocean. Before Madagascar, a major branch of the mass of water veers southwards to feed the Agulhas Current, most of which turns into the Antarctic Circumpolar Current by a process of retroflexion. Part, however, gets to the South Atlantic, in the form of warm eddies which are then drawn into the sub-tropical anticyclone circulation of the South Atlantic, joining a major supply of warm water to the South Atlantic. Some of these waters are drawn along by the South Equatorial Current, cross the equator north of Brazil and then move northwards in the Guyana Current, which joins the Gulf Stream. From there, they gradually take up the warmth of the tropical latitudes and leave it behind them again as they move northwards.
So it is perfectly possible for a particle of water to circumnavigate the globe. But it will change many times in the process. NADW starts cold and with medium salinity and warms up gradually as it moves south. It will take almost 1000 years to reach the Antarctic Circumpolar Current and be drawn into the thermocline waters. When in the intermediate Antarctic waters, its salinity and temperature go up when it mixes with the surface waters. The 40 000 km 'hot route' in the thermocline is much faster because, with speeds in the intermediate layers of 0.10 to 0.01m per second, the journey back to the deep water formation zones takes 13 to 130 years.
This potential journey of the waters formed in the North Atlantic points up the complexity and difficulty of understanding global ocean adjustment, because the time-scales are extremely large and anything may happen on the way. Two processes (which happen often when compared with thermohaline circulation) may affect it. Fluctuations lasting 10 years or so may occur in the sub-polar regions and strongly affect the quantity and quality of NADW. Circulation may be seriously disrupted or even reversed during inter-glacial cycles and the whole circulation of the ocean thermocline in the Indian, Pacific and Atlantic Oceans may be affected 1000 years later. Moreover, links between the oceans are very sensitive to changes in the distribution of atmospheric wind and pressure fields. Displacement of the wind vortex south of Africa can alter the movement of water into the Atlantic Ocean. Similarly, not much is known about the move through the maritime continent region and it is seemingly influenced by ENSO anomalies. Oceanic thermohaline flow, which is the result of the variations in temperature and salinity in the oceans, is still largely hypothetical. It is often thought of as being stable and permanent, though largely because it has not been observed for long rather than because of any proper investigation of its balance. A local change has general repercussions which we are in no position to assess today. But we do know that other states of circulation have existed and that there have been very strong variations in regional temperatures and sea-levels.
Over shorter time spans, clearly, the ocean-atmosphere pair may oscillate between various states of equilibrium. ENSO affects weather conditions in the tropical Pacific for more than a year and triggers variations of tens of centimetres in sea level. Amplification of a small disturbance leads to enormous energy and heat transport over a whole ocean. The difficulty of understanding a system involving a climatic pair comes from the many ways in which the two can relate and interrelate and from the variations in the time which responses may take. The ocean, in particular, leaves a delayed but inevitable stamp on the climatic system and it will take the intensive development of observation and digital simulation to grasp it. P.D.